Aeolian Sediments on the Northern Tibetan Plateau

Aeolian sediments on the northern Tibetan Plateau

Von der Fakultät für Georessourcen und Materialtechnik

der Rheinisch-Westfälischen Technischen Hochschule Aachen genehmigte

Habilitationsschrift

von

Dr. rer. nat. Georg Stauch

Gutachter: Univ.-Prof. Dr. rer. nat. F. Lehmkuhl

Univ.-Prof. Dr. rer. nat. R. Mäusbacher

Prof. Dr. L. Owen

Tag der Habilitation: 21. Juni 2016

Diese Habilitation ist auf den Internetseiten der Universitätsbibliothek online verfügbar.

In memory of Andreas Stauch (03.03.1977 – 16.09.2012).

Content

Summary ...... III 1 Introduction ...... 1 2 Current state of research ...... 5 2.1 The Asian monsoon system ...... 5 2.2 Aeolian sediments and palaeoclimate reconstructions ...... 11 3 Short summaries of the articles ...... 16 4 Local case studies from the northern Tibetan Plateau ...... 26 4.1 Environmental changes during the late Pleistocene and the Holocene in the Gonghe Basin, north-eastern Tibetan Plateau ...... 26 4.2 Aeolian sediments on the north-eastern Tibetan Plateau ...... 45 4.3 Interaction of geomorphological processes on the north-eastern Tibetan Plateau during the Holocene, an example from a sub-catchment of Lake Donggi Cona ...... 69 4.4 Landscape and climate during the late Quaternary on the northern Tibetan Plateau ...... 91 5 Synthesis of OSL ages from the Tibetan Plateau ...... 115 5.1 Geomorphological and palaeoclimate dynamics recorded by the formation of aeolian archives on the Tibetan Plateau ...... 115 5.2 Multi-decadal periods of enhanced aeolian activity on the north-eastern Tibet Plateau during the last 2ka ...... 138 6 Conclusions ...... 151 References ...... 156

Summary

In the frame of this thesis a new conceptual model was developed for the application of aeolian sediments for palaeoclimate reconstruction on the northern Tibetan Plateau (TP). It was highlighted that the previous state of the environment is an important parameter for the interpretation of aeolian sediments. Late glacial to early Holocene sediments indicate increasing moisture, while reactivation in the late Holocene is caused by a reduction in precipitation. The model was used to reconstruct the palaeoclimate in three basins on the northern TP (Gonghe Basin, Donggi Cona Basin and Heihai Basin). However, the complex sediment cascade with varying source areas through time, different topography, and frequent erosion requires a detailed basin-wide analysis of the transportation pathways and sediment deposits. Erosion is frequently occurring in the different terrestrial archives on the northern TP, resulting in the need for a relatively large number of sections. In the three different basins the entire sediment cascade of the aeolian system was reconstructed. Each system has a slightly different configuration, enabling a detailed understanding of the systems behaviour. The Gonghe Basin is dominated by large fluvial terraces of the Yellow River, providing vast amounts of material for aeolian processes. In combination with the width of the basin, high aeolian activity results in the formation of active dune systems and a well-developed spatial separation of different types of sediments. The Donggi Cona Basin is dominated by a large lake. Here, lake level variations have a profound influence on sediment availability. High accumulation rates of aeolian landforms are not directly related to climate variations in this case. However, the timing of sediment mobilisation and accumulation is a good indicator of climatic changes. A clear separation in the local geology in the Heihai Basin enabled a detailed reconstruction of different local transportation pathways. Furthermore, all three study areas are located on an east-west transect, following a reduction in present day precipitation in the area. An analysis of the aeolian sediments in all three areas showed a time-transgressive penetration of the Asian Summer monsoon (ASM) on the northern TP from the east during the late glacial and the early Holocene. A reverse movement was observed for the retreat of the ASM in the late Holocene.

Besides the local case studies, two meta-analyses for aeolian sediments were conducted. The first one is including all published OSL (optical stimulated luminescence) ages from the TP for the last 21 ka and identified several millennial scale climate changes. Dry and cold glacial times were followed by climate amelioration during the late glacial as a consequence of the post-glacial strengthening of the Asian summer monsoon. The mid Holocene was characterized by highest moisture conditions during a fully developed monsoonal system. Aeolian processes were strongly reduced and fluvial reworking occurred. During the late Holocene, aeolian activity increased again, first in the western part of the study area and later at the eastern side of the TP. A second study analysed multi-decadal climate fluctuation on the north-eastern TP. Six phases of enhanced aeolian activity with a duration of 80 to 200 years could be linked to phases of low temperature of the northern hemisphere, reduced total solar irradiance and a reduced ASM. The results from this study show that aeolian sediments are a suitable proxy for palaeoclimate reconstruction on the northern TP despite that most of the studied archives were influenced by erosional processes.

III

1 Introduction

Aeolian sediments are an important geo-archive and are frequently used to reconstruct past geomorphological processes and related climate changes. Large variations in the type, composition and preservation of aeolian sediments exist. In contrast to many other archives, they can be directly dated by optical stimulated luminescence. The dating methods captures the last exposure to sunlight and therefore, with some limitations, the last transportation process. This is particularly valuable in an area like the northern Tibetan Plateau (TP). Geo-archives are in many cases only partly preserved on the TP and in the surrounding dry lands. Large uncertainties regarding the interpretation of the proxies from these archives exist. However, the area was identified as especially sensitive to climate change (Colman et al., 2007; Y. Wang et al., 2010; An et al., 2012). The northern TP is located at the interplay between the Asian summer monsoon (ASM), the Asian winter monsoon (AWM) and the mid-latitude westerlies (Böhner, 2006) (Fig. 1.1). Despite intensive research during the last two decades, considerable differences occur among reconstructed palaeoclimate variations. These are for example, uncertainties in the first increase in perception related to the onset of the ASM after the dry and cold glacial times. In the Qinghai Lake Shen et al. (2005) reconstructed an early onset at around 17 ka, while An et al. (2012) assumed increased precipitation only at the beginning of the Holocene. Divergent interpretations also exist regarding the maximum strength of the ASM in the area. Several reconstructions indicate an early Holocene moisture maximum (Jin et al., 2015; Wang et al., 2015; Thomas et al., 2016) while other show a mid-Holocene maximum (Shen et al., 2005; Mischke et al., 2008; Liu et al., 2013a). As a consequence there is a considerable debate if the area was/is influenced by the Indian summer monsoon (ISM) or the East Asian summer monsoon (EASM), the two most important subsystems of the ASM (Y. Wang et al., 2010; Liu et al., 2015). Similar uncertainties exist regarding the onset of the late Holocene dry period as well as the timing of short term climate change during the last 2ka. However, the timing of the late Holocene climate changes are of special interest, as they presumably had an impact on Chinese civilizations (Zhang et al., 2008; Dong et al., 2016).

Fig. 1.1: The Tibetan Plateau and the main atmospheric systems (according to Böhner, 2006; Yao et al., 2012). The northern limit of ASM according to Chen et al. (2010; green line), An et al. (2012; red) and P. X. Wang et al. (2014; blue). The white box indicates the study area. 1

Some of the divergence in the interpretation of the archives might be related to different interpretation of the proxies. One of the main archives used on the TP are lake systems. Comparisons of different lakes on the TP showed considerable spatial and temporal variations (Wischnewski et al., 2011; Doberschütz et al., 2014). In many studies it is assumed that sedimentation occurred continuously through time, but with varying sedimentation rates. However, sedimentation rates can only be obtained between individual dating points. Recent studies with several cores from the NE TP show large variations of sedimentation within a lake, questioning the reliability of a single sediment core for palaeoclimate reconstructions (Wischnewski et al., 2011; Yan and Wünnemann, 2014). Additionally, the influence of local catchment related processes on lake sediments was highlighted (Wünnemann et al., 2012; Wang et al., 2015). Furthermore, dating of lake sediments are mainly based on radiocarbon ages. However, radiocarbon ages from lake sediments on the TP are hampered by the hard water effect. This effect describes the incorporation of old carbon into organic matter in the lake, leading to an overestimation of the true age. The actual hard-water effect can be determined by measuring the radiocarbon age of a living species. However, variations in the hard- water effect through time can be expected (Long et al., 2011; Hou et al., 2012; Mischke et al., 2013; Lockot et al., 2016). Using magnetic variations is another method to get an independent age control and is supposed to improve the chronology of lake sediments (Haberzettl et al., 2015; Henkel et al., 2016), but has not been widely used on the TP. Due to these limitations other archives have to be utilized. As peat deposits (Hong et al., 2003; H. Wang et al., 2010; B. Liu et al., 2014) and tree ring chronologies (Zhang et al., 2003; Shao et al., 2010; B. Yang et al., 2014) are rare in the semi-arid environment, terrestrial geomorphological archives are the only further indicator of local environmental changes. Aeolian sediments are especially useful since they can be directly dated. During the last years the number of OSL ages from the Tibetan Plateau has been increasing fast. Since the year 2000 roughly 500 OSL ages of aeolian sediments were published (Fig. 1.2). However, large spatial gaps exist. Additionally, individual studies often resulted in divergent palaeoclimate interpretations. This might be caused by insufficient data interpretation, e.g. an exact definition which process is dated by OSL, different local influences or anthropogenic effects.

Fig. 1.2: Published OSL ages from aeolian sediments from the Tibetan Plateau (red: number of OSL ages; blue: number of publications).

The aim of this habilitation thesis is:

a) To fill local data gaps regarding aeolian sediments on the northern and north-eastern Tibetan Plateau. b) Identify the local transportation processes and their significance for palaeoclimate reconstructions. c) To reconstruct the palaeoclimate evolution based on aeolian sediments of the TP with a specially focus on the north-eastern part.

To fill local data gaps on the northern TP, three catchment areas were intensively studied during several expeditions in the years 2006 to 2013. The three study areas, the Gonghe Basin, the Donggi Cona catchment and the Heihai catchment (Fig 1.3), are forming an east-west orientated transect. This transect reflects the reduction of the influence in the ASM influence from the east to the west. The Gonghe basin in the east is the lowest basin at an elevation of roughly 3000 m which receives the highest precipitation values. The Donggi Cona basin is located further to the south-west at 4000 m. The Heihai basin is situated at the western end of the transect at an elevation of 4500 m and exhibits the driest climate conditions. Each catchment area has a large variety of aeolian landforms ranging from fully developed dune system to local loess deposits. The configurations of the catchments enable a detailed reconstruction of the sediment transportation pathways and the analysis of the different variables influencing aeolian deposition. After identifying these variables, the palaeoclimatic significance of the deposits can be evaluated. The study is further extended to other catchment areas on the northern TP like the Qinghai Basin and the Qaidam Basin where an extensive set of OSL ages was previously published by other groups. This widens the base for an overall analysis of the palaeoclimatic drivers of aeolian sedimentation in the area.

Fig. 1.3: The study area on the northern TP; A: Gonghe Basin, B: Donggi Cona, C: Heihai.

For this thesis, six articles were combined. At first, results from three local studies in the Gonghe Basin (Stauch et al., 2017), the Donggi Cona (Stauch et al., 2012, 2014) and the Heihai catchment (Stauch et al., 2017) are presented in chapter 4. Chapter 5 contains two meta-studies based published OSL ages from aeolian sediments. The first one of these is an evaluation of all available OSL ages (2003-2015) since the late glacial for the whole TP (Stauch, 2015), while the second one is dealing with analysis of multi-decadal changes during the late Holocene (Stauch, 2016). A short summary and the most important results of each article are presented in chapter 3. Additionally, 14 co-authored publications which also contributed to the topic of this thesis are listed in this section.

2 Current state of research

2.1 The Asian monsoon system

The Asian monsoon system is one of the most important atmospheric systems on earth (Clift and Plumb, 2008; Cheng et al., 2012). It is characterized by seasonal reversals in the atmospheric circulation and the associated precipitation, resulting in wet summers and dry winters, the ASM and the AWM, respectively (Trenberth et al., 2000; Wang and Ding, 2008; P. X. Wang et al., 2014) (Fig. 1.1). The intensity, timing and extent of the Asian monsoon system is critical for billions of people in Asia (Webster et al., 1998). Variations in the monsoon intensity can result in large scale droughts or flooding (An et al., 2015). The ASM consist of basically two subsystems, the Indian summer monsoon (ISM) system and the East Asian summer monsoon (EASM) (Ding and Chan, 2005; P. Wang et al., 2005). While the ISM, sometimes also called South Asian monsoon, is a true tropical monsoon the EASM is the only subtropical monsoon in the global monsoon system (P. X. Wang et al., 2014; Yim et al., 2014). Beside these two subsystems, which affect large parts of south-east Asia, a third component in the ASM is often recognized, the Western North Pacific monsoon (WNP). However, this tropical monsoon is largely oceanic (Ding and Chan, 2005; P. X. Wang et al., 2014) and not considered further here. The ASM is driven by the land-sea thermal contrast due to seasonal variation of the solar radiation. Besides that, asymmetric diabatic heating and large-scale orography lead to substantial modifications (Wu et al., 2012, 2009). Especially the influence of the different mountain systems like the Himalayan on the southern side (Boos and Kuang, 2010; Molnar et al., 2010) and the Kunlun Shan on the northern side of the Tibetan Plateau has been discussed (Ramisch et al., 2016). Furthermore, the uplift of the Tibetan Plateau has a large influence on the intensity of the ASM, as it increases the thermal contrast (Flohn, 1957; Hahn and Manabe, 1975; Prell and Kutzbach, 1992). This results in an intensification of the ASM and AWM (Ruddiman and Kutzbach, 1989). Further to the north, subsistence results in the aridity of northern China (Manabe and Broccoli, 1990). Despite the immense importance of the ASM for large parts of Asia, considerable uncertainties exist regarding the future of the ASM. Since the late 1970s a weakening of the EASM in northern China has been observed (Yu et al., 2004; M. Xu et al., 2006; Liu et al., 2015). In eastern China, the Yangtze River Valley experienced higher rainfall while the Yellow River Valley became drier (Li et al., 2010). This pattern has been termed “southern China flood and northern China drought” (Ding et al., 2008; Zhou et al., 2009).

The ASM has developed at least since the last 8-9 Ma years (An et al., 2001) and probably even in the Miocene (Guo et al., 2002, 2008). Several long records covering orbital (105-104 yr) timescales are available which are assumed to reflect changes in the ASM. The longest records are from the Chinese Loess Plateau (CLP). Aeolian accumulation mainly occurs under a strong AWM while soil formation is used as an indicator of an enhanced ASM (An et al., 1990; Rutter and Ding, 1993; Ding et al., 1994; Sun et al., 1998; Wyrwoll et al., 2016). Additional records, spanning several hundreds of thousands of years, have been retrieved from the ocean floor and cave deposits. The most widely used marine records are from the Arabian Sea and are used as an indicator for the ISM (Clemens et al., 1991; Clemens and Prell, 2003; Caley et al., 2011). Speleothems were sampled on the Arabian Peninsula (Fleitmann et al., 2003, 2011) and northern India (Sinha et al., 2005; Kathayat et al., 2016) as indicator of the ISM and in central and south-eastern China (Wang et al., 2001, 2008; Dykoski et al.,

2005; Cheng et al., 2006; Zhang et al., 2008) related to the EASM. Variations in the δ18O content of the speleothems are used as an indicator of the strength of the ISM and EASM, respectively.

Since the beginning of the 1980s enormous progress was made in the understanding of the driving forces of the ASM (Clemens et al., 1991; P. X. Wang et al., 2014). In general, the ASM is driven by the summer insolation on the northern hemisphere. Higher insolation values resulted in a larger land-sea temperature contrast and therefore in a more intensive monsoon (Kutzbach, 1981; Joussaume et al., 1999; Braconnot et al., 2007). However, there are different opinions regarding the exact timing and drivers of the ASM (An et al., 2011; Sun et al., 2015). For the EASM, loess proxies indicate a strong coupling to ice volume changes, resulting in a pronounced 100 ka cycle (Ding et al., 1995; An et al., 2015) while oxygen isotope records from Chinese caves show a pronounced signal in the precession band. A direct, in phase link between precession-dominated northern hemisphere summer insolation (NHSI) and a strong EASM, reflected by low δ18O values in the caves, was assumed (Wang et al., 2001, 2008; Sun et al., 2015). However, this interpretation was questioned recently. Clemens et al. (2010) identified a phase lag of 2.9 ± 0.3 ka between NHSI and the δ18O record from Chinese caves and a 8 ± 1 ka lag between the marine ISM record and NHSI, indicating additional forcing mechanisms like winter temperature (Caley et al., 2011). Additionally, the interpretation of the δ18O as a pure climatic signal, especially in the EASM domain, is still debated (Maher, 2008; Clemens et al., 2010; Caley et al., 2011, 2014; Cheng et al., 2012; Maher and Thompson, 2012; Liu et al., 2015). In contrast, analysis of different speleothem records from the ISM and EASM domain revealed a high synchrony between the two regions, indicating a similar response to insolation forcing (Cai et al., 2010; Cheng et al., 2012; Kathayat et al., 2016; Rao et al., 2016a).

Beside orbital variations in the ASM suborbital to millennial scale climate variability has been frequently described. Weak monsoonal phases are linked with cold events on the northern hemisphere like Heinrich events and the related Dansgard-Oeschger cycles (Porter and An, 1995; Wang et al., 2001; Fleitmann et al., 2003; Cheng et al., 2009). For the Holocene, changes in the ASM were associated with changes in the North Atlantic region (Hong et al., 2003; Liu et al., 2009). A frequently proposed correlation was established between episodes of enhanced discharge of icebergs in the North Atlantic (Bond et al., 1997, 2001) and phases of a weak ASM (Fleitmann et al., 2003; Dykoski et al., 2005; Y. Wang et al., 2005; Wang et al., 2016). Additionally, several high resolution speleothem records show coherence with variations in solar variability (Y. Wang et al., 2005; Zhang et al., 2008; Berkelhammer et al., 2010).

However, these large scale changes have a profoundly different local impact. Early work by An (2000) proposed a south-east retreat of the maximum ASM precipitation following the decline in NHSI during the Holocene. Several more recent studies observed a divergent moisture evolution between southern and northern China since the late glacial. While the rainfall in northern China intensified during the early Holocene until a mid-Holocene maximum (Herzschuh, 2006; Y. Wang et al., 2010; Ran and Feng, 2013) and reduced afterwards, for southern China a gradual increase from the late glacial until the Holocene was observed (Liu et al., 2015). These results are incompatible with the suggestion that the well-dated speleothem record is a suitable indicator for monsoonal variations at the northern border of the ASM (Liu et al., 2015). Furthermore, this regional pattern, which is largely consistent with the above mentioned observation of “southern China flood and northern China drought”, was also observed in modelling studies (Z. Liu et al., 2014). Records from northern China indicate a strong EASM associated with higher precipitation during the middle Holocene. Holocene palaeosol formation on the Chinese Loess Plateau was especially intensive from 8.8 to 3.4 ka (H. 6

Wang et al., 2014). This is in accordance with proxies for precipitation from the CLP (Lu et al., 2013). Similar results (7.8 to 5.3 ka) were obtained from pollen records in northern China (F. Chen et al., 2015b) and reduced aeolian activity in the northern drylands (Mason et al., 2009; Yang et al., 2011; Li and Yang, 2016). A strong mid-Holocene EASM also occurred in recent models from the Palaeoclimate Modeling Intercomparison Project (PMIP) (D. Jiang et al., 2013). In contrast, regions influenced by the ISM experienced a Holocene humid period in the early Holocene (Fleitmann et al., 2003; Günther et al., 2015; Zhu et al., 2015), resulting in an asynchronous evolution of the two subsystems of the ASM (Y. Wang et al., 2010; Rao et al., 2016a). An asynchrony in the Holocene climate development was also observed between the regions influenced by the ASM and the regions influenced by the mid-latitude westerlies (Herzschuh, 2006; Chen et al., 2008).

The north-eastern Tibetan Plateau has been identified as an especially sensitive region to climate changes as it is located at the present monsoonal boundary. The area is influenced by three main atmospheric systems, the ASM, the mid-latitude westerlies and the AWM (Böhner, 2006; Morrill et al., 2006; An et al., 2012). Out of these the ASM can be regarded as the most important component as it is the main source of precipitation in the area (Henderson et al., 2010; An et al., 2012). However, the influence of the EASM and ISM in the area might have been variable thorough time (Y. Wang et al., 2010, 2014; Hudson and Quade, 2013). The area has been intensively studied during the last decades. A large number of records is available from (A) lacustrine deposits: Qinghai Lake (Lister et al., 1991; Yu and Kelts, 2002; Shen et al., 2005; Colman et al., 2007; Henderson et al., 2010; An et al., 2012; Jin et al., 2015; Liu et al., 2016), Lake Luanhaizi (Herzschuh et al., 2005, 2006, 2010a), Hala Lake (Wünnemann et al., 2012; Yan and Wünnemann, 2014; Wang et al., 2015), Donggi Cona (Aichner et al., 2012; Dietze et al., 2010, 2012, 2013, Mischke et al., 2010a, 2010b, Opitz et al., 2012, 2016; Y. Wang et al., 2014; Weynell et al., 2016), Chaka Salt Lake (X. Liu et al., 2008), Genggahai Lake (Song et al., 2012; Qiang et al., 2013b, 2014; Rao et al., 2016b), Kuhai (Wischnewski et al., 2011), Lake Koucha (Mischke et al., 2008; Herzschuh et al., 2009; Aichner et al., 2010a, 2010b; Wischnewski et al., 2011), Hurleg Lake (Zhao et al., 2007; C. Zhao et al., 2009; Zhao et al., 2013), Kusai Lake (Liu et al., 2009; Hou et al., 2014; X. Liu et al., 2014), (B) peat deposits: Gonghe Basin (Liu et al., 2013a; B. Liu et al., 2014), Zoige Basin (Hong et al., 2003; H. Xu et al., 2006; H. Wang et al., 2010; Zhao et al., 2011) and (C) aeolian sediments: Qinghai Lake (Lu et al., 2011b; Liu et al., 2011, 2012; Lu et al., 2015), Qaidam Basin (Yu and Lai, 2012, 2014; Zhou et al., 2012; Yu et al., 2013, 2015), Gonghe Basin (Liu et al., 2013b; Qiang et al., 2013a, 2016) (Fig. 2.1).

Last glacial maximum (30-15ka)

During the global last glacial maximum (gLGM) all records indicate arid and cold conditions in the area (Shen et al., 2005; Herzschuh, 2006; Herzschuh et al., 2006; Y. Wang et al., 2010, 2014; Jin et al., 2015; Thomas et al., 2016). Reconstructed precipitation values from Lake Luanhaizi (Herzschuh et al., 2010a), and Donggi Cona (Y. Wang et al., 2014) show arid climate conditions with precipitation values between 100 and 250 mm. The mid-latitude westerlies presumably had a strong influence on the north-eastern TP (Kapp et al., 2011; An et al., 2012; Thomas et al., 2016). The lakes on the north- eastern Tibetan Plateau had either low lake levels (e.g. Hala Lake (Wünnemann et al., 2012; Yan and Wünnemann, 2014), Qinghai (Lister et al., 1991; Jin et al., 2015), Lake Kuhai (Mischke et al., 2010c), Donggi Cona (Dietze et al., 2010)) or disappeared (Koucha (Mischke et al., 2008), Genggahai Lake (Song et al., 2012; Qiang et al., 2013b). 7

Fig. 2.1: Selected archives on the northern TP (blue: lakes; green peat deposits; yellow: aeolian sediments).

Late glacial (15-11.7ka)

During the late glacial many records indicate increasing moisture conditions, but with considerable differences in the timing of the onset. While pollen records from Lake Qinghai indicate a gradual warming since 17 to 14 ka (Shen et al., 2005) other reconstructions point to a rapid increase in precipitation at the beginning of the Holocene (An et al., 2012) and fluctuating lake levels during the late glacial (Jin et al., 2015). A leaf wax hydrogen isotope record from the Qinghai Lake shows extremely dry climate conditions from 15 to 14 ka (Thomas et al., 2016). At the nearby Lake Luanhaizi precipitation already increased from 15 ka onwards (Herzschuh et al., 2010a). A similar onset was recorded in Lake Kuhai (Wischnewski et al., 2011), the Donggi Cona (Y. Wang et al., 2014) and the Genggahai Lake (Song et al., 2012; Qiang et al., 2013b). The infilling of Hala Lake in the north of the TP at 14 ka was attributed to warmer climate and related glacier melt (Yan and Wünnemann, 2014). However, a cold and arid reversal during the Younger Dryas is also preserved in many proxy records at different sites (Yu and Kelts, 2002; Hong et al., 2003; Shen et al., 2005; X. Liu et al., 2008; Y. Wang et al., 2010; Yan and Wünnemann, 2014; Wang et al., 2015; Thomas et al., 2016).

Early Holocene (11.7 – 8 ka)

The early Holocene is characterized by a further increase in precipitation related to the strengthening of the ASM. At the Qinghai lake, highest moisture conditions were reconstructed from 11.5 to 8 ka (Jin et al., 2015) and from 11.2 until 7 ka based on leaf wax hydrogen isotopes (Thomas et al., 2016).

Colman et al. (2007) reported highest moisture values at 10 ka and steady decline until 4 ka. However, several dry periods during the early Holocene were identified in dune sediments at the Qinghai Lake (Lu et al., 2011b; Liu et al., 2012). While at Lake Luanhaizi the Holocene optimum occurred at the end of the early Holocene from around 9 to 7ka (Herzschuh et al., 2006). Similar wet climate conditions in the early Holocene were identified at the Hala Lake (Yan and Wünnemann, 2014) and the Donggi Cona (Mischke et al., 2010a; Opitz et al., 2012; Dietze et al., 2013; Y. Wang et al., 2014). In contrast, further to the south, Lake Koucha in the Bayan Har Shan area recorded arid conditions in ostracod data throughout the early Holocene (Mischke et al., 2008), while the pollen record indicated higher precipitation values from 14.6 to 6.6 ka, including a drier phases from 9.7 to 7.5 ka (Herzschuh et al., 2009; Aichner et al., 2010a). An early Holocene cold phase (11 – 9 ka) was recorded further to the north in the Hala Lake (Wang et al., 2015). However, this lake might be influenced by the westerlies and not the ASM (Wünnemann et al., 2012). From the Hongyuan peat section, high moisture values were reconstructed from 10.8 to 5.5 ka (Hong et al., 2003; H. Wang et al., 2010). At the end of the early Holocene several records indicate a short climate deterioration which was correlated with the 8.2 ka event (Shen et al., 2005; Mischke et al., 2008; Mischke and Zhang, 2010; Schlütz and Lehmkuhl, 2009; Yan and Wünnemann, 2014).

Mid-Holocene (8 – 4 ka)

For the Qinghai Lake different reconstructions for the mid-Holocene exist. While Shen et al. (2005) assumed a mid-Holocene maximum at around 6.5 ka, a stepwise decline in moisture since 7 ka was recorded by Thomas et al. (2016). In contrast, at Lake Luanhaizi present day climate conditions were probably already reached at 7 ka (Herzschuh et al., 2006). A peat section in the Gonghe Basin indicated optimal moisture conditions from 7.6 ka to 3.8 ka (Liu et al., 2013a; B. Liu et al., 2014). High lake levels in the Genggahai Lake in the Gonghe Basin were recorded from 7.4 to 5.5 ka (Qiang et al., 2013b). In contrast, Liu et al. (2013b) assumed dry climate conditions from 8.7 to 4.7 ka based on aeolian deposits. Qiang et al. (2013a, 2016) assumed that higher precipitation values during this time were offset by higher evaporation values. In the eastern Qaidam Basin aeolian deposits indicate highest moisture conditions from around 8.3 to 3.5 ka (Yu and Lai, 2012, 2014). At Lake Koucha, based on ostracod data, wetter climate conditions started at around 7.3 ka and persisted at least until 4.3 ka (Mischke et al., 2008), while for Lake Kuhai wetter but colder climate conditions from 7 to 6.3 ka were reconstructed from pollen data (Wischnewski et al., 2011). Significant cooler and drier climate existed at Lake Kuhai until 2.2 ka (Wischnewski et al., 2011). Interestingly at the Donggi Cona, after a decline in precipitation from 9.5 to 7.3 ka a second wet phase occurred in the mid-Holocene until 4.3 ka (Y. Wang et al., 2014). At the Hongyuan peat section drier climate conditions started already at 5.5 ka (Hong et al., 2003). Colluvial layers and palaeosols from the north-eastern TP indicate wetter climate conditions from 9 to 3.5 ka (Kaiser et al., 2007). In contrast, in the Qaidam Basin north of the TP a dry climate was recorded in the Hurleg Lake from 9.5 to 5.5 ka (Zhao et al., 2007).

Late Holocene (4 – 0 ka)

Most of the records indicate considerably drier conditions during the late Holocene (Y. Wang et al., 2010). Nevertheless, the onset of the late Holocene dry period varies considerably between the different studies. The stepwise decline in moisture at the Qinghai lake reached modern semi-arid

9 climate conditions at around 2.6 ka (Thomas et al., 2016) or probably already at ~ 4.5 ka (Ji et al., 2005; Shen et al., 2005) and 4 ka (Colman et al., 2007). The OSL dated shoreline deposits indicate a lake level decline since 3.6 ka (Liu et al., 2011). At the Hala Lake lake level dropped in response to a cooler climate at around 4.5 ka and reached a low stand already at 4.1 ka. However, afterwards a late Holocene lake level rise was recorded until 1.5 ka and a subsequent drop until today (Yan and Wünnemann, 2014). In the Gonghe Basin dry conditions were established from 5.5 ka in the Genggahai Lake (Qiang et al., 2013b), 5.2 ka in the Chaka Salt Lake (X. Liu et al., 2008) and since 3.8 ka in the peat section in the eastern part of the Gonghe Basin (B. Liu et al., 2014). At the Donggi Cona, the late Holocene dry phase started at 4.3 ka (Y. Wang et al., 2014).

Several studies from the northern and north-eastern TP indicate more humid climate conditions during the late Holocene. At the Qinghai Lake a return to wetter conditions was recorded at 2.3 ka (Ji et al., 2005). A similar trend was observed for the last 2.2 ka at Lake Kuhai (Wischnewski et al., 2011) and the last 2 ka at Lake Koucha (Mischke et al., 2008). At the Genggahai Lake in the Gonghe Basin, high lake levels occurred from 2.1 to 1.6 ka (Qiang et al., 2013b).

Beside these multi-millennial trends on the north-eastern TP, several multi-decadal to millennial climate fluctuations were described. Millennial-scale climate variations during the last 2.5 ka were described e.g. for the Qinghai Lake (Shen et al., 2005; Henderson and Holmes, 2009; Henderson et al., 2010) including e.g. the Little Ice Age (LIA) and the Medieval Warm Period (MWP). Additionally, climate fluctuations were linked to Bond events in the North Atlantic, e.g. Hurleg Lake (Zhao et al., 2013), Kusai Lake (X. Liu et al., 2014) and in the peat sections from the Gonghe (B. Liu et al., 2014) and Zoige Basin (Hong et al., 2003; H. Xu et al., 2006; H. Wang et al., 2010).

Moreover, the human influence on the landscape system was frequently discussed (e.g. Schlütz and Lehmkuhl, 2009; Yu et al., 2015).The human influence on the north-eastern TP might already have started at 7.2 ka with the first indicators of a human presence in the area (Schlütz and Lehmkuhl, 2009). Even older ages were reported from the northern edge of the TP, ranging back until the late glacial (Brantingham et al., 2003; Aldenderfer and Zhang, 2004; Madsen et al., 2006). An increasing human impact was assumed since the end of the mid-Holocene (Kaiser et al., 2009b; Miehe et al., 2009, 2014). Since 3.6 ka permanent settlements have been established on the north-eastern TP (F. Chen et al., 2015a). At present there is a strong human impact on the TP resulting in different stages of degradation and desertification of the landscape (Lehmkuhl et al., 1999; Feng et al., 2005; Cui and Graf, 2009; Yan et al., 2009; Hu et al., 2015).

2.2 Aeolian sediments and palaeoclimate reconstructions

In arid and semi-arid environments wind is an important factor in changing the earth surface (Wiggs, 1997). Aeolian sediments cover large parts of the earth (Thomas and Wiggs, 2008). Consequently, aeolian sediments are an important geo-archive for the reconstruction of past environments (Yang et al., 2011; Thomas, 2013; Újvári et al., 2016). A special emphasis was placed on the large mid-latitude loess deposits (Liu, 1988; Muhs, 2013) and aeolian landforms like dunes and sand seas (Yang et al., 2004, 2013; Roskin et al., 2011). The accumulation of aeolian sediments is controlled by a variety of influencing factors. The basic factors are sediment supply, sediment availability and the transport capacity of the wind (Kocurek and Lancaster, 1999). Sediment supply is defined as the sediment which is stored in the landscape and generally available for the aeolian processes, while sediment availability refers to the sediment which can be entrained by the wind under the actual environmental conditions. Changes in the environment can lead to higher sediment availability by utilizing larger amounts of the sediment supply for aeolian processes. The transport capacity describes the amount of material which can be transported by the wind (Kocurek and Lancaster, 1999). Sediment availability depends on numerous factors, especially the protection of the sediment e.g. by vegetation (Lancaster and Baas, 1998) or soil moisture (Fécan et al., 1998; Davidson-Arnott et al., 2008; Scheidt et al., 2010). Accumulation of aeolian sediments takes place when the transport capacity rate is reduced (e.g. Rodríguez-López et al., 2014). This can be caused e.g. by topographic obstacles (Mason et al., 1999; Bullard et al., 2000), enhanced surface roughness due to vegetation (Bagnold, 1941; Bullard, 1997; Lancaster and Baas, 1998) or a climate-induced reduction in wind speed (Roskin et al., 2011). Aeolian sediments accumulate when the sediment influx at a certain location exceeds the downwind sediment outflux, resulting in positive sediment budget at that point (Rodríguez-López et al., 2014). The parameters above regulate the aeolian transport from sediment storage to the accumulation area and the final preservation. Important factors for these are variations in the palaeoclimate, e.g. changes in wind regime and moisture. A large variety of possible sources for aeolian material was identified, ranging from aeolian archives, which are reworked, to fluvial or glacio-fluvial deposits and dry lake beds (Bullard and Livingstone, 2002; Prospero et al., 2002; Muhs et al., 2008; Prins et al., 2009; Bullard, 2013).

In general, aeolian sediment transport occurs under dry climate conditions (Bullard and Livingstone, 2002). However, due to the complex interaction between sediment supply, transport capacity, sediment accumulation and preservation not all dry periods result in the formation of aeolian archives (Bullard and Livingstone, 2002). In recent years, the influence of fluvial transport on the sediment availability was particular highlighted (Bullard and McTainsh, 2003; Visser et al., 2004; Field et al., 2009). Even a small increase in rainfall can significantly increase the fluvial transport to a sediment storage site where sediment is easily available for deflation (Lancaster, 1997; Clarke and Rendell, 1998). For the western Mu Us and the Chinese Loess Plateau a significant sediment transport from the floodplains of the Yellow River was discussed (Stevens et al., 2013; Nie et al., 2015). The spatial distribution of aeolian landforms in the northern foreland of the Qilian Shan is related to a sufficient sediment supply provided by local rivers. A strong coupling between sediment supply and aeolian archives was observed. Areas without a fluvial sediment supply only have a limited amount of aeolian landforms despite similar climate conditions (Nottebaum et al., 2014, 2015a).

A further problem in the interpretation of aeolian sediments is the frequent reworking of older sediments. This problem is obvious for presently active dunes which migrate downwind (Lancaster, 2008). However, reworking of sediments is not that obvious in fossil aeolian sections (Munyikwa, 2005; Telfer et al., 2010). The reworking of older aeolian sediments was also observed in loess deposits (Stevens et al., 2007; Licht et al., 2016). This further implies that a sedimentation gap in a single section cannot be directly interpreted as a phase without aeolian activity (Stone and Thomas, 2008; Thomas, 2013).

During aeolian transport sediment particles tend to be sorted. Depending on the speed of the wind (the wind stress) and the size of the particles several physical regimes are differentiated; I) long-term suspension ( < ~ 20 μm diameter), II) short-term suspension (~ 20 – 70 μm), III) saltation (70 – 500 μm) and IV) reptation and creep (> 500 μm) (Bagnold, 1941; Pye, 1987; Kok et al., 2012). The wind stress first mobilizes sand size particles when a threshold value (threshold friction velocity, Shao, 2008) is crossed (Bagnold, 1941). The value of the threshold depends on the properties of the transporting air as well as on the gravitational forces and the inter-particle cohesion (Kok and Renno, 2009). Additional influencing factors are e.g., soil moisture, mineralogy, soil texture, surface crusts, vegetation cover and roughness elements (Shao and Lu, 2000). Following lifting, saltation occurs. After a ballistic trajectory the particles hit the surface again, resulting in the release of further particles. These impacts are the main cause of the lifting of further particles (Bagnold, 1941; Shao, 2008), increasing the amount of particles subjected to saltation. The number and speed of the ejected particles depends on the size of the impacting particles, the impact speed and impact angle (Rice et al., 1995). The number of saltating particles is finally limited by the retardation of the wind stress by the drag of the saltating particles, which reduces the horizontal momentum of the wind. Therefore, the number of ejected sand particles is approaching a steady state with time until for each impacting particle only one other is ejected (Ungar and Haff, 1987; Werner, 1990). This is leading to a transport limited saltation process (Kok et al., 2012). From a geomorphological point of view, the above described saltation process is the main process for the formation of sand ripples and dunes (Bagnold, 1941). Smaller particles ( < 70 μm) are normally not directly lifted due to the cohesive inter-particle forces, which are relatively large, compared to the dragging forces of the air (Kok et al., 2012). The cohesive forces are overcome by impacts of the saltating particles. After the ejection from the surface, dust particles are susceptible to the turbulent flow and can enter long- and short-term suspension (Kok et al., 2012). However, recent studies from loess areas show, a considerable amount of particles is directly entrained from the surface (Sweeney and Mason, 2013; Újvári et al., 2016). This is in accordance with a significant dust production in temperate grasslands (Shinoda et al., 2011). A further important process for the emission of dust particles from a surface is the saltation of soil aggregates. These aggregates disintegrate during the impact and release the fine particles (Shao, 2008; Kok, 2011). Deposition of dust particles occurs either as wet or dry deposition. Wet deposition occurs when dust particles form cloud nuclei or are incorporated in raindrops. Dry deposition is caused by gravitational settling and absorption at surfaces (Shao, 2008). A recent summary of the dust deposition processes is given by Újvári et al. (2016). Due to their large inertia, particles larger than 500 μm are not able to saltate (Shao, 2008). These particles are either transported in a short hop of less than a centimeter, which is called reptation, or they are rolling across the surfaces (creep). These two processes can account for a substantial amount of the total aeolian sand flux (Bagnold, 1941).

During the above explained processes a sorting process of the sediment occurs. Variations in the grain size of deposits are frequently used for the reconstruction of palaeoclimate environments (Ding et al., 1995; Porter and An, 1995; Nugteren et al., 2004; Vriend et al., 2011; Wang and Lai, 2014). However, samples from aeolian deposits often do not show a simple uniform grain-size distribution (GSD). Instead, many GSDs have a bimodal or even multimodal distribution, due to the incorporation of several aeolian processes in one sample or post-depositional mixing (Dietze et al., 2012). Several mathematical models were proposed for the unmixing of these GSD, like Weibull functions (Sun et al., 2002) or end-member analysis (Weltje, 1997; Weltje and Prins, 2007; Dietze et al., 2012, 2014; IJmker et al., 2012a; Nottebaum et al., 2015b).

Aeolian sediments are widespread on the northern TP and are ranging from a thin loess cover to fully developed dune systems (e.g. Hövermann, 1987; Lehmkuhl, 1997; Lehmkuhl and Haselein, 2000; Sun et al., 2007; Lehmkuhl et al., 2014; Nottebaum et al., 2014). Sediment supply on the northern Tibetan Plateau is related to a variety of sources. Glacial outwash material (e.g. Sun et al., 2007; Smalley et al., 2014), fluvial deposits and exposed lake sediments (Lehmkuhl and Haselein, 2000; Kaiser et al., 2010; IJmker et al., 2012a; Yu et al., 2015; Qiang et al., 2016) are the main sediment sources. Due to the large amounts of sediments stored in the different basins of the TP, it can be assumed that the aeolian system is not supply limited. The boundary conditions are sediment availability and transport capacity. Climate data from the northern TP (Maussion et al., 2013) indicate that high wind speeds occur frequently on the northern TP (see also X. Yang et al., 2012). Therefore, the actual sediment availability can be considered as the most important factor for the initiation of aeolian transport. This was also recognized for the large sand seas in northern China (X. Yang et al., 2012).

Due to the frequent occurrence of aeolian sediments and the possibility to directly date the deposits, aeolian sediments were intensively studied on the northern (e.g. Stokes et al., 2003; Küster et al., 2006; Liu et al., 2012, 2013b, Yu and Lai, 2012, 2014; Yu et al., 2015; Qiang et al., 2013a, 2016; Lu et al., 2015) and southern TP (e.g. Lehmkuhl et al., 2000, 2002; Sun et al., 2007; Lai et al., 2009; Klinge and Lehmkuhl, 2015) during the last decades.

OSL dating is based on the accumulation of electrons in the crystal lattice of quartz or feldspar grains which is caused by ionizing radiation from radioactive isotopes within in the sediment and by cosmic rays (Duller, 2016). Theoretically, the energy is preserved until it is released in form of a luminescence signal, e.g., the stimulation by light. This occurs in nature when the sediment grains are exposed to light during transportation, thus resetting the luminescence signal to zero. The luminescence signal of an artificially stimulated sample can be measured in the laboratory. The time since burial of the sediment grains can be determined by dividing the measured luminescence signal (equivalent dose) by the radiation of the surrounding sediment and the radiation dose from cosmic rays (dose rate) (e.g., Singhvi et al., 2001; Duller, 2004; Roberts, 2008; Wintle, 2008). A recent summary of the different techniques of OSL dating was presented by Duller (2016). During aeolian transport the luminescence signal is reset due to the exposure to light, making aeolian sediments a prime target for OSL dating. In the last 16 years the number of OSL ages from aeolian deposits from the whole TP has exceeded 500.

However, despite this large amount of data, there is still no consensus reached regarding the interpretation of the different types of aeolian deposits, e.g. sand, loess and palaeosols, and the interpretation of the OSL ages from these deposits on the TP. As highlighted before, aeolian transport occurs under dry climate conditions when the sediment availability is high and high wind 13 speed occurs, enhancing the transport capacity of the wind. Therefore, aeolian sand deposits on the northern TP are often related to dry climate conditions (e.g. Lu et al., 2011b; Liu et al., 2012, 2013b; Yu and Lai, 2014). The deposition of loess is associated with intermediate climate conditions (Stokes et al., 2003; Lu et al., 2011b; Yu and Lai, 2012), while the formation of palaeosols is associated with a wet climate (Liu et al., 2012; Yu and Lai, 2014). For the formation of loess deposits the establishment of a sufficient vegetation cover is necessary (Lehmkuhl et al., 2000; Küster et al., 2006; Sun et al., 2007), requiring sufficient moisture. However, studies on the OSL ages from arid environments indicate that preserved sedimentary bodies from aeolian sands are mainly related to the end of a dry period (Chase and Thomas, 2007; Lancaster, 2008; Leighton et al., 2014). During full dry conditions a frequent reworking of aeolian sediments takes place while at the end of a dry period a stabilisation of the deposits occurs (Chase and Thomas, 2007). This was also observed in the sand seas of northern China. Sediments deposited during or shortly before an interglacial are more resistant to deflation (Li and Yang, 2016). The stabilisation can be related to a reduction in wind speed and/or the development of a sufficient vegetation cover, which, in turn, would require a sufficient moisture supply.

The interpretation of OSL ages from palaeosols is controversial (e.g. Mason et al., 2009). In many cases a deposition parallel to the pedogenesis is assumed. In this case the OSL age of the sediment reflects the timing of the palaeosol formation (e.g. Lu et al., 2011b; Liu et al., 2013b; Lu et al., 2015). This assumption was intensively used in the interpretation of the palaeosols in the Chinese Loess Plateau (CLP) (e.g. H. Wang et al., 2014). However, if no or only weak aeolian deposition occurred during palaeosol formation, the OSL age from the sediment in which the palaeosol developed reflects the age of the hosting sediment (Mason et al., 2009). The deposition of the sediment might occur a considerable time before the pedogenesis (Singhvi et al., 2001) and an OSL age of this sediment is not related to the age of the palaeosol. An additional blurring effect can be caused by bioturbation (Bateman et al., 2003a). This was recently reported for palaeosols from the CLP (Stevens et al., 2006). However, under intensive bioturbation it is theoretically possible that the sediment is sufficiently exposed to sunlight and that the luminescence signal is reset (Singhvi et al., 2001; Bateman et al., 2003a). Consequently, the OSL ages from such a layer might be close to the age of soil formation.

A further problem for the palaeoclimatic interpretation of aeolian sediments from the northern TP arises from the fact that the sediment cover is relatively thin and aeolian sections are often not continuous (Qiang et al., 2013a; Yu et al., 2015). This is indicated by very heterogeneous accumulation rates and a varying number of the observed palaeosols (Yu and Lai, 2014; Lu et al., 2015; Qiang et al., 2016). Palaeoclimate reconstructions based on interpolation between the individual dating points should be treated carefully. An alternative method is the interpretation of probability density functions (pdf) from OSL ages (Singhvi et al., 2001; Lang, 2003; Goble et al., 2004; Lai et al., 2009), beside other methods like frequency distributions or binary plots (Chase, 2009; Thomas and Burrough, 2016). A pdf is based on the mean age and the standard deviation of the OSL ages. From the summed up pdf a dimensionless cumulative pdf can be calculated, which can be used to interpret phases of enhanced aeolian accumulation. As the standard deviation is increasing with age, which results in a decrease in the height and an increase in the width of the peaks, the height of the peaks cannot be used as an indicator of the strength of the aeolian activity (Singhvi et al., 2001). However, each peak is an indicator of an accumulation phase. Several problems are related to the use of pdfs. The normally distributed standard deviation of the OSL ages does not reflect the measured variations in the equivalent dose (Galbraith, 2010; Fitzsimmons et al., 2013). A further

14 problem can be caused by unregularly distributed sampling points. A higher density of samples from specific stratigraphic units, e.g. marker horizons like palaeosols results in an overestimation of the aeolian accumulation at that point in time (Lang, 2003). As clear marker horizons are missing and many unconformities exist in the aeolian deposits this point is presumably not relevant on the northern TP. A general problem of all terrestrial sediments and especially of aeolian sediments is the negative correlation between age and preservation of sediments. The frequent reworking of sediments results in a lower number of older sediments (Telfer et al., 2010; Bailey and Thomas, 2014).

3 Short summaries of the articles

This chapter gives a short summary of each of the six manuscripts included in this thesis. Additionally, for each article the most important new results in the related field are highlighted. A list of the 14 articles which have been co-authored and which also contribute to the topic is located at the end of this chapter. The six full manuscripts are presented in chapter 4 and 5. In chapter 4, the local case studies from the northern TP are presented. Chapter 5 includes the two meta-studies summarizing the evaluation of all published OSL ages from the Tibetan Plateau.

Environmental changes during the late Pleistocene and the Holocene in the Gonghe Basin, north-eastern Tibetan Plateau

Georg Staucha, Zhongping Laib, Frank Lehmkuhla, Philipp Schultea

a Department of Geography, RWTH Aachen University, Templergraben 55, 52056 Aachen, Germany b School of Earth Sciences, China University of Geosciences, 388 Lumo Rd., Wuhan 430074, China

Palaeogeography, Palaeoclimatology, Palaeoecology (in press)

Contribution: >75%

Summary

In the Gonghe basin, a clear zonal differentiation between different types of aeolian sediments was observed. Aeolian sands and dunes are located in the central part of the basin, while fine sands and intercalated palaeosols are mainly preserved at the eastern side of the basin. In contrast, loess deposits are located on the mountain flanks in the north and the south. This zonal differentiation is caused by the topographic configuration of the basin and its orientation, which follows the main wind direction. However, loess sediments were also observed in the center of the basin below the active sand dunes. These sediments indicate a formerly larger extent of the loess cover, probably even covering the entire Gonghe Basin. Previous studies on the timing of the aeolian sediment showed a clear temporal separation, with the accumulation of aeolian sands during the early Holocene and loess deposition in the late Holocene. Based on 43 new OSL ages covering most of the regions in the basin, an early onset of the loess accumulation was detected. Loess accumulated started already at 14 ka, indicating the onset of the ASM in the area. Sand accumulation was especially strong during the YD period. During the mid-Holocene aeolian activity was very low due to maximum moisture levels. During the last 2.5 ka, aeolian activity resumed due to the weakening of the ASM. A comparison of the aeolian sediments in the Gonghe Basin with aeolian sediments in the neighbouring basins indicated a time-transgressive onset of the ASM from the east to the west from 17 ka to 10 ka.

New results

A new chronology for the aeolian sediments in the Gonghe Basin based on 43 new OSL ages was developed. This significantly improved the previous chronologies which were based on 39 OSL ages. The extent of the loess deposits in the basin was mapped and the sediments were dated. The remnants of loess in the central part of the basin were previously unknown. A new paleoclimate reconstruction based on aeolian sediments was developed. The shift in the onset of the aeolian accumulation in the different basin was observed for the first time.

Aeolian sediments on the north-eastern Tibetan Plateau

Georg Staucha, Janneke IJmkera, Steffen Pötschb, Hui Zhaoc, Alexandra Hilgersb, Bernhard Diekmannd, Elisabeth Dietzee, Kai Hartmanne, Stephan Opitzd, Bernd Wünnemanne,f, Frank Lehmkuhla

a Department of Geography, RWTH Aachen University, Templergraben 55, 52056 Aachen, Germany b Institute of Geography, University of Cologne, Albertus-Magnus-Platz, 50923 Cologne, Germany c Cold and Arid Regions Environmental and Engineering Research Institute, Chinese Academy of Science, Lanzhou 730000, PR China d Alfred Wegener Institute for Polar and Marine Research, Research Unit Potsdam, Telegrafenberg A43, 14473 Potsdam, Germany e Institute of Geographical Sciences, Interdisciplinary Center of Ecosystem Dynamics of Central Asia (EDCA), Freie Universität Berlin, Malteserstr. 74-100, 12249 Berlin, Germany f School of Geography and Oceanography, Nanjing University, 22 Hankou Road, Nanjing 210093, PR China

Quaternary Science Reviews 57, 71-84 (2012)

Contribution: >75%

Summary

Based on 51 new OSL ages a paleoclimate model for the accumulation of aeolian sediments was developed for the Donggi Cona catchment. Especially the synchronous ages of loess and sand in the early Holocene indicate that the OSL ages do not capture the full period of aeolian activity but only the final phase when the climatic conditions already changed to higher precipitation values. Based on this assumption a local palaeoclimate model was developed and compared to other proxy records. Aeolian accumulation started in the early Holocene indicating the onset of the ASM. During the mid- Holocene aeolian processes ceased and fluvial processes removed parts of the early Holocene deposits. During the late Holocene a reduced ASM with lower precipitation values resulted in again higher aeolian activity. Additionally, the spatial distribution of loess and sand in the basin can be explained by different components of long and short distance transport.

New results

A detailed palaeoclimate reconstruction was established based on different types of terrestrial sediments and 51 new OSL ages. A new concept was developed for the explanation of the distribution of OSL ages from aeolian sediments on the north-eastern TP.

Interaction of geomorphological processes on the north-Eastern Tibetan Plateau during the Holocene, an example from a sub-catchment of Lake Donggi Cona

Georg Staucha, Steffen Pötschb, Hui Zhaoc, Frank Lehmkuhla

a Department of Geography, RWTH Aachen University, Templergraben 55, 52056 Aachen, Germany b Institute of Geography, University of Cologne, Albertus-Magnus-Platz, 50923 Cologne, Germany; now: Institute of Geography and Geology, University of Greifswald, Friedr.-Ludwig-Jahn-Str. 16, 17487 Greifswald, Germany c Cold and Arid Regions Environmental and Engineering Research Institute, Chinese Academy of Science, Lanzhou 730000, PR China

Geomorphology 210, 23-35 (2014)

Contribution: >80%

Summary

The article describes in detail the sediment cascade on an alluvial fan east of the Donggi Cona. The different components of the sedimentary system ranging from glacial, fluvial to lacustrine and aeolian components were evaluated. Based on these results different transportation pathways through space and time were identified. Especially the influence of non-climatic factors on the sedimentary system was highlighted in the article. Changes in the lake level of the Donggi Cona had a profound influence on the sediment availability. The connection between low lake levels and high accumulation rates was described. Additionally, the seasonality of the transportation processes was highlighted for recent times. Fluvial activity dominates the area during summer time, leading to the erosion of aeolian archives and the relocation of aeolian sands, while during winter times the sediments are again transported up on the slopes of the mountains.

New results

This research is the first detailed, high resolution reconstruction of a sediment cascade on the whole TP. The influence of non-climatic factors on the sedimentary system was stressed. Different sedimentary storages released the material for the aeolian landforms through time. The reconstruction was used to validate previous lake level reconstructions for the Donggi Cona. Furthermore, the problem of seasonality in the geomorphic processes was first described for the aeolian sediments on the TP.

Landscape and climate during the late Quaternary on the northern Tibetan Plateau

G. Staucha, P. Schultea, A. Ramischb, K. Hartmannc, D. Hüllee, G. Lockotc, B. Diekmannd, V. Nottebauma, C. Müllerf, B. Wünnemannc,g, D. Yanc, F. Lehmkuhla

a Department of Geography, RWTH Aachen University, Templergraben 55, 52056 Aachen, Germany b GFZ German Research Centre for Geoscience, Telegrafenberg, 14473 Potsdam, Germany c Institute of Geographical Sciences, Interdisciplinary Center of Ecosystem Dynamics of Central Asia (EDCA), Freie Universität Berlin, Malteserstr. 74-100, 12249 Berlin, Germany d Alfred Wegener Institute for Polar and Marine Research, Research Unit Potsdam, Telegrafenberg A43, 14473 Potsdam, Germany e Institute of Geography, University of Cologne, Albertus-Magnus-Platz, 50923 Cologne, Germany f Institute of Geological Sciences, Freie Universität Berlin, Malteserstr. 74-100, 12249 Berlin, Germany g School of Geography and Oceanography, Nanjing University, 163 Xianlin Ave, Qixia District, 210023 Nanjing, China

Geomorphology 286, 78-92 (2017)

Contribution: >70%

Summary

At Lake Heihai the influence of alluvial fans on lake level variations and the aeolian system were studied in detail. 50 new OSL ages were obtained for the area. And end-member analysis was used to classify the aeolian sediments, resulting in four different classes. Additionally, geochemical properties of the sediments were examined using a factor analysis. Both classifications revealed large spatial differences between the different aeolian sediment types. Coarse-grained sediments accumulated mainly on the northern side of the lake, while loess deposits are exclusively found south and east of the lake. The factor analysis revealed three different source areas for the aeolian sediments. The loess sediments south of the lake originated from the limestone bearing Kunlun Mountains. Loess sediments east of the lake have an additional sediment source in exposed lake sediments around the lake. The sands on the northern side of the lake originate from nearby conglomerates. This spatial separation in the aeolian sediments results in different activation phases through time. However, sediment preservation is an important parameter in the high-altitude environment which is presently outside of the monsoon influenced area. Late glacial to early Holocene aeolian sands were only preserved when they were protected, either due to topographic effects or subsequent burial e.g. by fluvial sediments. It can be assumed that other sediments from this time period were deflated due to the high geomorphological activity in the area. Loess sediments were deposited in the mid-Holocene, indicating a phase of highest moisture conditions during the Holocene. During the late Holocene a high aeolian activity is emanated from the sands on the northern side. This observation is also

20 reinforced by high movement rates of barchans on the northern side of the lake. This later point was detected due to the evaluation of remote sensing images covering the last 50 years.

New results

The research in the Heihai Basin was the first detailed study on aeolian sediments on the northern Tibetan Plateau. It could be demonstrated that site site-specific processes have a profound influence on the distribution of different types of aeolian sediments. These processes are either the sediment sources but also the preservation of the sediments. It could be highlighted that the preservation of aeolian sediments in a protected position might lead to a different palaeoclimatic interpretation than aeolian sediments preserved in a more open environment. Furthermore, the studied showed that the ASM probably reached the area during the mid-Holocene.

Geomorphological and palaeoclimate dynamics recorded by the formation of aeolian archives on the Tibetan Plateau

Georg Stauch

Department of Geography, RWTH Aachen University, Templergraben 55, 52056 Aachen, Germany

Earth Science Reviews 150, 393-408 (2015)

Summary

This study reviewed all published OSL ages from aeolian sediments on the whole TP. It is the first summary of this kind. Large spatial gaps were identified in the western and central TP. Five different regions on the TP were analysed by the use of density probability functions. Furthermore, the OSL ages were classified according to the sediment type, e.g., loess, sandy loess, sand, and palaeosols. By doing so, environmental conditions for the accumulation and preservation of aeolian sediments were identified and resulted in a new model of the interpretation of OSL ages from aeolian sediments on the north-eastern TP. Aeolian sediments in the late glacial and the early Holocene are indicative of increasing precipitation after a period of extremely dry climate conditions with a constant reworking of aeolian sediments. In contrast, aeolian sediments during the late Holocene are caused by a return to a drier climate. This highlights the importance of the previous state of the environmental system for the interpretation of OSL ages from aeolian sediments. Besides that, the palaeoclimate reconstruction based on the dataset shows several phases: A dry glacial, an increasing monsoon in the late glacial and the early Holocene, a wet mid-Holocene and a dry late glacial. While the onset of the sediment accumulation is variable on north-eastern TP, the end of the first accumulation period at the beginning of the mid-Holocene is similar in all regions. In contrast, no consistent signal was observed in the data from the southern TP.

New results

The study is the first summary of all published OSL ages from the whole TP. A new conceptual model for the interpretation of OSL ages from aeolian sediments on the TP was developed. A consistent palaeoclimate reconstruction for the north-eastern TP based on aeolian sediments was established.

Multi-decadal periods of enhanced aeolian activity on the north-eastern Tibet Plateau during the last 2ka

Georg Stauch

Quaternary Science Reviews 149, 91-101 (2016)

Summary

In the study multi-decadal periods on the north-eastern TP were identified in a set of OSL ages from aeolian sediments for the last 2000 years. In a first step the use of density probability functions for the last 5ka were evaluated. However, only the signals for the last 2 ka were statistical robust. Six phase of enhanced aeolian accumulation occurred on then north-eastern TP during the last 2 ka. The first three phases happened during the dark ages cooling, while the last three phases occurred during the Little Ice Age. The induvial phases could also be linked to phases of low northern hemisphere temperature and phases of reduced solar insolation. During the last 800 years each grand solar minimum is synchronous with a phase of enhanced accumulation. Additionally, there is a strong correlation with precipitation records from the Qilian Shan and proxies of monsoon intensity from the Wanxiang Cave on the eastern TP. No explicit relationship was identified for the influence of the westerlies. However, this might be a result of inconsistent westerlies records.

New results

The research identified multi-decadal phases of climate variations from a proxy record without any data interpolation between individual dating points. Aeolian sediments are a suitable proxy for the reconstruction of multi-decadal climate changes, provided that a sufficient number of OSL ages are available, even from discontinues archives. Six multi-decadal phases were recognized. They correlate within errors with northern hemisphere temperature and the strength of the ASM.

Beside the above-mentioned six published articles, the following 14 articles, which also contribute to the topic of the thesis, were coauthored. For the clarity of the structure these articles are not included in the thesis as individual chapters.

Dietze, E., Wünnemann, B., Diekmann, B., Aichner, B., Hartmann, K., Herzschuh, U., IJmker, J., Jin, H., Kopsch, C., Lehmkuhl, F., Li, S., Mischke, S., Niessen, F., Opitz, S., Stauch, G., Yang, S., 2010. Basin morphology and seismic stratigraphy of Lake Donggi Cona, north-eastern Tibetan Plateau, China. Quaternary International 218, 131–142.

Dietze, E., Hartmann, K., Diekmann, B., IJmker, J., Lehmkuhl, F., Opitz, S., Stauch, G., Wünnemann, B., Borchers, A., 2012. An end-member algorithm for deciphering modern detrital processes from lake sediments of Lake Donggi Cona, NE Tibetan Plateau, China. Sedimentary Geology 243–244, 169–180.

Dietze, E., Wünnemann, B., Hartmann, K., Diekmann, B., Jin, H., Stauch, G., Yang, S., Lehmkuhl, F., 2013. Early to mid-Holocene lake high-stand sediments at Lake Donggi Cona, northeastern Tibetan Plateau, China. Quaternary Research 79, 325–336.

IJmker, J., Stauch, G., Dietze, E., Hartmann, K., Diekmann, B., Lockot, G., Opitz, S., Wünnemann, B., Lehmkuhl, F., 2012a. Characterisation of transport processes and sedimentary deposits by statistical end-member mixing analysis of terrestrial sediments in the Donggi Cona lake catchment, NE Tibetan Plateau. Sedimentary Geology 281, 166–179.

IJmker, J., Stauch, G., Hartmann, K., Diekmann, B., Dietze, E., Opitz, S., Wünnemann, B., Lehmkuhl, F., 2012b. Environmental conditions in the Donggi Cona lake catchment, NE Tibetan Plateau, based on factor analysis of geochemical data. Journal of Asian Earth Sciences 44, 176–188.

IJmker, J., Stauch, G., Pötsch, S., Diekmann, B., Wünnemann, B., Lehmkuhl, F., 2012c. Dry periods on the NE Tibetan Plateau during the late Quaternary. Palaeogeography, Palaeoclimatology, Palaeoecology 346–347, 108–119.

Lehmkuhl, F., Schulte, P., Zhao, H., Hülle, D., Protze, J., Stauch, G., 2014. Timing and spatial distribution of loess and loess-like sediments in the mountain areas of the northeastern Tibetan Plateau. CATENA 117, 23–33.

Nottebaum, V., Lehmkuhl, F., Stauch, G., Hartmann, K., Wünnemann, B., Schimpf, S., Lu, H., 2014. Regional grain size variations in aeolian sediments along the transition between Tibetan highlands and north-western Chinese deserts – the influence of geomorphological settings on aeolian transport pathways. Earth Surf. Process. Landforms 39, 1960–1978.

Nottebaum, V., Stauch, G., Hartmann, K., Zhang, J., Lehmkuhl, F., 2015b. Unmixed loess grain size populations along the northern Qilian Shan (China): Relationships between geomorphologic, sedimentologic and climatic controls. Quaternary International 372, 151–166.

Opitz, S., Ramisch, A., IJmker, J., Lehmkuhl, F., Mischke, S., Stauch, G., Wünnemann, B., Zhang, Y., Diekmann, B., 2016. Spatio-temporal pattern of detrital clay-mineral supply to a lake system on the north-eastern Tibetan Plateau, and its relationship to late Quaternary paleoenvironmental changes. CATENA 137, 203–218.

Opitz, S., Wünnemann, B., Aichner, B., Dietze, E., Hartmann, K., Herzschuh, U., IJmker, J., Lehmkuhl, F., Li, S., Mischke, S., Plotzki, A., Stauch, G., Diekmann, B., 2012. Late Glacial and Holocene development of Lake Donggi Cona, north-eastern Tibetan Plateau, inferred from sedimentological analysis. Palaeogeography, Palaeoclimatology, Palaeoecology 337–338, 159–176.

Ramisch, A., Lockot, G., Haberzettl, T., Hartmann, K., Kuhn, G., Lehmkuhl, F., Schimpf, S., Schulte, P., Stauch, G., Wang, R., Wünnemann, B., Yan, D., Zhang, Y., Diekmann, B., 2016. A persistent northern boundary of Indian Summer Monsoon precipitation over Central Asia during the Holocene. Scientific Reports 6, 25791.

Yu, K., Hartmann, K., Nottebaum, V., Stauch, G., Lu, H., Zeeden, C., Yi, S., Wünnemann, B., Lehmkuhl, F., 2016. Discriminating sediment archives and sedimentary processes in the arid endorheic Ejina Basin, NW China using a robust geochemical approach. Journal of Asian Earth Sciences 119, 128–144.

4 Local case studies from the northern Tibetan Plateau

4.1 Environmental changes during the late Pleistocene and the Holocene in the Gonghe Basin, north-eastern Tibetan Plateau

Georg Stauch, Zhongping Lai, Frank Lehmkuhl, Philipp Schulte

Abstract

Different types of aeolian sediments are widespread in Gonghe Basin on the north-eastern Tibetan Plateau. Despite that individual sections are often discontinuous, the analysis of a large number of individual sections distributed in the entire basin provides a valuable archive for the reconstruction of environmental changes since the late Pleistocene. This study presents 43 new OSL (optical stimulated luminescence) ages and combines them with 39 previously published ages from the Gonghe Basin for the last 16 ka. From around 15 to 14 ka onwards an increase in moisture facilitated the formation of a vegetation cover and the fixation of loess sediments on the northern and southern side of the basin. After enhanced mobility of aeolian sediments during the dry Younger Dryas greater than before precipitation fixated also the sands in the lower parts of the basin. The increase in moisture can be attributed to the strengthening of the Asian summer monsoon. A comparison with the timing of aeolian sediment accumulation in the neighbouring basins probably indicates a time- transgressive strengthening of the monsoon on the north-eastern Tibetan Plateau. Aeolian sands were first permanently stored from around 17 ka at the Qinghai Lake in the north-east of the area and at 10.5 ka at the Donggi Cona in the south-west. The Gonghe Basin is located in an intermediate position. Highest moisture values occurred during the mid-Holocene from around 7.5 until probably 2.5 ka in the Gonghe Basin. In most parts of the basin aeolian activity was low during this time. In the central part of the basin aeolian sand movement occurred sporadically. Since 2.5 ka aeolian activity in most parts of the basin resumed due to the late Holocene weakening of the Asian summer monsoon. However, the parallel accumulation of aeolian sands and silts indicates an increased human influence.

Keywords: Palaeoclimate; Monsoon; Tibetan Plateau; Aeolian sediments; Dating

Introduction

The Asian summer monsoon is one of the most important atmospheric systems on the northern hemisphere (e.g. P. Wang et al., 2005; Y. Wang et al., 2014; Zhang et al., 2011) and determines the water budget in large parts of Asia (Immerzeel et al., 2010). During the recent decades a southward

26 shift of the East Asian monsoon was observed (Ding et al., 2009) as well as a reduction in precipitation over the eastern Tibetan Plateau (TP) (K. Yang et al., 2014). However, there are still considerable uncertainties regarding the causes (Li et al., 2010), spatial variability (J. Chen et al., 2015) and future development of these changes (K. Yang et al., 2014). The evaluation of past changes in the different atmospheric systems is an essential prerequisite for the understanding of present and future climate changes. During the last decades, the north-eastern TP (Fig. 4.1.1) became a key area for the reconstruction of the palaeoclimate evolution in Central Asia. At present the climate is controlled by the interplay of the two subsystems of the Asian summer monsoon, the East Asian summer monsoon and the Indian summer monsoon as well as by the westerlies and the Asian winter monsoon (e.g. Böhner, 2006; Henderson et al., 2010; Maussion et al., 2013). Palaeoclimate reconstructions on the north-eastern TP are mainly based on proxy interpretation from lake sediments (Qinghai Lake: e.g. Lister et al., 1991; Shen et al., 2005; Henderson and Holmes, 2009; An et al., 2012; Donggi Cona: Dietze et al., 2010; Mischke et al., 2010a; Opitz et al., 2012; Hala Lake: Wünnemann et al., 2012; Yan and Wünnemann, 2014) peat bogs (e.g. Hong et al., 2003; Schlütz and Lehmkuhl, 2009; Liu et al., 2013a) and terrestrial sediments (Madsen et al., 2008; Lu et al., 2011b; Yu and Lai, 2012; Stauch, 2015; Qiang et al., 2016). For the late Holocene additional information from tree rings (Shao et al., 2010; B. Yang et al., 2014) and speleothems (Zhang et al., 2008) are available. Despite that, there is still some dissimilarity in the interpretation of key features of the paleoclimate like the onset of the Asian summer monsoon in the area or the strength of the monsoonal system in the early and middle Holocene.

Aeolian sediments are widespread on the north-eastern TP (Stauch, 2015) and have been proven to be valuable palaeoclimatic archive (e.g. Thomas, 2013; Thomas and Burrough, 2016). The time of the last aeolian transport can be directly dated by optical stimulated luminescence (OSL) dating (e.g. Singhvi et al., 2001; Duller, 2004; Preusser et al., 2008). Nevertheless, no consensus has been reached regarding the climatic interpretation of the OSL ages from aeolian sediments. The accumulation of aeolian sand was frequently related to dry climate conditions (e.g. Mason et al., 2009; Lu et al., 2011a; Liu et al., 2013b) while loess accumulation was assigned with a shift to wetter climate conditions (Sun et al., 2007; Yu and Lai, 2014). Other studies from the north-eastern TP suppose a shift to wetter climate conditions for the accumulation of both sediment types (e.g. Stokes et al., 2003; Stauch et al., 2012; Yu and Lai, 2012). Periods of high aridity in combination with high wind speed would result in a frequent recycling of the aeolian sediments. The formation of a vegetation cover results in a decline of the movement of the aeolian sediments and the fixation of the sediments (Chase and Thomas, 2007; Lancaster, 2008; Thomas and Wiggs, 2008). A recent summary of OSL ages from aeolian sediments from the TP advocates a combined interpretation (Stauch, 2015). After the cold and dry glacial times, with a constant sediment turnover, wetter climate conditions during the late Pleistocene and the early Holocene resulted in a fixation of the aeolian sediments. During the late Holocene dryer climate conditions caused again a remobilization of the aeolian sediments. Nevertheless, recent studies also highlight local effects on aeolian archives. Qiang et al. (2016) stressed the need for an ‘activated’ sediment supply for fine grained sediments in the Gonghe Basin and related the occurrence of these sediments to the incision of the Huang He (Yellow River). In the eastern Qaidam Basin fluvial erosion might have removed a large proportion of the pre-Holocene aeolian deposits (Yu et al., 2015), which would be another explanation for missing Pleistocene deposits. An additional hypothesis for early Holocene sand movement during times of a stronger monsoon was presented by (Qiang et al., 2013a) with a special focus on the TP. They suggested that higher insolation values offset higher precipitation values which in turn resulted in 27 reduced vegetation and the mobilization of aeolian sediments. In general, aeolian sediment sections on the north-eastern TP are often influenced by spatially heterogeneous accumulation rates and erosional events (Qiang et al., 2013a; Stauch et al., 2014; Yu et al., 2015), which complicates the palaeoclimatic interpretation from single sections or small datasets.

Fig. 4.1.1: Overview of the north-eastern Tibetan Plateau and the catchment of the Gonghe Basin.

During the last years, the number of OSL ages of aeolian sediments has been increasing rapidly. OSL ages are available from the eastern Qaidam Basin (Yu and Lai, 2012, 2014; Yu et al., 2015), the Qinghai Lake catchment (Madsen et al., 2008; Lu et al., 2011b, 2015, Liu et al., 2011, 2012), the Donggi Cona area (Stauch et al., 2012, 2014) and the Gonghe Basin (Liu et al., 2013b; Qiang et al., 2013a, 2016). An evaluation of all the published OSL ages from north-eastern Tibet indicates different accumulation phases of aeolian sediments since the late glacial (Stauch, 2015). Aeolian sands on the north-eastern TP accumulated predominantly between 14 and 7 ka. However, there are still some considerable differences between the different basins. At the Qinghai Lake, accumulation of aeolian sands started already at 17 ka, which is much earlier than in the neighbouring basins. Additionally, there is no consistent relationship in the timing of the loess accumulation in the different basins on the north-eastern TP (Stauch, 2015). At Qinghai Lake and Donggi Cona, a phase of loess accumulation occurred during the late glacial and the early Holocene whereas loess in the eastern Qaidam Basin was deposited from 8 ka to ~2 ka. From the Gonghe Basin, despite being located in the central part of north-eastern Tibet, no clear phase of loess accumulation could be identified. One of the reasons might be that most of the dated sections are located on the southern and eastern side of the basin which might bias the overall climatic analysis.

Beside aeolian deposits, previous palaeoclimate reconstructions in the Gonghe Basin (Fig. 4.1.2) are based on lake sediments (Chaka Salt Lake: X. Liu et al., 2008; Genggahai Lake: Song et al., 2012; Qiang

28 et al., 2013b, 2014, Rao et al., 2014, 2016b; Dalian Lake: Yan et al., 2002) and peat deposits (Liu et al., 2013a). Generally, during the late glacial an increase in temperature and precipitation was detected, followed by climate deterioration around the Younger Dryas period (e.g. X. Liu et al., 2008). Subsequently, the Asian summer monsoon strengthened again. Nevertheless, there is no consensus about the Holocene palaeoclimate evolution in the Gonghe Basin. Based on the organic content in the Chaka Salt Lake a warm and wet climate was assumed from 11 to 6 cal BP (X. Liu et al., 2008). Similar results were obtained for the palaeoproductivity, reduced aeolian influx and δDn-alkanes of aquatic plants in the Genggahai Lake (Song et al., 2012; Qiang et al., 2014; Rao et al., 2016b). In contrast, Liu et al. (2013b) assumed a relatively dry climate from 8.7 to 4.7 cal BP. Based on the peat deposits in the eastern Gonghe Basin, cold and dry climate conditions were reconstructed for the period from 8.5 to 7.6 cal. BP and a strong summer monsoon with high precipitation from 7.6 to 3.8 cal BP (Liu et al., 2013a). For the late Holocene most studies assume a return to dryer conditions, except a reconstruction based on aeolian sediments which assumed a warm and wet climate from 4.7 to 0.7 ka (Liu et al., 2013b).

The aim of this study is to provide new insights into the aeolian system in the Gonghe Basin since the late Pleistocene (16 ka). Sedimentological and geochronological analyses are extended to cover all the different sediment types available in the basin. Consequently, the factors governing aeolian accumulation in the basin are evaluated and compared to the neighbouring basins. Finally, the environmental factors and especially the influence of the Asian summer monsoon are assessed.

Study area

The Gonghe Basin (36°00’N, 100°28’E) has a size of ~20.000 km² and is located at a mean altitude of 3.000 m asl (Fig. 4.1.2). It is the largest intramontane basin on the north-eastern TP. The formation of the basin started at ~14-6 Ma ago with the uplift of the Ela Shan in the south-west and Qinghai Nan Shan on the northern side (Lu et al., 2012; H.-P. Zhang et al., 2012). The uplift of the ranges is still continuing. The last major earthquake occurred in 1990 with a magnitude of 6.9 (e.g. Hao et al., 2012). The basin is filled by several hundred meters of Neogene and Quaternary sediments, basically consisting of fluvial gravels and lacustrine sediments (Craddock et al., 2010; Perrineau et al., 2011). The Huang He has incised up to 500 m into these sediments (Métivier et al., 1998). There are still some uncertainties regarding the onset of the incision. Based on cosmogenic radionuclide dating the ages vary between 500 ka (Craddock et al., 2010) and 150 ka (Perrineau et al., 2011). Since 1985 the Huang He is artificially dammed at the north-eastern side of the basin (Yan et al., 2002). The Gonghe Basin is hydrological subdivided into four different catchments. The north-western part drains into the Chaka Salt Lake, the central part into several small endorheic sub-basins and the two south- eastern catchments are drained by the Huang He. The climate in the Gonghe basin is semiarid with a mean annual air temperature of 3.7°C and a mean annual precipitation of ~310 mm. 80% of the precipitation is occurring between May and September, the time when the region is influenced by the Asian summer monsoon (Qiang et al., 2013a). During winter and spring, strong winds from north- western and northern directions are conveying cold and dry air masses to the study area (Qiang et al., 2013a). During the recent years large parts of the basin are affected by desertification processes, which are mainly a consequence of increased life stock amounts since the mid of the 20th century and an attempt to cultivated farmland in the semi-arid basin in the late 1980ies (Yan et al., 2002, 2009). 29

Fig. 4.1.2: Map of the Gonghe Basin with the different catchment areas and location of the aeolian sections sampled for this study. Red circles in the background indicate OSL sampling points. Previously published sections are also marked (LG: Liu et al., 2013b; TGM, MGTA, MGTB and GMY: Qiang et al., 2014; QJ, LG, TX, ML, SG: Qiang et al., 2016; KE peat section: Liu et al., 2013a).

Material and method

During two field trips in 2012 and 2013, detailed geomorphological mapping of aeolian sediments was conducted. The extent of the dune-fields was mapped by Landsat 8 and Corona images (Fig. 4.1.2). To disentangle the spatial distribution of the sediment, 103 sections were sampled in the central and eastern part of the Gonghe Basin, resulting in 506 individual samples. A total of 43 OSL samples were taken from 15 sections (Tab. 4.1.1, 4.1.2).

Sedimentology

For grain-size measurements, samples were dried and aggregates were mechanically destroyed. Particles larger than 2 mm were removed by sieving. Further treatment followed the procedure of

Schulte et al. (2016). Organic matter was destroyed by H202 (30%). Before the measurement, sodium pyrophosphate (Na4O7P2) was used for dispersion. The grain-size measurement was conducted with a Beckmann Coulter LS 13320 laser diffraction analyser with 116 classes ranging from 0.04 to 2000 μm. Each sample was measured at least twice. The calculation of the frequency distribution was based on

30 the Lorenz-Mie theory (refractive index (RI) for water: 1.33, RI for the sample: 1.55, absorption coefficient of 0.1; Özer et al., 2010).

Samples were first classified in the field according to the underlying process of deposition (e.g. aeolian, fluvial, and lacustrine). A further classification was based on the 345 aeolian samples from 65 sections. Due to the large number of aeolian sediment samples a k-means cluster analysis was performed. However, polymodal grain-size distributions from sand sections were excluded, because they cannot be properly classified by the k-means algorithm.

OSL dating

OSL dating was conducted under red light in the luminescence laboratory. Samples were first treated with 10% HCl and 30% H2O2 to remove carbonates and organics, respectively, and then sieved to obtain the 38-63 μm and 90-125 μm grain fractions. To pure quartz extraction, 38-63 μm grains were etched using 35% H2SiF6 for about 10 days to remove feldspar. The 90-125 μm grains were etched by 40% hydrofluoric acid for 40 min. The purity of quartz grains was checked by infrared stimulation.

Samples with obvious infrared stimulated luminescence signals were re-treated with H2SiF6. Quartz samples were then mounted onto the central 0.7 cm of a 0.97 cm diameter stainless steel disc using silicone oil. OSL measurements were carried out on an automated Risø TL/OSL-DA-20 reader with stimulation by blue LEDs at 130 °C for 40 s, and detection was using a 7.5 mm thick U-340 filter. For all samples, preheat was using 260 °C for 10 s, and cut-heat 220 °C for 10 s, determined after preheat plateau and dose recovery tests. OSL signals of the first 0.64 s stimulation were integrated to construct growth curve after background subtraction (using the last 25 channels in the shine down curve).

Concentrations of uranium, thorium and potassium were obtained by neutron activation analysis in the China institute of Atomic Energy in Beijing. For the 38-63 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 elemental concentrations were converted into an annual dose rate according to Aitken (1998).

De was determined using the combination of the SAR (Murray and Wintle, 2000) and the standardized growth curve (SGC) (Roberts and Duller, 2004; Lai, 2006; Lai et al., 2007), i.e. SAR-SGC method. For each sample, six aliquots were measured by SAR to construct a SGC growth curve, then 12 additional aliquots were measured for natural luminescence (Ln) and luminescence induced by the test dose (Tn). Sensitivity-corrected natural OSL (Ln/Tn) was projected on the SGC to yield a De. De results determined by the SGC are in agreement with those by the SAR protocol within 10% for most samples. For all samples, the final De is the mean of all SAR Des and SGC Des measured by large and small aliquots. OSL signals are dominated by fast components, as expected for aeolian samples. Recuperation of all samples is negligible.

Table 4.1.1: Site information for sections with OSL ages.

No. OSL Section Northing Easting Altitude [m] samples GB032 36.449 100.544 3494 1 GB024 36.300 100.436 2985 1 GB029 36.114 100.416 2981 2 GB031 36.117 100.415 2988 7 GB032 36.113 100.416 2986 3 GB037 35.960 100.131 3178 2 GB038 35.925 100.061 3264 3 GB039 35.927 100.061 3263 2 GB044 36.200 100.639 2881 3 GB065 36.377 100.372 3164 3 GB073 36.255 99.770 3125 2 GB075 35.685 100.317 3190 4 GB082 36.673 101.097 3635 4 GB084 35.742 101.008 3391 3 GB091 36.014 100.563 2912 3

Table 4.1.2: Environmental radioactivity and OSL dating results.

Section Sample Depth Grain size K Th U Water Dose rate Aliquot De OSL age ID (m) (μm) (%) (ppm) (ppm) content (%) (Gy/ ka) Num. (Gy) (ka) GB023 GO-01 1.0 38-63 1.79±0.06 10.8±0.31 2.62±0.10 11.1±5 3.35±0.22 6a+12b 41.48±1.51 12.40±0.95 GB024 GO-02-02 0.30 38-63 1.62±0.05 9.87±0.30 3.41±0.12 15.9±5 3.11±0.21 6a+12b 39.62±1.17 12.73±0.93 GB029 GO-04-01 0.30 38-63 1.40±0.05 6.3±0.21 1.5±0.07 0.7±5 2.67±0.18 6a+12b 1.38±0.1 0.52±0.05 GB029 GO-04-02 0.55 38-63 1.40±0.05 6.45±0.22 1.5±0.07 3.3±5 2.62±0.18 6a+12b 9.08±0.27 3.47±0.26 GB031 GO-05-01 0.20 38-63 1.46±0.05 8.01±0.25 1.83±0.08 0.5±5 3.00±0.21 6a+12b 0.41±0.04 0.14±0.02 GB031 GO-05-02 0.45 38-63 1.43±0.05 7.18±0.24 1.65±0.08 0.3±5 2.86±0.22 6a+12b 0.61±0.07 0.21±0.03 GB031 GO-05-03 0.70 38-63 1.53±0.05 8.5±0.26 2.0±0.09 0.8±5 3.08±0.21 6a+12b 1.42±0.12 0.46±0.05 GB031 GO-05-04 0.90 38-63 1.43±0.05 6.73±0.22 1.48±0.07 0.4±5 2.72±0.19 6a+12b 2.17±0.12 0.80±0.07 GB031 GO-05-05 1.00 38-63 1.42±0.05 5.92±0.2 1.1±0.06 0.3±5 2.54±0.18 6a+12b 2.21±0.09 0.87±0.07 GB031 GO-05-06 1.45 38-63 1.54±0.05 8.53±0.26 1.96±0.09 4.3±5 2.97±0.20 6a+12b 6.27±0.22 2.11±0.16 GB031 GO-05-07 1.75 90-125 1.47±0.05 6.2±0.2 1.48±0.07 3.8±5 2.59±0.18 6a+12b 12.47±0.31 4.82±0.35 GB032 GO-06-01 0.30 38-63 1.52±0.05 8.33±0.26 1.92±0.08 1.4±5 2.99±0.21 6a+12b 4.17±0.18 1.39±0.11 GB032 GO-06-02 0.55 38-63 1.37±0.05 5.8±0.2 1.27±0.07 0.2±5 2.50±0.17 6a+12b 8.59±0.48 3.43±0.31 GB032 GO-06-03 3.30 90-125 1.38±0.05 6.42±0.22 1.18±0.06 4±5 2.21±0.15 6a+12b 7.22±0.33 3.26±0.27 GB037 GO-07-02 0.7 38-63 1.74±0.06 10.4±0.3 2.81±0.11 2.6±5 3.67±0.25 6a+12b 41.83±1.23 11.40±0.84 GB037 GO-07-03 1.15 38-63 1.77±0.06 9.41±0.28 3.02±0.11 17.9±5 3.03±0.2 6a+12b 45.11±0.91 14.88±1.04 GB038 GO-08-01 0.4 38-63 1.34±0.05 7.57±0.24 3.03±0.11 12.7±5 2.79±0.19 6a+12b 28.07±0.71 10.07±0.72 GB038 GO-08-02 0.75 38-63 1.82±0.06 11.1±0.31 3.19±0.11 2.7±5 3.91±0.26 6a+12b 37.01±1.59 9.46±0.76 GB038 GO-08-03 1.25 38-63 1.80±0.06 10.9±0.32 2.86±0.11 0.3±5 3.86±0.26 6a+12b 45.25±1.09 11.71±0.85 GB039 GO-09-01 0.35 38-63 1.88±0.06 12.4±0.35 2.32±0.09 12±5 3.45±0.24 6a+12b 15.31±0.6 4.44±0.35 GB039 GO-09-02 1.6 38-63 1.78±0.06 10.2±0.3 2.95±0.11 14±5 3.22±0.22 6a+12b 41.12±0.97 12.78±0.91 GB065 GO-10-01 0.23 38-63 1.82±0.06 11.3±0.32 2.36±0.09 5.4±5 3.60±0.25 6a+12b 2.23±0.1 0.62±0.05 GB065 GO-10-02 0.55 38-63 1.67±0.06 10.6±0.31 2.71±0.10 9.2±5 3.33±0.22 6a+12b 3.70±0.09 1.11±0.08 GB065 GO-10-03 0.95 38-63 1.81±0.06 10.6±0.32 2.92±0.11 5.4±5 3.65±0.25 6a+12b 37.58±0.94 10.29±0.74 GB073 GO-12-01 0.42 38-63 1.71±0.05 10.5±0.3 3.07±0.11 3.4±5 3.70±0.25 6a+12b 6.74±0.18 1.82±0.13 GB073 GO-12-02 1.0 38-63 1.96±0.05 11.8±0.33 3.17±0.11 6.2±5 3.92±0.26 6a+12b 27.72±0.69 7.08±0.51 GB075 GO-13-01 0.4 38-63 1.62±0.05 10.1±0.29 2.6±0.10 11.3±5 3.15±0.21 6a+12b 24.15±0.4 7.67±0.53 GB075 GO-13-02 0.75 38-63 1.60±0.05 9.66±0.29 2.96±0.11 5.4±5 3.40±0.23 6a+12b 37.58±0.86 11.04±0.78 GB075 GO-13-03 1.1 38-63 1.79±0.06 10.7±0.31 2.67±0.10 6.3±5 3.53±0.24 6a+12b 43.43±0.93 12.30±0.87 GB075 GO-13-04 1.5 38-63 1.85±0.06 10.8±0.31 3.22±0.11 7.5±5 3.67±0.25 6a+12b 49.59±1.14 13.51±0.97 GB082 GO-14-01 0.77 38-63 1.41±0.05 7.73±0.25 1.86±0.09 4.1±5 2.84±0.19 6a+12b 2.20±0.09 0.78±0.06 GB082 GO-14-02 1.65 38-63 1.47±0.05 8.21±0.25 1.73±0.08 6.1±5 2.78±0.19 6a+12b 5.80±0.18 2.09±0.16 GB082 GO-14-03 2.7 38-63 1.56±0.05 8.45±0.26 1.92±0.09 7.3±5 2.85±0.2 6a+12b 33.42±1.05 11.72±0.88

GB082 GO-14-04 3.35 38-63 1.61±0.05 9.69±0.29 2.06±0.09 9±5 2.94±0.2 6a+12b 34.44±1.65 11.70±0.98 GB084 GO-15-01 0.25 38-63 1.23±0.05 5.1±0.18 1.21±0.07 3.6±5 2.30±0.17 6a+12b 2.460.05 1.07±0.08 GB084 GO-15-02 0.5 38-63 1.35±0.05 7.28±0.24 1.4±0.07 7.1±5 2.52±0.17 6a+12b 6.03±0.2 2.39±0.18 GB084 GO-15-03 0.95 90-125 1.28±0.05 3.86±0.16 0.90±0.06 2.2±5 1.92±0.17 6a+12b 17.75±0.97 9.24±0.97 GB044 GO-16-01 0.25 38-63 1.40±0.05 6.96±0.23 1.73±0.08 3.9±5 2.72±0.19 6a+12b 5.09±0.14 1.87±0.14 GB044 GO-16-02 0.4 38-63 1.43±0.05 7.8±0.25 2.45±0.1 8±5 2.87±0.2 6a+12b 19.56±0.69 6.82±0.52 GB044 GO-16-03 1.15 38-63 1.83±0.06 10.5±0.30 2.85±0.11 3.8±5 3.69±0.25 6a+12b 36.12±0.72 9.79±0.7 GB091 GO-17-01 0.45 38-63 1.45±0.06 5.49±0.19 1.18±0.06 0.6±5 2.57±0.18 6a+12b 1.07±0.06 0.41±0.04 GB091 GO-17-02 0.7 90-125 1.46±0.06 7.18±0.24 1.67±0.08 3.1±5 2.74±0.19 6a+12b 3.02±0.14 1.10±0.09 GB091 GO-17-03 1.15 38-63 1.45±0.06 5.76±0.20 1.19±0.07 1±5 2.55±0.18 6a+12b 17.68±0.14 6.93±0.66 a: aliquot number using single aliquot regenarative-dose (SAR) protocol, b: aliquot number using standard growth curve (SGC) method. As the Des of SAR agree well with the Des of SGC, the De used for final age calculation are the mean of Des of both SAR and SGC.

Results

Sedimentology

The k-means cluster analysis of the 345 aeolian sediment samples revealed four distinct grain-size classes. Class I consist of 56 sediment samples. The median of the grain size is between 20 and 40 μm and the mean mode at around 45 μm (Fig. 4.1.3). According to the grain size parameters and the related aeolian transport the samples of class I were termed loess. Class II (n: 71) has a somewhat finer mode at around 36 μm and sometimes a secondary mode at 5 to 10 μm. Accordingly the median varies between 10 and 24 μm. Sediments of class II always occur in sections with loess sediments. All samples from soils and palaeosols in loess sections represent class II which was therefore termed loess soils. The grain-size distribution is a modification of the original loess. In many sections the grain-size modification continues below the macroscopically identifiable soil horizons.

Fig. 4.1.3: Sediment types (A: loess (I); B: loess soils (II); C: fine sand (III); D: sand (IV); E: bimodal sand(V)).

The 122 samples of class III have a distinct mode at around 110 μm. Median values have a large spread between 60 and 120 μm. The class was termed fine sand. Class IV (n: 96) has a coarse mode at 250 μm. However, there are several samples which are distinctively coarser. Median values are

35 between 150 and more than 300 μm. Class IV was labeled sand. Class V (n: 53) was not included in the k-means classification because these grain-size distributions are mainly bimodal. They represent a mixture of the loess group and the two sand classes, with one mode at 40 μm and another one at 110 μm or larger. Due to this characteristic the median values show large variations. The bimodal grain-size distribution indicates a parallel or seasonal accumulation of two different grain-size fractions. As this class is frequently associated with weak palaeosols in sandy deposits, a trapping of silty sediments by vegetation can be assumed. The slight increase in the fraction <10 μm might be related to pedogenic processes.

Spatial variations of aeolian sediments

Loess has been deposited particularly on the northern and southern slopes of the basin and on adjacent alluvial fans (Fig. 4.1.4). The thickness of the loess cover is between 80 and 180 cm. However, differences are apparent regarding the stratigraphic composition of the loess sections on the northern and southern side. On the northern side, a palaeosol is well developed in all six sections (e.g. section GB023, GB065; Fig. 4.1.4). In contrast, no well-expressed palaeosols were observed in the five longer sections on the southern side (e.g.GB037, GB038). At two sections in the central basin (GB019 and GB020) bimodal sediments with high loess content are covered by aeolian sand. At section GB044 (Fig. 4.1.4) loess is covered by reworked sediments and fine sand. The sections indicate a probably larger former extent of the surface loess cover than at present. In contrast, in none of our sections loess was observed above sandy aeolian deposits.

Sandy deposits are typical for the central and eastern part of the basin where large dune fields cover an area of 1.134 km², exhibiting a transition from dune fields to sand-sheets toward the north- western side (Fig. 4.1.2). At some locations bimodal grain-size distributions were found in the upper part of the sections. However, especially in the central basin (e.g. section GB032) they are again covered by sand. In section GB031, fluvial gravels are overlain by an alternating sequence of aeolian sand and two palaeosols. The uppermost part consists of fine sand and is considerably finer than the lower part. In the nearby section GB032, the fluvial gravels at the bottom are covered by 325 cm of dune sand and 40 cm of bimodal sand and again 10 cm of dune sand. Despite the close vicinity to section GB031 only a weak palaeosols in the upper part of the section was preserved.

The eastern side of the basin is dominated by fine sands. Two sections ~200 m apart from each other were sampled (GB082 and GB083) (Fig. 4.1.2). The deposits have thicknesses of 4.5 and 5 m, respectively, and include a well-developed palaeosol and up to three weak palaeosols (Fig. 4.1.5). The mode of all grain-size distributions is at ~111 µm. Section GB083 is identical to section GMY which has already been analyzed by Qiang et al. (2013a). As the OSL sampling locations were still visible in 2013, only samples for grain-size analysis were taken. In contrast to the classification scheme used by Qiang et al. (2013a) we classified the sediments as fine sands due to the median larger than 100 μm and not as loess. Fine sands dominate also the nearby section GB084 and section GB089 on the southern side of the eastern basin. In both section only one weak palaeosol was preserved.

Fig. 4.1.4: Transect through the Gonghe Basin showing main sedimentary units.

Fig.4.1.5: Selected sections in the eastern part of the Gonghe Basin (red circles: OSL ages from this study; green circles: OSL ages published by Qiang et al. (2013a), sample with an x was rejected). For the explanation see Fig.4.1.4. 37

OSL ages

43 OSL ages from 15 sections were analysed (Tab. 4.1.2). 13 of the ages are from loess, six are palaeosols in loess, 9 ages are from fine sands, another 9 from sands, five samples were taken from sandy palaeosols and one OSL age was obtained from reworked sediments.

The deposition of the loess started at least at 14.88 ± 1.04 ka. This age was obtained at section GB037 in 120 cm depth on the southern side of the Gonghe Basin (Fig. 4.1.4). Strong loess accumulation continued until 9.5 ka (Fig 4.1.6). Younger loess sediments were sporadically deposited at 7 to 7.5 ka, at 4.5, 1.8, 1.1 and 0.62 ka. Except for the youngest sample, all sediments younger than 8 ka were classified as loess with palaeosols. An additional loess palaeosol sample with an age of 10.07 ± 0.72 ka was dated to the early Holocene. In this case a pedogenetic modification of older sediments can be assumed. All sections with ages from loess sediments are located in the mountains or at the food-slopes on the northern and southern side of the basin in altitudes between ~3000 and 3500 m. An exception is one loess OSL sample (9.8±0.9 ka) from section GB044 which is located in the northern central part of the Basin in an elevation of 2880 m. Additionally, fine sand accumulated at the end of the Pleistocene in the eastern side of the Gonghe Basin (GB082). Aeolian accumulation of fine sand resumed in the late Holocene from around 2 ka onwards. In section GB083 only sediment samples were taken as this section has already been dated by Qiang et al., (2013a, GMY). Despite some diverged sedimentological interpretation the ages are nearly identical to the ages from section GB082 (Fig. 4.1.5). Nevertheless, we excluded sample GMY-10 (12.6 ± 1.4 ka) due to possible modification of the sediment by bioturbation, which would also explain the relatively high error compared to the other ages from the section. Additional sections with late Holocene fine sands are located a in the central part (e.g. GB031, GB032, GB044; Fig. 4.1.4), with ages ranging from 1.8 ± 0.1 to 0.14 ± 0.02 ka.

Fig. 4.1.6: OSL ages from the Gonghe Basin (circles: this study; boxes: previously published ages; Liu et al., 2013a; Qiang et al., 2013a, 2016). The lower bars denote phases of stronger sand and loess accumulation. 38

Sand ages date sporadically to the early and middle Holocene, but show a strong cluster in the late Holocene. The oldest age (9.24 ± 0.97 ka) is from the basal part of section GB084 east of the Huang He. On one of the lower terraces on the western side of the river an age of 6.93 ± 0.66 ka (section GB091) was obtained. The two other sections with sand ages (GB031 and GB032) of mid- to late Holocene age are located in the center of the Gonghe Basin. In section GB032 about 3 m of aeolian sand were accumulated at around 3.3 ka. Both sections also contain palaeosols. In section GB032 the age of the palaeosol is 1.39 ± 0.11 ka and in section GB031 ages of 2.11 ± 0.16 ka and 0.46 ± 0.05 ka were obtained. Additional sediments with sandy palaeosol were dated in section GB029 (3.47 ± 0.26 ka).

Discussion

Aeolian sediments in the Gonghe Basin

According to the results from this study loess was deposited at the end of the Pleistocene until the early Holocene while sand accumulation occurred mainly during the late Holocene, i.e the last 2ka. OSL ages from palaeosols developed in loess as well as ages from bimodal sands indicate a weak accumulation of silty particles also during the late Holocene. This difference in the timing of the accumulation of the different sediment classes is not in concordance with previous results from the Gonghe Basin (Liu et al., 2013b; Qiang et al., 2013a, 2016). Previously published OSL ages show an enhanced aeolian accumulation of sand from 8 to 12 ka and loess accumulation mainly during the last 5ka. This indicates a strong spatial variability in sediment distribution in the basins on the north- eastern TP and highlights the need for a basin wide analysis to decipher the environmental history of the area. While most of the previously dated sections are located on the southern side of the basin (QJ, TGM, LG) and in the eastern basin (TX, ML, MGTA, MGTB, SG, GMY; Fig. 4.1.2), ages from this study also include the central and northern part of the Gonghe Basin. The combined dataset shows an onset of the loess accumulation in the late Pleistocene and a nearly parallel accumulation of loess and sand in the Gonghe Basin during the early Holocene (Fig. 4.1.6). Two sections in the southern part (TGM: Qiang et al., 2013a; LG: Liu et al., 2013b) provide information about an early Holocene sand accumulation. At the same time loess was deposited in the nearby sections on the southern side as well as on the northern side. In the eastern part of the basin early Holocene aeolian sand and fine sand deposits are preserved in numerous sections (e.g. MGTA, MGTB, SG, GB082, GB083). Strongly developed palaeosols in loess sediments (e.g. section LG: Liu et al., 2013b; ML, QJ, SG: Qiang et al., 2016) are mainly younger than 5 ka. This is in accordance with the new ages from this study (section GB039, GB065, GB073, GB075). Additionally, five samples from palaeosols developed in sand or fine sand in the center of the Basin (GB029, GB031, GB032) have ages between 3.5 and 0.5 ka. Despite this, there are no age clusters of palaeosols evident during the late Holocene. This may be partly due to the fact that ages of palaeosols are difficult to assess precisely. Soils are prone to bioturbation, which might blur the original depositional age of the sediment. Additionally, it is impossible to be determined from the sedimentary body if the accumulation of the sediment occurred parallel to the soil formation or if older sediments were pedogenetically modified (Bateman et al., 2003a; Stevens et al., 2007; Stauch, 2015).

In general, all studies from aeolian sediments in the Gonghe Basin indicate a reduced aeolian accumulation from 7.5 ka to the late Holocene. Large breaks in the sedimentation are especially

39 obvious in the loess sections on the northern (GB065) and southern side (e.g. section QJ: Qiang et al., 2016) as well as in the sand section in the eastern part of the basin (GB082, GB084). A gap in sedimentation is also indicated by the missing early Holocene ages in the central part of the basin (GB029, GB031, GB032). Furthermore, the age of the reworked sediments in section GB044 of 6.82±0.5 ka is a first indicator of a mid-Holocene erosional phase.

At present there is a distinct spatial separation between the loess and the sand deposits in the Gonghe Basin. Loess is preserved at the northern and southern side of the basin (Fig. 4.1.4), whereas sands and active dunes are located in the center. This can be explained by different wind speeds due to topographic effects. Higher wind speeds transporting coarser grains prevail in the center, while the wind-speed at the sides of the basin is reduced resulting in absence of coarser grains. This spatial differentiation in the Gonghe Basin was previously recognized by Qiang et al. (2016). A similar spatial distribution was also observed in other basins on the north-eastern TP, e.g. in the Donggi Cona catchment (Stauch et al., 2012) or in the northern Qilian Shan (Nottebaum et al., 2014, 2015a). Besides differences in wind speeds different source areas and alternating sediment supply was proposed to explain the sediment distribution in the basin. Based on a more limited dataset Qiang et al. (2016) proposed that the accumulation of the sand and the loess occurred independently. The differences in the timing of the different archives were related to the activation of different source areas through time. While for the silty sand deposits a fluvial origin was proposed, phases of sand accumulation were related to changes in the local vegetation cover (Qiang et al., 2016). The silty sand is a little finer than fine sand investigated here. Nevertheless, to reduce the number of sediment classes we reclassified them into fine sand for figure 4.1.6. The now larger dataset from aeolian sediments in the Gonghe Basin indicates a partly parallel accumulation of fine grained and coarse grained sediments during the early Holocene. Therefore, climatic conditions in the area of accumulation might play a more important role than previously assumed. Despite that, there is an apparent late onset of aeolian sand accumulation in the central part of the basin in comparison to the sides of the basin. The oldest OSL age west of the Huang He has an age of 6.9 ka (GB091) while further to the west aeolian sands have an age of 4.8 ka (GB031). As the area is prone to deflation it can be assumed, that older aeolian sediments in the central part of the basin were removed. This might either be caused during aeolian reactivation periods or due to fluvial reworking. This assumption is further supported by the occurrence of loess in the central part of the basin. The loess below the sands indicates a probably larger extent of the loess cover during the early Holocene. In the northern central basin loess sediments with an early Holocene age have been found below aeolian sands (GB044) as well as undated bimodal sediments in section GB019 and GB020. Whether a larger loess cover existed in the early Holocene Gonghe Basin or there were sands deposits remains an open question.

Comparison with other areas

A special feature in the Gonghe Basin is the nearly 2000 years earlier onset of loess accumulation than the enhanced sand accumulation during the transition from the late Pleistocene to the early Holocene. In the other areas of the north-eastern TP, accumulation of sand started earlier than the accumulation of loess. In the Qinghai Basin north of the Gonghe Basin, enhanced accumulation of aeolian sand started already at 17 ka and lasted until 6.5 ka, with a short break at 11ka. In contrast to that, enhanced loess accumulation was recorded from 14.9 to 8.1 ka. Enhanced sand accumulation in 40 the Qaidam Basin ranged from 13 to 6.4 ka, while the loess accumulation started at around 8 ka. Enhanced aeolian accumulation of loess and sand in the Donggi Cona catchment did not start until the early Holocene (at around 10.5 ka) (Stauch, 2015). In contrast to that, the combined dataset from the Gonghe Basin shows an onset of the loess accumulation at around 14 to 15 ka (Fig. 4.1.6), while sand accumulation did not started until 12.5 ka. Enhanced loess accumulation ceased at around 8.5 ka, while the sand accumulation continued at least until 8.5 but probably until 7 ka. During the late Holocene, aeolian accumulation of sand resumed at around 3.6 ka in the Qaidam Basin, 3.2 ka at the Donggi Cona, at around 2 ka in Gonghe Basin and at 1.9 ka at the Qinghai Lake (Stauch, 2015).

Palaeoclimate interpretation

Late Pleistocene – Holocene transition

In the Gonghe Basin, OSL ages from loess date back to 15 to 14 ka and from aeolian sands at 12.5 ka. This timing reflects the gradual increase in moisture due to the strengthening of the Asian summer monsoon. An increase in precipitation leads to the formation of a vegetation cover and the trapping of the aeolian sediments. This effect is first evident at the loess sites located higher up in the mountains and therefore receive more precipitation due to orographic effects. This is also in accordance with the loess age in the section GB044, which is younger than the basal ages in the mountains. Additionally, silty sediments have a higher capability to retain water and consequently higher soil moisture than sandy sediments. The timing of the final trapping of the loess at around 14 to 15 ka is synchronous with a shift to wetter climate conditions as documented in other proxy records from the north-eastern TP: in the Genggahai Lake in the Gonghe Basin an increase in palaeoproductivity (Song et al., 2012) and rising lake levels (Qiang et al., 2013b) from 15.3 ka onwards were described. Pollen records from the nearby Qinghai Lake indicate an increase of humidity from around 14.1 ka. A similar timing for the onset of the Asian summer monsoon was recorded in d18O values from the lake (Colman et al., 2007) and recently at 13.6 ka from leaf wax hydrogen isotopes (Thomas et al., 2016). A likewise shift to more humidity was documented for the Donggi Cona at 13.5 ka (Y. Wang et al., 2014) and the Kuhai Lake (12.8 ka, Mischke et al., 2010c). This is also in accordance with the first decrease in d18O values from Dongge and Hulu cave, indicating an increase in the Asian summer monsoon intensity (Dykoski et al., 2005). This initial moisture increase was punctuated by the cold and dry Younger Dryas at 11.7 ka (e.g. Wang et al., 2001; Yu and Kelts, 2002; Herzschuh, 2006). The inception of enhanced sand accumulation and preservation at the margins of the Gonghe Basin at 12.5 ka reflects this last sand movement at the end of the late Pleistocene. Additionally, there was an increase in loess accumulation, as indicated by a higher number of OSL ages during this time. Subsequently, moisture supply increased and facilitated again the development of vegetation cover. Thus, the early Holocene moisture increase led to a reduction in sediment availability for the Gonghe Basin. This Asian summer monsoon enhancement after the Younger Dryas was documented in many records on the north-eastern TP, e.g. Genggahai Lake (Song et al., 2012; Rao et al., 2016b), Chaka Salt Lake (X. Liu et al., 2008), Lake Qinghai (Colman et al., 2007; An et al., 2012; Thomas et al., 2016), the Donggi Cona (Y. Wang et al., 2014) and the Hongyuan peat record (Hong et al., 2003). The increase in moisture during the late Pleistocene on the north-eastern TP can also be observed in the aeolian deposits in the neighbouring basins. Additionally, the onset of the enhanced sand accumulation started earliest at 17 ka at the Qinghai Lake on the edge of the TP, at 13 and 12.5 ka in the Qaidam and Gonghe Basin, respectively, while the latest onset is recorded at 41 the Donggi Cona at 10.5 ka, which is the westernmost basin. This trend might indicate the westward intrusion of the Asian summer monsoon on the north-eastern TP. The direction from the north-east to the south-west is in accordance with the present wind field during summer (Maussion et al., 2013) and reflects the topographic influence of the high mountains on the north-eastern TP.

This summary of the OSL ages from the early Holocene only captures the long-term trend. Frequent short term environmental changes might have a profound effect on the aeolian sedimentary system. Warmer and wetter climate might lead to the formation of a now buried palaeosol in the north-east of the basin at around 10 ka (section SG, Qiang et al., 2016). Nevertheless, sedimentary sections in the Gonghe Basin are very heterogeneous and a basin-wide comparison of short-term environmental changes in the early Holocene is not possible at the moment.

Mid-Holocene

During the mid-Holocene the accumulation of aeolian sediments in the Gonghe Basin was strongly reduced. There was only an occasional accumulation of aeolian sand from around 7.5 until 2.5 ka. However, in the central part of the basin aeolian accumulation occurred around 5 ka and an exceptionally strong accumulation was recorded at 3.3 ka, when nearly 3 m of sand where accumulated in section GB032 (Fig. 4.1.4). This might indicate the age of the first appearance of the dune field in the area. However, in all sections a subsequent reduction of aeolian activity is observed. Loess was probably continuously deposited as indicated by OSL ages from palaeosols, but the accumulation was low. This can be attributed to a further increase in the precipitation under a fully developed Asian summer monsoon. Sediment availability supposedly was markedly decreased due to a denser vegetation cover. At one section, fluvial reworking was recorded (GB044) during this phase. A similar phase was also documented in reworked sediments in the Donggi Cona catchment from around 9 to at least 6 ka (Stauch et al., 2012) and in other areas of the north-eastern TP (Kaiser et al., 2007). In the eastern Gonghe Basin, (B. Liu et al., 2014) reconstructed warm and wet climate conditions based on several proxies from peat deposits from around 7.1 – 3.8 ka. However, there is some conflicting evidence regarding a stronger Asian summer monsoon in the early or in the mid- Holocene. Several proxies from Lake Qinghai indicate an early Holocene moisture maximum (e.g. Lister et al., 1991; An et al., 2012; Jin et al., 2015; Thomas et al., 2016). A similar timing is recorded in the speleothems from different caves in southern and eastern China, but only a small decline during the mid-Holocene (e.g. Dykoski et al., 2005; Wang et al., 2008). However, studies from Lake Donggi Cona show a high moisture values from 7 to 4.5 ka (Y. Wang et al., 2014). A similar signal was also recorded in the Kuhai Lake (Wischnewski et al., 2011), the Hongyuan peat (Hong et al., 2003) pollen records (Y. Zhao et al., 2009; Zhao et al., 2011) and a loess-palaeosol sequence in the Anyemaqen Mountains (Lehmkuhl et al., 2014). Additionally many aeolian records from the north-eastern TP show a strong decline in aeolian activity during the mid-Holocene (Stauch, 2015) and development of palaeosols was strong on the Chinese Loess Plateau (H. Wang et al., 2014). Similarly, in the Horqin dunefield in north-eastern China a stabilisation of dunes was observed from 8 to 3 ka (L. Yang et al., 2012)

Late Holocene

From around 2.5 ka a general increase in aeolian accumulation started on the eastern side of the basin, including the reactivation of aeolian fine sand. Additionally, in one section (GB065) more than 50 cm of loess accumulated since 1.1 ka. In the Gonghe Basin, a declining lake level was documented for Genggahai Lake since 2.1 ka (Song et al., 2012). Proxy records from different lakes show a shift to dryer climate conditions at around 2.5 ka (Shen et al., 2005; Mischke et al., 2010a; Thomas et al., 2016). However, in many other records a more pronounced shift is recorded earlier at around 4 ka (Hong et al., 2003; Colman et al., 2007; Liu et al., 2013b; Y. Wang et al., 2014; F. Chen et al., 2015b). A similar timing was also described for the peat deposits in the eastern Gonghe Basin (Liu et al., 2013a). Despite being recognized in many records as a distinct reduction in moisture, the shift at 4 ka may not have had an explicit influence on the vegetation cover in the Gonghe Basin, except for some sections in the central part of the basin. Additionally, several sections in the eastern and central Gonghe Basin include several palaeosols which indicate short-term phases of wetter climate or reduced aeolian activity. In general the late Holocene reactivation of aeolian sediments is related to the late Holocene decline of the Asian summer monsoon and less precipitation on the north-eastern TP. During the late Holocene, aeolian accumulation of sand resumed at around 3.6 ka in the Qaidam Basin, 3.2 ka at the Donggi Cona, at around 2 ka in Gonghe Basin and at 1.9 ka at the Qinghai Lake (Stauch, 2015). Similar to the onset of the sand accumulation during the late Pleistocene, a spatial trend from the south-west to the north-east can be observed, exhibiting a lagged remobilization of aeolian sands in the late Holocene. Additionally, a stronger human impact during the late Holocene might have influenced the aeolian activity. F. Chen et al. (2015a) stated that permanent settlements were established on the north-eastern TP since 3.6 ka. Intensive grazing might lead to the destruction of the vegetation cover which would also enhance aeolian entrainment (Schlütz and Lehmkuhl, 2009). In the Gonghe Basin several sections in the central basin are covered by fine or bimodal sand which probably indicates the activation of different, previously stable, source areas. However, at present it is not possible to quantify this influence on the aeolian system in the basin.

Despite this overall climate reconstruction based on several studies from the Gonghe Basin, a number of geomorphological problems remain. There are some indications of a larger loess cover which was deposited during the early Holocene and which probably also reached the lower part of the basin, e.g. the plains which are now covered by dune sands. However, there is only evidence from three sections available. Additionally, up to now the oldest age from the sand in the central part has an age of around 7 ka. There are several interpretations possible: a former aeolian cover of loess or sand was removed, either by deflation or by fluvial processes. Strong aeolian reworking of older aeolian deposits was observed in other catchments on the north-eastern TP (Donggi Cona). Nevertheless, aeolian reworking of, e.g. an older sand cover, seems unlikely in the light of the palaeoclimatic interpretation presented above. Fluvial reworking might be a suitable process, but only few direct geomorphological indications were documented up to now.

Several sections in the Gonghe Basin contain palaeosols (Fig. 4.1.6) with a different degree of pedogenesis. However, there are no distinct clusters of ages evident. Local topographic effects, like subsurface water flow or the configuration of the upstream catchment, supposedly have an important effect, but further research into this topic is required. Furthermore, the precise dating of the timing of the soil formation is essential.

The results from the Gonghe Basin indicate the need for a relatively dense sampling strategy in such complex environment. Analyses of the dataset from the individual studies yielded different results than the synthesis of previously published data and respective interpretations presented here. This highlights, that the individual studies, including the data from this one, have not been representative for the environmental evolution in the overall basin. This is especially important as erosion of aeolian deposits is frequently occurring on the TP and many archives might be discontinuous (Yu et al., 2013; Stauch et al., 2014).

Conclusion

Aeolian sediments in the Gonghe Basin show a distinct spatial separation with loess deposited on the southern and northern side of the basin while sandy sediments accumulated in the central and eastern part. However, at several locations in the central part loess is preserved below aeolian sands, indicating a previously larger extent of the loess cover. In 15 sections different sediments were dated with OSL. Additional 39 OSL ages for the last 16 ka from aeolian sediments were previously published by other studies. The OSL ages show two separated clusters, during the late Pleistocene and during the last 2.5 ka. Aeolian accumulation of loess started at around 15 to 14 ka, indicating a first post- glacial increase in moisture in the Gonghe Basin due to the strengthening of the Asian summer monsoon. More precipitation resulted in a formation of a vegetation covers which acted as a sediment trap for the fine grained sediments. Aeolian sands did not accumulate until around 12.5 ka, indicating the last movement of aeolian sands during the cold and dry Younger Dryas. The subsequent return to wetter climate conditions and the development of vegetation also in the lower part of the basin resulted in the permanent fixation of these sands. Highest moisture values were probably achieved during the mid-Holocene. From around 7.5 ka aeolian activity was strongly reduced in most parts of the basin and there are indications from one section of fluvial reworking of aeolian sediments. However, in the central part of the Gonghe Basin aeolian activity probably occurred episodically, e.g. at around 3.3 ka. From around 2.5 ka a resumption of aeolian activity is detected due to again dryer climate conditions in the area. The parallel accumulation of silty and sandy sediments at the same sites points to the parallel activation of different sediment source areas. This might be an indicator for an increased human influence in the area. The comparison of the aeolian activity in the Gonghe Basin for the last 16 ka with deposits from the neighbouring basins indicates an intrusion of the Asian summer monsoon on the north-eastern TP. Aeolian sediments were first fixated in the Qinghai Basin and latest in the Donggi Cona area while the aeolian sediments in the Gonghe Basin represent an intermediate position.

Acknowledgements

The research in the Gonghe Basin was supported by the RWTH Aachen University with a grant for excellent science and teaching for G. Stauch. The OSL dating was founded by a grant of the German Science Foundation (DFG) for the project ‘Landscape and Lake-System Response to Late Quaternary Monsoon Dynamics on the Tibetan Plateau - Northern Transect’ in the framework of the SPP 1372 (Tibetan Plateau: Formation – Climate – Ecosystems) and China NSF grant 41290252. Baoliang Lu greatly supported the field work in the Gonghe Basin. The manuscript improved due to fruitful scientific discussions with Veit Nottebaum. 44

4.2 Aeolian sediments on the north-eastern Tibetan Plateau

Georg Stauch, Janneke IJmker, Steffen Pötsch, Hui Zhao, Alexandra Hilgers, Bernhard Diekmann, Elisabeth Dietze, Kai Hartmann, Stephan Opitz, Bernd Wünnemann, Frank Lehmkuhl

Abstract

Aeolian sediments on the Tibetan Plateau are an important archive of palaeoclimatic information. This study presents a detailed analysis of sediments from the Donggi Cona catchment on the north- eastern part of the Tibetan Plateau. Long- and short-distance sediment transport leads to a complex pattern of aeolian sediment deposition that depends on climatic changes as well as on the availability of sediments. Based on the largest dataset of OSL datings (51) from a single catchment on the Tibetan Plateau so far, different periods of increased sediment transport have been reconstructed. Increased aeolian deposition in this high elevation environment started in the early Holocene with the accumulation of sands from around 10.5 to 7 ka. Loess sediments have been preserved from a period from 10.5 to 7.5 ka. Both archives are related to the strengthening of the Asian summer monsoons characterized by wetter and warmer climate. This change in climate supported the trapping of aeolian sediments. Under full monsoon conditions from around 9 ka onwards fluvial processes resulted in erosion of the aeolian archives and the formation of colluvial sediments until 6 ka. A dry and cooler climate resulted in the reactivation of dune sands from 3 ka to present, possibly in combination with stronger human influence. Aeolian sediments on the Tibetan Plateau therefore indicate two different climatic modes. During the early Holocene wetter conditions were favourable to retain aeolian sediments. The reactivation of sediment in the late Holocene due to small-scale disturbances in the vegetation cover, points to a cooler and drier climate.

Keywords: Quaternary; aeolian; sedimentology; dating; Tibetan Plateau; palaeoclimate

Introduction

Aeolian sediments are widely used to reconstruct the climatic evolution in central Asia during the Quaternary. One of the major climate archives in Asia are the massive sedimentary deposits in the Chinese Loess Plateau east of the Tibetan Plateau (e.g. Kukla, 1987; Kukla and An, 1989; An et al., 1991; Sun, 2002; Stevens et al., 2007; Y. Sun et al., 2010). These deposits formed during the last 7 Ma (Sun et al., 1998; An, 2000). The formation of this sedimentary body has been linked to the uplift of the Tibetan Plateau and the reorganization of large scale atmospheric processes (An et al., 2000). Considerably less attention has been paid to the aeolian sediments in high mountain areas such as the Tibetan Plateau as a valuable climate archive. However, during recent years aeolian sediments on the Tibetan Plateau have gained increasing interest and the number of dating results is rising fast (Lehmkuhl et al., 2000; Küster et al., 2006; Sun et al., 2007; Buylaert et al., 2008; Kaiser et al., 2009a; Lai et al., 2009; Lu et al., 2011b; Vriend et al., 2011). The main focus has been placed on the

45 occurrence of loess and loess-like sediments. Most dating results yield late Pleistocene and early Holocene ages for the onset of the loess deposition. Nevertheless, as the number of numerical dating is increasing older ages for loess sediments beyond the Last Glacial Maximum (LGM, ca 21 ka) are reported (e.g. Lai et al., 2009). Less research focused on the timing of sand mobilization.

The widespread occurrence of aeolian sediments on the Tibetan Plateau has been already noticed early in the last century (see e.g. Lehmkuhl and Haselein, 2000). In north-eastern Tibet (Hövermann, 1987) described a loess cover between 3500 and 3900 m on the east-facing slopes of the A’nyêmaqên Shan. According to Lehmkuhl (1997) sandy loess forms a distinct belt above the loess area up to 4500 m. In the source area of the Huang He they reach a thickness of 40–60 cm at the slopes and partly more than 100 cm. These loess-like sediments are dominated by the coarse silt fraction. The basins are dominated by sand and dune fields. However, in many situations aeolian sediments on the Plateau are non-continuous and represent only small snapshots in time (e.g. Yang et al., 2011). Erosional processes and reworking of sediments led to complex sediment deposits and require a detailed analysis of related geomorphological processes.

The formation of aeolian deposits is often related to a cold and dry climate (e.g. Yang et al., 2004). During glacial times larger areas serve as a dust source (e.g. Mahowald et al., 1999; Kohfeld and Harrison, 2001; Nilson and Lehmkuhl, 2001). However, local sediment availability is an important factor as well. Several source areas have been discussed for the aeolian sediments on the Tibetan Plateau. Beside glacial outwash material (Sun et al., 2007) rivers and dried lake basins (Lehmkuhl and Haselein, 2000; Lehmkuhl et al., 2000) have been proposed as major source of aeolian sediments. Loess sediments on the Tibetan Plateau often have a larger median size than the sediments on the Loess Plateau in central China, indicating a shorter transport distance (Vriend and Prins, 2005; Sun et al., 2007; Prins et al., 2009). This was confirmed by the geochemical analysis of loess samples from southern Tibet (Kaiser et al., 2009a). Lehmkuhl et al. (2000) proposed mixed deposition for the coarse loess sediments on the eastern Tibetan Plateau. Finer materials in the archive are related to long-distance transport, while the coarser material is derived from local sources. Beside effective source areas, aeolian sediments and especially loess depend on suitable conditions for deposition (Pye, 1995; Smalley, 1995). Vegetation often serves as a dust trap (Tsoar and Pye, 1987; Lehmkuhl, 1997; Lehmkuhl et al., 2000; Sun et al., 2007). An analysis of different factors influencing dust transport on a large scale has been presented by Nilson and Lehmkuhl (2001).

So far, the study of aeolian sediments focused predominantly on the loess fraction. In this paper we present a detailed analysis of aeolian and reworked sediments in the catchment of Lake Donggi Cona on the north-eastern Tibetan Plateau, incorporating sand and sandy loess deposits. Geomorphological mapping and sedimentological analysis lead to the identification of different types of aeolian cover in relation to their geomorphological position. Taking into account all types of aeolian deposits including reworked sediments supports the understanding of the spatial distribution and modification of the archives and increases the completeness of the reconstruction. Age control is based on luminescence dating. The results are compared to the overall climatic evolution in the area and give insights in the different environmental conditions during the late Pleistocene and the Holocene.

Regional setting

The study area is situated on the north-eastern Tibetan Plateau in the catchment of the Lake Donggi Cona at 35°18’N and 98°32’E (Fig. 4.2.1). The catchment comprises an area of 3175 km². The lake with a size of 229 km² is located at an elevation of 4090 m asl in the western part of the catchment. The highest elevation is reached in the SE part of the catchment with 5230 m asl. This part is the western extension of the A’nyêmaqên Shan. Higher elevations also occur in the north-western part of the catchment, where the mountains reach elevations of up to 5050 m asl. However, most parts of the study area are below 4300 m asl. The higher parts in the NW and SE of the catchment are shaped by Quaternary glaciations. Related landforms are cirques, U-shaped valleys and terminal moraines. Presently the catchment is not glaciated. The main inflow into the Donggi Cona is the Dungqu River originating in the southern part of the catchment. The outflow of the lake is formed by the Tuosuo River and drains via the Jialu River into the endorheic Qaidam Basin. The river valleys and the large central basin are dominated by massive fluvial deposits and river terraces. Five former shorelines up to 18 m above the present lake level indicate past lake level fluctuations (Mischke et al., 2010b).

The overall catchment is dominated by the active Kunlun fault, which runs in WNW to ESE direction. The Kunlun Fault Zone has a length of 1600 km and is one of the major tectonic boundaries in northern Tibet (Tapponnier and Molnar, 1977; Yin and Harrison, 2000; Fu and Awata, 2007; Van Der Woerd et al., 2002; Kirby et al., 2007). The Donggi Cona area belongs to the Tuosuo Lake segment (Dongxi Co according to Van Der Woerd et al., 2002) of this fault (Fu and Awata, 2007). The basin with the lake is a pull-apart structure while the south-east located A’nyêmaqên Shan is the related pressure ridge, which is bound by left-lateral faults (Fu and Awata, 2007). The Kunlun Fault is also clearly traceable in the basin morphology of Lake Donggi Cona (Dietze et al., 2010). The last large earthquake occurred in 1937 (M 7.5) and produced a horizontal offset of 4.1 m (Guo et al., 2007). Similar values of 4.4 ± 0.6 m horizontal offset have been proposed by Van Der Woerd et al. (2002) for a section at the Nianzha He about 40 km east of the eastern shore of lake Donggi Cona. The vertical offset at this site was less than 1 m. Mole tracks related to the 1937 earthquake (Van Der Woerd et al., 2002; Guo et al., 2007) are easily identified in the field and on satellite images east and west of Donggi Cona. Limestone and sandstones of Triassic and Permian age dominate the catchment, as well as Permian to Neogene clastic rocks. The basins are filled with unconsolidated Quaternary sediments (Wang et al., 1998; Wang and Yang, 2004).

Fig. 4.2.1: Overview of lake Donggi Cona. The catchment is outlined in black. Inset shows the location of the map at the north-eastern edge of the Tibetan Plateau.

The climate in the study area is characterized by the cold and dry winter monsoon and the moisture bearing East Asian and South Asian summer monsoon. As the study area is located at the northern limit of the monsoon-influenced area, precipitation values reach 311 mm yr-1 at the climate station of Madoi, approximately 50 km to the south. Thunderstorms are frequent in summer leading to high runoff values in a short time. The mean air temperature for January is -15.8 °C and 7.9 °C for July. Mean annual temperature is -3.0 °C (Mischke et al., 2010a). Features of active permafrost are widespread in the catchment. The vegetation cover in the area consists of alpine meadows, alpine steppes and salix scrubs at the slopes (Huang, 1987; Kürschner et al., 2005).

Aeolian processes are an important transport mechanism in the catchment of the Donggi Cona. At present these are especially active during winter times. This has also been described for other catchments on the north-eastern Tibetan Plateau (Liu et al., 2009; Mischke et al., 2010c). The main wind direction during this time is from the west (IJmker et al., 2012c) as a result of the north-western winter monsoon (Böhner, 2006; Sun et al., 2008; X. Yang et al., 2012) and the local valley configuration. During summer times torrential rainfall leads to sheet wash and strong erosional processes on inclined slopes with unconsolidated sediments. Gully formation occurs frequently.

Material and methods

Geomorphological mapping and sampling

Geomorphological mapping is based on field investigations during four field trips in the catchment of Lake Donggi Cona in the years 2006, 2008 and two times in 2009. Field trips took place during different seasons and therefore, present day geomorphological processes have been observed during winter, spring and summer. During the fieldwork selected profiles of different sediment types were described according to their geomorphological position and their surrounding environment. The profiles were sampled for sedimentological and geochemical analyses and dating. Besides detailed sampling at selected key sections, single samples were taken to analyse the spatial distribution of sediments. Single samples were taken if representative for the whole section. In order to capture sedimentological differences in the catchment over 1000 samples were collected and analysed. This allows the reconstruction of spatial as well as temporally changes in a relatively small area. However, due to the climatic conditions with strong torrential rainfall and the partly high relief, the potential incompleteness of the archives is hampering this approach. Besides analysing the sediment traps, possible sediment sources have been identified and sampled.

Sedimentological analysis

Sedimentological analysis included measurements of grain-size distribution and geochemical analysis. Dried sample aggregates were mechanically destroyed before sieving and the fraction smaller than 2 mm was analysed using a LS 13320 Laser Diffraction Particle Size Analyzer (Beckman

Coulter). Before measuring, samples were treated with HCl (30%) and H2O2 (10%) to remove carbonate and organic material. For dispersion, sodium pyrophosphate was used. Geochemical analysis was conducted using X-ray diffraction (Niton, Xlt 700) (IJmker et al., 2012b). Sediments were classified according to the shape of the grain-size distribution curve and their secondary parameters (e.g. median). Furthermore, the FAO classification scheme was used to identify sediment types (FAO, 2006).

Dating

Luminescence dating was performed at the laboratory at the University of Cologne and at the Cold and Arid regions Environmental Research Institute, CAS in Lanzhou. Luminescence dating can be directly employed to aeolian deposits using quartz or potassium rich feldspar (K-feldspar) grains (Murray and Olley, 2002; Sun et al., 2006). The single aliquot regenerative-dose (SAR) technique, originally developed for coarse-grained quartz (Murray and Wintle, 2000) and then for coarse- grained K-feldspar (Wallinga et al., 2000), has successfully increased the accuracy of the dating results.

Sample preparation

The samples were taken in light tight steel cylinders from the profiles investigated. Samples were taken throughout the catchment at different depth to avoid repeated sampling of the same layer. Only the central part of the sediment core was used for luminescence analysis. Sample preparation and luminescence measurements were carried out under subdued red light. After sieving all samples were treated with hydrochloric acid (10%) and hydrogen peroxide (10% Cologne, 20% Lanzhou) to 49 remove carbonate and organic matters as well as to disperse the clay aggregates. Coarse-grained quartz and coarse-grained feldspar (100-150 µm and 100-200 µm) were separated by using heavy liquids of sodium polytungstate with densities of 2.68 g/cm3 and 2.58 g/cm3, respectively. Quartz grains were treated with 40% HF (45 min Cologne, 60 min Lanzhou) to remove the outer layer irradiated by alpha particles and any remaining feldspars grains. The K-feldspar grains were treated with 10% HF for 40 min to remove the outer layer which irradiated by alpha particles (Lanzhou). Then, all samples were treated with HCl (60 min Cologne, 10 min Lanzhou) to remove fluorides created during the HF etching. After re-sieving the material, the purified quartz and feldspar grains were mounted onto stainless-steel discs which measure ~9.8 mm in diameter each. Only the central 1 mm of each disc was covered with a few tens of grains and several of these sub-samples (‘1mm- aliquots‘) were finally measured per sample to obtain equivalent dose estimates.

Apparatus and measurements protocols

OSL measurements for equivalent dose (De) determination were carried out using automated Risø luminescence readers (types TL/OSL-DA-12, -15, or -20) equipped with 90Sr/90Y-β-sources for irradiation, blue light and infrared light-emitting diodes and EMI 9235 photomultiplier tubes for measurements of the luminescence signal (Bøtter-Jensen et al., 1999). The OSL signal was detected through a Hoya U-340 optical filter and the feldspar emission was optically filtered with a 410 nm interference filter. For quartz grains, in order to eliminate the influence from feldspar contamination, the post-IR single aliquot regenerative protocol was used to obtain De values from quartz extracts (Zhang and Zhou, 2007). The De values of K-feldspar extracts were determined using a SAR protocol of Wallinga et al. (2000).

Quartz coarse grains

In order to check whether quartz is suitable for determination of the palaeodose, LM-OSL (Linear Modulated – OSL) and IRSL measurements of quartz were carried out in Cologne. The presence of an IRSL-signal testifies for a feldspar contamination in most of the quartz fraction preventing accurate De determinations using quartz. Similar problems have been reported in other luminescence dating studies in Central Asia (Hülle et al., 2010).

The LM-OSL measurements showed that the fast component is only weakly developed in the quartz fraction (Fig. 4.2.2).The fast component is commonly easy and fast bleachable by optical stimulation (Bulur, 1996; Bulur et al., 2000; Singarayer and Bailey, 2003; Preusser et al., 2009). These signal characteristics are crucial for a successful palaeodose determination on quartz grains using the SAR procedure (Murray and Wintle, 2000, 2003; Wintle and Murray, 2006). In addition, for successful application of the SAR protocol the fast component of quartz should be about ten times higher than the medium and slow components (see Fig. 9 in Preusser et al., 2008).To verify the results sample DC09-07 was measured with quartz (standard SAR-protocol) and K-feldspar (Cologne). For K-feldspar an age of 8.5 ± 0.8 ka and for quartz an age of 6.0 ± 0.7 ka was calculated. This underestimation of at least 29% of quartz relative to the K-feldspars of the same sample underlines the difficulties in application of the SAR quartz protocol to material investigated in Cologne. Rather than quartz underestimating feldspars one could expect K-rich feldspars to yield younger ages due to fading of the luminescence signal over the time span of burial. This unexpected age underestimation of quartz most likely results from poor OSL characteristics of the quartz fraction in the study area. A reason for

50 the low luminescence intensities of the quartz fraction (Fig. 4.2.3) could be a low number of sedimentation cycles which the samples have run through (Pietsch et al., 2008; Preusser et al., 2009).

Based on these observations De determination were carried out on coarse-grained K-rich feldspars in Cologne, whereas the Lanzhou laboratory used the post-IR SAR protocol on coarse grains of quartz in order to circumvent any feldspar contamination.

Fig. 4.2.2: Deconvoluted LM-curve for loess (DC09-07) (left) and dune sand (DC09-30) (right). In both samples the fast component is only weakly developed. But, for successful application of the SAR protocol for OSL dating of sedimentary quartz the fast component should be about ten times higher than the medium and slow components, which is clearly not observed for the samples under study here.

Fig.4.2.3: Quartz shine–down curves of a quartz from NE-Germany with good OSL characteristics and a good signal-to- noise ratio (sample N6, see Hilgers (2007)) and a quartz sample from the Donggi Cona (DC09-07). Sample N6 with a depositional age of only 0.32 ka shows a fast signal depletion from the initial signal of cA 400 photon counts down to background level within the first few seconds of optical stimulation. The considerably older sample DC09-07 (~8.5 ka IRSL feldspar age or ~6.0 ka OSL quartz age) has a much lesser initial signal intensity which is almost difficult to differentiate from the background.

Potassium feldspar coarse grains

According to preheat tests with temperatures ranging in 30°C steps between 200 and 300°C, a preheat temperature of 210°C (De < 24Gy), 240°C (De = 24–47 Gy) and 270°C (De > 47 Gy) was chosen for the regeneration and test dose measurements. Stimulation of the IRSL signal was carried out with infrared LEDs for 300 s at an elevated temperature of 30°C. A minimum of 12 aliquots was measured for De determination. According to Murray and Wintle (2000) a 10% error range for the recycling ratio was accepted. The suggested threshold value of 5% for the recuperation (Murray and Wintle, 2000) was raised to 10% for younger samples (<3000 a), showing in general a poor recuperation.

Fading

Fading describes a loss of luminescence signal with time and results in a systematic age underestimation. This phenomenon was often observed for feldspars, but is uncommon for quartz. A test for the presence of fading has been proposed by Auclair et al. (2003) and was carried out for two samples (DC09-23-1, DC09-36-2). This test shows that the samples are affected by fading. The correction for fading followed the procedure after Huntley and Lamothe (2001), developed for samples, whose palaeodose is still in the linear section of the growth curve. De-values of the tested samples fulfil this criterion and therefore the fading correction is applicable. The mean g-value for the tested samples ranges between 2.1 ± 2.1% (DC09-23-1) and 2.3 ± 0.6% (DC09-36-2). This corresponds to an age underestimation of 17.4%–21.1%. As it cannot be ruled out that all dated feldspar samples are affected by fading to a similar amount, at this stage of the study all age estimates ought to be considered as minimum ages. However, for a concluding evaluation whether the fading corrected ages are correct, independent age estimations are necessary. Notwithstanding, at presence, they yield a good approach to the true depositional age.

Determination of the dose rate

The concentration of uranium, thorium and potassium were determined using gamma ray spectrometry in Cologne and neutron activation analysis (NAA) in Lanzhou (Table 4.2.2). Cosmic dose rates were calculated according to the actual burial depth of the samples, altitude and geographical position of the profile (Prescott and Hutton, 1988, 1994) (Table 4.2.1). To cope with fluctuations in the water content and their effect on the dose rate, the water content range for the dose rate calculation was set from 2 to 7.5 weight-% (weight water/dry sediment) for dune sand and 7.5 – 20 weight-% for loess or loess-like sediments. Furthermore an average internal K-content of potassium rich feldspar of 12.5 ± 0.5% (Huntley and Baril, 1997) and a decreased alpha efficiency of 0.07 ± 0.02 (Preusser et al., 2005) in respect to the effect of beta and gamma radiation are also used for dose rate calculations.

Table 4.2.1: Site information of the OSL dated sections.

Altitude Sediment Section Northing Easting [m a.s.l.] classification No. OSL samples

06-P16 35.409344 98.64199 4236 Loess 1

06-P21 35.23941 98.50851 4103 Dune sand 2

06-P23 35.41678 98.427135 4308 Loess 2

06-P24 35.27088 98.5035 4101 Loess / Sandy loess 2

P006 35.25933 98.74848 4099 Dune sand 1

P012 35.16555 98.81394 4179 Sandy loess 1

P019 35.20032 98.8708 4164 Dune sand 1

P021 35.24477 98.7762 4108 Reworked loess 1

P026 35.22505 98.82145 4118 Reworked loess 1

P027 35.22377 98.81076 4120 Dune sand 1

P028 35.2269 98.79793 4119 Dune sand 1

P030 35.2535 98.83316 4105 Dune sand 1

P063 35.35637 98.34821 4091 Reworked loess 4

P068 35.24113 98.80424 4109 Dune sand 1

P101 35.15997 98.91447 4195 Dune sand 1

P109 35.240366 98.85428 4123 Dune sand 1

P115 35.215789 98.58522 4119 Dune sand 2

P140 35.24704 98.47935 4164 Dune sand 8

P142 35.24726 98.48147 4157 Dune sand 1

P165 35.42086 98.19357 4091 Reworked loess 3

P173 35.23974 98.50448 4119 Dune sand 1

P204 35.22974 98.51993 4148 Dune sand 3

P209 35.24707 98.48171 4157 Dune sand 2

P218 35.32843 98.38128 4250 Loess 1

P221 35.22156 98.89403 4142 Dune sand 6

P237 35.20876 98.93463 4209 Dune sand 1

P238 35.2093 98.93837 4299 Sandy loess 2

P270 35.1929 98.89576 4285 Sandy loess 2

Table 4.2.2: OSL dating results (Lab.-No starting with T or C-L: Cologne, Lab.-No starting with L: Lanzhou).

Section Sample Lab.-No Depth U [ppm] Th [ppm] K [%] Dose rate Water content Equivalent Age [ka] [cm] [Gy/ka] [%] dose [Gy] 06-P16 DC06-01 C-L1623 45 2.62±0.12 9.88±0.82 1.62±0.04 2.73±0.12 4-20 34.7±2.1 12.7±1 T06-4 06-P21 DC06-02 C-L1624 45 2.23±0.17 8.13±0.42 1.27±0.03 2.38±0.09 4-20 19.4±2.1 8.14±0.93 T06-5 DC06-03 L2006-25 300 1.47±0.11 5.26±0.17 1.30±0.12 1.96±0.13 3.3 19.6±0.6 9.97±0.7 06-P23 DC06-04 C-L1853 40 2.57±0.13 10.48±0.55 1.62±0.04 2.77±0.12 4-20 22.9±1.3 8.28±0.59 T06-7 DC06-05 C-L1854 80 2.57±0.13 9.40±0.53 1.50±0.03 2.55±0.11 4-20 25.2±1.5 9.88±0.74 T06-6 06-P24 DC06-06 L2006-26 40 2.24±0.10 9.97±0.25 1.48±0.11 - 8.1 - no result DC06-07 L2006-27 157 3.53±0.12 12.3±0.31 1.92±0.11 3.07±0.24 16.7 13.9±0.9 4.45±0.38 P006 DC08-01 L2008-41 120 1.77±0.04 6.71±0.16 1.14±0.03 2.07±0.05 1.1 2.82±1.32 1.34±0.63 P012 DC08-03 C-L2443 52 ------no result P019 DC08-04 C-L2444 230 1.51±0.08 6.20±0.34 1.09±0.03 2.68±0.17 2-8 0.82±0.04 0.31±0.02 P021 DC08-05 L2008-43 36 2.86±0.08 10.3±0.27 1.86±0.04 3.26±0.08 2-8 9.32±0.86 2.86±0.28 P026 DC08-07 L2008-44 77 2.18±0.07 9.45±0.27 1.50±0.03 2.74±0.07 2-8 19.3±0.9 7.06±0.37 P027 DC08-08 C-L2446 87 1.79±0.09 7.43±0.41 1.24±0.04 3.04±0.19 2-8 1.73±0.09 0.57±0.05 P028 DC08-09 C-L2447 180 1.39±0.07 5.87±0.32 1.07±0.03 2.63±0.16 2-8 1.03±0.06 0.39±0.03 P030 DC08-10 C-L2448 116 1.73±0.09 6.93±0.38 1.25±0.04 2.98±0.19 2-8 3.25±0.18 1.09±0.09 P063 DC09-07 C-L2704 243 2.98±0.15 10.92±0.68 1.69±0.07 3.69±0.28 7.5-17.5 31.1±1.7 8.49±0.78 DC09-09 C-L2706 157 ------no result DC09-10 C-L2707 130 2.64±0.14 10.57±0.67 1.57±0.06 3.53±0.27 7.5-17.5 28.9±0.9 8.17±0.68 DC09-12 C-L2709 57 3.02±0.13 10.86±0.63 1.75±0.05 3.78±0.28 7.5-17.5 26.4±0.7 6.98±0.54 P068 DC09-13 C-L2710 57 1.94±0.10 8.16±0.54 1.20±0.05 3.11±0.21 2.5-7.5 1.03±0.05 0.33±0.03 P101 DC09-14 C-L2711 250 1.90±0.10 8.03±0.51 1.21±0.05 3.02±0.20 2.5-7.5 3.63±0.21 1.2±0.11 P109 DC09-19 C-L2716 203 1.66±0.09 6.82±0.44 1.04±0.04 2.74±0.18 2.5-7.5 1.38±0.07 0.5±0.04 P115 DC09-21 C-L2718 397 1.13±0.06 4.73±0.32 0.92±0.04 2.26±0.14 2.5-7.5 3.33±0.19 1.47±0.13 DC09-22 C-L2719 548 1.08±0.06 4.66±0.31 0.89±0.03 2.10±0.13 2.5-7.5 1.64±0.14 1.64±0.14 P140 DC09-23-1 C-L2720 572 1.32±0.07 5.38±0.35 0.98±0.04 2.36±0.15 2.5-7.5 6.87±0.13 2.91±0.19 DC09-23-2 C-L2721 522 1.37±0.07 5.81±0.33 1.11±0.04 2.47±0.15 2.5-7.5 6.41±0.37 2.6±0.22 DC09-23-3 C-L2722 472 1.24±0.07 5.33±0.34 0.91±0.04 2.23±0.14 2.5-7.5 2.73±0.15 1.22±0.1

DC09-23-5 C-L2724 384 1.29±0.07 5.47±0.35 0.97±0.04 2.41±0.15 2.5-7.5 2.92±0.16 1.21±0.1 DC09-23-8 C-L2727 240 1.22±0.07 5.29±0.35 0.96±0.04 2.34±0.16 2.5-7.5 1.84±0.09 0.79±0.06 DC09-23-9 C-L2728 193 1.24±0.06 5.40±0.31 0.98±0.04 2.39±0.16 2.5-7.5 1.36±0.09 0.57±0.05 DC09-23-10 C-L2729 140 1.18±0.06 5.34±0.31 1.00±0.04 2.41±0.16 2.5-7.5 0.83±0.04 0.34±0.03 DC09-23-11 C-L2730 82 1.19±0.07 5.24±0.34 0.94±0.04 2.46±0.016 2.5-7.5 0.75±0.04 0.31±0.03 P142 DC09-30 C-L2756 214 1.47±0.08 6.22±0.41 1.04±0.04 2.64±0.17 2.5-7.5 6.89±0.36 2.61±0.22 P165 DC09-24-1 C-L2731 227 2.80±0.14 10.47±0.66 1.59±0.06 3.54±0.27 7.5-17.5 41.5±2.3 11.7±1.1 DC09-24-3 C-L2733 184 3.07±0.16 10.48±0.66 1.81±0.07 3.81±0.29 7.5-17.5 24.4±1.5 6.40±0.62 DC09-24-7 C-L2737 52 2.70±0.14 10.04±0.65 1.62±0.06 3.59±0.28 7.5-17.5 16.8±1.3 4.68±0.51 P173 DC09-25-1 C-L2738 240 1.32±0.07 5.43±0.36 0.99±0.04 2.40±0.16 2.5-7.5 1.64±0.08 0.68±0.06 P204 DC09-29-1 C-L2751 276 1.28±0.07 5.37±0.36 0.94±0.04 2.30±0.17 3-11 4.14±0.25 1.8±0.17 DC09-29-3 C-L2753 187 1.24±0.06 5.40±0.31 0.98±0.04 2.36±0.17 3-11 1.14±0.10 1.14±0.1 DC09-29-5 C-L2755 83 1.12±0.06 4.75±0.31 0.90±0.04 2.25±0.16 3-11 1.08±0.06 0.48±0.04 P209 DC09-31-1 C-L2757 143 1.34±0.08 5.75±0.39 0.99±0.04 2.51±0.17 4-11 29.3±0.4 11.7±0.8 DC09-31-2 C-L2758 89 1.53±0.08 6.23±0.41 1.11±0.04 2.73±0.18 4-11 30.6±0.3 14.0±1.1 P218 DC09-35 C-L2766 78 2.71±0.14 9.66±0.62 1.45±0.06 3.49±0.27 5-15 31.5±1.9 9.02±0.88 P221 DC09-36-1 C-L2767 535 1.43±0.08 5.76±0.37 1.00±0.04 2.45±0.15 2.5-7.5 34.1±2.3 13.9±1.3 DC09-36-2 C-L2768 485 1.45±0.08 5.78±0.38 0.99±0.04 2.39±0.16 2.5-7.5 20.7±0.6 8.67±0.62 DC09-36-3 C-L2769 448 1.71±0.09 6.84±0.44 1.11±0.04 2.73±0.17 2.5-7.5 11.2±0.6 4.12±0.34 DC09-36-4 C-L2770 387 2.13±0.11 7.98±0.51 1.19±0.05 2.94±0.20 2.5-7.5 5.26±0.27 1.79±0.15 DC09-36-5 C-L2771 333 1.38±0.08 5.80±0.39 0.96±0.04 2.39±0.16 2.5-7.5 2.88±0.19 1.21±0.11 DC09-36-7 C-L2773 121 1.49±0.08 6.31±0.41 1.01±0.04 2.65±0.18 2.5-7.5 0.83±0.05 0.31±0.03 P237 DC09-37-1 C-L2774 257 1.31±0.07 5.70±0.33 0.98±0.04 2.42±0.16 2.5-7.5 0.94±0.05 0.39±0.03 P238 DC09-38-1 C-L2777 187 1.80±0.10 7.50±0.48 1.16±0.05 2.76±0.22 5-15 1.52±0.15 1.52±0.15 DC09-38-3 C-L2779 135 1.73±0.10 7.32±0.48 1.19±0.05 2.78±0.22 5-15 1.78±0.09 0.64±0.06 P270 DC09-40-1 C-L2781 91 2.30±0.12 8.82±0.58 1.30±0.05 3.02±0.24 10-20 5.96±0.35 1.97±0.19 DC09-40-3 C-L2783 53 2.32±0.12 9.67±0.57 1.50±0.06 3.27±0.25 10-20 1.11±0.11 1.11±0.11

Results

Aeolian sediments are widespread in the catchment of the Donggi Cona and have been found close to the present lake shore and up to more than 4800 m asl. According to their sediment characteristics and the geomorphological situation of the archive, four different types of sediments can be distinguished. Three of them are of primary aeolian origin (loess, sandy loess and dune sand, Fig. 4.2.4) and one type consists of remobilized loess and sandy loess (reworked loess). However, only reworked sediments with a relatively short transportation distance of approximately up to 1 km are considered in this study to get information on the erosion of the loess cover which are not influenced by other factors (e.g. changes in river courses). Therefore, longer transport e.g. by rivers is not discussed here. Sediments were classified according to their related archive in the field and grain- size composition. According to Pye (1995), loess is a terrestrial, clastic sediment formed by the accumulation of aeolian dust. It is composed mainly of silt-sized particles. A further differentiation has been made in this study. Sediment samples having a silt and clay content of 80% or more have been termed loess, while samples with a silt and clay content of 50–80% were termed sandy loess. The clay content in both datasets is around 10% (Fig. 4.2.5) and is probably related to the intensive periglacial weathering in this extreme environment. The class of the sandy loess can be further divided into two subgroups based on the grain-size distribution curve. Bimodal sandy loess has a major peak in the coarse silt (20-63 μm) and a minor peak in the medium to coarse sand fraction (200-1250 μm) while unimodal sandy loess has only one peak in the coarse silt fraction. Polymodal grain-size distributions are indicative of more than one transportation or sedimentary process (Weltje and Prins, 2003, 2007; Sun et al., 2004). Reworked loess was classified solely according to the underlying geomorphological process. Grain-size composition in these archives shows strong variations with partly high sand content. The sand content in the dunes is mostly well above 75%. Loess samples have a median of 10–30 μm while sandy loess sediments have a median between 30 and 100 μm. In contrast the median of aeolian sands varies between 100 and 300 μm.

Fig. 4.2.4: Mean grain-size distributions of loess, sandy loess and dune sand, as representative for the aeolian deposits in the Donggi Cona catchment.

Fig. 4.2.5: Ternary diagram of the textural classes of the aeolian sediments (classification according FAO, 2006). Black circles: loess, white boxes: sandy loess, white circles: dunes.

Spatial distribution

Loess

Loess deposits are predominantly located in the western part of the Donggi Cona basin (Fig. 4.2.6). The loess cover is thin with a maximum thickness of 2 m at protected sites and has a massive structure. However, in most situations it is less than 1 m, which is typical for many areas in Tibet (Lehmkuhl, 1997; Lehmkuhl et al., 2000; Kaiser et al., 2007, 2009a). The loess cover in the study area is only partly preserved on the top of small hills, the gentle slopes of the large glacial moraines and on the river terraces in elevations between 4130 and 4700 m asl. Loess is absent in the eastern part of the study area except from a situation in the NE part of the catchment. An overall number of 20 sections mainly consist of loess. In three sections the loess was underlain by coarser sediments. These bottom samples (6) were removed, resulting in a dataset of 46 loess samples.

Fig. 4.2.6: Location of aeolian sediment sections and main dune fields in the Donggi Cona catchment.

Sandy loess

Sandy loess, with a silt and clay content of 50–80%, is distributed over the entire catchment. However, most deposits are located in the higher parts of the mountains north of the lake and on the top of the hills in the eastern part of the large alluvial fan (Fig. 4.2.7). Additionally, five sections with sandy loess have been sampled on the hill tops on the southern side of the lake within a dune field and on higher lake terraces. At one of these higher lake terraces at approximately 10 m above the present lake level, the sandy loess was underlain by loess sediments (06-P24, Table 4.2.1, Fig. 4.2.8). This loess was not present at the base of a nearby profile on a lower terrace. Sections with sandy loess have a maximum thickness of 1.2 m and have been found in elevations between 4092 m close to the present lake level and 4812 m asl. Sediments in the southern and eastern part of the catchment have a unimodal grain-size distribution. Bimodal sandy loess was found mainly north of the lake. Sections in the eastern part of the study area in the surrounding of the large alluvial fan contain up to three weak fossil soils. 26 sections with 68 samples belong to the group of sandy loess.

Reworked loess

Reworked sediments can be found in the entire catchment, covering most of the slopes and forming massive colluvial deposits at the footslopes. Colluvial sections, formed by reworked loess of about 2.5 m depth have been found on the terraces at the outlet of the Donggi Cona (P63) and at the westernmost section on a terrace of the Tuosuo River (Fig. 4.2.7). Reworked sediments of 10–20 cm are covering the fluvial gravel of the highest terraces on the large alluvial fan and are covered again by aeolian sands. The class of the reworked loess contains 92 samples from 20 sections.

Fig. 4.2.7: Examples of different sediment archives in the Donggi Cona area (upper left: erosional cut in the southern dune field, upper right: sandy loess in the eastern part of the catchment, lower right: present day sand deposition, lower left: reworked aeolian sediments on the terrace of the Tuosuo River).

Aeolian sands

Sands are the most widespread aeolian sediments in the study area. 49 sections with 396 samples have been studied. They are especially distributed in a large dune field, which stretches from the south-western to the eastern part of catchment, having a length of more than 60 km. However, the dune field is not continuous and consists of several isolated patches. According to field observations and the remote sensing datasets, aeolian sands cover an area of at least 115 km². The thickest sand deposits can be found at the south-west orientated slopes in the southern and eastern part of the catchment, indicating the major direction of sand transport. Based on the geochemical composition of the sand the dune field can be divided into two parts. The eastern part gains its sediments primarily from the major inflow into the lake and former lake sediments, while sources in the southern part are located in the adjacent basins (IJmker et al., 2012b, 2012c) According to the distribution of the sand patches and the prevailing wind direction, the eastern dune field gains 59 additional sand input from the southern dune field. Most of the dunes are partly covered by vegetation.

In the eastern dune field 29 sections have been sampled resulting in an overall number of 175 samples. Dunes are widespread on the large alluvial fan but heights of over 2 m are rare. Most of the small dunes are presently eroded by deflation and surface runoff related to the torrential summer rainfall. However, present sand accumulation takes place at protected sites like the leeward side of the dunes. The thickest deposits of aeolian sand of over 6 m (e.g. P221) can be found on the edges of the fan and on the slopes of the neighbouring hills. This is also the area of the strongest present sand accumulation. The sand deposits in the eastern part of the dune field show first indications of sorting. Samples on the central and northern part of the fan have a smaller sand content (mean sand content 80.44% (n: 138)) than samples from location at the edges of the fan and especially at the eastern end of the dune field (mean sand content 94.36% (n: 37)). This differentiation can also be found when looking at the sand content alone. While the dune sediments on the fan have a mean fine sand content (fS: 63-<200μm; FAO, 2006) of 66% (of the sand fraction) and a mean medium sand content (mS: 200-<630μm) of 32%, the mean content of fS and mS on the edges is nearly the same (49.62% and 49.96%). The sources of the finer material are former lake sediments on the eastern side of the alluvial fan as revealed from their geochemical signature (IJmker et al., 2012b).

On the southern side of the lake an altitudinal differentiation of the aeolian sands has been observed. Dunes with a sand content of more than 90% have been found especially at the valley bottoms or near to the valley bottoms. Finer-grained sediments occur on the different slopes and on the top of small hills, which have a sand content between 80 and over 90%. Two sections on these hills belong to the sediment class of sandy loess. However, profiles with a length of around 3 m and often contain some parts with finer material indicating varying sand transport (e.g. P173, 06-P21, P115 and P140 (Fig. 4.2.8)).

Dating

At 28 sections (Table 4.2.1) a total number of 51 OSL ages (Table 4.2.2) were determined in the frame of this study resulting in the largest dataset on aeolian sediments on the Tibetan Plateau so far. Five of the results are obtained from loess deposits, four are related to sandy loess, eight are from reworked loess sediments and 34 are from dune sands. Three additional samples gave no results. The OSL signal from sample DC08-03 and DC09-09 was too weak to obtain reproducible results. The sample DC06-06 was obviously taken from modern deposits, as the sample did not yield a measurable luminescence signal.

The loess ages range from 12.7 ± 1 ka to 4.54 ± 0.38 ka. In contrast, all dates related to sandy loess deposits have a late Holocene age, spanning from 1.97 ± 0.19 ka to 0.64 ± 0.06 ka. Reworked loess sediments show an age range from 11.7 ± 1.1 ka to 2.86 ± 0.28 ka. The largest time span of deposition in our dataset is documented for dune sands, which range from 13.9 ± 1.3 ka to 0.31 ± 0.03 ka. However, the younger limit is due to the sampling strategy and not related to sediment transport processes. Sediments closer than 40 cm below the present surface have not been sampled, except for one sample (DC08-05, Table 4.2.2).

Fig. 4.2.8: OSL sampling sites. Numbers refer to the samples described in table 4.2.2.

Despite these large ranges in dating results, several periods of intensified deposition of the different sediments are apparent. In this study a simple definition of periods with intensified deposition is used. Enhanced deposition is eminent if three or more ages of the same archive show overlapping error bars. This approach easily excludes single ages which might not be representative for the overall sedimentary environment.

Loess sediments show a phase of intensified deposition (Fig. 4.2.9) during the early Holocene, ranging from 10.6 ka to 7.7 ka (including the error) [DC06-04, DC06-07, DC09-35]. Sandy loess sediments from the eastern part of the catchment do not show any clustering, while the reworked loess sediments have been mainly deposited from 9.3 ka to 5.8 ka [DC08-07, DC09-07, DC09-10, DC09-12, DC09-24-3]. A more complex picture emerges from the larger dataset of dune sediments with three main age clusters. The oldest phase of intensified aeolian deposition ranged from the early Holocene from 10.7 ka to 7.2 ka [DC06-02, DC06-03, DC09-36-2], the second phase lasted from 3.1 ka to 2.4 ka [DC09-23-01, DC09-23-02, DC09-30] and the third phase started at 2 ka and ended with the youngest date in our dataset at 0.3 ka [DC08-01, DC08-04, DC08-08, DC08-09, DC08-10, DC09-13, DC09-14, DC09-19, DC09-21, DC09-22, DC09-23-3, DC09-23-5, DC09-23-8, DC09-23-9, DC09-23-10, DC09-23-11, DC09-25-1, DC09-29-1, DC09-29-3, DC09-29-5, DC09-36-4, DC09-36-5, DC09-36-7, DC09-37-1]. However, as mentioned above, the end of this phase is, related to our sampling strategy and not due to sedimentation processes. One sample (DC08-01) with an age of 1.34 ± 0.63 ka has an unusually high error. If this sample is excluded from the analysis, the latest phase can be divided into four subgroups (2–1.3 ka; 1.3–1 ka; 0.9–0.4 ka and 0.4–0.3 ka). However, the difference between these subgroups is very small (up to 20 a). Three ages from dune deposits yielded late Pleistocene ages, but do not have overlapping error bars.

Fig. 4.2.9: Sediment ages and phases of sediment deposition and mobilization for the different types of aeolian sediments.

At two sections, a detailed chronology was established. At section P221 (Fig. 4.2.8) on the northern side of the alluvial fan six OSL dates were obtained, while in the southern dune field ten dates are available. The latter section consists of two sections (P140 and P209) which are only 2 m apart and were correlated based on the sedimentary sequence. Both sections show highest sedimentation rates during the late Holocene, especially during the last 2 ka (Fig. 4.2.10). In section P140 on the southern side of the lake, nearly 5 m of sand were deposited during this period, and nearly 4 m on the northern side.

Fig. 4.2.10: Deposition of dune sand at two selected sites south (P140/P209) and north-east of the Donggi Cona (P221). Section P209 is the downward extension of section P140.

Discussion

Sediment source

Aeolian sediments in the catchment of the Donggi Cona show a distinct spatial distribution, which can be related to different source areas and trapping environments. Loess sediments are preserved in the western part of the catchment. The median grain-size is between 10 and 30 μm (21 μm on average) which is remarkable low in comparison to other archives on the Tibetan Plateau. Generally, grain-size analyses from southern Tibet show coarser loess sediments with a higher proportion of the sand class (Sun et al., 2007; Kaiser et al., 2009a), while results from the Donggi Cona are more comparable to the loess on the Chinese loess plateau (Yang and Ding, 2004; Y. Sun et al., 2010). According to Pye (1987) dust particles finer than 20 μm are indicative for long-distance transport (see also Tsoar and Pye (1987). Despite being somewhat coarser, the fine loess might be indicative of long-distance transport probably due to a more southern transport path of loess sediments from the central Asia deserts during Lateglacial and early Holocene times. A possible source area is the Qaidam Basin, located just 100 km to the north-west. The depositional setting in the leeward side of the high mountains has been a suitable trapping environment for the early Holocene loess. A similar situation can be assumed for the loess deposits in the A’nyêmaqên Shan (Hövermann, 1987).

Sandy loess and sand have a more uniform spatial distribution, which points to a local sediment source. The bimodal sandy loess was deposited especially in the higher part of the mountains. A mixture of loess and local sand is most likely. While the source of the loess component might be also located in the Qaidam Basin, the sand component originates from local rivers and solifluction debris. This assumption is based on the geomorphological configuration of the valleys. However, up to now no ages on the bimodal sandy loess are available which would support the hypothesis of a synchronous deposition of the loess and the sandy loess and further research is necessary, e.g. analysis of the geochemical composition of the sediments. The unimodal sandy loess in the eastern part of the study area can be attributed to a different sediment source. Geochemical analyses of the 63 sediments show a high proportion of carbonate indicating a source in the dried lake sediments on the eastern shore of the lake (IJmker et al., 2012b). This is also supported by the young OSL ages of less than 2 ka. The source areas for the sand are more widespread. Two primary sources for the eastern dune field are supposed; the wide river beds of the inflow from the south, which leads to the depositions of extensive sand deposits in the river bed. These deposits are blown out and deposited on the large alluvial fan. A secondary source is the reactivation of older sand deposits, either nearshore deposits of a former higher lake stand or the reactivation of older dune deposits. For the southern dune field possible sand sources are the alluvial fans terminating in the neighbouring basin to the east. In contrast to previous work on the Tibetan Plateau (e.g. Lehmkuhl, 1997; Lehmkuhl et al., 2000) distinct belts of aeolian sediments have not been found in the catchment of the Donggi Cona. The distribution of sediments is more related to the location of the source areas and specific topographical conditions.

Chronology of aeolian and reworked sediments

Results from the Donggi Cona catchment show an intensified deposition of sand and loess during the early Holocene. These results are supported by a large number of Lateglacial to early Holocene ages which have been obtained from the eastern Tibetan Plateau. However, most datings are related to loess deposits. Owen et al. (2003) yielded an age of 12.3 ± 1.1 ka for a massive silt in the nearby A’nyêmaqên Shan. In the Qilian Shan at the north-eastern margin of the Tibetan Plateau loess accumulation started at the end of the Pleistocene and the early Holocene and continued until present (Stokes et al., 2003; Küster et al., 2006). In the Kunlun Shan, a loess sample was OSL dated to 8.6 ± 0.7 ka (Owen et al., 2006). Lehmkuhl et al. (2000) present 3 TL dates from eastern Tibet south of the study area which indicate a predominantly deposition of silty sediments in the early Holocene. A comprehensive study of aeolian sediments was undertaken at the Qinghai Lake about 200 km north of Donggi Cona (Lu et al., 2011b). 24 OSL ages were obtained for loess, sand and paleosols. The six loess ages range from 13.31 ± 1.26 to 3.17 ± 0.23 ka, with three OSL ages younger than 5 ka, showing mid-Holocene loess deposition. Additional ages on loess sediments at the Qinghai Lake are provided by Madsen et al. (2008) and Porter et al. (2001). Madsen et al. (2008) dated a loess deposit to 16.1 ± 1.25 ka. Porter et al. (2001) presented four OSL ages on loess ranging from 12.7 ± 1 to 5.1 ± 0.4 ka and one age of 16.9 ± 1.4 ka for sandy silt underlying the loess. Numerous ages are available for southern Tibet, especially in the area of Lhasa. Recent results from the Kyichu River where loess is the most widespread aeolian sediment (Kaiser et al., 2009a, 2010; Lai et al., 2009), indicate a massive loess deposition before the Last Glacial Maximum (ca 21 ka). In contrast, Sun et al. (2007) stated that most of the loess in southern Tibet was deposited after the last deglaciation. According to them, loess deposition started at 13–11 ka and lasted until 2.7 ka. Four TL ages of early Holocene loess were obtained by Lehmkuhl et al. (2000) ranging from 10.5 ± 2.2 to 7.8 ± 1.2 ka for central and southern Tibet. In the area of Bomi in the south-east of the Tibetan Plateau a sample with a TL age of 25.5 ± 4.0 ka was measured (Lehmkuhl et al., 2000).

Longer and older sections are known from the edge of the north-eastern Tibetan plateau where loess sediments have a thickness of up to 40 m and a TL age of 152 ± 12 ka at 20 m (Fang et al., 2003; Lu et al., 2004). Even longer sections are known for the Xining area which reach a loess thickness over 250 m (Lu et al., 2004; Vriend and Prins, 2005; Sun et al., 2011). Despite being often attributed to the

64 north-eastern Tibetan Plateau, these areas have a different geomorphological and environmental setting at lower altitude and are not considered further here.

Results from the Donggi Cona catchment fit well into the framework of the existing dataset of dating results on aeolian sediments on the north-eastern Tibetan Plateau, indicating a start of the loess deposition in the Lateglacial. Loess deposition was strongest, during the early Holocene. This is different from the timing of the main loess deposition on the Chinese Loess Plateau. While on the Loess Plateau loess accumulation was strongest during cold periods (Kukla and An, 1989; Kohfeld and Harrison, 2003; Vriend et al., 2011) main deposition on the Tibetan Plateau occurred under warmer climate. Several dates from the Qinghai Lake (Porter et al., 2001; Lu et al., 2011b) and from the Qilian Shan (Stokes et al., 2003; Küster et al., 2006) show loess deposition during the mid-Holocene and the late Holocene. This cannot be confirmed for our study area. Sample DC06-07 with an age of 4.45±0.38 ka is most likely not typical for the loess deposition and might be more related to local short-term processes. These differences in the loess deposition might be related to the more northern location of the Qinghai Lake and the Qilian Shan and the closer source areas of the sediments. Loess ages from southern Tibet (Lehmkuhl et al., 2000; Kaiser et al., 2009a, 2010; Lai et al., 2009) with loess deposition older than the Lateglacial might be caused by a better conservation of the sediments.

Sand deposits have achieved considerable less attention on the Tibetan Plateau compared to loess deposits. Lu et al. (2011b) reported 12 OSL dates on aeolian sands from the Qinghai Lake ranging from 10.24 ± 0.59 ka until 7.78 ± 0.74 ka and one age of 0.62 ± 0.07 ka. In the valley of the Kyichu River 15 OSL ages of aeolian sands range from 81.0 ± 7.0 ka to 2.9 ± 0.2 ka (Kaiser et al., 2009a). However, excluding the oldest one minimizes the upper range to 28.2 ± 3.1 ka. Additionally, Owen et al. (2006) reported an age of 14.9 ± 1.5 ka from the northern flank of the Kunlun Shan.

Comparing the results from this study with the ages of dune sand from the Qinghai Lake area shows an even older onset of the sand deposition at the Donggi Cona. While at the Qinghai Lake aeolian sands have a maximum age of around 10.5 ka, the oldest sand at the Donggi Cona an age of 14 ka and the early phase of intensified sand deposition started at around 10.5 ka. Ages in southern Tibet are again considerably older. This early phase of sand deposition ended at the Donggi Cona at around 7 ka. From 3 ka onto present there is again a phase of major sediment deposition (with two sub- phases). A similar phase of intensified sand deposition is neither documented for the Qinghai Lake in the north nor in southern Tibet. At the present stage it is not possible to asses if this is related to unique conditions in the Donggi Cona area or to different sampling strategies.

Colluvial loess deposits in north-eastern Tibet have been radiocarbon dated from 8988 ± 66 to 4145 ± 60 yr BP (uncalibrated) (Kaiser et al., 2009a). In the A’nyêmaqên Shan a colluvial loess has been dated to 7.1 ± 0.7 ka (Owen et al., 2003). Results from the Donggi Cona show intensified reworking of loess and loess-like sediments from 9 to 5.7 ka matching the existing record. Additional dates occur at 11.7 ka, 4.68 ka and 8.86 ka in the Donggi Cona area and might be related to single events.

Palaeoclimate implications

First indications of aeolian sedimentation in the Donggi Cona catchment did not appear before the late Pleistocene. However, with only two ages the evidence for this signal is relatively weak. 65

Sediment availability was high during Lateglacial times (Nilson and Lehmkuhl, 2001). In the nearby A’nyêmaqên Shan a glacier advance took place around 16 ka (Owen et al., 2003) resulting in high sediment mobility. Large amounts of detrital inflow are recorded in the lake sediments from 19 cal. ka BP onwards and lake levels were rising until 14 ka (Dietze et al., 2010; Opitz et al., 2012). Despite this high availability of sand and loess, no sufficient dust trap was present. The first start of warmer and wetter conditions after cold glacial times is recorded at Lake Qinghai about 14.1 cal. ka BP (Shen et al., 2005), followed by a period of again colder conditions at around 12 ka during the time of the Younger Dryas (Yu and Kelts, 2002; Shen et al., 2005). At Donggi Cona, a period of high aridity and weakening of the summer monsoon from 13.4 to 11.9 cal. ka BP is reported (Mischke et al., 2010a), while Opitz et al. (2012) inferred lower lake levels in this interval. Harsh conditions hampered the continuous fixation of sediments and sediments were continuously reactivated. Similar cold and dry climate conditions at the end of the Pleistocene are also known from other lakes in north-eastern Tibet (Mischke et al., 2008, 2010c; Herzschuh et al., 2009, 2010b).

Intensified aeolian deposition started in the very early Holocene with the deposition of aeolian sands and loess. The deposition only started after the formation of a sufficient vegetation cover to fix the sediments. Both sediment types are indicating the shift to more humid conditions after the dry and cool glacial period. Climate amelioration on the Tibetan Plateau and adjacent regions started again at around 11 ka with an increase in available moisture (Herzschuh, 2006; Cai et al., 2010; Y. Wang et al., 2010). At the Qinghai Lake climate amelioration following the Younger Dryas period is also documented from 11 ka onward showing the strengthening of the Asian summer monsoon (Lister et al., 1991; Yu and Kelts, 2002; Colman et al., 2007), probably in relation to insolation changes at low latitudes (e.g. Kutzbach, 1981). At the Donggi Cona warm climate conditions are recorded between 12.5 and 6.5 cal. ka BP (Mischke et al., 2010a; Opitz et al., 2012). Several records from north-eastern Tibet show a short climate deterioration around 8 ka which is often correlated with the 8.2 ka event (Shen et al., 2005; Mischke et al., 2008; Schlütz and Lehmkuhl, 2009; Zhang and Mischke, 2009; Mischke and Zhang, 2010). With the full establishment of the Asian summer monsoons the source areas for deflation decreased significantly and fluvial erosion became an important factor in this high altitude landscape. Loess was removed from the slopes and deposited at the footslopes and the river terraces from around 9 to 6 ka. The end of this phase of reworking is most probably not related to climate changes, e.g. reduced precipitation, but rather to a reduction in sediment availability. Most former loess deposits were eroded until this time. A reconstruction of aeolian processes on the north-eastern Tibetan Plateau from lake sediments of Lake Kuhai about 30 km east of the Donggi Cona by Mischke et al. (2010c) shows reduced aeolian flux to the lake from 12.8 BP to 7.1 cal. ka BP. In the following period until 6.1 cal. ka BP aeolian influx became stronger again, followed by a phase of maximum Holocene aeolian input until 5.4 cal. ka BP. This strong aeolian flux into Lake Kuhai cannot be confirmed by the data from the Donggi Cona and is more likely related to fluvial processes due to moister conditions.

A massive reactivation of sand in the study area started from around 3 ka and peaked from 2 ka to the present. Two sections from the dune fields show a sand accumulation of nearly 5 and 4 m since that time (Fig. 4.2.10), indicating dramatic changes in the environment. In contrast to the sediments from the early Holocene indicating the trapping of aeolian sediments, late Holocene dune activity is related to disturbances in the existing vegetation cover. This disturbance can be related to climatic changes or to an increased anthropogenic influence. Many palaeoclimatic records show a shift to cooler and drier conditions from around 5 ka to the present (Shen et al., 2005; Morrill et al., 2006;

Herzschuh, 2006; Colman et al., 2007). According to a recent summary by (Y. Wang et al., 2010) the different parts of the Asian summer monsoons show a slightly different behaviour during this phase. In the region of the Indian Summer Monsoon available moisture decreased from 6.5 ka onward after the early Holocene optimum while in regions influenced by the East Asian Monsoon the decrease of moisture was probably delayed by another 2 ka (Y. Wang et al., 2010). For the Donggi Cona lake records show a shift from saline to freshwater conditions around 4.3 cal. ka BP (Opitz et al., 2012). This has been attributed to reduced summer insolation and a resulting positive water budget, leading to an overspill of the lake. However, a tectonic influence in this highly active area is also possible. Afterwards lake level at the Donggi Cona dropped under dry and cool climate conditions. Additionally, several short-term climate fluctuations took place on the eastern Tibetan Plateau during the last 4 ka. Several lakes show cooler climate conditions around 2.5 ka (Shen et al., 2005; Mischke et al., 2010c) and in the last 600 years which might correspond to the timing of the Little Ice Age (Liu et al., 2009; Zhang and Mischke, 2009; Mischke and Zhang, 2010; Mischke et al., 2010c). Short-term climate fluctuations have also been identified in tree rings (Yang et al., 2003, 2009). Climate deterioration therefore played an important role in the reactivation of the aeolian sediments. However, there is an increasing amount of data indicating an intense human influence on the Tibetan Plateau during the late but probably also from the mid-Holocene onwards (Kaiser et al., 2009a; Miehe et al., 2009; Schlütz and Lehmkuhl, 2009). At present, the area of the Donggi Cona is used as grazing area by nomads during spring and autumn. Especially the utilization in autumn leads to a disturbance of the vegetation cover. As a result, the soil is more vulnerable to deflation during winter times. A comparison of satellite data from the late 1960s and the year 2003 shows a strong increase in campsites especially on the eastern alluvial fan. At the present stage we are not able to decipher climatic and human influence on the reactivation of sandy sediments during the late Holocene.

Conclusions

Aeolian sediments at Lake Donggi Cona reveal a complex pattern in spatial and temporal distribution. Loess and sandy loess sediments are preserved at protected sites while dunes cover large parts of the southern and western catchment. Grain-size analyses indicate a combination of long-distance transport and local material. Distinct phases of deposition have been identified for different sediment types. Aeolian deposition in the area was strongly enhanced during the early Holocene, and in phase with warmer and wetter climate conditions during the strengthening of the Asian summer monsoons. This is in contrast with the more intense deposition during cold climate on the Chinese Loess Plateau. Dune sands have been deposited from around 10.5 ka – 7 ka. Loess sediments have a slightly younger age. Under full monsoon conditions from around 9 ka onward aeolian archives were destroyed and colluvial deposits were formed at the footslopes and on the fluvial terraces. At around 6 ka reworking of sediment ceased due to the lack of erodible material indicating the importance of sediments sources in these environments. During the late Holocene from 3 ka to present, a drier and cooler climate, possibly in combination with an increased human impact, lead to the reactivation of dune sand and intensified aeolian activity. Furthermore, new source areas like former lake bottoms provided additional sediments.

The possibility of sediment trapping is the key factor for the onset of sediment deposition after the Last Glacial. However, during the mid to late Holocene aeolian sediments and especially aeolian sands are easily reactivated by small-scale disturbances in the vegetation cover. This leads to the 67 redeposition of the sediments. In this case aeolian sediments are indicative for an aridification of the climate and/or a human influence in the form of grazing. We conclude that the connection between climate and aeolian system response is not as straightforward as stated before (e.g. Lu et al., 2011b). However, this might be especially due to the particular environmental conditions in the high altitude study area, but also be true for other sites on the Tibetan Plateau.

Acknowledgement

We thank Marianne Dohms from the sedimentological laboratory of the RWTH Aachen University for measuring the grain-size properties and Christoph Schmidt from the University of Cologne for the analysis of the LM-OSL data. Furthermore, several Chinese and German students helped us in field. Special thanks are denoted David Loibl who supported the sampling during the field trip in winter times. This research has been done in the frame of the project ‘Landscape and Lake-System Response to Late Quaternary Monsoon Dynamics on the Tibetan Plateau - Northern Transect’ which has been funded by the German Science Foundation (DFG) as part of the SPP 1372: Tibetan Plateau – Formation Climate Ecosystems. Xiaoping Yang, Jef Vandenberghe and an anonymous reviewer are thanked for their helpful comments on an earlier version of this manuscript.

4.3 Interaction of geomorphological processes on the north-eastern Tibetan Plateau during the Holocene, an example from a sub- catchment of Lake Donggi Cona

Georg Stauch, Steffen Pötsch, Hui Zhao, Frank Lehmkuhl

Abstract

Geomorphological landforms on a large alluvial fan and neighboring areas on the eastern side of the Donggi Cona show a complex spatial pattern. Sediment availability is an important factor in the formation of these archives and is partly associated with lake level fluctuation. Different sedimentary archives therefore show a similar geomorphological signal while in some cases similar archives are related to different forcing. The chronology of the processes is based on 22 optical stimulated luminescence (OSL) ages obtained from coarse-grained quartz or potassium-feldspar. The base of the alluvial fan formed during the late Pleistocene but deposition of cover sediments on the fan only started at around 7 ka ago. Silty sediments form a thin cover on the uppermost terraces and have been available only during short timespans. Most of the sediments are sandy deposits which form an internal sediment cycle on the fan. They developed throughout the Holocene. Sandy deposits on the footslopes of the neighboring hills have been preserved since the Pleistocene-Holocene transition. Thicker aeolian deposits accumulated during the late Holocene. These sediments were washed into the Donggi Cona and formed sandy lake sediments which were susceptible to remobilisation during lower lake levels. In combination with drier climatic conditions during the late Holocene these sediments formed small dunes on the fan. Sediments on the alluvial fan are highly active, for which erosion and deposition vary in space and time. The interaction of fluvial and aeolian processes is an important part of this dynamic high mountain system.

Keywords: Holocene, geomorphology, sedimentology, dating, OSL, Tibetan Plateau, palaeoclimate

Introduction

Geomorphological landforms and processes are used widely to reconstruct the palaeoclimatic evolution, tectonic influence and human impact on the Tibetan Plateau. However, most studies regarding palaeoclimate reconstructions focus on one single archive, e.g. aeolian deposits (Lehmkuhl et al., 2000; Sun et al., 2007; Lu et al., 2011b), fluvial and colluvial sediments (Kaiser et al., 2007), lake sediments (Madsen et al., 2008; Mischke et al., 2008, 2010c; Opitz et al., 2012) or glacial sediments (Lehmkuhl, 1998; Owen et al., 2003; Heyman et al., 2011). Beside the effects of climate change, geomorphological archives often show evidence of tectonic activity and human influence. There is a wealth of publications reconstructing tectonic movements from geomorphological archives on the north-eastern Tibetan Plateau (e.g. Van Der Woerd et al., 2002; Fu and Awata, 2007; Guo et al., 2007; Kirby et al., 2007). For several decades an increased human influence on the environment and

69 geomorphological processes has been observed (Cui and Graf, 2009; Yan et al., 2009). However, recent publications also highlight an anthropogenic influence over the last several thousand years (Miehe et al., 2009; Schlütz and Lehmkuhl, 2009).

Studies focusing on larger sets of different geomorphological archives often deal with large areas and therefore have a low spatial resolution (e.g. Owen et al., 2006). Hardly any work in the area has tried to identify the entire sediment transport pathways on the catchment scale, both spatially and temporally. Information on these sediment cascades will deepen the understanding of the underlying driving forces, such as climatic changes, tectonic and human influences on these systems. This is especially important as geomorphological processes are interacting with each other and internal feedbacks in the system play an important role in the formation and preservation of geomorphological landforms (e.g. Field et al., 2009).

In this paper different geomorphological landforms, surface types and related geomorphological processes of a multicomponent system in the catchment of the Donggi Cona on the north-eastern Tibetan Plateau are identified. The area is a typical example for a complex pattern of sediment storage and transportation which is sensitive to environmental changes in a high altitude semi-arid climate. Aeolian, fluvial and lacustrine processes are interacting and result in a spatial and temporal different response of the geomorphological system to the climatic forcing. Timing of the processes during the Holocene is constrained by 22 optical stimulated luminesces (OSL) ages. The results are combined in a tentative landscape reconstruction and finally in a simple model of the different processes operating during different time-scales. A special focus is set on the complex interaction of sediment archives. Many palaeoclimatic reconstructions rely on the interpretation of lake sediments either from the lake itself or from onshore archives such as lake terraces. This study also provides basic information on terrestrial sediment activity, which is relevant for the interpretation of lake sediments.

Regional setting

Lake Donggi Cona is located in the Kunlun Shan on the north-eastern Tibetan Plateau (35°18’ N / 98°32’ E) at an elevation of 4090 m asl (above sea level) and is probably one of the best studied catchments on the Tibet Plateau. The catchment has a size of 3174 km² and is dominated by a WNW- ESE oriented pull-apart basin (Fig. 4.3.1). The basin is part of the Tuosu segment of the Kunlun Fault which is one of the major faults on the northern Tibetan Plateau (Yin and Harrison, 2000; Van Der Woerd et al., 2002; Fu and Awata, 2007; Guo et al., 2007). The last major earthquake in 1937 (M 7.5) caused several mole tracks which are still traceable at the surface east and west of the lake (Tapponnier and Molnar, 1977; Van Der Woerd et al., 2002; Guo et al., 2007). The horizontal offset has been estimated to be about 4 m, while the vertical movement was less than 1 m (Van Der Woerd et al., 2002; Guo et al., 2007). The surrounding mountains consist of Triassic and Permian lime- and sandstones, as well as of clastic rocks of Permian to Neogene age. The basins are filled by Quaternary sediments (Wang et al., 1998; Wang and Yang, 2004).

The highest elevation is located in the south-eastern part of the catchment (5230 m asl), which is the western extension of the A’nyêmaqên Shan. On the north-western side of the lake the mountains reach an elevation of up to 5050 m asl, while most areas in the catchment are below 4300 m asl. Remnants of Quaternary glaciations are frequent in the higher parts, consisting of cirques, U-shaped 70 valleys and terminal moraines. No modern glaciations exist in the higher part of the catchment. The nearest presently glaciated mountains are the A’nyêmaqên Shan, about 30 km south-east of the catchment with a maximum elevation of 6282 m asl. During the late Quaternary large valley glaciers developed in this mountain range (Wang, 1987; Lehmkuhl, 1998; Owen et al., 2003). Four terminal moraines have been mapped on the eastern side of the A’nyêmaqên Range. Three of them have been dated by CRN (cosmogenic radionuclides) dating of boulders from moraine crests to around 45 ± 5, 16 ± 3 and 9 ± 3 ka (Owen et al., 2003). The fourth moraine which has not been dated is only 1 km away from the present glacier margin (Owen et al., 2003). Terminal moraines of past glaciations are also preserved in the mountains north of the Donggi Cona. Field observations revealed terminal moraines less than one kilometer away from the present lake shore.

The Donggi Cona is mainly fed by the Dungqu River which is draining the mountains in the south-east of the catchment. Several ancient shore-lines have been recorded below (Dietze et al., 2010) and above the present lake level (Van Der Woerd et al., 2002; Mischke et al., 2010a; IJmker et al., 2012b; Dietze et al., 2013).

The area is located at the northern boundary of the Asian summer monsoon (e.g. Morrill et al., 2006). The climate is semi-arid, with mean annual precipitation of 311 mm y-1. More than 50% of the precipitation occurs in summer (June, July and August) and more than 80% from May to September. Thunderstorms resulting in high discharge events occur frequently during this time. Mean temperature for January and July are -15.8 and 7.9 °C. The mean annual temperature is -3.0 °C (Mischke et al., 2010b). Alpine meadows, alpine steppes and Salix scrubs on the slopes characterize the vegetation in the catchment (Huang, 1987; Kürschner et al., 2005).

Fig. 4.3.1: Overview of NE Tibet. The study area is marked with the black box in the main figure.

In this study, research focused on a large alluvial fan and the neighboring slopes in the eastern part of the pull-apart basin with an area of 180 km² (Fig. 4.3.2). The fan stretches for about 20 km in NNW to SSE direction and has a width of 10 km. Lake Donggi Cona is located on the western end of the alluvial fan, while the southern border is formed by a small mountain range consisting mainly of carbonate rocks of Permian age. The maximum altitude of this mountains is 4590 m asl. On the eastern side of the mountain range the Dongqu River enters the basin close to the village of Huashixia at an elevation of 4210 m. To the west an undulating surface borders the fan from the high mountain area. It consists of older gravels which might be related to a former west-to-east flowing palaeo-river which drained the whole basin prior to the development of the lake and the reorientation of the river system (Li et al., 2000). An elongated ridge (push-up fold, Van Der Woerd et al., 2002) with an elevation of up to 4280 m separates the NE part of the basin. This ridge is located directly on a fault line (Van Der Woerd et al., 2002). During the field work, seven steps were observed on its northern side. On the lower slopes small ponds and hills alternate. They might have formed during the Tuosuo Lake earthquake in 1937. The rupture zone is clearly traceable to the east and west of the mountain ridge. The north border of the fan is marked by a mountain range with maximum elevations of 4605 m asl.

Large parts of the fan are utilized as pastures, especially in spring and autumn. An exception is an area in the south-west of the fan which is only accessible during winter when the deep river courses are frozen.

Fig. 4.3.2: Study area (see Fig. 4.3.1) and sections mentioned in the text.

Two large catchments are connected to the eastern Donggi Cona basin. From the south the Dongqu River enters the basin. The river has a catchment area of 750 km² and drains the western A’nyêmaqên Range. The ranges on both sides of the river show clear traces of past glaciations. This

72 river builds up large parts of the alluvial fan. From the northern side a second inflow enters the area. The catchment of this river has an area of 607 km². However, its influence on the fan is only marginal and restricted to the northernmost part. Most sediments are retained in a neighboring basin before the river reaches the study area. Both catchments show distinctly different geomorphometric signatures which explain much of the different sediment delivery characteristics. The larger southern catchment of the Dongqu River has a mean elevation of 4701 m asl and an elevation range of 1033 m. 10% of the catchment area is situated above 4800 m asl. The northern catchment has a mean elevation of 4466 m asl and a range of 802 m and the area above 4480 m is only 10% of the overall catchment. A further interesting characteristic of the catchment is the slope angles. In the southern catchment 54% of the area has a slope of 10° or more, based on the SRTM (Shuttle Radar Topography Mission, Farr et al., 2007) grid, while in the north only 15.5% have such steep slopes.

Methods

A wide spectrum of methods has been used to get a detailed insight into the recent and past geomorphological processes in this relatively small area, including field investigations, multi- temporal remote sensing analysis as well as sedimentological and geochemical analysis. Age control on different sediment archives was provided by radiocarbon and OSL dating.

Geomorphological Mapping

Field investigations took place in spring 2008 and 2009 and in summer 2009 and included detailed geomorphological mapping, e.g. of river terraces, aeolian cover sediments and permafrost features. Selected sections were sampled for grain-size analysis and geochemical studies.

The multi-temporal remote sensing analysis is based on seven different satellite images covering a time-span from 1968 (Corona images) to 2009. However, for the geomorphological mapping we relied mainly on a Landsat ETM+ dataset from May 2003 with a spatial resolution of 15 m (panchromatic) and the Corona images from November 1968 with a spatial resolution of approximately 5 m. Images from 1976 (MSS), 1989 (TM), 2001 (ETM), 2005 (TM) and 2009 (TM) were used to obtain insights into the modern morphodynamic processes operating on different parts of the fan. Mapping was based on visual image interpretation. Spatial elevation data was obtained from SRTM data provided by the USGS. Generally SRTM data were in a good agreement with the measured GPS data sets from the field. Nearly 80% of our 321 GPS measurements differ by no more than 10 m from the related SRTM values, while the comparison of the measurements with the Aster GDEM produced by METI and NASA revealed that less than 30% fall in the same range.

Sedimentological analysis

A total number of 99 sections were sampled on and in the vicinity of the fan, resulting in more than 350 individual samples. Grain-size was determined using a Beckmann Coulter Laser Diffraction Particle Size Analyzer (LS 13320). Fractions larger than 2 mm were removed previously by sieving. For preparation, carbonates and organic material were removed with HCl (10%) and treated with H2O2 73

(20%) at 70°C for several hours. Sodium pyrosphate was used for dispersion. Geochemical interpretation to identify possible source areas of the sediments is based on the dataset analyses of (IJmker et al., 2012b).

OSL and radiocarbon dating

To obtain a chronology of the geomorphological processes, OSL dating was applied in the laboratories of the University of Cologne and Lanzhou. Light tight steel tubes with a diameter of 56mm were used for sampling and only the central part of the sample was taken from the tube for OSL measurement. Preparation of the samples included treatment with hydrochloric acid (10%) and hydrogen peroxide (10% in Cologne, 20% in Lanzhou). Heavy liquids of sodium polytungstate with densities of 2.68 g/cm3 and 2.58 g/cm3 were used to obtain the quartz and K-feldspar grains from the fraction of 100 to 250 μm. For the quartz grains, 40% HF (45 min Cologne, 60 min Lanzhou) was used to remove the outer layer irradiated by alpha particles of the quartz grains and any remaining feldspar grains. For the K-feldspar grains, the removal of the outer layer which was irradiated by alpha particles was achieved by treatment with 10% HF for 40 min. Finally HCl (60 min Cologne, 10 min Lanzhou) was used to remove fluorides created during the HF etching. After re-sieving the material, the fraction of 100 to 150 μm (quartz) and 100 to 200 μm (feldspar) were used for OSL measurements.

The equivalent dose (De) was determined with an automated Risø luminescence reader (types TL/OSL-DA-12, -15, or -20) equipped with 90Sr/90Y-β-sources for irradiation, blue light and infrared light-emitting diodes and EMI 9235 photomultiplier tubes for measurements of the luminescence signal. IRSL-tests (to check for feldspar contamination) and LM-OSL tests (to check which component is dominant within the quartz) where carried out. The results showed that the quartz is not suitable for a successful dating application with the standard SAR protocol (Stauch et al., 2012). Therefore, for the measurements of the quartz grains (Lanzhou), a post-IR OSL single aliquot protocol was applied to minimise the influence of any feldspar grains (Zhang and Zhou, 2007). Only the aliquots with low IRSL signals (read at 50°C and <5% of OSL signal), low recuperation (<5% of the sensitivity corrected natural signal) and recycling ratios within 10% of unity were selected for the De calculation (Wintle and Murray, 2006). A preheat of 260°C, a cut-heat to 220°C and a test dose of 3.12 Gy were used for all post-IR OSL measurements based on a preheat plateau test. The OSL signal was carried out with blue LED for 80s at 125°C. The De was calculated by the weight mean from all results of selected aliquots. For the three samples measured in Lanzhou, no unstable medium signals were noted in OSL signals from quartz as observed by the decreasing trend of De(t) (Li and Li, 2006). To determine the De values of K-feldspar (Cologne) the SAR protocol of Wallinga et al., (2000) was used. Here the IRSL stimulation was carried out at 30°C for 300s following a thermal pretreatment for 10s at 210°C (for samples with a De <24 Gy) and 240°C (for samples with a De between 24 and 47 Gy) (Table 4.3.3). Scattering of the individual De-values can arise from natural variations of the palaeodose, e.g. partial bleaching, differences in the microdosimetry within the sediment matrix or bioturbation. The De- values of sample DC09-36-1 and DC09-36-2 showed a broad scattering, therefore the age was calculated with the Finite Mixture Model (FFM) (Table 4.3.3). Reasons for this scattering can be one or a combination of the above mentioned factors. Because only aeolian sediments where sampled and analysed, the samples can be considered as well bleached (Stauch et al., 2012).

The environmental dose rate assessments involved several techniques. Gamma-ray spectrometry (Cologne) and neutron activation analysis (Lanzhou) were used to determine the concentration of U, Th and K. The dose rate contribution from cosmic rays was calculated by considering the burial depth, the altitude and the geographical position (Prescott and Hutton, 1994). The water content was determined by the ratio of the weight of water in the sediment to the weight of dry sediment and expressed as a percentage. Due to the uncertainties in water content during the burial, assumed water contents of the samples determined from the laboratory measured values were applied into dose rate calculations. The internal dose rate from 40K in K-feldspar grains was calculated with a K content of 13±1% (Lanzhou) (Zhao and Li, 2005) and 12±0.5% (Cologne) (Huntley and Baril, 1997).

A fading test was carried out for one sample (DC09-36-2) using K-feldspar fraction as proposed by Auclair et al., (2003) to check the possible IRSL signal anomalous fading. The resulting age underestimation of 21.1% corresponds to a mean g-value of 2.3 ± 0.6%. Since there are still not enough independent age estimations given, all IRSL ages are ‘fading uncorrected’ and ought to be considered as minimum ages. A detailed description of the methods used for the OSL dating including fading tests for the samples from the study area has been given by Stauch et al. (2012).

Additionally five mollusc shells from three sites have been analyzed by AMS (Accelerator Mass Spectrometry) radiocarbon dating method at the University of Erlangen.

Results

Geomorphology

The alluvial fan is distinctively separated into two parts (Fig. 4.3.2); a smaller western part (55 km²) with a mean slope angle of less than 0.1° and a larger eastern part (125 km²) with a mean slope angle of 0.5°. The apex of the fan is located close to the village of Huashixia in the south-east where the Dongqu River enters the basin. The eastern part of the fan is built up of fluvial gravels of at least two meters height according to section P026. The typical size of the gravels is about 4-5 cm. A borehole from the year 1990 revealed in the upper part of the fan close to the apex 2.5 m of gravels followed by at least 40 meters of sandy clays and clayey sands (Jin et al., 2007). On the northern side where small fans enter the basin from the neighboring mountains, layers of angular solifluction material and fine material alternate for several meters depth. While the fine layers in the upper one meter of these fan deposits are dominated by silt size particles, the lower part of the profiles consists mainly of sandy sediments. Fluvial gravels on the eastern part of the fan cover nearly 80 km². Several erosional terraces are incised into the gravels. Two main steps were mapped on the basis of remote sensing data and field results, resulting in three different terrace levels (T0-T2). The lowermost level (T0) is presently used by the river. The second level (T1) is only a few decimeters higher. Multi- temporal remote sensing analysis showed that this level was subject to most of the changes during the last 35 years, especially in the upper part of the fan. This indicates a relatively young age of formation. The uppermost terrace level (T2) is the most stable one showing only a slight loss in its spatial extent during the past decades. The height of the terrace level is up to 3 m above the present river level. The T2 level also forms the basement in the northern part of the fan north of the main river channel. At present the main river is deflected at the major fault line which crosses the alluvial fan (Fig. 4.3.2). Two terraces have also been mapped in the lower part of the Dongqu River upstream

75 of the village of Huashixia. In contrast only one terrace level has been observed at the northern periodic inflow.

The slightly inclined western part of the study area consists mainly of sands and clay. The river systems show a much higher dynamic state, and up to four different levels have been distinguished. The terraces are related to phases of high water discharge and frequent changes in the river course.

The present geomorphology is primarily controlled by fluvial and aeolian processes. The main periods of river discharge are spring and summer. Spring discharge is caused by snowmelt and is geomorphologically of minor importance compared to summer discharge which results from torrential rainfalls. These can lead to short discharge events of magnitudes several times higher than the average. Such torrential rainfall events have been observed, e.g. in summer 2009 leading to bankfull discharge of the Dongqu River. Generally rainfall events in the area are spatially and temporally highly variable. Presently, during most of the time sand-size particles are transported. Transport of gravels was only observed during bankfull discharge events. The transport of gravels generally ended at the western edge of the inclined part of the fan. On the northern side of the eastern part dunes of up to 2.5 m are preserved. Large aeolian sand deposits are found on the slopes of the neighboring hills. In the central and southern part of the western area only small dunes, up to 1.5 m in height, are developed.

On the basis of the Corona image from 1968, eight different surface types (Table 4.3.1, Fig. 4.4.3) have been mapped in the study area. The classification was supported by field observations and the use of Landsat ETM+ data.

Present fluvial channels, which are supposed to be at least partly active during bankfull discharge in summer, cover an area 31 km². They are identical to the T0 surface and are filled with gravel and sand.

Table 4.3.1: Surface types of the alluvial fan.

Type Area [km²]

Fluvial (recent) 31.49

Fluvial gravel surface 23.62

Fluvial sand 53.03

Transition zone 13.68

Clayey lake sediments 11.44

Sandy lake sediments 26.07

Beach 1.29

Dunes 28.03

Sum 188.65

Fig. 4.3.3: Surface types and dated sections (star: aeolian; circle: reworked; box: sandy lake sediments).

In an area of 24 km² fluvial gravels appear at the surface. These areas are easily identifiable in the Corona images by the large predominantly hexagonal ice-wedge pattern (Fig. 4.3.4) with a mesh size between 35 and 45 m. These patterns are only visible when the sediment cover on the gravels is thin or absent. They appear on the T1 and T2 levels. Close to section P026 two ice wedge casts in the river gravels have been observed at a river cut. The ice wedge casts have a depth of about 100 cm and a width of 70 cm and are 2.5 m apart from each other (Fig. 4.3.5). However, the top of the ice wedge casts is not identical to the top of the gravels. The ice wedge casts are covered by another 20 cm of gravels. Due to small size of the casts, the short distance between them and the gravel cover on top these ice wedge casts cannot be responsible for the large ice wedge pattern on the surface. This indicates the presence of several generations of ice wedge casts.

Fig. 4.3.4: Ice wedge pattern at the T2 surface and ice wedge casts in a river cut to section P026 representing different generations of cast formations.

By far the largest surface type is the area of the fluvial sands (53 km²). The thickness of the sand deposits is 80 to 190 cm and they cover parts of the T1 and T2 terraces. Typical sand content is between 55 and 95%. At present these sediments are mainly transported during times of torrential rainfalls. They originate from degrading dune fields on the fan. Thickness decreases to the western end of the eastern part of the alluvial fan where many small stones with a diameter of 2 to 4 cm are observed on the surface. These stones are related to frost-heave processes. On the highest terrace T2 the ‘sand wash’ sediments show an upward coarsening from silty sediments in the lower part to sandy deposits close to the surface (see P026 in Fig. 4.3.5 and 4.3.6 for example).

A small step marks the transition zone to the western part of the fan. It stretches for about 8 km in a NNE direction and has a width of around 2 km (13.7 km²). The transition zone is quite obvious during summer time due to a large number of seepages from which many small streams originate. These seepages are controlled by permafrost in the fine sediments. This confining stratum forces water, infiltrated into the gravels in the eastern part, up to the surface. Therefore the area is swampy and densely covered with vegetation in comparison to the relatively dry eastern part of the fan. The sediments have high clay (up to 15%) and silt (up to 87%) content and are generally rich in carbonates. They represent a mixture of fine lake sediments and fluvial sediments from the upper part of the fan. Small frost mounds are typical and occasionally small collapsed pingo-like features with a diameter of about 2 m have been developed on the surface. The permafrost table is around 60 cm below the surface in the western part of the transition zone.

Fig. 4.3.5: a) Fan surface on the northern side, b) T2 with silt and sand cover (P026), c) clayey lake sediments (P066 - near P006), d) sandy lake sediments (P221), e) dune on the northern footslope (P246).

Lake sediments cover the western part of the alluvial fan. According to the sediment properties and post-depositional modifications two different types of lake sediments have been mapped, clayey lake sediments and sandy lake sediments. Lake sediments with high clay content dominate on the northern side and in a narrowing belt close to the present lake shore (11.4 km²). Typically they are located at elevations 8 to 10 m above the present lake level and are covered by 10 to 40 cm of silty sediments. Those have high carbonate content (IJmker et al., 2012b) and are interpreted as lake sediments from relatively shallow water. A typical profile (P066) is shown in Fig. 4.3.5c. Especially on the northern side of the study area, frost mounds and collapsed pingo structures with diameters of up to 8 m are visible. However, parts of these elongated surface depressions indicate deflation processes. At the north-eastern side the surface is frequently covered by small gravels. These originate from the eastern fan area and are presently transported by sheet wash processes down onto the lake sediments. Under present climatic conditions the permafrost in the clayey lake sediments in the north-west of the study area is degrading. From 1968 to 2009 the area covered by small ponds gradually increased from 0.09 km² to 0.81 km².

The proportionally largest area in the west (26 km²) is covered by sandy and partly silty sediments (sandy lake sediments). In most parts the sands are dissected by fluvial channels. At few locations remnants of a former flat surface have been preserved. This surface is 5 to 6 m above the present

79 lake level. A typical section of this area is shown in Fig. 4.3.6 (P024). The sediments are horizontally layered and consist of alternating layers of sand and silt (Fig. 4.3.5). We counted 65 layers in the upper 380 cm of the section. The thickness of the layers gradually increases from around 2 cm at the top of the section to 10 cm at 380 cm. Below 380 cm the silt content reach values of more than 80% and the clay content is around 10%. Numerous layers of small limnic snails (Radix spec) are intercalated in the sands. Due to the layering of sands and snails, a deposition in the lake but close to the shoreline is assumed. The succession of sediments indicates a gradually reducing water column due to the increased infilling. Shells were sampled at 125 cm and 345 cm for radiocarbon dating. An OSL sample from the silty sediments was taken at a depth of 400 cm. In local depressions and on river cuts, clayey lake sediments have been observed at the base. These sediments mirror the clayey lake sediments described above.

The smallest area is the present lake shore with sandy beach sediments. These sediments are still mobile during periods of higher waves and will not be considered further.

Dunes have been developed at several locations across the whole fan. They are widespread on the northern and central part of the eastern alluvial fan area and reach elevations of up to 2.5 m. However, they are presently eroded by fluvial and aeolian processes. Thicker deposits of up to five meters are present on the sides of the fan and on the lower slopes of the neighboring hills (Fig. 4.3.5). South-west orientated slopes are covered by dune sediments up to 150 m above the alluvial fan, becoming progressively finer with increasing altitude. This is in accordance to the main direction of the aeolian transport from west to east. Dunes in the north-western part of the fan, e.g. covering the clayey lake sediments, show significantly smaller grain-sizes but are limited to a small area. Their geochemical composition indicates a mixing with material from the former lake bottom (IJmker et al., 2012a). Dunes on the fan and on the north-eastern side originate from the fluvial sediments on the alluvial fan. However, dunes on the north-eastern side also have relatively high carbonate content in the upper part, indicating an additional sediment input from the former lake floor (IJmker et al., 2012b). Present-day aeolian deposition is widespread on hillslopes and on the leeward side of the degraded dunes.

While most of the dune sediments have a yellow to light brownish color, dune sediments on the eastern side of the fan have a brownish color only. These dunes are fixed by Salix shrubs with deep roots and are growing with the accretion of the sediments. Hydrologically the shrubs are bound to a local seepage; roots can be observed down to its water table. Two sections have been studied (P019 and P101). Both sections consist of fine sands. According to the geochemical analysis they underwent weak pedogenesis (IJmker et al., 2012b), probably related to the constant water supply and the dense vegetation cover. In section P101, below 257 cm, four organic-rich layers of up to four centimeters are evident. Molluscs of the Radix type are preserved in this lowermost part. In conjunction, these features point to the former existence of seasonal or annual ponds in the area which were frequently filled by aeolian sands. Above 250 cm aeolian sediment forms the upper part of the section. Mean Grain-sizes are slightly increased here compared to the lower part. At 260 and 280 cm molluscs (Radix type) were sampled for radiocarbon analysis. At 250 cm an OSL sample was taken.

Timing of sediment deposition

From 16 locations (Table 4.3.2) a total of 22 OSL (Table 4.3.3) and five radiocarbon samples (Table 4.3.4) from different geomorphological archives were analyzed. Fifteen OSL samples originate from dune sediments, four from sandy loess deposits on the hill on the eastern side of the fan, two from reworked loess deposits, and one sample has been taken from the sandy lake deposits in the west of the study area. The ages cover a timespan from 14 to 0.3 ka, with a majority of the ages (18 samples) between 3 and 0.3 ka ago. Where reported OSL ages are based on coarse-grained quartz minerals it is stated in the text while all other OSL ages have been dated using coarse-grained K- feldspars and should be regarded as minimum ages.

Table 4.3.2: Site information of the OSL and radiocarbon dated sections.

Altitude Sediment Section Northing Easting [m a.s.l.] classification No. of samples Method

P006 35.25933 98.74848 4099 Dune sand 1 OSL

P019 35.20032 98.8708 4164 Dune sand 1 OSL

P021 35.24477 98.7762 4108 Reworked loess 1 OSL

Sandy lake 14C P023 35.24558 98.72812 4102 sediments 1

Sandy lake OSL/14C P024 35.23068 98.73406 4105 sediments 3

P026 35.22505 98.82145 4118 Reworked loess 1 OSL

P027 35.22377 98.81076 4120 Dune sand 1 OSL

P028 35.2269 98.79793 4119 Dune sand 1 OSL

P030 35.2535 98.83316 4105 Dune sand 1 OSL

P068 35.24113 98.80424 4109 Dune sand 1 OSL

P101 35.15997 98.91447 4195 Dune sand 3 OSL/14C

P109 35.240366 98.85428 4123 Dune sand 1 OSL

P221 35.22156 98.89403 4142 Dune sand 6 OSL

P237 35.20876 98.93463 4209 Dune sand 1 OSL

P238 35.2093 98.93837 4299 Sandy loess 2 OSL

P270 35.1929 98.89576 4285 Sandy loess 2 OSL

Table 4.3.3: OSL dating results (Lab.-No T or C-L: Cologne, L: Lanzhou).

Section Sample Lab.-No Depth U [ppm] Th [ppm] K [%] Dose rate Water content Equivalent dose Age [ka] [cm] [Gy/ka] [%] [Gy] P006 DC08-01 L2008-41 120 1.77±0.04 6.71±0.16 1.14±0.03 2.07±0.05 1,1 2.82±1.32 1.34±0.63 P019 DC08-04 C-L2444 230 1.51±0.08 6.20±0.34 1.09±0.03 2.68±0.17 2-8 0.82±0.04 0.31±0.02

P021 DC08-05 L2008-43 36 2.86±0.08 10.3±0.27 1.86±0.04 3.26±0.08 2-8 9.32±0.86 2.86±0.28

P024 DC08-06 C-L2445 400 1.74±0.09 7.01±0,38 1.28±0.04 2.91±0.18 2-8 0.85±0.05 0.29±0.03

P026 DC08-07 L2008-44 77 2.18±0.07 9.45±0.27 1.50±0.03 2.74±0.07 2-8 19.3±0.9 7.06±0.37

P027 DC08-08 C-L2446 87 1.79±0.09 7.43±0.41 1.24±0.04 3.04±0.19 2-8 1.73±0.09 0.37±0.05

P028 DC08-09 C-L2447 180 1.39±0.07 5.87±0.32 1.07±0.03 2.63±0.16 2-8 1.03±0.06 0.39±0.03

P030 DC08-10 C-L2448 116 1.73±0.09 6.93±0.38 1.25±0.04 2.98±0.19 2-8 3.25±0.18 1.09±0.09

P068 DC09-13 C-L2710 57 1.94±0.10 8.16±0.54 1.20±0.05 3.11±0.21 2.5-7.5 1.03±0.05 0.33±0.03

P101 DC09-14 C-L2711 250 1.90±0.10 8.03±0.51 1.21±0.05 3.02±0.20 2.5-7.5 3.63±0.21 1.2±0.11

P109 DC09-19 C-L2716 203 1.66±0.09 6.82±0.44 1.04±0.04 2.74±0.18 2.5-7.5 1.38±0.07 0.5±0.04

P221 DC09-36-1 C-L2767 535 1.43±0.08 5.76±0.37 1.00±0.04 2.45±0.15 2.5-7.5 34.1±2.3 13.9±1.3

DC09-36-2 C-L2768 485 1.45±0.08 5.78±0.38 0.99±0.04 2.39±0.16 2.5-7.5 20.7±0.6 8.67±0.62

DC09-36-3 C-L2769 448 1.71±0.09 6.84±0.44 1.11±0.04 2.73±0.17 2.5-7.5 11.2±0.6 4.12±0.34

DC09-36-4 C-L2770 387 2.13±0.11 7.98±0.51 1.19±0.05 2.94±0.20 2.5-7.5 5.26±0.27 1.79±0.15

DC09-36-5 C-L2771 333 1.38±0.08 5.80±0.39 0.96±0.04 2.39±0.16 2.5-7.5 2.88±0.19 1.21±0.11

DC09-36-7 C-L2773 121 1.49±0.08 6.31±0.41 1.01±0.04 2.65±0.18 2.5-7.5 0.83±0.05 0.31±0.03

P237 DC09-37-1 C-L2774 257 1.31±0.07 5.70±0.33 0.98±0.04 2.42±0.16 2.5-7.5 0.94±0.05 0.39±0.03

P238 DC09-38-1 C-L2777 187 1.80±0.10 7.50±0.48 1.16±0.05 2.76±0.22 5-15 4.19±0.23 1.52±0.15

DC09-38-3 C-L2779 135 1.73±0.10 7.32±0.48 1.19±0.05 2.78±0.22 5-15 1.78±0.09 0.64±0.06

P270 DC09-40-1 C-L2781 91 2.30±0.12 8.82±0.58 1.30±0.05 3.02±0.24 10-20 5.96±0.35 1.97±0.19

DC09-40-3 C-L2783 53 2.32±0.12 9.67±0.57 1.50±0.06 3.27±0.25 10-20 3.64±0.2 1.11±0.11

All samples were analyzed analyzed in 2010 and are not fading corrected. The assumed water content, the Central Age Model (CAM), Finite Mixture Model (FFM) and the Common Age Model (CoAM) were used for age calculations. An internal potassium content for the IRSL-ages was set to 12.5 ± 0.5% K (Huntley and Baril, 1997), the dose rate conversion factors used following Adamiec and Aitken (1998) and the source reference of the cosmic dose rate are after Prescott and Hutton (1994). The uncertainty on equivalent dose values are derived from the standard deviation about the mean (1σ standard error) and a systematic error of 5% associated with the beta source calibration are taken into account. Final ages have been rounded to two decimal places.

First indications of aeolian sedimentation are available from sand dunes at the northern side of the alluvial fan. At section P221 (Fig. 4.3.6) at the footslopes of a small hill, six OSL samples have been optained. Dune sands at a depth of 535 cm have an age of 13.9 ± 1.3 ka (DC09-36-1) (Table 4.3.3) which is the oldest age obtained in the vicinity of the alluvial fan. Ages of 8.67 ± 0.62 (DC09-36-2) and 4.12 ± 0.34 ka (DC09-36-3) were gained from depth of 485 and 448 cm, respectively. The upper 400 cm of the profile are younger than 2 ka (DC09-36-4, DC09-36-5) and the youngest age of 0.31 ± 0.03 ka (DC09-36-7) was obtained from a depth of 121 cm. All other dunes on the alluvial fan and the neighboring slopes have OSL ages less than 1.5 ka (Table 4.3.3); a dune on the northern slope has been dated to 0.39 ± 0.03 ka (DC09-37-1) while sand dunes on the eastern part of the alluvial fan have ages between 1.09 ± 0.09 ka (DC08-10) and 0.5 ± 0.04 ka (DC09-19) in the northern part, and around 300 to 400 years in the central part (DC08-08, DC08-09, DC09-13). The brownish dunes on the eastern edge of the alluvial fan show a wider age range. At P019 an OSL age of 0.31 ± 0.02 ka (DC08- 04) (Table 4.3.3) was obtained and a sample of P101 has an age of 1.2 ± 0.11 ka (DC09-14) at a depth of 250 cm. Additionally two radiocarbon dates are available for P101 (Table 4.3.4). A Radix shell from 260 cm yielded an age of 4776 ± 124 cal. yrs BP (P101-M1) while a Radix shell from 280 cm gave an age of 3976 ± 122 cal. yrs BP (P101-M2). Opitz et al. (2012) and Mischke et al. (2010a) suggested a reservoir effect for radiocarbon dating from lake Donggi Cona of 2000 years. As this part of the catchment has never been flooded by the lake, no indications about a potential reservoir effect are available. Beside changes in the reservoir effect, the age inversion in the radiocarbon ages might also indicate a remobilization of older sediments. However, the underlying process is not obvious from these deposits.

A dune covering the clayey lake sediments on the western part of the study area gave an OSL age of 1.34 ± 0.63 ka in a depth of 120 cm (DC08-01, quartz). The hills on the north-eastern side of the alluvial fan are covered by sandy loess. Dating results from two sections (P238, P270) yielded OSL ages of 1.97 ± 0.19 to 0.64. ± 0.06 ka (DC09-38-1, DC09-38-3, DC09-40-1, DC09-40-3) (Table 4.3.3).

The sandy lake sediments in the north-eastern part of the study area have been analyzed in section P024 (Fig. 4.3.6). An OSL sample was obtained from a depth of 400 cm and was dated to 0.29 ± 0.03 ka (DC08-06). However, the silty sediments have been deposited in an aquatic environment which imposes some uncertainties on the OSL age. To check the influence of a higher water content than the originally calculated content of 2 – 8% (Table 4.3.3), the OSL age was recalculated with a hypothetical water content of 100%. This shifts the OSL age of the sediments to 0.47 ± 0.04 ka. Contrasting results have again been gained from the radiocarbon samples. Radiocarbon ages from Radix shells in P024 gave an age of 2395 ± 86 cal. yrs BP at a depth of 125 cm (P24-M1) and 1414 ± 57 cal. yrs BP at 345 cm (P24-M2) (Table 4.3.4). The inverse radiocarbon results are best explained by a remobilization of older material but do not give the true age of the archive. A fossil Radix shell collected at the surface at P023 in the area of the sandy lake sediments about 2 km north of P024 was radiocarbon dated to 2786 ± 40 cal. yrs BP (P23-M1) (Table 4.3.4). The problems with older radiocarbon ages in comparison to the OSL ages, uncertainties in the reservoir effect, e.g. by changes in the lake size, and the age inversions cannot be resolved here. A recent summary of lake reservoir effects of lakes on the Tibetan Plateau with a special focus on the Donggi Cona has been presented by Mischke et al. (2013). Due to these uncertainties, radiocarbon dating will not be considered further in this paper. According to the stratigraphical position the sandy lake sediments are at least younger than the clayey lake sediments and, according to the OSL sample, younger than 0.5 ka.

Silty sediments covering the gravels of the T2 terrace have been dated by OSL to an age of 7.06 ± 0.37 ka (DC08-07 in Table 4.3.3, quartz) (Fig. 4.3.6). The sequence of gravels, silty sediments and sandy sediments on top are typical for this terrace level. A silty layer above the clayey lake sediments has been dated at section P021 close to the eastern end of the lake sediments. They are reworked sediments from the eastern part of the fan. The dating resulted in an OSL age of 2.86 ± 0.28 ka (DC08-05, quartz).

Table 4.3.4: Radiocarbon dating results. The ages have been calibrated using OxCal 4.1 (Bronk Ramsey, 2009) with the IntCal09 calibration curve (Reimer et al., 2011).

Section Sample Lab.-No Depth [cm] material Age [14C yrs BP] Age [cal yrs BP]

P023 P23-M1 Erl-12529 0 shell 2661 ± 47 2786 ± 40

P024 P24-M1 Erl-12527 125 shell 2347 ± 43 2395 ± 86

P24-M2 Erl-12528 345 shell 1516 ± 47 1414 ± 57

P101 P101-M1 Erl-14423 260 shell 4246 ± 79 4776 ± 124

P101-M2 Erl-14422 280 shell 3645 ± 85 3976 ± 122

Fig. 4.3.6: Selected sections (from the right: P026: sediments of the T2 terrace; P024: sandy lake sediments, P221: dune sediments on the northern footslope).

Discussion

The results from the study area show a complex pattern of sediment mobilisation and deposition. According to the geomorphological and sedimentological data a tentative reconstruction of the landscape evolution has been developed (Fig. 4.3.7) and is discussed below.

Fluvial gravels form the base of the fan. The existence of the large ice wedge casts within these gravels point to a formation under the cold climate conditions of the Pleistocene. Several dated ice wedge casts have been reported from the regions north of the study area. Porter et al. (2001) supposed a formation of ice wedges during the cold climate conditions of the late Pleistocene in the area of the Qinghai Lake based on a TL date from the fill of an ice-wedge cast. Even older ages from the same area are reported by Madsen et al. (2008) from dating the sandy infilling of the casts using OSL. Owen et al. (2006) used OSL to date an ice-wedge cast in the Qaidam Basin to 15 ka. Similar ages for ice-wedge casts have been obtained by radiocarbon dating in the Hexi Corridor (Nai’ang et al., 2003).

Fig. 4.3.7: Reconstruction of the landscape development at the Donggi Cona. 1. Gravel deposition, formation of ice wedges. 2. Deposition of loess, dunes and lake sediments. 3. Erosion of loess on hill slopes, transport to and deposition on the fan. 4. Incision of T1, erosion of lake sediments, formation of dunes on the lake sediments and the sides of the fan, sandy loess on the hills. 5. Lake sediment deposition, sand erosion, formation of on-fan dunes. 6. Incision of T0 and erosion of the sandy lake sediments.

Glaciers are efficient in providing sediments in this high mountain environment (e.g. Owen et al. (2009) especially during Pleistocene times. Glacial landforms which indicate the existence of valley 85

glaciers are frequent in the catchment of the Dongqu River. Owen et al. (2003) reported four terminal moraines of late Quaternary age in the nearby A’nyêmaqên Shan. The formation of the terraces at the Dongqu River might be related to the three older glacial stages and the subsequent meltwater pulse of the A’nyêmaqên Shan. The eastern part of the basin was filled by a large amount of fluvial gravels which reached down to the lake shore (Fig. 4.3.7-1). Detrital inflow in the lake was high during the late glacial (Opitz et al., 2012). The fluvial gravels in the Donggi basin are related to the formation and the erosion of the fluvial terraces of the Dongqu River. The lake level was about 24 m lower than the present lake level (pll) (Dietze et al., 2010, 39 m pll according to Opitz et al., 2012). The active surface covered the whole fan, preventing the continuous accumulation of finer material. This active fluvial phase presumably lasted until the early Holocene. The first accumulation of sediments which have been preserved until today occurred around 7 ka. A minor input to the fringe of the fan was caused by solifluction processes.

At the end of the Pleistocene and in the early Holocene sand accumulation started at the footslopes of the neighboring mountains but not on the alluvial fan (Fig. 4.3.7-2). The oldest OSL age from the dunes on the northern side of the alluvial fan (P221) is 13.9 ± 1.3 ka. The source of the sand was the large alluvial fan (IJmker et al., 2012b). An onset of the sand accumulation around the Pleistocene- Holocene boundary and a phase of enhanced sand deposition from around 10.5 to 7 ka is also documented by OSL ages from other dune fields in the area. A denser vegetation cover due to slightly warmer and wetter climatic conditions might be the cause for this accumulation (Stauch et al., 2012). Similar OSL ages for the deposition of sands have been obtained from the Qinghai Lake (Lu et al., 2011b). Parallel to the phase of enhanced sand deposition in the basins loess was deposited on the slopes and on the tops of the neighboring hills (Stauch et al., 2012) which is typical for north- eastern Tibet (Lehmkuhl, 1997; Owen et al., 2003; Porter et al., 2001; Lu et al., 2011b). Lake levels were rising during the late glacial and early Holocene and reached a maximum of 16.5 m above pll at 9.2 cal. ka BP according to radiocarbon ages from lake sediments above the present lake level (Dietze et al., 2013). However, no clear morphological landforms or sediments have been preserved in the area of the alluvial fan. A second lake level maximum was reached at around 7.5 cal. ka BP. Lake levels were slightly above 10 m above pll (Dietze et al., 2013). These lake levels higher than 10 m above pll correspond to the clayey lake levels on the western part of the alluvial fan.

At around 7 ka a layer of silty sediments was deposited upon the gravels of the alluvial fan (Fig. 4.3.7- 3). This fits well into a phase of intensified erosion of loess deposits in the catchment of the Donggi Cona from around 9 to 6 ka (Stauch et al., 2012). The silty layer on top of the fluvial gravels on the highest terrace (T2) is the reworked loess initially deposited on the slopes, which was eroded during a period of wetter climatic conditions and enhanced runoff. During later times the amount of silt in the system decreased considerably and was not available for fluvial deposition any more. Therefore no loess-like sediments can be found on the lower terraces. Today loess can only be found in geomorphologically stable positions such as flat areas in the mountains. In that position loess deposits are protected against fluvial and aeolian erosion (Stauch et al., 2012). The fact that the silty layer can only be found on the uppermost terrace level (T2) shows that its deposition occurred prior to the first phase of river incision, indicating that the terraces on the alluvial fan are considerably younger than the river terraces of the Dongqu River.

Aeolian sediments on the southern lake shore covering a lake terrace 10 m above pll show that the lake level began falling before 4.5 ka (DC06-07 in Stauch et al. (2012)). The lowering of the lake level might have triggered the first incision into the eastern alluvial fan. Assuming that the T2 level 86

remained active during times when the lake level was at least 10 m higher than at present and incision of the gravels started with or some times after the lowering of the lake level, the lake level drop can be tentatively narrowed down to the time between 7 and 4.5 ka. Changes in the ostracod assemblages at around 6.8 cal. ka BP have been associated with changes from a closed to an open lake system (Mischke et al., 2010a; Aichner et al., 2012) and are in accordance with this interpretation. Lake level lowering might be induced by a shift to cooler and drier conditions which is shown in several palaeoclimatic records from 5 ka onwards (e.g. Herzschuh, 2006; Morrill et al., 2006). Additionally there are indications of tectonic events in the area which might have also resulted in enhanced discharge of the lake. Van Der Woerd et al. (2002) found indication of seismic activity 40 km east of the lake. On the basis of radiocarbon ages they assumed an age of younger than 8.5 cal. ka BP. Guo et al. (2007) found indications of a seismic event between 5 and 7 cal. ka BP directly west of Donggi Cona. The silty sediments on top of the clayey lake sediments in P021 have been dated to 2.86 ± 0.28 ka. They are located at the eastern end of the lake sediments and they might represent a subsequent remobilization of sediments which is not directly related to the lake level lowering.

Low lake levels in the mid- to late Holocene (Fig. 4.4.7-4) provided an additional sediment source. First indications of a reactivation of the aeolian transport are sediments on the north-eastern side of the alluvial fan. Accumulation of sandy loess started at around 2 ka (P238 and P270). Interestingly these sediments have higher carbonate content than the older aeolian sediments. (IJmker et al., 2012b) supposed that these sediments originated partly from the former lake bottom on the western part of the alluvial fan. Sand deposition on the sides of the alluvial fan (P030, P101) and above the clayey lake sediments on the western part of the fan (P006) began around 1.34 ka. Increased aridity due to a reduction in the monsoonal strength might have triggered these sand movements from the exposed lake sediments.

From around 0.6 ka (P026), and especially from 0.4 ka onwards, small dunes accumulated in the center and north of the western part of the alluvial fan. The youngest age from dune sand is 0.3 ka (P019). However, as material close to the surface was not sampled for OSL dating, this age does not mark the end of the sand accumulation and continuous accumulation until the present is likely. The OSL sample from the sandy lake sediments in the west of the study area (P024) yielded an age of 0.29 ± 0.4 ka but might be older if a higher water content is assumed. These sandy lake sediments indicate higher lake levels about 6 m above the present level (Fig. 4.3.7-5). Due to the layering of the sediment complex, including the layers of snails, a near shore environment can be assumed for deposition. Most of the clayey sediments have been eroded between these two phases of high lake levels because of the wide areal distribution of the sandy lake sediments and the small area with clayey lake sediments. The younger phase of again higher lake level coincides with a phase of enhanced aeolian processes. Drier and colder environmental conditions during the last 0.7 ka have been reported also from other terrestrial archives in the Donggi Cona catchment, where the dust input was analyzed (IJmker et al., 2012c). The high amount of sand at the lake shore was washed into the lake due to the high availability of sand at the terrestrial surface. Dietze et al. (2013) described lake sediments from the western side of the Donggi Cona at a similar elevation. A following lake level drop has been tentatively aligned to changes in lake chemistry, e.g. the shift from an aragonite- to a calcite-dominated system at 4.3 cal. ka BP mentioned by Opitz et al. (2012). However, due to the weak chronology of the sandy lake sediments (only one OSL age, uncertain water content) further research regarding the timing of the deposition is necessary.

The subsequent lowering of the lake level leads to fast erosion of the sandy lake sediments (Fig. 4.3.7-6). Most of the sand was washed into the lake. The predominant west winds counteract the west directed fluvial transport on the fan, so that lake sediments are blown out and transported onto the alluvial fan, leading again to the formation of aeolian landforms. Additionally, on the northern side of the fan (P221) 120 cm of sand accumulated during the last 300 yrs. The lake level drop might have also triggered the incision to the present river level on the eastern part of the alluvial fan (T0) and the formation of up to four terrace levels in the sandy lake sediments. Terrace formation on this slightly inclined surface is closely related to heavy rainfall events, frequent river bed modifications and short term lake level fluctuations.

The younger dunes are partly on top and sometimes beside the ‘sand wash’ sediments, indicating a coexistence of aeolian accumulation and fluvial remobilization of the sand. Under present day climate conditions aeolian transport is most effective during the dry and cold winter times while fluvial transport occurs during the torrential rainfalls in summer. However, a general trend to more arid conditions is indicated by the formation of new dunes during the last 2 ka. Before that time denser vegetation cover most probably prevented or reduced aeolian sand transport due to more favorable climate conditions for vegetation (Stauch et al., 2012).

According to the presented results a simple model of the sediment pathways on the fan and in its vicinity has been developed (Fig. 4.3.8). The large alluvial fan serves as storage for sediments coming from the surrounding area. Sediments are only partly transported to the deeper part of lake which is the final sediment trap. Fluvial, aeolian, periglacial and limnic processes were and are influencing the area.

One of the main sediment input pathways is the Dongqu River which drains the A’nyêmaqên Mountains. Transported sediments vary from gravels with a diameter of several centimeters to the clay fraction. The accumulation of gravels was especially active during glacial times. Only a minor input of clastic sediments was delivered by solifluction processes from the neighboring slopes.

Fig. 4.3.8: Interaction of processes in the study area (white arrows: aeolian processes; black: fluvial).

Sandy material was subsequently blown out from the fan but continuous deposition started only in the early Holocene at the sides of the fan. During recent times the mobilization of sands from the Dongqu River plays only a minor role in the sediment cascade, due to limited surface area of the river course. In addition, other sediment sources like the dried lake area or sand from neighboring dune fields became available. Throughout the Holocene an internal sediment cycle developed on the alluvial fan and the neighboring slopes. Sand stored on the slopes is transported down to the fan by fluvial processes where it is forming the area of the fluvial sands. This process is active during the high rainfall in summer. The sand is subsequently blown out to the east following the main wind direction and is deposited on other areas of the fan as well as on the neighboring slopes. The small- scale sediment cycle on the alluvial fan itself is also manifested in the smaller mean grain-size of the sand deposits on the fan in comparison to those on the sides of the fan, which show indications of initial sorting processes. Larger grain-sizes are deposited on the lower part of the slopes while finer but well sorted sediments are transported further upwards. During winter large areas of the surface of the alluvial fan are affected by small-scale frost processes providing additional sediments. Those processes lead to the loosening of the surface and allow mobilization of sediments. Another important process during the latest Holocene is the anthropogenic influence in the area. The alluvial fan is used as pasture by the nomads especially during spring and autumn. This also leads to surface degradation and enhances aeolian erosion during cold winter times.

However, some sand is transported into the lake where it is stored over longer time-scales. These sands originate from two different sources: sand from the Dongqu River, which is directly transported into the lake, and sand from the alluvial fan. Especially during the latest lake level high- stand in the late Holocene, large amounts of sand must have been stored in the lake. These covered around 50 km² of the western fan area. The related sand deposits reach a height of more than 4 m. If only half of this area has been covered by sand deposits of that height an overall total of 1x108 m³ of sand was temporarily stored in the area. These sands must have originated mainly on the alluvial fan. During early high-stands of the lake, the amount of sand might have been considerably lower. Serving as a temporary sediment sink during high lake levels the area is nowadays an important sediment source.

In contrast to the sandy sediments silty deposits yield a different set of information. In addition to the short distance sediment input from the alluvial fan, silty sediments have been transported into the catchment by long distance transport during the early Holocene resulting in an additional archive on the surrounding hills. Under wetter climate conditions caused by an enhanced monsoon circulation the loess archives have been largely eroded forming the silty layer on top of the gravels. However, these loess archives had only a limited capability to release sediments because they originally consisted of a relative small amount of material. The end of the fluvial deposition of the silty sediments might not be identical with the end of a phase of wetter climate but is more likely caused by the reduced availability of the parent material. Only after new sediment sources became available, i.e. the former lake floor in the west of the alluvial fan, silty sediments were reactivated and played a minor role in the sediment cycle again. Under present climate conditions erosion of both types of lake sediments is fast. The dominant process is fluvial erosion which is presumably accelerated in the clayey lake sediments by the degradation of the underlying permafrost.

Conclusion

Sediment availability in the study area is an important factor for the formation of the landforms. While some sediments are only temporarily available, others have been active since at least the late Pleistocene. Sandy sediments are the major component in the investigated system. They form a local sediment cycle and are mainly mobilized by aeolian and fluvial processes. These two processes show strong interactions. Other depositional components like the silty sediments are only temporarily available. Here, the major phase of deposition occurred during the early Holocene but most of the sediment archives were subsequently eroded during wetter climate conditions. Silty sediments only became part of the system again with lake level lowering of the Donggi Cona. However, different factors such as tectonic movement or climatic changes might cause this drop. Therefore the terrestrial archives have to be interpreted carefully. Similar implications apply to other sedimentary archives, e.g. lake sediments. Transport into the lake is highly dependent on the availability of the terrestrial sediment sources. This is well documented in the sediments of the late Holocene lake level high-stand. These sediments have a large amount of sand. The sand accumulated during the Holocene at the side of the fan and probably also on the fan itself and has been deposited in the lake in a near-shore environment. While most of the sand has been transported into the lake, a considerable amount was remobilised by aeolian processes during lake level low-stands back onto the fan. This transport led to the formation of small dunes.

The alluvial fan east of the Donggi Cona forms a critical interface in the sediment cascade from the high mountain area to the lake. This interface has to be considered carefully especially in the case of a palaeoclimatic reconstruction. The (terrestrial) geomorphological archives mirror climatic changes if the suitable sediments, e.g. the loess, are available and sensitive to the designated process. Furthermore, some similar archives record different climatic conditions, such as the dune sediments in the Donggi Cona catchment. Dune sediments from the latest Pleistocene and the early Holocene are related to the onset of the monsoonal circulation whilst the late Holocene sediments indicate drier climate conditions. The identification of the sediment pathways enabled reconstruction of the palaeoclimate. However, it is important to note that the terrestrial archives may be biased due to times of reduced sediment availability.

Acknowledgments

We would like to thank Marianne Dohms from the sedimentological laboratory of the RWTH Aachen University for measuring the grain-size properties. Several Chinese and German students helped us in field. Two anonymous reviewers and Andrew Plater are thanked for their useful comments on an earlier version of this manuscript. This research has been done in the frame of the project ‘Landscape and Lake-System Response to Late Quaternary Monsoon Dynamics on the Tibetan Plateau - Northern Transect’ which has been funded by the German Science Foundation (DFG) as part of the SPP 1372: Tibetan Plateau – Formation Climate Ecosystems.

4.4 Landscape and climate during the late Quaternary on the northern Tibetan Plateau

G. Stauch, P. Schulte, A. Ramisch, K. Hartmann, D. Hülle, G. Lockot, B. Diekmann, V. Nottebaum, C. Müller, B. Wünnemann, D. Yan, F. Lehmkuhl

Keywords: geomorphology, OSL dating, paleoclimate, Tibetan Plateau,

Abstract

Palaeoclimate reconstruction on the northern Tibetan Plateau resulted in a large spectrum of different and partly divergent interpretations for the climate evolution during the late glacial and the Holocene. In some cases this is caused by incomplete understanding of the geomorphological processes influencing the different proxies used. To overcome these limitations and to enhance the understanding of the complex process interactions in a sensitive and highly dynamical environment a detailed analysis of different members of the sedimentary system at Lake Heihai on the northern Tibetan Plateau was conducted. Lake level variations during the late Pleistocene were influenced by sediment supply to an alluvial fan. This sediment surplus resulted in the temporary blocking of the outflow of Lake Heihai. High sediment supply presumably occurred during or shortly after large glaciations in the Kunlun Shan. The spatial distribution of aeolian sediments revealed a strong relationship to possible source areas. This resulted in a spatially heterogeneous distribution of the aeolian sediments. Furthermore, topographic effects have an important influence on the preservation of the sediments. Aeolian sediments deposited in sheltered positions might not be comparable with other archives with a similar grain size. Nevertheless, deposition of loess during the mid-Holocene indicates a shift to wetter climate conditions on the northern Tibetan Plateau. This might be caused by the intrusion of the East Asian Summer monsoon into the area. During the late Holocene, the Asian summer monsoon retreated and aeolian sediments were reactivated.

Introduction

Palaeoclimate changes on the northern and north-eastern Tibetan Plateau have been intensively studied in recent decades due to its location in the zone of interactions between the Indian summer monsoon (ISM), the East Asian summer monsoon (EASM), the Asian winter monsoon and the mid- latitude westerlies (An et al., 2012) (Fig. 4.4.1). However, considerable differences in the interpretation of the influence of different atmospheric systems during the Holocene still exist. A large variety of archives has been studied, ranging from lake sediments (Shen et al., 2005; Opitz et al., 2012; Wünnemann et al., 2012; Chen et al., 2016), pollen records (Herzschuh et al., 2006), tree- rings (B. Yang et al., 2014), speleothems (Zhang et al., 2008), ice cores (Thompson et al., 1997) and terrestrial sediments (Qiang et al., 2013a; Stauch, 2015; Yu et al., 2015). Special emphasis was placed on the evolution of the different phases of the Asian Summer monsoon (ASM) in general and on its

two components, i.e. the ISM and EASM (Herzschuh, 2006; Chen et al., 2008; Y. Wang et al., 2010; An et al., 2012; Chen et al., 2016). According to Herzschuh (2006) most parts of the Tibetan Plateau are dominated by the ISM, with a likely northern boundary along the Kunlun Mt. range (Ramisch et al., 2016), while the areas to the east (e.g. eastern margin of the Tibetan Plateau, the Chinese Loess Plateau) are influenced by the EASM. The ISM was strongest in the early Holocene, at around 10.9 to 7.0 ka BP, while the moisture maximum of the EASM occurred from 8.3 to 5.5 ka BP. In contrast to the Asian summer monsoon, the mid-latitude westerlies experienced no pronounced moisture maximum (Herzschuh, 2006). The asynchronous evolution of the ISM and the EASM was later also observed by Y. Wang et al. (2010). In contrast, Chen et al. (2008) proposed a synchronous evolution of both monsoonal subsystems, with a maximum in the early Holocene. Additionally, several different climatic interpretations of the influence of the mid-latitude westerlies exist. While Chen et al. (2008) assumed increasing moisture in the mid-Holocene due to the mid-latitude westerlies north of the Tibetan Plateau; An et al. (2012) proposed reduced precipitation during enhanced mid-latitude westerlies and a weaker ASM. A similar anti-phase relationship was reported in several studies from northern China (e.g. Herzschuh, 2006). In contrast, other studies highlight an influence of the Asian Winter Monsoon on the northern side of the Tibetan Plateau and an anti-correlation with ASM (B. Liu et al., 2014). Studies from Kusai Lake on the northern Tibetan Plateau indicate a synchronous strengthening of the monsoonal system and the Asian winter monsoon (Liu et al., 2009; X. Liu et al., 2014). However, the winter monsoon might just be a low level atmospheric system which does not influence the basins on the Tibetan Plateau (An et al., 2012; Zan et al., 2015). These conflicting palaeoclimate interpretations are at least partly related to difficulties in proxy interpretation and the varying local topographic effects in the different basins on the northern Tibetan Plateau (Stauch, 2015). Many reconstructions are based on lake sediments which are subject to a variety of processes. These influences are partly related to changes in the lake itself or in the catchment (Wischnewski et al., 2011; Yan and Wünnemann, 2014). One of the most widely discussed problems is the lake reservoir effect which significantly influences the radiocarbon chronology. The inflow of old carbon introduced to the lake system can result in an overestimation of the radiocarbon ages (Long et al., 2011; Hou et al., 2012; Mischke et al., 2013; Lockot et al., 2016). To overcome these limitations a better understanding of sediment transport processes is required. Reconstructing the terrestrial sediment cascade in a catchment would yield a better understanding of proxy variability and the involved mechanisms (e.g. Wünnemann et al., 2010).

This study reconstructs relevant geomorphological processes, sediment transport pathways and their chronological framework along the terrestrial sediment cascade represented by the catchment of Lake Heihai in the Kunlun Mountain range at the northern margin of the Tibetan Plateau. A special focus is placed on the analysis of aeolian sediments. Evaluating grain size and geochemical composition allows a spatio-temporal perspective on the climate-sensitive sediment cascade during the late Quaternary. A comparison with other palaeoclimate archives from the wider area is used to discuss the climate evolution.

Regional setting

Lake Heihai (36.00°N, 93.15°E) is located in the Kunlun Shan at an elevation of 4,420 m above sea level (asl) (Fig. 4.4.1). The catchment area is around 1,600 km², while the present day lake area is around 39 km². The water depth had a maximum of 22 m in 2011 (Lockot et al., 2016). The lake has 92

two major inflows which drain the southern glaciated area and one outflow on the eastern side. The outflow is a tributary of the Kunlun River (sometimes also named Golmud R.), a deeply incised river, which ends in the endorheic Qaidam Basin (Fig. 4.4.1).

Lake Heihai is situated in an intramontane basin with a WNW-ESE extent of around 50 km and a width of 15 km (Fig. 4.4.2). The basin is bordered by the main Kunlun Range in the south and the Burhan Buda Mountains in the north with elevation of up to 5700 m and 5400 m asl, respectively. The prominent Kunlun Fault, a large strike-slip fault at the northern side of the Tibetan Plateau, with a length of about 1600 km (Fu and Awata, 2007), is located on the southern side of the Kunlun Range (Fig. 4.4.1). The average slip rate during the Quaternary of the Kunlun Fault is estimated at around 10 mm/yr (Zhang et al., 2004; Li et al., 2005; Fu and Awata, 2007). Earthquakes occur frequently, e.g. at the Kusai Lake south of the Heihai in 2001 with a magnitude of 7.8 (Mw) and a horizontal displacement of 10 m (Fu and Awata, 2007). The earthquake resulted in a 400 km long rupture zone, one of the largest known in Asia (Lin et al., 2002; Li et al., 2005). Van Der Woerd et al. (2002) identified several large-scale offsets in the Kusai area and assumed a mainly post-LGM date of formation. Data from the Kunlun Pass area around 100 km to the east indicate a recurrence interval of large earthquakes of around 1000 yrs (Van Der Woerd et al., 2002), while west of Lake Heihai, at the Honshui Gou, a rate of even 300 ± 50 yrs was assumed for the last 6 ka (Li et al., 2005). However, despite the proximity of the Kunlun fault, no indications of young tectonic activity, like mole tracks or stream offsets were observed in the catchment of Lake Heihai.

Fig. 4.4.1: (left) Overview of the Tibetan Plateau and the main atmospheric systems, the green line represents the present day average limit of the ASM according to P. X. Wang et al. (2014), black box is the study area; (right) the central Kunlun Mountains with main fault lines according to Li et al. (2005), black rectangle denotes the outline of Fig. 4.4.2.

The geology of the Kunlun Mountains south of the lake is dominated by Permian to Triassic slate and schist and small areas of limestones, Oligocene to Miocene sandstones and conglomerates. Jurassic granite-type rocks occur in the south-eastern part of the catchment. The northern side of the lake is composed of Palaeozoic mafic rocks as well as Ordovician to Silurian sandstones and conglomerates. The basins are filled by Quaternary sediments (Kidd et al., 1988).

Fig. 4.4.2: The catchment of Lake Heihai and the described geomorphological landforms. MI to MIV mark terminal moraines, the arrow indicates the actual river flow direction for the SE catchment. The letters A to D are the location of the pictures in Fig. 3. For further explanation see text.

Up to now, geomorphological research in the central Kunlun area concentrated on the Golmud River valley and the Hongshui River (Fig. 4.4.1). Remnants of several Quaternary glaciations have been described for the Xidatan Valley in the upper reaches of the Kunlun River. Owen et al. (2006) dated remnants of three glacial advances in the mountains south of the Xidatan Valley. On the northern side of the mountains they formed during the penultimate or early last glacial cycle, the MIS 2 and in the Holocene. On the southern side the ages indicate a formation during the penultimate glacial cycle, the MIS 3 and sometimes between the LGM and the Holocene. Even older glaciations have been described by Wu et al. (2001) and Liu et al. (2006) with ages up to 700 ka. Four main terraces were identified in the Kunlun River valley (Liu et al., 2006; Chen et al., 2011). The accumulation of the valley fill was dated to 82-16 ka (Chen et al., 2011) or before 30 ka (Owen et al., 2006). Four tectonically controlled incision stages were identified at 16-13 ka, 13-11 ka, 11-5 ka and 5 ka until the present (Chen et al., 2011). In the Xidatan Valley, seven alluvial surfaces were mapped. Six of them have a Holocene age (Van Der Woerd et al., 2002). Haibing et al. (2005) recorded four terraces in the upper part of the Hongshui Valley (Fig. 4.4.1) out of which at least three have Holocene ages. At both study sites the formation of the Holocene terraces was related to Holocene climate changes at the northern side of the Tibetan Plateau (Van Der Woerd et al., 2002; Li et al., 2005).

The modern climate in the area is characterized by an annual precipitation of around 250 mm and high evaporation values (Liu et al., 2009). The region is presently outside of the influence of the ASM (An et al., 2012; P. X. Wang et al., 2014). The mean annual temperature is at around -8 °C (Maussion 94

et al., 2013). Consequently, permafrost features are widespread in the study area. Solifluction lobes are present at the slopes of the mountains on the northern side of the basin. In fine-grained sediments permanently frozen ground is occurring in a depth of around 120 cm. The vegetation in the catchment of the lake consists of dry to alpine steppe communities with distinct differences in the communities between the northern and southern side of the lake. On the northern side, which is dominated by sandy deposits, Kobresia robusta communities are widespread. They are indicative of a relatively dry environment. In contrast, Poa pachyantha communities are the main vegetation type on the southern side of the lake. The sediments are finer and the plant diversity is higher than on the northern side. Furthermore, the vegetation cover is denser on the southern side of the lake (Müller and Kürschner, 2013).

Methods

A multiproxy approach, comprising geomorphological mapping, detailed spatial sedimentological sampling, grain size and geochemical analysis of sediment samples and OSL dating, was applied for the reconstruction of the landscape evolution.

Geomorphological mapping

Geomorphological mapping was based on the analysis of remote sensing data, especially Landsat 8 Operational Land Imager (2013) and Corona (1970) images. Detailed mapping in the study area was realized during field work in summer 2011 and 2012, each lasting for several weeks. Mapping included alluvial fans surrounding the lake, terrace sequences at the outflow and modern and ancient glacial geomorphological landforms. Aeolian processes and archives were recorded in detail to identify aeolian transportation processes in the catchment and past changes in these processes. For the spatial reconstruction of the sediment cascade, surface samples were collected over large parts of the catchment. At selected sections of up to several meters depth, detailed sampling was conducted for the reconstruction of the geomorphological processes through time.

Sedimentology

Two hundred and forty-one individual terrestrial sections were sampled in the area of Lake Heihai, resulting in 560 samples. Out of these, 210 were surface or near surface samples with one or two samples up to a depth of 10 cm. Samples were dried and sieved to remove the fraction larger than 2 mm. They were treated with H2O2 for the removal of organic matter. Several tests showed that the removal of carbonates by HCl has no significant impact on the validity of the measurement and can even result in misleading grain size distributions (Schulte et al., 2016). Na4P2O7 was used for dispersion. For grain size measurement a Beckmann Coulter LS 13320 laser particle sizer with 116 classes ranging from 0.04 µm to 2000 µm was used. Each sample was measured twice with two different concentrations. A frequency distribution was calculated using the Lorenz-Mie theory with a refractive index for water of 1.33, for the sample of 1.55 and an absorption coefficient of 0.1 (Özer et al., 2010).

For the classification of the grain size distributions of the aeolian sediments, the sediments were decomposed into a set of end-members (see Weltje, 1997). We used the end-member modelling 95

routine introduced by Dietze et al. (2012) to estimate end-member loadings and sample scores from the previously weight transformed grain size distributions. The aeolian samples were afterwards classified according to their scores.

Geochemistry

To determine the element concentration of the fine-grained fractions, samples were sieved to the <63 µm fraction and dried at 105°C for 12 h. 8 g of the sieved material was mixed with 2 g Fluxana Cereox, homogenized and pressed to a pellet with a pressure of 20 tons for 120 s. All samples were measured twice with the XRF device (Spektro Xepos, 29 elements, cf. loading matrix). Mean values were calculated from the two measurements.

Factor analysis was applied to reduce the dimension of the geochemical dataset subdivided into fluvial (n=48), lacustrine (n=11) and aeolian sediments (n=258). Considering the number of elements, only the aeolian subset provides a reliable relation for derivation of the Eigenspace. In order to solve the problem of constant sum constraints in compositional data, we transformed the data matrix by isometric log-ratio transformation (ilr) according to Filzmoser et al. (2009).

Dating

OSL dating was done at the OSL laboratory of the University of Cologne. Fifty samples were prepared following standard procedures: After drying and sieving, all samples were treated with hydrochloric acid, sodium oxalate and hydrogen peroxide in order to remove carbonates, clay and organic material. To separate the quartz and the K-rich feldspar fraction for coarse-grain dating (100-200 µm), solutions of sodium polytungstate (2.68, 2.62, and 2.58 g cm-3) were used (e.g. Preusser et al., 2008; Hülle et al., 2010). The feldspar fraction was used because the quartz fraction was not suitable for dating: Stimulating each coarse-grained quartz sample with infrared light (cf Duller, 2003) demonstrated for a low signal and a feldspar contamination of the samples, inhibiting accurate palaeodose determinations. This is in accordance to luminescence dating studies in Central Asia showing that many samples exhibit feldspar contamination of the luminescence signal (Hülle et al., 2010; Lehmkuhl et al., 2011).

All luminescence measurements were carried out on automated Risø TL/OSL readers (TL-DA-15 or 20), which were equipped with 90Sr/90Y β-sources (dose rate of 0.11-0.15 Gy/s) for irradiation and with EMI 9235 photomultiplier tubes for luminescence detection. Based on the results of preheat tests, 230 °C for preheating of the regeneration dose as well as the test dose of the feldspar samples was chosen. For each feldspar sample 12 to 49 aliquots with 1 mm diameter (≈50-100 grains, Duller, 2008) were sufficient, depending on the data scatter.

A main problem of feldspar dating is anomalous fading. Many dating studies using feldspars showed that this signal loss due to leakage of electrons at room temperature can cause a severe underestimation of the true burial age of feldspars. Therefore, fading rates (“g-value”) were determined following Auclair et al. (2003). To calculate fading corrected De values the procedure proposed by Huntley and Lamothe (2001) was used.

For some samples, the postIR-IR180°C-protocol by Reimann et al. (2011) proved to be a more suitable procedure to meet the required quality criteria, using a preheat at 200 °C, an IR bleach at 50 °C and IR-stimulation at an elevated temperature of 180°C (followed by an IR-hotbleach at 280 °C only after the test–dose cycle). This post-IR IRSL signal using elevated stimulation temperatures has high potential to derive a dating result that is far more independent from fading-correction than the conventional IRSL50°C-signal. For this protocol, the applied dose could be recovered within ±3% and the general quality criteria (recycling ratio limit of ±10%, recuperation <5%, test dose error <10%) further indicate applicability of the protocol.

The environmental dose-rate is created by the radioactive elements existing in grains of the sample and surrounding sediment, with a contribution from cosmic rays. For the HEH-samples, laboratory high resolution gamma-spectrometry was conducted to obtain the contribution from uranium (U) and thorium (Th) decay chains and from potassium (K).

All measurements were converted to alpha, beta and gamma dose rates according to the conversion factors of Aitken (1998). For K-feldspars an internal potassium content of 12.5  0.5 (Huntley and Baril, 1997) and an α-efficiency factor of 0.070.02 (Preusser, 2003) was assumed. The dose rate from cosmic rays was calculated on the basis of sample burial depth and the altitude of the section (Prescott and Hutton, 1994). The dose rate was calculated taking variations in the water content of 5 to 15 wt-%, depending on the actual water content and the porosity, into account.

Results

Geomorphology

Only the high mountain area of the Kunlun Shan south of the lake is presently glaciated (Fig. 4.4.2). Small ice caps and valley glaciers cover an area of around 62 km² (western area: 38 km²; eastern area: 24 km²) in 2014. Former glaciations have been much larger as indicated by frequently observed glacial landforms, like moraines and cirques in the mountain ranges. About 2.5 km south-east of the present lake shore a large terminal moraine (M IV) has been deposited (Fig. 4.4.2). Similar landforms were mapped on the south-western alluvial fan 25 km west of the lake. Most parts of these moraines were destroyed by stream flow across the fan or buried by alluvial fan deposits. Lateral moraines are weakly preserved on the slopes of the south-eastern valley. According to their topographic position they appear to have formed simultaneously with the terminal moraine close to the lake. During this maximum glaciation, a large ice stream with a length of > 30km filled most parts of the valley. Additionally, glacial sediments are located on the upper part of the alluvial fan close to the mountain front (M III). Another set of terminal moraines is located a few tens of meters downstream of the present glacier margin (M II). The appearance of these moraines is still fresh but moraine crests are already collapsing and initial vegetation is appearing. Several fresh-looking moraines (M I) were deposited close to the present glacier margin (Fig. 4.4.3A). A comparison with Corona images from the year 1970 indicated a glacier retreat during the last decades of several tens to hundreds of meters.

The lower part of the catchment of Lake Heihai is dominated by large alluvial fans. The two most prominent ones originate from the main Kunlun Range south-east and south-west of the lake. Smaller active fans are also located on the southern side of the lake while alluvial fans on the 97

northern side are inactive and covered by aeolian deposits. The alluvial processes filled large parts of the Heihai basin with clastic sediments from the surrounding mountains. The two largest alluvial fans are connected to the glaciated mountain area.

The alluvial fan in the south-east of the catchment is an important feature for the formation of Lake Heihai. The propagation of the fan to the north resulted in the damming of the river and the development of the lake. Sediment supply into the basin was presumably high during glacial times, when a large piedmont-like glacier was developed (moraine M IV in Fig. 4.4.2). While at present the main outflow from the south-eastern valley is following the eastern margin of the alluvial fan (arrow in Fig. 4.4.2), during previous times, this river delivered large amounts of sediment into the basin of Lake Heihai. A hydrological connection between the glaciers and the lake would have increased the catchment area by around 200 km². This would have greatly altered the hydrological budget of the lake, as it would include the presently glaciated area. Similarly, the small lake west of Lake Heihai is presently disconnected from main lake by a small barrier of 10 m elevation difference. A higher lake level of the western lake would result in an overflow of the barrier and additionally increase the catchment area by 230 km². In this highly geomorphological active region, variations in catchment areas are an important aspect for variations in the sediment cascade and the hydrological budget of the lake. However, at present it is not possible to infer the timing of the last variations in the catchment area.

Hydrological variations are also recorded by fluvial terraces at the outflow east of the lake. The alluvial fan deposits are incised several meters by the outflow river. In the incised valley, a sediment infilling is preserved which was later incised again. While the formation of the alluvial fan deposits is related to the hydrological variation in the mountains at the southern border of the Heihai catchment, the incision and the formation of the fill terrace was caused by variations of the lake discharge.

Ancient lake sediments are located around the modern lake (Figs. 4.4.2, 4.4.3B) and cover an area of 28 km². The ancient lake sediments are uplifted up to 6m above the present lake level due to the formation of segregated ice. Segregation ice in form of ice lenses between layered lacustrine deposits can be observed around the lake basin. The uplift of the ice resulted in the formation of large local bulges with a diameter of several tens of meters. The eastern inflow of the lake is diverted to the north and south due to these uplifted lake sediments. The exposed sediments are subsequently eroded by fluvial processes and transported into the lake. Additionally, deflation leads to the relocation of the material into the lake and on the hills east of the lake.

Sand sheets and dunes are distributed on the northern and north-eastern side of the lake only. Sand sheets can reach a thickness of several meters. A longitudinal dune of around 10 to 15 m height and a length of 2.5 km formed behind a hill north of the lake (Fig. 4.4.3B). Additionally, climbing sand sheets are present in east of this dune in elevations up to 4800 m asl. Barchans are located in the neighbouring valley further to the east (Figs. 4.4.2, 4.4.3C). All these landforms indicate a predominant wind direction from WNW to ESE. However, during summer times observed modifications of the crest line of the barchans indicate a strong wind component from the east. The longitudinal dune and the barchans are highly active. While the longitudinal dune crest moved around 50 m in a northward direction in the past 30 years, the barchan crests moved between 40 and 80 m in an eastward direction, following the main wind direction based on the evaluation of the remote sensing images. 98

Fig. 4.4.3: A: Late Holocene terminal moraines (M I and II); B: Longitudinal dune and exposed lake sediments; C: Barchan dunes NE of the lake; D: dust entrainment across the SE alluvial fan (see Fig. 2 for location).

Grain size

Aeolian sediments are widespread in the catchment of Lake Heihai and form a partly continuous cover independent of the local topography. 284 samples from 100 sections were analysed from aeolian deposits in the whole catchment. The aeolian samples were classified with an end-member analysis (EMMA) according to their sample scores. Four robust end-members explained a total of 83% considering variations within the grain size spectrum and an explained variance of 84% in between samples. Including more robust end-members did not alter the explained variance considerably and leads to spurious overlap of several sub-modes in between the robust end-member loadings.

Ninety four aeolian samples revealed a score of > 50% to EM1 (Fig. 4.4.4A). The grain size distributions of these samples have modal values between 40 and 90 µm and, additionally, a considerable amount between 3 and 12 µm. The sample with the finest mode (40 µm, HH053) has been taken from a glacier surface in the south-eastern tributary in an elevation of 5029 m asl and can be regarded as an undisturbed reference sample for the present day aeolian dust transport. According to the grain size distribution and the classification during sampling the class can be termed loess as all samples are dominated by silt-sized particles.

A second class of 77 samples has an EM 2 score larger than 50%. Samples in this class have modal values between 90 and 190 µm in the fine sand fraction. Additional peaks occur in some samples in the fine silt and in the sand fraction. The sand fraction is generally larger than 50%. This class was named fine sand. The last two classes belong to the medium and coarse sand fraction. The medium 99

sand (Fig. 4.4.4C) includes 99 samples with a sample score of EM3 larger than 50%. Modal values are between 160 and 500 µm and the sand content is between 88 and 99%. The coarse sand class is comparably small with only 12 aeolian samples included based on their score of EM4. These samples have high modal values between 400 and 1140 µm. The four different groups reflect not only different transportation processes but also different local environmental settings, like sediment availability and topographic effects.

Fig. 4.4.4: Aeolian sediment classes (A: loess, B: fine sand, C: medium sand, D: coarse sand). The black curves are the respective end-members (EM1-EM4). Please note the different scaling of frequency in A.

Geochemistry

For the identification of the source areas the geochemical composition of different sediments was estimated. This included not only the aeolian sediments described above but also, fluvial and lacustrine sediments. Considering the number of elements, only the aeolian subset provides a good relation for derivation of the Eigenspace.

After iterative derivation of varimax rotated factors, five factors provide an eigenvalue of at least 1.6 (minimum normalized variance, cf. Kaiser, 1958). As expected from aeolian sediments, Si does not provide any correlation with elements related to feldspar, clay minerals or carbonate/sulphates/chlorides. Also, MnO, K2O, Rb2O, Ba, Nd show a uniqueness of at least 1/3 of their variance (Tab 4.4.1).

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Tab. 4.4.1: Results of the Factor analysis. Factor loadings >71% and uniqueness <30 are shown in bold.

Factor1 Factor2 Factor3 Factor4 Factor5 uniqueness

Na2O 0.00 -0.75 0.11 0.19 0.09 Na2O 0.36 MgO 0.47 0.35 0.45 -0.02 0.23 MgO 0.37

Al2O3 0.15 0.69 0.12 0.37 0.30 Al2O3 0.24

SiO2 -0.08 0.56 0.62 0.07 0.15 SiO2 0.25

P2O5 -0.72 -0.41 -0.24 -0.18 -0.10 P2O5 0.21

SO3 -0.62 -0.18 -0.39 -0.24 -0.30 SO3 0.27 Cl 0.68 0.33 0.57 0.12 0.25 Cl 0.04

K2O -0.03 -0.27 0.14 0.83 0.16 K2O 0.18 CaO 0.85 0.07 0.39 0.16 0.26 CaO 0.04

TiO2 0.75 -0.19 0.04 0.14 0.59 TiO2 0.03

V2O5 -0.06 0.11 0.01 0.06 0.96 V2O5 0.07

Cr2O3 0.83 0.17 0.22 0.17 -0.12 Cr2O3 0.18 MnO 0.97 0.18 0.02 0.11 0.01 MnO 0.02

Fe2O3 0.91 -0.18 0.16 0.05 -0.12 Fe2O3 0.10 NiO 0.97 0.09 0.02 0.12 0.00 NiO 0.03 CuO 0.88 0.29 0.31 -0.10 -0.12 CuO 0.03 ZnO 0.58 0.49 0.58 -0.08 -0.13 ZnO 0.07 Ga 0.77 0.31 0.11 0.00 -0.06 Ga 0.29

As2O3 -0.31 -0.49 -0.24 -0.01 -0.17 As2O3 0.54 Br 0.19 0.66 0.67 -0.08 -0.05 Br 0.08

Rb2O 0.13 0.13 -0.15 0.75 -0.09 Rb2O 0.36 SrO 0.73 0.22 0.41 0.22 -0.03 SrO 0.20 Y 0.21 0.58 0.26 0.53 -0.02 Y 0.26

ZrO2 0.15 0.03 0.80 0.06 -0.05 ZrO2 0.31

Nb2O5 0.34 0.34 0.26 0.62 0.10 Nb2O5 0.29 Ba 0.23 -0.08 0.60 0.22 0.02 Ba 0.50 Nd 0.23 0.46 0.19 0.22 -0.30 Nd 0.53 Hf 0.52 0.67 0.31 0.21 0.11 Hf 0.13 PbO 0.40 0.32 0.75 -0.08 0.00 PbO 0.16

SS 9.417 4.399 4.38 2.529 1.909 Proportion 0.325 0.152 0.151 0.087 0.066 Cumulative 0.325 0.476 0.627 0.715 0.78

With an eigenvalue of 10.7 the 1st factor reflects 47% of the variance represented by the considered five factors. The factor loadings within this dominating 1st factor show positive correlations with at least 50% declared variance (loadings > 0.71) concerning Cl, CaO, Cr2O2, MnO, Fe2O3, NiO, CuO, ZnO, Ga, Br, SrO, PbO. Here, we cannot interpret the remaining factors either in term of processes (e.g.

101

soil development and weathering) or provenance information. In order to estimate the spatial variances within the aeolian deposits, the scores of factor 1 in the ilr-space were estimated (Fig. 4.4.6).

Spatial sediment distribution

The spatial distribution of the aeolian sediments is distinct between the northern and southern side of the basin. Loess sediments (EM1) are only deposited south and east of the lake (Fig. 4.4.5). The thickness of the sediments ranges from a few centimeters to around 130 cm. The north is dominated by medium (EM3) and coarse aeolian sand (EM4), while no loess was recorded on that side. The sand deposits can reach a thickness of several meters and are draped across the overall landscape. In contrast, in the south only three sections on river terraces exhibited fine to medium aeolian sands. Fine sands (EM2) are particularly abundant east of the lake in the transition area between the loess and the medium to coarse sands. On the northern side of the basin all fine aeolian sands occurred in the lee of topographic obstacles. The occurrence of the northern medium to coarse sands indicate areas of much higher aeolian transport capacity or excessive sand supply in contrast to the southern part. Local topography supports the formation of longitudinal dunes and barchans. Additionally, the sand influences the local vegetation, due to a high infiltration rate, which limits water availability, resulting in self amplifying positive feedback cycle. An intermediate depositional environment is represented by the deposition of fine sand. Interestingly, most of the sections consist of one sediment type only. Loess sediments are covered by fine sands at only two locations (Fig. 4.4.7).

Fig. 4.4.5: Spatial distribution of the classified aeolian sediments (uppermost/surface sample).

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Similar to the spatial sediment distribution, the factor analysis showed geochemical differences between different regions around the lake (Fig. 4.4.6). High factor scores of factor 1 from the aeolian sediments occur in the sands north of the lake, while the loess south of the lake has slightly negative factor scores. Especially low factor scores occur only east of the lake. These geochemical differences are related to the different source areas of the sediments. The source for the sediments with the high factor loadings is located in the Ordovician to Silurian sandstones north of the lake. The sandstone outcrops are surrounded by moving sand, indicating a direct source of the sand from the weathered sandstone. Slightly negative scores are caused by the limestone south of the lake. Fine grained sediments are entrained by the wind from the surfaces of the alluvial fans. The low factor scores in the aeolian sediments east of the lake indicate a possible source from the exposed lake sediments.

Fig. 4.4.6: Factor scores of factor 1 of the aeolian sediments (see Tab. 4.4.1).

Dating

To decipher the landscape evolution of the Heihai area through time 50 OSL ages were obtained within the frame of this study (Fig. 4.4.7, Tab. 4.4.2). 43 OSL ages were taken from aeolian sediments, two from aeolian sediments below lake sediments and five date fluvial deposits. IRSL-measurements of the feldspar fraction met the required quality criteria. For all aliquots the regenerative growth curves show that (1) the recycling ratio is consistent with unity ±10%; (2) the OSL signal is not saturated at the level of the natural signal and (3) recuperation is lower than 5% for all samples, which indicates insignificant charge transfer during the measurements. These favorable luminescence characteristics indicate that credible equivalent dose values can be determined by the SAR protocol. 103

Table 4.4.2: Site information of the OSL dated sections.

Altitude Sediment No. OSL Section Northing Easting [m a.s.l.] classification samples

HH008 35.87835 93.19092 4725 Loess 2

HH012 35.87835 93.19092 4725 Loess 3

HH025 35.89476 93.26569 4574 Fine sand 4

HH044 35.98428 93.32642 4494 Fine sand / loess 5

HH046 35.96648 93.33446 4575 Loess 2

HH047 35.96592 93.33371 4558 Loess 1

HH054 36.01768 93.25953 4452 Sand / lacustrine 1

HH064 35.89443 93.33679 4501 Fine sand / loess 4

HH080 35.97146 93.37085 4458 Loess 2

HH081 36.02998 93.27098 4478 Sand 3

HH087 36.05121 93.19471 4456 Sand / lacustrine 6

HH108 36.00021 93.32070 4546 Sand 8

HH139 35.97093 93.24333 4439 Sand / lacustrine 1

HH143 35.95212 93.21556 4481 Loess 1

HH223 36.03270 93.30545 4609 Sand 2

HH224 36.03423 93.30759 4604 Sand 2

HH233 35.93465 93.36990 4425 Sand 1

HH240 35.94888 93.35461 4425 Sand 1

HH241 35.94888 93.35461 4435 Fluvial sand 1

Most of the measurements of 12-49 aliquots with 1mm diameter showed acceptable relative standard deviations <20% of the individual equivalent doses. Therefore, the mean De was calculated with the Common Age model according to (Galbraith et al., 1999). Nine samples (HEH 5-2, HEH 5-4, HEH 6-3, HEH 8, HEH 14-2, HEH 14-4, HEH 15-3, HEH 15-8, HEH 16-1) showed standard deviations >20%. But, as there are no further indicators for incomplete bleaching these ages were also calculated with the Central Age Model. All De values are given in Table 4.4.3.

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Fig. 4.4.7: OSL sample sites and sediment classification.

Most of the aeolian OSL samples have a Holocene age (n: 36) and only two ages are older than the global last glacial maximum (gLGM). The oldest aeolian sediments have been preserved at the north- western side of the lake. In section HH087 a natural cliff provided a 5 m section (Fig. 4.4.8). Dune sediments are located in a depth of around 3 m below the present surface. Ages of 97.5 ± 7.65 ka (HEH 14-1) and 83.1 ± 7.65 ka (HEH 14-2) were obtained. The dune sediments are overlain by layered fluvio-lacustrine deposits with many small pebbles. They indicate the inundation of aeolian deposits by fluvio-lacustrine sediments. The lower pale part of these fluvio-lacustrine deposits has an OSL age of 52 ± 2.8 ka (HEH 14-3) while the upper reddish part has an OSL age of 13.0 ± 1.14 ka (HEH 14-4). The hiatus thus represents a sedimentation gap of up to 43 ka. Dune sands at around 180 cm were dated to 12.3 ± 0.6 ka (HEH 14-5). The dune sands are again covered by fluvio-lacustrine deposits with an age of 11.6 ± 1.09 ka (HEH 14-6).

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Table 4.4.3: OSL dating results.

fading Lab.- Th K uncorr. Lum. n n g- Section Sample ID depth De error U U err Th K err Dr error error od corr. err code err (%) age technique meas used value age cm Gy Gy ppm ppm ppm ppm % % Gy/ka Gy/ka ka ka % % ka ka

HH008 HEH 1-1 C-L9743 60 20.8 1.2 3.12 0.14 10.70 0.54 1.89 0.03 3.67 0.39 5.7 0.5 pIR180 24 23 10.6 n.a. n.a. n.a. HEH 1-2 C-L9742 45 18.3 1.0 2.99 0.13 10.54 0.53 1.79 0.03 4.05 0.41 4.5 0.4 pIR180 24 24 6.9 n.a. n.a. n.a.

HH012 HEH 2-1 C-L3417 200 23.2 1.2 2.79 0.15 9.97 0.58 1.83 0.02 3.92 0.40 5.9 0.5 IRSL 20 20 4.5 3.2 7.5 0.9 HEH 2-2 C-L9758 160 17.7 0.9 2.79 0.15 9.97 0.58 1.83 0.02 3.94 0.40 4.5 0.4 IRSL 24 24 5.3 3.2 5.9 0.5

HEH 2-3 C-L9755 100 10.9 0.6 2.83 0.12 10.86 0.51 1.88 0.03 4.02 0.42 2.7 0.2 IRSL 47 25 8.4 3.2 3.6 0.3

HH025 HEH 5-1 C-L3423 380 10.7 0.6 2.12 0.09 8.75 0.42 1.68 0.03 3.49 0.37 3.1 0.3 pIR180 39 12 9.0 n.a. n.a. n.a. HEH 5-2 C-L9798 320 11.5 0.7 2.07 0.11 8.67 0.51 1.69 0.02 3.51 0.39 3.3 0.3 IRSL 30 20 16.2 4.3 0.4

HEH 5-3 C-L9794 240 10.7 0.6 2.10 0.12 8.61 0.53 1.74 0.02 3.58 0.38 3.0 0.3 pIR180 24 20 10.1 n.a. n.a. n.a.

HEH 5-4 C-L3420 135 8.2 0.6 2.19 0.12 8.86 0.52 1.66 0.02 3.58 0.45 2.3 0.2 pIR180 45 31 29.9 n.a. n.a. n.a.

HH044 HEH 6-1 C-L3428 340 20.1 1.1 2.46 0.16 9.04 0.64 1.69 0.09 3.60 0.29 5.6 0.3 IRSL 22 21 6.0 6.7 0.4 HEH 6-2 C-L3427 280 1.8 0.1 1.78 0.12 6.74 0.48 1.31 0.07 2.97 0.25 0.6 0.1 IRSL 15 14 11.0 0.7 0.1

HEH 6-3 C-L9835 220 1.3 0.2 1.74 0.09 7.16 0.43 1.32 0.01 3.02 0.57 0.4 0.1 pIR180 49 7 29.4 n.a. n.a. n.a.

HEH 6-4 C-L9833 180 1.0 0.1 1.55 0.08 6.19 0.37 1.19 0.01 2.83 0.40 0.4 0.0 pIR180 24 2 - n.a. n.a. n.a.

HEH 6-5 C-L3424 100 0.0 0.0 1.69 0.11 6.68 0.48 1.21 0.07 2.92 0.24 0.0 0.0 IRSL 2 2 0.0 0.0

HH046 HEH 7-1 C-L3431 140 17.6 0.9 2.62 0.14 10.14 0.60 1.77 0.02 3.87 0.39 4.5 0.4 IRSL 24 24 6.1 6.0 0.5 HEH 7-2 C-L9852 100 11.0 0.6 2.52 0.13 9.94 0.58 1.76 0.02 3.85 0.39 2.9 0.3 pIR180 40 21 6.6 n.a. n.a. n.a.

HH047 HEH 8 C-L3432 120 6.4 0.5 2.22 0.12 8.79 0.52 1.70 0.02 3.63 0.45 1.8 0.2 pIR180 48 28 27.4 n.a. n.a. n.a. HH054 HEH 9 C-L3433 121 1.3 0.1 1.86 0.10 7.62 0.43 1.61 0.06 3.37 0.25 0.4 0.0 IRSL 23 23 12.0 0.4 0.0 HH064 HEH 11-1 C-L3439 220 23.3 1.2 3.07 0.15 10.08 0.58 1.86 0.07 4.00 0.30 5.8 0.3 IRSL 24 22 3.0 2.0 6.7 0.4 HEH 11-2 C-L9897 160 17.0 0.9 2.82 0.15 9.77 0.59 1.75 0.02 3.79 0.39 4.5 0.4 IRSL 32 22 4.6 5.2 0.6

HEH 11-3 C-L3437 120 11.8 0.6 2.71 0.18 10.19 0.72 1.71 0.10 3.83 0.32 3.1 0.2 IRSL 24 22 6.0 3.6 0.2

HEH 11-4 C-L3436 93 3.8 0.2 2.27 0.15 8.34 0.59 1.59 0.09 3.51 0.29 1.1 0.1 IRSL 22 20 13.0 1.2 0.1

HH080 HEH 12-1 C-L3441 105 19.4 1.0 2.92 0.16 10.27 0.60 1.79 0.02 3.98 0.40 4.9 0.4 IRSL 24 24 3.7 6.0 0.5 HEH 12-2 C-L3440 60 13.6 0.7 2.60 0.14 9.71 0.57 1.67 0.02 3.75 0.38 3.6 0.3 IRSL 22 21 3.9 4.5 0.4

HH081 HEH 13-1 C-L3444 250 0.0 0.0 1.34 0.07 5.68 0.30 1.21 0.02 0.0 0.0 IRSL 4 4 HEH 13-2 C-L9942 200 0.0 0.0 1.45 0.08 6.17 0.37 1.30 0.01 0.0 0.0 IRSL 4 4

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HEH 13-3 C-L3442 150 0.0 0.0 1.49 0.08 6.23 0.37 1.29 0.01 0.0 0.0 IRSL 4 4

HH087 HEH 14-1 C-L3450 380 307.5 19.2 1.75 0.09 5.74 0.33 1.52 0.06 3.15 0.23 97.5 6.1 pIR180 12 12 12.0 n.a. n.a. n.a. HEH 14-2 C-L9972 320 274.4 17.5 1.30 0.07 5.24 0.32 1.81 0.02 3.30 0.37 83.1 7.7 pIR180 24 24 19.0 n.a. n.a. n.a.

HEH14-3 C-L3448 300 102.1 5.5 1.49 0.10 5.15 0.37 1.02 0.04 2.55 0.20 40.1 2.2 IRSL 24 24 9.0 52.1 2.8

HEH 14-4 C-L9965 250 34.0 1.8 1.49 0.08 5.54 0.35 1.06 0.01 2.62 0.27 13.0 1.1 pIR180 18 18 8.2 n.a. n.a. n.a.

HEH 14-5 C-L9961 180 33.0 2.2 1.15 0.07 4.74 0.29 1.43 0.02 2.84 0.36 11.6 1.1 pIR180 24 21 19.9 n.a. n.a. n.a.

HEH 14-6 C-L3445 150 24.4 1.3 1.42 0.07 5.18 0.30 1.02 0.04 2.58 0.19 9.4 0.5 IRSL 22 22 5.0 3.2 12.3 0.6

HH108 HEH 15-1 C-L3451 70 22.2 1.2 1.31 0.06 5.05 0.25 1.00 0.02 2.55 0.18 8.7 0.5 IRSL 22 22 6.0 3.2 11.5 0.6 HEH 15-2 C-L1099 125 26.4 1.5 1.39 0.08 5.36 0.32 1.22 0.01 2.78 0.29 9.5 0.8 IRSL 32 32 12.3 12.5 1.1

HEH 15-3 C-L1103 195 26.8 1.9 1.51 0.08 5.78 0.35 1.11 0.01 2.71 0.32 9.9 1.9 IRSL 24 21 21.5 13.1 2.4

HEH 15-4 C-L11107 270 24.0 1.5 1.33 0.07 5.63 0.36 1.20 0.01 2.71 0.30 8.9 0.8 IRSL 24 14 13.8 11.7 1.1

HEH 15-5 C-L1110 320 23.5 1.5 1.43 0.08 5.81 0.37 1.10 0.01 2.65 0.31 8.9 0.8 IRSL 24 18 17.8 11.7 1.1

HEH 15-6 C-L1113 380 24.0 1.5 1.39 0.08 5.63 0.36 1.27 0.01 2.75 0.31 8.7 0.8 IRSL 24 14 13.8 11.5 1.1

HEH 15-7 C-L1115 440 22.4 1.3 1.39 0.08 5.66 0.36 1.26 0.01 2.73 0.29 8.2 0.7 IRSL 24 24 13.3 10.8 1.0

HEH 15-8 C-L3458 560 1.4 0.1 1.24 0.06 4.81 0.24 1.08 0.02 2.45 0.16 0.6 0.0 IRSL 23 17 26.0 3.2 0.8 0.0

HH139 HEH 16-1 C-L3459 110 8.3 0.6 2.20 0.11 9.03 0.52 1.99 0.08 3.88 0.29 2.1 0.2 IRSL 24 24 27.0 2.7 0.2 HH143 HEH 19 C-L3461 55 26.3 1.4 2.79 0.15 10.61 0.62 1.81 0.07 3.99 0.30 6.6 0.3 IRSL 24 24 8.0 7.9 0.4 HH223 HEH 20-1 C-L3462 70 7.9 0.5 1.44 0.10 5.83 0.42 1.21 0.07 2.83 0.33 2.8 0.3 IRSL 24 19 14.3 3.4 0.3 HEH 20-2 C-L3463 210 7.2 0.4 1.46 0.08 5.84 0.33 1.25 0.05 2.81 0.31 2.5 0.2 IRSL 24 17 10.7 2.5 2.9 0.3

HH224 HEH 21-1 C-L3464 50 7.5 0.4 1.63 0.08 6.09 0.35 1.34 0.05 3.01 0.22 2.5 0.2 IRSL 24 21 7.0 3.0 0.3 HEH 21-2 C-L3465 110 7.8 0.4 1.62 0.08 6.07 0.35 1.34 0.05 2.98 0.22 2.6 0.1 IRSL 22 22 7.0 3.2 0.2

HH233 HEH 22 C-L3466 80 18.7 1.0 2.42 0.13 8.81 0.52 1.75 0.02 3.73 0.26 5.0 0.3 IRSL 24 24 5.0 2.3 6.0 0.6 HH240 HEH 23 C-L3467 90 20.1 1.1 2.20 0.10 7.41 0.38 1.57 0.03 3.41 0.34 5.9 0.5 IRSL 24 23 7.1 7.3 0.6 HH241 HEH 24 C-L3468 100 16.5 0.9 2.47 0.11 9.30 0.47 1.76 0.03 3.78 0.26 4.4 0.2 IRSL 24 23 7.0 5.3 0.3

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Fig. 4.4.8: Section HH087, Stratigraphy and OSL ages.

Aeolian sands with a late glacial to Holocene age (Fig. 4.4.9) were also recorded in a river cut on the north-eastern side of the lake (HH108). In this 580 cm long section eight samples were taken for OSL age determination. The seven upper OSL ages (HEH 15-1 to HEH 15-7) range from 10.8 ± 1.0 ka to 13.1 ± 2.4 ka. However, there is no clear age depth relation. We therefore assume a roughly late glacial/ early Holocene deposition age. This section is located in a small valley and the aeolian sediments are sheltered against aeolian deflation, which explains the relatively old ages. At the base of the section an age of 0.8 ± 0.0 ka (HEH 15-8) was determined. This young age is most probably related to a disturbance in the section, either due to bioturbation or to a not recognized undercut by the nearby episodic river. This age is not included in any further analysis. In general, all the OSL ages from aeolian sediments mentioned above were taken from coarse or medium sand deposits.

Mid Holocene ages of aeolian sands were only obtained at two sections east of the lake (HH233 and HH240). Fine aeolian sands cover the inactive surface of the alluvial fan. The sands exhibit deposition ages of 6.0 ± 0.6 ka (HEH 22) and 7.26 ± 0.63 ka (HEH 23).

Fig. 4.4.9: OSL ages of different sediments.

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Two aeolian sand phases were recorded for the late Holocene (Fig. 4.4.9). The first phase lasted around 2000 yrs from around 4.3 to 2.3 ka. The phase is represented by eight OSL ages (HEH 05-1 to HEH 05-4, HEH 20-1, HEH 20-2, HEH 21-1, HEH 21-2) from 3 sections (HH025, HH223, HH224). Section HH025 is located near the outflow of Lake Heihai at the flanks of a small hill, while the other two sections are on the northern side of the lake.

A second aeolian sand phase from 1.24 to 0.35 ka (HEH 06-2 to HEH 06-4 and HEH 11-4) is documented in two sections (HH044, HH064, Fig. 4.4.10). In both sections the aeolian sands cover loess deposits. Section HH044 is a dune located at the north-eastern side of the lake. The base consists of reworked sandy deposits with a small amount of coarse material which are overlain by fine aeolian sands. These sands are covered by loess deposits with an age of 6.72 ± 0.35 ka (HEH06- 19). The upper 310 cm consist of fine and medium sands. The fine sands date between 0.73 ± 0.08 ka (HEH06-2) and 0.43 ± 0.06 ka (HEH06-39). The base of the medium sands at around 180 cm was dated to 0.35 ± 0.05 ka (HEH06-4). At 100 cm depth already a recent OSL age was measured. In general, section HH044 indicates a rapid aeolian accumulation of > 3 m of aeolian sediments in the last 1 ka. Additionally, a small, vegetation covered aeolian sand ridge about 70 m south of the foot of active longitudinal dune was dated (HH081). However, three OSL ages between 150 and 250 cm below the present surface yielded present ages again. As any post-depositional disturbance can be excluded in this position, rapid accumulation must be assumed. This is consistent with the results from the geomorphological mapping. The comparison of the field measurements with Corona satellite images from 1970 shows a northward movement of the dune crest of 50 m.

Fig. 4.4.10: Selected sections from different archives. 109

Loess sediments were dated in eight sections. Beside the two already mentioned sections (HH044, HH064), two section are located in the mountains and the foreland south of the lake (HH008 and HH143) and four in the area east of Lake Heihai (HH012, HH046, HH047, HH080). They indicate an accumulation of loess sediments between 8 and 4 ka (Fig. 4.4.9, Tab. 4.4.2).

Fluvial sands of the fill terrace east of the lake have an age of 5.3 ± 0.3 ka (HEH24, section HH241, Fig. 4.4.10). As this terrace can be traced until the outflow of the lake a relation to higher lake levels can be assumed.

Aeolian sands between lacustrine deposits were dated at two sections. At the southern side of the lake in section HH054 lacustrine sands have an age of 0.4 ± 0.03 ka in a depth of 121 cm. The sands are covered by fine-grained lake sediments (Fig. 4.4.10). In section HEH139 lake sediments yielded an age of 2.7 ± 0.2 ka (HEH 16) in 110 cm.

Discussion

The discussion focuses on three major issues in the landscape evolution of Lake Heihai: (1) the influence of the alluvial fans on the lake level, (2) the aeolian sediments as a proxy for the palaeoclimate evolution, and (3) the palaeoclimate of the area.

The influence of the alluvial fans on the lake level and the development of Lake Heihai

The development of the Lake Heihai basin and the lake itself is strongly coupled to the discharge from the Kunlun Shan range in the south. The south-eastern alluvial fan controls the outflow of the lake. In case of high discharge and sediment routing directly to the north, the outflow will be blocked. In the present state with a main discharge to the east, the outflow can cut down the alluvial fan deposits. The formation of the alluvial fans on the southern side of the basin is related to the glaciations in the southern mountain range. Four sets of terminal moraines were mapped in the valleys south of the lake. While two of them (M I and M I) are located close to the present glacier margin, the other two (M III and M IV) were deposited further downstream in the valley. Owen et al. (2006) presented dating results for glacial advances for the northern side of the main Kunlun Shan ridge. According to these ages, glacial advances occurred during the penultimate or early in the last glacial cycle, in the MIS 2 and in the Holocene. At Lake Heihai only two far reaching terminal moraines (M III and M IV) were identified. Due to the strong degree of weathering a formation prior to the early Holocene seems to be reasonable. Therefore, the terminal moraines M III and M IV might correlate with the two older glaciations recognized by Owen et al. (2006), with a last major advance during the MIS 2. However, as no large boulders were found on the terminal moraines, dating by CRN (cosmogenic radionuclides) was not possible. Terminal moraine M II is most probably late Holocene in age, e.g. the Little Ice Age, based on the geomorphological appearance. Little Ice age glacier advances have been documented all over the Tibetan Plateau (Yang et al., 2008; Yi et al., 2008; Xu and Yi, 2014). However, numerical ages from the central Kunlun are absent. Based on the evaluation of the remote sensing data sets, moraine M I is related to the modern glacier retreat.

As stated before, the evolution of Lake Heihai is closely related to the formation of the south-eastern alluvial fan. The infilling of the Kunlun River Gorge about 100 km to the east (Fig. 4.4.1) was dated to 110

the last glacial (Owen et al., 2006; Chen et al., 2011). This would again indicate a main phase of the fan formation prior to the Holocene. Additionally, four incision stages of mainly Holocene age were recorded in the Kunlun River Gorge (Chen et al., 2011). Nevertheless, backward erosion from the Kunlun River has not reached the Heihai basin. Therefore, the terraces at the outflow of Lake Heihai reflect local environmental changes. A minimum age for the last activity of the uppermost level of the alluvial fan can be deduced from the aeolian cover sediment on the northern side of the fan with an age of around 7.3 ka. However, the strongest sediment supply for the alluvial fan occurred most probably at the end of the last glacial after the last main glacial advance during MIS 2.

The section north-west of the lake (HH087, Figs. 4.4.7 and 4.4.8) provides further information of the lake evolution. The fluvio-lacustrine sediments were deposited in a near-shore environment and indicate lake levels > 10 m above the present level at around 50, 13 and 11.6 ka. We assume that these high lake levels were caused by the temporally blocking of the eastern outflow of the lake. The timing of the blocking at the end of the last glacial is probably caused by the decay of the large ice body in the Kunlun Shan. The dune sands between the fluvio-lacustrine sediments indicate a temporary drop of the lake level. Within the uncertainties of the OSL ages this drop might correlate with the Younger Dryas event. Low lake levels and a generally drier environment were documented at several places on the Tibetan Plateau (Yu and Kelts, 2002; Shen et al., 2005; Yan and Wünnemann, 2014; Thomas et al., 2016). These preliminary data indicate considerable lake level variations previous to the onset of the Holocene. Previous lake cores from Lake Heihai did not include pre- Holocene sediments (Lockot et al., 2016; Ramisch et al., 2016). However, from the present geomorphological situation it can be assumed that the influence of the alluvial fans on the development of the lake during the Holocene was only minor. The main outflow from the alluvial fan south-east of the lake is presently not connected to the lake.

The fluvial and lacustrine sections around the lake provide only limited information about the lake development during the Holocene. The lower fill terrace at the outflow of the lake has an accumulation age of 5.3 ka. This is well in agreement with previous results from Lockot et al. (2016). They assumed a lake level high-stand until around 6 ka. The incision of the fill terrace presumably caused a subsequent drop of the lake level of Lake Heihai. Two ages were obtained from aeolian sands in-between lacustrine deposits north and south of the lake. These ages of 2.7 and 0.4 ka might indicate phases of short low stand of the lake. However, at present they are not in accordance with previous reconstruction from the lake sediments at the same sections. According to Lockot et al. (2016) the lake level declined after an early to mid-Holocene phase with high-lake levels.

Aeolian transport processes

Aeolian sediments at Lake Heihai show a distinct spatial and temporal differentiation. Aeolian sand deposits occur mainly on the northern and eastern side of the lake, while loess deposits are dominant on the southern side. Only at two locations both sediment types are deposited in the same section. These spatial differences point to an immense influence of the related source areas. Sediment source areas and sediment availability are important parameters for the palaeoclimatic interpretation of aeolian sediments (Pye, 1995; Lehmkuhl, 1997; Lancaster, 2008; Stauch, 2016). Occurrences and activity of aeolian sands have been genetically linked in regard of fluvial processes, e.g. in Donggi Cona basin (Stauch et al., 2014), the Qaidam Basin (Yu et al., 2015), the Gonghe Basin

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(Qiang et al., 2016) and the Hexi Corridor (Nottebaum et al., 2015a). According to these studies the availability of aeolian sands can be limited to periods following enhanced fluvial activity. However, the results of this study indicate a main sand supply from the local rocks. Therefore, a constantly high sand supply can be assumed for the northern side. In contrast, aeolian sands are absent on the southern side of the lake, except in areas close to the present river channel.

The importance of the local topographic position for the distribution of silty sediments on the north- eastern Tibetan Plateau was also shown in the Qilian Shan (Küster et al., 2006; Nottebaum et al., 2015b) and in the Donggi Cona catchment (Stauch et al., 2012). Nevertheless, while in these areas the contribution from far distant sources is considered much higher, silty sediments at Lake Heihai are predominantly from local sources. In both other areas, the loess is considerable finer (at ~ 30 µm) (Stauch et al., 2012; Nottebaum et al., 2015b). Source areas are the active alluvial fans on the southern side and the exposed lake sediments around the lake. After deposition and temporary desiccation of the transporting river systems fluvial storages provide an potential source for fine grained aeolian sediments (Langford, 1989; Smalley et al., 2014). These sources for the fine-grained sediments are not available on the northern side of the lake. Furthermore, under the present environmental conditions the denser vegetation cover on the southern side of the lake supports the trapping of the fine-grained sediments. In contrast, the high sand content on the northern side of the lake results in a constant reactivation of the fine-grained particles due to the impact of individual sand grains. The clear spatial differentiation highlights the importance of a basin wide analysis of terrestrial sediments in the high mountain environment of the Tibetan Plateau.

Different interpretations exist regarding the palaeoclimate factors leading to the permanent accumulation of aeolian sands on the Tibetan Plateau. Some studies assume an accumulation during dry periods (Lu et al., 2011b; Liu et al., 2012) which are important for the mobilisation of the aeolian sediments (Bullard and Livingstone, 2002). Other studies focus on the constant fixation of the aeolian sands, e.g. by the formation of a vegetation cover (Stauch et al., 2012; Yu and Lai, 2012; Stauch, 2015; Li and Yang, 2016). They assume, that the OSL ages from aeolian sands only capture the last phase of the mobilisation and therefore indicate increasing moisture availability (Lancaster, 2008; Leighton et al., 2014). At Lake Heihai aeolian sands are deposited at different topographic positions which complicates palaeoclimate interpretation. Aeolian sands deposited prior to the mid-Holocene are only preserved where the sediments are protected against deflation, either by topographic effects or the subsequent burial by other sediments. Therefore the permanent fixation of these sediments cannot be clearly related to a denser vegetation cover or reduced wind speed. Hence, the accumulation in the catchment presumably occurred under relatively dry climate conditions. This observation points to the immense control of local geomorphological configurations on aeolian deposition and preservation throughout different climatic phases. It can be further assumed that aeolian deposits from the late glacial and early Holocene, which were not protected against deflation, were remobilised during a later dry period. Late Holocene aeolian sands are frequently assigned to dry climate phases (Stauch et al., 2012; Qiang et al., 2014; Yu and Lai, 2014; Lu et al., 2015; Stauch, 2015, 2016). This assumption is further reinforced by the coarsening of the upper sediments since at least 350 yrs in section HH044 and the sand deposits above the finer loess sediments in section HH064 since 1.2 ka. Additionally, the modern aeolian activity is high as indicated by 4 OSL ages which gave recent ages and the evaluation of the dune movement from satellite images.

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In contrast, loess sediments on the Tibetan Plateau are frequently related to wetter climate conditions (Stokes et al., 2003; Lu et al., 2011b; Yu and Lai, 2012). Wetter climate conditions result in a denser vegetation cover, which in turn leads to a more effective trapping of the sediments (Stauch, 2015). In contrast to many classical loess areas, a constant sediment supply by rivers and exposed lake sediments is occurring on the northern Tibetan Plateau. Therefore, the environmental conditions in the area of deposition play an important role. In the Heihai catchment most OSL ages from loess deposits have a mid-Holocene age between 8 and 3.5 ka. No older loess sediments were observed. At two sites in the study area a transition from loess deposits to fine sands was observed (e.g.HH044 and HH064). In both sections the silty sediments at the bottom have ages of around 7 ka. The deposition of these sediments continued until 3.5 ka. Only two ages are younger than the mid- Holocene. These sections (HH046, HH047) are located in the hills east of the lake and indicate a constant sediment supply by the exposed lake sediments. In the lower parts of the basin these young ages are missing.

Palaeoclimate evolution of the catchment of Lake Heihai

At present, results regarding the pre-Holocene evolution of Lake Heihai are sparse. Indications of a higher lake level at around 50 ka were found at only one site. However, it remains uncertain whether this higher level was related to wetter climatic conditions or if it was caused by a complete blocking of the eastern outflow caused by increased sedimentation of an alluvial fan. This phase was preceded by a phase of high aeolian activity between 100 and 80 ka. During the late glacial another high lake level occurred. During the transition from the late glacial to the early Holocene aeolian activity was probably high on the northern site of the lake. Similar phases of a dry environment were documented at several locations on the northern Tibetan Plateau (Yu and Kelts, 2002; Hong et al., 2003; Lu et al., 2011b; Yan and Wünnemann, 2014; Thomas et al., 2016). The mid-Holocene loess deposits indicate wetter climate conditions from around 8 to 4 ka. Before 8 ka the environmental conditions in the basin were presumably too dry to support the formation of a sufficient sediment trap. However, Lockot et al. (2016) assumed a high-stand at Lake Heihai which started as early as 10 ka and lasted around 6 ka. Ramisch et al. (2016) interpreted enhanced clastic inflow from the Kunlun range during the early Holocene as an indicator of the maximum northward extent of the ISM. This clastic inflow declined during the mid-Holocene. However, the terrestrial record in the lower part of the Heihai Basin has no indications of an early Holocene moisture period. We therefore assume that the ISM only affected the Kunlun range. Higher precipitation values in the upper parts of the mountains resulted in enhanced fluvial activity and rising lake levels. This process was probably further reinforced by the melting of the residual glaciers in the Kunlun Shan. During the mid- Holocene the clastic inflow declined due to reduced fluvial activity in the higher mountains and the diminution of the clastic sources. Nevertheless, during this time the lower parts of the basin received higher moisture values, as indicated by the loess deposits. This secondary moisture pulse would also explain the high-stand of Lake Heihai until 6 ka. The mid-Holocene moisture maximum might be caused by the intrusion of the EASM in the area. This would be in accordance with results from aeolian deposits further to the east (Yu and Lai, 2012, 2014; Stauch 2015) as well as peat deposits (Liu et al., 2015).

During the late Holocene the climate in the catchment returned to dry conditions again. Aeolian sand activity was strong in the last 4 ka. This interpretation is supported by falling lake levels during this 113

time (Lockot et al., 2016). A reduced ASM during the late Holocene was also documented in numerous records throughout the northern Tibetan Plateau (Ji et al., 2005; Colman et al., 2007; Y. Wang et al., 2010; P. X. Wang et al., 2014; Liu et al., 2011).

Conclusion

The formation and development of the Lake Heihai on the northern Tibetan Plateau during pre- Holocene times is related to the geomorphological processes on the large alluvial fans originating in the Kunlun Shan south of the lake. Strong sediment supply after the last major glaciation, which probably occurred during the MIS 2, resulted in the blocking of the outflow east of the lake and high lake levels during late glacial times. During the Holocene the influence of the fans on the lake was presumable very small.

Aeolian sediments revealed a complex picture due to different source areas. Geochemical analysis showed a strict coupling between aeolian deposits and the local bedrock geology. Sandy deposits and the related geomorphological landforms only occur on the northern side of the lake. In contrast, silty sediments are mainly preserved on the southern side of the lake. They are related to weathering products from the Kunlun Shan Range. The sources for the silty sediments are the alluvial fans between the mountain front and the lake and the exposed lake sediments around the lake. However, the spatial distribution of the aeolian archives also indicates that basin wide sampling is essential to reconstruct the climate evolution in this geomorphological highly active environment. Only the combination of the archives on both sides of the lake reveals a complete picture of the climate evolution at Lake Heihai.

Strong aeolian activity during the early Holocene is assumed due to the poor preservation of archives from this period. Aeolian sediments with pre-Holocene age were only found at two sites north of the lake. At both sites the preservation of the aeolian sediments is caused by site specific processes. This hampers comparison with the younger aeolian archives in the catchment.

Nevertheless, the aeolian deposits at Lake Heihai can be used to reconstruct the palaeoclimate since the mid-Holocene. The accumulation of loess during the mid-Holocene indicates wetter climate conditions than during the early Holocene in the lower parts of the basin. During the late Holocene a dryer climate resulted in a reduced trapping of silty sediments and the reactivation of aeolian sands. These climate phases are probably caused by the inundation of the EASM into the area during the mid-Holocene. Under the present conditions the catchment of the lake is outside of the influence of the monsoonal system.

Acknowledgements

This research is part of the project ‘Landscape and Lake-System Response to Late Quaternary Monsoon Dynamics on the Tibetan Plateau - Northern Transect’ in the framework of the SPP 1372 (Tibetan Plateau: Formation – Climate – Ecosystems) funded by the German Science Foundation (DFG). Steffen Pötsch and Jannik Stephanus supported the work at the lake. Elisabeth Dietze helped with the end-member analysis.

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5 Synthesis of OSL ages from the Tibetan Plateau

5.1 Geomorphological and palaeoclimate dynamics recorded by the formation of aeolian archives on the Tibetan Plateau

Georg Stauch

Abstract

Aeolian sediments are an important archive for palaeoclimate reconstructions on the Tibetan Plateau. The accumulation of aeolian sediments and the formation of palaeosols are used as a proxy for dry and wet phases, respectively. During the last decade the number of OSL (optically stimulated luminescence) ages of aeolian sediments has been rapidly increasing. This study summarizes the results from more than 300 individual OSL ages. Despite the widespread occurrence of aeolian sediments on the Tibetan Plateau nearly all OSL samples are located on the north-eastern and the southern part. In most regions the strongest accumulation occurred during the late Glacial and in the early Holocene. This coincides with the strengthening of the Asian summer monsoon on the Tibetan Plateau and wetter climate conditions after the relatively dry glacial times. The development of a sufficient vegetation cover, acting as a sediment trap, seems to be an important requirement for the preservation of aeolian archives. Short dry periods like the Younger Dryas led to a more enhanced aeolian accumulation without destruction of previously formed archives in some regions. The formation of palaeosols occurred mainly during the wet early and mid-Holocene, but the number of ages is relatively small. In contrast, a second phase of aeolian accumulation during the late Holocene was related to the weakening of the Asian summer monsoon and drier climate conditions. This indicates that the reaction of aeolian accumulation on the Tibetan Plateau depends on the previous climate state. A comparison between the OSL ages from aeolian sands and loess deposits yields a different timing of deposition of these two sediment types. While the sand deposits in different basins resemble wet and dry phases, loess deposits show no common signal between the analysed areas. This indicates an influence of local factors, such as varying source areas and erosive processes on the slopes. In general, the aeolian archives in the high mountain environment are governed by numerous factors besides the climatic influence.

Keywords: palaeoclimate, aeolian sediments, palaeosols, monsoon, Tibetan Plateau, dating

Introduction

The Tibetan Plateau (TP) with its mean elevation of above 5000 m asl (above sea level) (Fielding et al., 1994) has a strong influence on the supra-regional and the global climate system (e.g. Ye and Wu, 1998; An et al., 2000; Zhang et al., 2015). It consequently plays an important role in the regional

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water budget (Immerzeel et al., 2010). The climate on the TP is predominantly controlled by the interplay of the Asian summer monsoon (ASM) and the westerlies (e.g. Böhner, 2006). Variations in the intensity of these two systems lead to changes in the hydrological cycle and alterations of the geomorphological processes. However, the palaeoclimatic development in the region is still a matter of debate. Besides a limited amount of terrestrial archives and ice cores, lake sediments are one of the main sources of information on palaeoclimate for the late Quaternary. Nevertheless, many archives on the TP suffer from severe dating problems. For example, radiocarbon dating is hampered by the reservoir effect due to local differences in the input of old carbon in the hydrological systems (Wischnewski et al., 2011; Mischke et al., 2013). During recent decades, aeolian sediments on the TP gained increasing interest for the reconstruction of local and regional environmental changes (Fig. 5.1.1). Loess and aeolian sands are widespread on the TP (Hövermann, 1987; Pewe et al., 1995; Lehmkuhl, 1997; Derbyshire et al., 1998; Lehmkuhl et al., 2000; Sun et al., 2007; Kaiser et al., 2009a). They can be dated directly using optically stimulated luminescence (OSL). An OSL age is obtained by measuring the luminescence signal emitted by electrons trapped in mineral grains which are exposed to radiation. The luminescence signal is reset to zero due to the exposure of the grains to sunlight and accumulates again after the burial of sediments. Therefore, the OSL ages are intended to reflect the actual timing of deposition due to the fast bleaching of aeolian sediments (Singhvi et al., 2001; Wintle and Murray, 2006; Wintle, 2008). A still rising number of OSL ages is now available, enabling a detailed analysis of the different parameters controlling the deposition of aeolian sediments on the TP and the surrounding area. Nevertheless, a recent summary of the OSL ages from the whole TP and their significance for evolution of the palaeoclimate is missing.

Fig. 5.1.1: Overview of the Tibetan Plateau and spatial distribution of OSL ages used in this study. White dots indicate sampling position of the OSL ages. White lines demark individual regions. The pie charts show the relative number of OSL ages (see Tab. 5.1.1) and the related sediment types. A kmz file with more information of the individual OSL ages is available in the supplement section.

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In general, the accumulation of aeolian sediments is related to i) sediment availability in the source area, ii) transportation processes and iii) local conditions in the area of deposition (e.g. Pye, 1995; Smalley, 1995; Lancaster, 2008; Chase, 2009). Suitable sources of aeolian sediments on the TP are glacial outwash material (Sun et al., 2007; Smalley et al., 2014) as well as fluvial/alluvial deposits and dried lake basins (Lehmkuhl and Haselein, 2000; Kaiser et al., 2010; Pan et al., 2013; Stauch et al., 2014). Besides local sources, long distance transport of loess deposits is discussed (Lehmkuhl et al., 2000; Nottebaum et al., 2014). Some of these sources are only temporarily available, either due to limited sediment availability in general or the fixation of possible source sediments by vegetation (Sun et al., 2007; Stauch et al., 2012). Suitable trapping conditions and the related reduction in wind speed are essential for the accumulation of aeolian sediments. The trapping of the sediments may be controlled by several, sometimes interacting factors: the local topographic position (e.g. Mason et al., 1999), increase in surface roughness, e. g. due to vegetation cover (Bagnold, 1941; Tsoar and Pye, 1987; Lancaster and Baas, 1998; Sun et al., 2007; Yu and Lai, 2014), and a climate-induced reduction in wind speed (e.g. Roskin et al., 2011). There are different interpretations regarding the climatic factors controlling the accumulation of aeolian sediments and the formation of palaeosols on the TP. Typically, accumulation of aeolian sands is associated with dry climatic conditions, while loess deposition is related to intermediate conditions. For example, Lu et al. (2011b) identified two phases of high dune activity at the Qinghai Lake (10.3-9.1 ka and 8.9-7.8 ka) and related this activity to dry climate conditions and low effective moisture. After 7.8 ka, the development of palaeosols indicates greater effective humidity. In the same area Liu et al. (2012) related the accumulation of aeolian sands from 14.6 to 7.4 ka to episodic short periods of dry climate conditions interrupted by two periods of soil formation (12.2-11 ka and 10-9 ka). In the neighboring eastern Qaidam Basin hyper- arid conditions are indicated in an aeolian sand accumulation from 12.4 to 11.6 ka (Yu and Lai, 2014). Similar results have been obtained for aeolian sands in the Gonghe Basin (Liu et al., 2013b) and southern Tibet (Pan et al., 2014). This is in accordance with studies from the Chinese deserts north and north-east of the Tibetan Plateau. Mason et al. (2009) reconstructed an enhanced aridity from 11.5 to 8 ka based on OSL dated sand dunes. Similar results have been presented by Lu et al. (2011a) and Yang et al. (2011, 2013). The formation of dunes at the northern border of the Chinese Loess Plateau (CLP) from 35 to 25 ka has also been related to dry climate conditions (Long et al., 2012b).

Several recent studies from the TP highlight the importance of a suitable sediment trap for the accumulation and stabilisation of aeolian sediments on the TP (Stauch et al., 2012; Yu and Lai, 2012). In the northern Qilian Shan loess deposition during the Holocene has been interpreted as an indicator for reduced aridity and the development of vegetation after very dry glacial times which prevented the accumulation of aeolian sediment (Stokes et al., 2003; Küster et al., 2006). Similar arguments have been used to explain the accumulation of loess in southern Tibet (Sun et al., 2007). In the Eastern Qaidam Basin, Yu and Lai (2012, 2014) supposed a strong increase of effective moisture in the early Holocene from 11.6 to 8.3 ka due to the synchronous accumulation of loess and sand. In the catchment of the Donggi Cona the formation of aeolian sand and loess deposits from 10.5 to 7 ka was associated with the development of a sufficient vegetation cover due to wetter climate conditions (Stauch et al., 2012). This is in accordance with recent studies on the interpretation of OSL ages from dunes of different desert margins worldwide which indicate, that the ages reflect a decline in sand movement and stabilisation of sedimentary bodies (Chase and Thomas, 2007; Lancaster, 2008; Leighton et al., 2014). During periods of high aridity and wind speed, the sediments are frequently reworked and the luminescence signal is reset (Chase and Thomas, 2007). The recycling rate of the sediments, however, varies between different dune types (Bateman et al., 117

2003b). Thus, the explanations for aeolian sand accumulation reach from wet to dry conditions. These different environmental explanations of aeolian sands are partly related to uncertainties regarding the meaning of OSL ages from different types of aeolian sediments. The formation of palaeosols as well as phases without aeolian deposition is generally related to wetter conditions. On the TP most studies suggest a parallel formation of the soil with the accumulation of the aeolian sediments. Lu et al. (2011b) identified two phases of wetter climate conditions at ~9 ka and 4-3ka at the Qinghai Lake based on OSL ages of palaeosols. Liu et al. (2012) documented four phases of palaeosol formation around the Qinghai Lake (~12.2-11 ka, 10-9 ka, 5.2-4 ka and 3.2-0.7ka). They supposed that especially during the Mid-Holocene optimum (~ 8-5 ka) accumulation rates of aeolian sediments where too low and soils mainly developed in the previously deposited sediments. In contrast, Yu and Lai (2014) reported especially high accumulation rates of the palaeosols during a phase from around 8 to 3 ka in the eastern Qaidam Basin. The dating of the formation of the palaeosols by OSL imposes some problems as the dating result refers to the hosting sediment and not to the soil itself. The deposition can either occur parallel to the formation of the soil or the pedogenetic process can affect an older sedimentary deposit. The latter process has been recently highlighted for soils of the Chinese loess plateau (CLP) (Stevens et al., 2007). Additionally, large variations in OSL ages have been reported from palaeosols on the CLP which might be related to reworking and/or mixing of the sediments due to bioturbation and or colluvial and mass wasting processes (Stevens et al., 2006).

The different interpretations of the climatic factors controlling the accumulation of the aeolian sediments on the TP, the local and regional variations for occurrence of the archives impose a major problem in the general climatic interpretation. A recent summary of all available OSL ages of aeolian sediments from the whole TP is lacking. Consequently, the palaeoenvironmental conditions, like wet and dry phases, leading to the accumulation and preservation of aeolian sediments are still not well constrained. This review evaluates the available OSL ages of loess, aeolian sand and palaeosols on the TP in relation to their spatial distribution and the available sediment classes in order to improve the understanding of factors controlling the accumulation of aeolian sediments in this region.

Study area

The Tibetan Plateau is the largest plateau on Earth, comprising an area of roughly 2.5 x 106 km². The study area is bound by the Himalaya in the south, the Pamir to the west, the Tarim Basin to the north and the Chinese Loess Plateau and the Sichuan Basin to the east (Fig. 5.1.2). The formation of the TP is related to the continent-continent collision between Asia and India which started at around 50 Myr (Yin and Harrison, 2000; Tapponnier et al., 2001; Royden et al., 2008; Hetzel, 2013). While the margins of the TP are formed by large mountain systems like the Kunlun Shan, the Qilian Shan or the Longmen Shan, the interior is relatively flat with a relief of usually less than 1 (Fielding et al., 1994). The formation of the TP and the Himalayas had a strong influence on the intensity of the regional monsoonal systems and the aridification of Central Asia (Hahn and Manabe, 1975; Ruddiman and Kutzbach, 1989; J. Sun et al., 2010). However, the influence on the monsoonal system strongly depends on the palaeoaltitude and the extension of the TP during its formation, which is still not well constrained (e.g. Spicer et al., 2003; Harris, 2006; Molnar et al., 2010). The traditional view that the high elevation of the TP serves as a heat source which is coupled to the intensity of the Asian Monsoon has been recently challenged by an alternative concept highlighting the importance of the 118

Himalayas as a topographic barrier to the monsoonal system (e.g. Boos and Kuang, 2010). In contrast, numerical model experiments show a regional intensification of the monsoon by the uplift of the northern TP (R. Zhang et al., 2012; Zhang et al., 2015).

Fig. 5.1.2: The Tibetan Plateau and the major atmospheric systems (solid lines: Asian summer and winter monsoon; dashed lines: westerlies; according to Böhner, 2006; Yao et al., 2012).

The present climate of the TP is controlled by the interaction of the mid-latitude westerlies and the Asian monsoon system (Indian monsoon and East Asian monsoon). There is a profound difference in the climate of the different parts of the TP. In the south-eastern part high summer temperatures and high precipitation prevail, while the north-western part exhibits low summer temperatures and low precipitation values (Böhner, 2006; Y. Wang et al., 2010; Herzschuh et al., 2011; Maussion et al., 2013). Most parts of the TP are influenced by the Indian monsoon during summer, with strongest effects at the southern and south-eastern margin. The eastern margin is influenced by the East Asian monsoon. In the western and northern part the westerlies play an important role (Böhner, 2006; Yao et al., 2012). During winter the East Asian winter monsoons develops due to the pressure gradient between the Siberian High and the Aleutian Low, resulting in strong north-westerly winds on northern TP (Chen et al., 2000; Böhner, 2006). The mid-latitude westerlies are shifted to the south and split into two branches east of the TP. The stronger southern branch is located south of the Himalayas and the weaker one north of the TP (Böhner, 2006). Aeolian transport mainly occurs during spring time when the area is under the influence of the Siberian high pressure system and the westerlies (Porter and An, 1995; Derbyshire et al., 1998; Sun et al., 2001; Zhang et al., 2001; Mao et al., 2011; Stauch et al., 2012). The aeolian transport during spring time is related to the breakdown of

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the Siberian High and the development of cold fronts, which leads to the generation of windstorms (Roe, 2009).

The climate of the TP experienced some major variations on orbital and millennial timescales. The monsoonal system is influenced by the insolation variation caused by changes in the orbital parameters and the internal feedback mechanisms of the earth (e.g. Kutzbach, 1981; Wang et al., 2008; Shi et al., 2011). Despite the broad consensus on the general influence of the monsoon in most parts of the TP, there are still many uncertainties regarding the timing of wet and dry periods. This is at least partly related to the remoteness of the area which results in a low number of studies and dating uncertainties. Studies of mountain glaciation on the TP and the Himalayas, for example, show only a few mountain systems exhibit a strong glacial advance during Marine Isotope Stage (MIS) 2. Instead, most glaciers in the monsoon influenced part of the TP advanced during MIS 3, the late Glacial and the early part of the Holocene, probably due to increased precipitation (Owen and Dortch, 2014). Despite this overall picture, the timing of glaciations in individual mountain systems still vary significantly (Lehmkuhl and Owen, 2005; Owen et al., 2008; Owen and Dortch, 2014). These variations have also been highlighted for other archives on the TP (e.g. Wischnewski et al., 2011). Nevertheless, a general climate trend has been deduced at least for parts of the TP. During much of MIS 2 cold and dry conditions with low lake levels seem to dominate (e.g. Herzschuh, 2006), owing to a profoundly weaker monsoon system and an approximately 10 degrees southward shift of the westerlies (Kapp et al., 2011). Climate amelioration on the TP started during the late Glacial from around 15 ka onwards (Y. Wang et al., 2010). This period was probably synchronous to the Bølling/Allerød in Europe (Herzschuh, 2006). Other archives record a shift to warmer and wetter conditions only during the Holocene (e.g. An et al., 2012). Many records indicate a shift to colder and drier conditions for about 1.5 to 2 ka during the Younger Dryas (Hong et al., 2003; Y. Wang et al., 2010, 2014). The early Holocene is characterized by an increase in moisture again (Hong et al., 2003; An et al., 2012). Several authors noted an asynchronous Holocene climate optimum throughout the TP, with wettest climate conditions during the early Holocene in areas dominated by the Indian monsoon and during the middle Holocene in areas influenced by the East Asian monsoon and the westerlies (e.g. Herzschuh, 2006; Y. Wang et al., 2010). In contrast, Cai et al. (2010) argued on the basis of oxygen isotope values from different caves in China for a synchronous start of the climate optimum in areas influenced by the Indian and East Asian monsoon. A later change to wetter climate conditions had also been described for the arid regions in northern China outside the monsoon influenced areas (Chen et al., 2008; Mason et al., 2009; Zhao et al., 2013). The late Holocene is characterized by drier conditions again (Herzschuh, 2006; Colman et al., 2007; An et al., 2012). During the last decades, a warming trend has been observed over much of the TP (You et al., 2010; K. Yang et al., 2014). Increasing temperature results in retreat of most glaciers (Kääb et al., 2012; Yao et al., 2012) and degradation of the widespread permafrost (Qiu, 2008; Xue et al., 2009; K. Yang et al., 2014).

Methods

For this study all prior published OSL dating results of aeolian deposits from the TP were compiled in one database. Aeolian sediments are very suitable for dating with OSL. It is assumed that the luminescence signal has been exposed to sufficient light and has been reset during the transportation process. After the burial of the sediment ionizing radiation from both naturally 120

occurring radioactive isotopes in the surrounding sediment and cosmic rays causes the excitation of atoms within the crystal lattice of the grains. The subsequent trapping of these electrons is linked to lattice defects in the crystal and is an ongoing process during the time of burial. The artificial stimulation of the sample in the laboratory releases the stored energy as a luminescence signal. The luminescence signal is proportional to the radiation dose received since burial (equivalent dose) and by dividing this by the rate of ionizing radiation dose from cosmic rays and the surrounding sediment, the time since burial can be estimated (e.g. Singhvi et al., 2001; Duller, 2004; Stevens et al., 2007; Preusser et al., 2008; Roberts, 2008; Singhvi and Porat, 2008; Wintle, 2008). Typically quartz or feldspar grains are used for the measurement of the luminescence signal. The use of the luminescence signal from feldspar grains to calculate the equivalent dose tends to underestimate the age of the sediment due to a loss of luminescence signal over time called anomalous fading (e.g. Auclair et al., 2003). However, the saturation of feldspar grains occurs later than in quartz grains. Hence, the luminescence signal from feldspar grains can be used to date older sediment samples (Buylaert et al., 2007). Additionally, quartz samples are sometimes contaminated by feldspar which impedes the accurate determination of the equivalent dose (Hülle et al., 2010; Stauch et al., 2012). The OSL method can be used to obtain ages for relatively young sediments with an age of a few years to several decades in suitable environments (Madsen and Murray, 2009) of up to 100 ka for feldspar ages and a lower limit of quartz ages (e.g. Buylaert et al., 2007; Wintle, 2008; Lai, 2010). Some studies extended this range to more than 200 ka (e.g. Buylaert et al., 2012; Li et al., 2014).

Despite being used since the 1980s, recent advancements in methodological approaches and in the measurement equipment used for OSL especially since 2000 (Stevens et al., 2007; Roberts, 2008; Wintle, 2008), complicates the comparison of ages produced in the early application phase of the method with the recent ones. To avoid methodological inconsistencies in the database, studies prior to 2003 were excluded, as well as studies based on radiocarbon ages and TL (thermoluminescence) samples. The reasons for the exclusion are that the TL signals on aeolian sands are not always completely reset by sunlight (Yu and Lai, 2012) and that radiocarbon ages might be problematic due to hard water effects and the possible remobilisation of organic material. Dates are also rejected where the studies in question present poorly record, or limited, stratigraphic information for the dated archives or the methodological descriptions are incomplete. Several OSL ages of aeolian infillings of sand wedges are additionally available (Porter et al., 2001; Owen et al., 2006; Madsen et al., 2008; Liu et al., 2010). These were also not used because of the specific formation of the archive.

An overall 353 OSL ages from 109 individual sections in eight regions are included in the database. A complete list of all OSL ages used in this study is presented in the supplement section of this paper (S1) as well as a kmz file for the location of the samples. From the north-west to the south-east these regions are (Fig. 5.1.1, Tab. 5.1.1): the northern Qilian Shan (Stokes et al., 2003; Küster et al., 2006; Pan et al., 2013); the area around the Qinghai Lake (Madsen et al., 2008; Rhode et al., 2010; Lu et al., 2011b; Liu et al., 2011, 2012); the Gonghe Basin (Liu et al., 2013b; Qiang et al., 2013a); the Qaidam Basin (Yu and Lai, 2012, 2014; Zhou et al., 2012; Yu et al., 2013, 2015); the Anyemaqen Shan (Owen et al., 2003; Lehmkuhl et al., 2014); the Donggi Cona catchment (Stauch et al., 2012, 2014); the Kunlun Shan (Owen et al., 2006) and southern Tibet (Sun et al., 2007; Lai et al., 2009; Pan et al., 2014). Nearly all OSL ages are based on quartz grains except for 40 OSL ages of feldspars from the Donggi Cona region and two samples from the eastern Qaidam Basin (Yu et al., 2015). The feldspar samples from the Donggi Cona are affected by anomalous fading resulting in a possible underestimation of the age of around 20% according to fading tests of two samples (Stauch et al.,

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2014). The feldspar ages from the Donggi Cona area are presented without fading correction and have to be regarded as minimum ages. All OSL ages were checked to avoid inconsistencies in the data set. Ages with a 1σ error higher than 20% of the mean ages were excluded (n: 4) as well as ages which were inverted in comparison to the neighboring ages in the same section (n: 11) and indistinct stratigraphic description (n: 2).

Subsequently the samples were analysed according to the respective sediment type (loess, sandy loess, sand and palaeosol). Based on the classification of the individual authors this study mainly focused on two classes, sand and silt. However, grain-size data in the studies is often not well documented. This is especially problematic as the term loess is often used in a very broad sense. Despite these limitations most authors followed the classical classification which defines loess as aeolian sediment which is mainly composed of silt-sized particles (e.g., Pye, 1995). The class of sandy loess (n: 12) was only used in studies from the Donggi Cona (Stauch et al., 2012) and the Anyemaqen Shan (Lehmkuhl et al., 2014). If a sedimentary classification of the palaeosols or additional grain-size data was available these sediments were also included. However, except for seven samples, the grain-size data of the palaeosols was not sufficiently documented. The degree of the palaeosol development (e.g. strong or weak) was not incorporated in the database, because it is a relative measure which differs to a certain degree even in neighboring regions.

Probability density functions (pdf) were calculated for each OSL age based on the mean and the standard deviation. . The individual pdfs then were summed as a cumulative pdf. Each OSL date is treated as an individual accumulation event (Singhvi et al., 2001; Lai et al., 2009; Yu and Lai, 2012). Aeolian deposits on the TP often include distinct phases of erosion (Yu et al., 2013; Stauch et al., 2014). Interpolation between individual dating points can lead to a substantial error in the ages of the sediments. Only the age of the sediment at the OSL sampling point is well known. However, there are substantial differences in the preservation of the different aeolian sediments and therefore in the interpretation of the OSL ages obtained from these sediments. As outlined before aeolian sands are prone to reworking during phases of active aeolian transport (Chase and Thomas, 2007; Telfer and Thomas, 2007; Lancaster, 2008). Hence, the OSL ages of aeolian sands are biased towards the end of an aeolian phase (Telfer et al., 2010). This problem should apply less to loess sediments. Studies from the CLP also indicate gaps in the loess sediments which are possibly be caused by deflation (Stevens et al., 2006, 2007; Buylaert et al., 2008). In turn, OSL ages from palaeosols can be affected by mixing processes, e.g. caused by bioturbation (Stevens et al., 2006). This would result in larger age variations (Lomax et al., 2007). It has to be pointed out that bioturbation especially caused by the plateau pika (Ochotona curzoniae) is frequent in nearly all loose sediments on the TP.

A further problem arises by different research strategies. Only a few studies attempted to date rather young aeolian sediments, resulting in an underrepresentation of late Holocene ages in some regions. Other studies used aeolian sediments to date the underlying geomorphological landforms and just sampled the basal part of the aeolian section, e.g. Pan et al. (2013) in the south-eastern Qilian Shan. Samples for the upper part of the section are therefore not available. By considering all available ages site and study specific disturbances are reduced and the general climate signal can be deduced.

Phases of enhanced accumulation were marked if at least three OSL ages exhibit an overlapping error bar. This approach excludes single OSL ages, which may represent a random accumulation or an outlier. Therefore these phases correspond to a higher probability in the pdfs. As each phase includes 122

the minimum and the maximum values of the standard deviation, phases are likely to represent the maximum duration of the depositional phase and can overestimate them. In contrast, as the OSL samples represent only a small subsample from an aeolian section the actual phases of deposition are generally longer than the reported mean ages, which might compensate for the aforementioned effect. This is especially important for aeolian sands which are frequently reworked during times of high aeolian activity. In this case only the end of the transportation phase is recorded. Additionally, older phases have a longer duration due to a larger standard deviation. Short term changes in the accumulation of aeolian sediments are easily blurred by this method. However, if a sufficient number of ages are available the cumulative pdf can also indicate shorter changes on the millennial scale.

Results

Spatial distribution of OSL ages

The overall spatial distribution of the data shows a strong concentration of OSL ages on the north- eastern TP including the Qilian Shan (Fig. 5.1.1). The greatest number of ages are available from the Qaidam Basin (116), the Qinghai Lake area (60), the catchment of the Donggi Cona (42), and the Qilian Shan (41). Only 9 ages originate from the high mountain area of the Anyemaqen Shan and one from the Kunlun Shan area close to the city of Golmud (Tab. 5.1.1).

Tab. 5.1.1: Spatial distribution of OSL ages.

Region Sections OSL Sand Loess Palaeosol Sandy

ages loess

Qilian Shan 20 41 0 41 0 0

Qinghai Lake 19 60 32 19 13 0

Gonghe Basin 5 30 16 5 12 0

Qaidam Basin 24 116 61 32 23 0

Anyemaqen 2 9 0 1 0 8

Donggi Cona 22 42 33 5 0 4

Kunlun Shan 1 1 0 1 0 0

Southern Tibet 15 37 19 18 1 0

108 336 161 121 49 12

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Timing of aeolian sediment deposition

The analysis of 336 OSL ages, including sand, loess, sandy loess, and palaeosols indicates a main phase of aeolian accumulation on the TP during the Holocene. 82% (275) of the OSL ages have a mean age of 11.7 ka or younger and only 23 OSL ages are older than the global last Glacial Maximum (gLGM, ~21 ka; Clark et al., 2009). Except for two ages these are either from southern Tibet or the Qilian Shan. The oldest age in the dataset is from the Qinghai Lake area and is older than 165 ka (Liu et al., 2012). Therefore, the analysis focuses on the time since 21 ka. Most of the mean ages are between 10 and 7 ka (82) and between 0 and 2 ka (75) while relatively few OSL ages are from 5 to 7ka (26).

A more detailed picture emerges from a regional analysis of the OSL ages in regard of the sediment types. As mentioned before, this analysis relies only on the classification of the individual studies.

Sand

The overall number of OSL samples from sandy deposits on the TP is 161. Out of these, 133 ages are of Holocene age and 6 are older than the gLGM. Except for one age from the Gonghe Basin, all ages beyond 21 ka are from southern Tibet.

In the Qinghai Lake area mean OSL ages from sand deposits range from 16.3 ± 1.4 ka to 0.15 ± 0.01 ka (Fig. 5.1.3) (32 OSL ages). Three phases of stronger sand accumulation are recorded. The first one was from the latest Pleistocene, ranging from 17.7 until 11.4 ka. A second phase is in the early Holocene, from 10.8 to 6.5 ka. According to the cumulative pdf aeolian accumulation is strongest at 13 and 9.3 ka. The third phase covers the time from 1.9 to 0.7 ka. Three ages are younger than the last phase. The dataset from the Gonghe Basin is much smaller and includes only 16 OSL ages. One of them is older than 21 ka. There is one period of enhanced sand accumulation from 12.4 ka to 8.6 ka. The phase extends until 7 ka if two OSL ages with non-overlapping error bars are included (difference 0.1 ka; light shading in Fig. 5.1.3). Four ages are between 6 ka and the present. Peaks in the pdf indicate strongest accumulation at 11.2 and 9.2 ka. Sand deposition in the Qaidam Basin has again a more complex picture. The database includes 61 OSL ages ranging from 12.4 ka to the present. Enhanced aeolian accumulation took place from 13.1 ka - 10.8 ka, 10.5 - 6.4 ka, 3.6 - 2.3 ka, 2.2 - 1.4 ka and 0.8 - 0.6 ka. The pdf indicates again two phases of stronger accumulation at 11.7 and 9.3 ka. In the Donggi Cona catchment a first phase of stronger aeolian accumulation was from 10.7 to 7.2 ka comprising only the minimum number of 3 OSL ages. The pdf expresses a single peak at 8.7 ka. Three further ages are 14 ka or younger. During the late Holocene 5 phases of enhanced aeolian accumulation occurred with short interruptions during the last 3.1 ka (3.1 – 2.4, 2.0 - 1.3, 1.3 - 1, 0.9 - 0.4, 0.4 - 0.3 ka). In contrast, enhanced aeolian sand accumulation in southern Tibet was much earlier from 31.6 ka until 12.7. Due to the large error, which ranges from 3.4 to 1.4 ka, only 8 ages cover the timespan of nearly 20 ka. A second phase took place from 9.2 to 6.2 ka, with a peak in the pdf at 7.5 ka. There are no OSL ages from sand in the Qilian Shan in the database, because the desert areas in the northern forelands were excluded from this analysis.

Summarizing the OSL ages from sandy deposits of the six regions indicates some regional trends. Archives of aeolian sand are mainly younger than 17 ka in the north-eastern part of the TP. According to the dataset, enhanced sand accumulation started first in the area of the Qinghai Lake. Sandy 124

archives in the three other regions of the north-eastern TP are several thousand years younger. In the Qaidam Basin enhanced accumulation started at 13.1 ka, at 12.4 ka in the Gonghe Basin, and at 12.5 ka in the Donggi Cona catchment. The end of this first phase of enhanced aeolian activity coincides closely in all regions on the north-eastern TP. The OSL ages range from 6.5 to 7 ka. However, the number of OSL ages already decreased from 8 ka onwards. In most of the regions on the north-eastern TP the cumulative pdf expresses two peaks of stronger accumulation in these phases at around 13 to 11.2 ka and at 9.3 ka. Sand accumulation in southern Tibet started much earlier and ended at around 13 ka. The end of the second phase in southern Tibet at 6 ka is comparable to north-eastern TP. Only 5 ages on the north-eastern TP range between 7 and 3.6 ka indicating a phase of strongly reduced sand mobility during the mid-Holocene. Enhanced sand accumulation started again in the Qaidam Basin at 3.6 ka, in the Donggi Cona catchment at 3.2 ka, and in the Qinghai Lake area at 1.9 ka. In these regions the late Holocene accumulation phase terminated between 0.3 and 0.7 ka. However, there are only few samples which have been taken from a depth of less than 50 cm below the surface. This appears to indicate a continuous accumulation until the present time. There is no indication of an enhanced late Holocene accumulation of aeolian sands in the Gonghe Basin. This might be related to the relatively small dataset and the sampling strategy in the respective studies. Based on the available OSL ages, two phases of stronger aeolian accumulation on the north-eastern TP can be identified. One started at the end of the Pleistocene and lasted until 6.5 ka and a second one started roughly at 3.6 ka. Individual regions have a slightly different timing of aeolian sedimentation. Up to now, the dataset from southern Tibet is relative small regarding Holocene ages. However, the early Holocene sand accumulation phase (9.2-6.2) terminated at a similar time as on the north-eastern TP.

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Fig. 5.1.3: OSL ages and probability density for aeolian sand from the Tibetan Plateau for the last 21 ka. Shaded areas are phases of enhanced aeolian accumulation.

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Sandy loess

The category of sandy loess has been used only in two studies (Anyemaqen Shan, Lehmkuhl et al., 2014; Donggi Cona, Stauch et al., 2012). 12 ages are available (not shown, see supplement S2). They date from 11 ka to the present. At the Anyemaqen Shan the youngest age is around 4 ka, while at the Donggi Cona sandy loess is observed only during the last 2 ka.

Loess

121 OSL ages in the dataset are originating from loess. The largest dataset is available for the northern Qilian Shan (n: 41). In the Qaidam Basin 32 OSL ages originate from loess sediments, which are exclusively from the eastern part of the basin and 19 ages are from the Qinghai area. A small number of loess ages is available from the Gonghe Basin (5), the Donggi Cona catchment (5) as well as the Kunlun Shan (1) and the Anyemaqen Shan (1). In southern Tibet 18 loess ages are published. If the whole dataset is taken into account, the accumulation of loess is relatively homogenous throughout the Holocene (Fig. 5.1.4), except a gap in OSL ages between 3.6 and 3.9 ka. Ages of loess deposits older than 15 ka are again mainly from southern Tibet and the Qilian Shan, except two ages from the Qinghai Lake area.

In the northernmost region, the Qilian Shan, loess accumulation started at 12.9 ka and ended at 3.7 ka (Fig. 5.1.4). 5 OSL ages are younger than this interval. Older ages come exclusively from the south- eastern Qilian Shan. The mean ages are between 76 ± 8.6 ka and 19.2 ± 1.5 ka. They concentrate in two phases from 84.6 ka to 48.8 and from 31.1 ka to 17.7 ka. Each of these older phases encompasses 4 OSL ages. In the Qinghai Lake area enhanced loess accumulation started at 14.9 ka. There is one age of 16.1 ± 1.2 ka which is older. This sample is not included due to the small overlapping 1σ error (50 a) and is regarded as an outlier. This phase terminated at around 8.1 ka. Three samples yielded ages between 4.6 ka and 2.9 ka. No phase of enhanced loess accumulation was identified in the dataset in the Gonghe Basin. There is one sample with an OSL age of 10.2 ± 1 ka and four ages younger than 1.1 ka. In the Qaidam Basin two OSL ages are from the beginning of the Holocene (10 ± 0.6 ka, 9.2 ± 0.5 ka). A phase of stronger accumulation of loess occurred from 8.3 ka until 3.9 ka and again from 3.5 ka until 1 ka. In the Donggi Cona area only one weak phase of loess accumulation (three ages) was identified. This phase lasted from 10.6 ka until 8.1 ka. One OSL age is younger (4.5 ± 0.4 ka) than this phase and one older (12.7 ± 1 ka). Two additional ages are available from other areas: in the Kunlun Shan further to the west loess has been dated to 8.6 ± 0.7 ka (Owen et al., 2006) and in the Anyemaqen Shan one OSL sample provided an age of 12.3 ± 1.1 ka (Owen et al., 2003). In southern Tibet there are 8 ages of loess deposits which are younger than 21 ka. Due to the wide distribution of ages throughout the Holocene no phase of enhanced loess accumulation could be identified. 5 OSL ages range from 90 ka to 56 ka.

At the present state, loess deposition on the TP seems to have a heterogeneous spatial pattern. Whilst in the Qinghai Lake area loess accumulation ranged from 14.9 ka until the early Holocene, other regions, such as the Qilian Shan and the Eastern Qaidam Basin, loess accumulation was predominantly a Holocene phenomenon. From the more central regions like the Gonghe or the Donggi Cona Basin none or very few OSL ages from loess deposits are available.

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Fig. 5.1.4: OSL ages and probability density for loess from the Tibetan Plateau since 21 ka. Shaded areas are phases of enhanced aeolian accumulation.

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Palaeosols

From aeolian sediments in which palaeosols were developed 49 OSL ages are available. Out of these, 13 ages originate from the Qinghai Lake area, 23 from the eastern Qaidam Basin, 12 from the Gonghe Basin and one from southern Tibet. The OSL ages span the time between 16.4 ± 1.3 ka and 0.7 ± 0.1 ka (Fig. 5.1.5). Analysis of the ages of palaeosol formation from OSL dating results is difficult as the hosting sediment can either be deposited a considerable time before the development of the soil or more or less synchronously. There are no clear indicators for the differentiation between both processes, which makes a paleaoclimatic interpretation of these deposits complicated.

Phases of enhanced accumulation of sediments in which palaeosols developed at the Qinghai Lake range from 11.8 to 8.5 ka and from 5.3 to 3.7 ka. Two samples are around 2.7 ka and another two have ages of 1.3 ± 0.1 ka and 0.9 ± 0.1 ka. In the Gonghe Basin two older OSL ages were obtained at 16.4 ± 1.3 ka and at 12.6 ± 1.4 ka. A phase of palaeosol sediment formation occurred from 10 to 7.1 ka whilst single OSL ages of 5.6 ± 0.3 ka to 0.7 ± 0.1 ka are also recorded. The eastern Qaidam Basin has a phase of palaeosol sediment formation from 12.9 until 6.6 ka, from 5.8 to 4.5 ka and one from 4.3.to 2.6 ka. Despite the differences in the duration and onset of the phases of enhanced accumulation of sediments hosting soils, two general phases can be identified; in the early Holocene and in the middle Holocene.

Fig. 5.1.5: OSL ages and probability density for palaeosols from the Tibetan Plateau since 21 ka. Shaded areas are phases of enhanced aeolian accumulation.

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Discussion

The regional distribution of OSL ages from the TP is biased, with a majority of ages from the north- eastern part of the TP and only few ages from southern Tibet. However, aeolian sediments such as aeolian sand and loess are widespread on the TP (Hövermann, 1987; Pewe et al., 1995; Liu and Zhao, 2001; Li et al., 2006; Lehmkuhl et al., 2000, 2002; Sun et al., 2007). Most areas on the western and southern TP are very remote and difficult to reach. Administrative and logistic problems due to the high altitude and sparse settlements might explain the biased spatial distribution of the ages. The large spatial data gaps prevent an analysis of the timing of aeolian sediment deposition on the entire TP. Further OSL dating is required to properly identify regional trends across the whole TP.

Phases of enhanced aeolian accumulation

Significant accumulation of aeolian sediments on the north-eastern TP started in the early Holocene (Fig. 5.1.6). Only in the Qinghai Lake area is there evidence for an earlier onset of enhanced aeolian accumulation (about 6 to 7 ka earlier than in other regions). Additionally, an earlier phase has been recorded for the easternmost region in the dataset, the eastern Qilian Shan. In southern Tibet aeolian sediments show a partial reversal to this pattern. A first accumulation phase started at around 30 ka and lasted until the end of the last Glacial. No OSL ages are reported for the very early Holocene, between 13 and 9 ka, a timespan when the regions in north-eastern Tibet indicate strong aeolian accumulation. The end of the early Holocene accumulation period shows remarkably similar ages at around 7 to 6.5 ka, followed by a phase of weak aeolian activity in four of the six studied regions. Only results from the Qilian Shan and the Qaidam Basin indicate enhanced aeolian activity by the deposition of loess from around 6.5 to 3.6 ka. During the late Holocene enhanced accumulation took place in the Eastern Qaidam Basin, in the Donggi Cona area, and for about 600 years at the Qinghai Lake. Continuous deposition until the present was proposed for the loess in the Qilian Shan (Küster et al., 2006) and the eastern Qaidam Basin (Yu and Lai, 2014).

The timing of enhanced phases of loess and sand deposition on the TP does not show similar trends in the different regions. While loess sedimentation in the area of the Qinghai Lake falls into a phase of sand accumulation but is shorter in time, in the eastern Qaidam basin loess accumulation during the middle Holocene is bracketed by phases of enhanced sand accumulation. The timing of onsets for sand and loess accumulation in the Donggi Cona catchment is similar but sand accumulation lasts longer than the loess phase. In general, it can be considered from the presented dataset that phases of enhanced loess accumulation are spatially and temporally heterogeneous while phases of sand accumulation are similar in the regions of the north-eastern TP. These differences may at least partly be explained by different source areas and erosional phases. In the Qilian Shan and the eastern Qaidam Basin, loess accumulation took place during most of the investigated time span which is different from all other studied regions on the TP. The accumulation in these areas might be caused by their proximity to major source areas. Major aeolian sediment input into the Qilian Shan comes from the adjacent desert in the north and west as well as the large alluvial fans in the foreland of the mountain system (Derbyshire et al., 1998; Nottebaum et al., 2014). The eastern Qaidam Basin is located downwind of the remnants of a large palaeolake, which provides an excellent source of silty material (Kapp et al., 2011).

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Aeolian sediments and palaeoclimatic dynamics in neighboring areas

Comparing the timing of the aeolian sediments on the TP with the areas to the east and north of the study area reveals some differences in the aeolian phases. The Chinese Loess Plateau (CLP) is located east of the study area with deposits of 150 to 300 meters thickness (An, 2000), covering an area of 440,000 km² (Liu, 1988). The formation of the loess-palaeosol sequences has been related to changes in the strength of the north-west winter monsoon and the Asian summer monsoon during the last 2.6 Ma (e.g. Kukla, 1987; Porter and An, 1995; An, 2000; Lu et al., 2004). The aeolian deposition of the underlying red clay started 7-8 Ma (An, 2000). Even older ages of 22 Ma are reported for the western part of the CLP (Guo et al., 2002). It was assumed that loess deposition on the CLP is continuous with highest deposition rates during glacial times and strongly reduced rates during the interglacial periods with soil formation (Porter, 2001; Kohfeld and Harrison, 2001, 2003; Vriend et al., 2011). This pattern is in contrast to the TP where only a few aeolian sediment archives from glacial periods are preserved. However, there are considerable changes in the deposition rate of interglacial periods on the CLP, with two to three times higher rates during MIS 1 than during MIS 5. The high rates during MIS 1 are related to strong accumulation during the late Glacial and the early Holocene while the lowest rates are from 7 to 3 ka (Kohfeld and Harrison, 2003). In turn, this is comparable to the timing in sediment accumulation on the north-eastern TP. However, the parameters guiding the accumulation rates during MIS 1 on the CLP are not fully understood (Kohfeld and Harrison, 2003). In general, highest accumulation rates appear in the north-west of the CLP and the lowest rates in the South-east (Kohfeld and Harrison, 2003). This is consistent with a large thickness (260 m) of loess deposits in the western most part of the loess plateau in the Huangshui valley near the city of Xining (Lu et al., 2004). Additionally, grain size analysis shows a decreasing trend from the north-west to the south-east, indicating a dominant sediment transport from the west and north-west (Lu and Sun, 2000; Nugteren et al., 2004; Yang and Ding, 2004; Vandenberghe et al., 2006). Recent detailed OSL sampling indicated that the loess deposits in the CLP are probably not always continuous (Stevens et al., 2006, 2007; Buylaert et al., 2008), which in turn questions the calculated accumulation rates for the CLP during different times.

Large desert areas with active and inactive dunes and sand sheets cover a significant portion of China north and north-east of the TP (e.g. X. Yang et al., 2012). High aeolian activity has been reported especially for the early Holocene (Li et al., 2002; Sun et al., 2006; Mason et al., 2009; Qiang et al., 2010; Lu et al., 2011a; Yang et al., 2011, 2013). The end of the aeolian phase varies between ~ 11.5 (Qiang et al., 2010; Liu and Lai, 2012) until 8 ka (Mason et al., 2009; Lu et al., 2011a), which is slightly earlier than on the north-eastern TP. Most studies from aeolian sections describe a strongly reduced aeolian activity, stabilisation of aeolian sediment and soil formation during the middle Holocene until about 5.6 ka (Sun et al., 2006) to 2.5 ka (Mason et al., 2009; Qiang et al., 2010). During the late Holocene a reactivation of aeolian archives is frequently documented (e.g. Mason et al., 2009; Yang et al., 2011). This pattern is again similar to the accumulation phases on the north-eastern TP. In contrast to the other studies Qiang et al. (2010) identified an additional dry period from around 7.2 to 5 ka at the southern margin of the Tengger Desert close to the CLP. Several studies from other archive support the hypothesis of a dry early Holocene and a wet mid-Holocene (Chen et al., 2008; Herzschuh, 2006; Long et al., 2012a; X. Yang et al., 2012). This pattern was also observed in climate modelling experiments (Jin et al., 2012).

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Aeolian phases, palaeoclimate development on the TP and implications for the formation of aeolian sediments on the TP

gLGM

Aeolian accumulation during glacial times was weak in most places on the TP, except at the borders of the eastern Qilian Shan and in southern Tibet. However, on the TP a high aeolian sediment flux can be assumed for the gLGM (~21ka). As such, a dominant aeolian input from 32 - 11.5 ka was proposed on the basis of the grain-size fractions larger than 25 μm of a core from Qinghai Lake (An et al., 2012). Cold and dry conditions were documented for many places on the TP. Many lakes desiccated (Herzschuh, 2006) or had low lake levels, for example Qinghai Lake (Lister et al., 1991; Rhode et al., 2010), Lake Kuhai (Mischke et al., 2010c), the Donggi Cona (Dietze et al., 2010; Opitz et al., 2012), Hala Lake (Wünnemann et al., 2012; Yan and Wünnemann, 2014) and Nam Co (Daut et al., 2010). Dust flux onto the CLP was also high during glacial times (An et al., 2000; Nugteren et al., 2004; Xiao et al., 1995; Stevens et al., 2006; Stevens and Lu, 2009; Lu et al., 2011b; Vriend et al., 2011). Large areas served as potential dust sources (Kohfeld and Harrison, 2001, 2003; Nilson and Lehmkuhl, 2001; Yang et al., 2004). In contrast to the aeolian sediments on the TP there is a large number of OSL ages on the CLP related to glacial phases.

Additional factors which might have prevented constant accumulation of aeolian sediments during the gLGM are the more widespread distribution of permafrost (Zhao et al., 2014) and the influence of glaciers and enhanced glaciofluvial outwash, especially at the end of the last Glacial. While it is difficult to assess the impact of permafrost, the direct influence of glaciers can be omitted for most of the study sites since these areas have not been glaciated during the gLGM or later. The influence of fluvial activity, probably caused by enhanced glaciofluvial outwash, has been recently highlighted (e.g. Sun et al., 2007; Yu et al., 2015). However, most of the glaciers on the TP had significant advances during the late Glacial (14 – 12 ka; Owen and Dortch, 2014), which is synchronous with the beginning of enhanced aeolian accumulation in most areas.

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Fig. 5.1.6: Palaeoclimatic comparison of the aeolian phases and selected records: (A) Phases of enhanced sediment accumulation and soil formation on the Tibetan Plateau, (B) June summer insolation at 30°N (Berger and Loutre, 1991), (C) δ18O record from Dongge Cave (Dykoski et al., 2005), (D) δ13C record from Hongyuan peat record (Hong et al., 2003), and (E) effective moisture for monsoonal central Asia (Herzschuh, 2006).

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Late Glacial and early Holocene

Enhanced aeolian accumulation of sand deposits started in north-eastern Tibet during the late Glacial and the early Holocene (Fig. 5.1.6A). During the late glacial warmer and wetter conditions than during the gLGM were reported from many study sites on the north-eastern TP. More humid conditions in the area are related to the strengthening of the ASM (e.g. Yu and Kelts, 2002) which is also shown in the shift towards lighter δ18O values at Dongge Cave (Dykoski et al., 2005). This strengthening of the ASM has been related to higher insolation at low latitudes (e.g. Kutzbach, 1981) (Fig. 5.1.6B). Y. Wang et al. (2014) reported increased precipitation values from the Donggi Cona from 13.5 ka onwards, while in the neighboring Lake Kuhai wetter conditions occurred especially from 12.8 ka (Mischke et al., 2010c). This is in accordance with the results from Qinghai Lake (Shen et al., 2005; Colman et al., 2007) and Lake Luanhaizi (Herzschuh et al., 2010a). At the Qinghai Lake, in the Gonghe Basin and especially in the Eastern Qaidam Basin, several OSL ages are from sediments with palaeosols. As these ages mirror the ages of the sediments without palaeosols a later formation of the palaeosols cannot be excluded. However, this assumption is not valid for sections where the palaeosols are covered by aeolian sediments. This applies to most of the sections in these three areas. Peak aeolian accumulation in three of the regions indicates stronger aeolian accumulation during the Younger Dryas period around 11.7 ka. Drier conditions during this time are indicated by the δ18O values at Dongge Cave (Dykoski et al., 2005) (Fig. 5.1.6C) and the effective moisture reconstruction for monsoonal central Asia (Herzschuh, 2006) (Fig. 5.1.6E). In most regions of the north-eastern TP, the aeolian archives have not been eroded, indicating the existence of at least a sufficiently stabilising vegetation cover. In contrast, at the Donggi Cona this drier period has resulted in complete recycling of the older aeolian sediments, which explains the late onset of the phases of enhanced accumulation. Aeolian accumulation continued during the full monsoonal conditions in the early Holocene. A second peak of aeolian accumulation was at around 9.2 ka, corresponding again with heavier δ18O values at Dongge Cave. This event has not been documented in the record from the Hongyuan peat or in the effective moisture reconstruction (Fig. 5.1.6D and 5.1.6E). This late Glacial – early Holocene accumulation phase on the TP ended at 7 to 6.5 ka.

According to these palaeoclimate reconstructions, the first phase of enhanced sand accumulation corresponds to a phase under more humid climate. This indicates a stabilisation of the aeolian sediments, most probably owing to the development of vegetation. The importance of a sufficient vegetation cover has been highlighted in several studies from the TP before (e.g. Stokes et al., 2003; Sun et al., 2007; Liu et al., 2012; Stauch et al., 2012; Yu and Lai, 2012) and has also been observed in other regions worldwide (Lancaster and Baas, 1998; Singhvi and Porat, 2008; Chase, 2009). Another parameter for the continuous accumulation of aeolian sediments might be a reduction in wind speed in the post-glacial period. Nevertheless, there are no independent proxies for the reconstruction of wind speed available. During the early Holocene accumulation phase, OSL ages in three regions of the north-eastern TP indicate stronger phases of palaeosol formation, further supporting the hypothesis of wetter climate conditions. Short periods of drier climate during this phase, like the YD, lead to the enhanced accumulation of aeolian sediments probably due to large source areas (e.g., exposed lake sediments).

In the deserts of northern China high aeolian activity has presumably acted until 8 ka, which is slightly older than on the TP. For these areas a formation under mostly dry condition was reconstructed for the early Holocene (Mason et al., 2009; Yang et al., 2011). This is in contrast to the early Holocene interpretation of the aeolian sediments from the TP. Several explanations for the 134

different climate evolution have been proposed, which included varying influences of the westerlies and the East Asian monsoon and enhanced evaporation rates (Herzschuh, 2006; Mason et al., 2009). Additionally, changes in wind speed due to variations in the location of the westerlies and the large amount of available sediment can explain differences in preservation of the aeolian archives in this areas. However, these remain relatively speculative and further analyses of the differences between the northern deserts and the TP, which are only evident during the early Holocene, are required.

The aforementioned aspects have some additional implications for the palaeoclimatic interpretation of aeolian sediments on the TP at the end of the last Glacial and the early Holocene. The majority of OSL ages from this period indicate a stabilisation phase due to higher humidity and denser vegetation cover. Prior to that loess and sand had high turnover rates which resulted in a frequent bleaching of the sediments. The presumably harsh climate conditions during glacial times led to erosion of most of the older aeolian archives on the TP. Such sediments have only been preserved at a few sites which were protected from erosion, explaining the low number of OSL results older than the gLGM. Three regions on the TP show a different temporal pattern of aeolian sediment accumulation. In the eastern Qilian Shan, loess sediments of gLGM and pre-gLGM ages have been reported (Pan et al., 2013). The loess deposits in this area are thicker (>10m) than in any other region on the TP. This area can be regarded as transition zone between the TP and the western CLP. Climate conditions in the area must have been more favorable for the preservation of sediment archives. At the Qinghai Lake, accumulation of aeolian sand started 3 to 6 ka earlier when compared to other TP areas. An increase in pollen concentration has been reported in Qinghai Lake from 16.9 ka onwards (Ji et al., 2005), which might indicate an earlier development of a vegetation cover in the surrounding of the lake despite still relatively cold and dry conditions. However, the palaeoclimatic proxies from the area yield no distinct explanation. The valleys in southern Tibet protected the aeolian sediments from erosion during the harsh climate conditions in glacial times (Lai et al., 2009). The accumulation phase at around 8 ka was related to a short time interval of colder conditions (Lai et al., 2009). The synchronous OSL ages from loess and sand phases and from palaeosols point to a parallel formation of the latter during the accumulation of the aeolian sediments.

Mid-Holocene

The phase of enhanced aeolian sedimentation of the sand decreased gradually from 8 ka onwards and ended at around 7 to 6.5 ka. This decrease presumably indicates a reduction in available source material as a consequence of denser vegetation cover. Only seven OSL ages of sand from the whole TP range between 6.5 and 3.5 ka. A reduction in wind speed can be assumed, possibly with synchronously reduced sediment source areas. A more frequent accumulation of loess is recorded by 25 OSL ages, with the majority of ages from the Qilian Shan and the eastern Qaidam Basin. In the Qinghai Lake catchment and in the eastern Qaidam Basin a phase of enhanced soil formation is recorded. There are still some uncertainties over whether the climate during the middle Holocene was wetter than the early Holocene or slightly drier and if so whether or not this is related to the influence of different monsoon systems (Y. Wang et al., 2010, 2014; Zhang et al., 2011). Synchronously, stronger erosion of slope material occurred on the north-eastern TP (Kaiser et al., 2007; Stauch et al., 2014). This erosive phase led to the reworking of many loess deposits on the slopes and might explain the low number of OSL ages from loess deposits in most regions on the north-eastern TP. Only in the eastern Qaidam Basin and in the Qilian Shan were loess deposits 135

preserved and accumulated continuously, probably due to the minimized fluvial erosion caused by the exceptionally dry climate in both areas in comparison to the other regions. The proximity to the major source areas of aeolian dust must also be considered.

Late Holocene

The accumulation of sand resumed at 3.6 ka in the Qaidam Basin, at around 3 ka at the Donggi Cona and at 2 ka in the Qinghai Lake area. No phases of enhanced aeolian sand accumulation are documented from the Gonghe Basin and from southern Tibet. These missing phases can be attributed to the relatively small datasets from these areas and may not necessarily be related to climate. The reactivation of sand deposits on the TP was caused by a return to drier conditions during the late Holocene which led to higher sediment availability. Several pollen and lake records from the TP indicate an onset of drier conditions between 5 and 4 ka caused by a reduction of the ASM (Hong et al., 2003; Shen et al., 2005; Colman et al., 2007; Zhang et al., 2011; Bird et al., 2014; Y. Wang et al., 2014). Other authors assumed an even later onset of drier conditions at 3.2 ka (Herzschuh, 2006). Dried lake beds have been an additional source for aeolian sediments (IJmker et al., 2012b; Stauch et al., 2014). The later onset of the aeolian accumulation might be caused by remobilisation of the earliest late Holocene aeolian archives. The impact of a drier climate has potentially been exacerbated by a stronger human influence on the landscape on the TP, at least during the late Holocene (e.g. Cui and Graf, 2009; Kaiser et al., 2009b; Schlütz and Lehmkuhl, 2009; Miehe et al., 2014).

Conclusions

This study summarises the published OSL ages from the Tibetan Plateau and their relation to palaeoclimatic changes since the gLGM. To date, the largest cluster of OSL ages is available for the north-eastern TP. Due to the large spatial data gaps the identification of regional trends for the whole TP is difficult. Further OSL ages between the north-eastern and southern parts of the TP are required, as well as ages from the central and western areas.

The OSL ages from aeolian sands from the north-eastern TP yield a coherent picture on a broad scale and are consistent with the influence of the reconstructed ASM. Only a few OSL ages are available from glacial times, dry and windy conditions during this time prevented the preservation of aeolian archives. Only at sites which are more protected, e.g. by local topography, are glacial age aeolian sediments preserved until today. In most regions the accumulation of aeolian sands started in the late Glacial to the early Holocene and ended at around 7 to 6.5 ka. From around 7 to 3.6 ka sand mobility was strongly reduced. The late Holocene is characterized by a widespread reactivation of aeolian sediments. Phases of enhanced aeolian activity of aeolian sand on the southern TP do not correspond with those on the north-eastern TP. However, the reasons still need to be identified. Possible causes are stronger protection against erosion and additionally, different source areas of the aeolian sediments. In contrast to the coherent picture regarding aeolian sands, the timing of loess accumulation shows strong regional differences. This can be attributed to different amounts of sediment in local source areas and erosional processes on hill slopes.

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Different climate forcing agents influence the aeolian processes in the early and late Holocene, depending on preceding climate conditions. In the early Holocene increased effective moisture after dry glacial times led to the accumulation and constant fixation of aeolian sediments. In contrast, drier climate conditions during the last 3.5 ka resulted in a reactivation of aeolian sediments following the relatively wet mid-Holocene. This indicates that there is no consistent relationship between aeolian accumulation and aridity on the TP and highlights that the local conditions in the area of deposition are of special importance in this high mountain environment. Aeolian sediments are not a simple palaeoclimate archive in the unique high mountain environment of the Tibetan Plateau. The number of influencing factors is higher than in the classical Chinese Loess Plateau and the accumulation is controlled by different parameters.

Acknowledgements

The presented study greatly benefited from stimulating discussions with numerous colleagues, especially Frank Lehmkuhl, David Loibl, Veit Nottebaum, Wolfgang Römer, Philipp Schulte and Christian Zeeden. Constructive comments from the editor Ian Candy as well as Thomas Stevens and one anonymous reviewer are gratefully acknowledged.

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5.2 Multi-decadal periods of enhanced aeolian activity on the north- eastern Tibet Plateau during the last 2ka

Georg Stauch

Abstract

The north-eastern Tibetan Plateau is regarded as key area for the understanding of the Holocene paleoclimate in central Asia. During the last decade a special emphasis has been placed on multi- decadal to millennial scale climate fluctuations, especially in the context of the recent climate change. However, most reconstructions are based on lake sediments, tree rings and speleothems whereas only little information from terrestrial archives is included. This study presents multi- decadal scale climate fluctuations based on optical stimulated luminescence (OSL) ages from aeolian sediments from three catchment areas. Six phases of enhanced aeolian accumulation during the last 2000 years, each lasting around 80 to 200 years were identified. The first three phases (1630-1725 CE, 1450-1530 CE and 1250-1350 CE) occurred during the Little Ice Age; the other three (750-950 CE, 390-540 CE, 50-225 CE) during the so-called dark ages cooling. Aeolian processes were strongly reduced during the medieval climate anomaly. A comparison with other proxy records indicates that the formation of aeolian archives on the north-eastern Tibetan Plateau during the late Holocene is facilitated by cool and dry climate conditions during times of weaker Asian Summer Monsoon and probably enhanced westerlies. The results show that short term climate fluctuations can be reconstructed from non-continuous and heterogeneous terrestrial archives in a semi-arid environment, provided a sufficient number of OSL ages from aeolian sediments is available.

Keywords: palaeoclimate; Holocene; north-eastern Tibet; aeolian sediments

Introduction

Multi-decadal to millennial scale climate variations have gained increasing interest in the recent years (e.g. Mayewski et al., 2004; Wanner et al., 2011; PAGES 2k Consortium, 2013). Frequent climate fluctuations during the Holocene have been recorded in marine sediments (Bond et al., 1997, 2001; Jennings et al., 2002; de Vernal and Hillaire-Marcel, 2006), lake sediments (Noren et al., 2002; Jones et al., 2006; Geirsdóttir et al., 2013; Hildebrandt et al., 2015), speleothemes (Fleitmann et al., 2003; Y. Wang et al., 2005; Dykoski et al., 2005; Zhang et al., 2008; Baker et al., 2015) or glacial deposits (Denton and Karlén, 1973; Solomina et al., 2015) in many regions around the world. The understanding of the causes and consequences of these short term climate fluctuations are crucial in the context of the recent warming. However, global reconstructions of temperature and precipitation revealed no coherent results for the entire Holocene (e.g. Mayewski et al., 2004; Wanner et al., 2011) and indicate substantial spatial and temporal differences for the last 2000 years (PAGES 2k Consortium, 2013).

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The north-eastern Tibetan Plateau (TP) has been identified as an especially climate sensitive region due to its location at the northern boundary of the Asian summer monsoon (Indian Summer Monsoon and East Asian Summer Monsoon) and the mid-latitude westerlies (e.g. Böhner, 2006; Chen et al., 2010; Henderson et al., 2010; An et al., 2012) (Fig. 5.2.1). Previous climate reconstructions mainly focused on the timing and strength of the monsoonal influence in the area during the entire Holocene (Herzschuh, 2006; An et al., 2012; Y. Wang et al., 2014). During the global last glacial maximum (gLGM) and late glacial times the influence of the moisture bearing Asian summer monsoon was weak. During the early and middle Holocene the Asian summer monsoon reached its maximum and declined during the late Holocene (Wang et al., 2008; Y. Wang et al., 2010; Liu et al., 2015). These long term climate changes are related to orbital variations leading to changes in summer insolation at low latitudes which in consequence drive the Asian monsoon system (Kutzbach, 1981; Shi et al., 2011; Sun et al., 2015). During times of a weaker monsoon the influence of the westerlies was presumably stronger (Porter and An, 1995; Stevens et al., 2007; Jin et al., 2015; Zhu et al., 2015). Beside these orbital induced changes several reconstructions of Holocene climate identified millennial scale variations in the strength of the Asian Summer monsoon in East Asia (Hong et al., 2003; Y. Wang et al., 2005; Wang et al., 2008; Yu et al., 2006; Li et al., 2015) and on the north- eastern TP (Mischke and Zhang, 2010; An et al., 2012; B. Yang et al., 2014). However, there is still considerable difference between the reconstructed timing of the different short-term events during the late Holocene. This is related to uncertainties in proxy interpretation (e.g. Wischnewski et al., 2011; Yan and Wünnemann, 2014) and problems with the established chronologies. Many reconstructions are based on proxies from lake sediments. Radiocarbon dating on lake sediments from the TP is frequently influenced by the reservoir effect due to the input of old carbon in the hydrological system. The amount of the reservoir effect varies between lakes and even between studies from individual lakes (Long et al., 2011; Hou et al., 2012; Mischke et al., 2013). Therefore further proxies have to be established for the reconstruction of Holocene climate variations on the north-eastern TP (e.g. Li et al., 2015).

Fig. 5.2.1: The north-eastern Tibetan Plateau, main atmospheric systems and the three areas with OSL ages for this study (Qinghai Lake, Eastern Qaidam Basin and Donggi Cona). Data source background image: Blue marble next generation – NASA Earth Observatory (Stöckli et al., 2005). 139

Aeolian sediments on the north-eastern TP are a suitable indicator for palaeoclimate reconstructions (Lehmkuhl et al., 2000; Madsen et al., 2008; Liu et al., 2012; Stauch et al., 2012; Yu and Lai, 2012; Stauch, 2015). Aeolian sediments can be dated directly by OSL (optical stimulated luminescence), which captures the timing of the accumulation and burial of the aeolian sediment grains (Wintle, 2008). Palaeoclimate reconstructions from aeolian sediments are mainly based on sedimentological proxies requiring interpolation between the dated parts of the individual sections (Yu et al., 2006; Liu et al., 2013b; Lehmkuhl et al., 2014). However, many terrestrial sections on the north-eastern TP are affected by varying accumulation rates and non-continuous accumulation (Qiang et al., 2013a; Stauch et al., 2014; Yu et al., 2015) leading to large age uncertainties for the sediments between the obtained ages. This also obstructs in many cases the determination of detailed accumulation rates for individual sections. The number of available OSL ages of aeolian sediments from the north- eastern TP has been rapidly increasing in the recent years. By now, the dataset is large enough to apply a different approach that relies solely on OSL ages, thus avoiding any data interpolation between individual dating points (Lang, 2003; Stauch, 2015).

This study analyses aeolian sediments from different regions of the north-eastern TP for the last 5000 years with a special focus on the last 2000 years. The dataset is evaluated regarding its potential to reveal multi-decadal to multi-centennial climate fluctuations. Subsequently the phases of high and low aeolian accumulation are interpreted in the context of the different forcing mechanisms, such as monsoonal intensity and the strength of the westerlies.

Study area

The north-eastern Tibetan Plateau (Fig. 5.2.1) is characterized by large intramontane basins at elevations between 3000 and 4000 m above sea level. The surrounding mountains reach elevations of more than 5000 m and are partially glaciated. The formation of the basins started at around 10 to 15 Ma ago (H.-P. Zhang et al., 2012; Hetzel, 2013). Major fault-lines like the Kunlun fault have a profound influence on the landscape (Van Der Woerd et al., 2002; Fu and Awata, 2007). Large earthquakes with magnitudes (M) of more than seven occur (Guo et al., 2007). The present-day climate in the basins is semi-arid with annual precipitation between 160 mm in the eastern Qaidam Basin and 300 to 400 mm at the Qinghai Lake (Liu et al., 2011; Yu and Lai, 2012), reflecting the reduced influence of the Asian Summer Monsoon from the east to the west. Aeolian sediments, like loess, sand sheets and active dunes are widespread in the basins (Lehmkuhl, 1997; Lu et al., 2011b; Stauch et al., 2012; Qiang et al., 2013a; Yu and Lai, 2014).

For this study, 75 OSL ages from the last five ka were evaluated. They cover three regions (Fig. 5.2.2); the Donggi Cona catchment (32 ages; Stauch et al., 2012), the Eastern Qaidam Basin (28; Yu and Lai, 2012, 2014; Yu et al., 2015), and the Qinghai Lake area (15; Rhode et al., 2010; Liu et al., 2011; Lu et al., 2011b, 2015). All OSL ages and sections used for this study are listed in the Supplement Table S1 and are available at the website of the journal.

The Qinghai Lake is located at the north-eastern side of the study area in an elevation of 3200 m. With a lake surface area of 4473 km² it is the largest lake without an outflow in China (Liu et al., 2011). OSL samples have been obtained from nine sections which are all located in a few kilometres distance from the lake shore (Fig. 5.2.2A). Seven sections are located in or nearby of a dune field on the eastern side of the lake. The OSL ages from aeolian sand and loess deposits consist of dune sand, aeolian sand covering raised shorelines and sections with varying loess and sand horizons. The 15

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OSL ages relevant for this study were taken from depth of 14 to 163 cm below surface. A clear age/depth relationship was not observed. Similar, the different geomorphological situations do not result in specific age clusters.

Fig. 5.2.2: The three regions and location with sections containing OSL samples younger than 5 ka (A: Qinghai Lake area; B: Eastern Qaidam Basin; C: Donggi Cona catchment).

Samples from the Eastern Qaidam Basin are from the transition between the eastern Tiekui Desert and the bordering mountain systems (Fig. 5.2.2B). They were obtained at elevations between 3000 and 3500 m. All samples are from sections with alternations of dune sand and loess. However, sediments classified as loess still have a significant content larger than 100 μm (Yu et al., 2015). Sand deposits were related to the eastward expansion of the Tiekui Desert, while the source of the loess is located the western Qaidam Basin where large areas of lake deposits are exposed (Yu et al., 2015).

Lake Donggi Cona is the southernmost catchment area in this study in an elevation of 4090 m. Aeolian sands and sandy loess are located east and south of the lake while loess deposits were found in the mountains in the north-western part of the catchment (Fig. 5.2.2C). A large alluvial fan is the main source of the sand east of the lake. Based on the geochemical signature and additional sediment source for the sandy sediments with an age younger than 300 years was identified. These young sediments originate from sandy deposits which were washed into the Donggi Cona and subsequently blown out when the lake level dropped (Stauch et al., 2014). Dune sediments south of the lake originate from smaller alluvial fans in the neighbouring basins while for the loess sediments a source in the Qaidam Basin was assumed (Stauch et al., 2012).

Material and methods

Database

Accumulation of aeolian sediments is influenced by several factors like sediment availability, sediment supply, transportation processes and environmental conditions in the area of deposition (e.g. Pye, 1995; Kocurek and Lancaster, 1999; Lancaster, 2008; Chase, 2009; Halfen et al., 2016). During the last 4 to 5ka drier conditions have been observed on the north-eastern TP (Shen et al., 2005; Y. Wang et al., 2010, 2014; Zhang et al., 2011; An et al., 2012), which leads to a substantial reactivation of aeolian processes (Qiang et al., 2014; Lu et al., 2015; Stauch, 2015). For the reconstruction of multi-decadal aeolian events on the north-eastern TP, all published OSL ages of aeolian sediments covering the last five ka were analysed (Supplement Table S1). In total, 75 ages

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from 36 sections were evaluated. The database comprises ages from sand and loess deposits from different geomorphological positions and varying sediment sources. To avoid the establishment of a third sediment class, four ages from the Donggi Cona area, which were previously classified as sandy loess (Stauch et al., 2012) were reclassified as sand. OSL ages of palaeosols were not considered because these might be prone to substantial bioturbation (e.g., Stevens et al., 2007). Two ages with a 1σ error larger than 20% were also not used for the calculation. The quality of OSL ages has been significantly improved during the last decade due to progress in measuring equipment and methodological approaches (Stevens et al., 2007; Roberts, 2008). Consequently, ages published prior 2003 were excluded to obtain a consistent database. A similar approach based on a likewise dataset was previously used to analyse accumulation in the entire Holocene (Stauch, 2015). This study focuses on short term changes with an updated database.

Cumulative probability functions

In this study, each OSL age is used as a proxy for an accumulation event. Therefore, a cluster of similar OSL ages represents a phase of strong aeolian accumulation, enabling the identification of phases of stronger and weaker aeolian accumulation. There are some critical points which have to be addressed. Higher sampling density of typical stratigraphic units would result in a higher number of ages from this unit and therefore an overestimated aeolian accumulation (Lang, 2003). However, sediments in the different basins on the north-eastern TP are affected by frequent erosion resulting in heterogeneous archives, with, for example, a varying number of intercalated paleosols (Yu and Lai, 2014; Lu et al., 2015; Qiang et al., 2016). Typical marker horizons are missing. These factors make the area very suitable for this method. Additionally, the relatively large number of OSL ages in contrast to the number of ages from regional studies reduces the influence of local effects which might be overestimated by the analysis of individual sections. From the individual ages a probability density function (pdf) was calculated based on the mean ages and the standard deviation and summed as a cumulative pdf (Singhvi et al., 2001; Lang, 2003; Goble et al., 2004; Qiang et al., 2013a; H. Wang et al., 2014; Stauch, 2015). Therefore, peaks in the cumulative pdf represent phases of enhanced aeolian accumulation. This method has also been applied for other archives, e.g radiocarbon chronologies of fluvial archives (Macklin et al., 2005; Johnstone et al., 2006; Hoffmann et al., 2008; Chiverrell et al., 2011; Jones et al., 2015). However, there are several problems related to the use of pdf from OSL ages, despite being far less problematic than the use of pdf from (calibrated) radiocarbon chronologies (Chiverrell et al., 2011). Pdfs from OSL ages do not reflect variations related to the distribution of the measured equivalent dose as these variations are not normally distributed (e.g. Galbraith, 2010; Galbraith and Roberts, 2012; Fitzsimmons et al., 2013; Dietze et al., 2016). Furthermore, as the standard deviation is much smaller for younger ages than for older ones, peak height decrease and peak width increase with age (Hesse, 2016). Consequently, peak height cannot be used as indicator of the strength of the aeolian activity (Singhvi et al., 2001). To avoid a bias towards the older end of the cumulative pdf (Lang, 2003), also ages up to six ka were included in the calculation (Supplement Table S1), but the analysis covers only the last five ka. However, reworking of aeolian sediments by wind is also occurring frequently, reducing the number of older sediment ages (Telfer et al., 2010; Bailey and Thomas, 2014). This is resulting in a positive curvelinear distribution of the temporal frequency of ages with a higher number of young ages and a lower number of older ages, which is sometimes called “taphonomic effect” (e.g Surovell and Brantingham, 2007; Surovell et al., 2009). However, this effect might not be so important for the identification of 142

short term variations (Surovell and Brantingham, 2007). Therefore, for this study only the peaks during the last 2000 years and not the height of the peaks are interpreted. Additionally, sampling for OSL ages normally starts in a depth of 30 cm or lower of the sedimentary body. Therefore, it is not possible to discriminate if the low number of OSL ages covering the last 250 yrs (n: 2) is related to a reduced aeolian activity or, more probably, to the sampling strategy. Consequently, the last 250 years are also excluded from the analysis. A detailed summary of the different problems in OSL dating of dune sediments which have to be considered for the reconstruction of aeolian activity in general and also the use of the pdf was recently presented by Hesse (2016) and Thomas and Burrough (2016). The 90 % and 10 % quantiles for the cumulative pdf were established on 10,000 artificial, randomly generated probability functions (Fig. 5.2.3). Each artificial pdf is based on the actual number and the mean error in the real dataset (Michczynska and Pazdur, 2004; Hesse, 2016). Therefore, 75 OSL ages with a one sigma uncertainty of 8.49% were generated. Based on the peaks above the 90% quantiles periods and phases of enhanced aeolian activity were identified. Their beginning and end is based on the nearest inflection points (rising or falling) of the pdf of all OSL ages of aeolian sediments on the whole north-eastern TP.

Fig. 5.2.3: Artificially generated probability functions. Calculation of the 90% and 10% quantile was based on 10,000 artificial generated probability functions. The black line is the cumulative probability function based on the real OSL ages, while the grey lines in the back are artificial generated pdfs. The upper and lower dashed line denote the 90% and 10% quantile respectively. Peaks in the black line above the 90% quantile are assumed to be phases of enhanced aeolian accumulation which are statistically robust. Artificially generated pdfs for the individual study areas are shown in the supplement section (S2).

Results

Aeolian accumulation on the north-eastern TP was analysed on two different temporal scales. The large scale is termed periods lasting between 500 to 900 yrs. Superimposed were short term phases of aeolian activity of up to 200 yrs. During the last five ka, four periods of stronger and four periods 143

of weaker aeolian accumulation were recorded on the north-eastern TP (Fig. 5.2.4, Tab. 5.2.1). All periods except the last one can be found in the complete dataset and in the individual datasets from the three catchment areas with a similar timing. The youngest period (I) started at 1250 CE and ended at around 1725 CE. The end of the period at 1725 CE is related to the low number of young ages in the dataset, preventing any more recent analysis. The second period (II) was from 50 to 950 CE, a third one (III) from 1400 to ~500 BCE and the oldest one (IV) from 2500 to ~2000 BCE. Periods I and II are based on 18 and 24 OSL ages, period III and IV on 13 and 5 OSL ages, respectively. The low number of OSL ages for periods III and IV are considered critical because they did not pass the 90% quantile threshold (Fig. 5.2.3). Reworking of the sediments might obfuscate the real timing of deposition. Consequently, the timing of the two older periods should be regarded as preliminary.

In the two younger periods, several short term phases with durations of 80 to 200 years were evident in the cumulative probability curve. They ranged from 1630–1725 CE (I.1), 1450–1530 CE (I.2) and 1250–1350 CE (I.3) in the first period, and 750–950 CE (II.1), 390–540 CE (II.2) and 50–225 CE (II.3) in the second period. All six phases in the cumulative PDF passed the 90% quantile threshold. In each of the three regions three phases were above the threshold. The periods and the short term phases were interpreted as periods of strong aeolian accumulation on the north-eastern TP. Short term phases were not evident in the dataset during the two older phases. This was caused by larger error bars which blur shorter events. For phase I and II the mean one sigma error was ~ 40 years (n: 19) and ~ 125 years (n: 27), respectively. For phase III and IV it was around 200 (n: 13) and 300 years (n: 5).

Tab. 5.2.1: Periods (marked in grey) and short term phase of enhanced aeolian accumulation on the north-eastern TP. The last two periods III and IV are below the 90% quantile.

periods/phases [CE]

I 1250 1725

I.1 1630 1725

I.2 1450 1530

I.3 1250 1350

II 50 950

II.1 750 950

II.2 390 540

II.3 50 225

III -1400 -500

IV -2500 -2000

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The timing of the periods and phases in the three basins is relatively similar. The larger periods are reflected in all basins, except period IV which is not expressed in the Qinghai Lake area. More regional differences are expressed in the short term phases. While most of the phases are evident in the ages from the Eastern Qaidam Basin and the Donggi Cona catchment, there is only a weak concurrence in the Qinghai Lake Basin. In the Qinghai Lake Basin the main accumulation event occurred at around 1000 CE. This accumulation event consist of four samples which are from aeolian sand covering shoreline deposits (Liu et al., 2011) and from dune sediments east of the Lake (Lu et al., 2015). In general in all three areas the number of ages from aeolian sand increases towards the late Holocene, especially during the last 2 ka. While during period IV four of five ages and during period III still six out of 13 ages are from loess deposits, aeolian sands are dominating period I and II. However, in most of the six phases loess and sand have a similar timing in the peaks of the pdf (Fig. 5.2.4). This is especially evident in the Eastern Qaidam Basin with the largest number of OSL ages from loess.

Fig. 5.2.4: OSL ages and cumulative probability functions from the north-eastern Tibetan Plateau. Black line: pdf for all OSL ages; white solid line: sand ages; white dashed line: loess ages. A: complete north-eastern TP, B: Qinghai Lake area, C: Eastern Qaidam Basin, D: Donggi Cona. Light gray shading indicates periods of enhanced aeolian accumulation; dark gray shading indicates short term phases of high aeolian accumulation. Dashed blocks mark the timespan 1750-2000 CE that was not analyzed due to the low number of ages. Periods before 0 CE (2 ka) did not passed the confidence test.

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Discussion

Paleoclimatic factors governing the aeolian accumulation on the north-eastern Tibetan Plateau during the last five ka

The increase in ages from aeolian sand and the relative decrease of OSL ages from loess deposits was frequently related to a decline in the strength of the Asian summer monsoon during the late Holocene. A previous summary indicates enhanced aeolian sand accumulation in the different basin during the late Holocene. In the Qinghai Lake Basin the enhanced accumulation started at around 2ka, in the Qaidam Basin at 3.6 ka and at the Donggi Cona since 3ka (Stauch, 2015). Similar results are available from the Chinese Loess Plateau, where reduced palaeosol formation indicates drier climate conditions since at least 3.4 ka (H. Wang et al., 2014). This timing was also recorded in other archives from northern China (e.g. Herzschuh, 2006; Chen et al., 2015). As documented by the mainly parallel timing of the sediment deposition of loess and sand both sediment types have a similar reaction on short term regional environmental changes. However, loess accumulation in the different basins is heterogeneous in time. This can be attributed to only temporally available sources for silty sediments on the north-eastern Tibetan Plateau (Stauch, 2015). In contrast, exposed lake sediments in the western Qaidam Basin are a continuous source for loess deposits in the eastern Qaidam Basin.

OSL ages from the north-eastern TP indicate two periods of stronger aeolian accumulation during the last two ka and probably four during the last five ka. Additionally, six multi-decadal scale phases of strongest accumulation events occurred during the last two ka. Period I matches to the period of the LIA (Little Ice Age, 1250–1850 CE, Wanner et al., 2011). The LIA is associated with a remarkable drop in temperatures especially on the northern hemisphere (Wanner et al., 2008). In many regions glaciers advanced (Solomina et al., 2015). Similarly, the second period covers the so-called dark ages cooling (300–600 CE, Ljungqvist, 2010). Low sediment accumulation occurred on the north-eastern TP from 950 to 1250 CE, the Medieval Climate Anomaly (MCA, Mann et al., 2009). Consequently, aeolian accumulation on the north-eastern TP during the last 2000 years fits relatively well to the reconstructed northern hemisphere temperatures (Moberg et al., 2005; Christiansen and Ljungqvist, 2012). In general, enhanced aeolian accumulation on the north-eastern TP occurred during longer phases of low extratropical northern hemisphere temperatures. Especially the accumulation phases I.2, I.3 and II.1 exactly match the temperature depressions (Fig. 5.2.5A). The temperature depression during phase I.1 started around 50 years earlier. However, an increase in the cumulative pdf is already evident during this time. Low aeolian activity during the MCA lags the peak temperature around 100 years. No consistent relationship can be found from around 0 to 500 CE. Phase II.2 covers several cold and warm phases and during phase II.3 extratropical northern hemisphere temperature is relatively high. Temperature is, beside other external and internal drivers, closely related to the total solar irradiance (TSI) (Moberg et al., 2005; Wanner et al., 2011). Periods of low TSI values during the last 1500 years (Steinhilber et al., 2012) match phases of higher aeolian accumulation, except at around ~850 CE. High aeolian activity is recorded during the gran solar minima at ~1700 CE (Maunder Minimum), ~1500 CE (Spörer Minimum), and ~1300 CE (Wolf Minimum) (Fig. 5.2.5A). Several authors highlight a correlation between insolation and the Asian Monsoon system (Fleitmann et al., 2003; Y. Wang et al., 2005; Tan et al., 2008; Steinhilber et al., 2012), the main source of precipitation on the north-eastern TP.

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Fig. 5.2.5: OSL ages and paleoclimate datasets. A: northern hemisphere extratropical temperature anomalies (Christiansen and Ljungqvist, 2012), grey: original data with 1 yr resolution, solid line: 100 yrs average), dashed line: pdf of aeolian sediments, black triangles mark grand solar minima (M: Maunder; S: Spörer; W: Wolf); B: δ18O record from the Wanxian Cave (Zhang et al., 2008), grey original data with 2.5 yrs resolution, red: 100 yrs average; C: reconstructed precipitation from tree rings in the Qilian Shan (B. Yang et al., 2014); D: stacked record of hematite stained grains from the North Atlantic (Bond et al., 2001).

Two high resolution precipitation records near the study area are used for a detailed comparison. The δ18O record from the Wanxian Cave (Zhang et al., 2008), about 400 km SE of the north-eastern TP, has been interpreted as a signal for the strength of the Asian summer monsoon (Fig. 5.2.5B). In contrast to other caves in southern and eastern China it is located relatively close to the north- eastern TP. The high resolution speleothem δ18O record showed a good correlation with measured precipitation data for the last 50 years (J. Liu et al., 2008). Low δ18O values are supposed to reflect phases of higher monsoon intensity while higher values (less negative) indicate a weakening of the Asian summer monsoon (Zhang et al., 2008). Low aeolian activity corresponds to phases of an enhanced Asian summer monsoon, especially at the beginning of the record from around 200 to 1000 CE. Additionally, an increase in the number of OSL ages from around 1300 to 1700 CE reflects the increased aeolian accumulation caused by reduced monsoonal precipitation. However, the aeolian record does not reflect the high frequency fluctuations of the speleothem record starting at around 1300 CE, but the general trend of a reduced monsoonal influence is reflected. In contrast to 147

the δ18O record from Wanxian, the reconstructed mean annual precipitation from tree rings in the Qilian Shan for the last 3.5 ka (B. Yang et al., 2014) is in concurrence with the aeolian record for the last 1300 years (Fig. 5.2.5C). From around 1800 to 1300 CE high aeolian accumulation is in accordance with low annual precipitation values. A cold and dry LIA and warm and wet MCA were also recorded in other archives from monsoonal central Asia (Holmes et al., 2009; Henderson et al., 2010; J. Chen et al., 2015) and in aeolian archives on the north-eastern Tibetan Plateau (IJmker et al., 2012c; Yu et al., 2015). However, for the central Qaidam Basin a dry and warm MCA and a wet and cold LIA was assumed based on the stabilisation of dune sediments (Yu et al., 2013). Reduced precipitation at 1150 CE based on the reconstruction from tree rings is not reflected in the OSL dataset. Additionally, phase II.1 is characterized by an increase in precipitation in the tree ring data, which is contrary to the data from Wanxian and the enhanced aeolian accumulation. The last two phases (II.2 and II.3) include high frequency variations of precipitation values in the Qilian Shan. These precipitation fluctuations are more rapid than in the younger times and are not fully captured by the aeolian system.

According to the results presented above, aeolian activity on the north-eastern TP during the late Holocene is related to the intensity of the Asian Summer Monsoon. Specifically, phases of high aeolian activity occurred during phases of a weaker monsoonal system with low precipitation values. The relationship between monsoonal intensity and aeolian activity during the course of the Holocene has been recognized in several prior studies (e.g. Liu et al., 2011; An et al., 2012; Yu and Lai, 2014; Qiang et al., 2014). However, recent studies highlighted, that the interpretation of OSL ages from aeolian sand is not straightforward. OSL ages from aeolian sands with an early to mid-Holocene age are related to the formation of a suitable sediment trap and therefore wetter climate conditions (Stauch et al., 2012; Yu and Lai, 2012), while late Holocene OSL ages are related to the remobilization of aeolian sediments (Stauch, 2015).

Climate variations on the TP have been frequently linked to changes in the North Atlantic regions via the influence of the westerlies (Henderson et al., 2010; An et al., 2012; Zhu et al., 2015). Aeolian activity in the area is strongest during spring times (Derbyshire et al., 1998; Ta et al., 2004; Qiang et al., 2007; Han et al., 2008; Mao et al., 2011). Roe (2009) linked the occurrence of dust-storms in spring to the breakdown of the Siberian anticyclone and the shift of the westerlies to the north of the TP. These effects result in an enhanced cyclogenesis and maximized aeolian transport at the surface. During times of a weaker monsoonal system this phase might be prolonged, leading to strong aeolian transport and enhanced accumulation. Additionally, less monsoonal precipitation and lower temperatures result in reduced vegetation cover and a shorter growing season, enlarging the surface which is susceptible for aeolian erosion. Strong monsoon would result in the opposite effect. Recent climatological studies show a close connection between the North Atlantic and the TP (e.g. Liu and Yin, 2001; Ding et al., 2005; Molg et al., 2014; Maussion et al., 2013; Curio et al., 2015) through propagation of wave trains via the atmospheric Rosby waves (Ding and Wang, 2005; Bothe et al., 2010, 2012). Therefore, enhanced westerlies might contribute to the formation of aeolian deposits on the north-eastern TP.

However, the connection to the North Atlantic in paleoclimate records is mainly based on correlation of different proxies with the so-called Holocene Bond events. These Bond events are based on an increased percentage of hematite-stained grains in marine cores from the North Atlantic and were interpreted as drift ice indices (Bond et al., 1997, 2001). During the course of the Holocene nine drift ice events were recorded. Events with a similar timing have been frequently described from the TP, 148

e.g., in glacial records from western TP (Owen, 2009; Seong et al., 2009), lake sediments (Ji et al., 2005; Liu et al., 2009; Qiang et al., 2014), peat bogs (Hong et al., 2003; B. Liu et al., 2014), and aeolian records (Yu et al., 2006; Liu et al., 2013b). The record of the drift ice indices during the last 2000 years (and the last 5000 years) matches the record of the aeolian activity remarkably well (Fig. 5.2.5D), despite the low resolution and the weak chronology of the drift ice indices (Trouet et al., 2012). The phases of enhanced aeolian accumulation are nearly identical to the Bond events, except peak I.1 and II.3. The good match between the two datasets might be at least partly accidentally. A comparison of the OSL ages with the reconstructed Atlantic Multidecadal Oscillation (AMO) by Mann et al. (2009) revealed only few commonalities. The AMO reflects sea surface temperature (SST) anomalies in the North Atlantic region which have been linked to changes in the Atlantic Meridional Overturning Circulation (AMOC) (e.g. Knudsen et al., 2014). A similar result was achieved when comparing the aeolian record with divergent North Atlantic Oscillation (NAO) reconstructions from Iceland (Geirsdóttir et al., 2013), Greenland (Olsen et al., 2012), speleothems from Scotland (Baker et al., 2015) and a tree-ring-based reconstruction from the North Atlantic area (Trouet et al., 2009). In contrast, recent climate models indicated a link between changes in the AMO and eastern China (Li and Bates, 2007). However, changes in the North Atlantic Region might be less effective for aeolian accumulation phases than monsoonal changes (J. Chen et al., 2015).

Evaluation of the paleo-climatic significance of the aeolian sediments from the north-eastern Tibetan Plateau

Summarizing the different reconstructions for temperature and precipitation mentioned above indicates that the aeolian activity on the north-eastern TP is well in accordance with other climate proxies on the regional up to the hemispheric scale for the last 2000 years. The best results are achieved for the last 1300 years, with only minor deviations from other records. There is probably a lag of around 100 years between the warmest and wettest conditions during the MCA and lowest aeolian activity on the north-eastern TP. From 0 to 700 CE the concurrence between the aeolian record and the different climate reconstructions are no longer explicit. This might be due to a high frequency of environmental changes as indicated by the tree-ring record from the Qilian Shan, which are not fully captured by the aeolian sediments, to the reworking of sediments during later periods or to local environmental changes. No confident signal was archived for the time from 3000 to 0 BCE due to the low number of OSL ages in combination with larger errors of the individual dating results. However, for the last 2000 years aeolian sediments can be used as proxy for multi-decadal scale climate variations if a sufficiently large number of OSL ages are available. This allows an analysis of aeolian sediments regrading palaeoclimate changes even in an area with frequent reworking. Additionally, the synchronous occurrence of phases of enhanced aeolian accumulation in the different regions indicates a common regional climatic signal in the aeolian sediments. Under the present environmental conditions the semi-arid environments of the north-eastern TP are not sediment supply limited and they are suitable to record such changes. It can be assumed that during full monsoon conditions until the end of the middle Holocene sediment supply was a limiting factor for aeolian transport due to denser vegetation cover, preventing an analysis of short term climate fluctuation. An increased human influence in the area (e.g. F. Chen et al., 2015a; Yu et al., 2015) might have influenced the intensity of the aeolian processes. However, the general agreement of timing of the phases and periods of enhanced aeolian accumulation with the regional climate signals indicates no significant human influence on the timing of the aeolian activity. 149

Conclusions

The aeolian record from the north-eastern TP based on the collection of OSL ages reflects multi- decadal to millennial-scale climatic changes on the north-eastern TP. During the last 5000 years, four periods of enhanced aeolian accumulation lasting between 500 and 900 years were identified. However, only for the two periods during the last 2000 years the results are robust. Six short term phases of strongest aeolian accumulation are superimposed on the last two periods. These phases last from 1630 to 1725, 1450–1530, 1250–1350, 750–950, 390–540 and 50–225 CE. They match well with phases of global and hemispheric temperature and local precipitation changes. Strong aeolian accumulation on the north-eastern TP occurred during periods of low temperature and a weak Asian Summer Monsoon. No clear concurrence could be found regarding the often highlighted influence from the North Atlantic area. While there is a good agreement to the drift ice indices from the North Atlantic, other high resolution records indicate no clear match on the multi-decadal to multi- centennial timescale. It can be assumed that the Asian Summer Monsoon is the main influencing factor for aeolian accumulation on the TP.

Acknowledgements

Intensive discussion with Frank Lehmkuhl, David Loibl, Veit Nottebaum, Wolfgang Römer and Philipp Schulte greatly improved the evaluation of the dataset. Christian Zeeden supported the statistical analysis of the OSL ages. Nick Lancaster and an anonymous reviewer provided helpful comments on an earlier version of this manuscript.

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6 Conclusions

The aim of the research presented here was to improve the understanding of the aeolian sedimentary system for palaeoclimate reconstructions during the late glacial to the Holocene on the northern TP. For this aim three lake catchment areas on the northern TP, the Gonghe, Donggi Cona, and Heihai catchment, were studied in detail. In contrast to previous work on the TP, sedimentary deposits and related landforms were mapped on a basin-wide scale. This enabled a detailed understanding of the sediment cascade in the different basins.

Each of the studied basins on the northern Tibetan Plateau is a unique sedimentary system with a large variety of influencing factors. Some of these factors are not directly related to the paleoclimatic changes but have a strong influence on aeolian accumulation. A typical example is sediment availability. In the Donggi Cona catchment a lake level drop in the late Holocene leads to a massive reactivation of sands which were previously deposited at the lake floor and subsequently exposed. In consequence, accumulation rates east of the lake increased by several magnitudes. Geochemical analysis of the aeolian deposits revealed the occurrence of a high carbonate content originating from the lake. Such signals of high accumulation can easily be misinterpreted without proper knowledge of the basin-wide processes. Similar pitfalls arise from the dunes on the northern side of Lake Heihai, which have a large component in the coarse sand fraction. However, this large component is not only related to the high wind speed in the area but also to the local geology. Large amounts of sands are released by neighbouring conglomerates with a matrix of coarse sands due to strong weathering processes.

Another main result of the local studies is that nearly all aeolian sections on the northern Tibetan Plateau are affected by erosion resulting in incomplete records at individual sites. This point is especially important as only few studies consider discontinuity of the sections as problematic. However, due to locally differentiated erosion, the combination of a large number of sections enables the reconstruction for most parts of the Holocene. Furthermore, local topography is an important boundary condition. Beside sediment availability, local topography is the most important factor controlling the distribution and timing of deposition of the different types of aeolian sediments. This was clearly demonstrated for all three study areas. It is especially obvious in the Gonghe Basin, where loess accumulation started earlier than the permanent sand accumulation due to an earlier increase in moisture in the upper parts of the basin. This is contrary to most other areas where permanent sand accumulation preceded the loess deposits. These factors highlight the need for basin-wide mapping and sampling of geo-archives to understand local and regional transportation processes and, in consequence, the palaeoclimatic evolution in the area.

Taking these limitations into account, the accumulation of aeolian sediments is a useful paleoclimate proxy on the northern Tibetan Plateau. This is especially important as additional proxies besides lake sediments are needed in the area. As outlined before, the majority of proxy records from the northern TP are derived from lake sediments, which are affected by a number of problems related to proxy interpretation and chronology. The OSL method provides a direct measure for the age of the accumulation of the sediments. Including the studies presented here more than 500 OSL ages from the TP were published since 2000 CE. However, meta-studies summarizing all available data were not conducted previously. Therefore, two meta-studies were conducted based on the compiled dataset; one for the time since the late glacial and the complete Holocene and one for the late Holocene. However, the interpretation of OSL ages from aeolian sediments is not that straightforward, 151

especially in a semiarid region like the TP and a new conceptual model for the interpretation of OSL ages from aeolian sediments on the northern TP was developed. In general, aeolian sediments are transported under dry climate conditions, provided sufficient wind speed and sediment availability. Therefore, aeolian sediments in general, and especially aeolian sands, are used as an indicator of an arid climate. Nevertheless, an obvious, but not trivial prerequisite for the analysis of aeolian sediments is that the sediments are preserved sampling. Under dry climate conditions in the area of deposition reworking of aeolian sediments is frequently occurring. In a semi-arid to arid environment like the TP aeolian sediments are mainly preserved from the end of an aeolian period. Increasing humidity is resulting in the initial formation of a vegetation cover, which, in turn, protects the aeolian archive. This is the main difference between the environmental conditions on the TP and the Chinese loess plateau. On the loess plateau, the climate conditions supported the formation of a vegetation cover even during glacial times (W. Jiang et al., 2013) and the influx rate of the aeolian sediments is basically governed by the sediment availability in the source areas. Aeolian transport on the TP, as shown in the three study areas, is dominated by short distance transport. Aeolian sediment supply, accumulation and preservation in the area are governed by the same local environmental conditions.

As a shift to more humid conditions is not instantaneous in the whole area, aeolian activity and related aeolian accumulation ceased with increasing humidity through time. This is especially important for phases at the end of long dry periods like glacial times. In contrast, OSL ages from the late Holocene indeed show the reactivation of aeolian sediments due to more arid climate conditions because the recent dry period is not as dry as the glacial times. Nevertheless, the curvelinear distribution of OSL ages with a lower number of ages from the beginning of the dry period and a strong increase during end of the late Holocene is a result of the reworking of the sediments. Consequently, the record from the beginning of the late Holocene arid phase will be removed sometimes in the future, assuming continuously dry climate conditions in the area. At the Heihai, this process is already nearly completed and only few late glacial and early Holocene aeolian sediments are preserved. Following this theoretical model, the interpretation of OSL ages from aeolian sediments depends on external knowledge about the previous climate state. However, this concept does not exclude the preservation of short term dry periods. This was shown either by the preservation of the Younger Dryas period in the Gonghe and Heihai Basin as well as several short term accumulation phases during the late Holocene. Additionally, differences exist between different sediment types. Sandy sediments are mainly located in the central parts of the basins, while loess sediments are deposited at the hills surrounding the basins. Due to the different environmental conditions, aeolian sands are more prone to aeolian remobilization due to higher wind-speeds. In contrast, loess sediments in the more humid west of the study area are predominantly eroded by fluvial processes as demonstrated in the Donggi Cona catchment.

The main results of the meta-studies and the detailed analysis of the sediment cascades in the three study areas resulted in a profound palaeoclimatic reconstruction for the northern TP. Aeolian sediments indicate a first increase in precipitation after extremely dry climate conditions during the glacial. However, a time-transgressive onset was observed across the northern TP. The oldest preserved aeolian sediments are from the Qinghai Lake Basin in the eastern part of the study area. In this region the strengthening of the ASM started already at around 17 ka, which is in accordance with pollen studies from the lake but not with many other studies. In the neighbouring Gonghe Basin loess sediments have OSL ages of up to 15 ka, while aeolian sand ages are around 12.5 ka. At the Donggi Cona, sand accumulation started in the early Holocene at around 10.5 ka. In contrast to that, only

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few aeolian sediments were preserved in the western-most study area, the Heihai catchment, due to more arid climate conditions. During the late glacial and the Holocene, environmental conditions were not suitable in the area for the constant preservation of aeolian sediments except at topographically shielded positions. At the end of the early Holocene and especially during the mid- Holocene, aeolian transport processes were strongly reduced on the north-eastern TP. Exceptions are the Heihai area and the eastern Qaidam Basin where loess accumulation mainly occurred during the mid-Holocene. However, both areas have a similar arid environment. The reduced aeolian activity in the east and trapping of the sediments in the west reflect the maximum moisture period on the northern TP, the so-called Holocene optimum. This is further supported by fluvial reworking of aeolian sediments in the Gonghe and Donggi Cona basin. During the late Holocene all areas on the northern TP experienced the reactivation of aeolian sands. Reactivation started earliest in the eastern Qaidam Basin and at the Heihai at around 3.5 to 3.4 ka, at around 3 ka at the Donggi Cona, at 2 ka in the Gonghe Basin and at 1.9 ka at the Qinghai Lake. Again a time-transgressive reduction across the northern TP is observed. The variable onset and reduction of the ASM was described for the first time in terrestrial archives on the northern TP. Due to the existing data gaps in the central and western TP, an analysis of whole TP is not possible at the moment. However, the summary of OSL ages from aeolian sediments on the southern TP already indicates a different trend in comparison with the northern TP.

Besides the millennial-scale analysis of Holocene trends in OSL ages from aeolian sediments, a multi- decadal analysis of the aeolian sediments of the last 2 ka was conducted for the first time. For this study with an unpreceded temporal resolution from non-continuous sediment sections probability density functions were adopted. Previous research in the area with a similar timeframe always relied on interpolations between individual dating points, which introduce severe uncertainties to the chronology and the timing of the individual phases. During the last 2 ka six phases of enhanced aeolian accumulation were identified on the north-eastern TP. These phases were linked to phases of low temperature of the northern hemisphere, reduced total solar irradiance and a reduced ASM. A human influence on the timing of the phases could not be confirmed. However, it can be expected that they influence the strength of the phases. The study showed that short term climate fluctuations with length of 80 to 200 years can be reconstructed from aeolian sediments. However, due to the relatively low number of OSL ages older than 2 ka and the increasing uncertainties the method is not suitable to detect older short term climate fluctuations at the moment.

Outlook

In general, aeolian accumulation on the northern TP is now relatively well understood. Due to the related case studies from the different basins and the now large number of OSL ages, the processes and the timing relevant for a paleoclimate interpretation can be reconstructed. First trends in the timing of the aeolian accumulation on different time scales are obvious. However, on large parts of the TP detailed studies of aeolian sediments and their timing are missing. This prevents the identification of trends across the plateau. It can be expected that differences in the timing of sediment accumulation occur to the south and the west of the TP. A first indication is the different timing of aeolian sediment accumulation in the northern and the southern part of the TP. However, whether there is a gradual or rapid shift remains speculative. Furthermore, the transition from the monsoon dominated area in the east to the more westerlies influenced regions in the west awaits 153

intensive studies. This is especially important as proxy interpretations from lake sediments resulted in divergent palaeoclimate interpretations. However, further field studies on the TP and especially in the Province of Tibet are hampered by administrative restrictions and the harsh environmental conditions. Nevertheless, up to now even detailed studies utilizing remote sensing analysis of aeolian landforms are missing for the large parts of the TP. Such studies can be used for the identification of possible source areas and would further improve the understanding aeolian transport processes. A further important task is the consideration of the human influence on the landscape system. The TP has been used by humans since several thousand years. The number of studies related to archaeological excavations on the TP is rising fast during the last five years. Further research is now able to analyse local influences on the scale of individual basins, at least on the north-eastern TP.

154

Acknowledgments

As always scientific research is not based on the work of one person alone. Many colleagues and quite often friends contributed to the results of this thesis.

I would like to thank Frank Lehmkuhl for introducing me to this wonderful topic of aeolian sediments on the Tibetan Plateau. I first started to admire the landscape in the area during an expedition in the summer of 2006. Since that year I spend a considerable time of each summer in the area. I also have to thank Frank for the ongoing support and intensive discussion in all those years.

Special thanks are devoted to all my colleagues at the chair of Physical Geography and Geoecology and only a few can be named here. I would like to start with the two PhD candidates I have been working together on the Tibetan Plateau, Janneke IJmker and Philipp Schulte. Despite intensive work in the remote areas for several weeks, we always had stimulating discussions and a generally good time. Many people supported me in giving me advice or by just listening to my ideas; many thanks to David Loibl, Veit Nottebaum, Wolfgang Römer and Christian Zeeden. I am especially in debt to Marianne Dohms from the sedimentological laboratory in Aachen and our secretary Anja Knops for doing all the administrative work which comes with big projects.

Additionally, many colleagues from the DFG Project ‘Landscape and Lake-System Response to Late Quaternary Monsoon Dynamics on the Tibetan Plateau - Northern Transect’ joined the field work. They were Bernhard (The Guitar) Diekmann, Elisabeth Dietze, Kai Hartmann, Carolina Müller, Stephan Opitz, Steffen Pötsch, Arne Ramisch, Bernd Wünnemann, Dada Yan, Hui Zhao and many more Chinese and German students.

Finally, I would like to thank my wife Cirsten and my daughter Mia for their unconditional support during all those times I spend on the Tibetan Plateau and in my office.

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References

Adamiec, G., Aitken, M., 1998. Dose-rate conversion factors: update. Ancient TL 16, 37–50. Aichner, B., Herzschuh, U., Wilkes, H., Schulz, H.-M., Wang, Y., Plessen, B., Mischke, S., Diekmann, B., Zhang, C., 2012. Ecological development of Lake Donggi Cona, north-eastern Tibetan Plateau, since the late glacial on basis of organic geochemical proxies and non-pollen palynomorphs. Palaeogeography, Palaeoclimatology, Palaeoecology 313–314, 140–149. doi:10.1016/j.palaeo.2011.10.015 Aichner, B., Herzschuh, U., Wilkes, H., Vieth, A., Böhner, J., 2010a. δD values of n-alkanes in Tibetan lake sediments and aquatic macrophytes – A surface sediment study and application to a 16 ka record from Lake Koucha. Organic Geochemistry 41, 779–790. doi:10.1016/j.orggeochem.2010.05.010 Aichner, B., Wilkes, H., Herzschuh, U., Mischke, S., Zhang, C., 2010b. Biomarker and compound- specific δ13C evidence for changing environmental conditions and carbon limitation at Lake Koucha, eastern Tibetan Plateau. Journal of Paleolimnology 43, 873–899. doi:10.1007/s10933-009-9375-y Aitken, M., 1998. An Introduction to Optical Dating: The Dating of Quaternary Sediments by the Use of Photon-stimulated Luminescence. Oxford University Press, Oxford. Aldenderfer, M., Zhang, Y., 2004. The Prehistory of the Tibetan Plateau to the Seventh Century A.D.: Perspectives and Research from China and the West Since 1950. Journal of World Prehistory 18, 1–55. doi:10.1023/B:JOWO.0000038657.79035.9e An, Z., 2000. The history and variability of the East Asian paleomonsoon climate. Quaternary Science Reviews 19, 171–187. doi:10.1016/S0277-3791(99)00060-8 An, Z., Clemens, S.C., Shen, J., Qiang, X., Jin, Z., Sun, Y., Prell, W.L., Luo, J., Wang, S., Xu, H., Cai, Y., Zhou, W., Liu, X., Liu, W., Shi, Z., Yan, L., Xiao, X., Chang, H., Wu, F., Ai, L., Lu, F., 2011. Glacial- Interglacial Indian Summer Monsoon Dynamics. Science 333, 719. doi:10.1126/science.1203752 An, Z., Colman, S.M., Zhou, W., Li, X., Brown, E.T., Jull, A.J.T., Cai, Y., Huang, Y., Lu, X., Chang, H., Song, Y., Sun, Y., Xu, H., Liu, W., Jin, Z., Liu, X., Cheng, P., Liu, Y., Ai, L., Li, X., Liu, X., Yan, L., Shi, Z., Wang, X., Wu, F., Qiang, X., Dong, J., Lu, F., Xu, X., 2012. Interplay between the Westerlies and Asian monsoon recorded in Lake Qinghai sediments since 32 ka. Scientific Reports 2, 619. An, Z., Kukla, G., Porter, S.C., Xiao, J., 1991. Late quaternary dust flow on the chinese Loess Plateau. CATENA 18, 125–132. doi:10.1016/0341-8162(91)90012-M An, Z., Kutzbach, J.E., Prell, W.L., Porter, S.C., 2001. Evolution of Asian monsoons and phased uplift of the Himalaya-Tibetan plateau since Late Miocene times. Nature 411, 62–66. doi:10.1038/35075035 An, Z., Liu, T., Lu, Y., Porter, S.C., Kukla, G., Wu, X., Hua, Y., 1990. The long-term paleomonsoon variation recorded by the loess-paleosol sequence in Central China. Quaternary International 7, 91–95. doi:10.1016/1040-6182(90)90042-3 An, Z., Porter, S.C., Kutzbach, J.E., Xihao, W., Suming, W., Xiaodong, L., Xiaoqiang, L., Weijian, Z., 2000. Asynchronous Holocene optimum of the East Asian monsoon. Quaternary Science Reviews 19, 743–762. doi:10.1016/S0277-3791(99)00031-1 An, Z., Wu, G., Li, J., Sun, Y., Liu, Y., Zhou, W., Cai, Y., Duan, A., Li, L., Mao, J., Cheng, H., Shi, Z., Tan, L., Yan, H., Ao, H., Chang, H., Feng, J., 2015. Global Monsoon Dynamics and Climate Change. Annu. Rev. Earth Planet. Sci. 43, 29–77. doi:10.1146/annurev-earth-060313-054623 Auclair, M., Lamothe, M., Huot, S., 2003. Measurement of anomalous fading for feldspar IRSL using SAR. Radiation Measurements 37, 487–492. doi:10.1016/S1350-4487(03)00018-0 Bagnold, R.A., 1941. The Physics of Blown Sand and Desert Dunes. Chapman & Hall, London. Bailey, R.M., Thomas, D.S.G., 2014. A quantitative approach to understanding dated dune stratigraphies. Earth Surf. Process. Landforms 39, 614–631. doi:10.1002/esp.3471

156

Baker, A., C. Hellstrom, J., Kelly, B.F.J., Mariethoz, G., Trouet, V., 2015. A composite annual-resolution stalagmite record of North Atlantic climate over the last three millennia. Scientific Reports 5, 10307. Bateman, M.D., Frederick, C.D., Jaiswal, M.K., Singhvi, A.K., 2003a. Investigations into the potential effects of pedoturbation on luminescence dating. Quaternary Science Reviews 22, 1169– 1176. doi:10.1016/S0277-3791(03)00019-2 Bateman, M.D., Thomas, D.S.., Singhvi, A.K., 2003b. Extending the aridity record of the Southwest Kalahari: current problems and future perspectives. Quaternary International 111, 37–49. doi:10.1016/S1040-6182(03)00013-2 Berger, A., Loutre, M.F., 1991. Insolation values for the climate of the last 10 million years. Quaternary Science Reviews 10, 297–317. doi:10.1016/0277-3791(91)90033-Q Berkelhammer, M., Sinha, A., Mudelsee, M., Cheng, H., Edwards, R.L., Cannariato, K., 2010. Persistent multidecadal power of the Indian Summer Monsoon. Earth and Planetary Science Letters 290, 166–172. doi:10.1016/j.epsl.2009.12.017 Bird, B.W., Polisar, P.J., Lei, Y., Thompson, L.G., Yao, T., Finney, B.P., Bain, D.J., Pompeani, D.P., Steinman, B.A., 2014. A Tibetan lake sediment record of Holocene Indian summer monsoon variability. Earth and Planetary Science Letters 399, 92–102. doi:10.1016/j.epsl.2014.05.017 Böhner, J., 2006. General climatic controls and topoclimatic variations in Central and High Asia. Boreas 35, 279–295. doi:10.1111/j.1502-3885.2006.tb01158.x Bond, G., Kromer, B., Beer, J., Muscheler, R., Evans, M.N., Showers, W., Hoffmann, S., Lotti-Bond, R., Hajdas, I., Bonani, G., 2001. Persistent Solar Influence on North Atlantic Climate During the Holocene. Science 294, 2130–2136. doi:10.1126/science.1065680 Bond, G., Showers, W., Cheseby, M., Lotti, R., Almasi, P., deMenocal, P., Priore, P., Cullen, H., Hajdas, I., Bonani, G., 1997. A Pervasive Millennial-Scale Cycle in North Atlantic Holocene and Glacial Climates. Science 278, 1257–1266. doi:10.1126/science.278.5341.1257 Boos, W.R., Kuang, Z., 2010. Dominant control of the South Asian monsoon by orographic insulation versus plateau heating. Nature 463, 218–222. doi:10.1038/nature08707 Bothe, O., Fraedrich, K., Zhu, X., 2012. Tibetan Plateau summer precipitation: covariability with circulation indices. Theoretical and Applied Climatology 108, 293–300. doi:10.1007/s00704- 011-0538-1 Bothe, O., Fraedrich, K., Zhu, X., 2010. The large-scale circulations and summer drought and wetness on the Tibetan plateau. Int. J. Climatol. 30, 844–855. doi:10.1002/joc.1946 Bøtter-Jensen, L., Mejdahl, V., Murray, A.S., 1999. New light on OSL. Quaternary Science Reviews 18, 303–309. doi:10.1016/S0277-3791(98)00063-8 Braconnot, P., Otto-Bliesner, B., Harrison, S., Joussaume, S., Peterchmitt, J.-Y., Abe-Ouchi, A., Crucifix, M., Driesschaert, E., Fichefet, T., Hewitt, C.D., Kageyama, M., Kitoh, A., Laîné, A., Loutre, M.- F., Marti, O., Merkel, U., Ramstein, G., Valdes, P., Weber, S.L., Yu, Y., Zhao, Y., 2007. Results of PMIP2 coupled simulations of the Mid-Holocene and Last Glacial Maximum – Part 1: experiments and large-scale features. Clim. Past 3, 261–277. doi:10.5194/cp-3-261-2007 Brantingham, P.J., Ma, H., Olsen, J.W., Gao, X., Madsen, D.B., Rhode, D.E., 2003. Speculation on the timing and nature of Late Pleistocene hunter-gatherer colonization of the Tibetan Plateau. Chinese Science Bulletin 48, 1510–1516. doi:10.1360/02wd0276 Bronk Ramsey, C., 2009. Bayesian Analysis of Radiocarbon Dates. Radiocarbon; Vol 51, No 1 (2009). Bullard, J.., Wiggs, G.F.., Nash, D.., 2000. Experimental study of wind directional variability in the vicinity of a model valley. Geomorphology 35, 127–143. doi:10.1016/S0169-555X(00)00033-7 Bullard, J.E., 2013. Contemporary glacigenic inputs to the dust cycle. Earth Surf. Process. Landforms 38, 71–89. doi:10.1002/esp.3315 Bullard, J.E., 1997. Vegetation and dryland geomorphology, in: Thomas, D.S.G. (Ed.), Arid Zone Geomorphology; Process, Form and Change in Drylands. Wiley, Chichester, pp. 109–131. Bullard, J.E., Livingstone, I., 2002. Interactions between aeolian and fluvial systems in dryland environments. Area 34, 8–16. doi:10.1111/1475-4762.00052

157

Bullard, J.E., McTainsh, G.H., 2003. Aeolian-fluvial interactions in dryland environments: examples, concepts and Australia case study. Progress in Physical Geography 27, 471–501. doi:10.1191/0309133303pp386ra Bulur, E., 1996. An alternative technique for optically stimulated luminescence (OSL) experiment. Radiation Measurements 26, 701–709. doi:10.1016/S1350-4487(97)82884-3 Bulur, E., Bøtter-Jensen, L., Murray, A.., 2000. Optically stimulated luminescence from quartz measured using the linear modulation technique. Radiation Measurements 32, 407–411. doi:10.1016/S1350-4487(00)00115-3 Buylaert, J.-P., Jain, M., Murray, A.S., Thomsen, K.J., Thiel, C., Sohbati, R., 2012. A robust feldspar luminescence dating method for Middle and Late Pleistocene sediments. Boreas 41, 435– 451. doi:10.1111/j.1502-3885.2012.00248.x Buylaert, J.P., Murray, A.S., Vandenberghe, D., Vriend, M., De Corte, F., Van den haute, P., 2008. Optical dating of Chinese loess using sand-sized quartz: Establishing a time frame for Late Pleistocene climate changes in the western part of the Chinese Loess Plateau. Quaternary Geochronology 3, 99–113. doi:10.1016/j.quageo.2007.05.003 Buylaert, J.P., Vandenberghe, D., Murray, A.S., Huot, S., De Corte, F., Van den Haute, P., 2007. Luminescence dating of old (>70 ka) Chinese loess: A comparison of single-aliquot OSL and IRSL techniques. Quaternary Geochronology 2, 9–14. doi:10.1016/j.quageo.2006.05.028 Cai, Y., Tan, L., Cheng, H., An, Z., Edwards, R.L., Kelly, M.J., Kong, X., Wang, X., 2010. The variation of summer monsoon precipitation in central China since the last deglaciation. Earth and Planetary Science Letters 291, 21–31. doi:10.1016/j.epsl.2009.12.039 Caley, T., Malaizé, B., Zaragosi, S., Rossignol, L., Bourget, J., Eynaud, F., Martinez, P., Giraudeau, J., Charlier, K., Ellouz-Zimmermann, N., 2011. New Arabian Sea records help decipher orbital timing of Indo-Asian monsoon. Earth and Planetary Science Letters 308, 433–444. doi:10.1016/j.epsl.2011.06.019 Caley, T., Roche, D.M., Renssen, H., 2014. Orbital Asian summer monsoon dynamics revealed using an isotope-enabled global climate model. Nat Commun 5. Chase, B., 2009. Evaluating the use of dune sediments as a proxy for palaeo-aridity: A southern African case study. Earth-Science Reviews 93, 31–45. doi:10.1016/j.earscirev.2008.12.004 Chase, B.M., Thomas, D.S.G., 2007. Multiphase late Quaternary aeolian sediment accumulation in western South Africa: Timing and relationship to palaeoclimatic changes inferred from the marine record. Quaternary International 166, 29–41. doi:10.1016/j.quaint.2006.12.005 Chen, F., Chen, J.-H., Holmes, J., Boomer, I., Austin, P., Gates, J.B., Wang, N.-L., Brooks, S.J., Zhang, J.- W., 2010. Moisture changes over the last millennium in arid central Asia: a review, synthesis and comparison with monsoon region. Quaternary Science Reviews 29, 1055–1068. doi:10.1016/j.quascirev.2010.01.005 Chen, F., Dong, G.H., Zhang, D.J., Liu, X.Y., Jia, X., An, C.B., Ma, M.M., Xie, Y.W., Barton, L., Ren, X.Y., Zhao, Z.J., Wu, X.H., Jones, M.K., 2015a. Agriculture facilitated permanent human occupation of the Tibetan Plateau after 3600 B.P. Science 347, 248–250. doi:10.1126/science.1259172 Chen, F., Wu, D., Chen, J., Zhou, A., Yu, J., Shen, J., Wang, S., Huang, X., 2016. Holocene moisture and East Asian summer monsoon evolution in the northeastern Tibetan Plateau recorded by Lake Qinghai and its environs: A review of conflicting proxies. Quaternary Science Reviews 154, 111–129. doi:10.1016/j.quascirev.2016.10.021 Chen, F., Xu, Q., Chen, J., Birks, H.J.B., Liu, J., Zhang, S., Jin, L., An, C., Telford, R.J., Cao, X., Wang, Z., Zhang, X., Selvaraj, K., Lu, H., Li, Y., Zheng, Z., Wang, H., Zhou, A., Dong, G., Zhang, J., Huang, X., Bloemendal, J., Rao, Z., 2015b. East Asian summer monsoon precipitation variability since the last deglaciation. Scientific Reports 5, 11186. Chen, F., Yu, Z., Yang, M., Ito, E., Wang, S., Madsen, D.B., Huang, X., Zhao, Y., Sato, T., John B. Birks, H., Boomer, I., Chen, J., An, C., Wünnemann, B., 2008. Holocene moisture evolution in arid central Asia and its out-of-phase relationship with Asian monsoon history. Quaternary Science Reviews 27, 351–364. doi:10.1016/j.quascirev.2007.10.017

158

Chen, J., Chen, F., Feng, S., Huang, W., Liu, J., Zhou, A., 2015. Hydroclimatic changes in China and surroundings during the Medieval Climate Anomaly and Little Ice Age: spatial patterns and possible mechanisms. Quaternary Science Reviews 107, 98–111. doi:10.1016/j.quascirev.2014.10.012 Chen, W., Graf, H.-F., Huang, R., 2000. The interannual variability of East Asian Winter Monsoon and its relation to the summer monsoon. Advances in Atmospheric Sciences 17, 48–60. doi:10.1007/s00376-000-0042-5 Chen, Y., Li, Y., Zhang, Y., Zhang, M., Zhang, J., Yi, C., Liu, G., 2011. Late Quaternary deposition and incision sequences of the Golmud River and their environmental implications. Quaternary International 236, 48–56. doi:10.1016/j.quaint.2010.05.028 Cheng, H., Edwards, R.L., Broecker, W.S., Denton, G.H., Kong, X., Wang, Y., Zhang, R., Wang, X., 2009. Ice Age Terminations. Science 326, 248–252. doi:10.1126/science.1177840 Cheng, H., Edwards, R.L., Wang, Y., Kong, X., Ming, Y., Kelly, M.J., Wang, X., Gallup, C.D., Liu, W., 2006. A penultimate glacial monsoon record from Hulu Cave and two-phase glacial terminations. Geology 34, 217–220. doi:10.1130/G22289.1 Cheng, H., Sinha, A., Wang, X., Cruz, F.W., Edwards, R.L., 2012. The Global Paleomonsoon as seen through speleothem records from Asia and the Americas. Climate Dynamics 39, 1045–1062. doi:10.1007/s00382-012-1363-7 Chiverrell, R.C., Thorndycraft, V.R., Hoffmann, T.O., 2011. Cumulative probability functions and their role in evaluating the chronology of geomorphological events during the Holocene. J. Quaternary Sci. 26, 76–85. doi:10.1002/jqs.1428 Christiansen, B., Ljungqvist, F.C., 2012. The extra-tropical Northern Hemisphere temperature in the last two millennia: reconstructions of low-frequency variability. Clim. Past 8, 765–786. doi:10.5194/cp-8-765-2012 Clark, P.U., Dyke, A.S., Shakun, J.D., Carlson, A.E., Clark, J., Wohlfarth, B., Mitrovica, J.X., Hostetler, S.W., McCabe, A.M., 2009. The Last Glacial Maximum. Science 325, 710–714. doi:10.1126/science.1172873 Clarke, M.L., Rendell, H.M., 1998. Climate change impacts on sand supply and the formation of desert sand dunes in the south-west U.S.A. Journal of Arid Environments 39, 517–531. doi:10.1006/jare.1997.0372 Clemens, S., Prell, W., Murray, D., Shimmield, G., Weedon, G., 1991. Forcing mechanisms of the Indian Ocean monsoon. Nature 353, 720–725. doi:10.1038/353720a0 Clemens, S.C., Prell, W.L., 2003. A 350,000 year summer-monsoon multi-proxy stack from the Owen Ridge, Northern Arabian Sea. Marine Geology 201, 35–51. doi:10.1016/S0025- 3227(03)00207-X Clemens, S.C., Prell, W.L., Sun, Y., 2010. Orbital-scale timing and mechanisms driving Late Pleistocene Indo-Asian summer monsoons: Reinterpreting cave speleothem δ18O. Paleoceanography 25, PA4207. doi:10.1029/2010PA001926 Clift, P.D., Plumb, R.A., 2008. The Asian monsoon : causes, history and effects. Cambridge University Press, Cambridge. Colman, S.M., Yu, S.-Y., An, Z., Shen, J., Henderson, A.C.G., 2007. Late Cenozoic climate changes in China’s western interior: a review of research on Lake Qinghai and comparison with other records. Quaternary Science Reviews 26, 2281–2300. doi:10.1016/j.quascirev.2007.05.002 Craddock, W.H., Kirby, E., Harkins, N.W., Zhang, H., Shi, X., Liu, J., 2010. Rapid fluvial incision along the Yellow River during headward basin integration. Nature Geosci 3, 209–213. doi:10.1038/ngeo777 Cui, X., Graf, H.-F., 2009. Recent land cover changes on the Tibetan Plateau: a review. Climatic Change 94, 47–61. doi:10.1007/s10584-009-9556-8 Curio, J., Maussion, F., Scherer, D., 2015. A 12-year high-resolution climatology of atmospheric water transport over the Tibetan Plateau. Earth Syst. Dynam. 6, 109–124. doi:10.5194/esd-6-109- 2015

159

Daut, G., Mäusbacher, R., Baade, J., Gleixner, G., Kroemer, E., Mügler, I., Wallner, J., Wang, J., Zhu, L., 2010. Late Quaternary hydrological changes inferred from lake level fluctuations of Nam Co (Tibetan Plateau, China). Quaternary International 218, 86–93. doi:10.1016/j.quaint.2010.01.001 Davidson-Arnott, R.G.D., Yang, Y., Ollerhead, J., Hesp, P.A., Walker, I.J., 2008. The effects of surface moisture on aeolian sediment transport threshold and mass flux on a beach. Earth Surf. Process. Landforms 33, 55–74. doi:10.1002/esp.1527 de Vernal, A., Hillaire-Marcel, C., 2006. Provincialism in trends and high frequency changes in the northwest North Atlantic during the Holocene. Global and Planetary Change 54, 263–290. doi:10.1016/j.gloplacha.2006.06.023 Denton, G.H., Karlén, W., 1973. Holocene climatic variations—Their pattern and possible cause. Quaternary Research 3, 155–205. doi:10.1016/0033-5894(73)90040-9 Derbyshire, E., Meng, X., Kemp, R.A., 1998. Provenance, transport and characteristics of modern aeolian dust in western Gansu Province, China, and interpretation of the Quaternary loess record. Journal of Arid Environments 39, 497–516. doi:10.1006/jare.1997.0369 Dietze, E., Hartmann, K., Diekmann, B., IJmker, J., Lehmkuhl, F., Opitz, S., Stauch, G., Wünnemann, B., Borchers, A., 2012. An end-member algorithm for deciphering modern detrital processes from lake sediments of Lake Donggi Cona, NE Tibetan Plateau, China. Sedimentary Geology 243–244, 169–180. doi:10.1016/j.sedgeo.2011.09.014 Dietze, E., Maussion, F., Ahlborn, M., Diekmann, B., Hartmann, K., Henkel, K., Kasper, T., Lockot, G., Opitz, S., Haberzettl, T., 2014. Sediment transport processes across the Tibetan Plateau inferred from robust grain-size end members in lake sediments. Clim. Past 10, 91–106. doi:10.5194/cp-10-91-2014 Dietze, E., Wünnemann, B., Diekmann, B., Aichner, B., Hartmann, K., Herzschuh, U., IJmker, J., Jin, H., Kopsch, C., Lehmkuhl, F., Li, S., Mischke, S., Niessen, F., Opitz, S., Stauch, G., Yang, S., 2010. Basin morphology and seismic stratigraphy of Lake Donggi Cona, north-eastern Tibetan Plateau, China. Quaternary International 218, 131–142. doi:10.1016/j.quaint.2009.11.035 Dietze, E., Wünnemann, B., Hartmann, K., Diekmann, B., Jin, H., Stauch, G., Yang, S., Lehmkuhl, F., 2013. Early to mid-Holocene lake high-stand sediments at Lake Donggi Cona, northeastern Tibetan Plateau, China. Quaternary Research 79, 325–336. doi:10.1016/j.yqres.2012.12.008 Dietze, M., Kreutzer, S., Burow, C., Fuchs, M.C., Fischer, M., Schmidt, C., 2016. The abanico plot: Visualising chronometric data with individual standard errors. Quaternary Geochronology 31, 12–18. doi:10.1016/j.quageo.2015.09.003 Ding, Q., Wang, B., 2005. Circumglobal Teleconnection in the Northern Hemisphere Summer. J. Climate 18, 3483–3505. doi:10.1175/JCLI3473.1 Ding, R., Li, J., Wang, S., Ren, F., 2005. Decadal change of the spring dust storm in northwest China and the associated atmospheric circulation. Geophys. Res. Lett. 32, n/a-n/a. doi:10.1029/2004GL021561 Ding, Y., Chan, L.J.C., 2005. The East Asian summer monsoon: an overview. Meteorology and Atmospheric Physics 89, 117–142. doi:10.1007/s00703-005-0125-z Ding, Y., Sun, Y., Wang, Z., Zhu, Y., Song, Y., 2009. Inter-decadal variation of the summer precipitation in China and its association with decreasing Asian summer monsoon Part II: Possible causes. Int. J. Climatol. 29, 1926–1944. doi:10.1002/joc.1759 Ding, Y., Wang, Z., Sun, Y., 2008. Inter-decadal variation of the summer precipitation in East China and its association with decreasing Asian summer monsoon. Part I: Observed evidences. International Journal of Climatology 28, 1139–1161. doi:10.1002/joc.1615 Ding, Z., Liu, T., Rutter, N.W., Yu, Z., Guo, Z., Zhu, R., 1995. Ice-Volume Forcing of East Asian Winter Monsoon Variations in the Past 800,000 Years. Quaternary Research 44, 149–159. doi:10.1006/qres.1995.1059 Ding, Z., Yu, Z., Rutter, N.W., Liu, T., 1994. Towards an orbital time scale for chinese loess deposits. Quaternary Science Reviews 13, 39–70. doi:10.1016/0277-3791(94)90124-4

160

Doberschütz, S., Frenzel, P., Haberzettl, T., Kasper, T., Wang, J., Zhu, L., Daut, G., Schwalb, A., Mäusbacher, R., 2014. Monsoonal forcing of Holocene paleoenvironmental change on the central Tibetan Plateau inferred using a sediment record from Lake Nam Co (Xizang, China). Journal of Paleolimnology 51, 253–266. doi:10.1007/s10933-013-9702-1 Dong, G., Liu, H., Yang, Y., Yang, Y., Zhou, A., Wang, Z., Ren, X., Chen, F., 2016. Emergence of ancient cities in relation to geopolitical circumstances and climate change during late Holocene in northeastern Tibetan Plateau, China. Frontiers of Earth Science 1–14. doi:10.1007/s11707- 015-0575-7 Duller, G.A.T., 2016. Challenges involved in obtaining luminescence ages for long records of aridity: Examples from the Arabian Peninsula. Quaternary International 410, Part B, 69–74. doi:10.1016/j.quaint.2016.01.028 Duller, G.A.T., 2008. Single-grain optical dating of Quaternary sediments: why aliquot size matters in luminescence dating. Boreas 37, 589–612. doi:10.1111/j.1502-3885.2008.00051.x Duller, G.A.T., 2004. Luminescence dating of quaternary sediments: recent advances. J. Quaternary Sci. 19, 183–192. doi:10.1002/jqs.809 Duller, G.A.T., 2003. Distinguishing quartz and feldspar in single grain luminescence measurements. Radiation Measurements 37, 161–165. doi:10.1016/S1350-4487(02)00170-1 Dykoski, C.A., Edwards, R.L., Cheng, H., Yuan, D., Cai, Y., Zhang, M., Lin, Y., Qing, J., An, Z., Revenaugh, J., 2005. A high-resolution, absolute-dated Holocene and deglacial Asian monsoon record from Dongge Cave, China. Earth and Planetary Science Letters 233, 71–86. doi:10.1016/j.epsl.2005.01.036 Fang, X., Lü, L., Mason, J.A., Yang, S., An, Z., Li, J., Zhilong, G., 2003. Pedogenic response to millennial summer monsoon enhancements on the Tibetan Plateau. Quaternary International 106–107, 79–88. doi:10.1016/S1040-6182(02)00163-5 FAO, 2006. Guidelines for Soil Description. Rome. Farr, T.G., Rosen, P.A., Caro, E., Crippen, R., Duren, R., Hensley, S., Kobrick, M., Paller, M., Rodriguez, E., Roth, L., Seal, D., Shaffer, S., Shimada, J., Umland, J., Werner, M., Oskin, M., Burbank, D., Alsdorf, D., 2007. The Shuttle Radar Topography Mission. Rev. Geophys. 45, n/a-n/a. doi:10.1029/2005RG000183 Fécan, F., Marticorena, B., Bergametti, G., 1998. Parametrization of the increase of the aeolian erosion threshold wind friction velocity due to soil moisture for arid and semi-arid areas. Annales Geophysicae 17, 149–157. doi:10.1007/s00585-999-0149-7 Feng, J., Wang, T., Qi, S., Xie, C., 2005. Land degradation in the source region of the Yellow River, northeast Qinghai-Xizang Plateau: classification and evaluation. Environmental Geology 47, 459–466. doi:10.1007/s00254-004-1161-6 Field, J.P., Breshears, D.D., Whicker, J.J., 2009. Toward a more holistic perspective of soil erosion: Why aeolian research needs to explicitly consider fluvial processes and interactions. Aeolian Research 1, 9–17. doi:10.1016/j.aeolia.2009.04.002 Fielding, E., Isacks, B., Barazangi, M., Duncan, C., 1994. How flat is Tibet? Geology 22, 163–167. doi:10.1130/0091-7613(1994)022<0163:HFIT>2.3.CO;2 Filzmoser, P., Hron, K., Reimann, C., Garrett, R., 2009. Robust factor analysis for compositional data. Computers & Geosciences 35, 1854–1861. doi:10.1016/j.cageo.2008.12.005 Fitzsimmons, K.E., Cohen, T.J., Hesse, P.P., Jansen, J., Nanson, G.C., May, J.-H., Barrows, T.T., Haberlah, D., Hilgers, A., Kelly, T., Larsen, J., Lomax, J., Treble, P., 2013. Late Quaternary palaeoenvironmental change in the Australian drylands. Quaternary Science Reviews 74, 78– 96. doi:10.1016/j.quascirev.2012.09.007 Fleitmann, D., Burns, S.J., Mudelsee, M., Neff, U., Kramers, J., Mangini, A., Matter, A., 2003. Holocene Forcing of the Indian Monsoon Recorded in a Stalagmite from Southern Oman. Science 300, 1737–1739. doi:10.1126/science.1083130 Fleitmann, D., Burns, S.J., Pekala, M., Mangini, A., Al-Subbary, A., Al-Aowah, M., Kramers, J., Matter, A., 2011. Holocene and Pleistocene pluvial periods in Yemen, southern Arabia. Quaternary Science Reviews 30, 783–787. doi:10.1016/j.quascirev.2011.01.004

161

Flohn, H., 1957. Large-scale aspects of the “summer monsoon” in south and east Asia. Journal of the Meterological Society of Japan 75, 180–186. Fu, B., Awata, Y., 2007. Displacement and timing of left-lateral faulting in the Kunlun Fault Zone, northern Tibet, inferred from geologic and geomorphic features. Journal of Asian Earth Sciences 29, 253–265. doi:10.1016/j.jseaes.2006.03.004 Galbraith, R., 2010. On plotting OSL equivalent doses. Ancient TL 28, 1–9. Galbraith, R.F., Roberts, R.G., 2012. Statistical aspects of equivalent dose and error calculation and display in OSL dating: An overview and some recommendations. Quaternary Geochronology 11, 1–27. doi:10.1016/j.quageo.2012.04.020 Galbraith, R.F., Roberts, R.G., Laslett, G.M., Yoshida, H., Olley, J.M., 1999. Optical dating of single and multiple grains of quartz from Jinmium Rock Shelter, Northern Australia: Part I, Experimental design and statistical models. Archaeometry 41, 339–364. doi:10.1111/j.1475- 4754.1999.tb00987.x Geirsdóttir, Á., Miller, G.H., Larsen, D.J., Ólafsdóttir, S., 2013. Abrupt Holocene climate transitions in the northern North Atlantic region recorded by synchronized lacustrine records in Iceland. Quaternary Science Reviews 70, 48–62. doi:10.1016/j.quascirev.2013.03.010 Goble, R.J., Mason, J.A., Loope, D.B., Swinehart, J.B., 2004. Optical and radiocarbon ages of stacked paleosols and dune sands in the Nebraska Sand Hills, USA. Quaternary Science Reviews 23, 1173–1182. doi:10.1016/j.quascirev.2003.09.009 Günther, F., Witt, R., Schouten, S., Mäusbacher, R., Daut, G., Zhu, L., Xu, B., Yao, T., Gleixner, G., 2015. Quaternary ecological responses and impacts of the Indian Ocean Summer Monsoon at Nam Co, Southern Tibetan Plateau. Quaternary Science Reviews 112, 66–77. doi:10.1016/j.quascirev.2015.01.023 Guo, J., Lin, A., Sun, G., Zheng, J., 2007. Surface Ruptures Associated with the 1937 M 7.5 Tuosuo Lake and the 1963 M 7.0 Alake Lake Earthquakes and the Paleoseismicity along the Tuosuo Lake Segment of the Kunlun Fault, Northern Tibet. Bulletin of the Seismological Society of America 97, 474–496. doi:10.1785/0120050103 Guo, Z.T., Ruddiman, W.F., Hao, Q.Z., Wu, H.B., Qiao, Y.S., Zhu, R.X., Peng, S.Z., Wei, J.J., Yuan, B.Y., Liu, T.S., 2002. Onset of Asian desertification by 22[thinsp]Myr ago inferred from loess deposits in China. Nature 416, 159–163. doi:10.1038/416159a Guo, Z.T., Sun, B., Zhang, Z.S., Peng, S.Z., Xiao, G.Q., Ge, J.Y., Hao, Q.Z., Qiao, Y.S., Liang, M.Y., Liu, J.F., Yin, Q.Z., Wei, J.J., 2008. A major reorganization of Asian climate by the early Miocene. Climate of the Past 4, 153–174. doi:10.5194/cp-4-153-2008 Haberzettl, T., Henkel, K., Kasper, T., Ahlborn, M., Su, Y., Wang, J., Appel, E., St-Onge, G., Stoner, J., Daut, G., Zhu, L., Mäusbacher, R., 2015. Independently dated paleomagnetic secular variation records from the Tibetan Plateau. Earth and Planetary Science Letters 416, 98–108. doi:10.1016/j.epsl.2015.02.007 Hahn, D.G., Manabe, S., 1975. The Role of Mountains in the South Asian Monsoon Circulation. J. Atmos. Sci. 32, 1515–1541. doi:10.1175/1520-0469(1975)032<1515:TROMIT>2.0.CO;2 Halfen, A.F., Lancaster, N., Wolfe, S., 2016. Interpretations and common challenges of aeolian records from North American dune fields. Quaternary International 410, Part B, 75–95. doi:10.1016/j.quaint.2015.03.003 Han, Y., Fang, X., Kang, S., Wang, H., Kang, F., 2008. Shifts of dust source regions over central Asia and the Tibetan Plateau: Connections with the Arctic oscillation and the westerly jet. Atmospheric Environment 42, 2358–2368. doi:10.1016/j.atmosenv.2007.12.025 Hao, M., Shen, Z.-K., Wang, Q., Cui, D., 2012. Postseismic deformation mechanisms of the 1990 Mw 6.4 Gonghe, China earthquake constrained using leveling measurements. Tectonophysics 532–535, 205–214. doi:10.1016/j.tecto.2012.02.005 Harris, N., 2006. The elevation history of the Tibetan Plateau and its implications for the Asian monsoon. Palaeogeography, Palaeoclimatology, Palaeoecology 241, 4–15. doi:10.1016/j.palaeo.2006.07.009

162

Henderson, A.C.G., Holmes, J.A., 2009. Palaeolimnological evidence for environmental change over the past millennium from Lake Qinghai sediments: A review and future research prospective. Quaternary International 194, 134–147. doi:10.1016/j.quaint.2008.09.008 Henderson, A.C.G., Holmes, J.A., Leng, M.J., 2010. Late Holocene isotope hydrology of Lake Qinghai, NE Tibetan Plateau: effective moisture variability and atmospheric circulation changes. Quaternary Science Reviews 29, 2215–2223. doi:10.1016/j.quascirev.2010.05.019 Henkel, K., Haberzettl, T., St-Onge, G., Wang, J., Ahlborn, M., Daut, G., Zhu, L., Mäusbacher, R., 2016. High-resolution paleomagnetic and sedimentological investigations on the Tibetan Plateau for the past 16 ka cal B.P.—The Tangra Yumco record. Geochem. Geophys. Geosyst. 17, 774– 790. doi:10.1002/2015GC006023 Herzschuh, U., 2006. Palaeo-moisture evolution in monsoonal Central Asia during the last 50,000 years. Quaternary Science Reviews 25, 163–178. doi:10.1016/j.quascirev.2005.02.006 Herzschuh, U., Birks, H.J.B., Mischke, S., Zhang, C., Böhner, J., 2010a. A modern pollen–climate calibration set based on lake sediments from the Tibetan Plateau and its application to a Late Quaternary pollen record from the Qilian Mountains. Journal of Biogeography 37, 752–766. doi:10.1111/j.1365-2699.2009.02245.x Herzschuh, U., Kramer, A., Mischke, S., Zhang, C., 2009. Quantitative climate and vegetation trends since the late glacial on the northeastern Tibetan Plateau deduced from Koucha Lake pollen spectra. Quaternary Research 71, 162–171. doi:10.1016/j.yqres.2008.09.003 Herzschuh, U., Kürschner, H., Mischke, S., 2006. Temperature variability and vertical vegetation belt shifts during the last ∼50,000 yr in the Qilian Mountains (NE margin of the Tibetan Plateau, China). Quaternary Research 66, 133–146. doi:10.1016/j.yqres.2006.03.001 Herzschuh, U., Mischke, S., Meyer, H., Plessen, B., Zhang, C., 2010b. Lake nutrient variability inferred from elemental (C, N, S) and isotopic (δ13C, δ15N) analyses of aquatic plant macrofossils. Quaternary Science Reviews 29, 2161–2172. doi:10.1016/j.quascirev.2010.05.011 Herzschuh, U., Ni, J., Birks, H.J.B., Böhner, J., 2011. Driving forces of mid-Holocene vegetation shifts on the upper Tibetan Plateau, with emphasis on changes in atmospheric CO2 concentrations. Quaternary Science Reviews 30, 1907–1917. doi:10.1016/j.quascirev.2011.03.007 Herzschuh, U., Zhang, C., Mischke, S., Herzschuh, R., Mohammadi, F., Mingram, B., Kürschner, H., Riedel, F., 2005. A late Quaternary lake record from the Qilian Mountains (NW China): evolution of the primary production and the water depth reconstructed from macrofossil, pollen, biomarker, and isotope data. Global and Planetary Change 46, 361–379. doi:10.1016/j.gloplacha.2004.09.024 Hesse, P.P., 2016. How do longitudinal dunes respond to climate forcing? Insights from 25 years of luminescence dating of the Australian desert dunefields. Quaternary International 410, Part B, 11–29. doi:10.1016/j.quaint.2014.02.020 Hetzel, R., 2013. Active faulting, mountain growth, and erosion at the margins of the Tibetan Plateau constrained by in situ-produced cosmogenic nuclides. Tectonophysics 582, 1–24. doi:10.1016/j.tecto.2012.10.027 Heyman, J., Stroeven, A.P., Caffee, M.W., Hättestrand, C., Harbor, J.M., Li, Y., Alexanderson, H., Zhou, L., Hubbard, A., 2011. Palaeoglaciology of Bayan Har Shan, NE Tibetan Plateau: exposure ages reveal a missing LGM expansion. Quaternary Science Reviews 30, 1988–2001. doi:10.1016/j.quascirev.2011.05.002 Hildebrandt, S., Müller, S., Kalugin, I.A., Dar’in, A.V., Wagner, M., Rogozin, D.Y., Tarasov, P.E., 2015. Tracing the North Atlantic decadal-scale climate variability in a late Holocene pollen record from southern Siberia. Palaeogeography, Palaeoclimatology, Palaeoecology 426, 75–84. doi:10.1016/j.palaeo.2015.02.037 Hilgers, A., 2007. The Chronology of Late Glacial and Holocene Dune Development in the Northern European Lowland Reconstructed by Optical Stimulated Luminescence (OSL) Dating. Ph.D. thesis, University of Cologne.

163

Hoffmann, T., Lang, A., Dikau, R., 2008. Holocene river activity: analysing 14C-dated fluvial and colluvial sediments from Germany. Quaternary Science Reviews 27, 2031–2040. doi:10.1016/j.quascirev.2008.06.014 Holmes, J.A., Cook, E.R., Yang, B., 2009. Climate change over the past 2000 years in Western China. Quaternary International 194, 91–107. doi:10.1016/j.quaint.2007.10.013 Hong, Y.T., Hong, B., Lin, Q.H., Zhu, Y.X., Shibata, Y., Hirota, M., Uchida, M., Leng, X.T., Jiang, H.B., Xu, H., Wang, H., Yi, L., 2003. Correlation between Indian Ocean summer monsoon and North Atlantic climate during the Holocene. Earth and Planetary Science Letters 211, 371–380. doi:10.1016/S0012-821X(03)00207-3 Hou, J., D’Andrea, W.J., Liu, Z., 2012. The influence of 14C reservoir age on interpretation of paleolimnological records from the Tibetan Plateau. Quaternary Science Reviews 48, 67–79. doi:10.1016/j.quascirev.2012.06.008 Hou, W., Dong, H., Li, G., Yang, J., Coolen, M.J.L., Liu, X., Wang, S., Jiang, H., Wu, X., Xiao, H., Lian, B., Wan, Y., 2014. Identification of Photosynthetic Plankton Communities Using Sedimentary Ancient DNA and Their Response to late-Holocene Climate Change on the Tibetan Plateau. Scientific Reports 4, 6648. Hövermann, J., 1987. Morphogenetic regions in Northeast Xizang (Tibet), in: Hövermann, J., Wang, W. (Eds.), Reports of the Qinghai-Xizang (Tibet) Plateau. Science Press, Beijing, pp. 112–139. Hu, G., Dong, Z., Lu, J., Yan, C., 2015. The developmental trend and influencing factors of aeolian desertification in the Zoige Basin, eastern Qinghai–Tibet Plateau. Aeolian Research 19, Part B, 275–281. doi:10.1016/j.aeolia.2015.02.002 Huang, R., 1987. Vegetation of the periglacial zone in the northeastern part of the Qinghai-Xizang Plateau, in: Hövermann, J., Wang, W. (Eds.), Reports of the Qinghai-Xizang (Tibet) Plateau. Science Press, Beijing, pp. 438–495. Hudson, A.M., Quade, J., 2013. Long-term east-west asymmetry in monsoon rainfall on the Tibetan Plateau. Geology 41, 351–354. doi:10.1130/G33837.1 Hülle, D., Hilgers, A., Radtke, U., Stolz, C., Hempelmann, N., Grunert, J., Felauer, T., Lehmkuhl, F., 2010. OSL dating of sediments from the Gobi Desert, Southern Mongolia. Quaternary Geochronology 5, 107–113. doi:10.1016/j.quageo.2009.06.002 Huntley, D.J., Baril, M.R., 1997. The K content of the K-feldspars being measured in optical dating or in thermoluminescence dating. Ancient TL 15/1, 11–13. Huntley, D.J., Lamothe, M., 2001. Ubiquity of anomalous fading in K-feldspars and the measurement and correction for it in optical dating. Canadian Journal of Earth Sciences 38, 1093–1106. doi:10.1139/e01-013 IJmker, J., Stauch, G., Dietze, E., Hartmann, K., Diekmann, B., Lockot, G., Opitz, S., Wünnemann, B., Lehmkuhl, F., 2012a. Characterisation of transport processes and sedimentary deposits by statistical end-member mixing analysis of terrestrial sediments in the Donggi Cona lake catchment, NE Tibetan Plateau. Sedimentary Geology 281, 166–179. doi:10.1016/j.sedgeo.2012.09.006 IJmker, J., Stauch, G., Hartmann, K., Diekmann, B., Dietze, E., Opitz, S., Wünnemann, B., Lehmkuhl, F., 2012b. Environmental conditions in the Donggi Cona lake catchment, NE Tibetan Plateau, based on factor analysis of geochemical data. Journal of Asian Earth Sciences 44, 176–188. doi:10.1016/j.jseaes.2011.04.021 IJmker, J., Stauch, G., Pötsch, S., Diekmann, B., Wünnemann, B., Lehmkuhl, F., 2012c. Dry periods on the NE Tibetan Plateau during the late Quaternary. Palaeogeography, Palaeoclimatology, Palaeoecology 346–347, 108–119. doi:10.1016/j.palaeo.2012.06.005 Immerzeel, W.W., van Beek, L.P.H., Bierkens, M.F.P., 2010. Climate change will affect the Asian water towers. Science 328, 1382–1385. doi:10.1126/science.1183188 Jennings, A.E., Knudsen, K.L., Hald, M., Hansen, C.V., Andrews, J.T., 2002. A mid-Holocene shift in Arctic sea-ice variability on the East Greenland Shelf. The Holocene 12, 49–58. doi:10.1191/0959683602hl519rp

164

Ji, J., Shen, J., Balsam, W., Chen, J., Liu, L., Liu, X., 2005. Asian monsoon oscillations in the northeastern Qinghai–Tibet Plateau since the late glacial as interpreted from visible reflectance of Qinghai Lake sediments. Earth and Planetary Science Letters 233, 61–70. doi:10.1016/j.epsl.2005.02.025 Jiang, D., Lang, X., Tian, Z., Ju, L., 2013. Mid-Holocene East Asian summer monsoon strengthening: Insights from Paleoclimate Modeling Intercomparison Project (PMIP) simulations. Palaeogeography, Palaeoclimatology, Palaeoecology 369, 422–429. doi:10.1016/j.palaeo.2012.11.007 Jiang, W., Cheng, Y., Yang, X., Yang, S., 2013. Chinese Loess Plateau vegetation since the Last Glacial Maximum and its implications for vegetation restoration. J Appl Ecol 50, 440–448. doi:10.1111/1365-2664.12052 Jin, H.J., Chang, X.L., Wang, S.L., 2007. Evolution of permafrost on the Qinghai-Xizang (Tibet) Plateau since the end of the late Pleistocene. J. Geophys. Res. 112, n/a-n/a. doi:10.1029/2006JF000521 Jin, L., Chen, F., Morrill, C., Otto-Bliesner, B.L., Rosenbloom, N., 2012. Causes of early Holocene desertification in arid central Asia. Climate Dynamics 38, 1577–1591. doi:10.1007/s00382- 011-1086-1 Jin, Z., An, Z., Yu, J., Li, F., Zhang, F., 2015. Lake Qinghai sediment geochemistry linked to hydroclimate variability since the last glacial. Quaternary Science Reviews 122, 63–73. doi:10.1016/j.quascirev.2015.05.015 Johnstone, E., Macklin, M.G., Lewin, J., 2006. The development and application of a database of radiocarbon-dated Holocene fluvial deposits in Great Britain. CATENA 66, 14–23. doi:10.1016/j.catena.2005.07.006 Jones, A.F., Macklin, M.G., Benito, G., 2015. Meta-analysis of Holocene fluvial sedimentary archives: A methodological primer. CATENA 130, 3–12. doi:10.1016/j.catena.2014.11.018 Jones, M.D., Roberts, C.N., Leng, M.J., Türkeş, M., 2006. A high-resolution late Holocene lake isotope record from Turkey and links to North Atlantic and monsoon climate. Geology 34, 361–364. doi:10.1130/G22407.1 Joussaume, S., Taylor, K.E., Braconnot, P., Mitchell, J.F.B., Kutzbach, J.E., Harrison, S.P., Prentice, I.C., Broccoli, A.J., Abe-Ouchi, A., Bartlein, P.J., Bonfils, C., Dong, B., Guiot, J., Herterich, K., Hewitt, C.D., Jolly, D., Kim, J.W., Kislov, A., Kitoh, A., Loutre, M.F., Masson, V., McAvaney, B., McFarlane, N., de Noblet, N., Peltier, W.R., Peterschmitt, J.Y., Pollard, D., Rind, D., Royer, J.F., Schlesinger, M.E., Syktus, J., Thompson, S., Valdes, P., Vettoretti, G., Webb, R.S., Wyputta, U., 1999. Monsoon changes for 6000 years ago: Results of 18 simulations from the Paleoclimate Modeling Intercomparison Project (PMIP). Geophys. Res. Lett. 26, 859–862. doi:10.1029/1999GL900126 Kääb, A., Berthier, E., Nuth, C., Gardelle, J., Arnaud, Y., 2012. Contrasting patterns of early twenty- first-century glacier mass change in the Himalayas. Nature 488, 495–498. doi:10.1038/nature11324 Kaiser, H.F., 1958. The varimax criterion for analytic rotation in factor analysis. Psychometrika 23, 187–200. doi:10.1007/BF02289233 Kaiser, K., Lai, Z., Schneider, B., Junge, F.W., 2010. Late Pleistocene genesis of the middle Yarlung Zhangbo Valley, southern Tibet (China), as deduced by sedimentological and luminescence data. Quaternary Geochronology 5, 200–204. doi:10.1016/j.quageo.2009.01.005 Kaiser, K., Lai, Z., Schneider, B., Reudenbach, C., Miehe, G., Brückner, H., 2009a. Stratigraphy and palaeoenvironmental implications of Pleistocene and Holocene aeolian sediments in the Lhasa area, southern Tibet (China). Palaeogeography, Palaeoclimatology, Palaeoecology 271, 329–342. doi:10.1016/j.palaeo.2008.11.004 Kaiser, K., Opgenoorth, L., Schoch, W.H., Miehe, G., 2009b. Charcoal and fossil wood from palaeosols, sediments and artificial structures indicating Late Holocene woodland decline in southern Tibet (China). Quaternary Science Reviews 28, 1539–1554. doi:10.1016/j.quascirev.2009.02.016

165

Kaiser, K., Schoch, W.H., Miehe, G., 2007. Holocene paleosols and colluvial sediments in Northeast Tibet (Qinghai Province, China): Properties, dating and paleoenvironmental implications. CATENA 69, 91–102. doi:10.1016/j.catena.2006.04.028 Kapp, P., Pelletier, J.D., Rohrmann, A., Heermance, R., Russell, J., Ding, L., 2011. Wind erosion in the Qaidam basin, central Asia: Implications for tectonics, paleoclimate, and the source of the Loess Plateau. GSA Today 21, 4–10. doi:DOI: 10.1130/GSATG99A.1 Kathayat, G., Cheng, H., Sinha, A., Spötl, C., Edwards, R.L., Zhang, H., Li, X., Yi, L., Ning, Y., Cai, Y., Lui, W.L., Breitenbach, S.F.M., 2016. Indian monsoon variability on millennial-orbital timescales. Scientific Reports 6, 24374. Kidd, W.S.F., Yusheng, P., Chengfa, C., Coward, M.P., Dewey, J.F., Gansser, A., Molnar, P., Shackleton, R.M., Yiyin, S., 1988. Geological Mapping of the 1985 Chinese--British Tibetan (Xizang-- Qinghai) Plateau Geotraverse Route. Philos Trans R Soc Lond A 327, 287. doi:10.1098/rsta.1988.0130 Kirby, E., Harkins, N., Wang, E., Shi, X., Fan, C., Burbank, D., 2007. Slip rate gradients along the eastern Kunlun fault. Tectonics 26, n/a-n/a. doi:10.1029/2006TC002033 Klinge, M., Lehmkuhl, F., 2015. Holocene aeolian mantles and inter-bedded paleosols on the southern Tibetan Plateau. Quaternary International 372, 33–44. doi:10.1016/j.quaint.2014.06.068 Knudsen, M.F., Jacobsen, B.H., Seidenkrantz, M.-S., Olsen, J., 2014. Evidence for external forcing of the Atlantic Multidecadal Oscillation since termination of the Little Ice Age. Nat Commun 5. Kocurek, G., Lancaster, N., 1999. Aeolian system sediment state: theory and Mojave Desert Kelso dune field example. Sedimentology 46, 505–515. doi:10.1046/j.1365-3091.1999.00227.x Kohfeld, K.E., Harrison, S.P., 2003. Glacial-interglacial changes in dust deposition on the Chinese Loess Plateau. Quaternary Science Reviews 22, 1859–1878. doi:10.1016/S0277- 3791(03)00166-5 Kohfeld, K.E., Harrison, S.P., 2001. DIRTMAP: the geological record of dust. Earth-Science Reviews 54, 81–114. doi:10.1016/S0012-8252(01)00042-3 Kok, J.F., 2011. A scaling theory for the size distribution of emitted dust aerosols suggests climate models underestimate the size of the global dust cycle. Proceedings of the National Academy of Sciences 108, 1016–1021. doi:10.1073/pnas.1014798108 Kok, J.F., Parteli, E.J.R., Michaels, T.I., Bou Karam, D., 2012. The physics of wind-blown sand and dust. Reports on Progress in Physics 75. doi:/10.1088/0034-4885/75/10/106901 Kok, J.F., Renno, N.O., 2009. A comprehensive numerical model of steady state saltation (COMSALT). J. Geophys. Res. 114, n/a-n/a. doi:10.1029/2009JD011702 Kukla, G., 1987. Loess stratigraphy in central China. Quaternary Science Reviews 6, 191–219. doi:10.1016/0277-3791(87)90004-7 Kukla, G., An, Z., 1989. Loess stratigraphy in Central China. Palaeogeography, Palaeoclimatology, Palaeoecology 72, 203–225. doi:10.1016/0031-0182(89)90143-0 Kürschner, H., Herzschuh, U., Wagner, D., 2005. Phytosociological studies in the north-eastern Tibetan Plateau (NW China). A first contribution to the subalpine scrub and alpine meadow vegetation. Botanische Jahrbuecher 126, 273–315. Küster, Y., Hetzel, R., Krbetschek, M., Tao, M., 2006. Holocene loess sedimentation along the Qilian Shan (China): significance for understanding the processes and timing of loess deposition. Quaternary Science Reviews 25, 114–125. doi:10.1016/j.quascirev.2005.03.003 Kutzbach, J.E., 1981. Monsoon Climate of the Early Holocene: Climate Experiment with the Earth’s Orbital Parameters for 9000 Years Ago. Science 214, 59–61. doi:10.1126/science.214.4516.59 Lai, Z., 2010. Chronology and the upper dating limit for loess samples from Luochuan section in the Chinese Loess Plateau using quartz OSL SAR protocol. Journal of Asian Earth Sciences 37, 176–185. doi:10.1016/j.jseaes.2009.08.003 Lai, Z., 2006. Testing the use of an OSL standardised growth curve (SGC) for determination on quartz from the Chinese Loess Plateau. Radiation Measurements 41, 9–16. doi:10.1016/j.radmeas.2005.06.031

166

Lai, Z., Kaiser, K., Brückner, H., 2009. Luminescence-dated aeolian deposits of late Quaternary age in the southern Tibetan Plateau and their implications for landscape history. Quaternary Research 72, 421–430. doi:10.1016/j.yqres.2009.07.005 Lai, Z.-P., Brückner, H., Zöller, L., Fülling, A., 2007. Existence of a common growth curve for silt-sized quartz OSL of loess from different continents. Radiation Measurements 42, 1432–1440. doi:10.1016/j.radmeas.2007.08.006 Lai, Z.P., Zöller, L., Fuchs, M., Brückner, H., 2008. Alpha efficiency determination for OSL of quartz extracted from Chinese loess. Radiation Measurements 43, 767–770. doi:10.1016/j.radmeas.2008.01.022 Lancaster, N., 2008. Desert dune dynamics and development: insights from luminescence dating. Boreas 37, 559–573. doi:10.1111/j.1502-3885.2008.00055.x Lancaster, N., 1997. Response of eolian geomorphic systems to minor climate change: examples from the southern Californian deserts. Geomorphology 19, 333–347. doi:10.1016/S0169- 555X(97)00018-4 Lancaster, N., Baas, A., 1998. Influence of vegetation cover on sand transport by wind: field studies at Owens Lake, California. Earth Surf. Process. Landforms 23, 69–82. doi:10.1002/(SICI)1096- 9837(199801)23:1<69::AID-ESP823>3.0.CO;2-G Lang, A., 2003. Phases of soil erosion-derived colluviation in the loess hills of South Germany. CATENA 51, 209–221. doi:10.1016/S0341-8162(02)00166-2 Langford, R.P., 1989. Fluvial-aeolian interactions: Part I, modern systems. Sedimentology 36, 1023– 1035. doi:10.1111/j.1365-3091.1989.tb01540.x Lehmkuhl, F., 1998. Extent and spatial distribution of Pleistocene glaciations in eastern Tibet. Quaternary International 45, 123–134. doi:10.1016/S1040-6182(97)00010-4 Lehmkuhl, F., 1997. The spatial distribution of loess and loess-like sediments in the mountain areas of Central and High Asia. Zeitschrift für Geomorphologie Neue Folgen, Supplementbände 111, 97–116. Lehmkuhl, F., Böhner, J., Haselein, F., 1999. Late Quaternary environmental changes and human occupation of the Tibetan Plateau. Man and Environment 24, 137–148. Lehmkuhl, F., Haselein, F., 2000. Quaternary paleoenvironmental change on the Tibetan Plateau and adjacent areas (Western China and Western Mongolia). Quaternary International 65–66, 121–145. doi:10.1016/S1040-6182(99)00040-3 Lehmkuhl, F., Hilgers, A., Fries, S., Hülle, D., Schlütz, F., Shumilovskikh, L., Felauer, T., Protze, J., 2011. Holocene geomorphological processes and soil development as indicator for environmental change around Karakorum, Upper Orkhon Valley (Central Mongolia). CATENA 87, 31–44. doi:10.1016/j.catena.2011.05.005 Lehmkuhl, F., Klinge, M., Lang, A., 2002. Late Quaternary glacier advances, lake level fluctuations and aeolian sedimentation in Southern Tibet. Zeitschrift für Geomorphologie Neue Folgen, Supplementbände 126, 183–218. Lehmkuhl, F., Klinge, M., Rees-Jones, J., Rhodes, E.., 2000. Late Quaternary aeolian sedimentation in central and south-eastern Tibet. Quaternary International 68–71, 117–132. doi:10.1016/S1040-6182(00)00038-0 Lehmkuhl, F., Owen, L.A., 2005. Late Quaternary glaciation of Tibet and the bordering mountains: a review. Boreas 34, 87–100. doi:10.1111/j.1502-3885.2005.tb01008.x Lehmkuhl, F., Schulte, P., Zhao, H., Hülle, D., Protze, J., Stauch, G., 2014. Timing and spatial distribution of loess and loess-like sediments in the mountain areas of the northeastern Tibetan Plateau. CATENA 117, 23–33. doi:10.1016/j.catena.2013.06.008 Leighton, C.L., Thomas, D.S.G., Bailey, R.M., 2014. Reproducibility and utility of dune luminescence chronologies. Earth-Science Reviews 129, 24–39. doi:10.1016/j.earscirev.2013.11.007 Li, B., Jacobs, Z., Roberts, R.G., Li, S.-H., 2014. Review and assessment of the potential of post-IR IRSL dating methods to circumvent the problem of anomalous fading in feldspar luminescence. Geochronometria 41, 178–201. doi:10.2478/s13386-013-0160-3

167

Li, B., Li, S.-H., 2006. Comparison of estimates using the fast component and the medium component of quartz OSL. Radiation Measurements 41, 125–136. doi:10.1016/j.radmeas.2005.06.037 Li, C., Yin, H., Yu, G., 2000. Evolution of drainage systems and its developing trend in connection with tectonic uplift of Eastern Kunlun Mt. Chinese Science Bulletin 45, 1904–1908. doi:10.1007/BF02886304 Li, H., Dai, A., Zhou, T., Lu, J., 2010. Responses of East Asian summer monsoon to historical SST and atmospheric forcing during 1950–2000. Climate Dynamics 34, 501–514. doi:10.1007/s00382- 008-0482-7 Li, H., Van der Woerd, J., Tapponnier, P., Klinger, Y., Xuexiang, Q., Jingsui, Y., Yintang, Z., 2005. Slip rate on the Kunlun fault at Hongshui Gou, and recurrence time of great events comparable to the 14/11/2001, Mw∼7.9 Kokoxili earthquake. Earth and Planetary Science Letters 237, 285– 299. doi:10.1016/j.epsl.2005.05.041 Li, H., Yang, X., 2016. Spatial and temporal patterns of aeolian activities in the desert belt of northern China revealed by dune chronologies. Quaternary International 410, Part B, 58–68. doi:10.1016/j.quaint.2015.07.015 Li, S., Bates, G.T., 2007. Influence of the Atlantic Multidecadal Oscillation on the winter climate of East China. Advances in Atmospheric Sciences 24, 126–135. doi:10.1007/s00376-007-0126-6 Li, S.-H., Sun, J.-M., Zhao, H., 2002. Optical dating of dune sands in the northeastern deserts of China. Palaeogeography, Palaeoclimatology, Palaeoecology 181, 419–429. doi:10.1016/S0031- 0182(01)00443-6 Li, X., Liang, J., Hou, J., Zhang, W., 2015. Centennial-scale climate variability during the past 2000 years on the central Tibetan Plateau. The Holocene 25, 892–899. doi:10.1177/0959683615572852 Li, X., Yi, C., Chen, F., Yao, T., Li, X., 2006. Formation of proglacial dunes in front of the Puruogangri Icefield in the central Qinghai–Tibet Plateau: Implications for reconstructing paleoenvironmental changes since the Lateglacial. Quaternary International 154–155, 122– 127. doi:10.1016/j.quaint.2006.02.006 Licht, A., Pullen, A., Kapp, P., Abell, J., Giesler, N., 2016. Eolian cannibalism: Reworked loess and fluvial sediment as the main sources of the Chinese Loess Plateau. Geological Society of America Bulletin. doi:10.1130/B31375.1 Lin, A., Fu, B., Guo, J., Zeng, Q., Dang, G., He, W., Zhao, Y., 2002. Co-Seismic Strike-Slip and Rupture Length Produced by the 2001 Ms 8.1 Central Kunlun Earthquake. Science 296, 2015–2017. Lister, G.S., Kelts, K., Zao, C.K., Yu, J.-Q., Niessen, F., 1991. Lake Qinghai, China: closed-basin like levels and the oxygen isotope record for ostracoda since the latest Pleistocene. Palaeogeography, Palaeoclimatology, Palaeoecology 84, 141–162. doi:10.1016/0031- 0182(91)90041-O Liu, B., Jin, H., Sun, L., Sun, Z., Su, Z., 2013a. Winter and summer monsoonal evolution in northeastern Qinghai-Tibetan Plateau during the Holocene period. Chemie der Erde - Geochemistry 73, 309–321. doi:10.1016/j.chemer.2013.03.006 Liu, B., Jin, H., Sun, L., Sun, Z., Su, Z., Zhang, C., 2013b. Holocene climatic change revealed by aeolian deposits from the Gonghe Basin, northeastern Qinghai–Tibetan Plateau. Quaternary International 296, 231–240. doi:10.1016/j.quaint.2012.05.003 Liu, B., Jin, H., Sun, Z., Miao, Y., Su, Z., Zhang, C., 2014. Evidence of Holocene millennial-scale climatic change from Gonghe Basin peat deposit, northeastern Qinghai-Tibet Plateau. Journal of Arid Environments 106, 1–10. doi:10.1016/j.jaridenv.2014.03.003 Liu, G., Zhang, X., Cui, Z., Wu, Y., Ju, Y., 2006. A review of glacial sequences of the Kunlun Pass, northern Tibetan Plateau. Quaternary International 154–155, 63–72. doi:10.1016/j.quaint.2006.02.004 Liu, J., Chen, J., Zhang, X., Li, Y., Rao, Z., Chen, F., 2015. Holocene East Asian summer monsoon records in northern China and their inconsistency with Chinese stalagmite δ18O records. Earth-Science Reviews 148, 194–208. doi:10.1016/j.earscirev.2015.06.004

168

Liu, J., Zhang, P., Cheng, H., Chen, F., Yang, X., Zhang, D., Zhou, J., Jia, J., An, C., Sang, W., Johnson, K.R., 2008. Asian summer monsoon precipitation recorded by stalagmite oxygen isotopic composition in the western Loess Plateau during AD1875–2003 and its linkage with ocean- atmosphere system. Chinese Science Bulletin 53, 2041–2049. doi:10.1007/s11434-008-0286- 5 Liu, K., Lai, Z.P., 2012. Chronology of Holocene sediments from the archaeological Salawusu site in the Mu Us Desert in China and its palaeoenvironmental implications. Journal of Asian Earth Sciences 45, 247–255. doi:10.1016/j.jseaes.2011.11.002 Liu, T., 1988. Loess in China. China Ocean Press/Springer, Beijing/Berlin. Liu, X., Dong, H., Rech, J.A., Matsumoto, R., Bo, Y., Yongbo, W., 2008. Evolution of Chaka Salt Lake in NW China in response to climatic change during the Latest Pleistocene–Holocene. Quaternary Science Reviews 27, 867–879. doi:10.1016/j.quascirev.2007.12.006 Liu, X., Dong, H., Yang, X., Herzschuh, U., Zhang, E., Stuut, J.-B.W., Wang, Y., 2009. Late Holocene forcing of the Asian winter and summer monsoon as evidenced by proxy records from the northern Qinghai–Tibetan Plateau. Earth and Planetary Science Letters 280, 276–284. doi:10.1016/j.epsl.2009.01.041 Liu, X., Lai, Z., Fan, Q., Long, H., Sun, Y., 2010. Timing for high lake levels of Qinghai Lake in the Qinghai-Tibetan Plateau since the Last Interglaciation based on quartz OSL dating. Quaternary Geochronology 5, 218–222. doi:10.1016/j.quageo.2009.03.010 Liu, X., Lai, Z., Madsen, D., Yu, L., Liu, K., Zhang, J., 2011. Lake level variations of Qinghai Lake in northeastern Qinghai-Tibetan Plateau since 3.7 ka based on OSL dating. Quaternary International 236, 57–64. doi:10.1016/j.quaint.2010.08.009 Liu, X., Lai, Z., Yu, L., Sun, Y., Madsen, D., 2012. Luminescence chronology of aeolian deposits from the Qinghai Lake area in the Northeastern Qinghai-Tibetan Plateau and its palaeoenvironmental implications. Quaternary Geochronology 10, 37–43. doi:10.1016/j.quageo.2012.01.016 Liu, X., Vandenberghe, J., An, Z., Li, Y., Jin, Z., Dong, J., Sun, Y., 2016. Grain size of Lake Qinghai sediments: Implications for riverine input and Holocene monsoon variability. Palaeogeography, Palaeoclimatology, Palaeoecology 449, 41–51. doi:10.1016/j.palaeo.2016.02.005 Liu, X., Yin, Z.-Y., 2001. Spatial and Temporal Variation of Summer Precipitation over the Eastern Tibetan Plateau and the North Atlantic Oscillation. J. Climate 14, 2896–2909. doi:10.1175/1520-0442(2001)014<2896:SATVOS>2.0.CO;2 Liu, X., Yu, Z., Dong, H., Chen, H.-F., 2014. A less or more dusty future in the Northern Qinghai- Tibetan Plateau? Scientific Reports 4, 6672. Liu, Z., Wen, X., Brady, E.C., Otto-Bliesner, B., Yu, G., Lu, H., Cheng, H., Wang, Y., Zheng, W., Ding, Y., Edwards, R.L., Cheng, J., Liu, W., Yang, H., 2014. Chinese cave records and the East Asia Summer Monsoon. Quaternary Science Reviews 83, 115–128. doi:10.1016/j.quascirev.2013.10.021 Liu, Z., Zhao, W., 2001. Shifting sand control in central Tibet. Ambio 30, 376–380. Ljungqvist, F.C., 2010. A new reconstruction of temperature variability in the extra-tropical northern hemisphere during the last two millennia. Geografiska Annaler: Series A, Physical Geography 92, 339–351. doi:10.1111/j.1468-0459.2010.00399.x Lockot, G., Ramisch, A., Wünnemann, B., Hartmann, K., Haberzettl, T., Chen, H., Diekmann, B., 2016. A Process- and Provenance-Based Attempt to Unravel Inconsistent Radiocarbon Chronologies in Lake Sediments: An Example from Lake Heihai, North Tibetan Plateau (China). Radiocarbon 57, 1003–1019. doi:10.1017/S0033822200058483 Lomax, J., Hilgers, A., Twidale, C.R., Bourne, J.A., Radtke, U., 2007. Treatment of broad palaeodose distributions in OSL dating of dune sands from the western Murray Basin, South Australia. Quaternary Geochronology 2, 51–56. doi:10.1016/j.quageo.2006.05.015

169

Long, H., Lai, Z., Fuchs, M., Zhang, J., Li, Y., 2012a. Timing of Late Quaternary palaeolake evolution in Tengger Desert of northern China and its possible forcing mechanisms. Global and Planetary Change 92–93, 119–129. doi:10.1016/j.gloplacha.2012.05.014 Long, H., Lai, Z., Fuchs, M., Zhang, J., Yang, L., 2012b. Palaeodunes intercalated in loess strata from the western Chinese Loess Plateau: Timing and palaeoclimatic implications. Quaternary International 263, 37–45. doi:10.1016/j.quaint.2010.12.030 Long, H., Lai, Z., Wang, N., Zhang, J., 2011. A combined luminescence and radiocarbon dating study of Holocene lacustrine sediments from arid northern China. Quaternary Geochronology 6, 1–9. doi:10.1016/j.quageo.2010.06.001 Lu, H., Mason, J.A., Stevens, T., Zhou, Y., Yi, S., Miao, X., 2011a. Response of surface processes to climatic change in the dunefields and Loess Plateau of North China during the late Quaternary. Earth Surf. Process. Landforms 36, 1590–1603. doi:10.1002/esp.2168 Lu, H., Sun, D., 2000. Pathways of dust input to the Chinese Loess Plateau during the last glacial and interglacial periods. CATENA 40, 251–261. doi:10.1016/S0341-8162(00)00090-4 Lu, H., Wang, E., Shi, X., Meng, K., 2012. Cenozoic tectonic evolution of the Elashan range and its surroundings, northern Tibetan Plateau as constrained by paleomagnetism and apatite fission track analyses. Tectonophysics 580, 150–161. doi:10.1016/j.tecto.2012.09.013 Lu, H., Wang, X., Ma, H., Tan, H., Vandenberghe, J., Miao, X., Li, Z., Sun, Y., An, Z., Cao, G., 2004. The Plateau Monsoon variation during the past 130 kyr revealed by loess deposit at northeast Qinghai–Tibet (China). Global and Planetary Change 41, 207–214. doi:10.1016/j.gloplacha.2004.01.006 Lu, H., Yi, S., Liu, Z., Mason, J.A., Jiang, D., Cheng, J., Stevens, T., Xu, Z., Zhang, E., Jin, L., Zhang, Z., Guo, Z., Wang, Y., Otto-Bliesner, B., 2013. Variation of East Asian monsoon precipitation during the past 21 k.y. and potential CO2 forcing. Geology 41, 1023–1026. doi:10.1130/G34488.1 Lu, H., Zhao, C., Mason, J., Yi, S., Zhao, H., Zhou, Y., Ji, J., Swinehart, J., Wang, C., 2011b. Holocene climatic changes revealed by aeolian deposits from the Qinghai Lake area (northeastern Qinghai-Tibetan Plateau) and possible forcing mechanisms. The Holocene 21, 297–304. doi:10.1177/0959683610378884 Lu, R., Jia, F., Gao, S., Shang, Y., Li, J., Zhao, C., 2015. Holocene aeolian activity and climatic change in Qinghai Lake basin, northeastern Qinghai–Tibetan Plateau. Palaeogeography, Palaeoclimatology, Palaeoecology 430, 1–10. doi:10.1016/j.palaeo.2015.03.044 Macklin, M.G., Johnstone, E., Lewin, J., 2005. Pervasive and long-term forcing of Holocene river instability and flooding in Great Britain by centennial-scale climate change. The Holocene 15, 937–943. doi:10.1191/0959683605hl867ft Madsen, A.T., Murray, A.S., 2009. Optically stimulated luminescence dating of young sediments: A review. Geomorphology 109, 3–16. doi:10.1016/j.geomorph.2008.08.020 Madsen, D.B., Haizhou, M., Brantingham, P.J., Xing, G., Rhode, D., Haiying, Z., Olsen, J.W., 2006. The Late Upper Paleolithic occupation of the northern Tibetan Plateau margin. Journal of Archaeological Science 33, 1433–1444. doi:10.1016/j.jas.2006.01.017 Madsen, D.B., Haizhou, M., Rhode, D., Brantingham, P.J., Forman, S.L., 2008. Age constraints on the late Quaternary evolution of Qinghai Lake, Tibetan Plateau. Quaternary Research 69, 316– 325. doi:10.1016/j.yqres.2007.10.013 Maher, B.A., 2008. Holocene variability of the East Asian summer monsoon from Chinese cave records: a re-assessment. The Holocene 18, 861–866. doi:10.1177/0959683608095569 Maher, B.A., Thompson, R., 2012. Oxygen isotopes from Chinese caves: records not of monsoon rainfall but of circulation regime. J. Quaternary Sci. 27, 615–624. doi:10.1002/jqs.2553 Mahowald, N., Kohfeld, K., Hansson, M., Balkanski, Y., Harrison, S.P., Prentice, I.C., Schulz, M., Rodhe, H., 1999. Dust sources and deposition during the last glacial maximum and current climate: A comparison of model results with paleodata from ice cores and marine sediments. Journal of Geophysical Research: Atmospheres 104, 15895–15916. doi:10.1029/1999JD900084

170

Manabe, S., Broccoli, A.J., 1990. Mountains and Arid Climates of Middle Latitudes. Science 247, 192– 195. doi:10.1126/science.247.4939.192 Mann, M.E., Zhang, Z., Rutherford, S., Bradley, R.S., Hughes, M.K., Shindell, D., Ammann, C., Faluvegi, G., Ni, F., 2009. Global Signatures and Dynamical Origins of the Little Ice Age and Medieval Climate Anomaly. Science 326, 1256–1260. doi:10.1126/science.1177303 Mao, R., Ho, C.-H., Shao, Y., Gong, D.-Y., Kim, J., 2011. Influence of Arctic Oscillation on dust activity over northeast Asia. Atmospheric Environment 45, 326–337. doi:10.1016/j.atmosenv.2010.10.020 Mason, J.A., Lu, H., Zhou, Y., Miao, X., Swinehart, J.B., Liu, Z., Goble, R.J., Yi, S., 2009. Dune mobility and aridity at the desert margin of northern China at a time of peak monsoon strength. Geology 37, 947–950. doi:10.1130/G30240A.1 Mason, J.A., Nater, E.A., Zanner, C.W., Bell, J.C., 1999. A new model of topographic effects on the distribution of loess. Geomorphology 28, 223–236. doi:10.1016/S0169-555X(98)00112-3 Maussion, F., Scherer, D., Mölg, T., Collier, E., Curio, J., Finkelnburg, R., 2013. Precipitation Seasonality and Variability over the Tibetan Plateau as Resolved by the High Asia Reanalysis. J. Climate 27, 1910–1927. doi:10.1175/JCLI-D-13-00282.1 Mayewski, P.A., Rohling, E.E., Curt Stager, J., Karlén, W., Maasch, K.A., David Meeker, L., Meyerson, E.A., Gasse, F., van Kreveld, S., Holmgren, K., Lee-Thorp, J., Rosqvist, G., Rack, F., Staubwasser, M., Schneider, R.R., Steig, E.J., 2004. Holocene climate variability. Quaternary Research 62, 243–255. doi:10.1016/j.yqres.2004.07.001 Métivier, F., Gaudemer, Y., Tapponnier, P., Meyer, B., 1998. Northeastward growth of the Tibet plateau deduced from balanced reconstruction of two depositional areas: The Qaidam and Hexi Corridor basins, China. Tectonics 17, 823–842. doi:10.1029/98TC02764 Michczynska, D., Pazdur, A., 2004. Shape analysis of cumulative probability density function of radiocarbon dates set in the study of climate change in the late glacial and Holocene. Radiocarbon; Vol 46, No 2 733–744. Miehe, G., Miehe, S., Böhner, J., Kaiser, K., Hensen, I., Madsen, D., Liu, J., Opgenoorth, L., 2014. How old is the human footprint in the world’s largest alpine ecosystem? A review of multiproxy records from the Tibetan Plateau from the ecologists’ viewpoint. Quaternary Science Reviews 86, 190–209. doi:10.1016/j.quascirev.2013.12.004 Miehe, G., Miehe, S., Kaiser, K., Reudenbach, C., Behrendes, L., La Duo, Schlütz, F., 2009. How old is pastoralism in Tibet? An ecological approach to the making of a Tibetan landscape. Palaeogeography, Palaeoclimatology, Palaeoecology 276, 130–147. doi:10.1016/j.palaeo.2009.03.005 Mischke, S., Aichner, B., Diekmann, B., Herzschuh, U., Plessen, B., Wünnemann, B., Zhang, C., 2010a. Ostracods and stable isotopes of a late glacial and Holocene lake record from the NE Tibetan Plateau. Chemical Geology 276, 95–103. doi:10.1016/j.chemgeo.2010.06.003 Mischke, S., Bößneck, U., Diekmann, B., Herzschuh, U., Jin, H., Kramer, A., Wünnemann, B., Zhang, C., 2010b. Quantitative relationship between water-depth and sub-fossil ostracod assemblages in Lake Donggi Cona, Qinghai Province, China. Journal of Paleolimnology 43, 589–608. doi:10.1007/s10933-009-9355-2 Mischke, S., Kramer, M., Zhang, C., Shang, H., Herzschuh, U., Erzinger, J., 2008. Reduced early Holocene moisture availability in the Bayan Har Mountains, northeastern Tibetan Plateau, inferred from a multi-proxy lake record. Palaeogeography, Palaeoclimatology, Palaeoecology 267, 59–76. doi:10.1016/j.palaeo.2008.06.002 Mischke, S., Weynell, M., Zhang, C., Wiechert, U., 2013. Spatial variability of 14C reservoir effects in Tibetan Plateau lakes. Quaternary International 313–314, 147–155. doi:10.1016/j.quaint.2013.01.030 Mischke, S., Zhang, C., 2010. Holocene cold events on the Tibetan Plateau. Global and Planetary Change 72, 155–163. doi:10.1016/j.gloplacha.2010.02.001

171

Mischke, S., Zhang, C., Börner, A., Herzschuh, U., 2010c. Lateglacial and Holocene variation in aeolian sediment flux over the northeastern Tibetan Plateau recorded by laminated sediments of a saline meromictic lake. J. Quaternary Sci. 25, 162–177. doi:10.1002/jqs.1288 Moberg, A., Sonechkin, D.M., Holmgren, K., Datsenko, N.M., Karlen, W., 2005. Highly variable Northern Hemisphere temperatures reconstructed from low- and high-resolution proxy data. Nature 433, 613–617. doi:10.1038/nature03265 Molg, T., Maussion, F., Scherer, D., 2014. Mid-latitude westerlies as a driver of glacier variability in monsoonal High Asia. Nature Clim. Change 4, 68–73. Molnar, P., Boos, W.R., Battisti, D.S., 2010. Orographic Controls on Climate and Paleoclimate of Asia: Thermal and Mechanical Roles for the Tibetan Plateau. Annu. Rev. Earth Planet. Sci. 38, 77– 102. doi:10.1146/annurev-earth-040809-152456 Morrill, C., Overpeck, J.T., Cole, J.E., Liu, K., Shen, C., Tang, L., 2006. Holocene variations in the Asian monsoon inferred from the geochemistry of lake sediments in central Tibet. Quaternary Research 65, 232–243. doi:10.1016/j.yqres.2005.02.014 Muhs, D.R., 2013. The geologic records of dust in the Quaternary. Aeolian Research 9, 3–48. doi:10.1016/j.aeolia.2012.08.001 Muhs, D.R., Bettis III, E.A., Aleinikoff, J.N., McGeehin, J.P., Beann, J., Skipp, G., Marshall, B.D., Roberts, H.M., Johnson, W.C., Benton, R., 2008. Origin and paleoclimatic significance of late Quaternary loess in Nebraska: Evidence from stratigraphy, chronology, sedimentology, and geochemistry. GSA Bulletin 120, 1378–1407. doi:10.1130/?B26221.1 Müller, C., Kürschner, H., 2013. Phytosociological and palynological studies of alpine steppe communities on the northern Tibetan Plateau, Qinghai Province, China. Feddes Repert. 124, 122–138. doi:10.1002/fedr.201400006 Munyikwa, K., 2005. The role of dune morphogenetic history in the interpretation of linear dune luminescence chronologies: a review of linear dune dynamics. Progress in Physical Geography 29, 317–336. doi:10.1191/0309133305pp451ra Murray, A.S., Olley, J.M., 2002. Precision and accuracy in the optically stimulated luminescence dating of sedimentary quartz: a status review. Geochronometria 21, 1–16. Murray, A.S., Wintle, A.G., 2003. The single aliquot regenerative dose protocol: potential for improvements in reliability. Radiation Measurements 37, 377–381. doi:10.1016/S1350- 4487(03)00053-2 Murray, A.S., Wintle, A.G., 2000. Luminescence dating of quartz using an improved single-aliquot regenerative-dose protocol. Radiation Measurements 32, 57–73. doi:10.1016/S1350- 4487(99)00253-X Nai’ang, W., Qiang, Z., Jijun, L., Gang, H., Hongyi, C., 2003. The sand wedges of the last ice age in the Hexi Corridor, China: paleoclimatic interpretation. Geomorphology 51, 313–320. doi:10.1016/S0169-555X(02)00243-X Nie, J., Stevens, T., Rittner, M., Stockli, D., Garzanti, E., Limonta, M., Bird, A., Andò, S., Vermeesch, P., Saylor, J., Lu, H., Breecker, D., Hu, X., Liu, S., Resentini, A., Vezzoli, G., Peng, W., Carter, A., Ji, S., Pan, B., 2015. Loess Plateau storage of Northeastern Tibetan Plateau-derived Yellow River sediment. Nature Communications 6, 8511. Nilson, E., Lehmkuhl, F., 2001. Interpreting temporal patterns in the late Quaternary dust flux from Asia to the North Pacific. Quaternary International 76–77, 67–76. doi:10.1016/S1040- 6182(00)00090-2 Noren, A.J., Bierman, P.R., Steig, E.J., Lini, A., Southon, J., 2002. Millennial-scale storminess variability in the northeastern United States during the Holocene epoch. Nature 419, 821–824. doi:10.1038/nature01132 Nottebaum, V., Lehmkuhl, F., Stauch, G., Hartmann, K., Wünnemann, B., Schimpf, S., Lu, H., 2014. Regional grain size variations in aeolian sediments along the transition between Tibetan highlands and north-western Chinese deserts – the influence of geomorphological settings on aeolian transport pathways. Earth Surf. Process. Landforms 39, 1960–1978. doi:10.1002/esp.3590

172

Nottebaum, V., Lehmkuhl, F., Stauch, G., Lu, H., Yi, S., 2015a. Late Quaternary aeolian sand deposition sustained by fluvial reworking and sediment supply in the Hexi Corridor — An example from northern Chinese drylands. Geomorphology 250, 113–127. doi:10.1016/j.geomorph.2015.08.014 Nottebaum, V., Stauch, G., Hartmann, K., Zhang, J., Lehmkuhl, F., 2015b. Unmixed loess grain size populations along the northern Qilian Shan (China): Relationships between geomorphologic, sedimentologic and climatic controls. Quaternary International 372, 151–166. doi:10.1016/j.quaint.2014.12.071 Nugteren, G., Vandenberghe, J., van Huissteden, J.K., Zhisheng, A., 2004. A Quaternary climate record based on grain size analysis from the Luochuan loess section on the Central Loess Plateau, China. Global and Planetary Change 41, 167–183. doi:10.1016/j.gloplacha.2004.01.004 Olsen, J., Anderson, N.J., Knudsen, M.F., 2012. Variability of the North Atlantic Oscillation over the past 5,200 years. Nature Geosci 5, 808–812. doi:10.1038/ngeo1589 Opitz, S., Ramisch, A., IJmker, J., Lehmkuhl, F., Mischke, S., Stauch, G., Wünnemann, B., Zhang, Y., Diekmann, B., 2016. Spatio-temporal pattern of detrital clay-mineral supply to a lake system on the north-eastern Tibetan Plateau, and its relationship to late Quaternary paleoenvironmental changes. CATENA 137, 203–218. doi:10.1016/j.catena.2015.09.003 Opitz, S., Wünnemann, B., Aichner, B., Dietze, E., Hartmann, K., Herzschuh, U., IJmker, J., Lehmkuhl, F., Li, S., Mischke, S., Plotzki, A., Stauch, G., Diekmann, B., 2012. Late Glacial and Holocene development of Lake Donggi Cona, north-eastern Tibetan Plateau, inferred from sedimentological analysis. Palaeogeography, Palaeoclimatology, Palaeoecology 337–338, 159–176. doi:10.1016/j.palaeo.2012.04.013 Owen, L.A., 2009. Latest Pleistocene and Holocene glacier fluctuations in the Himalaya and Tibet. Quaternary Science Reviews 28, 2150–2164. doi:10.1016/j.quascirev.2008.10.020 Owen, L.A., Caffee, M.W., Finkel, R.C., Seong, Y.B., 2008. Quaternary glaciation of the Himalayan– Tibetan orogen. J. Quaternary Sci. 23, 513–531. doi:10.1002/jqs.1203 Owen, L.A., Dortch, J.M., 2014. Nature and timing of Quaternary glaciation in the Himalayan–Tibetan orogen. Quaternary Science Reviews 88, 14–54. doi:10.1016/j.quascirev.2013.11.016 Owen, L.A., Finkel, R.C., Haizhou, M., Barnard, P.L., 2006. Late Quaternary landscape evolution in the Kunlun Mountains and Qaidam Basin, Northern Tibet: A framework for examining the links between glaciation, lake level changes and alluvial fan formation. Quaternary International 154–155, 73–86. doi:10.1016/j.quaint.2006.02.008 Owen, L.A., Finkel, R.C., Haizhou, M., Spencer, J.Q., Derbyshire, E., Barnard, P.L., Caffee, M.W., 2003. Timing and style of Late Quaternary glaciation in northeastern Tibet. Geological Society of America Bulletin 115, 1356–1364. doi:10.1130/B25314.1 Owen, L.A., Thackray, G., Anderson, R.S., Briner, J., Kaufman, D., Roe, G., Pfeffer, W., Yi, C., 2009. Integrated research on mountain glaciers: Current status, priorities and future prospects. Geomorphology 103, 158–171. doi:10.1016/j.geomorph.2008.04.019 Özer, M., Orhan, M., Isik, N.S., 2010. Effect of Particle Optical Properties on Size Distribution of Soils Obtained by Laser Diffraction. Environmental & Engineering Geoscience 16, 163–173. doi:10.2113/gseegeosci.16.2.163 PAGES 2k Consortium, 2013. Continental-scale temperature variability during the past two millennia. Nature Geosci 6, 339–346. Pan, B., Hu, X., Gao, H., Hu, Z., Cao, B., Geng, H., Li, Q., 2013. Late Quaternary river incision rates and rock uplift pattern of the eastern Qilian Shan Mountain, China. Geomorphology 184, 84–97. doi:10.1016/j.geomorph.2012.11.020 Pan, M., Wu, Y., Zheng, Y., Tan, L., 2014. Holocene aeolian activity in the Dinggye area (Southern Tibet, China). Aeolian Research 12, 19–27. doi:10.1016/j.aeolia.2013.10.005 Perrineau, A., Woerd, J.V.D., Gaudemer, Y., Liu-Zeng, J., Pik, R., Tapponnier, P., Thuizat, R., Rongzhang, Z., 2011. Incision rate of the Yellow River in Northeastern Tibet constrained by 10Be and 26Al cosmogenic isotope dating of fluvial terraces: implications for catchment

173

evolution and plateau building. Geological Society, London, Special Publications 353, 189– 219. doi:10.1144/SP353.10 Pewe, T.L., Liu, T., Slatt, R.M., Li, B., 1995. Origin and character of loesslike silt in the southern Qinghai-Xizang (Tibet) plateau, China. U.S. Geological Survey professional paper 1549, 1–55. Pietsch, T.J., Olley, J.M., Nanson, G.C., 2008. Fluvial transport as a natural luminescence sensitiser of quartz. Quaternary Geochronology 3, 365–376. doi:10.1016/j.quageo.2007.12.005 Porter, S.C., 2001. Chinese loess record of monsoon climate during the last glacial–interglacial cycle. Earth-Science Reviews 54, 115–128. doi:10.1016/S0012-8252(01)00043-5 Porter, S.C., An, Z., 1995. Correlation between climate events in the North Atlantic and China during the last glaciation. Nature 375, 305–308. doi:10.1038/375305a0 Porter, S.C., Singhvi, A., Zhisheng, A., Zhongping, L., 2001. Luminescence age and palaeoenvironmental implications of a late Pleistocene ground wedge on the Northeastern Tibetan Plateau. Permafrost Periglac. Process. 12, 203–210. doi:10.1002/ppp.386 Prell, W.L., Kutzbach, J.E., 1992. Sensitivity of the Indian monsoon to forcing parameters and implications for its evolution. Nature 360, 647–652. doi:10.1038/360647a0 Prescott, J.R., Hutton, J.T., 1994. Cosmic ray contributions to dose rates for luminescence and ESR dating: Large depths and long-term time variations. Radiation Measurements 23, 497–500. doi:10.1016/1350-4487(94)90086-8 Prescott, J.R., Hutton, J.T., 1988. Cosmic ray and gamma ray dosimetry for TL and ESR. International Journal of Radiation Applications and Instrumentation. Part D. Nuclear Tracks and Radiation Measurements 14, 223–227. doi:10.1016/1359-0189(88)90069-6 Preusser, F., 2003. IRSL Dating of K-rich feldspars using SAR protocol: Comparison with independent age control. Ancient TL 21, 17–23. Preusser, F., Andersen, B.G., Denton, G.H., Schlüchter, C., 2005. Luminescence chronology of Late Pleistocene glacial deposits in North Westland, New Zealand. Quaternary Science Reviews 24, 2207–2227. doi:10.1016/j.quascirev.2004.12.005 Preusser, F., Chithambo, M.L., Götte, T., Martini, M., Ramseyer, K., Sendezera, E.J., Susino, G.J., Wintle, A.G., 2009. Quartz as a natural luminescence dosimeter. Earth-Science Reviews 97, 184–214. doi:10.1016/j.earscirev.2009.09.006 Preusser, F., Degering, D., Fuchs, M., Hilgers, A., Kadereit, A., Klasen, N., Krebetschek, M., Richter, D., Spencer, J.Q.G., 2008. Luminescence dating: basics, methods and applications. Quaternary Science Journal 57, 95–149. doi:10.3285/eg.57.1-2.5 Prins, M.A., Zheng, H., Beets, K., Troelstra, S., Bacon, P., Kamerling, I., Wester, W., Konert, M., Huang, X., Ke, W., Vandenberghe, J., 2009. Dust supply from river floodplains: the case of the lower Huang He (Yellow River) recorded in a loess–palaeosol sequence from the Mangshan Plateau. J. Quaternary Sci. 24, 75–84. doi:10.1002/jqs.1167 Prospero, J.M., Ginoux, P., Torres, O., Nicholson, S.E., Gill, T.E., 2002. ENVIRONMENTAL CHARACTERIZATION OF GLOBAL SOURCES OF ATMOSPHERIC SOIL DUST IDENTIFIED WITH THE NIMBUS 7 TOTAL OZONE MAPPING SPECTROMETER (TOMS) ABSORBING AEROSOL PRODUCT. Rev. Geophys. 40, 2–1. doi:10.1029/2000RG000095 Pye, K., 1995. The nature, origin and accumulation of loess. Quaternary Science Reviews 14, 653–667. doi:10.1016/0277-3791(95)00047-X Pye, K., 1987. Aeolian Dust and Dust Deposits. Academic Press, London. Qiang, M., Chen, F., Song, L., Liu, X., Li, M., Wang, Q., 2013a. Late Quaternary aeolian activity in Gonghe Basin, northeastern Qinghai-Tibetan Plateau, China. Quaternary Research 79, 403– 412. doi:10.1016/j.yqres.2013.03.003 Qiang, M., Chen, F., Wang, Z., Niu, G., Song, L., 2010. Aeolian deposits at the southeastern margin of the Tengger Desert (China): Implications for surface wind strength in the Asian dust source area over the past 20,000 years. Palaeogeography, Palaeoclimatology, Palaeoecology 286, 66–80. doi:10.1016/j.palaeo.2009.12.005 Qiang, M., Chen, F., Zhou, A., Xiao, S., Zhang, J., Wang, Z., 2007. Impacts of wind velocity on sand and dust deposition during dust storm as inferred from a series of observations in the

174

northeastern Qinghai–Tibetan Plateau, China. Powder Technology 175, 82–89. doi:10.1016/j.powtec.2006.12.020 Qiang, M., Jin, Y., Liu, X., Song, L., Li, H., Li, F., Chen, F., 2016. Late Pleistocene and Holocene aeolian sedimentation in Gonghe Basin, northeastern Qinghai-Tibetan Plateau: Variability, processes, and climatic implications. Quaternary Science Reviews 132, 57–73. doi:10.1016/j.quascirev.2015.11.010 Qiang, M., Liu, Y., Jin, Y., Song, L., Huang, X., Chen, F., 2014. Holocene record of eolian activity from Genggahai Lake, northeastern Qinghai-Tibetan Plateau, China. Geophys. Res. Lett. 41, 589– 595. doi:10.1002/2013GL058806 Qiang, M., Song, L., Chen, F., Li, M., Liu, X., Wang, Q., 2013b. A 16-ka lake-level record inferred from macrofossils in a sediment core from Genggahai Lake, northeastern Qinghai–Tibetan Plateau (China). Journal of Paleolimnology 49, 575–590. doi:10.1007/s10933-012-9660-z Qiu, J., 2008. China: the third pole. Nature 454, 393–396. doi:10.1038/454393a Ramisch, A., Lockot, G., Haberzettl, T., Hartmann, K., Kuhn, G., Lehmkuhl, F., Schimpf, S., Schulte, P., Stauch, G., Wang, R., Wünnemann, B., Yan, D., Zhang, Y., Diekmann, B., 2016. A persistent northern boundary of Indian Summer Monsoon precipitation over Central Asia during the Holocene. Scientific Reports 6, 25791. Ran, M., Feng, Z., 2013. Holocene moisture variations across China and driving mechanisms: A synthesis of climatic records. Quaternary International 313–314, 179–193. doi:10.1016/j.quaint.2013.09.034 Rao, Z., Jia, G., Qiang, M., Zhao, Y., 2014. Assessment of the difference between mid- and long chain compound specific δDn-alkanes values in lacustrine sediments as a paleoclimatic indicator. Organic Geochemistry 76, 104–117. doi:10.1016/j.orggeochem.2014.07.015 Rao, Z., Li, Y., Zhang, J., Jia, G., Chen, F., 2016a. Investigating the long-term palaeoclimatic controls on the δD and δ18O of precipitation during the Holocene in the Indian and East Asian monsoonal regions. Earth-Science Reviews 159, 292–305. doi:10.1016/j.earscirev.2016.06.007 Rao, Z., Qiang, M., Jia, G., Li, Y., Dan, D., Chen, F., 2016b. A 15 ka lake water δD record from Genggahai Lake, northeastern Tibetan Plateau, and its paleoclimatic significance. Organic Geochemistry. doi:10.1016/j.orggeochem.2016.03.007 Reimann, T., Tsukamoto, S., Naumann, M., Frechen, M., 2011. The potential of using K-rich feldspars for optical dating of young coastal sediments – A test case from Darss-Zingst peninsula (southern Baltic Sea coast). Quaternary Geochronology 6, 207–222. doi:10.1016/j.quageo.2010.10.001 Reimer, P.J., Baillie, M.G.L., Bard, E., Bayliss, A., Beck, J.W., Blackwell, P.G., Ramsey, C.B., Buck, C.E., Burr, G.S., Edwards, R.L., Friedrich, M., Grootes, P.M., Guilderson, T.P., Hajdas, I., Heaton, T.J., Hogg, A.G., Hughen, K.A., Kaiser, K.F., Kromer, B., McCormac, F.G., Manning, S.W., Reimer, R.W., Richards, D.A., Southon, J.R., Talamo, S., Turney, C.S.M., van der Plicht, J., Weyhenmeyer, C.E., 2011. IntCal09 and Marine09 Radiocarbon Age Calibration Curves, 0- 50,000 Years cal BP. Radiocarbon; Vol 51, No 4 (2009). Rhode, D., Haizhou, M., Madsen, D.B., Brantingham, P.J., Forman, S.L., Olsen, J.W., 2010. Paleoenvironmental and archaeological investigations at Qinghai Lake, western China: Geomorphic and chronometric evidence of lake level history. Quaternary International 218, 29–44. doi:10.1016/j.quaint.2009.03.004 Rice, M.A., Willets, B.B., McEwan, I.K., 1995. An experimental study of multiple grain-size ejecta produced by collisions of saltating grains with a flat bed. Sedimentology 42, 695–706. doi:10.1111/j.1365-3091.1995.tb00401.x Roberts, H.M., 2008. The development and application of luminescence dating to loess deposits: a perspective on the past, present and future. Boreas 37, 483–507. doi:10.1111/j.1502- 3885.2008.00057.x

175

Roberts, H.M., Duller, G.A.T., 2004. Standardised growth curves for optical dating of sediment using multiple-grain aliquots. Radiation Measurements 38, 241–252. doi:10.1016/j.radmeas.2003.10.001 Rodríguez-López, J.P., Clemmensen, L.B., Lancaster, N., Mountney, N.P., Veiga, G.D., 2014. Archean to Recent aeolian sand systems and their sedimentary record: Current understanding and future prospects. Sedimentology 61, 1487–1534. doi:10.1111/sed.12123 Roe, G., 2009. On the interpretation of Chinese loess as a paleoclimate indicator. Quaternary Research 71, 150–161. doi:10.1016/j.yqres.2008.09.004 Roskin, J., Tsoar, H., Porat, N., Blumberg, D.G., 2011. Palaeoclimate interpretations of Late Pleistocene vegetated linear dune mobilization episodes: evidence from the northwestern Negev dunefield, Israel. Quaternary Science Reviews 30, 3364–3380. doi:10.1016/j.quascirev.2011.08.014 Royden, L.H., Burchfiel, B.C., van der Hilst, R.D., 2008. The Geological Evolution of the Tibetan Plateau. Science 321, 1054–1058. doi:10.1126/science.1155371 Ruddiman, W.F., Kutzbach, J.E., 1989. Forcing of late Cenozoic northern hemisphere climate by plateau uplift in southern Asia and the American west. J. Geophys. Res. 94, 18409–18427. doi:10.1029/JD094iD15p18409 Rutter, N., Ding, Z., 1993. Paleoclimates and monsoon variations interpreted from micromorphogenic features of the Baoji paleosols, China. Quaternary Science Reviews 12, 853–862. doi:10.1016/0277-3791(93)90024-G Scheidt, S., Ramsey, M., Lancaster, N., 2010. Determining soil moisture and sediment availability at White Sands Dune Field, New Mexico, from apparent thermal inertia data. J. Geophys. Res. 115, n/a-n/a. doi:10.1029/2009JF001378 Schlütz, F., Lehmkuhl, F., 2009. Holocene climatic change and the nomadic Anthropocene in Eastern Tibet: palynological and geomorphological results from the Nianbaoyeze Mountains. Quaternary Science Reviews 28, 1449–1471. doi:10.1016/j.quascirev.2009.01.009 Schulte, P., Lehmkuhl, F., Steininger, F., Loibl, D., Lockot, G., Protze, J., Fischer, P., Stauch, G., 2016. Influence of HCl pretreatment and organo-mineral complexes on laser diffraction measurement of loess–paleosol-sequences. CATENA 137, 392–405. doi:10.1016/j.catena.2015.10.015 Seong, Y.B., Owen, L.A., Yi, C., Finkel, R.C., 2009. Quaternary glaciation of Muztag Ata and Kongur Shan: Evidence for glacier response to rapid climate changes throughout the Late Glacial and Holocene in westernmost Tibet. Geological Society of America Bulletin 121, 348–365. doi:10.1130/B26339.1 Shao, X., Xu, Y., Yin, Z.-Y., Liang, E., Zhu, H., Wang, S., 2010. Climatic implications of a 3585-year tree- ring width chronology from the northeastern Qinghai-Tibetan Plateau. Quaternary Science Reviews 29, 2111–2122. doi:10.1016/j.quascirev.2010.05.005 Shao, Y., 2008. Physics and Modelling of Wind Erosion, 2. ed. Springer, Berlin. Shao, Y., Lu, H., 2000. A simple expression for wind erosion threshold friction velocity. J. Geophys. Res. 105, 22437–22443. doi:10.1029/2000JD900304 Shen, J., Liu, X., Wang, S., Matsumoto, R., 2005. Palaeoclimatic changes in the Qinghai Lake area during the last 18,000 years. Quaternary International 136, 131–140. doi:10.1016/j.quaint.2004.11.014 Shi, Z.G., Liu, X.D., Sun, Y.B., An, Z.S., Liu, Z., Kutzbach, J., 2011. Distinct responses of East Asian summer and winter monsoons to astronomical forcing. Climate of the Past 7, 1363–1370. doi:10.5194/cp-7-1363-2011 Shinoda, M., Gillies, J.A., Mikami, M., Shao, Y., 2011. Temperate grasslands as a dust source: Knowledge, uncertainties, and challenges. Aeolian Research 3, 271–293. doi:10.1016/j.aeolia.2011.07.001 Singarayer, J.S., Bailey, R.M., 2003. Further investigations of the quartz optically stimulated luminescence components using linear modulation. Radiation Measurements 37, 451–458. doi:10.1016/S1350-4487(03)00062-3

176

Singhvi, A.K., Bluszcz, A., Bateman, M.D., Rao, M.S., 2001. Luminescence dating of loess–palaeosol sequences and coversands: methodological aspects and palaeoclimatic implications. Earth- Science Reviews 54, 193–211. doi:10.1016/S0012-8252(01)00048-4 Singhvi, A.K., Porat, N., 2008. Impact of luminescence dating on geomorphological and palaeoclimate research in drylands. Boreas 37, 536–558. doi:10.1111/j.1502-3885.2008.00058.x Sinha, A., Cannariato, K.G., Stott, L.D., Li, H.-C., You, C.-F., Cheng, H., Edwards, R.L., Singh, I.B., 2005. Variability of Southwest Indian summer monsoon precipitation during the Bølling-Ållerød. Geology 33, 813–816. doi:10.1130/G21498.1 Smalley, I., 1995. Aeolian Sediments in the Quaternary RecordMaking the material: The formation of silt sized primary mineral particles for loess deposits. Quaternary Science Reviews 14, 645– 651. doi:10.1016/0277-3791(95)00046-1 Smalley, I., O’Hara-Dhand, K., Kwong, J., 2014. China: Materials for a loess landscape. CATENA 117, 100–107. doi:10.1016/j.catena.2013.11.016 Solomina, O.N., Bradley, R.S., Hodgson, D.A., Ivy-Ochs, S., Jomelli, V., Mackintosh, A.N., Nesje, A., Owen, L.A., Wanner, H., Wiles, G.C., Young, N.E., 2015. Holocene glacier fluctuations. Quaternary Science Reviews 111, 9–34. doi:10.1016/j.quascirev.2014.11.018 Song, L., Qiang, M., Lang, L., Liu, X., Wang, Q., Li, M., 2012. Changes in palaeoproductivity of Genggahai Lake over the past 16 ka in the Gonghe Basin, northeastern Qinghai-Tibetan Plateau. Chinese Science Bulletin 57, 2595–2605. doi:10.1007/s11434-012-5191-2 Spicer, R.A., Harris, N.B.W., Widdowson, M., Herman, A.B., Guo, S., Valdes, P.J., Wolfe, J.A., Kelley, S.P., 2003. Constant elevation of southern Tibet over the past 15 million years. Nature 421, 622–624. doi:10.1038/nature01356 Stauch, G., 2016. Multi-decadal periods of enhanced aeolian activity on the north-eastern Tibet Plateau during the last 2ka. Quaternary Science Reviews 149, 91–101. doi:10.1016/j.quascirev.2016.07.027 Stauch, G., 2015. Geomorphological and palaeoclimate dynamics recorded by the formation of aeolian archives on the Tibetan Plateau. Earth-Science Reviews 150, 393–408. doi:10.1016/j.earscirev.2015.08.009 Stauch, G., IJmker, J., Pötsch, S., Zhao, H., Hilgers, A., Diekmann, B., Dietze, E., Hartmann, K., Opitz, S., Wünnemann, B., Lehmkuhl, F., 2012. Aeolian sediments on the north-eastern Tibetan Plateau. Quaternary Science Reviews 57, 71–84. doi:10.1016/j.quascirev.2012.10.001 Stauch, G., Pötsch, S., Zhao, H., Lehmkuhl, F., 2014. Interaction of geomorphological processes on the north-eastern Tibetan Plateau during the Holocene, an example from a sub-catchment of Lake Donggi Cona. Geomorphology 210, 23–35. doi:10.1016/j.geomorph.2013.12.014 Steinhilber, F., Abreu, J.A., Beer, J., Brunner, I., Christl, M., Fischer, H., Heikkilä, U., Kubik, P.W., Mann, M., McCracken, K.G., Miller, H., Miyahara, H., Oerter, H., Wilhelms, F., 2012. 9,400 years of cosmic radiation and solar activity from ice cores and tree rings. Proceedings of the National Academy of Sciences 109, 5967–5971. doi:10.1073/pnas.1118965109 Stevens, T., Armitage, S.J., Lu, H., Thomas, D.S.G., 2006. Sedimentation and diagenesis of Chinese loess: Implications for the preservation of continuous, high-resolution climate records. Geology 34, 849–852. doi:10.1130/G22472.1 Stevens, T., Carter, A., Watson, T.P., Vermeesch, P., Andò, S., Bird, A.F., Lu, H., Garzanti, E., Cottam, M.A., Sevastjanova, I., 2013. Genetic linkage between the Yellow River, the Mu Us desert and the Chinese Loess Plateau. Quaternary Science Reviews 78, 355–368. doi:10.1016/j.quascirev.2012.11.032 Stevens, T., Lu, H., 2009. Optically stimulated luminescence dating as a tool for calculating sedimentation rates in Chinese loess: comparisons with grain-size records. Sedimentology 56, 911–934. doi:10.1111/j.1365-3091.2008.01004.x Stevens, T., Thomas, D.S.G., Armitage, S.J., Lunn, H.R., Lu, H., 2007. Reinterpreting climate proxy records from late Quaternary Chinese loess: A detailed OSL investigation. Earth-Science Reviews 80, 111–136. doi:10.1016/j.earscirev.2006.09.001

177

Stöckli, R., Vermonte, E., Saleous, N., Simmon, R., Herring, D., 2005. The Blue Marble Next Generation - A true color earth dataset including seasonal dynamics from MODIS. Stokes, S., Hetzel, R., Bailey, R.M., Mingxin, T., 2003. Combined IRSL-OSL single aliquot regeneration (SAR) equivalent dose (De) estimates from source proximal Chinese loess. Quaternary Science Reviews 22, 975–983. doi:10.1016/S0277-3791(03)00044-1 Stone, A.E.C., Thomas, D.S.G., 2008. Linear dune accumulation chronologies from the southwest Kalahari, Namibia: challenges of reconstructing late Quaternary palaeoenvironments from aeolian landforms. Quaternary Science Reviews 27, 1667–1681. doi:10.1016/j.quascirev.2008.06.008 Sun, D., Bloemendal, J., Rea, D.., Vandenberghe, J., Jiang, F., An, Z., Su, R., 2002. Grain-size distribution function of polymodal sediments in hydraulic and aeolian environments, and numerical partitioning of the sedimentary components. Sedimentary Geology 152, 263–277. doi:10.1016/S0037-0738(02)00082-9 Sun, D., Bloemendal, J., Rea, D.K., An, Z., Vandenberghe, J., Lu, H., Su, R., Liu, T., 2004. Bimodal grain- size distribution of Chinese loess, and its palaeoclimatic implications. CATENA 55, 325–340. doi:10.1016/S0341-8162(03)00109-7 Sun, D., Shaw, J., An, Z., Cheng, M., Yue, L., 1998. Magnetostratigraphy and paleoclimatic interpretation of a continuous 7.2 Ma Late Cenozoic Eolian sediments from the Chinese Loess Plateau. Geophys. Res. Lett. 25, 85–88. doi:10.1029/97GL03353 Sun, D., Su, R., Bloemendal, J., Lu, H., 2008. Grain-size and accumulation rate records from Late Cenozoic aeolian sequences in northern China: Implications for variations in the East Asian winter monsoon and westerly atmospheric circulation. Palaeogeography, Palaeoclimatology, Palaeoecology 264, 39–53. doi:10.1016/j.palaeo.2008.03.011 Sun, D., Zhang, Y., Han, F., Zhang, Y., Yi, Z., Li, Z., Wang, F., Wu, S., Li, B., 2011. Magnetostratigraphy and palaeoenvironmental records for a Late Cenozoic sedimentary sequence from Lanzhou, Northeastern margin of the Tibetan Plateau. Global and Planetary Change 76, 106–116. doi:10.1016/j.gloplacha.2010.12.006 Sun, J., 2002. Provenance of loess material and formation of loess deposits on the Chinese Loess Plateau. Earth and Planetary Science Letters 203, 845–859. doi:10.1016/S0012- 821X(02)00921-4 Sun, J., Li, S.-H., Han, P., Chen, Y., 2006. Holocene environmental changes in the central Inner Mongolia, based on single-aliquot-quartz optical dating and multi-proxy study of dune sands. Palaeogeography, Palaeoclimatology, Palaeoecology 233, 51–62. doi:10.1016/j.palaeo.2005.09.016 Sun, J., Li, S.-H., Muhs, D.R., Li, B., 2007. Loess sedimentation in Tibet: provenance, processes, and link with Quaternary glaciations. Quaternary Science Reviews 26, 2265–2280. doi:10.1016/j.quascirev.2007.05.003 Sun, J., Ye, J., Wu, W., Ni, X., Bi, S., Zhang, Z., Liu, W., Meng, J., 2010. Late Oligocene–Miocene mid- latitude aridification and wind patterns in the Asian interior. Geology 38, 515–518. doi:10.1130/G30776.1 Sun, J., Zhang, M., Liu, T., 2001. Spatial and temporal characteristics of dust storms in China and its surrounding regions, 1960–1999: Relations to source area and climate. J. Geophys. Res. 106, 10325–10333. doi:10.1029/2000JD900665 Sun, Y., An, Z., Clemens, S.C., Bloemendal, J., Vandenberghe, J., 2010. Seven million years of wind and precipitation variability on the Chinese Loess Plateau. Earth and Planetary Science Letters 297, 525–535. doi:10.1016/j.epsl.2010.07.004 Sun, Y., Kutzbach, J., An, Z., Clemens, S., Liu, Z., Liu, W., Liu, X., Shi, Z., Zheng, W., Liang, L., Yan, Y., Li, Y., 2015. Astronomical and glacial forcing of East Asian summer monsoon variability. Quaternary Science Reviews 115, 132–142. doi:10.1016/j.quascirev.2015.03.009 Surovell, T.A., Brantingham, P.J., 2007. A note on the use of temporal frequency distributions in studies of prehistoric demography. Journal of Archaeological Science 34, 1868–1877. doi:10.1016/j.jas.2007.01.003

178

Surovell, T.A., Byrd Finley, J., Smith, G.M., Brantingham, P.J., Kelly, R., 2009. Correcting temporal frequency distributions for taphonomic bias. Journal of Archaeological Science 36, 1715– 1724. doi:10.1016/j.jas.2009.03.029 Sweeney, M.R., Mason, J.A., 2013. Mechanisms of dust emission from Pleistocene loess deposits, Nebraska, USA. J. Geophys. Res. Earth Surf. 118, 1460–1471. doi:10.1002/jgrf.20101 Ta, W., Xiao, H., Qu, J., Xiao, Z., Yang, G., Wang, T., Zhang, X., 2004. Measurements of dust deposition in Gansu Province, China, 1986–2000. Geomorphology 57, 41–51. doi:10.1016/S0169- 555X(03)00082-5 Tan, L., Cai, Y., Yi, L., An, Z., Ai, L., 2008. Precipitation variations of Longxi, northeast margin of Tibetan Plateau since AD 960 and their relationship with solar activity. Clim. Past 4, 19–28. doi:10.5194/cp-4-19-2008 Tapponnier, P., Molnar, P., 1977. Active faulting and tectonics in China. Journal of Geophysical Research 82, 2905–2930. doi:10.1029/JB082i020p02905 Tapponnier, P., Zhiqin, X., Roger, F., Meyer, B., Arnaud, N., Wittlinger, G., Jingsui, Y., 2001. Oblique Stepwise Rise and Growth of the Tibet Plateau. Science 294, 1671–1677. doi:10.1126/science.105978 Telfer, M.W., Bailey, R.M., Burrough, S.L., Stone, A.E.S., Thomas, D.S.G., Wiggs, G.S.F., 2010. Understanding linear dune chronologies: Insights from a simple accumulation model. Geomorphology 120, 195–208. doi:10.1016/j.geomorph.2010.03.030 Telfer, M.W., Thomas, D.S.G., 2007. Late Quaternary linear dune accumulation and chronostratigraphy of the southwestern Kalahari: implications for aeolian palaeoclimatic reconstructions and predictions of future dynamics. Quaternary Science Reviews 26, 2617– 2630. doi:10.1016/j.quascirev.2007.07.006 Thomas, D.S.G., 2013. Reconstructing paleoenvironments and palaeoclimates in drylands: what can landform analysis contribute? Earth Surf. Process. Landforms 38, 3–16. doi:10.1002/esp.3190 Thomas, D.S.G., Burrough, S.L., 2016. Luminescence-based dune chronologies in southern Africa: Analysis and interpretation of dune database records across the subcontinent. Quaternary International 410, Part B, 30–45. doi:10.1016/j.quaint.2013.09.008 Thomas, D.S.G., Wiggs, G.F.S., 2008. Aeolian system responses to global change: challenges of scale, process and temporal integration. Earth Surf. Process. Landforms 33, 1396–1418. doi:10.1002/esp.1719 Thomas, E.K., Huang, Y., Clemens, S.C., Colman, S.M., Morrill, C., Wegener, P., Zhao, J., 2016. Changes in dominant moisture sources and the consequences for hydroclimate on the northeastern Tibetan Plateau during the past 32 kyr. Quaternary Science Reviews 131, 157–167. doi:10.1016/j.quascirev.2015.11.003 Thompson, L.G., Yao, T., Davis, M.E., Henderson, K.A., Mosley-Thompson, E., Lin, P.-N., Beer, J., Synal, H.-A., Cole-Dai, J., Bolzan, J.F., 1997. Tropical Climate Instability: The Last Glacial Cycle from a Qinghai-Tibetan Ice Core. Science 276, 1821. doi:10.1126/science.276.5320.1821 Trenberth, K.E., Stepaniak, D.P., Caron, J.M., 2000. The Global Monsoon as Seen through the Divergent Atmospheric Circulation. J. Climate 13, 3969–3993. doi:10.1175/1520- 0442(2000)013<3969:TGMAST>2.0.CO;2 Trouet, V., Esper, J., Graham, N.E., Baker, A., Scourse, J.D., Frank, D.C., 2009. Persistent Positive North Atlantic Oscillation Mode Dominated the Medieval Climate Anomaly. Science 324, 78– 80. doi:10.1126/science.1166349 Trouet, V., Scourse, J.D., Raible, C.C., 2012. North Atlantic storminess and Atlantic Meridional Overturning Circulation during the last Millennium: Reconciling contradictory proxy records of NAO variability. Global and Planetary Change 84–85, 48–55. doi:10.1016/j.gloplacha.2011.10.003 Tsoar, H., Pye, K., 1987. Dust transport and the question of desert loess formation. Sedimentology 34, 139–153. doi:10.1111/j.1365-3091.1987.tb00566.x

179

Újvári, G., Kok, J.F., Varga, G., Kovács, J., 2016. The physics of wind-blown loess: Implications for grain size proxy interpretations in Quaternary paleoclimate studies. Earth-Science Reviews 154, 247–278. doi:10.1016/j.earscirev.2016.01.006 Ungar, J.E., Haff, P.K., 1987. Steady state saltation in air. Sedimentology 34, 289–299. doi:10.1111/j.1365-3091.1987.tb00778.x Van Der Woerd, J., Tapponnier, P., Ryerson, F.J., Meriaux, A.-S., Meyer, B., Gaudemer, Y., Finkel, R.C., Caffee, M.W., Guoguang, Z., Zhiqin, X., 2002. Uniform postglacial slip-rate along the central 600 km of the Kunlun Fault (Tibet), from 26Al, 10Be, and 14C dating of riser offsets, and climatic origin of the regional morphology. Geophysical Journal International 148, 356–388. doi:10.1046/j.1365-246x.2002.01556.x Vandenberghe, J., Renssen, H., van Huissteden, K., Nugteren, G., Konert, M., Lu, H., Dodonov, A., Buylaert, J.-P., 2006. Penetration of Atlantic westerly winds into Central and East Asia. Quaternary Science Reviews 25, 2380–2389. doi:10.1016/j.quascirev.2006.02.017 Visser, S.M., Sterk, G., Ribolzi, O., 2004. Techniques for simultaneous quantification of wind and water erosion in semi-arid regions. Journal of Arid Environments 59, 699–717. doi:10.1016/j.jaridenv.2004.02.005 Vriend, M., Prins, M.A., 2005. Calibration of modelled mixing patterns in loess grain-size distributions: an example from the north-eastern margin of the Tibetan Plateau, China. Sedimentology 52, 1361–1374. doi:10.1111/j.1365-3091.2005.00743.x Vriend, M., Prins, M.A., Buylaert, J.-P., Vandenberghe, J., Lu, H., 2011. Contrasting dust supply patterns across the north-western Chinese Loess Plateau during the last glacial-interglacial cycle. Quaternary International 240, 167–180. doi:10.1016/j.quaint.2010.11.009 Wallinga, J., Murray, A., Wintle, A., 2000. The single-aliquot regenerative-dose (SAR) protocol applied to coarse-grain feldspar. Radiation Measurements 32, 529–533. doi:10.1016/S1350- 4487(00)00091-3 Wang, B., Ding, Q., 2008. Global monsoon: Dominant mode of annual variation in the tropics. Dynamics of Atmospheres and Oceans 44, 165–183. doi:10.1016/j.dynatmoce.2007.05.002 Wang, H., Chen, J., Zhang, X., Chen, F., 2014. Palaeosol development in the Chinese Loess Plateau as an indicator of the strength of the East Asian summer monsoon: Evidence for a mid-Holocene maximum. Quaternary International 334–335, 155–164. doi:10.1016/j.quaint.2014.03.013 Wang, H., Hong, Y., Lin, Q., Hong, B., Zhu, Y., Wang, Y., Xu, H., 2010. Response of humification degree to monsoon climate during the Holocene from the Hongyuan peat bog, eastern Tibetan Plateau. Palaeogeography, Palaeoclimatology, Palaeoecology 286, 171–177. doi:10.1016/j.palaeo.2009.12.015 Wang, P., Clemens, S., Beaufort, L., Braconnot, P., Ganssen, G., Jian, Z., Kershaw, P., Sarnthein, M., 2005. Evolution and variability of the Asian monsoon system: state of the art and outstanding issues. Quaternary Science Reviews 24, 595–629. doi:10.1016/j.quascirev.2004.10.002 Wang, P.X., Wang, B., Cheng, H., Fasullo, J., Guo, Z.T., Kiefer, T., Liu, Z.Y., 2014. The global monsoon across timescales: coherent variability of regional monsoons. Clim. Past 10, 2007–2052. doi:10.5194/cp-10-2007-2014 Wang, R., Zhang, Y., Wünnemann, B., Biskaborn, B.K., Yin, H., Xia, F., Zhou, L., Diekmann, B., 2015. Linkages between Quaternary climate change and sedimentary processes in Hala Lake, northern Tibetan Plateau, China. Journal of Asian Earth Sciences 107, 140–150. doi:10.1016/j.jseaes.2015.04.008 Wang, W., 1987. Surveying of glacier variations in the northeastern part of the Qinghai–Xizang Plateau., in: Hövermann, J., Wang, W. (Eds.), Reports of the Qinghai–Xizang (Tibet) Plateau. Science Press, Beijing, pp. 22–37. Wang, X., Chu, G., Sheng, M., Zhang, S., Li, J., Chen, Y., Tang, L., Su, Y., Pei, J., Yang, Z., 2016. Millennial-scale Asian summer monsoon variations in South China since the last deglaciation. Earth and Planetary Science Letters 451, 22–30. doi:10.1016/j.epsl.2016.07.006

180

Wang, Y., Cheng, H., Edwards, R.L., He, Y., Kong, X., An, Z., Wu, J., Kelly, M.J., Dykoski, C.A., Li, X., 2005. The Holocene Asian Monsoon: Links to Solar Changes and North Atlantic Climate. Science 308, 854–857. doi:10.1126/science.1106296 Wang, Y., Cheng, H., Edwards, R.L., Kong, X., Shao, X., Chen, S., Wu, J., Jiang, X., Wang, X., An, Z., 2008. Millennial- and orbital-scale changes in the East Asian monsoon over the past 224,000[thinsp]years. Nature 451, 1090–1093. doi:10.1038/nature06692 Wang, Y., Herzschuh, U., Shumilovskikh, L.S., Mischke, S., Birks, H.J.B., Wischnewski, J., Böhner, J., Schlütz, F., Lehmkuhl, F., Diekmann, B., Wünnemann, B., Zhang, C., 2014. Quantitative reconstruction of precipitation changes on the NE Tibetan Plateau since the Last Glacial Maximum – extending the concept of pollen source area to pollen-based climate reconstructions from large lakes. Clim. Past 10, 21–39. doi:10.5194/cp-10-21-2014 Wang, Y., Liu, X., Herzschuh, U., 2010. Asynchronous evolution of the Indian and East Asian Summer Monsoon indicated by Holocene moisture patterns in monsoonal central Asia. Earth-Science Reviews 103, 135–153. doi:10.1016/j.earscirev.2010.09.004 Wang, Y., Yang, H., 2004. Middle Permian palaeobiogeography study in East Kunlun, A’nyêmaqên and Bayan Har. Science in China Series D: Earth Sciences 47, 1120–1126. doi:10.1360/02yd0450 Wang, Y., Zhang, K., Gong, Y., Zhang, Z., Luo, M., 1998. The discovery of early Permian reef belt in east Kunlun and its significance. Chinese Science Bulletin 43, 947–950. doi:10.1007/BF02884620 Wang, Y.J., Cheng, H., Edwards, R.L., An, Z.S., Wu, J.Y., Shen, C.-C., Dorale, J.A., 2001. A High- Resolution Absolute-Dated Late Pleistocene Monsoon Record from Hulu Cave, China. Science 294, 2345–2348. doi:10.1126/science.1064618 Wang, Z.-T., Lai, Z.-P., 2014. A theoretical model on the relation between wind speed and grain size in dust transportation and its paleoclimatic implications. Aeolian Research 13, 105–108. doi:10.1016/j.aeolia.2014.04.003 Wanner, H., Beer, J., Bütikofer, J., Crowley, T.J., Cubasch, U., Flückiger, J., Goosse, H., Grosjean, M., Joos, F., Kaplan, J.O., Küttel, M., Müller, S.A., Prentice, I.C., Solomina, O., Stocker, T.F., Tarasov, P., Wagner, M., Widmann, M., 2008. Mid- to Late Holocene climate change: an overview. Quaternary Science Reviews 27, 1791–1828. doi:10.1016/j.quascirev.2008.06.013 Wanner, H., Solomina, O., Grosjean, M., Ritz, S.P., Jetel, M., 2011. Structure and origin of Holocene cold events. Quaternary Science Reviews 30, 3109–3123. doi:10.1016/j.quascirev.2011.07.010 Webster, P.J., Magaña, V.O., Palmer, T.N., Shukla, J., Tomas, R.A., Yanai, M., Yasunari, T., 1998. Monsoons: Processes, predictability, and the prospects for prediction. Journal of Geophysical Research: Oceans 103, 14451–14510. doi:10.1029/97JC02719 Weltje, G.J., 1997. End-member modeling of compositional data: Numerical-statistical algorithms for solving the explicit mixing problem. Mathematical Geology 29, 503–549. doi:10.1007/BF02775085 Weltje, G.J., Prins, M.A., 2007. Genetically meaningful decomposition of grain-size distributions. Sedimentary Geology 202, 409–424. doi:10.1016/j.sedgeo.2007.03.007 Weltje, G.J., Prins, M.A., 2003. Muddled or mixed? Inferring palaeoclimate from size distributions of deep-sea clastics. Sedimentary Geology 162, 39–62. doi:10.1016/S0037-0738(03)00235-5 Werner, B.T., 1990. A Steady-State Model of Wind-Blown Sand Transport. The Journal of Geology 98, 1–17. Weynell, M., Wiechert, U., Zhang, C., 2016. Chemical and isotopic (O, H, C) composition of surface waters in the catchment of Lake Donggi Cona (NW China) and implications for paleoenvironmental reconstructions. Chemical Geology 435, 92–107. doi:10.1016/j.chemgeo.2016.04.012 Wiggs, G.F.S., 1997. Sediment mobilisation by the wind, in: Thomas, D.S.G. (Ed.), Arid Zone Geomorphology. Process, Form and Change in Drylands. Wiley, Chichester, pp. 351–372. Wintle, A.G., 2008. Luminescence dating: where it has been and where it is going. Boreas 37, 471– 482. doi:10.1111/j.1502-3885.2008.00059.x

181

Wintle, A.G., Murray, A.S., 2006. A review of quartz optically stimulated luminescence characteristics and their relevance in single-aliquot regeneration dating protocols. Radiation Measurements 41, 369–391. doi:10.1016/j.radmeas.2005.11.001 Wischnewski, J., Mischke, S., Wang, Y., Herzschuh, U., 2011. Reconstructing climate variability on the northeastern Tibetan Plateau since the last Lateglacial – a multi-proxy, dual-site approach comparing terrestrial and aquatic signals. Quaternary Science Reviews 30, 82–97. doi:10.1016/j.quascirev.2010.10.001 Wu, G., Liu, Y., Dong, B., Liang, X., Duan, A., Bao, Q., Yu, J., 2012. Revisiting Asian monsoon formation and change associated with Tibetan Plateau forcing: I. Formation. Climate Dynamics 39, 1169–1181. doi:10.1007/s00382-012-1334-z Wu, G.X., Liu, Y., Zhu, X., Li, W., Ren, R., Duan, A., Liang, X., 2009. Multi-scale forcing and the formation of subtropical desert and monsoon. Ann. Geophys. 27, 3631–3644. doi:10.5194/angeo-27-3631-2009 Wu, Y., Cui, Z., Liu, G., Ge, D., Yin, J., Xu, Q., Pang, Q., 2001. Quaternary geomorphological evolution of the Kunlun Pass area and uplift of the Qinghai-Xizang (Tibet) Plateau. Geomorphology 36, 203–216. doi:10.1016/S0169-555X(00)00057-X Wünnemann, B., Demske, D., Tarasov, P., Kotlia, B.S., Reinhardt, C., Bloemendal, J., Diekmann, B., Hartmann, K., Krois, J., Riedel, F., Arya, N., 2010. Hydrological evolution during the last 15 kyr in the Tso Kar lake basin (Ladakh, India), derived from geomorphological, sedimentological and palynological records. Quaternary Science Reviews 29, 1138–1155. doi:10.1016/j.quascirev.2010.02.017 Wünnemann, B., Wagner, J., Zhang, Y., Yan, D., Wang, R., Shen, Y., Fang, X., Zhang, J., 2012. Implications of diverse sedimentation patterns in Hala Lake, Qinghai Province, China for reconstructing Late Quaternary climate. Journal of Paleolimnology 48, 725–749. doi:10.1007/s10933-012-9641-2 Wyrwoll, K.-H., Wei, J., Lin, Z., Shao, Y., He, F., 2016. Cold surges and dust events: Establishing the link between the East Asian Winter Monsoon and the Chinese loess record. Quaternary Science Reviews 149, 102–108. doi:10.1016/j.quascirev.2016.04.015 Xiao, J., Porter, S.C., An, Z., Kumai, H., Yoshikawa, S., 1995. Grain Size of Quartz as an Indicator of Winter Monsoon Strength on the Loess Plateau of Central China during the Last 130,000 Yr. Quaternary Research 43, 22–29. doi:10.1006/qres.1995.1003 Xu, H., Hong, Y., Lin, Q., Zhu, Y., Hong, B., Jiang, H., 2006. Temperature responses to quasi-100-yr solar variability during the past 6000 years based on δ18O of peat cellulose in Hongyuan, eastern Qinghai–Tibet plateau, China. Palaeogeography, Palaeoclimatology, Palaeoecology 230, 155–164. doi:10.1016/j.palaeo.2005.07.012 Xu, M., Chang, C.-P., Fu, C., Qi, Y., Robock, A., Robinson, D., Zhang, H., 2006. Steady decline of east Asian monsoon winds, 1969–2000: Evidence from direct ground measurements of wind speed. Journal of Geophysical Research: Atmospheres 111, n/a-n/a. doi:10.1029/2006JD007337 Xu, X., Yi, C., 2014. Little Ice Age on the Tibetan Plateau and its bordering mountains: Evidence from moraine chronologies. Global and Planetary Change 116, 41–53. doi:10.1016/j.gloplacha.2014.02.003 Xue, X., Guo, J., Han, B., Sun, Q., Liu, L., 2009. The effect of climate warming and permafrost thaw on desertification in the Qinghai–Tibetan Plateau. Geomorphology 108, 182–190. doi:10.1016/j.geomorph.2009.01.004 Yan, C.Z., Song, X., Zhou, Y.M., Duan, H.C., Li, S., 2009. Assessment of aeolian desertification trends from 1975’s to 2005’s in the watershed of the Longyangxia Reservoir in the upper reaches of China’s Yellow River. Geomorphology 112, 205–211. doi:10.1016/j.geomorph.2009.06.003 Yan, D., Wünnemann, B., 2014. Late Quaternary water depth changes in Hala Lake, northeastern Tibetan Plateau, derived from ostracod assemblages and sediment properties in multiple sediment records. Quaternary Science Reviews 95, 95–114. doi:10.1016/j.quascirev.2014.04.030

182

Yan, P., Shi, P., Gao, S., Chen, L., Zhang, X., Bai, L., 2002. 137Cs dating of lacustrine sediments and human impacts on Dalian Lake, Qinghai Province, China. CATENA 47, 91–99. doi:10.1016/S0341-8162(01)00193-X Yang, B., Bräuning, A., Dong, Z., Zhang, Z., Keqing, J., 2008. Late Holocene monsoonal temperate glacier fluctuations on the Tibetan Plateau. Global and Planetary Change 60, 126–140. doi:10.1016/j.gloplacha.2006.07.035 Yang, B., Bräuning, A., Liu, J., Davis, M.E., Yajun, S., 2009. Temperature changes on the Tibetan Plateau during the past 600 years inferred from ice cores and tree rings. Global and Planetary Change 69, 71–78. doi:10.1016/j.gloplacha.2009.07.008 Yang, B., Bräuning, A., Shi, Y., 2003. Late Holocene temperature fluctuations on the Tibetan Plateau. Quaternary Science Reviews 22, 2335–2344. doi:10.1016/S0277-3791(03)00132-X Yang, B., Qin, C., Wang, J., He, M., Melvin, T.M., Osborn, T.J., Briffa, K.R., 2014. A 3,500-year tree-ring record of annual precipitation on the northeastern Tibetan Plateau. Proceedings of the National Academy of Sciences 111, 2903–2908. doi:10.1073/pnas.1319238111 Yang, K., Wu, H., Qin, J., Lin, C., Tang, W., Chen, Y., 2014. Recent climate changes over the Tibetan Plateau and their impacts on energy and water cycle: A review. Global and Planetary Change 112, 79–91. doi:10.1016/j.gloplacha.2013.12.001 Yang, L., Wang, T., Zhou, J., Lai, Z., Long, H., 2012. OSL chronology and possible forcing mechanisms of dune evolution in the Horqin dunefield in northern China since the Last Glacial Maximum. Quaternary Research 78, 185–196. doi:10.1016/j.yqres.2012.05.002 Yang, S.L., Ding, Z.L., 2004. Comparison of particle size characteristics of the Tertiary ‘red clay’ and Pleistocene loess in the Chinese Loess Plateau: implications for origin and sources of the ‘red clay.’ Sedimentology 51, 77–93. doi:10.1046/j.1365-3091.2003.00612.x Yang, X., Li, H., Conacher, A., 2012. Large-scale controls on the development of sand seas in northern China. Quaternary International 250, 74–83. doi:10.1016/j.quaint.2011.03.052 Yang, X., Rost, K.T., Lehmkuhl, F., Zhenda, Z., Dodson, J., 2004. The evolution of dry lands in northern China and in the Republic of Mongolia since the Last Glacial Maximum. Quaternary International 118–119, 69–85. doi:10.1016/S1040-6182(03)00131-9 Yang, X., Scuderi, L., Paillou, P., Liu, Z., Li, H., Ren, X., 2011. Quaternary environmental changes in the drylands of China – A critical review. Quaternary Science Reviews 30, 3219–3233. doi:10.1016/j.quascirev.2011.08.009 Yang, X., Wang, X., Liu, Z., Li, H., Ren, X., Zhang, D., Ma, Z., Rioual, P., Jin, X., Scuderi, L., 2013. Initiation and variation of the dune fields in semi-arid China – with a special reference to the Hunshandake Sandy Land, Inner Mongolia. Quaternary Science Reviews 78, 369–380. doi:10.1016/j.quascirev.2013.02.006 Yao, T., Thompson, L., Yang, W., Yu, W., Gao, Y., Guo, X., Yang, X., Duan, K., Zhao, H., Xu, B., Pu, J., Lu, A., Xiang, Y., Kattel, D.B., Joswiak, D., 2012. Different glacier status with atmospheric circulations in Tibetan Plateau and surroundings. Nature Clim. Change 2, 663–667. doi:10.1038/nclimate1580 Ye, D.-Z., Wu, G.-X., 1998. The role of the heat source of the Tibetan Plateau in the general circulation. Meteorology and Atmospheric Physics 67, 181–198. doi:10.1007/BF01277509 Yi, C., Chen, H., Yang, J., Liu, B., Fu, P., Liu, K., Li, S., 2008. Review of Holocene glacial chronologies based on radiocarbon dating in Tibet and its surrounding mountains. J. Quaternary Sci. 23, 533–543. doi:10.1002/jqs.1228 Yim, S.-Y., Wang, B., Liu, J., Wu, Z., 2014. A comparison of regional monsoon variability using monsoon indices. Climate Dynamics 43, 1423–1437. doi:10.1007/s00382-013-1956-9 Yin, A., Harrison, T.M., 2000. Geologic Evolution of the Himalayan-Tibetan Orogen. Annu. Rev. Earth Planet. Sci. 28, 211–280. doi:10.1146/annurev.earth.28.1.211 You, Q., Kang, S., Pepin, N., Flügel, W.-A., Sanchez-Lorenzo, A., Yan, Y., Zhang, Y., 2010. Climate warming and associated changes in atmospheric circulation in the eastern and central Tibetan Plateau from a homogenized dataset. Global and Planetary Change 72, 11–24. doi:10.1016/j.gloplacha.2010.04.003

183

Yu, J.Q., Kelts, K.R., 2002. Abrupt changes in climatic conditions across the late-glacial/Holocene transition on the N. E. Tibet-Qinghai Plateau: evidence from Lake Qinghai, China. Journal of Paleolimnology 28, 195–206. doi:10.1023/A:1021635715857 Yu, L., Lai, Z., 2014. Holocene climate change inferred from stratigraphy and OSL chronology of aeolian sediments in the Qaidam Basin, northeastern Qinghai–Tibetan Plateau. Quaternary Research 81, 488–499. doi:10.1016/j.yqres.2013.09.006 Yu, L., Lai, Z., 2012. OSL chronology and palaeoclimatic implications of aeolian sediments in the eastern Qaidam Basin of the northeastern Qinghai-Tibetan Plateau. Palaeogeography, Palaeoclimatology, Palaeoecology 337–338, 120–129. doi:10.1016/j.palaeo.2012.04.004 Yu, L., Lai, Z., An, P., Pan, T., Chang, Q., 2015. Aeolian sediments evolution controlled by fluvial processes, climate change and human activities since LGM in the Qaidam Basin, Qinghai- Tibetan Plateau. Quaternary International 372, 23–32. doi:10.1016/j.quaint.2014.09.043 Yu, L., Lai, Z.-P., An, P., 2013. OSL chronology and paleoclimatic implications of paleodunes in the middle and southwestern Qaidam Basin, Qinghai–Tibetan Plateau. Sci. Cold Arid Regions 5, 0211–0219. Yu, R., Wang, B., Zhou, T., 2004. Tropospheric cooling and summer monsoon weakening trend over East Asia. Geophys. Res. Lett. 31, n/a-n/a. doi:10.1029/2004GL021270 Yu, Y., Yang, T., Li, J., Liu, J., An, C., Liu, X., Fan, Z., Lu, Z., Li, Y., Su, X., 2006. Millennial-scale Holocene climate variability in the NW China drylands and links to the tropical Pacific and the North Atlantic. Palaeogeography, Palaeoclimatology, Palaeoecology 233, 149–162. doi:10.1016/j.palaeo.2005.09.008 Zan, J., Fang, X., Yan, M., Zhang, W., Zhang, Z., 2015. Magnetic variations in surface soils in the NE Tibetan Plateau indicating the climatic boundary between the Westerly and East Asian summer monsoon regimes in NW China. Global and Planetary Change 130, 1–6. doi:10.1016/j.gloplacha.2015.03.008 Zhang, C., Mischke, S., 2009. A Lateglacial and Holocene lake record from the Nianbaoyeze Mountains and inferences of lake, glacier and climate evolution on the eastern Tibetan Plateau. Quaternary Science Reviews 28, 1970–1983. doi:10.1016/j.quascirev.2009.03.007 Zhang, H.-P., Craddock, W.H., Lease, R.O., Wang, W., Yuan, D.-Y., Zhang, P.-Z., Molnar, P., Zheng, D.- W., Zheng, W.-J., 2012. Magnetostratigraphy of the Neogene Chaka basin and its implications for mountain building processes in the north-eastern Tibetan Plateau. Basin Research 24, 31– 50. doi:10.1111/j.1365-2117.2011.00512.x Zhang, J., Chen, F., Holmes, J.A., Li, H., Guo, X., Wang, J., Li, S., Lü, Y., Zhao, Y., Qiang, M., 2011. Holocene monsoon climate documented by oxygen and carbon isotopes from lake sediments and peat bogs in China: a review and synthesis. Quaternary Science Reviews 30, 1973–1987. doi:10.1016/j.quascirev.2011.04.023 Zhang, J.F., Zhou, L.P., 2007. Optimization of the ‘double SAR’ procedure for polymineral fine grains. Radiation Measurements 42, 1475–1482. doi:10.1016/j.radmeas.2007.06.007 Zhang, P., Cheng, H., Edwards, R.L., Chen, F., Wang, Y., Yang, X., Liu, J., Tan, M., Wang, X., Liu, J., An, C., Dai, Z., Zhou, J., Zhang, D., Jia, J., Jin, L., Johnson, K.R., 2008. A Test of Climate, Sun, and Culture Relationships from an 1810-Year Chinese Cave Record. Science 322, 940–942. doi:10.1126/science.1163965 Zhang, P.-Z., Shen, Z., Wang, M., Gan, W., Bürgmann, R., Molnar, P., Wang, Q., Niu, Z., Sun, J., Wu, J., Hanrong, S., Xinzhao, Y., 2004. Continuous deformation of the Tibetan Plateau from global positioning system data. Geology 32, 809–812. doi:10.1130/G20554.1 Zhang, Q.-B., Cheng, G., Yao, T., Kang, X., Huang, J., 2003. A 2,326-year tree-ring record of climate variability on the northeastern Qinghai-Tibetan Plateau. Geophys. Res. Lett. 30, n/a-n/a. doi:10.1029/2003GL017425 Zhang, R., Jiang, D., Liu, X., Tian, Z., 2012. Modeling the climate effects of different subregional uplifts within the Himalaya-Tibetan Plateau on Asian summer monsoon evolution. Chinese Science Bulletin 57, 4617–4626. doi:10.1007/s11434-012-5284-y

184

Zhang, R., Jiang, D., Zhang, Z., Yu, E., 2015. The impact of regional uplift of the Tibetan Plateau on the Asian monsoon climate. Palaeogeography, Palaeoclimatology, Palaeoecology 417, 137–150. doi:10.1016/j.palaeo.2014.10.030 Zhang, X.Y., Arimoto, R., Cao, J.J., An, Z.S., Wang, D., 2001. Atmospheric dust aerosol over the Tibetan Plateau. J. Geophys. Res. 106, 18471–18476. doi:10.1029/2000JD900672 Zhao, C., Liu, Z., Rohling, E.J., Yu, Z., Liu, W., He, Y., Zhao, Y., Chen, F., 2013. Holocene temperature fluctuations in the northern Tibetan Plateau. Quaternary Research 80, 55–65. doi:10.1016/j.yqres.2013.05.001 Zhao, C., Yu, Z., Zhao, Y., Ito, E., 2009. Possible orographic and solar controls of Late Holocene centennial-scale moisture oscillations in the northeastern Tibetan Plateau. Geophys. Res. Lett. 36, n/a-n/a. doi:10.1029/2009GL040951 Zhao, H., Li, S.-H., 2005. Internal dose rate to K-feldspar grains from radioactive elements other than potassium. Radiation Measurements 40, 84–93. doi:10.1016/j.radmeas.2004.11.004 Zhao, L., Jin, H., Li, C., Cui, Z., Chang, X., Marchenko, S.S., Vandenberghe, J., Zhang, T., Luo, D., Guo, D., Liu, G., Yi, C., 2014. The extent of permafrost in China during the local Last Glacial Maximum (LLGM). Boreas 43, 688–698. doi:10.1111/bor.12049 Zhao, Y., Yu, Z., Chen, F., 2009. Spatial and temporal patterns of Holocene vegetation and climate changes in arid and semi-arid China. Quaternary International 194, 6–18. doi:10.1016/j.quaint.2007.12.002 Zhao, Y., Yu, Z., Chen, F., Ito, E., Zhao, C., 2007. Holocene vegetation and climate history at Hurleg Lake in the Qaidam Basin, northwest China. Review of Palaeobotany and Palynology 145, 275–288. doi:10.1016/j.revpalbo.2006.12.002 Zhao, Y., Yu, Z., Zhao, W., 2011. Holocene vegetation and climate histories in the eastern Tibetan Plateau: controls by insolation-driven temperature or monsoon-derived precipitation changes? Quaternary Science Reviews 30, 1173–1184. doi:10.1016/j.quascirev.2011.02.006 Zhou, J., Zhu, Y., Yuan, C., 2012. Origin and lateral migration of linear dunes in the Qaidam Basin of NW China revealed by dune sediments, internal structures, and optically stimulated luminescence ages, with implications for linear dunes on Titan. Geological Society of America Bulletin 124, 1147–1154. doi:10.1130/B30550.1 Zhou, T., Gong, D., Li, J., Li, B., 2009. Detecting and understanding the multi-decadal variability of the East Asian Summer Monsoon Recent progress and state of affairs. Meteorologische Zeitschrift 18, 455–467. doi:10.1127/0941-2948/2009/0396 Zhu, L., Lü, X., Wang, J., Peng, P., Kasper, T., Daut, G., Haberzettl, T., Frenzel, P., Li, Q., Yang, R., Schwalb, A., Mäusbacher, R., 2015. Climate change on the Tibetan Plateau in response to shifting atmospheric circulation since the LGM. Scientific Reports 5, 13318.

185