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Climatic and tectonic controls on the fluvial morphology of the northeastern Tibetan Plateau (China) Wang, X.

2014

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Download date: 02. Oct. 2021 CONTENTS

1. Introduction 4

1.1 BACKGROUND 4 1.2 STUDY AREA 5 1.3 THE TECTONIC FRAMEWORK 6 1.4 OBJECTIVES 9 1.5 OUTLINE OF THE THESIS 9 1.6 REFERENCES 10

2. Late Miocene uplift of the NE Tibetan Plateau inferred from basin filling, planation and fluvial terraces in the Huang Shui catchment 17 2.1 INTRODUCTION 18 2.2 GEOLOGICAL AND GEOMORPHOLOGICAL BACKGROUND 20 2.3 METHODS 22 2.4 RESULTS 23 2.5 DISCUSSION 28 2.6 CONCLUSIONS 34 2.7 REFERENCES 35

3. Differential impact of small-scaled tectonic movements on fluvial morphology and sedimentology (the Huang Shui catchment, NE Tibet Plateau) 44 3.1 INTRODUCTION 45 3.2 TECTONIC SETTING OF THE STUDY REGION AND CONSEQUENT GENERAL MORPHO- LOGICAL EXPRESSION 48 3.3 METHODS 50 3.4 RESULTS 52 3.5 TERRACE DEVELOPMENT 70

3.6 DISCUSSION 75 3.7 CONCLUSIONS 78 3.8 REFERENCES 79

4. Late Quaternary paleoclimatic and geomorphological evolution at the interface between the Menyuan basin and the Qilian Mountains, northeastern Tibetan Plateau 87 4.1 INTRODUCTION 88 4.2 GEOMORPHOLOGICAL SETTING 90 4.3 METHODS 91 4.4 RESULTS 94 4.5 DISCUSSION 105 4.6 CONCLUSIONS 112 4.7 REFERENCES 113

5. Differential tectonic movements in the confluence area of the Huang Shui and Huang He rivers (Yellow River), NE Tibetan Plateau, as inferred from fluvial terrace positions 121 5.1 INTRODUCTION 123 5.2 GEOLOGICAL AND GEOMORPHOLOGICAL SETTING 124 5.3 METHODS 125 5.4 RESULTS 129 5.5 DISCUSSION 137 5.6 CONCLUSIONS 142 5.7 REFERENCES 143

6. SYNTHESIS 149 6.1 MORPHO-TECTONIC AND GEOMORPHOLOGICAL EVOLUTION OF THE NORTHEASTERN TIBETAN PLATEAU 149 6.2 MORPHOLOGY OF TERRACE STAIRCASE AND ITS RESPONSE TO TECTONIC MOVEMENT AND CLIMATIC CHANGE 149 6.3 QUARTZ OPTICALLY STIMULATED 2

LUMINESCENCE DATING 151 6.4 FLUVIAL GEOMORPHOLOGICAL PROCESSES AND TERRACE FORMATION IN RELATION TO A NEARBY GLACIAL ENVIRONMENT 152 6.5 LOCAL TECTONIC MOVEMENTS AND THE COUPLED IMPACTS OF TECTONIC AND CLIMATE CHANGES ON FLUVIAL EVOLUTION SINCE THE LAST INTERGLACIAL 153 6.6 CONCLUSIONS AND FUTURE OUTLOOK 154 6.7 REFERENCES 155

SUMMARY 157

SAMENVATTING 161

-Chapter 1- Introduction

1.1. Background Rivers play a dominant role in the earth surface process system on the continents, providing the major pathways for both the water and sediment fluxes, causing material transport from land to ocean. The evolution of river systems is controlled by internal processes and external forces induced by climate change, tectonic movements and anthropogenic influences. Unravelling the relative contributions of these factors is one of the most challenging goals of fluvial geomorphological research (Schumm, 1965; Vandenberghe, 1995a; Bridgland et al., 2000; Bridgland and Westaway 2008). Alternations of aggradation and incision, and the formation of fluvial terraces may reflect the different impacts of individual driving forces. As concerns climate forcing, morphological cyclicity in temperate and cold regions has generally been attributed to cold (glacial)-warm (interglacial) alternations in settings with a background tectonic uplift (e.g., Büdel, 1977; Vandenberghe, 1993, 1995b; Van den Berg, 1996; Maddy et al., 2001; Starkel., 2003; Lewin and Gibbard, 2010). But, similar climate-driven terrace staircases do exist also in other climatic environments (e.g., Bridgland and Westaway, 2008; Pan et al., 2009; Wang, this study). The climate-driven model fails to explain the progressive Quaternary valley incision, which is suggested to be the result of fluvial system adjustment to long-term regional uplift. The degree of regional uplift is a key issue in the generation and subsequent preservation of terrace flights (Maddy et al., 2001; Bridgland and Westaway, 2008), as it can force a river system to incise during each climate cycle to separate terrace levels adequately. Terraces can be generated only when the regional uplift rate is sufficiently high (e.g., Pan et al, 2009). Based on the assumption that terrace surfaces record net incision driven by tectonic uplift, the vertical separation and the longitudinal profiles of terrace surfaces have widely been used to infer tectonic uplift rate (e.g., Van Balen et al., 2000; Maddy et al., 2001; Van den Berg and Van Hoof, 2001; Peters and Van Balen, 2007).

In contrast to large-scale tectonic movements, tectonic differentiation at small scale may lead to a variegated fluvial morphology. For example, in the Huang Shui catchment (Northeastern Tibetan Plateau) areas of fluvial incision and -aggradation spatially alternate and erosion and accumulation terraces simultaneously develop at relatively short distances indicating effects of relative uplift and subsidence on relatively small spatial scales (Wang et al., 2010; Vandenberghe et al., 2011). The relation between forcing intensity by climate and tectonics and fluvial response is not linear due to the preponderant role of delay effects and thresholds (Schumm, 1979; Knox, 1972; Vandenberghe, 2002). For instance, fluvial incision may lag behind uplift as the river may be dependent on a climatic change to enable incision to take place (Maddy et al., 2001). Tectonic movements and climate are independent forcing factors and their interplay may result in conflicting, or at least complex, effects on the fluvial morphology. Until now, both forcing factors have been approached as separate as possible. However, it is challenging to investigate the complex result of their interference.

1.2. Study area The study area in this thesis is situated in the north-eastern part of the Tibetan Plateau (NETP), defined as the region bounded by the Kunlun fault to the south and the Altyn and the Haiyuan faults to the north (Fig. 1.1). This area is thought to be the tectonically most active part of the Tibetan Plateau and a key region for the transfer of deformation by continental extrusion to the east, caused by the collision of India with Asia (e.g., Meyer et al., 1998; Tapponnier et al., 2001). The crustal deformation primarily due to these relative plate motions causes a complex tectonic system with folds, strike slip faults, thrust faults, normal faults, and related high mountains and basins. The general uplift of the Tibetan Plateau during the Neogene and Quaternary has led to drastic climate changes, for example several glacial advances in the Qilian Shan and the upper Huang He catchment during the Quaternary. Thus the river catchments in this region represent an ideal area for studying the impacts of the coupling of tectonics and climate changes on fluvial morphological development and the resulting sedimentary architecture of fluvial deposits. 5

1.3. The tectonic framework Large tectonic intra-continental basins were formed since the Mesozoic in the NETP. For example, the Longzhong basin (Fig. 1.2) may be considered as a fault-bounded structure in the NETP during the Late Jurassic to Early Cretaceous and subsequently reactivation of subsidence created further accommodation space for sediment accumulation from the Paleocene to the Middle Miocene (Horton et al., 2004; Dai et al., 2006). The starting time and process of deformation and uplift of the NETP are still debated. The results from different studies (e.g., Tapponnier et al., 1990, 2001; Burchfiel et al., 1991; Meyer et al., 1998; Lu et al., 2004; Fang et al., 2005; Dai et al., 2006; Lu and Xiong, 2009; Yin, 2010) based on the dating of basin sediment sequences and thermo-chronological histories in uplifted ranges are quite incompatible. Some studies indicate that the NETP undergoes a most recent uplift; the area is referred to as the „Pliocene-Quaternary Tibet (Tapponnier et al., 2001). But others suggest that the NETP began to uplift much earlier, in the middle and late Miocene (e.g., Lu et al., 2004; Zheng et al., 2006; Lu and Xiong, 2009) or even from the Eocene-Oligocene onward (e.g., Yin, 2009). It was argued that Neogene northward motion of the Tibetan Plateau was accomplished by thrust- and strike-slip faulting combined with uplift resulting in the segmentation of the large basins (Horton et al., 2004; Dupont-Nivet et al., 2008; Abels et al., 2011; Xiao et al., 2010, 2012). Consequently, sub-basins were formed, such as the Xining-, the Linxia- and Xunhua Basins (Fig. 1.2) (e.g., Garzione et al., 2005; Dai et al., 2006; Hough et al., 2011). These sub-basins, as the older large basins, are also filled with thick continental clastic red beds, while simultaneously the surrounding highs were bevelled (as in the areas surrounding the Xining basin; Wang, this study). Subsequently, the area was incised by rivers in response to regional tectonic uplift of the Tibetan Plateau and local deformation affected the red basin fills, probably during the Pliocene and Pleistocene (Li et al., 1997; Lu et al., 2004; Fang et al., 2005; Gao et al., 2008; Stroeven etal., 2009; Craddock et al., 2010).

Figure 1.1 Simplified map of major tectonic boundaries and Tertiary faults in northeastern Tibetan Plateau (box) (modified after Dai et al., 2006)

Figure 1.2 Topographic map of the NE Tibetan Plateau showing the location of the Cenozoic sedimentary basins (Longzhong-, Xining- and other basins). The insets show the position of the studied region on the Tibetan Plateau (grey line is the boundary of Longzhong basin).

Morphological analyses (especial fluvial morphology) in the NETP may contribute substantially to unravel the tectonic history of the area (Craddock et al., 2010; Pan et al., 2012; Wang, this study) like in other areas (e.g., Davis 1911; Ollier and Pain, 2000; Burbank and Anderson, 2001; Garcia-Tortosa et al., 2011). In addition, rivers are very sensitive to changes in gradient and respond relatively rapid to tectonic movements (e.g., Burbank and Anderson, 2001; Westaway et al., 2002). For example, river terraces are now widely used as morphotectonic markers to study the magnitude and rate of tectonic uplift (e.g., Li et al., 1996; Maddy et al., 1997; Van Balen et al., 2000; Burbank and Anderson, 2001; Lu et al., 2004a; Peters and Van Balen, 2007; Viveen et al., 2012). Thus, not only should fluvial development in the NETP be explained by

8 tectonic movements, but in addition the fluvial morphology can provide information on the (rates of) tectonic movement.

1.4. Objectives Bearing in mind the previous considerations, the aims of this thesis are defined as follows: 1) to reconstruct paleoclimatic changes and landscape evolution using OSL dating; 2) to infer the uplift and deformation processes in the study area, the NETP; 3) to investigate in detail the correlation of fluvial terrace and morphological evolution near the confluence region of the Huang Shui and Huang He rivers, and based on this, to study the local tectonic movements; 4) to investigate fluvial incision and sedimentation and terrace morphology in the Huang Shui catchment and their relationship with climate changes and tectonic movement in the NETP.

1.5. Outline of the thesis This thesis contains four journal papers (chapters 2 to 5), and a synthesis (Chapter 6). The journal papers are presented in their original form, except for some lay-out changes. Consequently, some repetition in the text with respect to regional setting, methods and references is inevitable. Chapters 2 deals with morpho-tectonic characters and geomorphological evolution of Huang Shui catchment, concentrating primarily on a morphotectonic analysis of the Huang Shui catchment using a digital elevation model and field observations, reconstruction and dating of the geomorphological history using the biochronology, fluvial terraces and the depositional record of the basin fill. Chapter 3 presents the sedimentation and morphology of terrace staircases at different temporal scales (from Neogene to late Quaternary) and spatial scales (from regional to individual blocks) in the Huang Shui catchment. In addition, the sedimentation and terrace formation processes in different blocks

9 are discussed to distinguish the tectonic and climatic impacts on fluvial evolution. Chapter 4 presents depositional and morphological connections in a system consisting of connected glacial-, fluvio-glacial, and fluvial environments, situated in the Qilian Mountains and the Menyuan basin. It concentrates primarily on distinguishing facies units, dating of sedimentation and terrace sequences, and the reconstruction of landscape evolution and climate changes from the penultimate glacial onwards. Chapter 5 deals with dating and correlation of terrace staircases in the confluence region of the Huang Shui and Huang He rivers. The results are combined with previous results to infer tectonic motions in different blocks in the Huang Shui and Huang He catchments since the last interglacial, The specific coupled impacts of tectonic and climate forcings on fluvial process since the last interglacial are also discussed. Chapter 6 provides a synthesis of the results presented in Chapter 2 to 5, dealing specifically with the response of fluvial processes to climatic change and tectonic movements in the NETP and proposes an outlook for future research.

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Van den Berg, M., 1996. Fluvial sequences of the Maas: a 10 Ma record of neotectonics and climatic change at various time-scales. Ph.D. Thesis, University Wageningen, The Netherlands, 180 p. Van den Berg, M., Van Hoof, T., 2001. The Maas terrace sequence at Maastricht, SE Netherlands: evidence for 200 m of late Neogene and Quaternary surface uplift. In: Maddy, D., Macklin, M., Woodward, J. (Eds.), River Basin Sediments Systems: Archives of Environmental Change. Balkema, Rotterdam, The Netherlands, 45-86. Vandenberghe, J., 1993. Changing fluvial processes under changing periglacial conditions. Zeitschrift fur Geomorphologie, Supplement Band 88, 17-28. Vandenberghe, J., 1995a. The role of rivers in palaeoclimatic reconstruction. In Frenzel, B., Vandenberghe, J., Kasse, C., Bohncke, S. and Gläser, B. (Eds.), European river activity and climatic change during the Lateglacial and early Holocene, Paläoklimaforschung 14, 11-19. Vandenberghe, J., 1995b. Timescales ， climate and river development. Quaternary Science Reviews 14, 631-638. Vandenberghe, J., 2002. The relation between climate and river processes, landforms and deposits during the Quaternary. Quaternary International 91, 17-23. Vandenberghe, J., Wang, X.Y., Lu, H.Y., 2011. The impact of differential tectonic movement on fluvial morphology and sedimentology along the northeastern Tibetan Plateau. Geomorphology 134, 171-185. Viveen, W., Van Balen, R.T., Schoorl, J.M., Veldkamp, A., Temme, A., Vidal- Romani, J.R., 2012. Tectonic geomorphology and neotectonic activity of the NW Iberian Atlantic Margin with examples from the Miño river system. Tectonophysics 544-545, 13-30. Wang, X.Y., Lu, H.Y., Vandenberghe, J., Chen, Z.Y., Li, L.P., 2010. Distribution and forming model of fluvial terrace in Huangshui catchment and its tectonic indication. Acta Geologica Sinica 84, 415-423.

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-Chapter 2-

Late Miocene uplift of the NE Tibetan Plateau inferred from basin filling, planation and fluvial terraces in the Huang Shui catchment Xianyan Wang, Huayu Lu, Jef Vandenberghe, Shaohua Zheng and Ronald van Balen Abstract: The geomorphological evolution of the marginal areas of the Tibetan Plateau may provide valuable information for reconstructing the tectonic movements of the region. This study reports on a morpho-tectonic analysis of the Huang Shui catchment (tributary of the Yellow River), in the Northeastern Tibetan Plateau using a digital elevation model and field observations. One prominent bevelled surface, preliminarily interpreted as a peneplain surface, is recognized at around 2750 m altitude. It corresponds with the top of the relict sedimentary fill of large tectonic basins, and the adjacent summits. After formation of this peneplain, a terrace sequence was formed along the Huang Shui river. The transition of peneplain surface formation to incision was dated as older than 10-6 Ma using the biochronology of micromammalian assemblages from fluvial terraces and the depositional record of the basin fill. The river incision into the former peneplain is attributed to an important uplift event around 10-17 Ma.

Based on: Global and Planetary Change 2012, 88-89, 10-19.

2.1. Introduction The northeastern margin of the Tibetan Plateau, defined as the region bounded by the Kunlun fault to the south and the Altyn and the Haiyuan faults to the north (Fig. 2.1A), is thought to be the tectonically most active part of Tibet and a key region to transfer the deformation caused by the collision of India with Asia to the east by continental extrusion (e.g., Meyer et al., 1998; Tapponnier et al., 2001). Different aspects of the Cenozoic uplift of the northeastern Tibetan Plateau (NTP), have attracted wide attention over the last two decades (e.g., Tapponnier et al., 1990, 2001; Burchfiel et al., 1991; Meyer et al., 1998; Fang et al., 2005; Dai et al., 2006; Lu and Xiong, 2009; Yin, 2009). The timing and the process of deformation and uplift of the NTP are important in order to understand the mechanics of how and when the plateau has grown in time. They are, however, still under debate. Some studies indicate that the NTP undergoes a most recent uplift; the area is referred to as the „Pliocene- Quaternary Tibet‟ (Tapponnier et al., 2001). But others suggest that the NTP began to uplift much earlier, in the middle and the late Miocene (e.g., Zheng et al., 2006; Lu and Xiong, 2009) or even during the Eocene-Oligocene (e.g.,Yin, 2009). Already nearly one century ago, W.M. Davis (1911) used morphological characteristics to study the tectonic history of landscapes. His theory of cyclic landscape denudation in mountain belts involves rejuvenation of a previously beveled surface (peneplain) by tectonic uplift exceeding denudation rates. Since peneplains were believed to form near to oceanic base level, their surviving remnants were used as evidence in support of posterior vertical uplift. Although the idea that peneplains formed at oceanic base level was questioned in special cases (e.g., Babault et al., 2005; Nielsen et al., 2009), the raised peneplain theory remains the default hypothesis for many mountain ranges (e.g., Ollier and Pain, 2000; Burbank and Anderson, 2001; Garcia-Tortosa et al., 2011) and high peneplain remnants indicate a post-tectonic uplift. Thus the geomorphic

18 evolution recorded by the spatial distribution of “relict bevelled landscapes” may provide valuable information for understanding a region‟s tectonic movement (e.g., Epis and Chapin,1975; Gregory and Chase, 1994; Spotila and Sieh, 2000; Clark et al., 2004, 2006; Schoenbohm et al., 2004; Westaway, 2008, 2009; Pan et al., 2010a, b). River terraces are another, important example of morphotectonic markers (e.g., Peters and Van Balen, 2007; Srivastava et al., 2008; Vandenberghe et al., 2011).

strike-slipe fault Thrust fault (A) ult n Fa Normal fault 0 100 200 km Alty Alashan Block Ordos M. Qilia n sh n an si Block Ba Q im a ar i Qinhai H T da a m Lake B. i y B Fig.1b ua a n si n XINING Fa u lt

Kunlun Fault Siberia N .Qi nglin Fig.2.1a North g F. Ba China ya Tarim n Ha r u sh tea Qi an Pla an n South gta et a China ng Song ib Ji pan-G T Bl ns an zi oc sh Block k a F au India Indouchina lt

(B)

Q in Hu ha ang i Shui La ke Ri Xining ver

Quat. Sed. Tert. Rock ExtrusiveRock er Minhe iv R Plutonic Prec.-Mesoz . Fault Lanzhou Rock Rock w o ll 0 15 30 km e Y

Figure 2.1 (A) Simplified map of major tectonic boundaries and Tertiary faults in northeastern Tibetan Plateau (box) (modified after Dai et al., 2006; Vandenberghe et al., 2011). (B) Simplified geological map of the Huang Shui catchment in northeastern Tibetan Plateau. Stars indicate the location of the sites east of Xining (Fig. 2.6A) and east of Minhe (Fig. 2.6B) (modified from Lu et al., 2004; Vandenberghe et al., 2011).

Around the margin of the Tibetan Plateau, most studies interpret the high relict surfaces with low relief and similar altitudes as peneplain-like surfaces. Fluvial erosion has dissected the margins of these large, regionally uplifted surfaces. For example, the combination of rapid fluvial incision into bedrock and relict geomorphic surfaces has been interpreted to reflect a tectonic uplift on the southern Tibetan Plateau margin with a magnitude equal to the incision depth (Burbank et al., 1996). At the southeastern margin of the Tibetan Plateau relatively flat, highly elevated surfaces of postulated pre-uplift age were heavily dissected, creating steep fluvial valleys, and indicating the relative lowering of base levels due to tectonic uplift (Clark et al., 2004, 2005, 2006; Schoebohm et al., 2004, 2006). The existence of peneplain remnants has also been reported in NTP (e.g., Pan et al., 2003, 2010a, b; Craddock et al., 2010; Vandenberghe et al., 2011) and is further investigated in the present paper. In this study, we combine field observations and topographic data to analyze the geomorphological development of the Huang Shui catchment. The micromammalian fossils from sediments overlying the river terraces are used to provide a minimum age of the widespread beveled surfaces. Finally, the integrated results are used to infer the uplift and deformation processes at the NTP.

2.2. Geological and geomorphological background The study region is situated in the NTP (Fig. 2.1A) where coeval crustal shortening and left-lateral strike-slip faulting are related to the northeastward growth of the plateau margin as a result of the collision of the Indian and Asian plates. The crustal deformation primarily due to these relative plate motions is expressed by folding and faulting of the Paleozoic to Cenozoic bedrock. Within the NTP, also syn-orogenic basins were formed, such as the Gonghe basin (Craddock et al., 2010), Lingxia basin (Li et al., 1997), Baode basin (Pan et al., 2010b) and Xining basin (Fig. 2.1A). The basins have been filled with thick

20 continental clastic red beds, while simultaneously the surrounding highs were bevelled (as in the areas surrounding the Xining basin, see further discussion below). Subsequently, the area was incised by rivers in response to regional tectonic uplift of the Tibetan Plateau and local deformation affecting even the red basin fills, probably during the Pliocene and Pleistocene (Li, 1997; Fang et al., 2005; Gao et al., 2008; Stroeven et al., 2009; Craddock et al., 2010). The deformation in the study area was controlled by two major large sinistral strike-slip faults, the Altyn-Haiyuan Fault in the north and the Kunlun fault in the south (Fig. 2.1A; Tapponnier et al., 2001; Dai et al., 2006; Yin, 2009). The offset along these faults decreases gradually in eastward direction, which is accommodated by internal compressive deformation (Tapponnier et al., 2001; Dai et al., 2006). This deformation has resulted in a tectonic mosaic, consisting of open folds, and reverse and normal, dextral- and sinistral strike- slip faults, leading to the formation of alternating subsided and uplifted blocks of more limited extent than the large Mesozoic- Early Cenozoic basins, resulting in a fragmentation and compartmentalize- -tion of the former basins and inter-basin areas. The present-day Huang Shui river follows a series of those small-scale uplifted and subsided blocks that results from the fragmentation of the Xining and other basins. The basement of the Xining Basin consists of Proterozoic gneisses and schists, Cambrian gray limestones and green basalts and Mesozoic clastic sediments (Fig. 2.1B) (Qinghai Bureau of Geology and Mineral Resources, 1991; Dai et al., 2006). Cenozoic successions of playa to fluvio- lacustrine environments, concordantly overlying the Cretaceous clastic sediments or discordantly older basement rocks, have been subdivided into the Xining and Guide Groups. In addition, 'fanglomeratic' rocks unconformably overlie the Guide Group (Qinghai Bureau of Geology and Mineral Resources, 1991; Dai et al., 2006). The Huang Shui cut ~600 m into the Cenozoic strata and the Proterozoic-Mesozoic basement rocks, and meanwhile built a well-

21 shaped staircase of terraces in the Xining basin (Fig. 2.2) (Vandenberghe et al., 2011). The youngest deposits consist of aeolian sediments that blanket the landscape. In this paper we extend the age control of the higher terraces and their tectonomorphic implications.

2.3. Methods Large-scale landforms of the NTP were mapped using remote sensing data from various sources. The main dataset was the ASTER Global Digital Elevation Model (ASTER GDEM), with a horizontal resolution of 30 m. This data set was analysed in ArcGIS 9.3, which was also used to create shaded relief images and slope models.

Figure 2.2 Terrace sequence along a schematic section in the Xining basin (modified after Lu et al., 2004, and Vandenberghe et al., 2011). The Huzhu 1 and Huzhu 2 terraces are equivalent to terraces YS and CJB, respectively, in Vandenberghe et al., (2011).

The highest terraces along the Huang Shui valley, occurring just below the peneplain-like surface, were mapped and investigated. In the field, the sedimentary structures, lithological properties and thicknesses of the fluvial deposits were investigated. The altitude of the top of the fluvial gravel deposits has been measured using GPS with 5-10 meter resolution, taking the floodplain of the present Huang Shui as a reference level. Micromammalian fossil teeth in the sediments from the two topmost terraces at two sections, Huzhu 1 (36.72628°N, 101.82048°E) and Huzhu 2 (36.71617°N, 101.82744°E) (Fig. 2.3), are used to determine the age of the fluvial deposits and surfaces.

2.4. Results 2.4.1 The peneplain surface and the higher fluvial terraces The Huang Shui catchment is situated in the Xining basin, which is filled with c. 1500 m of fluvio-lacustrine sandstone, conglomerate, mudstone, clay and playa sediments from the Cretaceous until the middle Miocene, c. 17 Ma ago (according to Dai et al., 2006). Eastward the Xining basin gently grades into the Lanzhou basin (Fig. 2.3A). The depositional top of the basin shows a weakly undulating topography (Fig. 2.4), at a mean altitude around 2750 m (Fig. 2.3). This surface has been deeply incised (up to 600 m) by the Huang Shui and its tributary rivers (Fig. 2.3C), and it has been covered by loess. Field investigations combined with inspection of published geological- (Fig. 2.1B) and morphological maps (Fig. 2.3A) demonstrate that the surface locally cuts off the Proterozoic-Mesozoic basement (Fig. 2.6), but for its major part it is constituted by the top of the Cenozoic filling of the basins. Thus, the surface is erosional on top of the Proterozoic-Mesozoic basement, while depositional in the Cenozoic basins (Fig. 2.5). This kind of surface has also been documented at several other places (e.g., Clark et al., 2006; Pan et al., 2007, 2010a, b; Craddock et al., 2010; Fig. 2.3B). Indeed, two planation surfaces on the top of folded strata from the eastern Qilian Moutains, around one hundred km north of

23 the Huang Shui catchment (location in Fig. 2.3A), have been reported by Pan et al., (2007): an older 'main surface' and a younger 'erosion surface' (1.4 Ma). According to its age (see below), the peneplain-like surface in the Huang Shui catchment might be equivalent to the 'main surface' in the eastern Qilian Mountain. However, descriptions of relict surfaces in NTP are often vague, although they may be a quite common morphological characteristic of the region (Fig. 2.3B).

Elevation(m) 101 102 103 4988 (A) EQL") !( Menyuan 4500

37 B 4000 C Hz1 # #Hz2 3500 (! Xining

A ") !( WLZ 3000 D Minhe

GD !( ") !( Hualong !( Lanzhou 2500 Guide 36

") HL 2000 ") LX !( Linxia 1500 TD") !( Tongde

1000 0 50 100 150 KM Sections in this study City Sections studied before River

100 101 102 103 (B)

Elevation Slope 1500 5800(m) 0 84( )

( C ) A B

2 828000 0

i u

2 626000 0 h

a u

2 424000 0 H

) 2 222000 0

m (

00 1100 2200 3300 4400 5500 6600 e

d C u D

t 2800 i

t 2 8 0 0

l A

2600 i

2 6 0 0 u

2400 n

2 4 0 0 a

u H

2200 2 2 0 0

2000 2 0 0 0 0 20 40 60 80 100 0 2 0 4 0 Distance6 0 (km) 8 0 1 0 0

Figure 2.3 (A) Morphologic characteristics of the NTP derived from ASTER Global Digital Elevation Model (ASTER GDEM) data, the location of the Huzhu 1 and Huzhu 2 sections in the Xining basin, and the studied sections in other basins. EQL-easter Qilian basin; WLZ-western Lanzhou basin; LX- Linxia basin; HL-Hualong basin; TD-Tongde basin. GD-Guide basin. (B) The relict surfaces separated by mountain ranges in the NTP. The map displays a semitransparent grey-scale slope image draped over a brown DEM. The relict surfaces with low relief have a bright and“hazy” appearance whereas steep slopes have a darker and clearer hue in mountain ranges. (C) Topography along A-B and C-D sections, derived from ASTER GDEM data illustrating the character of the peneplain surface (equivalent with the top surface of the relict basin fill; dashed line) in the Huang Shui catchment.

Below this relict surface, the Huang Shui valley has a well-developed fluvial terrace staircase (Fig. 2.3; Vandenberghe et al., 2001; Lu et al., 2004). Around Xining, a few high terrace-like flat plains (Fig. 2.2, 2.5, 2.7, 2.8a) occur

25 at altitudes of 566 m, 596 m and 616 m above the present water level of the Huang Shui river, which is tens of meters below the peneplain surface (Fig. 2.9). Below, we describe and interpret the sediments underlying these terrace- like flat plains at the Huzhu 1 and Huzhu 2 sections, which are equal to the YS and CJB terraces, respectively, in Vandenberghe et al.,( 2011). The gravel layers below the peneplain remnants mainly consist of quartzite, quartz and sandstone pebbles. Small amounts of conglomerate, and crystalline and metamorphic pebbles also occur. The mean diameter of the pebbles is in the range of 60-100 mm. The pebbles are mainly rounded to sub- rounded and strongly weathered. The 5-6 m thick gravel layers are generally massive at the base (Gm facies), grading towards the top into finer-grained, coarsely planar bedding (Gp). Imbrication and cross-bedding are typical sedimentary structures (Fig. 2.8B). These gravels are topped by a clayey deposit and loess (Fig. 2.8C). These gravel deposits could have two genetic origins: (1) 'fanglomeratic' fan deposits like the Linxia Group, (2) fluvial terrace sediments. We prefer the latter interpretation because: (1) The roundness of the gravels indicates fluvial dynamics rather than debris flow deposition; (2) The massive gravel, grading upward to finer-grained planar bedding with imbrication and cross-bedding, and then to medium to fine sand represents an evolution from channel bedload, via lateral and longitudinal bars, to the waning flow phase, which altogether represents a complete fluvial sequence rather than the intermixed structure of a debris flow fan deposit; (3) The so-called Linxia Group 'fanglomeratic' deposits in the Xining basin are very scarce . In general, these Linxia Group 'fanglomeratic' deposits seem to occur only in limited areas near to the foot of the Laji Shan and Datong Shan (according to Dai et al., 2006) and at the footwall of the uplifted mountain ranges in the Linxia and Guide basins in the NTP according to Fang et al. (2003, 2005). Our studied sites are near to the center of the basin, far from the foots of the surrounding mountains, and at least several to tens of km from the nearest

26 possible 'fanglomeratic' deposits indicated by Dai et al., (2006); (4) The areas around our sites are relatively flat (Fig. 2.7), resembling terrace surfaces. Thus, the fanglomeratic origin may be excluded, and it is more logic that these gravel and sand deposits are parts of a fluvial terrace system. Due to the topographical position of these gravel layers, which is slightly below the peneplain remnants (several tens of m), the associated terraces are slightly younger than that peneplain. The gravel characteristics and sedimentary structures point to fluvial erosion in wide and shallow valleys after the planation of the relief. The fluvial deposits of the terraces are covered with more than 70m of fine-grained sediments. The upper c. 40 m consist of loosely packed, Quaternary loess-paleosoil sequences, while the lower c. 30 m consist of Tertiary Red Clay, characterized by a finer grain size and stronger compaction than the overlying loess (Fig. 2.8C, 2.9). There is a clear erosion surface between the loess and Red Clay. Grain-size distribution patterns and geochemical composition of the Red Clay are similar to those of the aeolian loess, but different from those of fluvial and lake deposits (Lu et al., 2001; Guo et al., 2002) indicating their wind-blown origin at the Huzhu 1 and Huzhu 2 sections (Wang et al., 2006). Since these terraces are the two highest ones, their age could be only slightly younger than the peneplain surface. See Fig. 2.9 for a complete stratigraphical description of these sedimentary series. 2.4.2. Micromammalian assemblage from the terrace at Huzhu 2 More than 50 micromammalian fossil teeth at 6 levels in the Red Clay from the terrace at the Huzhu 2 section were sampled and identified as 6 species (Table 2.1, Fig. 2.9). These fossils are characteristic of late Miocene- early Pliocene time (Zheng, 1994; Qiu and Qiu, 1995; Flynn et al., 1997; Deng, 2006). Especially, Prosiphneus licenti from 11.5m, 21.1m, and between 22.5 and 23 m is correlative to MN12 (Zheng,1994; Zheng et al., 2004) in the European Neogene Mammal Chronology (Mein, 1990). The recently reported

27 magnetostratigraphy and biostratigraphy of the Miocene-Pliocene (since 22 Ma) loess in Northwest China also shows that the distribution of Prosiphneus licenti dates from the late Miocene, 8-6 Ma ago (Guo et al., 2002; Hao et al., 2007; Zhang and Liu, 2005). In addition, according to the biostratigraphic and magnetostratigraphic correlation, the sediment sequence from Lingtai (Gansu in North China) with Nannocricetus mongolicus and Pseudomeriones abbreviatus covers the late Miocene to Early Pliocene (Zheng and Zhang, 2001). The latter species were also discovered in other typical late Miocene-Pliocene fauna such as the Eremte fauna and Bilike fauna from Nei Mongo, and the Baode fauna and Gaozhuang fauna from Shanxi (e.g., Qiu and Qiu, 1995; Zheng and Zhang, 2001; Zhang and Liu, 2005). The Hansdebruijnia sp. has been discovered in the Baode fau-na, which is very similar to Hansdebruijnia pusillus from Ertemte (Storch, 1987; Storch and Ni, 2002). Thus, the micromammalian fossil assemblage from the Red Clay covering the terraces Huzhu 1 and Huzhu 2 probably has a Baodean age (late Miocene) corresponding with the Turolian of the European land mammal stratigraphy (Deng, 2006).

Figure 2.4 The weakly undulating topography in the Xining basin as an example of the peneplain-like surface.

2.5. Discussion 2.5.1 Age of the terraces and the peneplain surface

The provenance areas of the aeolian deposits on the northeastern Qinghai- Tibetan Plateau are the middle and western parts of the Tibetan Plateau and the dry area of inland Asia (Lu et al., 2004; Wang et al., 2006; Nugteren and Vandenberghe, 2004; Vriend et al., 2011). Investigations indicate that these regions were dry and produced continuously dust at least since early Miocene times (Guo et al., 2002; Vandenberghe et al., 2004). Thus, it is reasonable to assume that as soon as terraces were formed after floodplain abandonment, the dry terrace surfaces started to accumulate dust (Porter et al., 1992; Pan et al., 2003; Lu et al., 2004; Sun, 2005). The micromammalian assemblage in the aeolian Red Clay from the terrace at the Huzhu 2 section shows that the aeolian sequence was deposited since the late Miocene. By extrapolation, the higher terrace at Huzhu 1 could be slightly older, corroborating our previously published preliminary opinion (Lu et al., 2004). According to the geomorphological evolution in the Huang Shui catchment (Fig. 2.5), the peneplain surface perching tens of meters above these terraces (Fig. 2.9) is thus older than these terraces, i.e. late Miocene (10 - 6 Ma), but younger than the final filling of the tectonic basins, that is around 17Ma according to Dai et al. (2006). Erosion surfaces have served as important markers of tectonic uplift and deformation (e.g., Epis and Chapin, 1975; Gregory and Chase, 1994; Spotila and Sieh, 2000; Clark et al., 2004). It has been shown that, for instance, European and African, continent-scale land masses were stripped off by surface erosion leading to the development of wide peneplain-like morphologies up to Pliocene times, while river incision and terrace formation started somewhere at the end of the Pliocene or the beginning of the Pleistocene as a result of intensified tectonic uplift (e.g., Van den Berg, 1996; Bridgland and Westaway, 2008; Gibbard and Lewin, 2009). In addition, intensified tectonic uplift around 2 Ma has been held responsible for a rapid increase of the fluvial incision rate

29 near Xining that coincided with a drastic change in drainage direction in the same region (Vandenberghe et al., 2011).

II (ⅰ)

(ⅱ)

III

basement rock basin fill aeolian deposit erosion

accumulation peneplain-like surface river terrace staircase Figure 2.5 Schematic diagram illustrating the geomorphological evolution of the Xining basin and the formation of the relict surface. Firstly, the tectonic basin was formed before the Mesozoic (I). Then, the surrounding mountains eroded and the basin was filled (II, i). As a last stage in this step, limited and very local erosion in the basin might have slightly lowered the surface formed in II (i), ultimately resulting in the formation of a peneplaine surface (II, ii). Finally, tectonic uplift caused the Huang Shui to incise vertically into the surface, producing a terrace sequence (III).

2.5.2. Uplift of the Huang Shui catchment at the NE Tibetan Plateau prior to the late Miocene Theoretically, land degradation without tectonic uplift could have been induced by lowering of the local base level due, for instance, to headward erosion and catchment capture by an adjacent, topographically lower river (as described by Craddock et al. (2010) and Pan et al. (2011) in adjacent regions). However, in the Huang Shui catchment, the river incised from a peneplain-like surface without the indication of a (migrating) knickpoint or past river capture. Thus, the landscape transition from large-basin filling and peneplanation to fluvial incision probably indicates an intensive uplift starting just before the late Miocene (10-6Ma), but after the period of final basin filling around 17 Ma (Fig. 2.3A) (Qiu et al., 2001; Dai et al., 2006). During the same period, erosion surfaces at the SE Tibetan plateau have also been entrenched as a consequence of increased uplift (Westaway, 2009), indicating a nearly synchronous uplift event at large regional scale (at least NE and SE Tibetan Plateau) during the late Miocene. Thus, the opinion that the Tibetan plateau uplift process transfers from southwest gradually to the north and east step by step with formation of the northeastern and southeastern parts of the plateau until the Pliocene- Quaternary (Tapponnier et al., 2001) may be questioned. Similar studies in regions near to the Xining basin in the NTP show that rivers started entrenching from a bevelled (erosional) surface while landscape morphology changed from basin filling to denudation at ages of c. 3.7 Ma and 1.8 Ma along the Yellow River (according to Pan et al. (2010a, b) and Craddock et al. (2010) respectively), which is much later than the similar landscape transition period in the Huang Shui catchment studied here. This discrepancy may be explained in several ways. First, the older landscape records at the other sites might have been eroded. Secondly, the relict bevelled surfaces may have experienced a complex history with different ages of devel-

()A ()B

erosional boundary loess Cenozoic deposit Precambrian-Mesozoic basement Figure 2.6 Examples of the undulating beveled surface (below aeolian deposits) on top of the Cenozoic fill east of Xining (A) and Precambrian- Mesozoic basement east of Minhe (B).

Figure 2.7 Photo showing the flat plains at 500-600 meters above the Huang Shui water level, approximately 1.5 km southeast of the Huzhu 1 section.

(A) ( C )

(B)

Figure 2.8 (A) Overview of section and topography of the Huzhu 2 terrace, looking north. (B) Lower parts of the fluvial gravel layer at the Huzhu 2 setion. (C) Lithology of the Red Clay deposit with yellow-brown clay layers interbedded with reddish, strongly developed soil layers at Huzhu 2 section. The Red Clay deposit lies between the fluvial sediments and Quaternary loess.

Peneplain surface 2750

Huzhu 2 Huzhu 1

2730 Mean size (μm) 5.500 1100.00 1155.00 2020.00

Mean size (μm) CJB(um) 5 5 1010 1515 2020 E

2710 o

(

)

2690

sandstone gravel sand

Red Clay loess fossil position 2670

Figure 2.9 Stratigraphy of the two topmost terraces with the fossil position and grain size of the Red Clay at the Huzhu 1 and Huzhu 2 sections.

-opment in different areas (Clark et al., 2006). The late entrenching of the relict surface reported by the former authors could have been caused by headward erosion of rivers flowing at a lower level and not by a river properly flowing on the peneplain and the related basin as the Huang Shui in the Xining basin (Fig. 2.5). Thus a certain lag is possible because of time needed by the headward erosion. In our opinion, both the infilling of the large tectonic basins and the peneplain formation should have occurred largely in the same period on the NTP. But, the effective duration of these different processes, and thus their time of termination, may have been different at different locations. Possibly the younger phases of uplift in the NTP, as reported by the former authors, may correspond with the period of intensified uplift at 2Ma found by Vandenberghe et al. (2011; Fig. 2.2). More research is needed to clarify this hypothesis. Also, the hypothesis put forward by Craddock et al (2010) of a lag between formation of the bevelled surface and river entrenchment requires further investigation.

Depth (m) Fossils

11.5 Prosiphneus licenti 12.7 Pseudomeriones abberviatus

15.0 Pseudomeriones abberviatus

20.5 Pseudomeriones abberviatus

21.1 Prosiphneus licenti, Nannocricetus mongolicus, Dipus cf. D. fraudatou

23.0 Prosiphneus licenti, Nannocricetus mongolicus, Pseudomeriones abberviatus, Hansdebruijnia sp.

Table 2.1 Micromammalian Fossils from the Red Clay deposits at the Huzhu 2 Section.

2.6. Conclusions Micromammalian fossils show that the aeolian deposit sequences overlying the high fluvial terraces along Huang Shui river valley were deposited during the late Miocene (10-6 Ma). Thus, the peneplain-like surface a few tens of meters above these terraces and the related landscape transition from basin filling to incision happened at least before the late Miocene. This indicates that an intense uplift event with morphological significance happened around 10-17 Ma in the NTP.

Acknowledgements This research was supported by the National Natural Science Foundation of China (Grant no. 40901002, 40930103 and 41021002), and the Dutch- Chinese CAS-KNAW PhD project (05-PhD-10) and Exchange Programs (08CDP028, 09CDP033 and 10CDP036). Thanks are extended to Prof. Z. T. Guo, Dr. J.Y., Ge and Prof. M. J. Dekkers for helpful discussions. The comments by two anonymous reviewers are highly appreciated.

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-Chapter 3- Differential impact of small-scaled tectonic movements on fluvial morphology and sedimentology (the Huang Shui catchment, NE Tibet Plateau) Jef Vandenberghe, Xianyan Wang and Huayu Lu Abstracts : The aim of this paper is to investigate the morphological implications of the interplay between tectonic movements at different rates, timescales, and spatial extent in a catchment of intermediate size, the Huang Shui River (a main tributary of the Yellow River in the NE Tibet Plateau). River incision started from a peneplain-like surface that developed in post- Miocene-Pliocene times. At the end of the Tertiary, a general tectonic uplift of the Tibetan Plateau initiated river incision. This general uplift also caused local fragmentation of the Huang Shui catchment into blocks (kilometers or maximally, a few tens of kilometers) of local extent that have subsided and/or uplifted relative to one another. Fluvial deposition of >30 m in the subsiding blocks contrasts with erosion and formation of gorges in the uplifted blocks. Incision into the former peneplain was not continuous but a staircase of terraces developed under climatic influences. Terrace deposits are sometimes capped by interglacial soils or soil-derived material, Apparently, terrace incision occurred at the transition to the next cold period. The different rates of uplift and subsidence of the individual blocks resulted in the simultaneous development of erosion and accumulation terraces of different sizes within the same catchment, even within the same tectonic block. This makes it impossible to connect the terraces of the different blocks, except for the three youngest terraces (representing the last 300,000 years), thus illustrating the uniform tectonic history of the catchment since that time.

Based on: Geomorphology 2011, 3-4, 171-185.

3.1. Introduction River systems in uplifted regions are usually characterized by the development of a series of terraces where erosion is dominant; albeit the traction load or the „sediment in transport‟ within the channels (up to several meters thick) may be preserved (they are also referred as strath terraces in this paper). Such terrace sequences represent repeated stability-incision cyclicity which is a response to climate changes. The resulting staircase morphology of river systems has traditionally been attributed to cold (glacial)-warm (interglacial) cycles (e.g., Büdel, 1977; Van den Berg, 1996; Starkel, 2003). Their global occurrence has recently been investigated in International Geological Correlation Projects 449 and 518 (Bridgland and Westaway, 2008b). The original concept was modified by several authors (Vandenberghe, 1995; Bridgland and Allen, 1996; Maddy, 1997; Vandenberghe, 2008) and confirmed by field data at many locations (e.g., Antoine et al., 2000; Mol et al., 2000; Cordier, 2006). In contrast, terraces are generally not formed in gradually subsiding regions. Instead, sediment accumulation continues to fill such tectonic basins; and complex, overlapped to onlapped terraces continued to develop. In comparison with the interruption of incision by tectonic stability in uplifted regions, accumulation may also be interrupted, giving rise to soil formation in subsiding regions. In both cases, alternations of deposition (respectively incision) and soil formation in subsiding (respectively uplifted) regions may also be triggered climatically. Based on the assumption that terrace surfaces record net incision driven by tectonic uplift, vertical separation and longitudinal profiles of terrace surfaces have been used to infer tectonic uplift rates (e.g., Van Balen et al., 2000; Maddy et al., 2001; Van den Berg and Van Hoof, 2001; Peters and Van Balen, 2007). Tectonic uplift patterns and -rates may be inferred, using the along- course changing vertical separation of terraces as well as their longitudinal

45 profiles (e.g., Maddy, 1997; Van Balen et al., 2000; Peters and Van Balen, 2007) on condition that eustatic influence is absent and rivers reached similar longitudinal gradients after each uplift phase (see discussion by Kiden and Tornqvist, 1998). Likewise, the thickness of the sedimentary fill may provide qualitatively an idea of the amount of tectonic subsidence (e.g., Amorosi et al., 1999; Nádor et al., 2003). Usually, one of both tectonic scenarios (uplift or subsidence) has prevailed in any particular region, so that river catchment evolution has been progressive in one or the other direction (terrace formation or basin fill, respectively). In such simple tectonic settings, reconstructions of tectonic impact on fluvial systems are relatively easy. In particular tectonic settings such as that of the Huang Shui, however, both tectonic scenarios alternate at relatively short distances. A river catchment occurring in such a complex setting may thus show morphological and sedimentological characteristics that accordingly change over a short distance. Furthermore, we may expect that in such a setting the individual tectonic blocks may each have their own tectonic history and thus have their individual uplift or subsidence rates. As a result, terrace correlation (for instance) between different tectonic blocks may be problematic or impossible, even within a single catchment. The present paper aims to investigate fluvial incision and sedimentation and terrace morphology in a complex setting of different tectonic movements within one catchment, the Huang Shui, a tributary of intermediate size of the Yellow River (Huang He) (Fig. 3.1). Tectonic impact on river terrace formation is especially well understood in a relatively simple and modest tectonic context (Leeder, 1993; Burbank and Anderson, 2001). The question arises about the effects of tectonics in a relatively complex setting as drivers for the morphological development and the sedimentary architecture as, for instance, also occurs in SE Spain (Stokes and Mather, 2003). The study region provides a framework of regional tectonic uplift overprinted by tectonic movements of

46 more local extent and changing over time. It is the aim to study the reaction of the fluvial system to such a spatially and temporal dynamic forcing. Until now the impact of complicated differential tectonic motions on fluvial development is unknown. The Huang Shui catchment is a study area where, due to the active transpressive tectonic setting, the motions are highly variable along the river‟s course. In addition, there is the overarching question of how climate and tecton-

Siberia

F ig.4.1a North China Tarim au te la n P South ta be China lt (A) Ti a u yn F Alt Alashan Block Or d o s India M Indouchina .Q i 38 lian shan Blo ck Q a H id Qinhai a a i y m u B Lake B. XINING a n a Fig.4.1b 36 si F n a u Kunlun Fault l t

Strike-slipe fault Thrust fault 34 Normal fault 0 100 200 km

86 90 94 98 102 106

101 102 103 (B)

Q in Huang ha shui i L ak Riv e er L Xining ao ya Pingan Ledu Minhe r Quat. Rock Tert. Rock Infracrustal Rock ve Ri

w Lanzhou Effusive Prec.-Meso . Fault. llo Rock Rock e Y 36 101 102 103 Figure 3.1 (A) Simplified tectonic map of the Tibetan Plateau including major Cenozoic faults (modified after Dai et al., 2006). (B) Simplified geological map of the Huang Shui catchment in northeastern Tibet, illustrating the locally subsided blocks (consisting of Tertiary and Quaternary rocks) separated by upthrusted blocks (outcropping Precambrian to Mesozoic rocks). Stars indicate the location of the sections at Xining (Fig. 3. 3) and LaoYa-Minhe (Fig. 3.7).

-ic movement are expressed individually as external drivers in fluvial development. Specifically in the Huang Shui catchment, it is intended to disentangle tectonic and climatic forcing (in a monsoonal setting) at different temporal and spatial scales (cf. Olszak, 2011). The combination of relatively well dated and well pronounced differential vertical motions under alternating paleoclimatic conditions is unique, and makes this river system ideal for studying the impact of complex tectonics on river development. Sea level changes took place at a much too large distance to be influential for fluvial development. The results in this paper are part of an ongoing project and therefore are preliminary. They will be supplemented in future by more absolute ages of the terraces, whereas this paper focuses on the relation between a complex setting of tectonic movements and fluvial morphology.

3.2. Tectonic setting of the study region and consequent general morphological expression The tectonic evolution of the Plateau at a large geological timescale is constrained by diverse datasets (see references below), but the more recent Plio-Pleistocene evolution is less well known, and can be addressed by using fluvial terraces, such as those of the Huang Shui river. The study region is situated in the northeastern corner of the Tibetan Plateau and is characterized by the occurrence of folded and faulted bedrock of Proterozoic to Cenozoic age. Large tectonic basins formed during the Paleogene as a response to early India- Asia collision probably between 55 and 52.5 Ma, for instance the Xining and Lanzhou basins, which both are sub-basins of the Longzhong basin (Horton et al., 2004; Dai et al., 2006; Dupont-Nivet et al., 2008). They are similar to other syn-orogenic basins exist in northeast of the Tibetan Plateau, such as the Gonghe basin (Craddock et al., 2010), Lingxia basin (Li, 1991) and Baode basin (Pan et al., 2010, 2011). The Huang Shui catchment is situated within the

Xining subbasin. The large tectonic basins are syntectonically filled with continental, often folded and faulted clastic red beds of upper to middle Miocene age (Dai et al., 2006; Dupont-Nivet et al., 2008). Other basins were completely filled only at a later stage (e.g., Fang et al., 2005). In general, the fill deposits are of lacustrine, fluvio-lacustrine, fluvial and aeolian origin (e.g., Lu et al., 2001; Guo et al., 2002; Vandenberghe et al., 2004; Alonso-Zarza et al., 2009). Since the Miocene, the tectonic deformation in the study area has mainly been controlled by two major large sinistral strike-slip faults, the Altun Tagh- Haiyuan fault in the north and the Kunlun fault in the south (Fig. 3.1A; Tapponnier et al., 2001; Dai et al., 2006). The northeastern part of the Tibetan Plateau, defined as the region bounded by these two faults, is thought to be the most tectonically active part of Tibet and a key region to transfer the deformation caused by the collision of India with Asia to the east by continental extrusion. The lateral offset along these faults decreases gradually in an eastward direction, which is accommodated by internal compressive deformation. In the study region coeval crustal shortening and left-lateral strike-slip faulting are related to the northeastward growth of the plateau margin. Tectonically, deformation between these two major fault zones has resulted in a tectonic mosaic, consisting of open folds and of reverse and normal, dextral- and sinistral strike-slip faults, leading to the formation of alternating subsided and uplifted blocks of more limited extent than the Mesozoic-Early Cenozoic basins, resulting in a tectonic fragmentation and compartmentalization of the former basins and inter-basin areas. At the same time, the study area was subjected to a general, large-scale uplift of the Tibetan Plateau. After completion of the filling of the large tectonic basins, the landscape was planed off by erosion and, apparently, ended up as a composite of beveled (erosional) surfaces of peneplain or pediplain type (Li, 1991; Gao et al., 2008;

Stroeven et al., 2009). After the development of these beveled surfaces the entire region has been affected by accelerated general uplift simultaneously with continued regional deformation (Li, 1991; Horton et al., 2004). Subsequently, rivers started entrenching from about 14 Ma onward in the Huang Shui catchment (according to Lu et al., 2004; Wang, 2008) as described below and from 4.9, 3.7 and 1.8 Ma along the Huang He (according to Craddock et al., 2010; Pan et al., 2010, 2011, respectively). From data of the Huang Shui catchment and from the eastern Tibetan Plateau, presented by Clark et al., (2006), apparently the relict planation surfaces have experienced a complex history with different ages of development. Against a background of general uplift of the Tibetan Plateau and accompanied formation of large tectonic basins, tectonic motions of local extent have segmented the NE part of the Plateau and in particular, the Huang Shui catchment into small uplifted areas and elongated stretches that show relative subsidence. We can easily recognize this tectonic fragmentation on the geological map (Fig. 3.1B), which has resulted in the spatially small-scaled alternation of subsided blocks, filled with siliciclastic deposits (as in the large tectonic basins: see above), and in upthrusted, metamorphic (Proterozoic and Palaeozoic) blocks comparable with those surrounding the large tectonic basins. The consequent morphological expression in the Huang Shui catchment consists of narrow gorges with erosional terraces in the uplifted blocks, alternating with wide depressions in the relatively subsided block characterized by cut-and-fill terraces.

3.3. Methods The tectonic blocks have been mapped using satellite imagery (GOES Satellite Data-Imager (GVAR_IMG) from NOAA) and verified in the field. Individual terraces have been identified and their regional distribution specified over the various tectonic units. A program of selected field mapping was

50 carried out by GPS three-dimensional positioning after analyzing 1:50,000 topographic maps. As an example, Fig. 3.2 demonstrates the existence of many flat surfaces at different altitudes above the present floodplain. The primary genesis of those surfaces is that of terraces of the Huang Shui River as they are underlain by widely occurring fluvial beds, containing gravel from that river. However, determination of the elevation of the terrace surface (by definition, the morphological top of the fluvial deposits) required in all cases a field measurement as that fluvial deposit was always covered by varying thicknesses of aeolian deposits. Posterior erosion of the uppermost fluvial deposits was never observed (see below), thus not obstructing terrace correlation. A number of typical transects were selected from different tectonic blocks and are discussed below in terms of their morphological and sedimentological properties. The individual outcropping fluvial deposits were described with regard to their sedimentary structures (using lithofacies according to Miall, 1978), lithological properties, and thickness as far as they were accessible in order to derive the sedimentary processes. The altitude of the top of the gravel deposits has been used for correlation, as the base of the gravels was often not visible, taking the floodplain of the present Huang Shui as a reference level. Absolute dating of the terrace series remains a persistent problem. Because of their age ranges, optically stimulated luminescence (OSL) and C14 methods could only date the youngest terraces. In addition, the fluvial deposits only rarely contain any organic material for radiocarbon dating, while they are often too coarse-grained for luminescence dating. In contrast, the overlying loess gives better possibilities, especially for the ages of the older terraces. This procedure assumes that there was no time hiatus between fluvial and aeolian deposition. It appears that the source regions of aeolian dust in inland Asia were very dry and had produced continuously dust at least since early Miocene times even during interglacials (Pan et al., 2003; Lu et al., 2004; Vandenberghe

51 et al., 2004; Vriend, 2007; Wang, 2008). Thus, it is reasonable to assume that as soon as terraces were formed after floodplain abandonment, the dry terrace surfaces started to accumulate dust (Porter et al., 1992; Pan et al., 2003; Lu et al., 2004; Sun, 2005). In addition, at several occasions a gradual transition could be observed between fine-grained, silty-sandy floodplain deposits and overlying, slightly sandy aeolian silt. The age of the loess may be determined by OSL, palaeosol stratigraphical correlation and palaeomagnetic dating (e.g., Pan et al., 2003; Zhang et al., 2009, for which often marine isotope terminology (MIS or marine isotope stages) is applied). All dates are, however, „ante quem‟. Absolute dating of the deposits for fluvial research in the study region is still in a premature stage, but we discuss available data where possible and useful.

0 0 4 2

23 00

22 00

hu an gs hu i

riv Loess er Yunjiakou Gravel 2300 1 km Red base 0

Fiure 3.2 Example of geomorphological terrace mapping in the Huang Shui valley near Yunjiakou village slightly east of Xining.

3.4. Results A selection is made from the tectonic units in the Huang Shui catchment

52 that are most illustrative (composite) to describe their morphological development as a function of their specific tectonic movements. For each tectonic unit a schematic profile of terraces is compiled or summarized. The different terraces show similar rock and sediment sequences: Cenozoic fluviolacustrine, clastic red beds in the depressions and Paleozoic metamorphic rocks in the gorges are covered by fluvial gravels of various thicknesses, intermingled with lenses of sands, silts, and clays and finally covered with aeolian loess at the top. The gravels mainly consist of quartzite, quartz, and sandstone, with small numbers of conglomerate and crystalline and metamorphic pebbles; they are mainly rounded to subrounded and strongly weathered. The gravel layers are generally massive at the base (Gm facies), while otherwise planar bedding is dominant (Gp facies), frequently including imbrication and cross-bedding as a result of lateral and longitudinal bar development and, occasionally, typical pothole fills. 3.4.1. Results 1: The terrace staircase near Xining The town of Xining occurs at the bottom of a wide depression (c. 2240 m altitude); the topmost planation level (peneplain) surrounding this depression is at some 2950 to 3100 m altitude (Fig. 3.1). The depression at Xining is oval- shaped, c. 55 kmwide and c. 42 kmlong. Its substratum consists of reddish Neogene fluviolacustrine mudstone, siltstone, and sandstone and occasionally intercalated with saline beds. The Huang Shui River, following the W-E trending axis of that depression, has built a well-formed staircase of terraces. This is in contrast with the upstream and downstream areas of the depression where the Huang Shui has cut a gorge through resistant Precambrian to Mesozoic bedrock, within which (because of the constraint on lateral migration caused by the hardness of these rocks) it has generally not been able to form terraces. The depression is a local downwardly mobile tectonic block within the otherwise uplifted Tibetan Plateau. The areas in which the gorges have been formed may be considered as local upthrust blocks. From this tectonic setting

53 of general uplift and local subsidence, it follows that the Huang Shui River was able to form erosional but rather wide terraces, especially on its northern side. On the other hand, these terraces had a considerable preservation potential, as they were developed in a wide open landscape. As a result, a large series of terraces could be mapped; they are exemplified in Fig. 3.2, while the complete sequence is represented schematically in Fig. 3.3. The Huang Shui system is particularly characterized by its long history of river erosion and terrace formation when compared with most terrace sequences elsewhere. The oldest terraces occur at about 500 to 600 m above the present floodplain. A series of 15 to 20 younger terraces may be distinguished at lower levels. However, it is possible that the uppermost terraces belonged to a proto-Huang Shui that was draining westward (opposite to the present eastward direction) to an intramountainous tectonic basin (see below). Three low terraces are distinguished (T1a, T1b, and T1c), respectively, at 1 to 1.5, 7-8, and 11-13 m above the present floodplain level (=apf). Their ages can only be estimated in general terms at this moment. Terraces 1b and 1c are covered by loess, while terrace 1a is the only terrace without any loess mantle. The two lowermost terraces (T1a-T1b) have a relatively thin gravel thickness (1-2 m) and are covered by a thin alluvial loam (0.5-1.5 m). They are widespread along the present Huang Shui and are well expressed in the morphology. Terrace T1c has a slightly greater gravel thickness (3 m) and is covered by a thick cover of displaced or undisturbed loess. The top of the gravel bed of the lowermost terrace (T1a) is at about 1-1.5 m apf, while the top of the gravels of the present-day floodplain occurs at about 1 m below the present floodplain. Therefore, terrace 1a is about 2-2.5 m above the present river deposits and thus cannot be considered as part of the present floodplain, although present-day peak discharges cause flooding of that terrace (with deposition of fine alluvium) every 1 or 2 years. In addition, the gravels of T1a are dissected in many places by the present river. A Holocene soil is

54 formed at the top of the gravel layer. The latter soil excludes that this terrace would be part of the present-day floodplain and its dissection should only be due to intrinsic rive dynamics. On the other hand, the terrace is not covered by loess, in contrast to all other terraces, and therefore formed contemporaneously with the last main loess depositional phase (MIS 4-2 or last glacial). This means that the Huang Shui incised slightly into that terrace during the Lateglacial-Holocene.

T(HW) T(CJB) T(YS) T(PZS) T(DDL) 660

600

540

480 u

2 . 0 M a d

o 420 v

1 . 4 M a

360 -

T8 a

300 d

T7 a

(

240 ) T6

T5 180 T4 T3 120

T 2 a 60 T 1 c T 1 b T 1 a Gravel Loess Red-earthy deposit 0

Figure 3.3 Terrace sequence along a schematic section in the Xining basin (location in Fig. 3.1B).

The top of the gravel bed of terrace 1b is at some 7 m apf and is covered by about 2 m of loess. Regarding the age of the terrace, only one TL-date from overlying loess in a certain position is known (17.3±1.5 ka according to Lu et

55 al., 2004) and no interglacial soil (S1) has been observed. Despite the scarcity of age indications at present and the presumed large hiatus between the TL date of 17.3 ka and the terrace gravel that can be supposed to have been deposited under cold climate conditions (see below), we provisionally attribute a penultimate glacial age (corresponding with MIS 6) to terrace 1b. The age of the third terrace (T1c) has been inferred from the litho- pedostratigraphy of the overlying loess (Vandenberghe et al., 2006) and OSL dates at Xining TXD. At that site, the 27-m-thick last-glacial (Malan) loess (20 OSL-dates between 14 and 60 ka according to Buylaert et al., 2008) is underlain by a well-developed soil complex of last-interglacial age (S1) (Chen et al., 1997; Vandenberghe et al., 2006; Vriend, 2007), which overlies about 60 m of poorly exposed loess. Taking into account the 60 m of loess below S1, we assume that this c. 60 m of loess represents at least one glacial period. Therefore, we estimate the age of the terrace deposits of T1c as equivalent to MIS 8. An MIS 7 age cannot be excluded, but interglacial terrace formation is less probable as will be discussed below. This dating strengthens the assumed age of MIS 6 of the next younger terrace T1b. The next oldest terrace (T2a), at 26 m apf, appears to be different, as the gravel thickness is c. 10 m. In explanation of this relatively large thickness, in comparison with the max. 3m thick gravel deposits of the younger terraces, tectonic subsidence of the basin is inferred; a climatic origin would require the occurrence of such a terrace over the entire length of the river which is not the case (further discussion see below). The next series of older terraces (T3a-b-c, T4, T5, and T6a-b) occurs at considerably higher elevations above the present- day floodplain (96, 118, 140, 160, 196, 240, and 264 m, respectively) (Fig. 3.3; Table 3.1). Despite the large steps in elevations between them, the individual terraces can readily be recognized over large distances within the Xining basin (Fig. 3.2). With gravel thickness of 2-6 m, they fit well in the general pattern of the predominantly erosional terrace staircase in this region. All of those terraces

56 are covered by an incomplete series of loess that, unfortunately, obstructs palaeomagnetic and soil-stratigraphic dating of the terraces. These loesses are also too old for OSL dating. In contrast, the 232- and 182-m-thick loess-palaeosol layers on top of terraces T7 and T9, respectively have been preserved completely and been drilled at the sites of Dadungling (DDL; 36.657° N., 101.787° E.) and Panzhishan (PZS; 36.649° N., 101.844° E.), just north of the town of Xining (location Fig. 3.1B). The fluvial gravels occur at about 299 and416 m apf, respectively, and are only a few meters thick. A combination of loess-palaeosol stratigraphy and palaeomagnetic analysis enabled Lu et al. (2007) to date the base of the loess series at 2.0 and 1.4 Ma at the PZS and DDL sites, respectively. However, the latter date is questionable because of reworking of the lower part of the loess, although the Jaramillo event is detected with certainty and thus the basal loess is certainly older than 1 Ma. Rates of river entrenchment reflect the general morphological development and may help to infer the associated speed of tectonic uplift. Time-averaged incision rate between the Yanshan and T9 (PZS) terraces was c. 17 m/Ma (about 200 m in 12 Ma) or c. 34 m/Ma (assuming the Yanshan terrace to be 8 Ma instead of 14 Ma), while between T9 and now it was c. 195 m/Ma (about 390 m in 2 Ma). This phase of obviously accelerated uplift between 1.4 and 2 Ma initiated the development of the terrace staircase of mostly strath terraces of the modern Huang Shui since T7. Noteworthy is that an „erosion surface‟ from where river incision started is described from the eastern Qilian Shan, a few hundred kilometers north of the Huang Shui catchment, with an age of 1.4 Ma by Pan et al. (2007). That erosion surface is situated c. 650 m lower than an undated „main surface‟ which is also a planation surface on top of folded Neogene strata (Pan et al., 2007). The latter „main surface‟ might be equivalent to the „planation surface‟ in the Huang Shui region which is about 350 to 450 m above T7. Furthermore, the start of river incision mentioned by

Craddock et al., (2010) in the nearby Gonghe basin might correspond with the same initiation o accelerated fluvial incision in the Huang Shui.

Xining depression Ledu depression Laoya gorge Minhe depression Terrace Height Terrace Height Terrace Height Terrace Height Apf Apf Apf Apf(m (m) (m) (m) ) T1a 1-2 T1a 2 T1a 2 T1a 2 T1b 7.5 T1b 7.5 T1b 7 T1b 7-9 T1c 11-13 T1c 13 T1c 15 T1c 11-14 T1d 20 T1d 22 T1d 22 T2a 26 T2b" 35 T2c 40 T2c" 44 T2d" 54 T2e" 60 T3a 96 T3a' 100 T3a" 97- 100 T3b 118 T3b' 124 T3c 140 T3c' 144 T4a 160 T4a 167 T4b' 181 T5 196 T5' 197 T5" 200 T6a 240 T6b 264 T6c' 271 T7(DDL) 295 T8 332 T(PZS) 416 T(HW) 566 T(CJB) 596 T(YS) 616 Table 3.1 Distribution and elevation of terraces of the Huang Shui River in the Xining, Ledu, and Minhe depressions and the LaoYa gorge.

To summarize, we deduced the following sequence of tectonic events with their morphological expression (Fig. 3.4). Toward the end of the filling of the tectonic depression around Xining, which also marked the end of peneplanation

58 around the depression, the whole region was gradually uplifted. Rivers responded by slowly entrenching their valleys, while draining toward the deepest point of the depression. In that period the oldest terraces (Yanshang, Houwan, Caijiabao) were formed. Accordingly, Miao et al. (2008) found that these three oldest terraces near Xining and T9 are terraces belonging to such internal drainage, representing a proto-Huang Shui River. The modern drainage southeastward to the Yellow River would seem to have prevailed since the formation of T7. Weak uplift persisted until about 2 Ma, when it accelerated dr- astically. Thus, the change to rapid incision also coincided with a drastic change in drainage direction. A logical explanation for these changes could be an intensified uplift of the Tibetan Plateau that was grad ually declining to the Huang He catchment in the east (east of Lanzhou). In other regions, for instance in SE Spain (Stokes and Mather, 2003), changes in tectonic movements are also associated with drainage diversions as in the Xining Subbasin. 3.4.2. Results 2: Terraces and fluvial deposits in the Ledu tectonic depression The depression is located around the town of Ledu (Fig. 3.1B). It is about 15 km wide and 17 km long. Gorges at its upstream and downstream ends are again incised into resistant Precambrian to Mesozoic rocks, whereas the substratum in the depression consists of Neogene sands similar to those in the Xining depression. The surface of the up-thrust blocks is at c. 3100 melevation, while the Huang Shui flows now at c. 1960 m in the deepest part of the depression. The covering loess layer is of variable thickness (up to more than 25 m), but no dating is yet available. The youngest terrace development is equal to that in the Xining block. The lowermost terraces (T1) are recognized at the same elevations apf (see Table 3.1) and show the same sedimentary characteristics. Thus they may be correlated with the respective terraces in the Xining depression and will not be discussed again here.

Although the origin of the Ledu depression is similar to that of the Xining depression, it has a quite different visual character. The topography within the Ledu depression is dominated by gentle slopes, while the number of terraces is relatively small. A terrace with its gravel top at 38-40 m is highly prominent in the Ledu basin, extending several kilometers in length and a few kilometers in width. However, it cannot be correlated with terrace levels in other blocks on the basis of elevation and, moreover, is completely different from the terraces around Xining. In fact, its main characteristic is its gravel thickness of >30 m in the deepest part of the basin (c. 15 km east of the town of Ledu). The thickness gradually diminishes toward both basin sides (for instance, c. 20 m at Ledu), giving the base of the terrace deposits a bowl shape in the longitudinal direction. Therefore, this depositional form is interpreted as an accumulation terrace, partly covering the subsided block, and not as an alluvial fan. In general, horizontal bedding is dominant, although the structure is rather massive at the base. Large blocks (up to 80 cm diameter) occur in the basal part, while the gravel shows a gradual fining upwards. Their often quartzitic composition points to a provenance from the nearby outcropping metamorphic Paleozoic rocks. In the deepest part of the depression many sand lenses are present, reflecting low-energy conditions. The general horizontal stratification is interrupted in places by low-angle cross-bedding (Fig. 3.5A). At such places, and especially toward the top of the deposit, many small-scale and shallow channels occur (Fig. 3.5B) with relatively steep sides (a few meters wide and a few decimeters deep, facies Gt). They alternate with sandy and silty, horizontal planar beds (facies Gp). Despite their small dimensions, these channels may have been eroded laterally into pre-existing sediments (Fig. 3.5C, D). After abandonment they were filled with silty deposits sporadically interbedded with lenses of finegrained gravel during occasional flooding (described as facies A by Williams and Rust (1969)). Scour features (potholes) have also been observed at their base. The upper part of the fluvial sequence consists locally of

60 a succession of 1.5 m of reddish silt, up to 2 m gravel, and a 1 m reddish silt layer. This succession is interpreted as a flood deposit of varying energy conditions, marking the end of the fluvial activity at that place.

Figure 3.4 Schematic diagram illustrating the main tectonic events and corresponding morphological evolution around Xining since the late Miocene.

Thus far no terraces have been found at a higher elevation. This points to considerable tectonic subsidence. As the elevation of the thick fluvial gravel layer is at a topographically lower position than terrace T7 at Xining, which is 1.4 Ma old, this tectonic subsidence should be younger than that date. As terrace T1d is formed in the thick gravel deposit of Ledu, the tectonic subsidence is older than that terrace. No exact age is available yet for T1d but it is clearly older than MIS 8 (c. 250 ka). This means that the period of subsidence in the Ledu depression must be somewhere between ~1.4 and 0.25 Ma. This is in a period of general uplift in the Xining depression, where erosional terraces are the main feature. We should stress that subsidence of the Ledu block is a relative concept in comparison with the continuous uplift of the whole region. Large-scale fluvial deposition (like the terrace at 38 m) took place only where this local subsidence exceeded the widespread uplift. 3.4.3. Results 3: Terraces in the gorges: The staircase at LaoYa The subsided blocks, such as those of Xining and Ledu, are separated by up-thrust blocks that the Huang Shui has traversed by excavating gorges. Fluvial deposits are only rarely preserved in the gorges, except for the lowermost terraces (T1). This is easy to understand as, since the start of the general uplift of the region, the subsequent entrenchment of the Huang Shui is expressed in the hard basement between subsided blocks. Vertical erosion was thus largely dominant. However, at the transition from such a gorge toward a basin with a corresponding widening of the topography, terraces may have formed. An example of such a situation exists a few kilometers west of the town of Minhe, where a steep staircase of very narrow terraces has been formed at the transition from the LaoYa gorge (located in the Paleozoic bedrock between Ledu and Minhe, Fig. 3.1B) downstream toward the Minhe subsided block.

(A) (B)

Figure 3.5 Characteristic sedimentary structures in the upper part of the main aggradational terrace of the Ledu depression (40 m): (A) Planar sheets of gravel alternating with finergrained, shallow, trough cross-bedded channels; (B) planar sheets of fine-grained gravel alternating with silty floodplain deposits toward the top of the sediment series; (C) clearly erosive character of the basal boundary of gravelly channels; (D) small, gravel-filled channel invaded in former floodplain deposits at the top of the sediment series.

Apart from the lowest terraces, individual terraces of very limited extent are found at altitudes of 100, 124, 144, 167, 181, 197, and 271 m apf (Fig. 3.7). All these terraces have similar sedimentary properties: gravel thicknesses of 1 to 3 m, matrix-supported massive gravel (Gm facies) at the base with large blocks often reaching 1 m in diameter (this facies shows apparent characteristics of debris flows), and Gp alternating with Gt facies in the upper part, with smaller-sized pebbles and imbrication. The gravels are weakly to strongly cemented. Often, the terraces are covered with poorly sorted,

63 homogenized, blocky silt deposits that are interpreted as reworked, weathered material mixed with loess. No dating of these terraces was possible.

Figure 3.6 Terrace with gravel of 4 m thick on top of Cenozoic red beds and overlain by c. 40 m of loess in the Minhe subsided block (gravel top at an altitude of 100 m apf). Remark the strikingly flat topography that mimics the terrace surface at 40 m below the loess cover.

Theoretically, this terrace staircase at LaoYa can be compared with the similar sequence below T8 in the Xining block since both series are mainly representative of predominant uplift. However, apart from this similarity, we found also intrinsic differences. In particular, correlating these terraces with other areas in the Huang Shui catchment by their altitude above the present floodplain level (apf) appears difficult since the uplift history and rates are different (Table 3.1). The terraces of the LaoYa block occur in an upthrust position (Fig. 3.7), whereas the terrace staircase near Xining (Fig. 3.3) occurs in an originally subsided block. This results in different substratum beneath the two terrace series: Neogene sands at Xining and Precambrian-Mesozoic rocks

64 in the LaoYa block. In addition, terrace formation in the Xining area is, albeit rather rarely, intermittently characterized by accumulation terraces, whereas the terraces in the LaoYa gorge are strictly erosional., The logical consequence is that any terrace correlation on the basis of elevation apf appears impossible (see Table 3.1). 3.4.4. Results 4: Terraces and fluvial deposits in the tectonic depression of Minhe The Minhe depression extends over a long distance (several tens of kilometers) up to the confluence of the Huang Shui with the Huang He, but is only about 3.5 km wide (Fig. 3.9). We limit our analysis mainly to the most upstream (western) part around the town of Minhe. At this place, the beveled surface is about 2700m high, whereas the present floodplain is at c. 1760 m. All terraces in the depression are underlain by soft reddish clays and sands of Cenozoic age (Fig. 3.8). The surrounding areas consist of hard Precambrian to Mesozoic rocks (Fig. 3.1B). The lowermost terraces (T1) occur at the same elevation above the present floodplain as in the other tectonic blocks and may thus be easily correlated. As with the Ledu depression, the Minhe depression is characterized by the occurrence, above the level of the lowermost terraces, of impressive accumulation terraces. A notable difference is that in the Ledu depression only one accumulation terrace is found, whereas in the Minhe depression at least three of these terraces are present, the tops of their gravels occurring at about (197 to) 200, 60 (57-65), and 44 (48) m apf. Possibly the latter terrace might be correlated with the impressive accumulation terrace at 40 m in the Ledu depression. The terraces at 60 and 200 m in the Minhe depression have no equivalent in the Ledu depression. Isolated patches of erosional terraces are found at 54, 35, and 22 m toward the east in the Minhe depression. The 200-m terrace has the largest thickness of fluvial deposits: c. 30-35 m. The gravel thickness at the 60-m terrace varies from 18 m in the westernmost part of the

65 depression, at the town of Minhe, to 11 m some 20 km east of that town and to 4 m to the east (near the confluence with the Huang He). The +44-m terrace has a gravel thickness of >12 m in the western part of the Minhe depression, but seems to become much thinner (3 m) toward the east. 300 NW

271m

250

200m 197m 200

181m A

167m d

150 o 144m v

124m e

97m l

100m

(

100 m

)

60m 50 44m

T1d T1c T1b T1a SE 0 Figure 3.7 Terrace sequence at the exit of the LaoYa gorge and the transition to the western most part of the Minhe basin (location in Fig. 3.1B).

In between these terraces, a morphologically well-expressed terrace level occurs, with the top of its gravel at 97-100 m apf. The gravel unit is only c. 4 m thick and is therefore considered as an erosional terrace in contrast to the terraces immediately below and above (Fig. 3.6). It is covered by 5-6 m horizontally and finely laminated, fine-grained floodplain deposits. The final overburden consists of c. 40 m of loess, which is also the case for the 60 m terrace. The sedimentary structures of the accumulation terraces at 200 m, 60 m and 44 m are very similar. The lowest part of the terrace deposits consists of

66 massive gravel (Gm) with blocks up to 1 m diameter, grading toward the top into finer-grained, coarsely planar bedding (Gp) as a result of alternating coarse- and fine-grained gravel. Shallow channels occur with increasing frequency toward the top of the terrace deposits and are filled with cross- bedded fine gravel or sands (planar to low-angle cross-bedding). The gravels are generally matrix-supported, but in the channels sometimes clastsupported. Imbrication is a common characteristic. The distance of the terraces near Minhe (in the western most part of the tectonic block) to the terraces at the edge of the LaoYa gorge is only a few kilometers. Therefore, the terraces of both the gorge and the Minhe depression are combined into one schematized section (Fig. 3.7). The terrace sequence in the western Minhe depression occurs mainly south of the Huang Shui, while the terraces at the exit of the LaoYa gorge (described above) occur in particular to the north of the Huang Shui. The lateral transition between the very erosion-resistant Palaeozoic rocks of the upliftted block to the much younger and less resistant Tertiary and Mesozoic sands, sandstones and clays in the depression corresponds with a very prominent, almost vertical escarpment in the topography near the town of Minhe (Fig. 3.8A, B). It is clear that this morphological escarpment may be interpreted as a fault. Sometimes the boundary is more gradual at places where the depression is bounded by many small faults. In the combined terrace section of Fig. 3.7, vertical erosion (with formation of the strath terraces) obviously was always the main fluvial activity NW of the Huang Shui, as a result of the regional tectonic uplift, while SE of the Huang Shui vertical erosion was interrupted by accumulation at specific periods as a result of local subsidence in that part of the Minhe depression. We like to stress that the decrease of the gravel thickness of both the 44 and 60 m terraces toward the east suggests tectonic tilting of the subsiding Minhe block at the time of formation of those terraces (high rate of subsidence in the west, slight or no subsidence in the east).

(A)

(B)

Figure 3.8 (A) Fault scarp at the boundary between the LaoYa block composed of Proterozoic and Paleozoic hard rock and the Minhe depression underlain by Cenozoic soft sediments.(B) Same, near-vertical fault visible in exposure of Datong tributary near its confluence with the Huang Shui (gray metamorphic rocks at left side of white line-red Tertiary sandstone at right side of white line).

Reconstruction of the tectonic history of the Minhe depression, in combination with the LaoYa up-thrusted block, since the formation of the

68 highest terrace at 271 m can only be made in general terms (Fig. 3.7). In a first phase, incision occurred from 271 to 197 m, due to regional tectonic uplift. Uplift then continued in the LaoYa block, while the Minhe block subsided. This subsidence coincided or was followed by accumulation of some 30 m of river gravel (terrace at 200 m). Subsequently (second phase), uplift was dominant again everywhere, resulting in the formation of a series of erosion terraces in the hard bedrock of the LaoYa block north of the Huang Shui between 197 and 100 m; while only one erosional terrace is preserved in the tectonic depression (a prominent terrace at 100 m). The very sharp morphological scarp north of the Huang Shui below the 100 m-erosion terrace (Fig. 3.8) demonstrates recent fault activity (third phase), which resulted in sharp vertical erosion north of the Huang Shui (even no erosion terraces were formed). This uplift may have amounted to up as 80 m (the difference between the terraces at 100 m and T1d), confirming that the tectonic uplift here was different from that of the Xining depression (with low or no uplift). The tectonic history and the consequent terrace morphology in the Minhe depression were more complex in that third phase. In the western part, subsidence prevailed, resulting in the formation of two accumulation terraces at 60 and 44 m apf, which shows that the tectonic activity was probably episodic rather than instantaneous or gradual., In the eastern part, however, uplift was dominant, resulting in the formation of five erosion terraces below 100 m (at 60, 54, 44, 35, and 22 m). No solid information on the age of the accumulation terraces is available. With reference to the ages of the lower most terraces in the Xining depression, all terraces at a higher elevation than 11-13 m apf are older than the T1c terrace at that elevation (c. 250 ka BP).This is thus a minimum age. In addition, the 100 m terrace may preliminarily be correlated with the T4 terrace of the Huang He near their confluence west of Lanzhou. This is possible since the Minhe depression is directly connected with the Huang He valley without any

69 upthrusted area in between (Fig. 3.9). That T4 terrace of the Huang He was dated at c. 860 ka by Pan et al. (2009), as it is covered by the S8 soil and L9 loess layer. This is thus a maximum age for the terraces below the 100 m terrace. Between these two brackets this might lead to the as-yet hypothetical ages intermediate between 250 and 860 ka for the 22, 35, 44, 54 and 60 m terraces. No solid information on the age of the accumulation terraces is available. With reference to the ages of the lower most terraces in the Xining depression, all terraces at a higher elevation than 11-13 m apf are older than the T1c terrace at that elevation (c. 250 ka BP).This is thus a minimum age. In addition, the 100 m terrace may preliminarily be correlated with the T4 terrace of the Huang He near their confluence west of Lanzhou. This is possible since the Minhe depression is directly connected with the Huang He valley without any upthrusted area in between (Fig. 3.9). That T4 terrace of the Huang He was dated at c. 860 ka by Pan et al. (2009), as it is covered by the S8 soil and L9 loess layer. This is thus a maximum age for the terraces below the 100 m terrace. Between these two brackets this might lead to the as-yet hypothetical ages intermediate between 250 and 860 ka for the 22, 35, 44, 54 and 60 m terraces.

3.5. Terrace development The present-day Huang Shui may be considered as a wandering gravel-bed river (terminology by Miall, 1996) characterized by intermediate sinuosity, bank-attached point bars and a variable amount of channel paths depending on the river stage. This river type is in contrast with the shallow gravel-bed braided river that we reconstruct from the sedimentological descriptions as dominated by varying (often cyclic) assemblages of gravel traction-current deposits (mainly the Gp facies). The grain size in such assemblages often decreases upward as the track aggrades and becomes inactive (Rust, 1978).

Occasionally, this deposition pattern is interrupted by channel abandonment, in which case lenses of sand or silt may be deposited at low stage (Gt facies). Such a river type requires certainly higher energy conditions at times, as is also confirmed by the larger blocks that were transported (up to 1 m diameter) in comparison with the maximum size of the blocks that are transported now (about 30 cm).

Elevation (m) 2674

LaoYa

Minhe

0 1 2 3 4 KM 1688 Figure 3.9 DEM of the Minhe subsided block with the transition to the LaoYa gorge in the left upper corner.

Two distinct classes of terraces can be distinguished. In strath terraces, erosion and accumulation are alternating processes keeping the thickness of bed deposits rather constant. Low-angle, planar and trough crossbedding are dominant except near to the base, where massive gravel may occur. Internal scouring and imbrication structures are very common. The ensemble of (sub)facies clearly points to deposition at the level of the channel bed, while the reduced thickness of these bed sediments (a few meters) cannot be the consequence of successive erosion and aggradation („cut and fill‟) as no large- scaled truncation of previously aggraded sediments has been observed in the sedimentary sections. Instead a near-equilibrium state is derived, in which the river was mainly in the stage of sediment transport (without net deposition or

71 erosion). In contrast, in accumulation terraces gravel deposits may reach thicknesses of up to >30 m. The basal part is deposited in large, wide channels with many extensive internal bars. The very large clasts (>1 m diameter) testify to the occasionally very energetic conditions. In general, the average grain size is slightly decreasing toward the tops, while silt and fine-sand lenses are common. The dominant structure (especially in the upper part) consists of extensive horizontal gravel sheets with common planar cross-bedding and occasionally internal, shallow channel structures (max. 1 m deep, up to some 20 m wide). Erosion of these small channels locally interrupts the general accumulation trend (Fig. 3.5). The fine-grained lenses that fill these channels interfinger with the gravel sheets. In the upper part the river became shallower (a few decimeters) and somewhat less energetic, ending up in a multichannel system with smaller individual channels (5-10 m wide). In the uppermost part, flooding and silting up of abandoned channels are preserved because the channel activity was much reduced. In general, sedimentary structures do not differ from those in strath terraces. After gravel deposition ceased, the river incised slightly in response to tectonic uplift or climatic change. Occasionally, at peak discharges, the river flooded over the former channel deposits and covered them with fine-grained flood loams, this occurring despite the fact that the new channel bed was now at a lower level and the older channel gravels were morphologically already in a terrace position. Such a situation was advantageous for the preservation of the flood loams. It is an evolution that has been found at many occasions in stepwise incising rivers (e.g., Vandenberghe, 1993, 1995; Antoine et al., 2000; Cordier, 2006). At several locations, interglacial soil formation has been recognized at the top of fluvial depositional sequences (e.g., Pan et al., 2003) as is the case in other fluvial systems (e.g., Antoine et al., 2000). However, this requires

72 sufficient time for soil formation between the deposition of the uppermost fluvial sediments and of the overlying loess cover. Mostly, a gradual transition from flood loam to loess is present (Fig. 3.10A). In the studied sections, soil formation was not well developed, but the flood loams are often red-colored, pointing to upstream removal or local reworking of soil material (Fig. 3.10B). Deposition of flood loams at peak discharges means also that the river was not deeply entrenched at that time. A significant difference between the Huang Shui catchment and all other mentioned cases is the great age of the oldest terraces found along the Huang Shui (late Miocene to Pliocene). Those terraces were separated only by small differences in elevation and were possibly related to an interior drainage system that pre-dated the present-day Huang Shui. They record the regionally extended, relatively slow dissection of a previously planed-off, peneplain-like surface, irrespective of the different ages of those individual beveled surfaces in the regions adjacent to the Huang Shui catchment (Chapter 2). Terrace formation in most European catchments started distinctly later than in China (end of the Pliocene), except in some cases as the Maas and Dnjepr catchments (van den Berg and van Hoof, 2001; Matoshko et al., 2004; Bridgland and Westaway, 2008b). Interestingly, in Europe and China tectonic uplift accelerated afterward: between 1.4 and 2 Ma near Xining in the Huang Shui, and in Middle Pleistocene times (roughly between 0.5 and 1 Ma) in Europe (e.g., Van Balen et al., 2000; Van den Berg and van Hoof, 2001). According to Bridgland and Westaway (2008a) and Westaway et al. (2009) it should be a response to Middle Pleistocene climate change. Thus, morphological development was in general quite similar in both regions, only the timing was different (earlier in the Huang Shui catchment than in Europe) and may be due to localized tectonic effects.

(A)

(B)

Figure 3.10 (A) Slightly dipping sheets of silty, fine-grained sand with intercalated megaripples of medium sand, overlain by flood loams in which several soil layers may be observed (location: c. 25 km east of Xining); (B) reworked soil material in flood loam deposits between fluvial gravel and loess (c. 30 km west of Lanzhou, 60 m terrace).

3.6. Discussion: Tectonic and climatic impact on terrace morphological and sedimentological development Rivers were only able to entrench and construct a terrace staircase during or after tectonic uplift, but the fluvial processes did that under specific climate constraints, as has been proposed in models of climate driven fluvial development (references cited above). Indeed, climate change would only cause vertical incision if there had been uplift; otherwise the river would probably not have been able to cut significantly downwards. While staircases have been commonly well described for the Pleistocene epoch in temperate and periglacial conditions, as a response to glacial-interglacial succession, a similar alternation of stability and incision as a result of climate changes has not firmly been established for monsoonal catchments. Specific case studies have shown that similar staircases formed also in other climates. Conditions of vegetation cover, slope stability, sediment supply and peak-runoff cyclicity were also valid for terrace development in other, even tropical, climatic regions (e.g., Krook, 1979; Bridgland and Westaway, 2008b). Thus, a similarly climate-driven cyclicity in the monsoonal environment of China should not surprise. We have shown that the fragmentation of tectonic blocks was spatially differentiated. However, tectonic movements also varied within the temporal frame. For instance, at certain times and in certain areas sudden and drastic reductions in slope gradient by tectonic movements were responsible for continuing deposition, sometimes resulting in accumulation of fluvial sediments of considerable thickness. Such areas are intricately linked to tectonic subsidence, so that climatic forcing can be excluded here. In contrast, in the most upstream part of the Huang Shui catchment and in the gorges, tectonic uplift explains the removal and transport of substantial and coarse- grained fluvial sediments and river entrenchment without climatic forcing. Such a spatial differentiation of terrace development is absent for the lower most terraces that may be correlated by elevation easily over the entire

75 catchment (Table 3.1). This enabled us to conclude that the young tectonic history was also uniform over the entire catchment. Of course, climate changes also played a role in uplifted and subsiding basins, but the tectonic impact may have been by far more preponderant during specific periods. In fact, the combination and interplay of climate influences and relative movements between uplifted and subsided areas have determined, at different temporal and spatial scales, fluvial activity and its morphological expression. Climate proxies are mostly absent within the Huang Shui terrace deposits, which makes it difficult to prove the link between climate and fluvial processes. Nevertheless, the majority of the (coarse grained) fluvial deposits seem to date from cold periods, while the fluvial succession is terminated by fine-grained deposits and ultimately interglacial soil formation, as illustrated in Fig. 3.8. In accordance with the reconstructed character of the river morphology as a shallow gravel-bed braided river (Miall, 1996) and in line with general models of fluvial development derived for temperate and periglacial environments (last update by Vandenberghe, 2008), the main body of the (coarse-grained) terrace sediment dates probably from a glacial period. Subsequently, we assume that the fine-grained sandy gravels and sands near to the top of the sediment series were deposited during the waning of the same glacial period as a result of the beginning incision of the river (possibly in combination with reduced precipitation). Rivers incised slightly at the transition from the glacial to the interglacial, due to less peaked river discharge and the reduction of sediment supply to the river. As a consequence of that new, lower position in the next interglacial, the previous coarse grained channel deposits were flooded only during peak discharges resulting in episodic fine-grained floodplain deposition, while most of the time soil formation took place on the floodplain. Analogous to the present-day river, the interglacial river was probably less energetic and dominantly meandering. The main incision took place at the next interglacial- glacial transition, when fluvial energy considerably increased and sediment

76 transport was still limited (Vandenberghe, 2008). At that time, general loess deposition resumed, covering the former floodplain deposits and soil without any further reworking by the river. This timing is somewhat different from that proposed by Pan et al. (2009) who claimed that the main incisions took place at the transition from glacials to interglacials. They base their conclusions on the occurrence of interglacial soils in the lower most loess that covers the fluvial deposits coinciding with the abandonment of the floodplain due to river entrenchment, in contrast to our observations that show the soil formation within the top of the fluvial overbank deposits. Also Bridgland (2000) and Bridgland and Westaway (2008a) favor main incision at the beginning of interglacial periods. In either case, we can conclude that the general sequence of terrace deposits reflects a climatic impact although the exact timing needs further investigation. The temporal and spatial differentiation of tectonic movements resulted in a three-fold geomorphological subdivision of the Huang Shui catchment. Most obvious are the gorges with entrenched river sections in outcropping, resistant, uplifted Paleozoic-Cretaceous rocks with terrace remnants, if any, that are solely erosional., In the second place, compartments where overall tectonic uplift of the NE Tibetan Plateau was dominating but affected by moderate local subsidence resulting in the formation of a series of well-developed, wide, mostly erosional terraces (e.g., Xining block). Thirdly, some compartments underwent at specific time intervals local subsidence of such intensity that it exceeded the regional Tibetan uplift resulting in the formation of accumulation terraces at those times in those compartments (e.g., Ledu basin). However, it should be stressed that in all of these blocks, tectonic impact was overprinted by climate-induced terrace staircase development at a more restricted (glacial- interglacial) timescale than the phases of tectonic uplift or subsidence (see Table 3.1).

3.7. Conclusions (i) From late Miocene to Pliocene times, the Huang He and its tributaries incised from a peneplain-like surface that was formed on top of Cenozoic and Palaeozoic folded and fractured bedrock. This geomorphic evolution was induced by general uplift of the Tibetan Plateau. (ii) The general tectonic uplift of the Tibetan Plateau also caused regional fragmentation of the Huang Shui catchment into separate tectonic blocks that alternatively subsided or were uplifted. The geomorphological result is a succession of smallscale depressions in which the Huang Shui has deposited thick layers of sediment and of uplifted blocks through which this river has cut narrow gorges. (iii) Erosional terraces (fluvial deposits of <3 m) were formed during phases of uplift; accumulation terraces (fluvial deposits>30 m) during subsidence. (iv) Unequal rates of uplift and subsidence make it impossible to correlate terraces of the individual tectonic blocks, except for the three lowermost terraces, which have been formed since about 250 ka. Thus differential movements within the Huang Shui catchment have not been important since that time. (v) The general, large-scale incision was superseded by alternations of fluvial stability and erosion of shorter duration, induced by climatic changes and resulting in the formation of a terrace staircase. Fluvial layers generally consist of gravelly channel fill, while finer-grained floodplain deposits are only preserved from the final phase of deposition just before rivers renewed incision. The floodplain deposits sometimes show traces of interglacial type soil formation.

Acknowledgments This work is supported by the National Natural Science Foundation of

China (NSFC) (projects 40901002, 40325007 to the second and third author); by the Chinese Academy of Sciences (CAS) and the Royal Netherlands Academy of Arts and Sciences (KNAW) in their joint PhD Training Programme (05-PhD-10) to the second author; and continuous support by the Dutch-Chinese cooperation programme. Ronald van Balen is thanked for his constructive discussions. The comments by David Bridgland and two other anonymous reviewers are highly appreciated.

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-Chapter 4- Late Quaternary paleoclimatic and geomorphological evolution at the interface between the Menyuan basin and the Qilian Mountains, northeastern Tibetan Plateau XianyanWang, Dimitri Vandenberghe, Shuangwen Yi, Jef Vandenberghe, Huayu Lu, Ronald Van Balen and Peter Van den haute

Abstract: The Tibetan Plateau is regarded as an amplifier and driver of environmental change in adjacent regions because of its extent and high altitude. However, reliable age control for paleoenvironmental information on the plateau is limited. OSL appears to be a valid method to constrain the age of deposits of glacial and fluvial origin, soils and periglacial structures in the Menyuan basin on the northeastern Tibetan Plateau. Dating results show glaciers advanced extensively to the foot of the Qilian Mountains at ~21 ka, in agreement with the timing of the global Last Glacial Maximum (LGM) recorded in Northern Hemisphere ice cores. Comparison with results from the eastern Tibetan Plateau suggests that the factor controlling glacial advance in both regions was decreased temperature, not monsoon-related precipitation increase. The areas of the Menyuan basin occupied by glacio-fluvial deposits experienced continuous permafrost during the LGM, indicated by large cryoturbation features, interpreted to indicate that the mean annual temperature was ≥7 °C lower than at present. Glacio-fluvial systems in the Menyuan basin aggraded and terraces formed during cold periods (penultimate glaciation, LGM, and possibly the Younger Dryas) as a response to increased glacial sediment production and melt water runoff then.

Based on: Quaternary Research 2013,80, 534-544.

4.1. Introduction The Tibetan Plateau is thought to be a critical factor in the environmental evolution in Asia, especially as an amplifier and driver of climatic change due to its high altitude over a vast area (e.g., Pan et al., 1995; Liu and Chen, 2000). However, paleoenvironmental conditions (especially air temperature) and landscape evolution interlinked with climate change there are not clearly understood. This is due to insufficient paleoenvironmental indicators and reliable age control, especially in the northeastern part of the region. Glacier fluctuations (e.g., Derbyshire, 1987; Shi, 2002; Zhou et al., 2002; Yi et al., 2008; Owen, 2009) and periglacial structures (formed by thaw modification of frozen ground; e.g., Porter et al., 2001; Wang et al., 2003; Vandenberghe et al., 2004; Liu et al., 2013) have been used to reconstruct former paleoenvironmental conditions in the study region and its surrounding areas. However, there is an ongoing debate on the timing of Quaternary glaciations and their relationship to changes of precipitation and temperature (e.g., Shi, 2002; Zhou et al., 2002; Lehmkuhl and Owen, 2005; Owen et al., 2009; Zhou et al., 2010; Ou et al., 2012), which is partly due to a lack of suitable datable materials (for radiocarbon and terrestrial cosmogenic nuclides exposure-age dating). Although the timing and succession of glaciations are ill- defined as a consequence of poor chronological control in most parts of the Tibetan Plateau, it was previously concluded that glacial advances there were mainly influenced by precipitation, controlled by the South Asian monsoon (e.g., Owen et al., 2009). However, it was also proposed that in the eastern Tibetan Plateau glacial advances during the last glacial period were mainly controlled by temperature and that the precipitation is a secondary factor (Ou et al., 2012). In addition, it was pointed out that the glacial advances in the southeastern Tibetan Plateau occurred during Marine Oxygen Isotope Stage (MIS) 3 and the early Holocene, while they were scarce during the global Last Glacial Maximum (LGM) (e.g., Owen et al., 2009). Some authors concluded

88 also that in accordance with the glacial advances, rivers aggraded and filled valleys during MIS 3 in the southern Tibetan Plateau (e.g., Goodbred, 2003; Dutta et al., 2012). In the Qilian Mountains of the northeastern Tibetan Plateau, several glacial advances have been reported, dating from 470-460 ka, penultimate glaciation, the last glaciation, and the Little Ice Age (Kang et al., 1992; Guo et al., 1995; Zhao et al., 2001; Zhou et al., 2002). Age constraints in these studies are limited to a few radio-carbon and electron spin resonance (ESR) dates (Kang et al., 1992; Guo et al., 1995; Zhao et al., 2001; Zhou et al., 2002). Likewise, the age control on periglacial conditions is poor. Thus far, aeolian sands in sand wedges have been dated as LGM near Qinghai Lake, in the Hexi Corridor and on the Mu Us desert by using thermoluminescence (TL) and optically stimulated luminescence (OSL), supplemented by a few radiocarbon ages of fluvial and lacustrine deposits (Porter et al., 2001; Wang et al., 2003; Vandenberghe et al., 2004; Liu and Lai, 2013). These results show the presence of continuous permafrost during the LGM, indicating that at that time the mean annual temperatures were more than ~12 °C lower (Wang et al., 2003; Vandenberghe et al., 2004) or ~3 °C lower (Porter et al., 2001) than nowadays. In accordance with the morphological evolution in glacial environments, a few case studies (e.g., Pan et al., 2009) have shown that fluvial aggradation took place during periglacial conditions on the Tibetan Plateau during the Pleistocene, caused by changes in precipitation, cessation of soil development and decrease of vegetation cover as a response to climate deterioration, which is in agreement with general observations (e.g., Vandenberghe, 1995, 2003; Bridgland and Westaway, 2008; Srivastava et al., 2008; Vandenberghe, 2008). In addition, investigations have shown that fluvial landforms in the northeastern Tibetan Plateau were controlled by both climate change and tectonic uplift (Lu et al., 2004; Vandenberghe et al., 2011; Wang et al., 2012). However, a few studies from the southern Tibetan Plateau have reported fluvial accumulation

89 and valley fill, mainly in interstadial conditions (e.g., in MIS 3) and due to enhanced precipitation associated with a strong insolation controlled South- Asian monsoon (e.g., Goodbred, 2003; Thiede et al., 2004; Gibling et al., 2005; Dutta et al., 2012). The study area is located at the foot of the Qilian Mountains in the northeastern Tibetan Plateau, where glacial valleys and moraines, out-wash fans, periglacial structures, and fluvial terraces are extensively present and are morphologically connected. Thus, it is an ideal area to unravel the relationship between the glacial and fluvial processes and their detailed response to climate changes. The objective of this study is to refine and extend the paleoclimatic database and to specify reconstruction of landscape evolution using OSL ages in the glacial and fluvial setting of the Menyuan basin (Fig. 4.1).

4.2. Geomorphological setting The Menyuan basin is a topographic depression at the northeastern margin of the Tibetan Plateau, bounded by the Daban Mountains to the south and the Qilian Mountains to the northeast (Fig. 4.1). The subsurface of the basin consists of reddish Neogene fluvio-lacustrine mudstone, siltstone, and sandstone, overlain by Pleistocene deposits. The altitude of the study area is between 2750 m and 5200 m. The mean annual temperature and precipitation are ~0.8 °C and ~520 mm respectively. Preci-pitation falls mainly in the summer, caused by invasions of the East-Asian and south-Asian monsoons. In winter, the climate of this region is controlled by the Mongolian high-pressure cell. The present snowline is at 4400 m. Modern glaciers cover high mountain areas in the study area and the melt water from high-altitude snow is an important source of fluvial discharge. The Datong River flows through the basin from northwest to southeast.

Altitude 101°10′ 101°40′ (m) 5219 Q il ia L n a r d i t e u t a i 37°40 s c y la e M l g l o R A a u i v n ′ v C# # t e HC a r GSK in s 4400 d e D t a i s c y la e g ll Da a to v ng

R QSZ iv # er E

B Menyuan !. #DZ 37°20 D aba F fig.4.2 n

′ Tibetan Plateau

M ou nta ins 2750 Figure 4.1 Digital Elevation Model of the Menyuan basin derived from ASTER Global Digital Elevation Model (ASTER GDEM), with geomorphologic characteristics and localities mentioned in the text (white triangles: the studied sections; white dashed lines: the extension of outwash fans; white solid lines with letter: selected topographical profiles). The insets show the position of the studied region (Menyuan basin) on the Tibetan Plateau (star: Menyuan basin; circle: Mu Us desert; rectangle: Hexi Corridor; triangular: Qinghai Lake).

The morphology of the southern edge of the Qilian Mountains is characterized by glacial valleys, whereas moraines and outwash fans occur just below the mountains in the Menyuan basin. Tributaries to the Datong River have terraces. The fans and terraces are connected to or correlatable to the fluvial terraces of the Datong River in the basin (Fig. 4.1).

4.3. Methods 4.3.1. Field work

Large-scale landforms were mapped using remotely sensed data from various sources. The principal dataset was the Advanced Spaceborne Thermal Emission and Reflection Radiometer (ASTER) Global Digital Elevation Model (ASTER GDEM), with a horizontal resolution of 30 m (http://gdem.ersdac.jspacesystems.or.jp). Field work was undertaken on the south slope of the Qilian Mountains and in the Menyuan basin. Firstly, detailed studies were carried out in the Datong catchment to determine the Datong terrace sequence and its spatial distribution. Then, typical topographic sections composing glacial valleys at the southern side of the Qilian Mountains and outwash fans, glacial and fluvio-glacial sediments in the Menyuan basin were morphologically correlated with the terrace sequence of the Datong River. These sections (Gangshika (GSK), Qingshizui (QSZ), Huangcheng (HC), and Dazhuang (DZ); locations in Fig. 4.1) in combination with topographic analysis in between allowed examination of the relationship between the glacial valleys and moraines in the Qilian Mountains following the outwash fan successions towards the terrace staircase and into the basin. The individual outcropping deposits were described for their sedimentary structures (using lithofacies according to Miall, 1996; Benn and Owen, 1998), lithological properties, and thickness as far as they could be measured, in order to derive the sedimentary processes. The moraines, outwash fans, and terraces were examined and correlated morpho-stratigraphically, and mapped using a Garmin Vista GPS and Magellan Mobile Mapper CX, integrated with real-time, sub-meter accuracy using SBAS. For correlation purposes, we used the relative height of the terraces above the present-day floodplain of the Datong River. 4.3.2. Optically stimulated luminescence (OSL) dating Sediment samples for OSL dating were collected by hammering light-tight tubes into freshly exposed sediments. The tubes remained sealed until processed in safe light conditions. Within the frame of this work, a total of eleven samples were analyzed. Sample preparation as well as the determination

92 of the dose rate were done at Nanjing University (China); the OSL measurements were undertaken at Ghent University (Belgium) and Nanjing University. Very fine sand-sized (63-90 μm) quartz grains were extracted from the inner material of the sampling tubes in the standard manner (10% HCl, 30% H2O2, wet sieving, 40% HF, and dry sieving). Initial screening suggested that the OSL signals of the samples were not bright and that their luminescence sensitivity was low, similar to other OSL dating studies on the Tibetan Plateau (Lai et al., 2009; Liu et al., 2010; Lu et al., 2011). We therefore analyzed all samples using large aliquots (8 mm diameter). We also analyzed two relatively bright samples (GSK-1 and -2) using small aliquots (2 mm diameter) and compared the measurement results from small aliquots and large aliquots to assess the degree of partial bleaching, as the glacio-fluvial sediments are likely to suffer from poor or inhomogeneous bleaching. Luminescence measurements were made using automated Risø-readers equipped with blue (470 ± 30 nm) LEDs and an IR laser diode (830 nm). The luminescence emissions were detected through a 7.5 mm Hoya U-340 UV filter. Details on the measurement apparatus can be found in Bøtter-Jensen et al. (2003). A single-aliquot regenerative-dose (SAR) protocol (Murray and Wintle, 2000) was used to determine the equivalent dose (De). Preheating of natural and regenerative doses was for 10 s at 240 °C, while the response to the test dose was measured after a cut-heat to 200 °C. Optical stimulation with the blue diodes was for 40 s at 125 °C. The initial 0.32 s of the decay curvewas used in the calculations, minus a background evaluated from the following 0.64 s. As the samples were not particularly sensitive, a relatively high test dose in the range of ~50-100 Gy was used for all but sample HC-1; for the latter sample, a test dose of ~6 Gy was used. After measuring the response to the test dose, a high-temperature bleach was performed by stimulating with the blue diodes for 40 s at 280 °C (Murray and Wintle, 2003). For each aliquot, the dose response

93 was obtained by measuring the response to four regenerative doses. This was followed by three additional measurements to obtain estimates of recuperation and recycling (Murray and Wintle, 2000), and purity (OSL IR depletion ratio; Duller, 2003). To assess the reliability of the laboratory measurement procedure, a dose recovery test was performed (Murray and Wintle, 2003). Fresh aliquots were bleached for two times 250 s using the blue diodes at room temperature; the two bleaching treatments were separated by a 10,000 s pause. A regenerative dose was then given, chosen to be equal to the estimated natural dose, and measured using the SAR protocol outlined above. The material from the outer ends of the sampling tubes was retained for radioisotope and water content analysis. The in-situ water content and that at saturation were determined by drying sub-samples in an oven at 50 °C (Long et al., 2011). A subsample of ~20 g dried sediment was ground to a fine powder to determine concentrations of U, Th, and K by neutron activation analysis (NAA). The elemental concentrations were converted into external beta and gamma dose rates using the conversion factors of Adamiec and Aitken (1998) and beta attenuation factors of Mejdahl (1979). The water content may have varied over the entire burial period of the samples, but it is difficult to determine the degree of such changes; we have assumed 50% of the saturation water content as representative for the time-averaged degree of wetness and assigned a relative uncertainty of 25% to this value to allow for possible fluctuations. An internal dose rate in quartz grains of 0.013 ± 0.003 Gy ka was assumed (based on Vandenberghe et al., 2008), and the contribution from cosmic rays was calculated following Prescott and Hutton (1994).

4.4. Results Firstly, the terrace sequence in the basin is summarized as derived from the DEM analysis and field investigations. Then, three typical sedimentary sections of different ages are described and morphologically correlated with the

Datong terrace sequence. Finally, the sedimentological interpretations and the OSL dating results are reported and discussed. 4.4.1. The Datong River terrace staircase in the Menyuan basin From the DEM topographic analysis (Fig. 4.2) and field morphological investigations (Fig. 4.3A), several extended flat surfaces (or patches) in the Datong valley were distinguished at different altitudes above the present floodplain. Field investigations show that these surfaces are usually underlain by widely occurring fluvial beds, containing gravel from the Datong River, and covered successively by alluvial and Aeolian loess (Fig. 4.3C). In total, six terraces are distinguished (T1 to T6), at around 2 m, 7-8 m, 13-14 m, 21-23 m, 43-45 m and 78-80 m above the present floodplain level (m apf) (Fig. 4.3C). The different terraces show similar sedimentary properties: fluvial gravels of various thicknesses, interbeded with lenses of sands, silts, and clays and finally topped by horizontally laminated silt containing gravel strings of limited extent. The gravels consist mainly of quartzite, sandstone, and granite with small amounts of conglomerate and crystalline and metamorphic pebbles, which are mainly rounded to sub-rounded. The lowest part of the gravel deposits consists of massive gravel (Gm) representing channel bedload, grading towards the top into finer-grained, coarsely planar bedding (Gp). The latter deposits frequently show imbrication and small-scale cross-bedding, which indicate lateral and longitudinal bar development during deposition. Small and shallow channels occur with increasing frequency towards the top of the terrace deposits; they are filled with cross-bedded fine gravel or sands (planar to low-angle trough cross-bedding). The gravels are generally matrix-supported, but in the channels they are sometimes clast-supported. The upper laminated silts are flood loam deposits that complete the fluvial sequence prior to or simultaneously with the abandonment of the concerned floodplain as a result of renewed river incision. They may even have been deposited at peak flooding during the initial phase of that renewed river incision.

The three lower terraces (T1, T2, and T3) are widespread along the Datong River and are well expressed in the morphology, whereas the three higher terraces (T4, T5, T6) are distributed as isolated patches. The thickness of the exposed terrace gravel is 2 m, 7 m, 6-9 m, 14 m, 36 m, and 3-5 m for T1, T2, T3, T4, T5 and T6 respectively (Fig. 4.3C). These thicknesses are mostly minima as the base of the gravel deposit is not always exposed. Based on the thicknesses of the gravel deposit, it seems that at least the fifth and fourth terraces (T5 and T4) could be interpreted as accumulation terraces (in this interpretation, we consider that the terrace surface formed due to continuous aggradation of the river). But it is also possible that the complete gravel thickness belongs to the sixth terrace and the later terraces just incised in, and removed part of the former gravel deposit of T6. Additional field data are required to obtain a more detailed insight into the aggradation and incision cyclicity recorded by the higher terrace sequence. The alluvial loam at the top of the terraces is 0.5-1 m thick except on the sixth terrace, where it is around 3.2 m. The loess overlying the second (T2), third (T3) and fourth terrace (T4) is 0.5-1.5 m thick. The loess overlying the fifth (T5) and sixth terrace (T6) is around 25-40 m thick. The first terrace is not covered by loess. In general, the main sediment bodies of these terraces are too coarse- grained for luminescence dating. Two samples (DZ-23 and DZ-28) from the flood loam of T6 were taken and one sample (DZ-1) from the top of the overlying loess (DZ section) and one sample (QSZ-1) from the flood loam of T3 at QSZ near the GSK section were taken (Fig. 4.3C).

4.4.2. Description of the sections and geomorphological correlation Three typical sites near the footwall of the Qilian Mountains, Gangshika (GSK), Huangcheng (HC) and Dazhuang (DZ) (Fig. 4.1), were selected for

96 detailed sedimentological description, morphological correlation with the Datong terraces and OSL dating. 1. At the GSK section, a large U-shaped tributary valley, a morainic hill- plain and an extensive glacial outwash fan are present (Fig. 4.1, 4.4). The top of the studied sedimentary sequence is at around 50 m above the present floodplain of the Baishui River, and this surface gradually declines to connect with the T5 terrace (43-45 m apf) of the Datong River (Fig. 4.5). The sediments arewell exposed over ~6 m depth in road cut, and situated below the surface of the moraine that was formed during the last glaciation according to Kang et al. (1992). The series is composed of four sedimentary units (Fig. 4.4B, C). The first unit, at 0-0.5 m depth, is the present black soil layer. The second unit, at a depth of 0.5-2.8 m, is a gray diamicton. It consists mainly of very poorly sorted, matrix-supported, disorganized, ungraded, sub- angular to angular gravel (3-20 cm diameter), interfingering with few layers of sand and silt. It contains a few big cobbles (>60 cm diameter), some of which are glacially striated. Therefore, this layer is interpreted to be a till. The third unit, at 2.8-3.7 m depth, is a black-brown sandy soil layer. The fourth unit, at 3.7-6 m depth, consists of gravel with sand lenses. The lower boundary of the latter layer is not exposed. This unit mainly consists of sub-rounded to sub- angular, poorly sorted pebbles (2-20 cm diameter) in a sandy matrix, with very crude horizontal stratification, a fining upward structure and occasionally shows clast imbrication. These sedimentary properties suggest that this layer is deposited as bedload and in longitudinal bars by glacio-fluvial systems. Three OSL samples from this sediment sequence (Fig. 4.4B) were taken for dating.

T6 T6 T6 T5 T4 T3 T2 T6 T5 T4 T4 T4 N T3

4km

Figure 4.2 Example 3D-view of a combined slope-elevation map (based on the Aster GDEM) in the eastern part of theMenyuan basin (see the location in Fig. 4.1), showing terraces T2, T3, T4, T5 and T6, along the Datong valley.

2. The HC section is located at a terrace ~7 m above the present floodplain (altitude 3120 m) of the tributary Lairitu River (Fig. 4.6). The surface of this terrace declines gradually to connect with the T4 terrace (21-23 m apf) of the Datong River (Fig. 4.6). The section is composed of a fining-upward fluvial sediment series, covered by a black-brown sandy soil; four OSL samples (HC-1 from covering soil, HC-2, 3 and 4 from fluvial sand and silt) were taken from this section (Fig. 4.7). The fluvial deposits consist of gravel and coarse sand, which are interpreted to be channel deposits deposited under conditions of waning flow by accretion of progressively smaller clasts and overbank fine sand and silt. The fluvial deposits are strongly involuted and subvertical microjoints are commonly present, which probably originated within a perennially frozen sediment (e.g., Mol et al., 1993). The amplitude of these involutions is >4 m. The involutions include upward pointed and lobate protrusions of gravels into fine sands and silts, and downward protrusions of sands and silts into gravels. Elongate clasts

98 with the long axes nearly vertical in the gravel layer and fine vertical laminations in sand and silt layer commonly occur (Fig. 4.7). By analogy to similar structures (e.g., Vandenberghe, 1988, 1992), the involutions of this size are interpreted as structures that are due to periglacial deformation in degrading permafrost (type 2 cryoturbations).

(A) (B)

flood loam wide shallow channel with sand dunes T3 T2 scour T1

T6 ( C ) 180 ± 46 ka DZ-1 OSL sample 105

Gravel 90 200 ± 27 ka Loess DZ-23 DZ-28 131 ± 25 ka 75 Altitude above water level (m) level water above Altitude Sand and silt ?

T5 60

30 T4 11 ± 1 ka T3 QSZ-1 15 T2 Datong River ? T1 ? 0 Figre 4.3 (A) Photos showing topographic distribution and (B) sediment structure of the terraces; (C) terrace sequence of the Datong River along a schematic section in the Menyuan basin.

(A) (B)(b)

GSK-1 21 ± 2 kaGSK-1

GSK-2GSK-2 13 ± 1 ka

((b)©(b) C)

GSK-3 GSK-3 98 ± 15 ka

Figure 4.4 (A) Morphologic character (U-shaped valley and moraine hill surface) of the GSK area. The black dot is the site of the studied sediment sequence. (B) Till and underlying soil layer. (C) Fluvioglacial deposits at the lower part of the GSK section. The white dots show the position of the OSL samples. The white coarse line is a hammer, ~20 cmlong, as scale. The sample sites for GSK-1 and GSK-2 are ~4 m horizontal distance.

3400

GSK site in pla ash utw

) O 3200

Fa n tude (m tude i t l

A 3000 Glacio-fluvial deposit

T5 Datong River

2800 A0 5 10 15 20 25 B Distance (km)

Figure 4.5 Topography along A-B profile derived from the ASTER Global Digital Elevation Model (ASTER GDEM) shows the position of outwash fan and its deposits on top of the glaciofluvial deposits. The profile shows also the connection between the glaciofluvial deposits and terrace T5.

100

3150

Lairitu river HC site

3100 T4 river Datong Top of fluvial gravel

Altitude (m) 3050

3000 C0 1000 2000 3000 4000 5000 (m) D

Figure 4.6 Topography along the C-D profile, derived from the ASTER Global Digital Elevation Model (ASTER GDEM) illustrating the fluvial deposits at the HC section and the connection of these deposits with the T4 terrace of the Datong River.

3. Near the Dazuang (DZ) section, four fluvial terraces (T1, T4, T5 and T6) of the Datong River and onemoraine ridge occur (Fig. 4.8). The surface of the ridge declines gradually, and connects with an outwash plain at a sharp slope change. The outwash plain declines gradually to connect with terrace T6 (80 m apf of the Datong River)(Fig. 4.8). The ridge is mainly composed of diamicton gravels and sands with similar sedimentary structure as in the second unit of the GSK section which is interpreted to be a till. Below this moraine, the section consists of poorly sorted, ungraded, occasionally imbricated, sub-angular to sub-rounded gravels (5-10 cm diameter). Interfingering sand layers occasionly occur. These sediment properties are in agreement with a glacio-fluvial origin similar to the fourth unit of the GSK section. Thick floodloams (3.2 m) with silt and brown soil layers were deposited on the sixth terrace (T6), covered by 34 m loess; sediments were sampled for OSL-dating from the loess (DZ-1) and from the floodloams (3 samples) (Fig. 4.8).

101

(A) (B)(b)

2.7 ± 0.4 ka HC-1

22 ± 2 ka HC-2 HC-3 23 ± 2 ka HC-4 26 ± 2 ka

Fiure. 4.7 (A) Sediment sequence and some parts of the large cryoturbations with clear upturned sedimentary structures at Huangchen (HC) section. The dotted white line shows the boundary between gravel layer and sand layer. The white circle shows a person for scale in (B). (B) One wave of the involutions and the position of the OSL samples (white dots) at HC section.

4.4.3. Dating results From all samples, a small but sufficient amount of very fine-sand quartz could be extracted to permit standard SAR-OSL analysis. The dose response can be well represented by either a single saturating exponential or the sumof a single saturating exponential and a linear component (Fig. 4.9). In general, sensitivity changes occurring during the measurement procedure were accurately corrected for (as indicated by recycling ratios within the range 0.90- 1.10) and the growth curves pass very close to the origin (recuperation values well below 1% of the corrected natural OSL signal). Of 163 large aliquots measured in total, ten were rejected on account of a recycling ratio and three on account of an OSL IR depletion ratio >10% from unity. For most samples, an acceptable-to-good dose recovery was achieved (Table 5.1; measured to given dose ratios within 10% from unity); it may not be possible to measure given doses in excess of a few hundred Gy with the same degree of accuracy (Table 5.1, dose recovery data for samples DZ-23 and DZ-28). The average large aliquot quartz doses are summarized in Table 4.1. The results of the measurements using small aliquots of samples GSK-1 and -2 are shown as histograms in Fig. 4.10. The small amount of quartz

102 available limited the analysis to 30 small aliquots for each sample; six aliquots of sample GSK-1 and 14 aliquots of sample GSK-3 were rejected on account of a recycling ratio or an OSL IR depletion ratio not within 10% from unity. The distributions are broad, with relative standard deviations in the range of ~16 to 26%, and display little asymmetry. The unweighted average De's are somewhat lower than the values obtained using large aliquots (Table 4.1), but it cannot be excluded that this relates to the limited number of small-aliquots. Broad distributions, showing little or no asymmetry, have previously been reported for well-bleached samples (see e.g., Buylaert et al., 2009; Derese et al., 2009; Vandenberghe et al., 2009); the distributions observed for samples GSK-1 and - 2 are thus not incompatible with a well-bleached nature. Table 4.1 summarizes the analytical data and OSL ages. The uncertainties on the OSL ageswere calculated following the error assessment system proposed by Aitken and Alldred (1972) and Aitken (1976). As the ages were obtained using large aliquots, they should at least in principle be considered as maximum ages. The dose distributions, although based on a small number of observations, do not hint incomplete resetting as a significant source of error, however. Also, it has been argued (see e.g., Murray and Olley, 2002; Jain et al., 2003) that incomplete resetting is unlikely to give rise to significant age overestimations for samples older than a few ka, even in glacio-fluvial environments. Thus, like common practice, the mean equivalent doses (ED) and associated standard error were used to estimate the paleodose. The OSL ages are generally consistentwith the stratigraphic position of the samples, except at the GSK section. From top to bottom the deposit sequence is composed of a till on top of glacio-fluvial sediments, separated by a palaeosol at the GSK section (Fig. 4.4B, C). This sequence may represent a glacial- interglacial-glacial cycle. The sample GSK-1 from the till is dated at ~20 ka, indicating a glacial advance during the LGM. In accordance with this result, the underlying soil could be of last interglacial age and the underlying glacio-

103 fluvial sediments could date from the penultimate glaciation. However the OSL dating for GSK-2 is around 13 ka.

(A)

( C ) DZ-1 180 ± 46 ka 3500 DZ-23 (B) 200 ± 27 ka 131 ± 25 ka 3300 DZ-28 de (m) de

u 3100 ltit A

2900 T6

T5 River Datong

2700 2.5 57.5 10 12.5 E02.557.51012.5F Distance (km) Figure 4.8 Topography along the E-F profile, derived from the ASTER Global Digital Elevation Model (ASTER GDEM) illustrating moraine hill, glaciofluvial deposit and fluvial terraces connection at Dazhuang (DZ) section,with photos of the (A) moraine-hill morphology and (B) sedimentology of the glaciofluvial gravel sheet and (C) overlying flood-loam and loess of terrace (T6).

These two OSL-dates (GSK-1 and GSK-2) are difficult to reconcile on the basis of physical arguments. Nevertheless, we are inclined to accept the 21 ka date for two reasons. First, the age of sample GSK-1 is consistent with the field deposit sequence (last glacial) and 14C ages, reported by Guo et al., (1995), in which one sample of disseminated organic carbon was dated at ~25 14C ka BP and one sample using inorganic carbon from calcium carbonate from sandy silt was dated at ~31 14C ka BP. Second, this moraine occurs at the most exterior position of the nested moraines downstream in the glacial valley, coinciding with the largest extent of the glacier. Usually this position is reached near the maximum of the glacial period. In contrast, the tills deposited at ~13 ka (well

104 after the LGM, 19-23 ka) usually occur in a much more interior position in the glacial valley in many regions in the Tibetan Plateau (e.g., Lehmkuhl and Owen, 2005; Liu et al., 2006; Kong et al., 2009; Zhou et al., 2010). Therefore, although we have no real physical argument to reject the dating of sample GSK-2 (13 ka), we favour the age of GSK-1 (21 ka) because it fits better in the context of a general glacial evolution and the published 14C ages. Further work is needed to clearly resolve the discrepancy. The dating result of sample GSK-3 seems to be somewhat underestimated when considering that it was collected from a body of sediment underlying the last-interglacial soil. It has been shown that OSL ages derived from the high- dose region of the dose response curve can be inaccurate, with Des being underestimated from about 150-200 Gy onwards (e.g., Murray and Funder, 2003; Buylaert et al., 2007). For the same reason, the results of the samples from the T6 terrace at the DZ section should also be interpreted with caution. In any case, these OSL ages demonstrate that these samples are quite old. Since these samples were taken from the loess and floodloam overlying the fluvial gravel deposits, it indicates that the Datong River could have aggraded already before the penultimate glacial (Table 4.2). The sample (QSZ-1) from floodloam on top of the T3 terrace is dated at around 11 ka, indicating that the Datong River possibly also aggraded during the Younger Dryas stadial., Three samples (HC-2, -3 and -4) from fluvial sands involuted in the fluvial gravel which correlates to Datong terrace T4 are dated at around 20-25 ka. This may point to fluvial aggradation of the Datong River during the LGM. Thus, these OSL dating results at three sections are generally consistent with their geomorphologic position (see the correlation of the terrace sequence as summarized in Table 4.2).

4.5. Discussion 4.5.1. Morphological response to climate conditions

105

Based on our OSL results and their interpretation,we conclude that a moraine was constructed at ~21 ka, indicating a considerable glacial advance during the LGM in the Qilian Mountains, and possibly correlating to LGM glaciations on the Tibetan Plateau itself (Lehmkuhl and Haselein, 2000). Any MIS 3 advance in the Qilian Mountains must have been smaller than the LGM advance. This chronology differs from that of the southern Tibetan Plateau as established by Owen et al. (2008) in that there the LGM advances were smaller than the MIS 3 advances. A glacial advance during MIS3 when Asian monsoon enhanced in the southern Tibetan Plateau has been considered as the response to changes in insolation-controlled monsoon precipitation (Owen et al., 2008; Owen, 2009). But glaciers advanced extensively in the southern Qilian Mountains at ~21 ka, at the same time as in the northern hemisphere including the eastern Tibetan Plateau, reaching their maximum extent since the late Pleistocene (Ou et al., 2012). This would suggest that the factor controlling glaciation in the Qilian Mountains is similar to that in the eastern Tibetan Plateau. It means that the general temperature decline in the Northern Hemisphere during MIS-2 seems to have been more important than any precipitation increase during MIS-3 in causing glaciers to grow (Ou et al., 2012). However, this conclusion remains to be confirmed through additional absolute dating of tills that belong to the maximum glacier extent. Fluvial sediments deposited during LGM at the HC section (T3) and penultimate glacial at the GSK section (T5) indicate that fluvial aggradation in the Menyuan basin occurred mainly during cold periods. This link between fluvial deposition and cold environment might also be supported by the aggradation of T4 that occurred at ~11 ± 1 ka, perhaps during the Younger Dryas. Because this is based on a single age with associated uncertainty, this last conclusion is not definitive. Besides, paleo-precipitation conditions in the

106 northeastern Tibetan Plateau during the Younger Dryas are not well known (Zhou et al., 2001; Shen et al., 2005).

Figure 4.9 Representative growth curve for a single aliquot of very fine quartz sand (63-90 μm) extracted from samples GSK-1 (A) and HC-4 (B). The inset shows the natural and regenerated OSL decay curves.

107

6 8 (A) (B)

3 4 netit (Gy) Uncertainty

0 Uncertainty (Gy)0 NJU-588 NJU-589 RSD: 26% RSD: 16% n=24 n=16 4 4

2 2 Number of aliquots Number of aliquots

0 0 0 50 100 02550 Equivalent dose (Gy) Equivalent dose (Gy)

Figure 4.10 Histogram of De distribution: (A) for GSK-1 and (B) for GSK-2 using small aliquots.

The aggradation of the younger terraces (T1 and T2) during the Holocene is exceptional and could be affected by intrinsic-dynamic evolution or tectonic activity as no strong climate fluctions occured during the Holocene. The relationship between fluvial activity and climatic conditions is similar to results from other regions as summarized by Vandenberghe (1995, 2002) and Bridgland and Westaway (2008). In contrast, on the southern Himalayan front, it has been argued tht relatively high precipitation caused river aggradation during MIS 3 and 2 since the ITCZ migrat-ion to somewhere on the lower heights of the Lesser Himalayan zone should have induced strong precipitation there even during the LGM (Ray and Srivastava, 2010). Prell and Kutzbach (1987) modeled rainfall over the Indian subcontinent during the LGM and found that precipitation was reduced to ~70% of present-day values, yet even that lowered amount is thought to have been enough to mobilize sediment and

108 cause aggradation in the valleys (Srivastava et al., 2008; Ray and Srivastava, 2010). Precipitation (in the form of snow) might also have been sufficient during cold periods in the northeastern Tibetan Plateau (Yang and Scuderi, 2010) to affect the fluvial aggradation in the Qilian Mountains, as in the southern Himalayas, especially during the summer with large amounts of melt- water. We conclude that fluvial aggradation took place in the Menuyan basin especially during cold climates. 4.5.2. Paleoenvironmental reconstruction At the GSK section the sequence consists of a moraine, a soil (of interglacial type), and glacio-fluvial sheet deposits (sandrs). This sequence probably reflects a temporal evolution from cold and probably moist conditions during the penultimate glaciation, to warm and wet climate during the last interglaciation, and finally to very cold climate with a glacial advance during the LGM. In addition, the fluvial sediment with very large cryoturbation structures around 24-21 ka pointed to the presence of continuous permafrost during the LGM, supporting the opinion that at elevations higher than 1200 m the southern limit of continuous permafrost occurred south of 38°N then (Vandenberghe et al., 2004). By analogy to similar structures in coarse-grained deposits, the involutions of this size were interpreted as type 2 cryoturbations, indicating a mean annual temperature of maximum −8 °C (e.g., Vandenberghe, 1988, 1992, 1993) during the LGM. Based on these estimates, the mean annual temperature difference between the LGM and now could be estimated to be at least 7.2 °C in the Menyuan basin. This was consistent with the early view that the temperature was 6-9 °C lower during the LGM than at present in the Tibetan Plateau as derived from changes of the snow-line (Shi, 2002, and references in there). Although this temperature difference was smaller than 13 °C as reconstructed from sand wedges in adjacent regions (Wang et al., 2001; Vandenberghe et al., 2004), the value of 7.2 °C was only a minimum value which is in any case much higher than the 2.8 °C difference assumed at

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Sample Sample Depth U Th K W.C. dose rate D Age D.R. e lab-code field-code (cm) (ppm) (ppm) (%) (%) (Gy ka-1) (Gy) (ka)

NJU-719 HC-1 80 2.26 ± 0.04 12.80 ± 0.03 2.08 ± 0.03 25 ± 6 2.99 ± 0.02 1.04 ± 0.03 (6) 8.0 ± 0.2 (24) 2.70 ± 0.4

NJU-720 HC-2 230 1.91 ± 0.04 8.24 ± 0.03 1.71 ± 0.03 15 ± 4 2.55 ± 0.03 1.09 ± 0.02 (6) 55 ± 1 (18) 22 ± 2

NJU-721 HC-3 320 1.72 ± 0.05 7.25 ± 0.03 1.60 ± 0.03 12 ± 3 2.39 ± 0.03 1.06 ± 0.02 (6) 55 ± 2 (17) 23 ± 2

NJU-722 HC-4 430 1.92 ± 0.04 8.54 ± 0.03 1.63 ± 0.03 12 ± 3 2.52 ± 0.03 1.16 ± 0.09 (6) 67 ± 1 (22) 26 ± 2

NJU-723 QSZ-1 110 3.20 ± 0.04 11.90 ± 0.03 1.70 ± 0.03 18 ± 5 2.97 ± 0.03 0.95 ± 0.05 (6) 34 ± 2 (17) 11 ± 1

NJU-588 GSK-1 250 2.78 ± 0.04 12.80 ± 0.03 2.03 ± 0.03 13 ± 3 3.33 ± 0.02 1.00 ± 0.06 (6) 69 ± 3 (19) 21 ± 2

NJU-589 GSK-2 350 3.06 ± 0.04 12.50 ± 0.03 1.92 ± 0.03 15 ± 4 3.19 ± 0.02 0.99 ± 0.04 (6) 40 ± 1 (16) 13 ± 1

NJU-590 GSK-3 450 2.35 ± 0.05 12.10 ± 0.03 2.19 ± 0.02 12 ± 3 3.29 ± 0.02 0.88 ± 0.07 (6) 321 ± 42 (8) 98 ± 15

NJU-699 DZ-1 70 2.36 ± 0.04 10.80 ± 0.04 1.86 ± 0.03 15 ± 4 2.96 ± 0.03 0.97 ± 0.10 (3) 533 ± 127 (3) 180 ± 46

NJU-710 DZ-23 2570 2.86 ± 0.04 12.60 ± 0.03 2.02 ± 0.03 15 ± 4 3.03 ± 0.03 1.40 ± 0.12 (3) 604 ± 48 (3) 200 ± 27

NJU-713 DZ-28 2800 3.09 ± 0.04 11.80 ± 0.03 1.93 ± 0.03 15 ± 4 2.95 ± 0.03 1.40 ± 0.06 (3) 387 ± 60 (3) 131 ± 25

※ NJU-588 GSK-1 250 2.78 ± 0.04 12.80 ± 0.03 2.03 ± 0.03 13 ± 3 3.33 ± 0.02 62 ± 3 (24) 19 ± 1

※ NJU-589 GSK-2 350 3.06 ± 0.04 12.50 ± 0.03 1.92 ± 0.03 15 ± 4 3.19 ± 0.02 34± 2(16) 11 ± 1 Table 4.1 Summary of U, Th and K concentrations, estimates of past water content (W.C.), calculated dose rates, ratios of measured to given dose (dose recovery test; D.R.), equivalent doses (De). The number of aliquots used to obtain D.R. and De data is given between brackets in the subscript. The samples with asterisk mark were measured using small aliquots and others were measured using large aliquot.

110 the Qinghai Lake by Porter et al. (2001). The presence of cryoturbations and absence of sand wedges in the Menyuan basin contrasted with the general presence of sand-wedges and absence of cryoturbations in adjacent regions (Wang et al., 2003; Vandenberghe et al., 2004). This indicated that it was more humid in the Menyuan basin than in adjacent regions such as the Mu Us desert and Hexi Corridor during the LGM.

Terrace in the basin Sediments near mountain MIS

sequence Altitude (m) GSK HC HTS

Soil Soil MIS 1

Younger T3 + 14 Dryas

Involuted + 21-23 MIS 2 T4 Till fluvial Till?

sands

Soil? MIS 5?

Glacio- fluvial T5 + 44-46 MIS 6 gravel and sands

Glacio- fluvial T6 + 80 >MIS 6 gravel and sands

Table 4.2 Terrace sequence in the basin, related sediments in the mountain footwall, and their correlation to global chronostratigraphy as represented by the Marine Oxygen Isotope Stages (MIS).

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4.6. Conclusions Quartz-based SAR-OSL analysis yielded ages that are generally consistent with the stratigraphic and geomorphological position of the samples. The results demonstrate that the method could provide robust age control for glacial and fluvial deposits in the Menyuan basin in the northeastern Tibetan Plateau, which can be a vital solution for the lack of suitable materials for radiocarbon and terrestrial cosmogenic nuclide exposure dating in this region. Although further work at more localities is required, our finds suggest that glaciers advanced during LGM, indicating that the temperature decline was the critical factor for glacial fluctuations in this region as in the eastern Tibetan Plateau, rather thanmonsoon-related precipitation. Because fluvial aggradation occurred during the penultimate glacial, the LGM and the Younger Dryas, as the general model in temperate and cold regions (last modified by Vandenberghe, 2008), the increase of sediment supply in periglacial conditions was a controlling factor rather than monsoon-related precipitation raise in the Menyuan basin the northeastern Tibetan Plateau. Periglacial structures in the Menyuan basin show that the mean annual temperature was >7 °C lower during the LGM than at present and that conditions there were more humid than in adjacent regions. Extensive quartz-based SAR-OSL datings of till, fluvial and fluvio-glacial deposits in the Tibetan Plateau where it is sensitive to global climate could provide vital clues to unravel the detail relationship between earth surface processes (glacial fluctuations, fluvial erosion and deposition) and climate changes.

Acknowledgments This researchwas supported by the National Natural Science Foundation of China (grant nos. 41271001, 40901002 and 41230526), PhD Programs Foundation of Ministry of Education of China (20090091120032), and the Dutch-Chinese CAS-KNAW PhD Project (05-PhD-10) and Exchange

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Programs (08CDP028, 09CDP033 and 10CDP036). Prof. Pan Baotian and Prof. Yang Dayuan are thanked for constructive discussions. Miss. Zhai Min, Mr. Zhang Wenchao, Mr. Yang Chuanbin and Dr. StefanVasiliniuc are thanked for field and laboratory assistance. DV is a Post-doctoral Fellow of the Research Foundation Flanders (FWO-Vlaanderen). The comments by Prof. Alan R. Gillespie and Prof. Lai Zhongping and two anonymous reviewers are highly appreciated.

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- Chapter 5- Differential tectonic movements in the confluence area of the Huang Shui and Huang He rivers (Yellow River), NE Tibetan Plateau, as inferred from fluvial terrace positions Xianyan Wang, Ronald van Balen, Shuangwen Yi, Jef Vandenberghe and Huayu Lu

Abstract: In the Northeastern Tibetan Plateau, the courses of the Huang Shui river and the Huang He (Yellow river) near their confluence are characterized by alternating gorges and wide depressions, segmenting the fluvial systems. The river valleys have developed terrace staircases which are used to infer relative tectonic motions between the segments. The terrace staircases in the different segments are correlated by means of their relative height and OSL dating. At least eight fluvial terraces are present, two of which have been dated by OSL (the sixth and the third ones; -70 ka and -24 ka respectively). The correlated longitudinal terrace profiles show that within the confluence area no distinct relative tectonic movements have taken place since the seventh terrace demonstrating that this area behaved as one large tectonic block. The terrace staircase correlation of this block to areas further upstream (Xining area) and downstream (eastern Lanzhou area) indicates relative tectonic motions with respect to those areas, which must therefore represent different tectonic blocks. The average river incision rate since -70 ka was much higher in the confluence region block than in the up and downstream blocks, possibly indicating relative uplift of the confluence region. This relatively strong uplift gives more space for differentiation within the terrace staircase as a result of climatic changes, leading to 6 terraces formed as a response to small climatic fluctuations (103- 104 year timescale) since the last inter-glacial., This may indicate that the stronger the tectonic movement the better the climatic imprint is expressed in the terrace development. On a smaller time scale, two accumulation terraces

121 with thick stacked fluvial deposits (>18 m) which are not related to changes of sediment supply due to other than climatic impact, may indicate relative subsidence in the confluence area relative to the upstream and downstream blocks, occurring at some time between 20 and 70 ka. This indicates changes in relative vertical crustal motions at timescales of tens of thousands of years. We speculate that the inferred tectonic motions are related to the transpression movements in the Northeastern Tibetan Plateau as a result of the collision of the Indian and Asian plates.

Based on: Boreas, 2013, DOI: 10.1111/bor.12054.

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5.1. Introduction As rivers are very sensitive to changes in gradient and respond very rapidly to tectonic movements, river morphology and sedimentology are widely used to study neotectonic deformation (e.g., Burbank and Anderson 2001; Westaway et al., 2002). For example, river terraces, which are typical morphological features of fluvial landscapes, are widely used as morphotectonic markers to study the magnitude and rate of tectonic uplift (e.g., Li et al., 1996; Maddy 1997; Van Balen et al., 2000; Burbank and Anderson, 2001; Lu et al., 2004a; Peters and Van Balen, 2007; Viveen et al., 2012). Extensive terrace staircases are apparently formed exclusively in uplifted regions (e.g., Bridgland et al., 2007; Bridgland and Westaway, 2008). However, climate changes are thought to be a vital factor in controlling fluvial incision and aggradation cycles and, therefore, terrace formation (e.g., Vandenberghe 1995; Li et al., 1996; Lu et al., 2004a; Lewin and Gibbard, 2011). Still, differential tectonic motions may complicate this simple pattern. For example, it is recently reported that areas of fluvial incision and -aggradation spatially alternate in the Huang Shui catchment (northeastern Tibetan Plateau; NETP, Fig. 5.1A, B, 5.2), indicating the effects of relative uplift and subsidence on relatively small spatial scales (Wang et al., 2010; Vandenberghe et al., 2011). In this paper we extend this work. We use terrace dating and correlation to infer tectonic motions between different blocks along the Huang Shui and part of the upper reaches of the Huang He (Yellow River; downstream of their confluence). At a large scale, the incision of the Huang He has been used to study timing of crustal uplift and subsequent erosion in the northeastern Tibetan plateau (e.g., Li et al., 1996; Lu et al., 2004a; Craddock et al., 2010; Wang Erqi et al., 2011, Wang et al., 2012). In this study, however, we focus on specific parts of the Huang He system and use field observations and topographic data, combined with Optically Stimulated luminescence (OSL) dating, to determine the fluvial evolution in relation to local tectonic movements since the late

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Quaternary. More specifically, we correlate and compare the morphological evolution of the Huang Shui catchment with that of the Huang He in its confluence area near Lanzhou (Fig. 5.1). The new results are integrated with published data from the adjacent regions (Wang et al., 2008a, b; Vandenberghe et al., 2011; Pan et al., 2009).

5.2. Geological and geomorphological setting The study region is situated in the NETP, defined as the region bounded by the Kunlun fault to the south and the Altyn and Haiyuan faults to the north (Fig. 5.1A). In the NETP coeval crustal shortening and left-lateral strike-slip faulting are related to the northeastward growth of the plateau margin as a result of the collision of the Indian and Asian plates. Crustal deformation primarily due to these relative plate motions is expressed by folding and faulting of the Paleozoic to Cenozoic bedrock and inversion of syn-orogenic basins, such as the Xining basin (Fig. 5.1). The deformation at the NETP and the activation of the major strike-slip faults are supposed to have started just a few million years after initial India- Asia continental collision despite the >3000 km distance from the plate boundary at c. 45-50 Ma (e.g., Yin and Harrison, 2000; Dupont-Nivet et al., 2008; Clark et al., 2010; Duvall et al., 2011). The larger syn-orogenic basins were filled with thick continental clastic red beds. Subsequently, the secondary strike-slip faults, such as the Riyueshan fault (Fig. 5.1), and related basement rooted anticlines started to form, possibly following the early Miocene, due to continuous activation of these major boundary faults and the related clockwise rotation of the basins (e.g., George et al., 2001; Zheng et al., 2006; Dupont- Nivet et al., 2008; Zhang et al., 2009; Craddock et al., 2011). Local-scale uplift compartmentalized the larger syn-orogenic system. Accelerated general uplift accompanied with widespread deformation during the late Miocene resulted in river incision into a regional peneplain and the formation of alternating units, of

124 limited extent, displaying relative subsidence and relative uplift (e.g., Lu et al., 2004a; Pan et al., 2010; Wang et al., 2012). Subsequently, the deeply incised rivers constructed well-formed staircases of terraces (clearly distinguished on the basis of their elevation and widely developed steps in the morphology) (Fig. 5.2) (e.g., Lu et al., 2004a; Wang et al., 2010; Vandenberghe et al., 2011). The basement in all the discussed regions consists of Proterozoic gneiss and schists, Cambrian grey limestones and basalts, and Mesozoic clastic sediments (Fig. 5.1B) (Qinghai Bureau of Geology and Mineral Resources 1991). Cenozoic successions of playa to fluvio-lacustrine environments overlie Cretaceous clastic sediments concordantly or older basement rocks discordantly. They have been subdivided into the Xining and Guide Groups (Qinghai Bureau of Geology and Mineral Resources 1991). In addition, ‘fanglomeratic’ rocks unconformably overlie the Cenozoic deposits (Qinghai Bureau of Geology and Mineral Resources 1991). The Quaternary sediments are fluvial deposits of the Huang Shui and Huang He rivers and loess blanking the landscape (e.g., Lu et al., 2001).

5.3. Methods 5.3.1. Mapping and field work First, large-scale landforms were mapped using remote sensing data from various sources. The principle dataset was the ASTER Global Digital Elevation Model (ASTER GDEM) with a horizontal resolution of 30 m. Next, a programme of targeted field mapping was carried out using GPS and 1:50,000 topographic maps. Field work was undertaken to identify tectonic structures (faults), and to map individual terraces and their distribution (e.g., Fig. 5.3). The altitude of the top of the terrace gravel deposits with respect to the present floodplain has been used for correlation. In fact, terraces are typically covered and thus preserved by loess deposits (see below). Where accessible, outcropping fluvial deposits were described with regard to their sedimentary

125 structures (using lithofacies according to Miall 1978), lithological properties, and thickness, in order to determine the sedimentary processes responsible for their emplacement.

5.3.2. Optically stimulated luminescence dating In order to compare the terrace ages, we have taken a set of five samples for OSL dating from fine sand layers and overlying loess, from the T3ML and T6ML terraces at three locations: Renhe (103.30737°E, 36.15883°N), Dazhuang (102.84357°E, 36.33042°N), and Tuanjie (102.87001°E, 36.31875°N) (see locations in Fig. 5.5A). Sediment samples were collected by hammering light-tight tubes into freshly exposed sediments. The tubes remained sealed until processed in safe light conditions. One sample could not be measured due to the small grain size and lack of quartz. All measurements were performed at Nanjing University (China). An additional set of published OSL data is also used in this study (Wang et al., 2007, 2008a, b). Fine sand-sized (63-90 µm) quartz grains were extracted from the inner material of the sampling tubes in the conventional manner (10% HCl, 30%

H2O2, wet and dry sieving, density separation using heavy liquids, 40% HF). Initial screening suggested that the luminescence sensitivity was low. OSL dating results using different size aliquots for fluvial sands in the studied region show no apparent difference, which indicates that incomplete resetting is unlik- ly to be the cause of significant age overestimations in the region (Wang et al., 2013). Thus, in common with other OSL dating studies in the Tibetan Plateau (Lai et al., 2009; Liu et al., 2010), we analyzed all samples using large aliquots (5 mm diameter). A monolayer of quartz grains was spread out on the inner 5 mm of 9.7 mm diameter stainless steel discs using silicone spray oil as adhesive. Luminescence signals were measured on a Risø TL/OSL-DA-20C/D reader fitted with blue-green diodes (470±30nm; 40 m W cm-2) and an IR laser diode (830 nm). The luminescence signal was detected through a 7.5mm Hoya U-340

126

Siberia

Fig.4.1A North Ta r i m China au t e la P South an et China lt Tib a u yn F ( C ) A lt Alashan Block Ordos India M Indouchina . Qi 38 lian shan 1 Block (A) Q a H id Qinhai a a iy m u B Lake B. XINING a n a Fig.5.1B 36 si F 2 n a u Kunlun Fault lt 3

Strike-slipe fault Thrust fault 34 0 100 200 km Normal fault 4 86 90 94 98 102 106 (B)

37 Q inh H 6 ai uang S La hui ke River XN Fig.4 7 Quat. sed. Tert. rock PA Extrusive rock LD e MH H Plutonic P rec.-Mesoz . Fault rock rock . LZ a 0 15 30 km u N 36 H Fig.1C 101 102 103 0 5 10 km

127 filter. Irradiation was carried out using a 90Sr/90Y beta source with a dose rate at 0.149 Gy s-1 built into the reader. A single-aliquot regenerative-dose (SAR) protocol (Murray and Wintle, 2000) was used to determine the equivalent dose (De). Preheating of natural and regenerative doses was for 10s at 260°C, while the response to the test dose was measured after a cut-heat of 220°C. Optical stimulation with the blue diodes was for 40 s at 125°C. The initial 0.8 s of the decay curve was used in the calculations, minus a background evaluated from the following 0.8 s. After measuring the response to the test dose, a high-temperature bleach was performed by stimulating with the blue diodes for 40 s at 280°C (Murray and Wintle, 2003). For each aliquot, the dose response was obtained by measuring the response to three regenerative doses. This was followed by two additional measurements to obtain estimates of recuperation and recycling (Murray and Wintle, 2000).To assess the reliability of the measurement procedure, a dose recovery test was performed (Murray and Wintle, 2003). Fresh aliquots were bleached for two times 250 s using the blue diodes at room temperature; the two bleaching treatments were separated by a 10,000 s pause. A regenerative dose was then given, chosen so as to be equal to the estimated natural dose. The material from the outer ends of the sampling tubes was retained for radioisotope and water content analysis. The in situ water content and that at saturation were determined by drying sub-samples in an oven at 50 °C. A subsample of ~20 g dried sediment was ground to a fine powder to determine concentrations of U, Th, and K by neutron activation analysis (NAA). The elemental concentrations were converted into external beta and gamma dose rates using the conversion factors of Adamiec and Aitken (1998) and the beta attenuation factors of Mejdahl (1979). The water content may have varied over the burial period of the samples, but it is difficult to determine the degree of such changes. We assumed 25% of the saturation water content as representative for the time-averaged degree of wetness. In addition, we assigned a relative uncertainty of 25% to this value to allow for possible fluctuations. An internal dose rate in quartz grains of 0.013±0.003 Gy ka−1 was assumed (based on Vandenberghe et al., 2008), and the contribution from cosmic rays was calculated following Prescott and Hutton (1994).

Figure 5.1 (A) Simplified map of major tectonic boundaries and Cenozoic faults in the northeastern Tibetan Plateau (modified after Dai et al., 2006; Vandenberghe et al., 2011). (B) Simplified geological map of the Huang Shui catchment (the dashed boxes are the boundaries of Fig. 5.1C and Fig. 5.4). (C) Morphologic characteristics of the confluence region of Huang Shui and Huang He rivers derived from ASTER GDEM data, illustrating the wide depressions and separating gorges (Dashed lines: boundary of the depression; 1: Minhe depression; 2: Bapan gorge; 3: Hekou depression; 4: Hutou gorge; 5: west Lanzhou depression; 6: Jingchengguan gorge; 7: east Lanzhou depression).

128

NW SE 300

271 m

250

200 m 197 m 200

181 m apf)m Altitude (

167 m

150 144 m

124 m

97 m 100 m 100

60 m 50 44 m

T1d T1c T1b T1a 0 Figure 5.2 Terrace sequence (a composite cross-profile) at the very west of the Minhe depression and the exit of the upstream gorge (LaoYa gorge), showing the relative tectonic movement in these small area. Three accumulation terraces with thick fluvial gravels at -200m, -60m and -40m above the present

5.4. Results Firstly, the terrace sequences in the Minhe depression and in the confluence area of the Huang Shui and Huang He are presented and inter- correlated by means of DEM analysis and field investigations. Subsequently, the OSL dating results are reported. Finally, the studied terraces are morphologically correlated with dated terraces in the West Lanzhou depression (downstream) and in the Xining depression (upstream).

5.4.1. Morphostratigraphy of terraces Based on the DEM topographic analysis (Fig. 5.5) and field morphological investigations, several extended flat surfaces (or terrace remnants) along the Huang Shui and Huang He river valleys in the studied region have been identified at different altitudes above the present floodplain. Field investigations show that these surfaces are usually underlain by fluvial beds, containing gravel from the Huang Shui and Huang He rivers. The gravel layers are parallel to the surfaces, and they are covered by alluvial loams and aeolian loess (Fig. 5.3). The heights mentioned below are the top of the gravel layers with respect to the present-day floodplains.

129

14m

Figure 5.3 (A) A fault below T7ML , at 74m apf at Liuquan (LQ)(see location in Fig. 5.5A )(dashed line: terrace strath plain; solid line: fault), the terrace is not faulted. (B) Terraces at the RenHe (RH) site (see location in Fig. 5.5) showing the the flat plain of T7ML and the small terrace T3ML). (C) Aggradation sequence of T6ML at Tuanjie (TJ) site (see location in Fig. 5.5), showing sediment thickness of >8m in the western part of the Minhe depression (see location in Fig. 5.5). (D) Aggradation sequence of T5ML at section 6, at - 40m apf, with gravel thickness of 14 m, at Jiaojia (JJ) site, in the eastern part of the Minhe depression.

In total, eight regional terraces have been identified in the Minhe-West Lanzhou depressions (the confluence block), at altitudes of about 2, 7-8, 13-14, 20-24, 41-45, 60-66, 76-100 and 130 m above the present floodplain level (apf) (Fig. 5.5). They are labeled T1ML to T8ML. Six additional, local terraces of limited extent are present in the Laoya gorge near the western border of the Minhe depression, which indicate fau-lt movement between the confluence block (Minhe-West Lanzhou depressions) and the upstream Laoya block (Vandenberghe et al., 2011). The terraces T1ML to T7ML in the Minhe depression have been documented before by Vandenberghe et al. (2011). In that paper they have been labeled as terraces T1a, T1b, T1c, T1d, T2c", T2e" and T3a. The different terraces show similar sedimentary sequences. From bottom to top: fluvial gravels of various thickness, intebedded with lenses of sands, silts, clays, and finally topped by a horizontally laminated silt, occasionally containing gravel stringers of limited extent (Fig. 5.3C, D, 5.4). The lowest parts are usually planar sheets of gravel, while sometimes sand-filled shallow

130 channels and lateral accretion sand lenses are interbedded with the gravels (Fig. 5.4D, E). These configurations are suggestive of active channels divided by sand bars. The lower parts are usually covered by multi-storey wide channels with scouring erosional bases; they are filled with trough-cross-stratificatied gravels (Fig. 5.4A, B). The average thickness of the terrace gravel bodies is around 3-6 m, except for terraces T5ML and T6ML, in which the gravel layers are over 15 m thick. The gravels consist mainly of sandstone and crystalline rock (granite) with small amounts of conglomerate and metamorphic rocks; the pebbles are mainly rounded to sub-rounded. The lowest part of the gravel deposits consists of massive gravel (Gm), grading towards the top to become finer-grained and coarsely planar bedding (Gp) (Fig. 5.4). The latter deposits usually show imbrication and small-scale planar or trough cross bedding, indicating deposition in lateral and longitudinal bars. Small and shallow channels occur with increasing frequency towards the top of the terrace deposits; they are filled with cross-bedded fine gravel or sands (planar to low- angle trough cross bedding). The upper laminated silts are sometimes interbedded with small, gravel-filled, very shallow channels (Fig. 5.4C, F). They are interpreted as floodloam deposits that complete the fluvial sequence as formed prior to or simultaneously with the abandonment of the floodplain as a result of renewed river incision. They may have been deposited at peak flooding during the initial phase of subsequent, renewed river incision (Vandenberghe et al., 2011). The gravels of the terraces are uncemented, except for the highest terrace, T8ML in the West Lanzhou depression, the gravels of which are strongly cemented by carbonate. Loess of different thicknesses (related to the age of the terrace) covers these fluvial deposits, except for the lowest terrace. Locally, alluvial fan deposits overlie the T5ML terrace in the West Lanzhou depression. The highest terrace, T8ML, is only present in the West Lanzhou depression. Terrace T7ML is morphologically the most pronounced, being characterized by very wide and continuous flat surfaces, which therefore can be correlated with certainty over large distances. The top of the gravel of T7ML is at 93-100m apf upstream of the Renhe site, while it is at 76-83 m apf downstream of this locality. Terrace T6ML occurs mainly in the very western part of the Minhe depression and in the West Lanzhou depression, whereas T5ML occurs as a small terrace mainly developed and preserved at the edges of T7ML remnants. Together, T7ML and T5ML form a very characteristic morphological double feature. The fourth terrace, T4ML, is mainly present in the Minhe depression. The three lower terraces, T3ML, T2ML and T1ML are widespread along the valley and are well expressed morphologically. The discontinuous presence of T8ML, T6ML and T4ML is probably the result of their initial limited extent and/or erosion.

131

5.4.2. Dating Results In the West Lanzhou depression, two terraces at the Fanjinping (FJP, 103.66743°E, 36.08229°N) and Shajingyi (SJY, 103.60954°E, 36.14058°N) sites (see locations in Fig. 5.4A) have been dated previously by means of OSL (Wang et al., 2007, 2008a ). These two sites occur at T6ML and T3ML, respectively (Fig. 5.4A). At FJP, one sample from sands in a shallow channel is dated at 81±6.2 ka, and five samples from floodplain deposits (laminated fine sands and silts) gave ages of 74-70 ka. The covering loess deposits (sometimes interfingering with colluvial deposits) are dated at 57-3 ka from bottom to top (Wang et al., 2007; Wang et al., 2008a). The sequence of dates is consistent. In terrace SJY, two samples from the floodloam are dated at 24-20 ka, and two samples from covering loess are dated at 13-12 ka (Wang et al., 2008a). This shows T3ML to have been formed during MIS2. At the Renhe section (RH), the sediment sequence of T3ML is made up of 2 m of gravel including one interbedded sand layer near the top of the gravel sequence. The gravel is overlain by 1 m brown-reddish soil, 3 m of silt with small gravel stringers of limited extent and finally 1.5 m of loess (Fig. 5.6). Two samples for OSL dating, RH-1 and RH-2, were taken from the soil and the sand layer respectively (Fig. 5.6). At the Tuanjie section (TJ), the sediment sequence of T6ML is made up of at least 8 m of gravel (the lower boundary is not exposed), overlain by 2 m of fine sand, and 24 m of loess. Two samples, TJ-1 and TJ-2, have been taken for OSL dating from the middle part of the loess and the fine sand layer, respectively (Fig. 5.6). At the Dazhuang section (DZ), the sediment sequence of T6ML is composed of 12 m gravel, overlain by 1 m of floodloam, which in turn is covered by 16 m of loess. One sample, DZ-1, was taken from the lower part of the covering loess, 1.5 m above the floodloam (Fig. 5.6). From all samples except RH-1, a small but sufficient amount of fine-sand sized quartz could be extracted to permit standard SAR-OSL analysis. The sample RH-1 was too fine grained and did not contain enough quartz grains for OSL measurement. The dose response can be well represented by either a single saturating exponential or the sum of a single saturating exponential and a linear component (Fig. 5.7). In general, sensitivity changes occurring during the measurement procedure were accurately corrected for (as indicated by recycling ratios within the 0.90-1.10) and the growth curves pass very close to the origin (recuperation values well below 1% of the corrected natural OSL signal). For most samples, an acceptable to good dose recovery could be achieved (Table 5.1; measured to given dose ratios within 10% from unity).

132

(A) (B)

( C) (D)

(E) (F)

Fiure. 5.4 Characteristic sedimentary structures of terrace deposits (A) planar sheets of gravel covered by finer grained, shallow, trough cross-bedded channels (T6ML at Tuanjie site in the western part of the Minhe depression); (B) sedimentrary sequence of planar sheets of gravel, shallow sand-filled channel which is cut by trough cross-bedded channels, and overlying planar gravel sheets (T6ML at west Lanzhou depression; solid line: boundary between sand filled channel and planar sheets of gravel; dashed line: low boundary of trough cross bedded gravel); (C) horizontally laminated silts occasionally containing gravel stringers of limited extent (T4ML at east Minhe depression); (D) small-sand filled shallow channel (bounded by dashed lines) alternating with planar sheets of gravel (T3ML at Renhe site); (E) shallow sand-filled channel (bounded by dashed lines) between thick planar gravel sheets (T5ML at Jiaojia site); (f) small, gravel-filled channel (bounded by dashed lines) in former floodplain deposits at the top of the sediment series (T3ML at Renhe site).

133

(A)

DZ a TJ 1 2

d RH

5 SJY T2 DZ Dazhuang River c T3 TJ Tuanjie 1 7 T4 RH Renhe Topography section JJ 8 T5 JJ Jiaojia 6 Fault LQ T6 LQ Liuquan 9 Mentioned sites T7 SJY Shajingyi e FJP T8 FJP Fanjiaping 0 5 10 km (B) 1875 T7 1850 1900 T7 T6 2 1825 1 3 OSL-DZ-1 T4 T6 1750 T5 T5 1775 T3 1800 T3 T4 T3 T3 T3 ? ? ?

Altitude1725 (m a.s .l ) 1650

1700 0 1000 2000 3000 4000 5000 m 0 2000 4000 6000 m 0 1000 2000 3000 4000 m 1900

1800 1800 1700 T7 6 4 5 T7 T7 T7 T5 1700 1700 T5? 1600 T5 ? T3 T3 T3 T3 T2 Altitude (m a.s.l) (m Altitude T2 1600 1600 1500 0 2000 4000 6000 m 0 1000 2000 3000 4000 m 0 1000 2000 3000 m

T7 OSL smaple 1675 1675 Loess 1900 1625 T6 T5 8 T5 9 Sand and silt 1800 7 1625 T7 134 T3 1575 Gravel T5 T3 T3 1700 T2 T2 T3 1575 1525 Fan deposit

Altitude (m a.s.l) (m Altitude 1600 0 2000 4000 6000 m 0 2000 4000 6000 m 0 2000 4000 6000 m Substract Table 5.1 summarizes the analytical data and optical ages. The uncertainties on the optical ages were calculated following the error assessment system proposed by Aitken and Alldred (1972) and Aitken (1976). As most dates were obtained using large aliquots, they should, at least in principle, be considered as maximum ages. However, incomplete resetting is unlikely to cause significant age overestimations for samples older than a few ka, even in fluvial environments (e.g., Murray and Olley 2002; Jain et al., 2003). In addition, previous experience also indicates that incomplete resetting is unlikely to cause significant age overestimations in the studied region (Lu et al., 2011; Wang et al., 2013). Sample DZ-1, from the lowest part of the loess sequence covering T6ML, is dated at 45 ka. Sample TJ-1, from the middle part of the loess sequence covering the same terrace, is dated at about 23 ka and sample (TJ-2), from the overbank fine sand, is dated at about 70 ka. This compares well to the dating of T6ML in the West Lanzhou basin, where the floodloam was dated at 74-70 ka. In addition, Last Glacial loess sits directly upon terrace deposits of T6ML at both locations. This confirms the correlation of these two terraces based on relative height (Fig. 5.5A, 5.8A) Sample RH-2 from fluvial sands of T3ML in the Minhe depression is dated at about 24 ka and the overlying thin gravel (thickness <20 cm) (Fig. 5.6) is directly covered by a soil layer. This may indicate that the channel filling and floodloam were deposited shortly after 24 ka. This age is in agreement with the age of floodloam of this terrace at the SJY site (20-24 ka; Wang et al., 2008a) in the West Lanzhou depression. This again seems to confirm the terrace correlation between both locations based on relative height (Fig. 5.5A, 5.8B). Thus T3ML was probably formed during MIS2. Previously, the three lowest terraces in the Minhe depression (T3ML, T2ML and T1ML) have been correlated to the three lowest terraces upstream in the Xining depression (T1c, T1b and T1a), based on height above the present floodplain (Vandenberghe et al., 2011). In the Xining depression, T1c is estimated to have been deposited during MIS8. This is because it is overlain by two loess sequences and an interbedded interglacial soil. Twenty OSL dates from the upper 27 m of loess gave ages between 14 and 60 ka according to Buylaert et al. (2008). This loess layer is underlain by a well-developed soil complex of last-interglacial age (S1, MIS5) (Chen et al., 1997; Lu et al., 2001,

Fiure. 5.5 Terrace distribution in the confluence area of Huang Shui and Huang He rivers. One of the prominent morphological features is the flat wide surface of T7ML, with a narrow terrace (T5ML) next to it. This feature is useful for correlation of the terraces. The faults are based on Yuan et al. (2008), Zhang et al. (2009) and Liu et al. (2007). (B) Topographic sections. The data is from ASTER DEM and field GPS measurement. T6ML at section 8 and T3ML at section 9 were dated by Wang et al. (2008) at about 70-74 ka and 22 ka, respectively.

135

2004b; Vandenberghe et al., 2006; Vriend et al., 2011). Below the soil is another sequence of loess, about 60 m thick, which should date from MIS6. The next lower terrace, T1b, is assumed to date from MIS6 (Fig. 5.8B) (Vandenberghe et al., 2011). Terrace T1c has also been dated by OSL by Wang et al., (2008b). According to them, the overlying loess has an age of 74-52 ka and one sample from underlying fluvial sands is dated at 98.9±11.8 ka. However, the dating by Wang et al., (2008b) is less reliable, because the luminescence saturation and the condition of the loess have not been described in that work. Thus, we argue that the dating result (Vandenberghe et al., 2011) for T1c in the Xining depression based on loess-soil stratigraphy and detailed OSL dating is more reliable. At present it is not possible to decide whether the lowest terraces, T1ML in the Minhe depression and T1a in the Xining depression (assigned to MIS2 by Vandenberghe et al., 2011), are of the same age as absolute dating is not available yet.

Total Samp Sam Dep U Th Dose le ple K w.c. D Age th (mg (mg rate D.R. e lab- field- (%) (%) (Gy) (ka) (m) kg-1) kg-1) (Gy ka- code code 1) 1.05 1.8 NJU- 2.84±0 10.9±0 14± 3.02±0 ± 144.03 47.7± DZ-1 15 7±0 698 .1 .3 4 .2 0.02 ±5.6 4.0 .06 ( (13) 12) 1.13 2.0 79.58 NJU- 3.02±0 11.2±0 16± 3.26±0 ± 24.4± RH-2 5.7 4±0 ±3.3 715 .1 .3 4 .2 0.02( 2.0 .06 (14) 12) 1.04 81.07 1.9 NJU- 3.57±0 12.8±0 12± 3.39±0 ± ± 23.9± TJ-1 11 1±0 716 .1 .4 3 .3 0.02 4.3 2.2 .06 ( 12) (11) 1.08 1.5 185.7 NJU- 24. 2.36±0 10.3±0 11± 2.66±0 ± 70.0± TJ-2 6±0 8±5.7 717 5 .1 .3 3 .2 0.02( 5.8 .05 (13) 12)

Table 5.1 Summary of U, Th and K concentrations, estimates of past water content (W.C.), calculated dose rates, ratios of measured to given dose (dose recovery test; D.R.), equivalent doses (De). The number of aliquots used to obtain D.R. and De data is given in brackets in the subscripts.

136

Lithology Lithology Lithology RH site DJ site DZ site 0 0 0

6 8 2 TJ-1 23.9± 2.2 ka

) 16 DZ-1

th (m 4 47.7± 4.0 ka p De 18 RH-1 24 TJ-2 70.0± 5.8 ka RH-2 6 24.4 ± 2.0 ka 24

Loess Silt Sand Soil Gravel Figure 5.6 Stratigraphy of the sediment sequences of T3ML, T6ML and T6ML, at sites Renhe (RH), Dazhuang (DZ) and Tuanjie (TJ), respectively, and OSL dates.

5.5. Discussion 5.5.1. Spatial and temporal variations of tectonic motions In the upstream part of the Huang Shui river catchment, it has been demonstrated by the different terrace staircases in the adjoining areas that the morphological subdivision in alternating depressions (Xining, Pinglian and Ledu depressions, see location in Fig. 5.1B) and gorges is a response to different tectonic movements (Wang et al., 2010; Vandenberghe et al., 2011). Each morphological depression seems to correspond to a different tectonic unit. The bedrock lithology in the depressions consists mostly of Mesozoic-Cenozoic red beds, whereas the gorges are mostly formed by harder Paleozoic rocks (metamorphic and crystalline). Their spatial distribution is the result of tectonic deformation, which is still ongoing as evidenced by the terrace correlation. The resulting large-scale morphology is thus due to a combination of differential erosion and differential tectonic movements. The large-scale morphology in the present study area is also characterized by successive depressions and gorges. Three wide depressions (Minhe, Hekou and West Lanzhou depressions) are

137 separated by two narrow gorges (the Bapan and Hutou gorges) (Fig. 5.1C). However, the terrace distribution illustrated by the longitudinal profile (Fig. 5.8A, B ) shows that no distinct relative tectonic movements have taken place within the area from the Minhe to the West Lanzhou depression, at least since the formation of T7ML, which is older than 70 ka (the age of T6ML). Thus, in the study area, which is centered at the confluence of the Huang Shui and Huang He, the large-scale morphology is the result of differential erosion operating on a tectonically controlled heterogeneous bedrock distribution. Correlation of our results to the area in the upstream part of the Huang Shui (Xining- to Ledu depressions) (Li et al., 1996; Lu et al., 2004a; Vandenberghe et al., 2011) and the area immediately downstream of our study area, the East Lanzhou area (Pan et al., 2009), shows that these three areas have different incision histories (Fig. 5.8C), and thus should be considered as different tectonic blocks. First, the number of terraces and their altitudes are different. Second, terraces at the same apf have different ages (Fig. 5.8C). This concerns also the younger terraces (except maybe the lowermost terrace), indicating that the tectonic movements continued until Holocene time or are still ongoing. Thirdly, the long-term incision rates derived from terrace ages are different for the three areas. In the Xining block, the terrace at 14 m apf is dated to MIS8 (Fig. 5.8C) (Vandenberghe et al., 2011; see discussion above). Thus, the time-averaged incision rate of the Huang Shui river since that time is 0.046-

Figure 5.7 Representative OSL growth curve for a single aliquot of fine-sandy (63-90µm) quartz of sample TJ-1. The inset shows the natural OSL decay curves.

138

(A) (C) Xining basin This studied region East Lanzhou basin

) late Miocene . l .

s 555 . a 455 (m (m

e 2000 ka d a

b 355 tu i t

l 1400 ka

A 255 water level T2 1240 ka T3 T4 T5

c d 155

T6 T7 T8 e 150

1500 1600 1700 1800 20 40 60 80 100 120 km

1050 ka (B) 125

960 ka 100 860 ka ) Altitude (m apf) (m Altitude . l . s . a 75 130 ka (m e

d 70 ka tu i t

l 50 A

50 ka Fp2 Fp3 Fp4 25 Fp5 Fp6 Fp7 MIS 8? 21 ka

1500 1600 1700 1800 1900 20 40 60 80 100 120 km 0 Figure 5.8 (A) Longitudinal profiles of terraces, showing absence of deflections and offsets (dashed black line: faults; see locations of a, b, c, d, e sections in Fig. 5.4A). (B) The distribution of flat wide plains (FP), identified using DEM. These surfaces are underlain by loess with different thickness covering different terraces. The levels FP2, FP3, FP4, FP5, FP6 and FP7 correspond with terraces T2ML, T3ML, T4ML, T5ML, T6ML and T7ML, respectively. (C) Comparison of terrace sequence in the studied confluence region with the upstream Xining-Pingan-Ledu area (Wang et al., 2010; Vandenberghe et al., 2011) and the downstream East Lanzhou area (Pan et al., 2009).

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0.057 m ka-1. In the confluence region, the incision rate is c. 0.64 m ka-1 since about 24 ka, based on the age of T3ML (also at 14m apf) and c. m ka-1 since about 73 ka, based on the age of T6ML (at 60-66 m). The average incision rate in the East Lanzhou block is c. 0.346 m/ka since the last interglacial (130 ka) (Fig. 5.8B) (Pan et al., 2009). Thus, the long-term incision rate in the confluence area is much higher compared to the adjoining areas. Rates of river entrenchment may help to infer the associated speed of tectonic uplift (Maddy, 1997). In theory, the increase of discharge in the confluence area may also explain the increase of the incision rate. However, stream incision rates are not just higher in the confluence site but in the whole studied region, which extends upstream for more than 80 km and downstream for at least 40 km from the confluence site. Thus, the increased discharge cannot explain the higher incision rates. On a shorter timescale the confluence area also experienced relative subsidence (slower uplift rates compared to upstream and downstream areas), as evidenced by two terraces, T5ML and T6ML, extending over more than one hundred kilometers, with anomalously thick fluvial deposits (>18 m) (Fig. 5.2, 5.3C, D). Evidence for increased sediment supply caused by local landslides has not been found. Coeval thick terrace deposits do not occur in the Xining and East Lanzhou depressions (maybe they appear in the Ledu depression; Vandenberghe et al., 2011). This indicates that the long-term relative uplift of our study area was interrupted by two periods of relative subsidence. Thus, the tectonic motions not only differ between the three blocks, but also vary in time. 5.5.2. Climate change and terrace formation In general, fluvial terraces are thought to be formed in cold periods and as a response to glacial-interglacial climate changes, set against background tectonic uplift (Vandenberghe 1995). The causal relation between fluvial terrace formation and climate change also can be demonstrated in some parts of the Huang Shui and Huang He catchments (Li et al., 1996; Lu et al., 2004a; Pan et al., 2009; Vandenberghe et al., 2011). However, in our study area, six terraces (T1ML to T6ML) have been formed since the last interglacial. This number is higher than expected, based on the models that attribute terrace formation to cold-climatic conditions. This may indicate that terrace formation in the study region is additionally controlled by smaller climatic fluctuations (103-104 year timescale) or relatively small tectonic movements of limited temporal and spatial significance in a setting of more general and well- expressed uplift. The sedimentary sequences are dominated by varying cyclic assemblages of gravel traction-current deposits and sand lenses with rough fining-upward

140 grain size, showing a multi-storage of active channels separated by sand bars. This points to braided river systems, which contrast with the present-day wandering or meandering pattern of the Huang Shui as discussed previously (Vandenberghe et al., 2011). This does not of course exclude a tectonic origin for the initiation of the observed terraces. A cessation or retardation of the general tectonic uplift might invoke lateral migration instead of incision and result in (strath) terrace formation. Thick gravel deposits, as in terraces T5ML and T6ML, would be more difficult to explain in that way. Taking into account the OSL dates (and their error bars) that are available at present (see above; Table 5.1), we may provisionally attribute T6ML (accepting the date of the coarse-grained deposits at (c. 81±6.2 ka) to a cold phase within MIS 5 (for instance stage 5b, when a strong Asian winter monsoon supplied coarse aeolian grains into the loess sequence in the northeastern Tibetan Plateau (Lu et al., 2004b; Vriend et al., 2011), possibly corresponding to Greenland stage 22). The terrace was probably abandoned while occasionally locally flooded during the next warm phase (floodplain deposits dated 74-70 ka: stage 5a, when a relatively stronger Asian summer monsoon decreased the aeolian influx, and soil formation was stimulated in the northeastern Tibetan Plateau (Lu et al., 2004b; Vriend et al., 2011)). The anomalous thickness of T6ML is however attributed to a temporary change in the tectonic movements, see above, which indicates an important tectonic control on terrace aggradation and abandonment. According to the OSL ages (24.42±2.0 ka), terrace T3ML was formed during the cold MIS 2 when the temperature was at least 7 °C lower than at present in the northeastern Tibetan Plateau (Wang et al., 2013). Unfortunately, until now no OSL dating was available for T4ML and T5ML. Interpolating between T3ML and T6ML, we assume provisionally that T5ML is of MIS 4 age and T4ML of MIS 3 age. If a tectonic origin (for which presence there is thus far no evidence) for the formation of T4ML may be excluded, climatic forcing might look strange within the relatively warmer MIS 3. But, we have to take into account that several very cold phases occurred during MIS 3, for instance the Hasselo stadial (van Huissteden, 1990; Kasse et al., 2003) around 42 ka in western Europe and an increase of aeolian influx caused by a stronger Asian winter monsoon during this period (Vriend et al., 2011). In many regions (such as Europe), these cold intervals were of too low magnitude or too short duration to induce sufficiently impoverished vegetation to initiate general changes in the fluvial morphology (Van Huissteden et al., 2001; Van Huissteden and Kasse, 2001; Vandenberghe and Woo, 2002). However, the threshold to increase upstream supply of sediment to the river and subsequent deposition in a braided

141 river in the confluence region of Huang Shui and Huang He, thus creating a terrace, could have been crossed in a setting of continuously strong tectonic uplift. An alternative interpretation would be to explain T4ML as a result of internal evolution (incision) after the formation of a cold-based terrace T5ML in a situation of strong uplift without (or in combination with) climate interference (Bull, 1991). It must be kept in mind, however, that these hypotheses are based on a linear interpolation between T3ML and T6ML. Like T6ML, the excessive thickness of T5ML is explained by changes in the rates and direction of tectonic motions (see above). Therefore, like T6ML, the aggradation and formation of T5ML might be caused by tectonic effects. In that case T4ML might be of MIS 4 age, and the climatic impact on the terrace formation would be more in line with glacial-interglacial climatic cycles (Vandenberghe 1995). It is clear that for further progress the chronology of the terraces needs to be confirmed and specified by additional OSL dating. The formation of T2ML and T1ML may result from intrinsic fluvial evolution within MIS2 and MIS1, respectively. Alternatively, they might be related to the effect of renewed climatic cooling during the Younger Dryas. Here again, more age constraints are required.

5.6. Conclusions In the confluence region of the Huang Shui and Huang He rivers, from the Minhe depression to the West Lanzhou depression, eight fluvial terraces are present, of which the sixth and the third terrace (T6ML, T3ML) are dated at 81- 73 ka and 22 ka, respectively, using SAR-OSL analysis. Longitudinal terrace profiles show that no different tectonic movements have taken place within this area since the formation of the seventh terrace. Comparison to areas immediately upstream and downstream of confluence area indicates differential tectonic movements, implying these three areas are situated on different tectonic blocks. The evidence is based on different terrace staircases, different ages of the terraces, and different long-term incision rates. The last shows that our study area experienced the highest uplift rate. As a result, this relatively strong uplift generally overprints the climatic impact on terrace formation. However, two terraces with thick fluvial deposits (>18 m) (T6ML and T5ML), indicate relative subsidence (or relative retardation in the general uplift) at those times. In all, our results show variable tectonic motions in space and time in the upper reaches of the Huang He and the Huang Shui rivers. Terrace formation in the study region appears to be controlled by small climatic fluctuations (103-104 year timescale) against a background of strong uplift.

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Acknowledgements This research was supported by the National Natural Science Foundation of China (Grant no. 40901002, 41271001 and 41021002), and Dutch–Chinese CAS-KNAW PhD project (05-PhD-10) and Exchange Programs (08CDP028, 09CDP033 and 10CDP036). The comments by David Bridgland, Darrel Maddy, Manfred Frechen and Jan A. Piotrowski are highly appreciated. Thanks are also extended to Baotian Pan for helpful discussions. We also appreciate Dr. Dimitri Vandenberghe for constructive suggestions on OSL dating.

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-Chapter 6- Synthesis

6.1. Morpho-tectonic and geomorphological evolution of the Northeastern Tibetan Plateau (NETP) Several stages can be discriminated in the geomorphological evolution of the Huang Shui catchment. During the first stage, the landscape consisted of mountains and filled-up basins that were leveled synchronously, ultimately resulting in the formation of a peneplain. This peneplain thus consists partly of a bedrock surface and partly of the eroded top of the basin fills. During the second stage, the Huang Shui incised into the peneplain surface, producing a terrace sequence. The transition from peneplain formation to incision in the study region was dated as older than 10-6 Ma using the biochronology of micromammalian assemblages from fluvial terraces. In addition, the depositional record of the basin fill indicates that the river incision into the former peneplain should be younger than 17 Ma. The start of the incision may thus be attributed to an important uplift event starting at around 10-17 Ma (chapter 2, 3). Former studies near the Huang Shui catchment, by Pan et al. (2011, 2012) and Craddock et al. (2010), show that rivers started entrenching from a bevelled (erosional) surface at much younger ages, c. 3.7 Ma and 1.8 Ma along the Yellow River. The discrepancy may be explained in several ways (chapter 2). First, the older landscapes at the other sites may not have been preserved; they might have been eroded. Secondly, the relict beveled surfaces may have experienced a complex history with different ages of development in different areas (Clark et al., 2006). The late entrenching of the relict surface reported by the former authors could have been caused by headward erosion of rivers flowing at a lower level and not by a river properly flowing on top of the peneplain, like in the Huang Shui catchment. Thus, a certain lag is possible because of time needed by the headward erosion. In our opinion, the infilling of the large tectonic basins and the peneplain formation should have occurred largely in the same time period on the NETP. But, the effective duration of these different processes, and thus their time of termination, may have been different at different locations.

6.2. Morphology of terrace staircases and its response to tectonic movement and climatic change The general uplift at 10-17 Ma caused local fragmentation of the Huang Shui catchment into blocks of small extent (scale of kilometers or maximally, a

149 few tens of kilometers) (chapter 3). These blocks have undergone relative subsidence and/or uplift. We suggest that the inferred tectonic motions are related to the transpression movements in the NETP as a result of the collision of the Indian and Asian plates. Fluvial deposition (>30 m thick) in the subsided blocks, resulting in accumulation terraces, contrasts with entrenching (resulting in erosional terraces with thin fluvial deposits, i.e. strath terraces) and formation of gorges (where terraces are rare or completely absent) in the uplifted blocks. The different rates of uplift and subsidence in the individual blocks resulted in the simultaneous development of erosion and accumulation terraces in different blocks within the same catchment. In other words, fluvial aggradation may have occurred in specific blocks at the same time rivers incised in other (adjacent) blocks (chapter 5). The entrenchment rate of the Huang Shui was not stable. It accelerated for instance from 2 cm/ka to 19.5 cm/ka at around 2 Ma, indicating a specific phase of accelerated uplift. In contrast with the generally persisting uplift of the NETP, at the scale of individual blocks uplift sometimes alternated with relative subsidence, leading to laterally alternating accumulation and erosion terraces in those blocks (see also 6.4). The sedimentary series of the different terraces are similar (chapter 3). From bottom to top: fluvial gravels of various thickness, interbedded with lenses of sands, silts, and clays, and finally topped by a horizontally laminated silt, occasionally containing gravel strings of limited extent. The lowest part of the gravel deposits consists of massive gravel (Gm) representing channel bedload, grading towards the top into finer-grained, coarsely planar bedding (Gp). The latter deposits frequently show imbrication and small-scaled cross- bedding, which indicate deposition in lateral and longitudinal bars. Small and shallow channels occur with increasing frequency towards the top of the terrace deposits; they are filled with cross-bedded fine gravel or sands (planar to low- angle trough cross-bedding). The upper laminated silts are interpreted as floodloam deposits that complete the fluvial sequence prior to or simultaneously with the abandonment of the floodplain as a result of renewed river incision. These sedimentary characteristics of the fluvial gravel-sand deposits, dominated by varying (often cyclic) assemblages of gravel traction- current deposits, clearly point to the shallow gravel-bed braided river systems that contrast with the present-day wandering or meandering pattern of the Huang Shui (chapter 3, 5). At several locations, soil formation has been recognized at the top of fluvial depositional sequences, although not being well developed. Mostly, a gradual transition from flood loam to loess is present and

150 the flood loams are often red-colored, pointing to upstream removal or local reworking of soil material., Incision into the former peneplain was not continuous but a staircase of terraces (consisting of at maximum 15 to 20 individaul levels in some blocks), developed as a result of climatic influences. Climate proxies are mostly absent within the Huang Shui terrace deposits, which makes it difficult to prove the link between climate and fluvial processes. Nevertheless, in accordance with the reconstructed character of the river morphology as a shallow gravel-bed braided river (Miall, 1996) that contrasts with the Holocene meandering river, the majority of the (coarse grained) fluvial deposits of the terraces seems to date from cold periods. Subsequently, we assume that, in line with the general model of fluvial development derived for temperate and periglacial environments (last update by Vandenberghe, 2008), the fine-grained sandy gravels and sands near to the top of the sediment series were deposited during the waning of the same glacial period as a result of the beginning incision of the river (possibly in combination with reduced precipitation). Ultimately the fluvial sediments were capped by interglacial soil formation. Rivers incised slightly at the transition from the glacial to the interglacial, due to less peaked river discharge and the reduction of sediment supply to the river. As a consequence of that new, lower position in the next interglacial, the previous coarse grained channel deposits were flooded only during peak discharges resulting in episodic, fine-grained floodplain deposition, while most of the time soil formation took place on the floodplain. Analogous to the present-day river, the interglacial river was probably less energetic and dominantly meandering. Renewed incision took place at the next interglacial-glacial transition, when fluvial energy considerably increased and sediment transport was still limited. At that time, loess covered the former floodplain deposits and soil without any further reworking by the river.

6.3. Quartz Optically Stimulated Luminescence (OSL) dating Datings are the prerequisite to unravel the details of the relationship between the fluvial events (as terrace formation, processes of erosion and deposition) and climate changes, and they are also significant to examine the neotectonic process in the Northeastern Tibetan Plateau. It is difficult to establish a reliable chronological control by 14C dating for the last-glacial events in the study region because of the dating limit (<45 ka) and the lack of suitable dating materials in this arid and cold area. In this study, the typical fluvial channel sands, floodloam sediments, outwash fans and glacial- and fluvio-glacial sediments were dated using quartz optically stimulated

151 luminescence (OSL) with the single aliquot regenerative-dose (SAR) protocol. From all samples, a small but sufficient amount of fine-sand quartz could be extracted to permit standard SAR-OSL analysis. The measurements using small aliquots (or single grains) might be the best to assess the degree of partial bleaching and date glacial and glacio-fluvial sediments, as they are likely to suffer from poor or inhomogeneous bleaching. However, our measurements show that OSL signals of the samples in the Tibetan Plateau were not bright and that their luminescence sensitivity was low. Thus, in most cases, we can only use large aliquots to date our samples, which are assumed to possibly mask any inhomogeneity in the sample and provide overestimated ages. The results of the measurements using small aliquots for two selected relatively bright samples show the distributions of the equivalent doses are broad, with relative standard deviations in the range of ~16 to 26%, and display little or no asymmetry. In addition, the un-weighted average De’s using small aliquots are not different from the values obtained using large aliquots. These results indicate a well-bleached nature of glacial- and fluvial sediments in the NETP, and do not hint at incomplete resetting as a significant source of error. Also, it has been argued (see e.g., Murray and Olley, 2002; Jain et al., 2003) that incomplete resetting is unlikely to give rise to significant age overestimations for samples older than a few ka, even in glacio-fluvial environments. Thus the Quartz-based ASR-OSL analysis could provide robust age control for glacial and fluvial terrace deposits in the NETP, although they should, at least in principle, be considered as maximum ages.

6.4. Fluvial geomorphological processes and terrace formation in relation to a nearby glacial environment (chapter 4) The Menyuan Basin, located in the north of the Huang Shui catchment, is an area of tectonic subsidence, bounded by the Qilian mountains with modern glaciers covering high areas. The Datong river, an important tributary of the Huang Shui, drains this basin occurring at relatively high altitude. The morphology of the southern border zone of the Qilian mountains is characterized by glacial valleys, while moraines and outwash fans occur in the adjacent border zone of the Datong valley. The Qilian mountains have experienced intensive glaciation during the last glacial cycle, with glacial melt-water systems draining in to the Datong river. The areas occupied by glacio-fluvial deposition experienced continuous permafrost in the Menyuan basin during LGM as indicated by mega- cryoturbations. These large cryoturbations show that the mean annual temperature was at least 7 °C lower than at present. The presence of

152 cryoturbations and absence of sand wedges in the Menyuan basin contrast with the general presence of sand-wedges and absence of cryoturbations during the LGM in adjacent regions like the Mu Us desert and Hexi corridor, indicating that it was more humid in the Menyuan basin than in those adjacent regions. During glacial periods, such as the LGM, glaciers advanced in the Qilian mountains, probably as a response to temperature decline rather than monsoon- related precipitation increase. In accordance with the morphological evolution in glacial environments, fluvial systems in the Menyuan basin aggraded and terraces formed during cold periods (penultimate glacial, LGM, and possibly the Younger Dryas), likely as a response to high glacial sediment production in combination with high peak discharges. Down cutting after deglaciation is probably due to relatively decreased sedimentation supply (e.g., Marren and Toomath, 2013). In total, six fluvial terraces have been identified in the Menyuan basin since the penultimate glacial.

6.5. Local tectonic movements and the coupling impacts of tectonic and climate changes on fluvial evolution since the last interglacial (chapter 5) In the confluence region of the Huang Shui and Huang He rivers, from the Minhe depression to the West Lanzhou depression, eight fluvial terraces are present, of which the sixth and the third terrace (T6ML, T3ML) are dated at - 70-81 ka and -24 ka, respectively, using SAR-OSL analysis. Longitudinal terrace profiles show that no distinct different tectonic movements have taken place within this area since the formation of the seventh terrace (older than - 70ka). The average river incision rate since -70 ka was much higher in the confluence region block than in the up- and downstream blocks, indicating relative uplift of the confluence region. This relatively strong uplift gives more space for differentiation within the terrace staircase as a result of climatic changes, leading to 6 terraces formed as a response to small climatic fluctuations (103-104 year timescale). Assuming that the OSL-age of terrace T6ML is reliable, it leads to the striking conclusion that the stronger the tectonic movement the better the climatic imprint is expressed in the terrace development. At a smaller time scale, somewhere between 20 and 70 ka, two accumulation terraces with thick stacked fluvial deposits (>18 m) indicate a phase of subsidence in the confluence region relative to the up- and downstream blocks. This points to relatively short phases of tectonic stability or even subsidence during a period of general tectonic uplift.

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6.6. Conclusions and future outlook Quartz-based SAR-OSL dating yielded ages that were generally consistent with the stratigraphic and geomorphological position of the samples, showing that OSL dating is well suitable for dating not only in aeolian and fluvial, but also in glacial (morainic) and glacio-fluvial deposits in the NE Tibetan Plateau. In the NE Tibetan Plateau, fluvial aggradation occurred during cold periods in general, resulting from the increase of sediment supply both in periglacial and glacial conditions. In this monsoonal climate region, the resuming vegetation cover reduced sediment supply to the rivers so that rivers progressively incised, reaching only sporadically the previous floodplain at the transition to inter-glacials. This change in fluvial activity as a response to climatic impact is reflected in the general sedimentary sequence on the terraces from high-energy (braided) channel deposits (at full glacial) to lower-energy deposits of small channels (towards the end of the glacial), mostly separated by a rather sharp boundary from overlying flood-loams (at the glacial-interglacial transition) and overall soil formation (interglacial). Probably, renewed erosion at the next interglacial-glacial transition was responsible for the absence of interglacial channel deposits. During the next glacial period this fluvial sedimentary series is covered by purely aeolian loess. Transpression movements, related to the collision of the Indian and Asian plates, lead to the fragmentation of the NETP into local blocks of small extent during middle and late Miocene. Since that time, rivers started to incise into the peneplain and formed a staircase of terraces. Since late Miocene, in spite of generally persisting uplift of the whole region, the neighbouring tectonic blocks had different uplift rates, and the uplift rates could change on a time scale of 10 Ka. This tectonic differentiation causes a complicated fluvial response with accumulation terraces alternating with erosion terraces at a small spatial and temporal scale. In addition, it is hypothesized that in some strongly uplifted blocks energy thresholds could be crossed to allow terrace formation as a response to small climatic fluctuations (103-104 year timescale). These results show that commonly applied assumptions in fluvial terrace studies concerning tectonic motions are too simplistic. For example, it is in general commonly assumed that tectonic uplift rates are similar over relatively large areas, and that tectonic motions are more or less constant with time. Our results show spatial and temporal dynamics of tectonics occurring in an active mountain belt. Our results also imply that in this kind of setting terraces cannot be correlated based on altitude and/or their aggradational- or strath characteristics alone. Detailed mapping and dating are the most useful ways to solve this problem.

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Although the thesis contributes to a better understanding of the morphotectonic and geomorphological evolution of the NETP, and to improved insights on the impacts of tectonic motions and monsoonal climate changes on fluvial processes, a number of questions and tasks remain to be answered and carried out in the future: 1) How far are the peneplain and the related morphological features, such as oldest terraces and wide and shallow valleys, preserved, and can they be correlated over large distances? 2) Morphological features as terraces and peneplains at different blocks have to be dated more precisely to make better correlations and, therefore, firmer hypotheses on morphological evolution and the link in different blocks and/or in different tectonic settings. As OSL-analysis has been proven to provide reliable ages, an extensive programme of OSL-dating of fluvial and fluvio-glacial deposits should be established. 3) Are there different geomorphological processes and evolutions in the different parts of the NETP because of local geology and tectonic setting? 4) Are the response styles of fluvial systems to small climatic fluctuations (103-104 year timescale) similar in different tectonic blocks with different uplift rates? 5) Can we specify in more detail to what extent and intensity tectonic movements may influence the crossing of climatic thresholds, leading to terrace development?

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Murray, A.S., Olley, J.M., 2002. Precision and accuracy in the optically stimulated luminescence dating of sedimentary quartz: a status review. Geochronometria 21, 1-16. Pan, B., Hu, Z., Wang, J., Vandenberghe, J., Hu, X., 2011. A magnetostratigraphic record of landscape development in the eastern Ordos Plateau, China: Transition from Late Miocene and Early Pliocene stacked sedimentation to Late Pliocene and Quaternary uplift and incision by the Yellow River. Geomorphology 125, 225-238. Pan, B., Hu, Z., Wen, Y., Vandenberghe, J., Wang, J., Hu, X., Gao, H., Guan, Q., Li, Q., 2012. The approximate age of the planation surface and the incision of the Yellow River. Palaeogeography, Palaeoclimatology, Palaeoecology 356-357, 54-61. Vandenberghe, J., 2008. The fluvial cycle at cold-warm-cold transitions in lowland regions: a refinement of theory. Geomorphology 98, 275-284.

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