Sedimentology and magnetostratigraphy of the Tierekesazi Cenozoic section in the foreland region of south West Tian Shan in Western China

Tectonophysics 654 (2015) 156–172

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Sedimentology and magnetostratigraphy of the Tierekesazi Cenozoic section in the foreland region of south West Tian Shan in Western China Xinwei Chen, Hanlin Chen, Xiaogan Cheng ⁎, Zhongyue Shen, Xiubin Lin Department of Earth Sciences, Zhejiang University, Hangzhou 310027, China Research Center for Structures in Oil- and Gas-Bearing Basins, Ministry of Education, Hangzhou 310027, China

a r t i c l e

i n f o

Article history: Received 30 October 2014 Received in revised form 15 April 2015 Accepted 20 May 2015 Available online 30 May 2015 Keywords: Magnetostratigraghy Tian Shan Tibetan Plateau Tarim Basin Talas–Fergana fault

a b s t r a c t The geology of Tian Shan provides an excellent example for understanding the intracontinental orogeny in the context of Indian–Eurasian convergence. Previous studies leave much space in basinfill deposition process to be assessed in the regions west to the Talas–Fergana fault (TFF). We improve the understanding by conducting new investigations on sedimentology and magnetostratigraphy in the Tierekesazi section of the foreland region of south West Tian Shan. Four lithofacies have been identified, (i) marine lithofacies from the Aertashi to Bashibulake Formations, (ii) lacustrine to fluvial (plain) lithofacies from the Keziluoyi to the middle Pakabulake Formations, (iii) alluvial sand–gravel sheet lithofacies in the upper Pakabulake Formation, and (iv) conglomerate lithofacies from the Atushi to Xiyu Formations. Magnetostratigraphic analysis, accompanied with biostratigraphic correlation, indicates that four lithofacies cover age intervals of ca. 65 Ma to 34 Ma, ca. 22.1 Ma to 12 Ma, 12 Ma to 5.2 Ma, and 5.2 Ma to approximately present (?), with the sediment accumulation rates increasing from ca. 2.4/3.3–3.5 (compacted/decompacted) cm/ka in the lithofacies (i), to 12.3/16–17 cm/ka in the lithofacies (ii), to 16.3/19.5–20.6 cm/ka in the lithofacies (iii), and finally to N 22.8/N22.8 cm/ka in the lithofacies (iv). These results suggest three episodes of sedimentary events. Combined with previous results, these episodes of sedimentary events are attributed to tectonic activities that are widespread along south Tian Shan. We speculate that the Oligo-Miocene boundary event more directly and likely marks the initial underthrusting of the Tarim block beneath the south Tian Shan. The mid-Miocene and MioPliocene boundary events, although approximately synchronous between the regions east and west to the TFF, have different structural expressions in the two regions. Such difference is proposed to cause the dextral slipping of the TFF, and more fundamentally, likely be driven by the northward indentation of the Pamir at this time. © 2015 Elsevier B.V. All rights reserved.

1. Introduction The Indian–Eurasian convergence significantly reorganized the tectonic framework and topography in Asia, resulting in the uplift of the Tibet Plateau and the Himalaya Mountains, and built up the mountain ranges in Central Asia. Among these ranges, the Tian Shan (Fig. 1), which is an approximately 2500-km-long intracontinental range with an average altitude of 2500 m and summits reaching up to 7000 m, accommodates approximately 40% of the India-Eurasian convergence (Avouac and Tapponnier, 1993; Abdrakhmatov et al., 1996; Reigber et al., 2001; Wang et al., 2001; Chen et al., 2007). The erosion of the Tian Shan results in the thick Cenozoic sediments deposited in adjacent basins with the thickness of kilometers scale in the foreland regions of Tarim basin (Sobel and Dumitru, 1997), which provide an opportunity

⁎ Corresponding author at: 38# Zheda Road, Hangzhou 310027, Zhejiang Province, China. Tel./fax: +86 571 87952791. E-mail addresses: [email protected] (H. Chen), [email protected] (X. Cheng), [email protected] (X. Lin).

http://dx.doi.org/10.1016/j.tecto.2015.05.009 0040-1951/© 2015 Elsevier B.V. All rights reserved.

to reveal the orogenic history and to understand the Cenozoic intracontinental orogeny within the context of Indian–Eurasian convergence. Despite its significance, the Cenozoic orogenic processes of south Tian Shan (sTS, a geographical terminology defined here as south portion of the Tian Shan, including the Chinese South Tian Shan (CSTS), the Kyrgyz South Tian Shan (KSTS) and south part of the West Tian Shan (WTS)) have not been completely established. Most previous works have focused on dating the Cenozoic sediments in foreland regions in front of Chinese North Tian Shan (CNTS), the Chinese portion of the KSTS and CSTS (Charreau et al., 2005, 2006; Chen et al., 2002, 2007; Heermance et al., 2007; Huang et al., 2006; Li et al., 2011; Sun et al., 2004, 2009) east to the Talas–Fergana Fault (TFF, defined here as boundary between the Tian Shan and WTS and between the Tarim and West Tarim blocks (e.g., Burov and Molnar, 1998; Sobel, 1999; Sobel et al., 2013)) (Fig. 1), leaving the study of sedimentological records in the foreland regions of south WTS west to the TFF being rarely addressed. An exception of this is the magnetostratigraphic study conducted by Yang et al. (2015), who dated the Miocene sediments but still left much space for the sediments to be more comprehensively dated.

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Fig. 1. Digital elevation model (DEM) of the Tian Shan and its adjacent areas showing tectonic setting of the studied region (modified after Lei et al., 2013; Xiao et al., 2013; Loury et al., in press), and localities of previous magnetostratigraphic studies in the foreland regions of the Tian Shan which have been cited for comparison in this study (Charreau et al., 2005, 2006; Chen et al., 2002, 2007; Heermance et al., 2007; Huang et al., 2006; Li et al., 2011; Sun et al., 2004, 2009). Series of foreland fold-and-thrust belts (FTBs) develop in the regions east to the Talas–Fergana fault (TFF), in contrast to their absence in the regions west to the TFF. MPT: Main Pamir Thrust; TFF: Talas–Fergana Fault; WTS: West Tian Shan; KNTS: Kyrgyz North Tian Shan; KCTS: Kyrgyz Central Tian Shan; KSTS: Kyrgyz South Tian Shan; CNTS: Chinese North Tian Shan; CCTS: Chinese Central Tian Shan; CSTS: Chinese South Tian Shan; FTB: fold-and-thrust belt.

We here aim to improve the understanding of Cenozoic intracontinental orogeny with new investigations of Cenozoic sedimentologic records, comprising sedimentology and magnetostratigraphy of the Neogene foreland basinfills in the Tierekesazi section of south WTS west to the TFF. This work, combined with previous results, allows for comprehensively establishing the Cenozoic orogenic process of sTS. 2. Geological setting The Tian Shan was initially established through complex accretions of island arcs and amalgamations of continental lithospheric blocks during late Paleozoic (Allen et al., 1993; Carroll et al., 1990, 1992, 1995; Charvet et al., 2007, 2011; Coleman, 1989; Gao et al., 1998; Glorie et al., 2010; Shu et al., 2003; Watson et al., 1987; Windley et al., 1990, 2007; Xiao et al., 1992, 1994), and then underwent strike–slip deformation and transpression in the Carboniferous–Permian, transtension in the late Triassic to early Jurassic (Rolland et al., 2013), and compression from late Jurassic to early Cretaceous (e.g., Allen et al., 1991; Bullen et al., 2003; Sobel, 1999). More recently, within the context of the Indian– Eurasian convergence, the Tian Shan reactivated since Oligo-Miocene boundary (Allen and Natal'in, 1995; Allen et al., 1999; Avouac and Tapponnier, 1993; Burchfiel and Royden, 1991; Burchfiel et al., 1999; Dumitru et al., 2001; Heermance et al., 2008; Jolivet et al., 2010; Lu et al., 1994; Molnar and Tapponnier, 1975; Sobel et al., 2006; Tapponnier and Molnar, 1979; Tapponnier et al., 1982; Yin et al., 1998) or earlier (Glorie et al., 2010; Jolivet et al., 2010; Wang et al., 2009; Yang et al., 2014; Zhang et al., 2014), creating terrestrial basinfills of thickness of kilometers scale in the foreland regions (e.g., Sobel and Dumitru, 1997). The Tian Shan is now mainly located in China and Kyrgyz and has been described by various terminologies. In the Kyrgyz part, the Tian Shan has been divided into Kyrgyz North Tian Shan (KNTS), Kyrgyz Middle Tian Shan (KMTS) and KSTS separated by Terksey and KSTS sutures respectively (e.g., Loury et al., in press) (Fig. 1). In the Chinese portion, the Tian Shan has been divided into CNTS, Yili block, Chinese Central Tian Shan (CCTS) and CSTS (Burtman, 2006, 2010; Charvet et al., 2007; Loury et al., in press; Wang et al., 2007, 2008; Xiao et al., 2013) (Fig. 1). In the study, we define the sTS, as noted above, to be the geographical domain including CSTS, KSTS and south portion of WTS (Fig. 1). The TFF, a dextral strike–slip fault with episodic activities since Carboniferous–Permian (e.g., Rolland et al., 2013), defines the boundary between the CSTS–KSTS and the WTS (e.g., Burov and

Molnar, 1998) (Fig. 1), with its reactivation during the Cenozoic Era initiating at about mid- to late-Miocene (e.g., Burtman et al., 1996). Our studied Tierekesazi section is located in the foreland region of south WTS west to the TFF (Fig. 1). Tectonically, it is closely bounded by the KSTS/WTS to the north and the Pamir to the south. However, interpretation of seismic profile indicates that the Tierekesazi region is structurally dominated by south-vergent sTS thrust system, implying the predominate impact from the sTS (Luo, 2005). The speculation is consistent with the detrital zircon U–Pb age results of the Cenozoic sediments in the Ulugqat tens of kilometers west to the studied Tierekesazi section, suggesting a predominant provenance from the sTS (Yang et al., 2014). Predominant north-to-south paleocurrent measured from the Cenozoic sedimentary records in the Kashi (e.g., Chen et al., 2002; Heermance et al., 2007; Scharer et al., 2004) to the east of studied region provides additional supportive evidence for the speculation. The Tierekesazi section generally appears to be a monocline dipping SSW, except for an anticline–syncline complex developing in the middle section (Fig. 2c). The section is cropped out by a seasonal stream perpendicular to the strike of strata which has exposed the complete Cenozoic succession that overlies the Cretaceous gypsum beds conformably (Figs. 2a and b). The Cenozoic sediments in this section include marine facies deposits in the lower section overlain disconformably by the terrestrial sediments, generally forming an upward coarsening succession. Regionally, these Cenozoic sediments have been divided into various stratigraphic frameworks based on specific research purpose. Initially, Sobel (1999) divided the sediments into Paleogene Kashi Group, Miocene Wuqia Group, Pliocene Atushi Formation and Quaternary Xiyu Formation. Chen et al. (2007) and Heermance et al. (2007) divided the upper sequences into Xiyu, Atushi and Pakabulake Formations downward section with emphasis on the latest Cenozoic deformation in Kashi region. Bosboom et al. (2011) and Bershaw et al. (2012), focusing on the Paleogene sequence to understand the Paratethys Sea retreat from the Tarim basin and related reorganization of atmospheric circulation, divided the Cenozoic deposits into Paleogene Aertashi, Qimugen, Kalataer, Wulagen and Bashibulake Formations and overlying Wuqia Group. Another school of stratigraphic framework includes Paleogene Kashi Group (including Aertashi, Qimugen, Gaijitage, Kalataer, Wulagen and Bashibulake Formations), Miocene Wuqia Group, Pliocene Atushi Formation and Quaternary Xiyu Formation (Wang et al., 2014). A more detailed framework has been proposed to divide the Cenozoic sediments into Paleogene Aertashi, Qimugen, Kalataer, Wulagen and Bashibulake Formations,

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Fig. 2. Simplified geological map of the studies area (a) (modified from Yang et al., 2014) and the measured Tierekesazi section (b), and the details of structural cross section established on the basis of field measurement (c). Locations of the magnetostratigraphic and pollen samples are labeled in the geological map and cross section, respectively. The anticline–syncline complex has been leapt over during the sampling.

and Miocene Keziluoyi, Anjuan and Pakabulake Formations, and Pliocene Atushi Formation and Quaternary Xiyu Formation (Liu et al., 2012; Yang et al., 2014). The last stratigraphic framework will be employed in this study in an attempt to present our results with a more detailed stratigraphic framework and an up-to-date geological work. In particular, the terminological reference of the Xiyu Formation in this study is based on the traditional division, from which the formation is identified on the basis of sediment color variation from light gray to dark gray (Burchfiel et al., 1999; Huang et al., 1947; Jia et al., 2004; Liu et al., 1996; Molnar et al., 1994; Sun et al., 2004), rather than another school of division which is based on the emergence of conglomerate (Charreau et al., 2005, 2006, 2009a,2009b; Chen et al., 2002, 2007; Heermance et al., 2007). 3. Sedimentology of the Cenozoic deposits in the Tierekesazi section The N 4294 m Cenozoic sequences have been investigated and measured in the field within the stratigraphic framework described above. The Paleogene marine facies sediments, which conformably overlie the Upper Cretaceous, are divided into five stratigraphic units, the Aertashi, Qimugen, Kalataer, Wulagen and Bashibulake Formations, upwardly (Fig. 3), with the division lithologically consistent with previous work in the region (Bosboom et al., 2014; Wang et al., 2014; Yang et al., 2014). The lowermost 255 m Aertashi Formation (thickness level of 0–255 m) is composed of thick gypsum beds interbedded with layered dolomite (Figs. 3 and 4a), representing a restricted shallow marine to lagoon setting. The Qimugen Formation (thickness level of 255–427 m) overlies the Aertashi Formation conformably (Figs. 3 and 4b), attaining a thickness of 172 m. It begins with a limestone layer, progressing upwardly into interbedded green sandstone and mudstone

packages in the lower portion and into red thick-bedded mudstone at the upper section (Fig. 4b). It is interpreted as typical shallow marine circumstances. The section passes upwardly into the Kalataer Formation (thickness level of 427–444.1 m) with a thickness of 17.1 m (Fig. 3), which is dominated by shelly limestone with abundant Ostracoda fossils present throughout the formation (Fig. 4c). These features suggest that the Kalataer Formation inherited from the shallow marine setting (Yang et al., 2014). The sedimentary circumstances then underwent a regression into a tidal-flat setting, indicated by yellowish green thickbedded mudstone intercalating brownish red thin-bedded sandstone with abundant shelly fragments in the overlying Wulagen Formation (thickness level of 444.1–488.6 m) (Fig. 3). The Bashibulake Formation (thickness level of 488.6–756.6 m), overlying the Wulagen Formation in conformity (Fig. 3), attains a thickness of 268 m. This formation consists of massive brownish red mudstone intercalating thin-bedded red siltstone and gray gypsum (Figs. 3 and 4d), implying a sedimentary setting of restricted shallow marine circumstance. The studied region then entered into a terrestrial setting in later periods recorded by the overlying Neogene sediments. The 152 m thick Keziluoyi Formation (thickness level of 756.6– 908.6 m), overlying the Bashibulake Formation in a disconformity (Fig. 3) indicated by the first presence of sandstone body (interpreted as underwater channel) (Fig. 4e) and absence of the gypsum layer, consists of lacustrine sediments, dominated by thick-bedded brownish red mudstone intercalating thin-bedded brick red sandstone in the lower formation and by interbedded thick-bedded muddy sandstone/ siltstone–sandstone packages in the upper portion (Fig. 3). The significant retreat of marine sequences and the onset of overlying terrestrial facies sediments imply that the disconformity between the Bashibulake and Keziluoyi Formations should represent a striking sedimentary

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Fig. 3. The lithologic logs of the Tierekesazi section, showing a generally upward coarsening section. The vertical axis suggests the thicknesses, and the horizontal axis represents the sediment particle size. L: limestone; M: mudstone; S: sandstone; P: pebbly sandstone; C: conglomerate.

hiatus. The Anjuan Formation (thickness level of 908.6–1423.6 m), consisting of 515 m thick sandstone–mudstone complexes (Fig. 3), gradually shifts from a lacustrine to a fluvial (plain) setting. The sandstone beds increase their thickness significantly and change from thin-bedded into lens-shaped beds upwardly, which possibly represents channel deposition (Fig. 4f). In contrast, the mudstone/siltstone layers with decreased thickness may represent overbank deposition. Sedimentary structures, presented in this formation, include ripples (Fig. 4f-1) and cross-beddings (Fig. 4f-2). The overlying Pakabulake Formation (thickness level of 1423.6–3108.6 m), attaining a thickness of 1780 m, consists of sandstone–mudstone–siltstone packages and conglomerate (Fig. 3). The sandstone, usually moderate- to thick-bedded with abundant cross-beddings (Fig. 4g-1), changes its patterns from lens-shape in the lower to sheet form in the upper formation (Fig. 4g) with particle size increasing upwardly. The mudstone/siltstone forms thick sequences in the lower portion and changes into sheets in the upper portion (Fig. 4g), with ripples, mud cracks and burrows evident (Fig. 4g-2 to g-4). It is notable that the siltstone component significantly increases upwardly. The conglomerate, dominated by well-sorted round pebbles, first appears in the central formation and becomes more frequent upward in the section (Fig. 4g-5). The conglomerate is arranged in sheet form or thin layer hosted in the sandstone beds. These features suggest that the sedimentary facies change gradually from fluvial (plain) to alluvial (fan) setting. The section then passes upwardly into the 540 m thick Atushi Formation (thickness level of 3108.6–3719.6 m). This formation forms several upward thinning sequences, with massive pebbles to cobble gravels in the lower portion (Fig. 4h-1) and sandstone–siltstone–mudstone packages (Fig. 4h) in the upper part of each sequence (Fig. 3). The gravel is sub-angular, moderately to poorly sorted, and the sandstone, siltstone and mudstone

are arranged in sheet forms. The evidence suggests that the Atushi Formation deposits in the central to the proximal alluvial fan. The Xiyu Formation (thickness level of 3719.6–ca. 4294.6 m) (Fig. 3), forming the uppermost formation of the Tierekesazi succession, is predominated by cobble-size gravel, changing from the underlying Atushi Formation with gradual increase in particle size and a sudden shift of sediment color from light gray to dark gray (Fig. 4i). The gravel in the Xiyu Formation is sub-angular to angular and poorly sorted. These facts indicate that the Xiyu Formation deposits in a setting of proximal alluvial fan. In summary, within a general upward coarsening section in particle size, three striking events are evident in the Tierekesazi section: (1) the first appearance of underwater channel sandstone at the bottom of the Keziluoyi Formation, (2) the first appearance of a pebble gravel sheet in the middle of the Pakabulake Formation, and (3) the first appearance of a massive cobble conglomerate at the bottom of the Atushi Formation. These events divide the Cenozoic succession into four lithofacies: (i) Paleogene marine lithofacies from the Aertashi to the Bashibulake Formations, (ii) lacustrine to fluvial (plain) lithofacies from the Keziluoyi to the middle Pakabulake Formations, (iii) alluvial sand–gravel sheet lithofacies in the upper Pakabulake Formation, and (iv) alluvial conglomerate lithofacies from the Atushi to the Xiyu Formations. These findings suggest three pulses of particle size and sedimentary energy increase throughout the Tierekesazi region. 4. Magnetostratigraphic sampling, laboratory processing and rock magnetic analysis To constrain the ages of the sedimentary events, we have conducted a magnetostratigraphic investigation on the Neogene succession (thickness

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Fig. 4. Representative outcrops of the described Tierekesazi section. (a) Thick gypsum beds in the Aertashi Fm., with details of the gypsum zoomed out in (a-1); (b) The Qimugen Fm. overlies conformably above the underlying Aertashi Fm.; (c) Shelly limestone in the Kalataer Fm., with details of the Ostracoda fossils zoomed out; (d) Massive brownish red mudstone intercalating thin-bedded red siltstone and gray gypsum in the Bashibulake Fm.; (e) Massive sandstone beds at the bottom of the Keziluoyi Fm., interpreted to be underwater channels which mark the disconformity between the Keziluoyi Fm. and the underlying Bashibulake Fm.; (f) Lentoid sandstone beds in the Anjuan Fm., with ripples (f-1) and cross-bedding (f-2) shown in details; (g) Sandstone sheets in the Pakabulake Fm., with cross-bedding (g-1), ripples (g-2), mud cracks (g-3) and burrows (g-4) evident, and the conglomerate appearing in the middle of the formation (g-5); (h) Sandstone-siltstone-mudstone packages and (h-1) cobble to pebble gravels in the Atushi Fm.; (i) The Xiyu Fm. overlies the Atushi Fm., with particle size increasing gradually and sudden variation of sediment color from light to dark gray.

level of 756.6–3824.8 m) in the Tierekesazi section, with beginning and ending coordinates 39°50′N, 74°35′E and 39°48′N, 74°31′E, respectively. The lower Paleogene succession (thickness level of 0–756.6 m), dominated by limestone, dolomite, gypsum and strongly weathered

mud/siltstone, is practically impossible for sampling and therefore was excluded from our magnetostratigraphic investigation. Sampling density is dependent on fresh outcrops and lithology. In particular, the syncline and anticline have been left out by tracing corresponding

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beds in the two limbs, which ensures the continuity of the sampling (Figs. 2b and c). The sample orientations and the associated strata attitudes have been measured for each sample for the paleomagnetic data correction. Totally 708 sites have been sampled in the field, with at least 3 orientated samples for each site. In total, 2027 samples have been collected in which 100 samples in the thickness interval 756.6–805.6 m (covering lower Keziluoyi Formation), 1580 samples in the thickness interval 908.6–2341 m (covering Anjuan to middle Pakabulake Formation), 86 samples in the thickness interval 2341–3108.6 m (covering upper Pakabulake Formation), 228 samples in the thickness interval 3108.6– 3719.6 m (covering Atushi Formation), and 33 samples in the thickness interval 3719.6–3824.8 m (covering lower Xiyu Formation). The outcrops of the upper Keziluoyi Formation (thickness interval 805.6– 908.6 m) are strongly weathered and absent for sampling, while the Xiyu Formation (thickness interval 3719.6 to top section) is dominated by massive cobble conglomerate in which samples are sporadically distributed in the intercalated sandstone lens. These suggest high sampling density, one sample for less than 2.7 m thickness in thickness intervals 756.6–805.6 m, 908.6–2341 m and 3108.6–3719.6 m, in contrast to lower density of 1 sample per N 3.2 m in the rest section. All the drilled specimens have been cut into standard cylinder samples with a diameter of 25 mm and a length of 22 mm for paleomagnetic analyses. Among all of these samples, 17 representative samples have been selected for isothermal remanent magnetization (IRM) analyses in Paleomagnetic lab of Zhejiang University to determine the magnetic minerals, with an applied magnetic field from 0 to 2600 mT. From these 17 analyzed samples, typical results of 6 samples have been listed in Fig. 5 and Table 1. Most of the results show similar features, including that the Mr/Ms values (ratio between the induced magnetic field intensity of the analyzed sample and applied field) reach 0.6 when the B (applied magnetic field intensity) values are at 500 mT, and the magnetic field intensities of the samples (Mr) are not saturated when B reaches 2600 mT (Fig. 5). When imposing a reverse magnetic direction, we are able to obtain the coercive force values of the samples, which are larger than 250 mT in most samples (Fig. 5). These results indicate a typical feature of hematite. The exception is Sample 329-3, in which the Mr/Ms value reaches 0.8 when the B value increases to 200 mT and Mr gets saturated when B reaches 2600 mT (Fig. 5). The coercive force value of this sample is only 82 mT (Fig. 5). These facts indicate that its magnetic mineral is mainly magnetite. Within all the 2027 samples, 792 cylindrical samples, at least one from each sampled site, were subjected to progressive demagnetization by a TD-48 thermal demagnetizer in Paleomagnetic laboratories of Nanjing University and Zhejiang University. Demagnetization was operated in temperature steps of 50–80 °C under 500 °C and 10–30 °C from 500 °C to 700 °C. 584 of the samples were measured by a 2G-superconducting

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Table 1 Coercive force values of representative samples shown in Fig. 5. Sample

Thickness level (m)

coercive force (mT)

019-1 161-3 329-3 419-3 595-3 657-1

798.4 1240.9 1725.2 1953.3 3262.4 3719.6

449 351 82 475 294 234

magnetometer in the lab of Nanjing University; the others were measured using a JR-6A rotating magnetometer in the lab of Zhejiang University. Stepwise demagnetization results were evaluated on stereographic projections and Zijderveld diagrams (Zijderveld, 1967), and components of characteristic remanent magnetization (ChRM) were calculated by principle component analysis proposed by Kirschvink (1980). These components have been corrected with sample and stratum orientation data, respectively. Within the demagnetization of the 792 samples, most of the samples present simple demagnetization behavior with stable components at temperatures lower than 680 °C, which reveals two clear components (in both geographic and strata coordinate systems): a high-temperature component (HTC) that decays toward origin and a low-temperature component (LTC) that does not. The order of natural remanent magnetization (NRM) is between 10−2 and 10−4 A/m. LTCs are isolated at temperature of 220–300 °C. HTCs are obtained at approximately 600–700 °C and can yield opposite paleomagnetic polarities (Fig. 6). Three typical samples are presented in Fig. 6, showing representative demagnetization behavior and LTC and HTC results. In Sample 015-1 (Fig. 6a), remanence reduces slightly in low temperature, yielding an LTC with D = 337.1°, I = 53.5°, α95 = 3.5 (D is declination, I is inclination and α95 is the radius that the mean direction lies, within 95% confidence, in geographic coordination), which is similar to the present geomagnetic field in the site; decreases obviously at 580 °C and is unblocked at 680 °C, returning an HTC with the normal polarity direction of D = 11.9°, I = 54.1°, α95 = 4.3 (in strata coordination). These features are consistent with the previous determination of magnetic minerals dominated by hematite and subordinate magnetite. In contrast, Sample 319-2 (Fig. 6b) returns an HTC (620 °C to 700 °C) of reversal component direction with D = 149°, I = −52.8°, α95 = 1.6 (in strata coordination). Some specimens show uniform LTC and HTC directions. Sample 155-1 provides an example for this case, yielding consistent LTC and HTC directions of D = 0.7°, I = 60.8°, α95 = 3.4 (LTC and HTC directions calculated in geographic and strata coordination systems, respectively) (Fig. 6c). The rest samples return complex HTCs. Taking Sample 490-3 for example (Fig. 6d), it

Fig. 5. Diagrams of remanence (right diagram) and coercive force curves of remanence (left diagram) for representative samples, with the sample information and coercive data shown in Table 1.

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Fig. 6. Thermal orthogonal demagnetization diagrams of representative samples from the Tierekesazi section (directions are shown in geographic coordination). Empty/solid circles represent vertical/horizontal projections; intensities are given in mA/m. NRM, natural remanent magnetization. Sample ID and thickness level are noted in each diagram.

yields LTC of D = 32.2°, I = 63.8°, α95 = 0 (0–100 °C) (in geographic coordination) and potential HTC of D = 19.2°, I = −4.8°, α95 = 17.7 (100–350 °C) (in strata coordination), with components chaotic in temperatures over 350 °C (Fig. 6d). The rest 130 samples with similar chaotic components are therefore excluded for determining polarities. Totally 662 samples revealed stable LTCs and HTCs. The LTCs have been correlated under geographic coordination, with an average direction D = 343.2°, I = 62.1°,α95 = 1.3 and κ = 19.8, consistent with the present-day geomagnetic field (Fig. 7a). The HTCs fall into normal and reversal polarities, which have been calculated their mean direction of D = 6.6°, I = 7.4°, α95 = 2.9, κ = 8.7 and D = 6.1°, I = 5.2°, α95 = 3.6, κ = 7.9 under geographic coordination (Fig. 7b) in contrast to D = 353.2°, I = 47.2°, α95 = 2.5, κ = 12.6 and D = 175.2°, I = − 46.3°, α95 = 3.2, κ = 10.6 under strata coordinate system (Fig. 7c). The reversal test (Mcfadden and McElhinny, 1990) against these results returns critical gamma of 4.2°at 95% confidence level, passing the test with the critical gamma value falling into ‘A’ classification. The fold test with

the method described by McElhinny (1964) returns κs/κg = 1.19, greater than 1.14 (F at 5%) and less than 1.20 (F at 1%), suggesting that these results pass the fold test at 95% confidence level. All the 662 samples yield 387 normal polarities and 275 reversal polarities (in strata coordination). Among these, 25, 517, 35, 79 and 6 reliable polarities have been obtained within the thickness intervals 756.6–805.6 m, 908.6–2341 m, 2341–3108.6 m, 3108.6–3719.6 m and 3719.6–3824.8 m, respectively, returning sample densities 2.0, 2.8, 22.0, 7.7 and 17.5 m/n (meters per sample) for respective segment. These results suggest that thickness intervals 756.6–805.6 m and 908.6–2341 are of high resolution, less than 3 m/n. Thickness interval 3108.6–3719.6 m with density of 7.7 m/n seems of acceptable resolution, when taking the potential high deposition rate into consideration (as shown in later results, Fig. 8). However, thickness intervals 805.6– 908.6 m, 2341–3108.6 m and 3719.6–3824.8 m obviously dissatisfy the resolution requirement of magnetostratigraphic study and therefore the results in the intervals should be regarded as indicative.

Fig. 7. Upper hemispherical equal-area projection of the low-termperature and high-temperature component (LTC and HTC respectively) directions (circles) and their average directions (stars). (a) LTC directions (solid circles) are shown in geographic coordination; the average direction (blue star) of the LTCs is consistent with that of present geomagnetic field. (b) HTC directions (circles) and their average directions (red stars) are shown in geographic coordination. Downward/upward directions are shown in solid/open circles. (c) HTC directions (circles) and their average directions (red stars) in strata coordination. Downward/upward directions are shown in solid/open circles.

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Fig. 8. Magnetostratigraphic results of the sediments in the Tierekesazi section, showing the correlations between measured magnetic polarities to the global polarity time scale (GPTS) of Gradstein et al. (2012) and the sediment accumulation rates (SAR) (compacted) of typical lithologic associations (blue line) and the described four lithofacies (gray boxes). The sediment particle size labels are the same with those in Fig. 3, and the exception Q/S represents (quartz) sandstone. SUT/A: stratigraphic unit terms/ages of the units. Dec.: declination; Inc.: inclination. Thickness intervals 0–756.6 m, 805.6–908.6 m, 2341–3108.6 m and 3719.6–3824.8 m are poorly constrained. In particular, the chronologic constraint for thickness interval 0–756.6 m is referred to Wang et al. (2014) with less confidence, and therefore the SAR in the interval is shown in dashed line.

5. Magnetostratigraphic interpretation 5.1. Biostratigraphic and lithostratigraphic constraints Bio- and litho-stratigraphic correlation would significantly help in correlating our measured paleomagnetic polarities to the geomagnetic polarity time scale (GPTS) (Gradstein et al., 2012). We here present our

own paleontological pollen analyses, complemented with a summary of regional bio- and litho-stratigraphic results, which provide reliable constraints for correlation between our measured polarities to the GPTS. The lower part of the measured section (thickness 756.6 to 0 m), terminologically referred as Bashibulake to Aertashi Formations (Fig. 8), lies out of the reach of magnetostratigraphic analyses due to difficulty in sampling. Fortunately, distinct lithological features of these formations,

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as noted above, provide solid evidence for lithostratigraphic correction of our described section to other section with detailed biostratigraphic constraint. These published biostratigraphic results provide approximate dates for the formations. As summarized by Wang et al. (2014), traditional biostratigraphic findings, accompanied by mineral and geochemical evidence, suggest that the Aertashi Formation was deposited during early Paleocene (roughly 65–62 Ma) (Guo, 1994; Guo et al., 2000; Hao et al., 2001; Jia et al., 2004; Tang et al., 1989; Wang et al., 2014; Yong et al., 1989). This correlation is supported by the widespread thick gypsum and dolomite beds, providing a distinctive mark for the correlation. In light of the Bashibulake Formation, detailed biostratigraphic analyses suggest that the deposition of the formation ended at approximately 34 Ma (see Wang et al., 2014 and Bosboom et al., 2014 for details). All these discussions suggest that this lower part of the measured section spans age interval of ca. 65–34 Ma.

In particular, three samples were obtained in our study from mud/ siltstone beds in this part of the section for paleontological pollen analyses, including Sample 1 from the thickness level of 367 m in the Qimugen Formation, Sample 2 (thickness level of 453 m) and Sample 3 (thickness level of 480 m) in the Wulagen Formation (Fig. 3 and Table 2). Sample 1 presents typical dinocyst assemblages Hystrichokolpoma–Fibrocysta paleocenica–Impletosphaeridium densum–Deflandrea oebisfeldensis (Table 2), which are widespread in Tarim basin during late Paleocene (e.g., He, 1991), suggesting the Qimugen Formation being late Paleocene in age. Samples 2 and 3 return similar dinocyst, chlorophyta and pollen assemblages, dominated by dinocyst fossils. The emergence of dinocyst assemblages Rhombodinium–Adnatosphaeridium–Glaphyrocysta, combined with the chlorophyta Crassophaera tuberculata–Pterosphermella packages (Table 2), provides diagnostic evidence to assign an Eocene age to the Wulagen Formation (e.g., He, 1991). This age assignment is

Table 2 Results of the analyzed paleontological pollen samples. Sample Name Formation Longitude Latitude depth (m)

Sample information

Palynological results

Pollen-1 Qimugen 74°35′07″ 39°54′54″ 367

Pollen-2 Wulagen 74°34′20″ 39°50′58″ 453

Pollen-3 Wulagen 74°34′06″ 39°50′54″ 480

Pollen-4 Pakabulake 74°33′12″ 39°49′55″ 1808

Pollen-5 Pakabulake 74°33′11″ 39°49′55″ 1819

Pollen-6 Pakabulake 74°33′08″ 39°49′50″ 1865

Pollen-7 Pakabulake 74°33′10″ 39°49′48″ 1889

Dinoflagellate Choratedinocysts Hystrichokolpoma Cordosphaeridium Fibrocystapaleocenica Hystrichosphaeridium Cleistosphaeridium Impletosphaeridiumdensum Deflandreaoebisfeldensis Lingulodinium Glaphyrocysta Adnatosphaeridium Rhombodinium

N100 3 5 1 10 5 2 1 0 0 0 0

N100 19 12 0 0 2 0 0 2 13 18 2

39 13 2 0 0 3 0 0 0 12 28 1

0 0 0 0 0 0 0 0 0 0 0 0

0 0 0 0 0 0 0 0 0 0 0 0

0 0 0 0 0 0 0 0 0 0 0 0

0 0 0 0 0 0 0 0 0 0 0 0

Chlorophyta Crassophaeratuberculata Pterospermella Pediastrum

0 0 0

N100 3 0

3 1 0

0 0 2

0 0 0

0 0 0

0 0 0

Pteridophyta Toroisporis Polypodiaceoisporites Leptolepidites

1 1 1

0 0 0

0 0 0

0 0 0

0 1 0

0 0 0

0 0 0

Needle-leaved Tree Pinuspollenites Piceaepoolenites Abiespollenites

1 0 0

0 0 0

0 0 0

65 0 7

74 0 0

29 2 2

38 6 18

Board-leaved tree Alnipollenites Quercoidites Juglanspollenites Ulmipollenites

0 0 0 1

0 0 0 0

0 0 0 0

1 3 0 12

0 0 1 0

0 0 0 0

0 0 0 0

Bush Ephedripites (D.) E. (E.) Elaeangnacites Betulaepollenites

7 1 1 0

26 0 0 0

13 0 0 0

2 0 0 10

6 0 0 3

1 0 0 0

0 0 0 3

Herbaceous plant Euphorbiacites Qinghaipollis Nitrariadites Labitricolpites Graminidites Chenopodipollis Cichorieacidites

0 0 0 0 0 0 0

6 4 5 1 0 0 0

0 2 1 1 0 0 0

0 1 0 0 2 16 2

0 0 0 0 3 9 1

0 0 0 0 0 3 0

0 0 0 0 1 0 0

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consistent with the finding of contemporaneous pollen assemblages Euphorbiacites–Qinghaipollis–Nitrariadites (Table 2) (Zhang and Zhan, 1991). Our own paleontological findings and age assignments are consistent with previous age placement ca. 65–34 Ma for this part of the section (thickness interval 756.6 to 0 m). The lowermost portion of our magnetostratigrapic section (thickness interval 756.6–908.6 m), terminologically referred as the Keziluoyi Formation (Fig. 8), has low resolution in sampling. However, published Foraminifera Pullenia quinqueloba–Cibicidoides ovaliformis–C. amygdaliformis (Hao et al., 1982) and Ostracods Mediocypris ordinate assemblages (Hao et al., 2002) suggest that the formation is most likely to be deposited in earliest Miocene, consistent with the finding of disconformity between the Keziluoyi and underlying Bashibulake Formations which may represent ca. 12 m.y. sedimentary hiatus. The lower part of the magnetostratigraphic section (thickness interval 908.6–2341 m), terminologically referred as Anjuan to midPakabulake Formations (Fig. 8), has been densely sampled. In addition, four paleontological samples were obtained in the mud/siltstone beds in the middle Pakabulake Formation, including Sample 4 (thickness level of 1808 m), Sample 5 (thickness level of 1819 m), Sample 6 (thickness level of 1865 m) and Sample 7 (thickness level of 1889 m) (Fig. 3 and Table 2). The four pollen samples (Samples 4 to 7) in the Pakabulake Formation present assemblages dominated by Pinuspollenites– Chenopodipollis–Betulaepollenites–Ulmipollenites–Ephedripites (D.), and also decreasing Chenopodipollis component from Samples 4 to 7 upward section (Table 2). The typical pollen assemblages are similar to the Chenopodipollis–Ephedripites (D.)–Tsugaepollenites identified in middle Miocene sequences in the adjacent Qaidam basin (Zhu et al., 1985). More recently, Wang and Shu (2013) summarized the Cenozoic xeromorphic vegetation in China and identified a significant increase of Chenopodipollis component during middle Miocene, which correlates well to the Middle Miocene Climate Optimum (MMCO) at ca. 14 Ma (Zachos, 2001). This increase is followed by a significant fall of Chenopodipollis component thereafter (Wang and Shu, 2013). Such a decreasing trend has also been obtained from Kuche region closer to our section (Zhang et al., 2011), with the beginning age of ca. 13.3 Ma dated by magnetostratigraphy (Sun et al., 2009). Our four pollen samples, although limited by sample amount and temporal interval, show such a similar profound fall in Chenopodipollis component from Samples 4 to 7, suggesting a likely placement of the stratigraphic level (thickness level of 1808 m), where the Sample 4 has been derived, into an age of 13–14 Ma, possibly corresponding the MMCO. This date claim is consistent with the paleontologic findings of the Pakabulake Formation in the Tarim basin, suggesting mid- to late Miocene age for the formation (Hao et al., 2002). The upper portion of our magnetostratigraphic section (thickness of 2341 m 3108 m) is characterized by scattered samples and therefore lower resolution (Fig. 8). This problem can be overcome, at least partly, by regional correlation of the distinct Xiyu Formation, a widespread stratigraphic formation (e.g., Liu et al., 1996) featured by its darker color compared to the underlying conglomerates (Huang et al., 1947). In spite of the widespread distribution of the Xiyu conglomerate formation, its depositional age has been a controversial issue for decades, which can be attributed to differentiate definitions of the Xiyu Formation. Initially, the Xiyu Formation was regarded as chronostratigraphic unit and its deposition was believed to begin at early Pleistocene (Huang et al., 1947; Jia et al., 2004; Liu et al., 1996), which was consistent with the regionally and synchronously sudden variation in sediment color—providing unique mark for regional correlation—and also supported by the regional discovery of the mammalian fossil Equus sanmeniensis within the lower formation (Huang and Ji, 1984; Jia et al., 2004). Molnar et al. (1994) and Burchfiel et al. (1999), on the basis of such paleontologic finding, followed this timing assignment and proposed the formation to be time-equivalent strata. More recently, the ‘Xiyu conglomerate formation’, regarded as a tectonostratigraphic unit arranged in gravel wedge forms, was suggested to be diachronous

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(time-transgressive) from 0.7 Ma to as old as 15 Ma (Charreau et al., 2005, 2006, 2009a,2009b; Chen et al., 2002, 2007; Heermance et al., 2007) based on magnetostratigraphic analyses. This stratigraphic division is valid in understanding the orogenic propagation by defining the Xiyu Formation to be ‘gravel wedge’ (possibly involving conglomerates previously being assigned into Atushi and even Pakabulake Formations), however, on the other hand neglects the obvious color variation from light gray to dark gray, which has been traditionally and chronostratigraphically defined as the boundary between the Xiyu Formation and underlying strata (e.g., Huang et al., 1947; Liu et al., 1996). We identified the Xiyu Formation in the field strongly based on the color change, therefore the traditional chronostratigraphic division (e.g., Huang et al., 1947; Liu et al., 1996) should be invoked in our study. In addition, based on this traditional terminology of the Xiyu Formation, high-resolution magnetostratigraphy and palynological analyses suggested that the boundary between the Xiyu Formation and underlying strata is ca. 2.6 Ma in age (e.g., Sun and Zhang, 2008; Sun et al., 2004), with their age claims correlating well to global climate events (Sun and Zhang, 2008; Zhang and Sun, 2011). We here follow this claim and assign the beginning deposition of the Xiyu Formation in our studied section to be ca. 2.6 Ma. The above discussions on bio- and litho-stratigraphic correlation, combined with the following correlation between our measured paleomagnetic polarities to the GPTS, would firmly constrain the ages of the stratigraphic units and observed sedimentary events as well. In the following Section 5.2, the similarity between measured polarity assemblages and the GPTS, accompanied with the bio- and lithostratigraphic correlation results, will be discussed to establish the magnetostratigraphic ages. 5.2. Magnetostratigraphic correlation Generally, the thickness intervals 3108.6–3719.6 m (Atushi Formation) and 908.6–2341 m (Anjuan to mid-Pakabulake Formations) have been sampled in high resolution and therefore can be expected to give reliable magnetostratigraphic ages. We begin the correlation between measured polarities and the GPTS (Gradstein et al., 2012) from these intervals. As previously discussed, the boundary between the Xiyu and Atushi Formations (thickness level of 3719.6 m) has been assigned to be ca. 2.6 Ma in age. This constraint leads to correlate the N1 magnetozone to the C2A.1n of the GPTS (Fig. 8). Downward until the thickness level of 3108.6 m (lower limit of the Atushi Formation), high-resolution samples force us to confidently and strictly correspond the measured polarity to the GPTS one by one, resulting into the correlation of N1–N7 to the C2A.1n–C3.4n (Fig. 8). These correlations suggest an age of ca. 5.2 Ma for the thickness level of 3108.6 m (lower limit of the Atushi Formation). Downward section, thickness interval 3108.6–2341 m provides scattered samples and low-resolution measured magnetozones (Fig. 8), which fails to give a reliable correlation with the GPTS. Therefore, we took a necessary jump to the lower section (thickness interval 908.6–2341 m), where dense samples have been analyzed which could be expected to be reliably correlated to the GPTS (Fig. 8). Within the thickness interval 908.6–2341 m (Anjuan to midPakabulake Formations), the above reliable biostratigraphic results indicate that the Pakabulake Formation (thickness interval 1423.6– 3108.6 m) is mid- to late-Miocene in age. This is also supported by speculative constraints of our own paleontological results which suggest that the thickness level 1808 m (Sample 4) is aged ca. 13–14 Ma. In this thickness interval, we first correlate the measured mangetozones to the GPTS based on their interval similarities, and then the biostratigraphic constraints are applied to test the correlation. The most distinct feature of the measured magnetozones in this interval is the R21–R24, characterized by 2 long reversed polarities (R21 and R24) with a zone punctuated by 3 normal polarities (N21, N22 and N23) and 2 reversed

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polarities (R22 and R23) (Fig. 8). This polarity assemblage can be easily correlated to C5B.2r to C5C.3r which shares the same interval pattern (Fig. 8). The rest of the measured polarities are then correlated to the GPTS one by one and these correlations yield upper and lower ages of 9.8 Ma and 20.4 Ma for this thickness interval 908.6–2341 m, respectively (Fig. 8). The correlations are supported by corresponding long normal polarity C5.2n of the GPTS, covering an age interval of ca. 1 Ma, to the measured long N12 polarity. In particular, these correlations provide an age interval of 9.8–17.1 Ma for the lower Pakabulake Formation and another age of ca. 13.4 Ma for the thickness level (1808 m) of paleontological Sample 4. It is consistent with the above biostratigraphic constraints (13–14 Ma in age), which independently testifies the validity of our correlations. Downward section, the lowermost magnetostratigraphic section (thickness interval 756.6–805.6 m) have been sporadically sampled which is separated from the above section by the interval 805.6– 908.6 m where there are no paleomagnetic samples. We speculatively assign N29–N31in the thickness interval 756.6–805.6 m to C6A.6n– C6B.2n of the GPTS which has given ca. 22.1 Ma for the lower boundary of the Keziluoyi Formation (Fig. 8). These correlations, although tentative, is supported by biostratigraphic results which suggest that the formation is earliest Miocene in age. The above discussions suggest a chronological framework for the sediments in the Tierekesazi section with some age assignments reliable while others remain speculative. The Paleogene sediments (including Aertashi to Bashibulake Formations with thickness interval 0–756.6 m) cover speculative age interval ca. 65–34 Ma. The earliest Miocene Keziluoyi Formation (thickness interval 756.6–908.6 m) spans from ca. 22.1 Ma to 20.4 Ma, with the upper limit more reliably constrained while the lower limit tentative (Fig. 8). The early Miocene Anjuan Formation is 20.4–17.1 Ma in age, with both boundary ages being reliable. The mid- to late-Miocene Pakabulake Formation spans the age interval 17.1– 5.2 Ma (Fig. 8), with the lower boundary age reliable in contrast to the upper limit age speculative. However, the field investigation indicates that the Atushi Formation overlies the underlying Pakabulake Formation conformably, which implies that the reliable lower boundary age 5.2 Ma can be assigned to represent the upper limit age of the Pakabulake Formation. The Pliocene Atushi Formation covers ages from 5.2 Ma to 2.6 Ma (Fig. 8), with the upper limit age speculated to be the lower boundary age of the overlying Xiyu Formation, which is consistent with the conformable contact between the two formations as suggested by the field investigation. The Quaternary Xiyu Formation begins its deposition at 2.6 Ma, with the upper limit age unknown in the Tierekesazi section. This chronological framework provides age constraints for the observed four lithofacies, with lithofacies (i) ca. 65–34 Ma, (ii) ca. 22.1– 12 Ma, (iii) 12–5.2 Ma, and (iv) 5.2 Ma–present (?) (Fig. 8). 6. Interpretation of sediment accumulation rate The sediment accumulation rate (SAR) (Fig. 8) can be determined from magnetostratigraphic ages and measured sediment thicknesses by using sediment thickness divided by age interval. The SAR results can be primarily applied to evaluate the validity of the magnetostratigraphic dates. Generally, sediment associations with coarse particle size always present higher SARs compared to those with finer particle size. In addition, the SAR results can also be used to assess the deposition rates of the lithofacies that we have divided in the Tierekesazi section. We follow this strategy to firstly calculate the SARs of the sediment associations with similar particle size and then calculate the SARs of the four lithofacies. As shown in Fig. 8, SARs of the sediment associations are strictly correlated to the sediment particle sizes, especially in the reliable magnetostratigraphic intervals 908.6–2341 m and 3108.6–3719.6 m. Taking the thickness interval 965–1213.6 m for example, the SAR of the mudstone association is 7.3 cm/ka, increasing to 15.7 cm/ka in the siltstone association contrasting to 19.3 cm/ka in the sandstone association

(Fig. 8). This correlation between SAR and particle size is also valid in the upper reliable magnetostratigraphic interval, 13.3 cm/ka in the 3295.9–3536.9 m sandstone association in contrast to 90.4 cm/ka in the 3536.9–3676.5 m conglomerate association (Fig. 8). This strict correlation between SAR and particle size primarily testifies the validity of the magnetostratigraphic results. The SARs of the four lithofacies are based on the date and thickness data discussed above: lithofacies (i) 0–756.6 m with age interval ca. 65– 34 Ma; lithofacies (ii) 756.6–1997 m with age interval ca. 22.1–12 Ma; lithofacies (iii) 1997–3108.6 m with age interval 12–5.2 Ma; and lithofacies (iv) 3108.6–N 4294 m with age interval 5.2 Ma–present (?). The calculations of sediment thickness divided by age interval yield SARs of ca. 2.4 cm/ka for the lithofacies (i), 12.3 cm/ka for the lithofacies (ii), 16.3 cm/ka for the lithofacies (iii) and N22.8 cm/ka for the lithofacies (iv), respectively (Fig. 8). These results indicate that accompanied with the upward coarsening section, the SAR increases significantly. The SARs can be affected by the compaction effect. Ideally, the compaction effect would be well constrained by acoustic logging data from boreholes. Lacking of borehole logging data, the compaction effect on SARs can be roughly accessed by using the depth vs. porosity curve of the strata, with which the porosity decrease with the increase of burial depth can be determined. Zhang et al. (2011) reported the depth vs. porosity curve in the Kuche region east of the studied area, which was employed here to determine the decompacted SARs in the Tierekesazi section. The calculation indicates that the decompacted SARs, ca. 3.3– 3.5 cm/ka in the lithofacies (i), 16–17 cm/ka in the lithofacies (ii) and 19.5–20.6 cm/ka in the lithofacies (iii), vary slightly compared with the compacted ones. The compaction effect of the uppermost lithofacies (iv), which dominantly comprises gravels, can be ignored (e.g., Prothero and Schwab, 1996). The discussion suggests that the significant SAR increases upward section remain valid when taking the decompaction into consideration.

7. Discussions 7.1. Regional correlation of the observed sedimentary events Our sedimentologic investigation in the Tierekesazi section in the foreland region of south WTS has identified three striking events, including the first appearances of (1) underwater channel sandstone at the bottom of the Keziluoyi Formation (thickness level of 756.6 m), (2) pebble gravel sheet in the middle of the Pakabulake Formation (thickness level of 1997 m), and (3) massive cobble conglomerate at the bottom of the Atushi Formation (thickness level of 3108.6 m) (Fig. 9). These events, combined with the facts of sedimentary facies variation, allow us to simply divide the whole Tierekesazi succession into four lithofacies, Paleogene marine lithofacies (i) from the Aertashi to Bashibulake Formations (0–756.6 m), lacustrine to fluvial (plain) lithofacies (ii) from the Keziluoyi to middle Pakabulake Formations (756.6–1997 m), alluvial sand–gravel sheet lithofacies (iii) in the upper Pakabulake Formation (1997–3108.6 m), and conglomerate lithofacies (iv) from the Atushi to Xiyu Formations (3108.6–N 4294 m) (Fig. 9). Magnetostratigraphic analyses, combined with biostratigraphic results, provide chronological constraints and SAR estimations for the lithofacies, suggesting lithofacies (i) aged ca. 65–34 Ma with SAR of ca. 2.4/3.3–3.5 (compacted/decompacted) cm/ka, lithofacies (ii) aged ca. 22.1–12 Ma with SAR of 12.3/16–17 cm/ka, lithofacies (iii) aged 12–5.2 Ma with SAR of 16.3/19.5–20.6 cm/ka, and lithofacies (iv) aged 5.2–present (?) with SAR of N22.8/N22.8 cm/ka. All the above results collectively suggest three individual sedimentary events at ca. 22.1 Ma, 12 Ma and 5.2 Ma respectively, as evidenced by significant variation in sedimentary facies, particle size and SAR. The sedimentary events can be linked to local/regional tectonic activity or, alternatively, to global climatic or eustatic change (e.g., An et al., 2001; Herman et al., 2013; Zhang et al., 2001). This argument can be partially

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Fig. 9. Summaries of the sediment particle size, magnetostratigraphic ages, and sediment accumulation rates (SAR) (compacted) in the Tierekesazi profile. The SAR of Paleogene sequences, out of the reach of our magnetostratigraphic samples, is calculated from age estimations by Wang et al. (2014) as described in the text, shown in dashed line. Sediment particle size labels are the same with those in Fig. 8.

solved by integrating basinfill records with mountain range uplift and structural deformation constraints (e.g., Lin et al., 2011). The ca. 22.1 Ma sedimentary event, represented by disconformable contact (possibly marking a ca. 12 m.y. hiatus), increasing particle size, shifting from marine to terrestrial facies and significant sediment accumulation acceleration in basinfill records of the Tierekesazi section (Fig. 9), is also independently recorded by thermochronologic results from samples in mountain ranges. Hendrix et al. (1994) conducted

apatite fission track analysis with the Mesozoic strata samples in the northern flank of the Chinese Tian Shan and their results indicated that this region underwent approximately 4–5 km of unroofing beginning at ca. 24 Ma. Similar results were also reported by Yang et al. (2003), in which the Late Paleozoic granite samples in Kuche regions of CSTS presented ages of 17–25 Ma, consistent with the cooling paths starting at ca. 25 Ma indicated by the thermal history modeling with track age and length data (Du and Wang, 2007; Wang et al., 2009).

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Apatite (U–Th)/He analysis of Pre-Mesozoic sandstone samples in the Kalpin and north Bachu Uplift showed significant age clusters at 22– 20 Ma (Chang et al., 2012). This claim was testified by apatite fission track results in the south of sTS (including both KSTS and WTS portions), which suggested an initial exhumation at the Oligo-Miocene boundary (ca. 24–22 Ma) (Chen et al., 2008; De Grave et al., 2012; Sobel and Dumitru, 1997; Sobel et al., 2006), with the localities more close to our studied section. Low-temperature thermochronological results indicated that this event may also distribute in KCTS (e.g., Macaulay et al., 2014). Moreover, facies transition and fold-and-thrust belt deformation study suggested that the sTS underwent initial activity at 24–21 Ma (Yin et al., 1998), consistent with results from structural deformation interpretation of seismic profiles in the Kuche region (e.g., Wang et al., 2002). The coincidence between the uplift in the mountain ranges and the basinfill sedimentary event forces us to link the 22.1 Ma sedimentary event to the uplift of the sTS. The ca. 12 Ma sedimentary event is represented by the first appearance of pebble gravel sheets, sedimentary facies shifting from lacustrine–fluvial (plain) to alluvial sand–gravel sheet lithofacies and a slight SAR increase (Fig. 9). We noticed that the depositional acceleration has been widely reported along the sTS. In the northern flank of the Kashi foreland basin, Heermance et al. (2007) observed that the SARs increased from approximately 15 ± 3 cm/ka to about 43 ± 1 cm/ka at 13.5 Ma in the region south of the Tashipishake Anticline (Fig. 10). Accompanying their growth strata and deformation propagation observations, they attributed this acceleration event to the southward stepping propagation of the foreland fold-and-thrust belt and the uplift of the Tashipishake

Anticline (Heermance et al., 2007). Synchronous depositional acceleration and structural deformation have also been observed in the Kuche fold-and-thrust belt (Li et al., 2012a,2012b; Wang et al., 2011; Zhang et al., 2014). In the Yaha section, located in the Kuche region on the southern flank of the CSTS, Charreau et al. (2006) dated an acceleration of SAR from 20 ± 7 to 43 ± 16 cm/ka at ca. 11 Ma (Fig. 10), along with the average grain shape shifting from spherical to oblate, which was linked to the acceleration in erosion and uplift of the sTS. Reinterpretation of Huang et al.'s (2006) magnetostratigraphic results in another section approximately 10 km north to the Yaha section also suggest a contemporaneous accumulation acceleration event at ca. 11–10 Ma (Charreau et al., 2008) (Fig. 10). Such event has been also indicated by thermochronological records in south WTS and KCTS, suggesting significant cooling in ca. 15–8 Ma (e.g., De Grave et al., 2012; Macaulay et al., 2014; Sobel and Dumitru, 1997; Sobel et al., 2006, 2013). Again, the coincidence between the uplift in the mountain ranges, the deformation of the foreland fold-and-thrust belts and the basinfill sedimentary event leads to an interpretation that the observed sedimentary event in the study results from the uplift of the sTS. The latest ca. 5.2 Ma sedimentary event is recorded by the first appearance of massive cobble conglomerates, facies transition from alluvial sand–gravel sheet to conglomerate lithofacies and SAR acceleration in the Tierekesazi section (Fig. 9). This event is consistent with the particle size and magnetic susceptibility change in the Ulugqat section approximately 10 km west to our studied section (Wang et al., 2014). Regional correlation suggests that this event marks another episode of regional tectonic activity. In the Kuche region, Huang et al. (2006)

Fig. 10. Sediment accumulation rate (SAR) (compacted) comparison between the Tierekesazi section in this study and other sections in the foreland regions of south Tian Shan, with the locations of the sections shown in the right-lower sketch tectonic map. The SARs in the Kashi region are the averages of results from Heermance et al. (2007); SARs in the Kuche region are from Sun et al. (2009), Charreau et al. (2006) and the reinterpreted results by Charreau et al. (2008) from original data of Huang et al. (2006).

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observed SAR increasing from approximately 13 cm/ka to 23 cm/ka at ca. 7 Ma (Fig. 10), temporally close to the 6.5–5.2 Ma occurrence of the syntectonic growth strata in the Qiulitag and Yaken Anticline and the SAR rising from 32.5 cm/ka to 40.3 cm/ka at ca. 6.5 Ma (Fig. 10) (Wang et al., 2002; Sun et al., 2009). The SAR increasing event is consistent with the 5.2 Ma coarse conglomerate in the Yaha section, which prevented Charreau et al. (2006) from continuously magnetostratigraphic sampling. In the Kashi basin, with its locality closer to our studied section, the SAR increased from an average of approximately 43 cm/ka to about 70 cm/ka at ca. 5.3 Ma (Fig. 10), accompanied with continuous thrusting propagation indicated by growth strata in the Kektamu Anticline (Heermance et al., 2007). Synchronous thrusting propagation toward the basinward has also been reported in the Kuche fold-and-thrust belt (Li et al., 2012a,2012b; Wang et al., 2011; Zhang et al., 2014). This episode of tectonic activity likely established significant mountain ranges around the Tarim basin, leading to increased aridity in both the Tarim and Junggar basins at ca. 5.3 Ma which was inferred to result from rain shadow effect of mountain uplift (e.g., Sun et al., 2013). The speculation has been supported by the mountain uplifting event observed in the WTS (e.g., De Grave et al., 2012). Again, the coincidence between the sedimentary event, fold-and-thrust belt deformation and mountain building suggests the attribution of sedimentary event to the tectonic activity along sTS. 7.2. Tectonic implications The Oligo-Miocene boundary event has been primary interpreted as the initial response of intracontinental deformation to the Indian– Eurasian convergence (e.g., Hendrix et al., 1994; Sobel et al., 2006; Soble and Dumitru, 1997). However, such a far-field response should be expressed as more direct and local driving force. Sobel et al. (2013) attributed the Oligo-Miocene boundary thermochronological ages observed in the margins of southwest Tarim basin to the initial underthrusting of the Tarim block beneath the surrounding orogens, which has been testified by geophysical data in the sTS, Pamir and West Kunlun (e.g., Matte et al., 1996; Gao et al., 2000; Kao et al., 2001; Li et al., 2001; Buslov et al., 2007; Li et al., 2012; Jiang et al., 2004; Wittlinger et al., 2004; Schurr et al., 2014; Sippl et al., 2013a, 2013b; Schneider et al., 2013). The widespread thermochronological records of this episode in West Kunlun and also sTS (Hendrix et al., 1994; Yang et al., 2003; Du and Wang, 2007; Wang and Du, 2009; Chang et al., 2012; Sobel and Dumitru, 1997; Sobel et al., 2006; Chen et al., 2008; De Grave et al., 2012; Macaulay et al., 2014) imply that the onset of underthrusting of the Tarim block during the this episode was not only limited in southwest margin of the Tarim basin, but can also take place in the south and north margins of the basin. The mid-Miocene and Pliocene events are generally synchronous along the sTS (Fig. 10). They may represent expressions of regional geodynamics (e.g., Sobel et al., 2013, 2011). However, these episodes of event more directly express as basinward propagation of the foldand-thrust belt in the sTS region east to the TFF, i.e., the Kashi, Keping and Kuche fold-and-thrust belts, although the accelerated mountain building may also exist (Fig. 1). Taking the Kashi fold-and-thrust belt, which is close to the TFF, for example, it episodically propagated to the basinward in a maximum distance of ca. 62 km, with the active timing of different rows of structure varying from as early as 18.9– 16.3 Ma to present (Heermance et al., 2007; Sobel et al., 2006). The Kashi fold-and-thrust belt absorbed a total shortening of 66–76 km (Heermance et al., 2007). In contrast to such expressions along the southern foreland region of KSTS–CSTS east to the TFF, the foreland region of the south WTS has no such significant expression of fold-andthrust belt propagation. Taking the southernmost thrust that controls the anticline–syncline complex in the studied Tierekesazi section for example, it propagated to the basinward by ca. 8 km (Luo, 2005), which is magnitude-scale less than those in the region east to the TFF. Although

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the shortening amount of the thrust and related anticline–syncline complex remains unclear, very limited shortening amount, less than 10 km, can be expected. The contrast in structural expression and shortening amount between the regions east and west to the TFF implies that the TFF should activate since mid-Miocene and have dextrally slipped by ca. 60–70 km, comparable with the estimation of ca. 100 km proposed by Burtman et al. (1996). It is notable that the north-vergent Pamir Frontal Thrust (belonging to the Pamir domain) has activated since mid-Miocene as well (Coutand et al., 2002). This result implies that the Pamir underwent significant northward indentation since that time. We tentatively attribute the activity of the TFF and the structural contrast in opposite regions of the TFF since mid-Miocene to the northward indentation of the Pamir.

8. Conclusions Our new sedimentologic and magnetostratigraphic investigations on the Tierekesazi section in the foreland regions of the south West Tian Shan provide results for a more comprehensive understanding of the foreland deposition process in south Tian Shan (sTS). Within a section of generally upward increasing particle sizes, three striking events have been identified, represented by the first appearances of (1) underwater channel sandstone at the bottom of the Keziluoyi Formation, (2) pebble gravel sheet in the middle of the Pakabulake Formation, and (3) massive cobble conglomerate at the bottom of the Atushi Formation. These events divide the Cenozoic succession into four lithofacies: (i) Paleogene marine lithofacies from the Aertashi to Bashibulake Formations, (ii) lacustrine to fluvial (plain) lithofacies from the Keziluoyi to middle Pakabulake Formations, (iii) alluvial sand–gravel sheet lithofacies in the upper Pakabulake Formation, and (iv) conglomerate lithofacies from the Atushi to Xiyu Formations. These new findings suggest three episodes of particle size and sedimentary energy increase in the Tierekesazi region. Our magnetostratigraphic results, combined with biostratigraphic correlations, provide the chronologic constraints for each lithofacies and also the sediment accumulation rates (SAR). These results indicate lithofacies (i) aged ca. 65–34 Ma with SAR of ca. 2.4/3.3–3.5 cm/ka (compacted/decompacted), lithofacies (ii) aged ca. 22.1–12 Ma with SAR of 12.3/16–17 cm/ka, lithofacies (iii) aged 12–5.2 Ma with SAR of 16.3/19.5–20.6 cm/ka, and lithofacies (iv) aged 5.2–present (?) with SAR of N22.8/N 22.8 cm/ka. The SAR accelerations during these episodes are well consistent with the particle size and sedimentary energy increase. Regional correlation has suggested that these three episodes of sedimentary events are consistent with synchronous mountain building and/or foreland fold-and-thrust belt deformation events, implying a linkage of the sedimentary event to the tectonic activity along sTS. We propose that the earliest Miocene event may represent the initial response of the far-field effect of Indian–Eurasian convergence, but more directly and likely marks the initial underthrusting of the Tarim block beneath the sTS. The mid-Miocene and Mio-Pliocene boundary events have different structural expression in the opposite regions east and west to the Talas–Fergana fault (TFF). The events are likely resulted from significant basinward propagation of foreland foldand-thrust belts along the sTS in the regions east to the TFF. In the studied Tierekesazi region west to the TFF, the events are tentatively related to accelerated mountain building in the south West Tian Shan. The difference of the shortening amounts in opposite foreland regions east and west to the TFF is proposed to cause the dextral slipping of the TFF, with its dextral slipping amount at least 60–70 km. More fundamentally, such structural contrast and the activity of the TFF are likely driven by the northward indentation of the Pamir at this time.

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Acknowledgments This research is supported by the National Natural Science Foundation of China (grant nos. 41330207, 41102128, 41472181 and 41072154), the National S&T Major Project (grant no. 2011ZX05009-001), the Scientific Research Fund of Zhejiang Provincial Education Department (no. Y201019040 and Y201224590), and the Fundamental Research Funds for the Central Universities (nos. 2014FZA3007 and 2015QNA3015). Thank Geoffrey E. Batt, Karl-Heinz Wyrwoll, Karen Lee Wyrwoll, Kang Li, Dan Liu, Biao Bi and Xiaogen Fan very much for their effective help in sampling and field work. We also acknowledge Yongxiang Li, Junpeng Li, Bin Wen and Xiaoqing Pan for their Laboratory support. We are grateful for the comments from an anonymous reviewer and Yann Rolland and Editor Laurent Jolivet that significantly improve the manuscript.

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