Clay mineralogical and geochemical constraints on late Pleistocene weathering processes of the Qaidam Basin, northern Tibetan Plateau

Journal of Asian Earth Sciences 127 (2016) 267–280

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Clay mineralogical and geochemical constraints on late Pleistocene weathering processes of the Qaidam Basin, northern Tibetan Plateau WeiLiang Miao a,b,⇑, QiShun Fan a, HaiCheng Wei a, XiYing Zhang a, HaiZhou Ma a a b

Qinghai Institute of Salt Lakes, Chinese Academy of Sciences, Xining 810008, China University of Chinese Academy of Sciences, Beijing 100049, China

a r t i c l e

i n f o

Article history: Received 31 December 2015 Received in revised form 9 June 2016 Accepted 14 June 2016 Available online 27 June 2016 Keywords: Clay minerals Weathering process Major-element geochemistry Late Pleistocene Qarhan Salt Lake The Qaidam Basin

a b s t r a c t At the Qarhan Salt Lake (QSL) on the central-eastern Qaidam Basin, northern Tibetan Plateau, Quaternary lacustrine sediments have a thickness of over 3000 m and mainly composed of organic-rich clay and silty clay with some silt halite and halite. In this study, a 102-m-long sediment core (ISL1A) was obtained from the QSL. Combining with AMS 14C and 230Th dating, clay minerals and major-element concentrations of ISL1A were used to reconstruct the weathering process and trend of the QSL since late Pleistocene. The results reveal that the clay mineral from <2 lm fraction in ISL1A is composed of illite (47–77%), chlorite (8–27%), smectite (including illite-smectite mixed layers, 3–29%) and kaolinite (2–11%). Such clay mineral assemblages in ISL1A derived primarily from felsic igneous rocks, gneisses and schists of Eastern Kunlun Mountains on the south of the QSL. The abundance of illite mineral displays an opposite fluctuation trending with that of smectite, chlorite and kaolinite mineral in ISL1A, which is significantly different from the monsoon-controlled regions. Moreover, higher values of illite, kaolinite/chlorite and illite/chlorite ratios, and lower values of smectite, chlorite and kaolinite minerals occurred in 83–72.5 ka, 68.8–54 ka, 32–24 ka, corresponding to late MIS 5, late MIS 4, early MIS 3 and late MIS 3, respectively. These three phases were almost similarly changed with oxygen isotopes of authigenic carbonates and pollen records in ISL1A, which implies that stronger chemical weathering corresponds to higher effective moisture periods of source region in the Qaidam Basin. Based on chemical weathering index and (Al2O3-(CaO + Na2O)-K2O) diagram, chemical weathering degree in this study area takes a varying process from low to intermediate on the whole. Ó 2016 Elsevier Ltd. All rights reserved.

1. Introduction The uplift of the Tibetan Plateau (TP) and its environmental effects during the Cenozoic have been widely studied since tectonic uplift
weathering hypothesis was proposed by Raymo et al. (1988). On the one hand, the uplift of TP fostered Asian monsoon climate systems and the evolution of river basins in East and South Asia; On the other hand, weathering or erosion process from these river basins on the TP plays a key role in surface processes and geochemical cycles in earth supergene environments, including global carbon cycle and chemical composition of the oceans (Berner et al., 1983; Raymo and Ruddiman, 1992; Kump et al., 2000). Previous studies have focused on main drainage basins around the TP, such as the Mekong drainage basin (Liu et al., 2004; Noh et al., 2009; Borges et al., 2008), the Salween ⇑ Corresponding author at: Qinghai Institute of Salt Lakes, Chinese Academy of Sciences, Xining 810008, China. E-mail address: [email protected] (W. Miao). http://dx.doi.org/10.1016/j.jseaes.2016.06.013 1367-9120/Ó 2016 Elsevier Ltd. All rights reserved.

drainage basin (Noh et al., 2009; Borges et al., 2008), the Yangze River catchment (Yang et al., 2006; Wang and Yang, 2012; Shao and Yang, 2012) and the Yellow River drainage basin (Li, 2003; Yang et al., 2004). And the results indicate that there is a strong relationship between weathering/erosion and the climate change during glacial and interglacial periods (Liu et al., 2004; Yang et al., 2006; Borges et al., 2008). However, the research on weathering process in the source region of the Qaidam Basin on the northern TP is still limited and unclear due partly to limited numbers of long continuous paleoclimatic records. Most previous records in the Qaidam Basin extend only as far back as the late glacial or early Holocene (Zhao et al., 2007, 2008; Liu et al., 2008), although there are a few low resolution paleoclimate records extending back to the last glacial or the late Pleistocene (Chen and Bowler, 1986; Chen et al., 1990; Huang and Chen, 1990; Zhang et al., 1993; Yang et al., 1995). In recent works, high resolution pollen and oxygen isotope records of lacustrine sediments in ISL1A were obtained spanning the last glacial period (Fan et al., 2014a; Wei et al., 2015). The results show

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that three drier climatic intervals at 90
80 ka, 52
38 ka and 10
9 ka, correspond to the late MIS 5, middle MIS 3 and early Holocene, respectively, were identified based on the oxygen isotope record (Fan et al., 2014a). Fan et al. (2014a) also argued that ‘‘a uniform Qarhan mega-paleolake” spanning the period of 44–22 ka mentioned by predecessors may not have existed because a relatively dry climate with high d18O values corresponded to this period. Furthermore, the pollen results reported that the identification of the primary wetter periods of late MIS 5, early MIS 4, early and late MIS 3 in the Qarhan Salt Lake (QSL) region (Wei et al., 2015). Here we present a 500-yr-resolution clay-mineral record in ISL1A to develop a better understanding of the paleoclimate information recorded in deposits of QSL in the Qaidam Basin. Because clay minerals in sediments present a record of weathering conditions in the watershed, they have often been used as a proxy in paleoclimate reconstructions (e.g., Chamley, 1989). In this study, we attempted to establish the relationship between weathering/ erosion and the climate in QSL region since the late Pleistocene. 2. Regional setting The Qaidam Basin is the largest intermontane endorheic basin on the northern TP (Fig. 1). It located at the northern TP with an average elevation of 2800 m a.s.l., and influenced by both of the East Asian summon monsoon, the Westerlies and the Siberian high (Fig. 1A). The basin area is 120,000 km2 and its catchment covers approximately 250,000 km2. It is enclosed by three large mountain belts: the Aljun Mountains to the northwest, the eastern Kunlun Mountains to the south and the Qilian Mountains to the northeast. It contains 27 salt lakes, many of which were linked via thick salt beds and brines, and a very thick (>3000 m) sequence of Quaternary sediments deposited in the central Qaidam Basin (Chen and Bowler, 1986; Zhang, 1987; Liu et al., 1998) (Fig. 1B). As the largest playa of the Qaidam Basin, QSL region is the depocenter of the Qaidam Basin during the Quaternary, covering an area of approximately 5856 km2 (Chen and Bowler, 1985, 1986; Liu et al., 1998) (Fig. 1B). It has a surface elevation of 2675 m a.s.l. at the lowest point in the depression (Chen and Bowler, 1985). In contrast, the mountains surrounding the QSL reach an average elevation of 3500
4500 m a.s.l. (Zhang, 1987), which results in the basin-and-range topography in the study area (Fig. 1B). From west to east, the QSL is divided into four sections: Bieletan, Dabuxun, Qarhan and Huobuxun (Fig. 1C). There are 9 residues brine lakes with different sizes scattered in the playa, and field investigation indicates that there are 18 rivers originating from the Kunlun Mountain to feed into the QSL (Yuan et al., 1995) (Fig. 1C). The QSL area is also one of the driest places in the world with a mean annual temperature of 5.2 °C, the mean annual precipitation level is approximately 24 mm, while the mean annual evaporation level reaches 3564 mm (Yang et al., 1993; Fan et al., 2014a). However, the mean annual precipitation on the northern slopes of the Kunlun Mountain is slightly greater than that of QSL. Additionally, the average wind speed is 4.3 m/s and relative moisture is 27.7% (Yu et al., 2009). The landscapes of the QSL area can be divided into 6 major types (Yang et al., 1993) (Fig. 1C): (1) Mountains are present primarily on the north and south sides of the lake. The Eastern Kunlun Mountains, located along the south side of the lake, are underlain primarily by Proterozoic gneiss and schist, and lower Paleozoic greenschist and carbonate rocks that were widely intruded by polyphasic granitic rocks of the Caledonian, Variscan, Indosinian and Yanshanian orogenies. The Xitieshan and Amunike Mountains, located on the north side of the lake, are underlain primarily by Proterozoic gneiss, schist and marble and Paleozoic biotite schist with intermediate
mafic to felsic volcanic rocks and volcaniclastic

rocks that were intruded by polyphasic ultrabasic to felsic igneous rocks of the Caledonian, Variscan and Indosinian orogenies. Fissures in the bedrock produce local bedrock aquifer. (2) Gobi zones are located primarily on the south side of the piedmont and consist of ancient alluvial
pluvial fans. The lithology of this region is dominated by glutenite and gravel strata, which constitute the main phreatic aquifer of the piedmont clinoplains. (3) Alluvial
pluvial plains are primarily distributed as east
west
trending belts with widths of 10
15 km on the south side of QSL. These plains are used for farming and grazing and contain numerous springs. Sand, gravelly sand, clayey sand and clay are the primary lithologic types in this area and constitutes the overlapped artesian aquifer of the alluvial
lacustrine plains. (4) Lacustrine plains are primarily along the east, south and west sides of QSL and have widths of several to tens of kilometers. Centripetal radial networks of dry gullies extend across this area, and the lithology is primarily composed of silty sand, clayey sand and clay soil. (5) Playas dominate the flat, saline desert landscapes of the QSL region. The surfaces of playas are covered with various types of salt crust composed primarily of halite. There are also intercrystalline brine deposits. (6) Brine lakes dominate the western, southern and eastern edges of the playas and formed primarily in the convergence areas of rivers (including intercrystalline brine deposits) and lakes. 3. Materials and methods 3.1. Sediment core and dating The core ISL1A (37°30 5000 N, 94°430 4100 E) was obtained from the Bieletan section of Qarhan playa (Fig. 1B). The lithostratigraphy of the core alternates primarily between evaporite layers and silt-clay sediment layers from 0 to 51.1 m, and between silt and organic-rich clay sediments from 51.1 to 102 m (Fig. 2). Twelve clay samples containing dark organic matter were collected from the upper 54.5 m of the core for accelerator mass spectrometry (AMS) 14C dating, and eight halite samples were collected from the upper 46.0 m of the core for 230Th dating. An assessment and comparison of these 230Th and AMS 14C ages have been discussed in detail by Fan et al. (2014b). In addition, three carbonate samples were collected from 64.5 to 98.9 m of the core for isochron 230Th dating. The impure carbonate samples were dissolved with 0.1 M HCl, 1 M HCl and HF-HClO4, respectively, to obtain isochron 230Th ages. Chemical procedures followed those described by Ma et al. (2004, 2010). 230Th ages of carbonate deposits were determined using an OctêteÒ plus alpha spectrometer, with a vacuum of 20 mT and an energy resolution (FWHM) of approximately 25 keV at 5.15 MeV. Analyses were conducted at the U-series Dating Laboratory of The Institute of Geology and Geophysics, Chinese Academy of Sciences. 3.2. Clay mineralogy analysis A total of 154 samples were taken at intervals of approximately 50 cm from the core for clay mineral analysis. Of these, 53 samples were taken from the evaporation sequence, and the rest were taken from lacustrine silt-clay sedimentary strata. The clay fraction (<2 lm) were isolated from deflocculated suspensions using Stoke’s Law in a setting beaker. Relative clay mineral contents in the core samples were determined with X-ray diffraction (XRD) according to the procedures and methods described in detail by Petschick et al. (1996) and Liu et al. (2004). All clay samples were processed on oriented mounts of non-calcareous clay-sized (<2 lm) particles (Holtzapffel, 1985). And three XRD runs were performed, following air-drying, ethylene-glycol salvation for 24 h, and heating at 490 °C for 2 h. The measurements were conducted on PANalytical X’ Pert PRO diffractometer with CuKa

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269

Fig. 1. Schematic geological map of the Qaidam Basin and location of the study area. In (A), yellow square denotes the location of Qaidam Basin, the purple line marks the modern Asian summer monsoon limit (modified from Gao, 1962), and the arrows display the dominant circulation systems of Westerlies, East Asian summer monsoon and East Asian winter monsoon. In (B), the DEM map reveals the geomorphology of the Qaidam Basin. The yellow square indicates the Qarhan area. (C), Geological map showing the location of Qarhan Salt Lake, the lithofacies of the surrounding mountains and the river-lake systems of the Qarhan Salt Lake (modified from Qinghai province 1: 1,000,000 scale geological map). (For interpretation of the references to color in this figure legend, the reader is referred to the web version of this article.)

radiation (40 kV, 30 mA). The samples were X-rayed in the range 3–35°2h with a scan speed of 0.02°2h/s. Identification of clay minerals was made mainly according to the position of the (0 0 1) series of basal reflections on the three XRD diagrams. Semi-quantitative estimates of peak areas of the basal reflections for the main clay mineral groups of smectite (17 Å) (e.g., Fig. 3A), illite (10 Å), and kaolinite/chlorite (7.1 Å) were carried out on the glycolated curve (Fig. 3) using the MacDiff software application (Petschick, 2000). Relative proportions of kaolinite and chlorite were determined based on the ratio from

the 3.57/3.54 Å peak areas. Furthermore, swelling clays are evidenced by expansion after EG solvation, but in most of the samples, there is no well-defined peak but rather a shoulder at the 14 Å reflection, and it expands toward the lower angles (e.g., Fig. 3B–D). This behavior indicates the occurrence of illite within swelling random illite—smectite mixed layers (Thorez, 1976). Following the laboratory routine, the weighting factors introduced by Biscaye (1965) or Holtzapffel (1985) are not used when generating the relative weight percentages of each clay mineral because such regular weighting factors actually do not exist by systemic

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Fig. 2. Lithology and age-depth model based on U/Th and AMS

experiences (Gensac, 2008). No effort was made to quantify mixedlayer clay minerals. The illite crystallinity (IC) was obtained from the half height width of the 10 Å peak on the glycolated curve. Lower values represent the higher crystallinity, characteristic of weak hydrolysis in continental sources and arid and cold climate conditions (Chamley, 1989; Krumm and Buggisch, 1991). Replicate analyses of 10 randomly selected samples yielded a precision of ±2% (2r). Based upon the XRD method, the semiquantitative evaluation of each clay mineral has an accuracy of 5%. 3.3. Geochemical analysis Fifteen samples were selected for major-element determinations in bulk (<63 lm) particles based on the results of clay mineral analysis. The measurement was performed by Inductively Coupled Plasma-Optical Emission (ICP-OES) using an IRIS

14

C dating of core ISL1A (modified from Wei et al., 2015).

Advantage at the State Key Laboratory of Marine Geology, Tongji University. Carbonate-free particles (<63 lm) were wet sieved to remove potential local coarse grains to minimize the grain size effect on chemical compositions (Loring and Asmund, 1996; Datta and Subramanian, 1998). Our experiments indicate only a few particles >63 lm in the original argillaceous samples. Carbonate was removed using 0.5% HCl to purify the granule silicate particles. The fine fraction (<63 lm) was wet sieved from core sediments in deionized water, dried at 60 °C in a clean oven. Approximately 30–40 mg of pre-prepared sediments were heated under 600 °C to obtain the loss of ignition (LOI) and then were dissolved using a mixture solution of HNO3 + HF on a hotplate. The eluted sample was diluted by 2% HNO3 for the major element and trace element measurement. This preparation procedure does not allow Si to be measured directly from ICP-OES. Concentrations of SiO2 were then obtained by using 100% to minimize all other

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271

Fig. 3. Multiple X-ray diffractograms of typical samples from the core ISL1A. Curves B, C and D show that these samples contain predominantly illite with minor amounts of chlorite and random illite—smectite mixed layers and very scarce kaolinite. Curve A shows that the sample contains predominantly neoformed smectite with little illite, chlorite and kaolinite. See Table 1 for clay mineral assemblages of the core.

major-element concentrations and LOI. Duplicate measurements of the geostandards GSR-5, GSR-6 and GSD-9 show that relative deviations between measured and certified values are generally less than 5%.

Additionally, 163 samples in core ISL1A were selected for Ti abundance, which was measured with an X-ray fluorescence spectrometer (XRF) at Qinghai Institute of Salt Lakes.

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4. Results

4.3. Clay mineralogical results

4.1. Chronology

The clay mineral assemblages of core ISL1A dominantly consist of illite, chlorite, smectite (including illite-smectite mixed layers), neoformed smectite (Fig. 3A) and kaolinite. The clay mineral record is dominated by illite (47–77%), followed by significantly less of smectite (including illite-smectite mixed layers) (3–29%), chlorite (8–27%) and traces of kaolinite (2–11%). Most of the clay assemblage changes, in part, are controlled by the opposite evolution between smectite and illite (r2illite-smectites = 0.85) and between illite and chlorite (r2illite-chlorite = 0.39). A decrease in illite is partly counterbalanced by an increase in smectite (including illite-smectite mixed layers) and chlorite, with the changes in other clay species being minor (Fig. 4). The illite crystallinity is more variable and relatively higher in lacustrine clastic sediments, ranging between 0.27 and 0.84°D2h. In order to aid the interpretation of the clay profiles (Fig. 4), we identify 7 units or clay zones (CZ) based mainly on variation trends of illite and chlorite abundances. These zones are labeled from 1 to 7 from the base of the core at 102.03 m (94.7 ka) to the top at 1.36 m (9.4 ka) (Table 1, Fig. 4). Unit CZ1 (102.03–89.23 m, ca. 94.7–83.7 ka) is marked by the opposite trend and relatively large fluctuation of illite (mean of 61%) and smectite (14%) abundances. The chlorite abundance remains stable and high (mean value of 19%), whereas the kaolinite (mean of 7%) shows a gradual upward decreasing trend. A sharp change at 89.23 m (ca. 83.7 ka) also exists in illite, smectite and kaolinite abundance curves. The IC value varies between 0.29 and 0.51°D2h, with relatively low values and an upward decreasing trend in this stage. In unit CZ2 (89.23–72.03 m, ca. 83.7–68.9 ka), the illite abundance is still dominant remains relatively high fluctuates between 58 and 77% (mean of 66%), and then shows a sharp decrease in the late stage. Smectite (mean of 14%) shows a slight increasing trend. Chlorite (mean of 14%) and kaolinite (mean of 5%) are present in similar fluctuations and low values. Moreover, they also show a concomitant increase with illite sharply decrease within this interval. As with illite, the IC value also remains stable and high (mean of 0.45°D2h) and decreases in the late stage. Unit CZ3 (72.03–54.49 m, ca. 68.9–55.4 ka) is characterized by upward increase of illite (mean of 67%), chlorite (mean of 13%) and kaolinite (mean of 6%), and corresponds to the decline of smectite (mean of 15%). There is an opposite behavior between illite and smectite, but the relationship is not distinct (Fig. 4). The IC value remains stable and reaches the highest value (mean of 0.55°D2h) compared with other clay units. Unit CZ4 (54.49–46.51 m, ca. 55.4–49.7 ka) is defined by a significant decrease of illite (from 72% to 55%, mean value of 64%) with a moderately increasing trend of smectite (mean of 14%) and chlorite (mean of 16%). The kaolinite abundance is highly variable from one sample to another in this short interval (4% < kaolinite < 8%), but the five-point running average remains relatively high (mean of 6%). Like illite, the IC value also shows a significant upward decreasing trend. The next unit, CZ5 (46.51–31.24 m, ca. 49.7–32.8 ka), is characterized by significant large fluctuations of all clay species and parameters, close to the lithological limit between clayey and halite silt sediments. On the five-point-averaged curves, illite and chlorite show an upward increasing trend, whereas smectite and kaolinite are the opposite. We note the low mean values of IC (mean of 0.41°D2h) within this interval. In unit CZ6 (31.24–18.23 m, ca. 32.8–22.3 ka), kaolinite and chlorite increase slightly, whereas illite decreases within a short period at 27.28–22.56 m (ca. 28.7–25.4 ka) in the late part of CZ6, after which it rapidly shifts to a rising trend. The smectite shows a perfectly opposite behavior to that of illite. The IC

Twelve AMS 14C ages of total organic carbon (TOC) and eight Th ages of halite in the upper 54.5 m of the sediment core of ISL1A were previously reported by Fan et al. (2014b). Three isochron 230Th ages of lake carbonates from 64.5 to 98.9 m in ISL1A were measured using the isotopic ratios of U and Th among three fractions of leachates (L), residuals (R) and whole-sample (W). We employed this simple model to correct the initial 230Th values of carbonate samples and calculated their ages by utilizing the ISOPLOT program (Ludwig, 1991). The details of 230Th dating results are presented in Supplementary Table 1. In this study, the age-depth framework of core ISL1A was established based on the selected ages in stratigraphic order as Wei et al. (2015) reported (Fig. 2). By comparing 230Th ages of halite with AMS 14C ages of TOC in ISL1A, we found that the upper four AMS 14 C ages from 4.65 to 30.29 m are in good agreement with the upper six 230Th ages from 0.35 to 32.29 m in this core. As presented in Fig. 2, the AMS 14C age (8661 ± 45 cal a BP) of TOC at a depth of 4.65 m is close to 230Th ages (8.2 ± 0.3 ka) of halite at a depth of 0.35 m. Thus, we chose the 230Th ages (8.2 ± 0.3 ka) of halite at a depth of 0.35 m to establish the age of the top of the ISL1A core. Similarly, as the AMS 14C age 19,087 ± 65 cal a BP at the depth of 13.01 m were in agreement with 230Th age 19.4 ± 0.8 ka of halite at a depth of 12.01 m, we chose the AMS 14C age for the ISL1A chronology framework. However, the AMS 14C age of TOC from 22.18 m and 230Th age of halite from 37.21 m in this core was obviously older than that from other ages between 20.22 and 46.00 m. Thus, we rejected those two ages. AMS 14C ages at 30.29–54.50 m were constant or younger with increasing depth. Therefore, we argue that AMS 14C ages of TOC and 230Th ages of halite in the upper 30 m of lacustrine sediments are more reliable, while other AMS14C ages of TOC between 30.29 and 54.50 m of lacustrine sediments are underestimated due, in part, to the contamination by underground water (Fan et al., 2014b). Due to the lowest U content in L fraction of a sample and the obvious deviation from other samples in dating results, we selected two isochron 230Th ages of lake carbonates at 64.54 and 98.90 m to establish the age model. In general, the age model in ISL1A is more reasonable. Based on the age model, the sedimentation rate of core ISL1A from the top to the bottom were 0.87, 2.27, 1.29, 1.45, 0.77, 2.04, 0.84, 1.41 and 1.16 m/1000 yrs respectively (Fig 2). 230

4.2. Lithology We describe the core sediment starting from the sediment surface using a centimeter scale. It mainly consists of six lithological units which present a gradually changed transitional relation (Fig. 2). The top 17.73 m (unit 1, 22 ka) of the core sediments are mainly composed of halite and halite with silt. The unit 2 sediments between 17.73 and 21.73 m (22–25.3 ka) corresponds to interbedded halite with a silt layer and silt. Unit 3 (21.73–31.57 m, 25.3–32.5 ka) is marked by laminated gray silt—clay with two intercalated halite silt layers. The sediments consist of silt-rich halite and interbedded layers of halite silt in unit 4 (31.57–47.36 m, 32.5–50.2 ka). Between 47.36 and 74.29 (unit 5, 50.2–71.6 ka), the sediments consist of laminated gray/green silty clay, and a halite-gypsum silt layer is embedded in this section between 49.93 and 51.59 m (52.1–53.2 ka). The lowest lithological section between 74.29 and 102.03 m (71.6–94.6 ka, unit 6) is composed primarily of dark organic-rich clay, containing several thin grayish green silt and dark clayey silt layers.

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Fig. 4. Clay mineral percentage diagram for core ISL1A, at Qarhan Salt Lake, Qaidam Basin, NE TP. The bold curves correspond to the five-point running average. The light blue bars show the relatively humid period in the Qarhan Salt Lake region during late Pleistocene. The boundary of clastic sedimentary stage and the salt-forming stage of core ISL1A are also shown in the figure.

Table 1 Mean and standard deviation of clay mineral abundance and clay parameters for the defined clay units in Core ISL1A. Core ISL1A

Relative clay mineral abundance (%) Smectite

⁄

Kaolinite

Illite

Clay mineral parameters Illite crystallinity (°D2h)

Chlorite

Unit

Depth(m)

Age (ka)

Mean

STD

Mean

STD

Mean

STD

Mean

STD

Mean

STD

CZ 1

102.03
89.23

54.49
46.51

CZ 5

46.51
31.24

CZ 6

31.24
18.23

CZ 7

18.23
1.36

0.41 0.29
0.51 0.45 0.31
0.73 0.55 0.35
0.84 0.43 0.32
0.54 0.41 0.30
0.60 0.41 0.33
0.51 0.38 0.27
0.55

0.06

CZ 4

19 14
26 14 8
24 13 8
18 16 11
22 18 11
26 18 11
25 14 4
27

3.0

72.03
54.49

61 47
67 66 58
77 67 57
74 64 55
72 61 51
73 65 54
71 41 15
74

5.2

CZ 3

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

0.8

89.23
72.03

14 8
29 14 9
23 15 8
25 14 9
21 16 11
24 11 3
23 41 3
77

5.3

CZ 2

94.7
83.7 Min
max 83.7
68.9 Min
max 68.9
55.4 Min
max 55.4
49.7 Min
max 49.7
32.8 Min
max 32.8
22.3 Min
max 22.3
9.4 Min
max

3.5 3.3 4.3 4.5 5.1 31.9

0.8 0.9 1.0 1.6 1.2 1.9

4.6 3.8 5.3 7.5 5.7 23.5

4.0 2.4 3.5 4.1 3.4 8.7

0.09 0.12 0.08 0.10 0.06 0.08

STD: standard deviation; Smectite⁄ including mixed-layers.

still maintains a low value (mean of 0.41°D2h) with a weak fluctuation. The last unit, CZ7, (18.23–1.36 m, ca. 22.3–9.4 ka) is defined by dramatic changes in the abundance of clay minerals. The highest amount of neoformed smectite (67–77% with mean value 73%) is observed at 4.12–4.45 m (ca. 12.9–12.6 ka), thus leading to a sharp decline in the abundance of other clay species. IC reaches

its lowest value of ca. 9.4 ka, with an overall low mean value of 0.38°D2h. 4.4. Geochemical data The fine fraction (<63 lm) carbonate-free lacustrine sediments from ISL1A core of QSL are characterized by high contents of SiO2,

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Table 2 Major element composition (%) of bulk (<63 lm) sediments in core ISL1A. Depth (m)

Age (ka)

Al2O3

CaO

Fe2O3

K2O

MgO

MnO

Na2O

P2O5

TiO2

SiO2

LOI

CIA

2.83 4.27 7.87 18.23 30.24 45.02 55.04 65.59 72.48 77.51 81.56 90.03 93.54 96.51 101.48

11.1 12.7 16.0 22.3 32.3 48.1 55.7 63.4 69.3 73.6 77.1 84.4 87.4 90.0 94.2

16.31 14.46 14.42 17.62 16.62 14.52 17.22 18.21 16.08 16.75 16.79 15.26 15.28 15.88 13.16

0.83 0.71 1.21 0.56 0.79 1.34 0.73 0.60 1.04 0.70 0.68 0.84 0.86 0.83 1.28

5.68 4.34 4.67 6.28 5.64 4.95 5.57 5.62 5.34 5.31 5.08 5.35 5.13 5.33 4.12

3.36 2.59 2.92 3.75 3.60 3.04 3.72 3.91 3.40 3.59 3.62 3.19 3.16 3.38 2.71

2.61 1.77 2.04 2.60 2.30 2.52 2.40 2.81 2.65 2.45 2.53 2.57 2.79 2.69 2.12

0.05 0.04 0.05 0.05 0.05 0.05 0.05 0.05 0.05 0.04 0.05 0.05 0.05 0.05 0.05

2.13 1.59 2.35 1.96 1.94 2.20 1.99 1.70 1.94 1.83 2.01 1.99 2.06 2.07 2.43

0.16 0.18 0.07 0.11 0.13 0.17 0.16 0.19 0.24 0.40 0.11 0.13 0.13 0.12 0.15

0.81 0.79 0.83 0.81 0.77 0.76 0.70 0.82 0.76 0.81 0.78 0.75 0.79 0.76 0.64

61.50 68.51 65.90 60.59 63.13 65.14 62.28 59.02 62.60 61.85 66.43 65.72 68.45 65.62 71.45

6.56 5.03 5.55 5.68 5.04 5.32 5.19 7.08 5.91 6.27 1.93 4.16 1.31 3.28 1.88

65.3 68.3 60.9 68.0 66.1 60.8 66.6 69.1 64.7 67.2 66.5 64.9 64.6 64.9 58.6

15.17

4.19

4.49

3.39

2.20

0.07

3.89

0.20

0.50

65.89

–

46.1

UCC

LOI = loss of ignition. The samples were decarbonated before the chemical analysis. UCC data from Taylor and McLennan (1985). SiO2% = 100%
(All other major elements% + LOI), see text of analytical methods for more information.

Fig. 5. Variation diagrams of major elements of different clay zones from core ISL1A. Dashed lines indicate correlations with a coefficient r2 for core sediments. Black arrow indicate the anomalous data during 12.6–12.9 ka. See Table 2 for major element compositions.

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Al2O3, and Fe2O3, and low concentrations of CaO, K2O, MgO, MnO, Na2O, P2O5 and TiO2 (Table 2). Diagrams of these elements versus Al2O3 indicate relatively strong negative correlations of Al2O3 with SiO2, CaO and Na2O and positive correlations with K2O, Fe2O3 and MgO (Fig. 5). The former implies that along with an increase in Al2O3—i.e., the formation of clay minerals (Al-rich mostly)—quartz particles (mostly Si-rich) decrease gradually, and contents of mobile alkaline elements (the most leachable elements Na and Ca) decrease accordingly. Enrichment of Fe2O3 during the formation of clay minerals suggests that along with the strengthening of chemical weathering, Fe is easily enriched in weathering products. There is no significant correlation for other elements in fine fraction element compositions—e.g., MnO, TiO2 and P2O5 (Fig. 5). The enrichment of K2O and MgO contradicts the mobility of the leachable alkaline element but is consistent with the majority of illite (K-rich) and is especially important in maintaining consistency with the low values of illite chemistry index (Mg-Fe-rich illite) of these samples in ISL1A core (Table 1). Furthermore, mineralogical analysis shows that the main components of the core lacustrine sediments are quartz, muscovite (K-rich), chlorite (Mg-Fe-rich) and feldspar (An, 2012). This is also consistent with the enrichment of K and Mg in weathering products. The comparison of major element results among the different clay zone units show that the samples of CZ2, CZ3 and CZ6 contain higher Al2O3, K2O, Fe2O3, MgO and P2O5 and lower SiO2, CaO and Na2O than those of other clay zone units (Fig. 5). Moreover, the sample point at 4.27 m (ca. 12.7 ka) always show abnormal dispersion in the relationship diagram of the major elements (Fig. 5), which means the significant difference of chemical composition between these sediments.

5. Discussion 5.1. Sediment sources The predominance of illite throughout core ISL1A suggests lithology- and climate-controlled clay mineral assemblages. Illite and chlorite could be derived from physical erosion of metamorphic and granitic parent rocks (Liu et al., 2004). The field observations and petrographic features indicate that the granites, gneisses and schists which widely exposed in surrounding mountains of the QSL region supplied primarily source materials to the lake by runoff and other agents (Fig. 1C). Minerals identification of the core sediments suggest a mineral components of quartz, feldspar, muscovite and Mg-Fe-rich chlorite (Yang et al., 1993; An, 2012), while the feldspar and muscovite are easy to form illite by exfoliation under moderate hydrolysis conditions (Liu et al., 2004; Kuwahara et al., 2010). Variations in illite and illite crystallinity of core ISL1A reflect the strengthening of physical erosion/chemical weathering during CZ1–CZ4 and CZ6 (94–51.2 ka and 32.5–24 ka), and the weakening of erosion/chemical weathering during CZ5 and CZ7 (51.2–32.5 ka and 24–9.3 ka) in the source of QSL region. Such pattern clearly differs from the weathering patterns of monsoon-controlled regions. Generally, illite and chlorite are regarded as primary minerals that reflect a decrease in hydrolysis as part of continental weathering and an increase in direct rock erosion under cold, arid climatic conditions (Liu et al., 2004). In the Qaidam Basin, the significant elevation differences between basin and surrounding mountains have greatly fostered the physical erosion rates and prevented further chemical weathering to form secondary clay minerals as the illite and chlorite soil were continuously removed from rock surfaces by rainfall or glacier meltwater, especially in humid climate periods. Thus, this process caused the unique concordant variations between chemical weathering and physical erosion in the QSL region. Indeed, such

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weathering process was also confirmed in Taiwan area with extremely high physical erosion rates and a moderate chemical weathering under orogenesis effect and subtropical East Asian monsoon climate conditions (Selvaraj and Chen, 2006). The differential resistance to chemical weathering between illite and chlorite (Bain, 1977; Zhao et al., 2003) and the limited distribution of schist may have led to the significant differences in the variation trends shown in Fig. 4. Therefore, the abundances of illite and chlorite in core ISL1A were mainly controlled by factors related to the lithology, tectonic setting and climate. Kaolinite is readily found in soils of warm, humid regions with good drainage conditions (Chamley, 1989; Robert, 2004). Parent rocks enriched in alkali and alkaline elements (e.g., granites) and parent clay minerals (e.g., chlorite) are rapidly weathered to form kaolinite (Liu et al., 2012). In core ISL1A sediments, kaolinite, which is the least abundant clay mineral, displays a clear positive correlation with the chlorite (Fig. 4), suggesting that most of the kaolinite could be derived from active erosion of inherited clays. In addition, the widely distribution of the fracturedbedrock aquifer, phreatic aquifer and overlapped artesian aquifer in the piedmont plain of the surrounding mountains around QSL region could also provide kaolinite to the lake by continuous water–rock interactions. Therefore, the kaolinite contents of core ISL1A could not reflect the contemporary climates directly due in part to the very low content and the uncertain contribution of climate-derived kaolinite. The inverse relationship between the amounts of smectite (including mixed layers) and illite in core ISL1A could simply reflect a mixture of illite-rich and smectite-rich sources (Fig. 4). However, such a source-control hypothesis is not supported by any lithologic variations in the sediments (Fagel and Boës, 2008). Even the mineralogic data remain relative, such a trend rather favors a genetic link between these two clays (Fagel and Boës, 2008). The smectite, therefore, should be regarded as a secondary clay mineral that was derived from pedogenesis of parental illite in the QSL region. Although its origin is one of chemical weathering, the smectite does not reflect contemporary climate-controlled chemical weathering due to the greater resistance of illite to hydrolysis (Zhao et al., 2003). Additionally, an unusually high smectite content (mean of 73%) with a relatively broad half-height width at the 17 Å peak (Fig. 3A) was observed in the core interval of 4.12–4.45 m (ca. 12.6–12.9 ka). Fagel and Mackay (2008) argued that such neoformed smectite with abnormally high amounts may be due to freeze
thaw weathering under periglacial conditions. And the rapidity of volcanic rock weathering often induces a high abundance of smectite that forms easily on rhyolitic materials rather than as Al species (Chamley, 1989). Fairly-poor-crystallized smectite forms abundantly on volcanic rocks indicate stronger chemical weathering conditions (Liu et al., 2009), which is consistent with the clay species and CIA values of the core interval of 4.12–4.45 m (Fig. 6). Therefore, the neoformed smectite in core ISL1A could have formed under such periglacial conditions as local volcanic rocks were eroded retrogressively by glacier melt water and/or precipitation during a short warm period (Liu et al., 2015) (Fig. 1C), which also suggested a stronger in-situ chemical weathering process. The clay mineral (illite, chlorite, smectite, kaolinite) contents and assemblages of the whole core are similar, implying similar sediment sources in QSL region since 94 ka. While the well developed alluvial
pluvial plains in the piedmont of Eastern Kunlun Mountains and the configuration of the river systems suggesting that the bedrocks of Eastern Kunlun Mountains may be a primary source of QSL region (Fig. 1C). In addition, although eolian processes are an important component of terrigenous input to QSL region as shown by the extreme variations in the clays and

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Fig. 6. Comparison of clay mineral records with other proxy records from core ISL1A. (A, B, C, D) Clay mineral records of core ISL1A. (E) Mean grain size of core ISL1A (Wei, 2011). (F) The chemical index of alteration (CIA) of core ISL1A, calculated based on the measurement of major elements (see Table 2). (G) Ti content of clastic sediments of core ISL1A measured by XRF. (H) d18O of lacustrine carbonates of core ISL1A sediments (Fan et al., 2014a). (I, J, K, L) Pollen records of core ISL1A (Wei et al., 2015). The light blue bars show the relatively humid period in the Qarhan Salt Lake region during late Pleistocene. The bold curves of A—D, G, H correspond to the five-point running average. The boundary of clastic sedimentary stage and the salt-forming stage of core ISL1A are also shown in the figure.

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other proxies (An et al., 2012) (Fig. 6), the good correspondence between the Ti content and CIA values, as well as the d18O values of the authigenic carbonates of the same core can obviously confirm the predominance of fluvial input, particularly during wet climate periods (Fan et al., 2014a) (Fig. 6F–H). This conclusion was also supported by the mean grain size and pollen analysis of the core sediments (Wei, 2011; Wei et al., 2015). 5.2. Chemical weathering and climate evolution The sediments of core ISL1A present a record of the primary parent silicate rocks in the area (e.g., granites, gneisses and schists), the clay minerals and major element compositions can therefore be used to reconstruct the weathering/erosion conditions of the Qaidam Basin during the past 94 ka. In general, multiple sources and transport processes, as well as the dilution of individual clay minerals and elements, make it difficult to assign variations in the downcore record of a single component to the changes in a paleoclimatic factor (Wehausen and Brumsack, 2002; Liu et al., 2004). However, comparisons of the ratios of two components offers the advantage of eliminating dilution effects by other components (Liu et al., 2004). We adopted the ratios of kaolinite/chlorite and illite/chlorite as mineralogic indicators based on the differential resistance of these clay species to chemical weathering. In addition, being effective indicators of terrigenous input (Murray and Leinen, 1996; Schroeder et al., 1997) and the degree of chemical weathering (Nesbitt and Young, 1982; Mclennan, 1993), the titanium content and CIA values, which is defined as Al2O3/(Al2O3 + CaO⁄ + Na2O + K2O)  100 (molar concentrations, with CaO⁄ being the CaO content in the silicate fraction of the sample) (Table 2), were also used in the comparative analysis. 5.2.1. Period of 94–54 ka (late MIS 5, MIS 4 and early MIS 3) The illite abundance and clay mineralogical indicators exhibit significant changes during late MIS 5 (94–75 ka), specifically, lower values during MIS 5b (94–83 ka) and higher values during MIS 5a (83 to 75 ka) (Fig. 6B–D). The CIA and Ti values also show the same trend during this period (Fig. 6F and G). This pattern indicates weaker chemical weathering and physical erosion during MIS 5b and subsequent strengthening during MIS 5a under the climate changes from drier to wetter. The mean grain size and d18O value of the core sediment, which represent the strength of eolian input and evaporation, respectively (An et al., 2012; Fan et al., 2014a), also suggest the change from strong to weak eolian input and evaporation with the change from drier to wetter climate conditions in the QSL region (Fig. 6E and H). Pediastrum, a sensitive indicator of low-salinity water and a warm, wet climate (Wei et al., 2015), is also more abundant in the MIS 5a portion than the MIS 5b portion of the core (Fig. 6I). However, there still existed a relatively moist climate with relatively stronger chemical weathering/erosion condition during MIS 5b, according to the higher values of CIA and Ti % during this period than other drier periods in the Qaidam Basin. In addition, an organic-rich silty clay, a widespread distribution of arboreal and broad-leaved pollen and a relatively high A/C ratio support this conclusion (Figs. 4 and 6J–L) (Wei et al., 2015), which is also consistent with the high level records of Qinghai Lake (northeastern Qaidam Basin, Fig. 1B) during the period of 110–75 ka (Madsen et al., 2008; Rhode et al., 2010). Early MIS 4 (75–72.5 ka) continues the varying trend of MIS 5a (83–75 ka), and then turned to low values in the illite and clay indicators at 72.5–68.8 ka, which match well with the CIA and Ti % curves (Fig. 6B–D, F and G). Moreover, a slight increase could also be observed in the d18O curve of the authigenic carbonates (Fig. 6H), suggesting a decrease in weathering/erosion intensity and drier climate conditions during the period of 72.5–68.8 ka.

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Afterward, the relevant indicators all rapidly rebound, indicating the strengthening of weathering/erosion with the climate conditions turning wetter during late MIS 4 (68.8–57.5 ka). These results differ little from the pollen records: a high percentage of Pediastrum, arboreal and broad-leaved pollen with a high value of A/C ratio during early MIS 4 suggest cool, wet climatic conditions and high lake levels in the central eastern Qaidam Basin, and a reversal during late MIS 4 (Wei et al., 2015) (Fig. 6I–L). Studies in the Gahai Lake region (northeastern QSL in the Qaidam Basin) suggest that two periods of higher lake levels occurred at 85
72 and 63
55 ka, consistent with our clay mineral findings (Fan et al., 2010, 2012). However, the falling trend, the illite abundance and the clay indicator ratios remain high during early MIS 3 (57.5–54 ka), correspond to the CIA and Ti % curves (Fig. 6) and suggest relatively stronger chemical weathering/erosion processes under wetter climate conditions in the Qaidam Basin. The d18O values of the authigenic carbonates remain low with a slight upward trend (Fig. 6H), indicating a trend of slightly enhanced evaporation and reduced effective humidity in the Qarhan region during MIS 3c, which correlates well with the lithologic change from silty clay to halite with clayey silt (Fig. 4). Pediastrum was also continuously present during this time, implying lower salinity and higher lake levels (Wei et al., 2015). This conclusion is also consistent with geomorphic and chronometric evidences obtained at Gahai Lake, where the high lake levels were recorded at 65–55 ka (Fan et al., 2010, 2012). 5.2.2. Period of 54–24 ka (middle and late MIS 3) The middle of MIS 3 (54–32 ka) is marked by low illite abundance and clay indicator ratios and extreme variations in their curves, which match well with the variations in the CIA and Ti % in the core sediments and also coincides with a salt-bearing deposit and the gradually disappearance of Pediastrum in the core (Figs. 4 and 6I). These results suggest relatively dry climate conditions with the lowstand of the lake and weakened weathering/ erosion processes in the QSL during 54–32 ka. The d18O value is abnormally high in the core sediments of this period, which can be attributed to strong evaporation under the dry climate conditions (Fan et al., 2014b) and is consistent with the strong eolian input indicated by the mean grain size curve (Fig. 6E and H). Clearly, the widespread development of the alluvial
pluvial plains further affected the clay mineral assemblages and geochemical compositions of the core sediments by way of high winds during this period. The illite abundance and clay indicator ratios all rose again during late MIS 3 (32–24 ka), which matches well with the CIA and Ti % in the core sediments (Fig. 6B–D, F and G). The deposition of an organic-rich silty clay also corresponded with the lower d18O values and higher A/C ratios, as well as the increased abundance of arboreal and broad-leaved pollen (Figs. 4 and 6H–L). These results indicate that the chemical weathering strengthened again during this period as the climate became warmer and wetter. This conclusion is consistent with the geomorphic and chronometric evidence obtained at Toson Lake (northeastern QSL), which was at a high lake level at 31 ka (Fan et al., 2012). However, the effective humidity during this period was still lower than that during the period of 94–54 ka in the Qarhan region based on the claymineral and pollen results (Wei et al., 2015). In addition, a decline of illite abundance and clay mineralogical ratios was observed during the short period of 28.7–25.4 ka, which corresponding to the changes of Ti %, d18O values and mean grain size curve and suggesting a drier climate conditions and weaker weathering/erosion processes in the Qaidam Basin. The presence of thin layer of halite silt in core ISL1A sediments during this period is also good to support this conclusion (Fig. 4).

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5.2.3. Period of 24–9.3 ka The period of 24–9.3 ka cannot be described in detail because of the low resolution of this stage. However, the extreme variations in the illite and clay indicator curves match well with the low CIA and Ti % values (Fig. 6B–D, F and G) and the higher mean grain size and d18O values (Fig. 6E and H), which were likely a signal of a cold, dry climate with high winds and weak chemical weathering during this stage. These findings also correspond with the significant presence of halite in the core sediments during this period. Nevertheless, a short period of 12.6–12.9 ka exhibiting higher CIA values and abundantly neoformed smectite was noted in the core, which corresponding to the lithology of dark organic-rich clay. The low Ti % value in these core sediments suggest a fractionation between Ti-rich heavy minerals and clastic sediments, which could be caused by glacial meltwater under weak hydrodynamic conditions. We tentatively conclude that these features represent stronger in-situ chemical weathering under periglacial conditions during this short warm period (Fagel and Mackay, 2008; Liu et al., 2015). Simultaneously, a lower input of terrigenous siliciclastic sediments was the result of a drying up of the streams; consequently, QSL evolved into a dry playa by the end of MIS 2. 5.3. Trends in chemical weathering and intensity For more-detailed investigation of chemical weathering in the QSL region since the late Pleistocene, geochemical data from core ISL1A were used to quantitatively evaluate the chemical weathering intensity based on the mobility of various elements during chemical weathering (Nesbitt et al., 1980; Singh et al., 2005). The CIA reflects the proportion of Al2O3 versus labile oxides and thereby indicates the extent of silicate weathering, e.g., weathering of plagioclase to clay (Nesbitt and Young, 1982; Fedo et al., 1995). Consequently, the CIA values of approximately 45
55 indicate virtually no weathering, and the average UCC has a CIA value of approximately 47; a value of 100 indicates intense weathering with the complete removal of alkali and alkaline earth elements from the parent rocks (McLennan, 1993). In addition to the CIA, a A–CN–K (Al2O3–(CaO⁄ + Na2O)–K2O) diagram was developed to indicate the trends and degree of silicate weathering and to evaluate the parent rock composition (Nesbitt and Young, 1989; Fedo et al., 1995). On the A–CN–K diagram, the primary trend of silicate weathering in the Qarhan region indicates preferential leaching of CaO and Na2O and then K2O and enrichment in Al2O3 (Fig. 7). Plagioclase is more susceptible to weathering than potassium feldspar, and thus, Ca and Na are leached preferentially over K (Nesbitt and Young, 1982). Overall, the weathering trend of the samples parallels the A–CN line, which implies the removal of the Ca- and Na-bearing silicate minerals from the parent rocks, whereas the K-bearing minerals were attacked less. The weathering trend toward the illite reference-segment in Fig. 7 partly reflects the predominance of illite in the clay mineral assemblages in the core. The distribution of the data points appear to indicate weathering of similar parent rocks of felsic igneous origin (e.g., granite, granodiorite). The deviation of data point of a depth of 4.27 m (12.7 ka) from the main weathering trends (Fig. 7), as well as the chemical composition (Fig. 5), suggests the difference between sample at 4.27 m and others in source rocks, which is consistent with our clay mineral results. According to the classification of weathering intensity developed by Nesbitt and Young (1989), the samples from core ISL1A exhibit degrees of weathering ranging from weak to intermediate (Fig. 7). Samples of CZ5, CZ7 and part of CZ1, which correspond to MIS 2, the middle of MIS 3 and MIS 5b, respectively, exhibit weaker silicate weathering, whereas samples CZ2, CZ3 and CZ6, which correspond to late MIS 3, MIS 4 and MIS 5a, respectively,

Fig. 7. Weathering trends from the A–CN–K ternary diagram for sediment averages of core ISL1A and comparisons of the chemical index of alteration (CIA). A = Al2O3; C = CaO; N = Na2O; K = K2O. UCC (Taylor and Mclennan, 1985) was also plotted as a reference. Note that all samples indicate weathering trends in parallel with the A–CN line.

exhibit stronger weathering/erosion process. These findings are also consistent with the clay mineral results. 6. Conclusions The clay mineral from <2 lm fraction and major-element concentrations in a 102-m-long sediment core (ISL1A) were analyzed to reconstruct the relationship between weathering process and paleoclimatic change on the northern TP. The results show that: (1) clay mineral from <2 lm fraction in ISL1A is composed of illite (47–77%), chlorite (8–27%), smectite (including illite-smectite mixed layers, 3–29%) and kaolinite (2–11%). The variation in abundance of clays and ratios between clay minerals (illite/chlorite, kaolinite/chlorite) can be used as good indicators to build up paleoclimatic changes; (2) higher values of illite, kaolinite/chlorite and illite/chlorite ratio, and lower values of smectite, chlorite and kaolinite minerals occurred in 83–72.5 ka, 68.8–54 ka, 32–24 ka, corresponding to stronger chemical weathering/physical erosion processes, and warm and wet periods, respectively, which are coincided with higher effective moisture phases documented by oxygen isotopes of authigenic carbonates and pollen records in ISL1A; (3) combining with chemical weathering index and (Al2O3-(CaO + Na2O)-K2O) diagram, chemical weathering degree in this study area takes a varying process from low to intermediate on the whole. Acknowledgments Thanks would go to Dr. X.J. Liu, Dr. H.D. Chen and Dr. F.Y. An for their helpful discussions on the contents and figures, and Y.S. Zhang for laboratory work. We specially thank Prof. M.F. Zhou and two anonymous reviewers for their helpful comments and suggestions to improve the manuscript. This work was financially supported by the National Natural Science Foundation of China (41303029, 41272274 and U1407107). Appendix A. Supplementary material Supplementary data associated with this article can be found, in the online version, at http://dx.doi.org/10.1016/j.jseaes.2016.06. 013.

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