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Late Quaternary landscape evolution in a small catchment on the Chinese Loess Plateau
Quaternary International 234 (2011) 159–166
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Late Quaternary landscape evolution in a small catchment on the Chinese Loess Plateau Brigitta Schu¨tt a, *, Manfred Frechen b, Philipp Hoelzmann a, Georg Fritzenwenger a a b
¨t Berlin, Department of Earth Sciences, Physical Geography, Malteserstraße 74-100, Haus H, 12249 Berlin, Germany Freie Universita Leibniz Institute for Applied Geophysics (LIAG), Section S3: Geochronology and Isotope Hydrology, Stilleweg 2, 30655 Hannover, Germany
a r t i c l e i n f o
a b s t r a c t
Article history: Available online 6 January 2010
Late Quaternary landscape evolution was analysed with reference to a small (0.08 km2) drainage basin system in southeastern Gansu Province, southwestern Chinese Loess Plateau. In this region, V-shaped valleys with broad valley floors are prominent landscape features whose basic shape already existed prior to MIS 2. During MIS 2 the valley bottoms were infilled with fluvially reworked loess, today forming accumulation terraces. Infrared optically stimulated luminescence (IRSL) dating of these loess-like deposits assigned them to the time span ranging from 22.5 to 17.7 ka BP. During the late Holocene, valley fills were partly eroded, and the receiving stream channel was deepened by at least 25 m close to the confluence of the study site’s channels. These alternating erosion and accumulation phases were triggered by climatic factors or human impact. In the most recent past, the high landscape sensitivity of the Chinese Loess Plateau has led to increasing soil conservation measures that have effectively confined surface erosion. However, the lack of drainage on the artificial terraces induces piping processes. Ó 2010 Elsevier Ltd and INQUA. All rights reserved.
1. Introduction In China, loess deposits, reaching several hundred metres in thickness, are widespread over an area of approximately 440,000 km2 between latitudes 33 N and 47 N and longitudes 75 E and 125 E (Liu et al., 1985). North of Lanzhou, in the western part of the Loess Plateau, such deposits attain a maximum thickness of 505 m (Derbyshire and Meng, 2000), decreasing to the south and east (Sun, 2002). Liu et al. (1985) describe the stratification of a 140 m thick loess profile at Luochuan in central Shaanxi province, identifying four major units of loessial deposits with interlayered palaeosols (Ding et al., 1993). The lowest unit is the Wucheng Loess, which was deposited in the early Pleistocene (2.4–1.15 Ma). Up to 100 m thick, it is composed of a reddish to light-brown, wellcemented and fine-grained silt. The overlying Lishi Loess (1.15– 0.1 Ma) ranges between 40 and 70 m in thickness (rarely 200 m); it is poorly cemented and of a greyish brown colour. The late Pleistocene Malan Loess (100–2 ka) is 10–30 m thick. This loess type is greyish yellow, loosely cemented, and coarsely grained. It is overlain by a 2–3 m thick loess layer of Holocene age. This basic stratigraphy has been assigned and adapted to many sites of the
* Corresponding author. Tel.: þ49 30 838 70479. E-mail address: [email protected] (B. Schu¨ tt). 1040-6182/$ – see front matter Ó 2010 Elsevier Ltd and INQUA. All rights reserved. doi:10.1016/j.quaint.2009.12.018
Chinese Loess Plateau by Kukla (1987), Derbyshire et al. (1995) and Sun (2002), among many others. Previous work has shown that since the early Pleistocene there have been numerous alternations between periods of aridisation, accompanied by loess deposition, and more humid environmental conditions with (palaeo-)soil formation (An et al., 1991; Ding et al., 1995; Rost, 2000). In consequence, loess stratigraphy has been widely used to reconstruct Cenozoic palaeoenvironmental conditions of the Chinese Loess Plateau (Ding et al., 1993). The well-known high sensitivity of loess to erosion, soil erosion processes (Casali et al., 1999; Vanwalleghem et al., 2003; Porter and An, 2005; Ionita, 2006), and erosion by piping processes (Bariss, 1977; Bocco, 1991; Bryan and Jones, 2000; Derbyshire, 2001) has led to many studies on present-day morphodynamics (among others: Wu and Cheng, 2005; Hessel and van Asch, 2003; Cheng et al., 2007). However, only few authors (e.g., Bork et al., 2001) have investigated fluvial processes and soil erosion as landform-shaping modes and their reflection in a morphogenetic context. The present study investigates the relief evolution of a small catchment in the southwestern Chinese Loess Plateau. The aim is to enhance understanding of late Quaternary relief shaping by studying the chronology of fluvial erosion and deposition phases in an area of high landscape sensitivity due to the high erodibility of primary loess and its fluvially redeposited sediments under the influence of climate change and tectonic activity. Luminescence
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dating of the sediment allows the age determination of the corresponding erosion and fluvial deposition phases. 2. Regional setting The study site is a small catchment area (0.08 km2), located in the southwestern Chinese Loess Plateau, about 60 km southeast of Lanzhou, the capital city of Gansu Province. This region is known as the Longxi Plateau and is situated at the northeastern margin of the Tibetan Plateau, where crustal deformations resulted in uplift and major faulting during the Quaternary (Dijkstra, 2001). In the south and southwest, the study area is bounded by the active WNW–ESE Qinling fault zone and in the east by the active NNW–SSE striking faults of the Liupan Mountains (Derbyshire et al., 1995). In the area of the Longxi Plateau, Quaternary loess was deposited on a slightly hilly relief. The underlying bedrock shows evidence of major tectonic disturbance and exhibits a variety of folds and faults, creating an extremely unstable subsurface for the overlying loess formation. The resulting late Cenozoic uplift of the area caused increasing relief and significant erosion by mass movements and fluvial processes (Derbyshire and Meng, 2000). The present climate of the southwestern Chinese Loess Plateau is highly seasonal, with cold, dry winters and warm, humid summers. During winter, cold and dry air masses move between eastern Sibiria and northern China, causing temperature and precipitation to drop (Porter, 2001), while the Siberian cold continental high-pressure zone builds up and expands. Air pressure at ground level on the adjoining Tibetan Plateau is lower owing to its latitude and stronger altitude-induced insolation. The resulting northerly winds are cold and dry (Shahgedanova, 2003). By contrast, differential summer heating between the Tibetan Plateau and the south- and eastwardly adjoining Indian Ocean and the East Chinese Sea results in the flow of moist, southeasterly air to the hot continental low-pressure system of the Tibetan Plateau (Rokosh et al., 2002). These monsoonal currents also affect the fringe of the Tibetan Plateau and adjacent lowlands, locally overlapping with thunder cells due to heating where wetlands occur (Lehmkuhl and Haselein, 2000). At present, the progressive expansion of the summer monsoon, with its predominantly southern to southeastern winds, reaches the study area in late July each year and gradually declines from late August onwards. Accordingly, the rainy season generally lasts 2–3 months. The expansion and the strength of the summer monsoon depend on the position and development of the Northern Pacific subtropical high-pressure zone (Domroes, 2001). As the study area is located at the northern margin of the monsoonal realm, precipitation is reliable during the summer rainy season. However, total annual precipitation is highly variable. The mean annual precipitation measured at the neighbouring climate stations totals 328 mm a1 at Lanzhou and 327 mm a1 at Yu Zhong; more than 70% of this rainfall occurs between July and September. More than 60% of the precipitation falls during rainstorm events, whose intensity often exceeds soil infiltration capacity and in consequence causes surface runoff (Feng et al., 2007). 3. Methods 3.1. Chronology The samples were dated using the infrared optically stimulated luminescence dating technique (IRSL). Five samples of fluvially reworked loess were extracted at different river terrace levels (accumulation terraces) along the study site’s channel and its receiving stream (Fig. 2, Table 1).
Table 1 Palaeodose values and loss deposition ages using infrared optically stimulated luminescence. Sample field No.
Gamma spectrometry K (%) Th (ppm)
U (ppm)
OSL0807-1 OSL0807-2 OSL0807-3 OSL0807-4 OSL0807-7
1.66 1.70 1.72 1.82 1.72
3.4 0.03 3.3 0.03 3.4 0.02 3.0 0.02 3.1 0.02
13.2 0.07 12.2 0.06 12.8 0.04 10.8 0.06 11.6 0.06
Cosmic dose Palaeodose Age (ka) rate (mGy/yr) (Gy) 219.0 11.0 107.6 2.5 22.5 1.5 113.0 5.7 93.8 2.9 20.7 1.4 182.0 9.1 84.7 2.5 17.8 1.2 171.0 8.6 78.1 2.0 17.7 0.8 160 8.0 85.2 3.8 19.1 1.5
Luminescence dating of aeolian deposits has proved to be successful where radiocarbon and other dating methods are not applicable (Frechen, 1999; Roberts, 2008). The basic principle of luminescence dating is solid state dosimetry of ionising radiation (Aitken, 1998; Lian and Roberts, 2006). Luminescence is the light emitted from crystals such as quartz, feldspar or zircon when they are stimulated with heat or light after receiving a natural or artificial radiation dose. The equivalent dose (De) is a measure of the past radiation energy absorbed in natural dosimeters like quartz and feldspar minerals and, in combination with the dose rate (the rate of radiation absorbed per unit time), yields the time elapsed since the last exposure to sunlight. Infrared optically stimulated luminescence (IRSL) dating was carried out to determine the deposition age (the time elapsed since the last exposure to daylight or sunlight) for the reworked loess sediments from the sections under study. Experimental details were applied following the IRSL dating approach of Frechen et al. (2009) and so are only summarised here. IRSL measurements were carried out on five samples. Polymineral fine-grained material (4–11 mm) was prepared to determine the equivalent dose, as described by Frechen et al. (1996). The samples were irradiated by a 90Sr/90Y beta source (0.172 Gy s1) in at least seven dose steps with five discs each and a maximum radiation dose of 930 Gray (Gy). All discs were stored at room temperature for at least four weeks after irradiation. The irradiated samples were preheated for 1 min at 230 C before infrared stimulation. A Schott BG39/Corning 7-59 filter combination was placed between photomultiplier and aliquots for IRSL measurements. Each aliquot was kept at a temperature of 50 C during 10 s of IR decay. The equivalent dose was obtained by integrating the 1–5 s region of the IRSL decay curves. Alpha efficiency was estimated to a mean value of 0.08 0.02 for all samples. Dose rates for all samples were calculated from potassium, uranium and thorium contents, as measured by gamma spectrometry (N-type high purity Germanium (HPGe) detector with 25% relative efficiency) in the laboratory, assuming radioactive equilibrium for the decay chains. Cosmic dose rate was corrected for the altitude and sediment thickness, as described by Aitken (1998) and Prescott and Hutton (1994). The natural moisture content of the sediment was estimated at 10 3%. In fluvially reworked sediments, partial bleaching prior to deposition is one of the major causes of significant age overestimation (Frechen et al., 2008). The existence of scatter in De value determination is an indication for incomplete zeroing of the OSL signal but was not observed for the samples under study. As polymineral fine-grain material was used for dating, statistical methods could not be applied to study partial bleaching. Although there is no evidence for partial bleaching, the results should be considered as maximum deposition ages. Wintle (1973) describes the problem of anomalous fading, which is a loss of charge with time results in age underestimation. A delay of at least four weeks between artificial irradiation and measurement
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was applied to minimise the age underestimation caused by anomalous fading. 3.2. Geomorphological mapping Landform unit identification and mapping were based on the LANDSAT-7 ETMþ satellite image (p130/r035, 18.07.1999). The digital elevation model (DEM) was generated from the data provided by the ASTER satellite images (Advanced Spaceborne Thermal Emission and Reflection Radiometer, images from 14.07.2001) with a spatial resolution of 30 m 30 m. A geomorphological field survey was carried out in August 2007 following the mapping guidelines of Leser and Sta¨blein (1975). High resolution topography was measured along the sampling sites with a measuring tape and a clinometer. 4. Results 4.1. Relief The study site is a small drainage basin with an area of 0.08 km2 and a vertical extent of 150 m; the highest point of the divide is at 1710 m above sea level (asl). The total drainage basin length is 220 m; the main channel is 95 m long and drains from east to west. The study site’s streams are ephemeral, bearing runoff only after rainfall events. The receiving stream is also ephemeral and flows into a tributary of the Huang River (Huang He) (Fig. 1). The receiving stream’s catchment, which contains the study site, covers 2.32 km2 (Fig. 1). Loess is the parent material in the whole catchment. The summit altitudes of the divides reach 1800–2000 m asl. Total drainage basin length is 600 m, with a stream length of 520 m and a most recent channel incision of about 120 m. Correspondingly, slope gradients widely exceed 25 and the relative relief (ratio of maximum difference in altitude and drainage basin length) ranges between 0.15 and 0.35. Depending on slope gradient and slope aspect, the vegetation cover density varies between 0 and 50%, with an average of 28.5%; the lowest vegetation cover is on the south-facing slopes. Along the lower course of the receiving stream, fluvial accumulation terraces are visible at three different levels along the main drainage way (Fig. 1). At the confluence of the study site’s channel and the receiving stream, the latter’s channel is boxshaped, and its floor is located approximately 25.5 m below the valley bottom of the tributary channel draining the study site. The lowest accumulation terrace is located 15 m above the valley floor of the receiving stream, followed by a second terrace 8 m upslope (23 m above the receiving stream’s valley floor), and a third accumulation terrace after another 2.5 m (25.5 m above the receiving stream’s valley floor). The study site’s drainage pattern is dendritic, with three secondorder channels (after Horton, 1945) draining the headwater area and feeding the main channel. In the upper part of the study site, channels indicate the temporary occurrence of concentrated runoff, and the high frequency of riffle-pool sequences indicates turbulent runoff (Clifford, 1993) as an expression of high flow velocity (Leopold et al., 1992); both factors are also reflected in the occurrence of V-shaped valleys (Fig. 2). The lower part of the study site is characterised by a gully-like, box-shaped valley. This valley section has a total length of 95 m; the valley bottom is infilled, and the channel is up to 65 m wide (Figs. 2 and 3). The longitudinal profile of this valley section is step-like, and its floor is adjusted to the uppermost terrace level of the receiving stream: in four steps the channel of the box-shaped valley overcomes 30 m difference in altitude. Sediments exposed along the steps are layered and composed of silt. These sediments are termed ‘loess-like’ as the origin of these
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valley infills is loess that has been removed by hill wash and fluvial erosion (Wenske et al., 2011). Valley cross-profiles (Fig. 2) are measured for the headwater area (A), the confluence of the three major receiving gullies (B) and the lower course before it drains into the main valley (C). Crossprofile A shows two V-shaped valleys with slopes up to 45 . Typical steppe vegetation such as Artemisia and Stipa is present all along the cross-profiles. Canopy cover varies from 5% on steep slopes to more than 50% on artificial terraces, where slope amounts to less than 10 . The north-facing slope is shaped by up to 4 m wide terraces, created for soil conservation. However, at present the site is not used for agriculture. These terraced slopes lack linear erosion features. Instead, several piping forms indicate subsurface runoff; they occur predominantly at the terrace edges. Correspondingly, piping occurs throughout the drainage basin of the receiving stream, wherever terraces have been built as soil conservation measures. A pipe inventory shows that 30% of the pipes recorded appear in artificial terraces, which cover about 15% of the total catchment area. Cross-profile B is located immediately downstream of the confluence of the three channels draining the headwater area. The maximum slope is 45 in this transition zone between the upper headwater area and its lower part. The north-facing slopes have been shaped by artificial terraces built for soil conservation purposes. In the downslope parts, terrace width increases up to 8–9 m, corresponding to the slope angle of 25 . The flat topography of the channel indicates an infilled valley. The vegetation cover resembles that of cross-profile A: Artemisia and Stipa dominate, with the densest canopy cover (more than 50%) along the flat valley floor. The flat topography in the central part of cross-profile C again marks the valley infill. Maximum channel width is 30 m; adjacent slopes incline by as much as 70 . Small mass movements occur all along the steep slopes. Valley infills are riddled by piping.
4.2. Chronology In the downstream section of the study site the main drainage way (Figs. 3 and 4) is divided into four levels (L1–L4) by a step-pool sequence with three steps. All levels are almost planar, owing to the fluvially infilled valley floor and separated from each other by steep, cascade-like steps. The valley fills are exposed at the steps and sediments show a clear layering with occasional cross beds, confirming deposition by fluvial processes. To reconstruct valley evolution, four samples were extracted from the valley fills exposed in the steps (Figs. 3 and 4). OSL0807-1 was sampled from the valley fill of section L2 at 1.1 m below the present-day valley bottom; IRSL dating yielded an age of 22.5 1.5 ka BP. OSL0807-2 and OSL0807-3 were extracted from section L3, with OSL0807-2 sampled from the valley fill at 5.4 m, and OSL0807-3 at 2.0 m below the present-day valley floor. IRSL dating yielded an age of 20.7 1.4 ka BP for OSL0807-2 and 17.8 1.2 ka BP for OSL0807-3. OSL0807-4 was sampled from the valley fills of section L4, extracted 2.0 m below the valley floor; IRSL sediment ages are 17.7 0.8 ka BP (Table 1). In addition, sediments from the accumulation terraces of the study site’s receiving stream were sampled for dating immediately downstream of the confluence (Fig. 2). Four samples were taken from the three recorded accumulation terraces of the valley fill. OSL0807-7 was taken from the terrace level L2, 7.0 m below surface and about 16 m above the present-day valley floor. The sediment structure indicates that accumulation processes were primarily fluvial; the silty texture underlines that the deposits derive from fluvially transported loess in the headwater area. IRSL dating of OSL0807-7 yielded an age of 19.1 1.5 ka BP (Table 1).
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Fig. 1. Location of the study site. The overview map (left) shows the location in the southwestern Chinese Loess Plateau (a) with regard to the Yellow River as the major regional drainage system (b; map source: Fu et al., 2006). The detailed map (c) gives an overview of the geomorphology of the receiving stream’s drainage basin; the location of the study site at the right lower course is marked by the black frame (elevation data are based on Aster DEM, 2001).
5. Discussion 5.1. Present-day processes According to Derbyshire (2001) the study site is located in the highest collapsibility zone of the Loess Plateau. It displays severe erosional damage such as small mass movements, gully erosion and piping (Figs. 2 and 4). This is mainly due to the texture of loess, which causes high soil erodibility (Planchon et al., 1987). Throughout the headwater area, erosion features dominate along the slopes and control present-day morphodynamics, whereas accumulation features are absent. In their studies of the Chinese Loess Plateau, Valentin et al. (2005) and Guanglu et al. (2004) observe that slope steepness and absence of vegetation cover are the major factors contributing to rill and gully erosion. These results compare to some degree with the study site: rill erosion is initiated at all slopes with poor vegetation cover and a slope gradient of more than 5 , progressing to gully erosion with increasing slope gradient. The construction of artificial terraces increases slope stability at the headwater area of the study site and explains the absence of linear erosional forms. Most artificial terraces are bounded by low
earth benches to retain runoff. Consequently, the water infiltrates in small desiccation cracks at the edge of the terraces, where the hydraulic gradient is highest, and enters the subsurface drainage system. Piping develops especially in these positions. Thus, piping forms at the study site are frequently caused by the lack of drainage at terraces created as soil conservation measures (Hudson, 1995). Accordingly, the slope surface is not exposed to erosion processes, whereas resulting subsurface runoff causes severe erosion damage, superficially visible where pipe systems collapse and at worst build gully heads (Derbyshire, 2001). With regard to gully formation, Bocco (1991) observes that 60% of the gullies in semi-arid areas were formed as a result of piping; because the subterranean runoff systems lead to instability and subsidence of the soil, the resulting small depressions are starting points for further gully formation. 5.2. Late Quaternary landscape history The lower part of the study site is dominated by a gully-like valley. The sediment character of the outcrops along the cascadelike steps clearly indicates that present-day valley shaping results from the very recent dissection of the fluvially deposited infill of an
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Fig. 2. Geomorphological map of the study site. Black squares show sampling locations for luminescence dating with samples 1–4 extracted from the valley infills of the study site and samples 5–7 extracted from accumulation terraces of the receiving stream. Diagrams on the right show valley cross-profiles A–C.
older, more strongly incised valley. The valley fill consists of layered silts, originating from erosion of the loess of the tributary slopes and headwater divides. The dated samples scatter between 23 and 17 ka and are in stratigraphic order with the oldest sample in the deposits exposed downstream and the youngest in the deposits exposed upstream. Ages of the valley fill document that the infill process occurred during MIS 2, around the Last Glacial Maximum. Investigation of palaeoenvironmental proxies (by, among others, Frenzel, 1994; Schlu¨tz, 1999; Yan et al., 1999; Zhou et al., 2001; Hong et al., 2003) indicate alternating, but overall cold and semi-
arid conditions changing to dry-subhumid conditions between 18.5 and 17.0 ka BP. Feng et al. (2007) verified this statement for the investigated area. For the Chinese Loess Plateau, Chen et al. (1997) found that the post-glacial climate not far from the study site was slightly drier between 17.2 and 15.4 ka BP. The shift to the modern atmospheric circulation pattern with an intensified Asian monsoon circulation with increased summer rainfall set in around the Bølling-Allerød interstadial (e.g., Wang et al., 2001). It is assumed that the trigger for this change in monsoon circulation was the warming of the North Atlantic Ocean (Overpeck et al., 1996).
Fig. 3. Longitudinal profile of the main thalweg. Luminescence samples derive from different levels. Further details in Table 1.
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Fig. 4. Sample location. The photo shows the tributary catchment with 4 sample sites. (a) marks sample OSL0807-1, (b) OSL0807-2, (c) OSL0807-3 and (d) OSL0807-4.
Given these overall climatic conditions, dominant fluvial processes during MIS 2 corresponded to the processes known today from drylands: periodic to erratic rainstorm events generate concentrated surface runoff, whereas events with low rainfall intensity are not runoff-effective because of the high infiltration capacity of loess (Le Bissonais, 1990). The ubiquitous fines of the parent material in the Chinese Loess Plateau are easily eroded and cause high suspended load in the concentrated runoff. Downstream infiltration loss, which is typical for dryland rivers, causes a sediment overload and the deposition of valley fills (Schwanghart and Schu¨tt, 2008). This process is facilitated by the downstream decreasing slope of the channel. The same process occurres at different drainage basin scales. A similar formation process is likely for the accumulation terrace systems of the receiving stream. In view of the high erosional sensitivity of loess sediments in this area, accumulation rates of up to 6 m are consistent with short runoff periods. The time frame of the valley fills and terrace accumulations preclude a derivation from soil erosion processes resulting from human activity. As the Asian monsoon circulation intensified during the Holocene, humidity increased and the vegetation cover became denser (Zhou et al., 2001). Thus, it may be concluded that intensive rainfall events characterised the rainfall-runoff regime both during MIS 2 and during the subsequent Holocene (MIS 1). Such storm events were erratic during MIS 2 but became more periodic with the change to MIS 1. During the Holocene, increased annual precipitation and more reliable rainfall occurrence facilitated vegetation growth, which in turn confined erosion processes (Hudson, 1995). These relationships are controlled by the decreasing ratio of annual mean precipitation to effective annual mean precipitation (runoffeffective precipitation) with increasing vegetation cover (Langbein and Schumm, 1958; Douglas, 1967). Hence, the area’s relief was stable as long as a natural vegetation cover was present. Landscape destabilisation and soil erosion started after clearing and cultivation. With reference to the piedmont of the Qinling Mountains,
Huang et al. (2006) showed that significant and sustained soil erosion processes – as documented in the deposition of colluvial deposits – started approximately during the Bronze Age (Xia and Shang Dynasties). By contrast, early Neolithic cultures destabilised the landscape system only temporarily (Huang et al., 2006). Accordingly, it is assumed that the LGM valley fills were dissected during the early Holocene. However, correlative sediments are lacking. The data set obtained for the fluvial terrace (OSL0807-7) is not sufficient for further chronological interpretation. With no independent age control, IRSL age estimates from single samples must be regarded with caution (Lehmkuhl et al., 2007). However, the present IRSL ages are considered to be maximum deposition ages and so provide a rough chronological frame for the sediments under study. Data validation is limited by the lack of any data on soils, palaeosols or organic material. Nevertheless, the dated age of the second fluvial terrace (19.1 1.5 ka BP) is in accordance with results from the tributary catchment. 6. Conclusion This study of a small drainage basin shows that the strong relief, characterised by deeply incised V-shaped valleys, originated during MIS 2 at the latest. During this time the level of the receiving stream channel was approximately 25 m higher in altitude than today (w1700 m asl), documented by alluvial deposits at this elevation. The study site’s drainage was aligned to this level. Prior to MIS 2, the tributary valley was strongly V-shaped. This may have been caused by climate-triggered incision processes as well as by tectonic activity (Dijkstra, 2001); both are high-probability factors likely to have affected the shaping of the valley. During MIS 2, incision processes stopped, and hill wash and fluvial accumulation processes created valley fills and accumulation terraces. These processes are typical for cold, dry climates in areas with erratic rainfall, high sediment supply, and downstream decreasing runoff
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due to infiltration loss. However, since the study site is located in an area of high tectonic activity, this factor cannot be excluded as an additional influence affecting relief decline. During the early Holocene, enhanced humidity and a denser vegetation cover stabilised the landscape, coinciding with limited erosion. It is assumed that dissection of the valley fills is due to human impact and, thus constitutes a soil erosion process. The onset of soil erosion processes coincides with intensified land use during the late Holocene; prior to this, the natural vegetation has been disturbed only occasionally. At the same time, the receiving stream channel was deepened by at least 25 m close to the confluence of the study site’s channels. Here too, the possible erosion-triggering impact of tectonic activity has to be taken into account as a controlling factor, in addition to human-induced changes in runoff behaviour. Today, the land surface is again stable: artificial terraces reduce the local relief and promote infiltration. Bush planting offers additional protection against soil erosion. However, insufficient drainage of the artificial terrace systems entails severe piping processes which locally have even developed into gully erosion. Dense vegetation cover on the slopes mitigates soil erosion, so rill erosion is rare. Acknowledgments The authors would like to thank especially Silke Wilhelm from Freie Universita¨t Berlin for her help during field work and all the Chinese, Mongolian and German students of the Eurasia Network for their assistance during the Summer School on Landscape Sensitivity in 2007. We thank Anne Beck for English-language editing. References Aitken, M.J., 1998. An Introduction to Optical Dating. Oxford University Press, Oxford, pp. 280. An, Z., Kukla, G.J., Porter, S.C., Xiao, J., 1991. Magnetic susceptibility evidence of monsoon variations on the Loess Plateau of central China during the last 130,000 years. Quaternary Research 36, 29–36. Bariss, N., 1977. Gullying in a semi-arid loess region of North America. Great PlainsRocky Mountains Geographical Journal 6 (2), 125–132. Bocco, G., 1991. Gully erosion: processes and models. Progress in Physical Geography 15, 392–406. Bork, H.-R., Li, Y., Zhao, Y., Zhang, J., Shiquan, Y., 2001. Land use changes and gully development in the Upper Yangtze River Basin, SW-China. Journal of Mountain Research 19 (2), 97–103. Science Press, Chengdu. Bryan, R.B., Jones, J.A.A., 2000. The significance of soil piping processes: inventory and prospect. Geomophology 20, 209–218. Casali, J., Lopez, J.J., Giraldez, J.V., 1999. Ephemeral gully erosion in southern Navarra (Spain). Catena 36, 65–84. Chen, F., Bloemendal, J., Wang, J., Oldfield, F., 1997. High-resolution multi-proxy climate records from Chinese loess: evidence for rapid climatic changes over the last 75 kyr. Palaeogeography, Palaeoclimatology, Palaeoecology 130, 323–335. Cheng, H., Zou, X., Wu, Y., Zhang, C., Zheng, Q., Jiang, Z., 2007. Morphology parameters of ephemeral gully in characteristics hillslopes on the Loess Plateau of China. Soil and Tillage Research 94, 4–14. Clifford, N.J., 1993. Formation of riffle-pool sequences: field evidence for an autogenetic process. Sedimentary Geology 85, 39–51. Derbyshire, E., 2001. Geological hazards in loess terrain, with particular reference to the loess regions of China. Earth-Science Reviews 54, 231–260. Derbyshire, E., Meng, X., 2000. Loess as a geological material. In: Derbyshire, E., Meng, X., Dijkstra, T.A. (Eds.), Landslides in the thick loess terrain of North-West China, Chichester, pp. 47–90. Derbyshire, E., Van Asch, T., Billard, A., Meng, X., 1995. Modelling the erosional susceptibility of landslide catchments in thick loess: Chinese variations on a theme by Jan de Ploey. Catena 25, 315–331. Dijkstra, T.A., 2001. Geotechnical thresholds in the Lanzhou loess of China. Quaternary International 76/77, 21–28. Ding, Z., Rutter, N.W., Liu, T., 1993. Pedostratigraphy of Chinese loess deposits and climatic cycles in the last 2.5 Myr. Catena 20, 73–91. Ding, Z., Liu, T., Rutter, N.W., Yu, Z., Guo, Z., Zhu, R., 1995. Icevolume forcing of Asian winter monsoon variations in the past 800,000 years. Quaternary Research 44, 149–159. Domroes, M., 2001. Ra¨umliche und zeitliche Variabilita¨t der Sommerniederschla¨ge in China. Geographische Rundschau 10, 36–41. Douglas, I., 1967. Man, vegetation and sediment yield of rivers. Nature 215, 925–928.
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Late Quaternary landscape evolution in a small catchment on the Chinese Loess Plateau Brigitta Schu¨tt a, *, Manfred Frechen b, Philipp Hoelzmann a, Georg Fritzenwenger a a b
¨t Berlin, Department of Earth Sciences, Physical Geography, Malteserstraße 74-100, Haus H, 12249 Berlin, Germany Freie Universita Leibniz Institute for Applied Geophysics (LIAG), Section S3: Geochronology and Isotope Hydrology, Stilleweg 2, 30655 Hannover, Germany
a r t i c l e i n f o
a b s t r a c t
Article history: Available online 6 January 2010
Late Quaternary landscape evolution was analysed with reference to a small (0.08 km2) drainage basin system in southeastern Gansu Province, southwestern Chinese Loess Plateau. In this region, V-shaped valleys with broad valley floors are prominent landscape features whose basic shape already existed prior to MIS 2. During MIS 2 the valley bottoms were infilled with fluvially reworked loess, today forming accumulation terraces. Infrared optically stimulated luminescence (IRSL) dating of these loess-like deposits assigned them to the time span ranging from 22.5 to 17.7 ka BP. During the late Holocene, valley fills were partly eroded, and the receiving stream channel was deepened by at least 25 m close to the confluence of the study site’s channels. These alternating erosion and accumulation phases were triggered by climatic factors or human impact. In the most recent past, the high landscape sensitivity of the Chinese Loess Plateau has led to increasing soil conservation measures that have effectively confined surface erosion. However, the lack of drainage on the artificial terraces induces piping processes. Ó 2010 Elsevier Ltd and INQUA. All rights reserved.
1. Introduction In China, loess deposits, reaching several hundred metres in thickness, are widespread over an area of approximately 440,000 km2 between latitudes 33 N and 47 N and longitudes 75 E and 125 E (Liu et al., 1985). North of Lanzhou, in the western part of the Loess Plateau, such deposits attain a maximum thickness of 505 m (Derbyshire and Meng, 2000), decreasing to the south and east (Sun, 2002). Liu et al. (1985) describe the stratification of a 140 m thick loess profile at Luochuan in central Shaanxi province, identifying four major units of loessial deposits with interlayered palaeosols (Ding et al., 1993). The lowest unit is the Wucheng Loess, which was deposited in the early Pleistocene (2.4–1.15 Ma). Up to 100 m thick, it is composed of a reddish to light-brown, wellcemented and fine-grained silt. The overlying Lishi Loess (1.15– 0.1 Ma) ranges between 40 and 70 m in thickness (rarely 200 m); it is poorly cemented and of a greyish brown colour. The late Pleistocene Malan Loess (100–2 ka) is 10–30 m thick. This loess type is greyish yellow, loosely cemented, and coarsely grained. It is overlain by a 2–3 m thick loess layer of Holocene age. This basic stratigraphy has been assigned and adapted to many sites of the
* Corresponding author. Tel.: þ49 30 838 70479. E-mail address: [email protected] (B. Schu¨ tt). 1040-6182/$ – see front matter Ó 2010 Elsevier Ltd and INQUA. All rights reserved. doi:10.1016/j.quaint.2009.12.018
Chinese Loess Plateau by Kukla (1987), Derbyshire et al. (1995) and Sun (2002), among many others. Previous work has shown that since the early Pleistocene there have been numerous alternations between periods of aridisation, accompanied by loess deposition, and more humid environmental conditions with (palaeo-)soil formation (An et al., 1991; Ding et al., 1995; Rost, 2000). In consequence, loess stratigraphy has been widely used to reconstruct Cenozoic palaeoenvironmental conditions of the Chinese Loess Plateau (Ding et al., 1993). The well-known high sensitivity of loess to erosion, soil erosion processes (Casali et al., 1999; Vanwalleghem et al., 2003; Porter and An, 2005; Ionita, 2006), and erosion by piping processes (Bariss, 1977; Bocco, 1991; Bryan and Jones, 2000; Derbyshire, 2001) has led to many studies on present-day morphodynamics (among others: Wu and Cheng, 2005; Hessel and van Asch, 2003; Cheng et al., 2007). However, only few authors (e.g., Bork et al., 2001) have investigated fluvial processes and soil erosion as landform-shaping modes and their reflection in a morphogenetic context. The present study investigates the relief evolution of a small catchment in the southwestern Chinese Loess Plateau. The aim is to enhance understanding of late Quaternary relief shaping by studying the chronology of fluvial erosion and deposition phases in an area of high landscape sensitivity due to the high erodibility of primary loess and its fluvially redeposited sediments under the influence of climate change and tectonic activity. Luminescence
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dating of the sediment allows the age determination of the corresponding erosion and fluvial deposition phases. 2. Regional setting The study site is a small catchment area (0.08 km2), located in the southwestern Chinese Loess Plateau, about 60 km southeast of Lanzhou, the capital city of Gansu Province. This region is known as the Longxi Plateau and is situated at the northeastern margin of the Tibetan Plateau, where crustal deformations resulted in uplift and major faulting during the Quaternary (Dijkstra, 2001). In the south and southwest, the study area is bounded by the active WNW–ESE Qinling fault zone and in the east by the active NNW–SSE striking faults of the Liupan Mountains (Derbyshire et al., 1995). In the area of the Longxi Plateau, Quaternary loess was deposited on a slightly hilly relief. The underlying bedrock shows evidence of major tectonic disturbance and exhibits a variety of folds and faults, creating an extremely unstable subsurface for the overlying loess formation. The resulting late Cenozoic uplift of the area caused increasing relief and significant erosion by mass movements and fluvial processes (Derbyshire and Meng, 2000). The present climate of the southwestern Chinese Loess Plateau is highly seasonal, with cold, dry winters and warm, humid summers. During winter, cold and dry air masses move between eastern Sibiria and northern China, causing temperature and precipitation to drop (Porter, 2001), while the Siberian cold continental high-pressure zone builds up and expands. Air pressure at ground level on the adjoining Tibetan Plateau is lower owing to its latitude and stronger altitude-induced insolation. The resulting northerly winds are cold and dry (Shahgedanova, 2003). By contrast, differential summer heating between the Tibetan Plateau and the south- and eastwardly adjoining Indian Ocean and the East Chinese Sea results in the flow of moist, southeasterly air to the hot continental low-pressure system of the Tibetan Plateau (Rokosh et al., 2002). These monsoonal currents also affect the fringe of the Tibetan Plateau and adjacent lowlands, locally overlapping with thunder cells due to heating where wetlands occur (Lehmkuhl and Haselein, 2000). At present, the progressive expansion of the summer monsoon, with its predominantly southern to southeastern winds, reaches the study area in late July each year and gradually declines from late August onwards. Accordingly, the rainy season generally lasts 2–3 months. The expansion and the strength of the summer monsoon depend on the position and development of the Northern Pacific subtropical high-pressure zone (Domroes, 2001). As the study area is located at the northern margin of the monsoonal realm, precipitation is reliable during the summer rainy season. However, total annual precipitation is highly variable. The mean annual precipitation measured at the neighbouring climate stations totals 328 mm a1 at Lanzhou and 327 mm a1 at Yu Zhong; more than 70% of this rainfall occurs between July and September. More than 60% of the precipitation falls during rainstorm events, whose intensity often exceeds soil infiltration capacity and in consequence causes surface runoff (Feng et al., 2007). 3. Methods 3.1. Chronology The samples were dated using the infrared optically stimulated luminescence dating technique (IRSL). Five samples of fluvially reworked loess were extracted at different river terrace levels (accumulation terraces) along the study site’s channel and its receiving stream (Fig. 2, Table 1).
Table 1 Palaeodose values and loss deposition ages using infrared optically stimulated luminescence. Sample field No.
Gamma spectrometry K (%) Th (ppm)
U (ppm)
OSL0807-1 OSL0807-2 OSL0807-3 OSL0807-4 OSL0807-7
1.66 1.70 1.72 1.82 1.72
3.4 0.03 3.3 0.03 3.4 0.02 3.0 0.02 3.1 0.02
13.2 0.07 12.2 0.06 12.8 0.04 10.8 0.06 11.6 0.06
Cosmic dose Palaeodose Age (ka) rate (mGy/yr) (Gy) 219.0 11.0 107.6 2.5 22.5 1.5 113.0 5.7 93.8 2.9 20.7 1.4 182.0 9.1 84.7 2.5 17.8 1.2 171.0 8.6 78.1 2.0 17.7 0.8 160 8.0 85.2 3.8 19.1 1.5
Luminescence dating of aeolian deposits has proved to be successful where radiocarbon and other dating methods are not applicable (Frechen, 1999; Roberts, 2008). The basic principle of luminescence dating is solid state dosimetry of ionising radiation (Aitken, 1998; Lian and Roberts, 2006). Luminescence is the light emitted from crystals such as quartz, feldspar or zircon when they are stimulated with heat or light after receiving a natural or artificial radiation dose. The equivalent dose (De) is a measure of the past radiation energy absorbed in natural dosimeters like quartz and feldspar minerals and, in combination with the dose rate (the rate of radiation absorbed per unit time), yields the time elapsed since the last exposure to sunlight. Infrared optically stimulated luminescence (IRSL) dating was carried out to determine the deposition age (the time elapsed since the last exposure to daylight or sunlight) for the reworked loess sediments from the sections under study. Experimental details were applied following the IRSL dating approach of Frechen et al. (2009) and so are only summarised here. IRSL measurements were carried out on five samples. Polymineral fine-grained material (4–11 mm) was prepared to determine the equivalent dose, as described by Frechen et al. (1996). The samples were irradiated by a 90Sr/90Y beta source (0.172 Gy s1) in at least seven dose steps with five discs each and a maximum radiation dose of 930 Gray (Gy). All discs were stored at room temperature for at least four weeks after irradiation. The irradiated samples were preheated for 1 min at 230 C before infrared stimulation. A Schott BG39/Corning 7-59 filter combination was placed between photomultiplier and aliquots for IRSL measurements. Each aliquot was kept at a temperature of 50 C during 10 s of IR decay. The equivalent dose was obtained by integrating the 1–5 s region of the IRSL decay curves. Alpha efficiency was estimated to a mean value of 0.08 0.02 for all samples. Dose rates for all samples were calculated from potassium, uranium and thorium contents, as measured by gamma spectrometry (N-type high purity Germanium (HPGe) detector with 25% relative efficiency) in the laboratory, assuming radioactive equilibrium for the decay chains. Cosmic dose rate was corrected for the altitude and sediment thickness, as described by Aitken (1998) and Prescott and Hutton (1994). The natural moisture content of the sediment was estimated at 10 3%. In fluvially reworked sediments, partial bleaching prior to deposition is one of the major causes of significant age overestimation (Frechen et al., 2008). The existence of scatter in De value determination is an indication for incomplete zeroing of the OSL signal but was not observed for the samples under study. As polymineral fine-grain material was used for dating, statistical methods could not be applied to study partial bleaching. Although there is no evidence for partial bleaching, the results should be considered as maximum deposition ages. Wintle (1973) describes the problem of anomalous fading, which is a loss of charge with time results in age underestimation. A delay of at least four weeks between artificial irradiation and measurement
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was applied to minimise the age underestimation caused by anomalous fading. 3.2. Geomorphological mapping Landform unit identification and mapping were based on the LANDSAT-7 ETMþ satellite image (p130/r035, 18.07.1999). The digital elevation model (DEM) was generated from the data provided by the ASTER satellite images (Advanced Spaceborne Thermal Emission and Reflection Radiometer, images from 14.07.2001) with a spatial resolution of 30 m 30 m. A geomorphological field survey was carried out in August 2007 following the mapping guidelines of Leser and Sta¨blein (1975). High resolution topography was measured along the sampling sites with a measuring tape and a clinometer. 4. Results 4.1. Relief The study site is a small drainage basin with an area of 0.08 km2 and a vertical extent of 150 m; the highest point of the divide is at 1710 m above sea level (asl). The total drainage basin length is 220 m; the main channel is 95 m long and drains from east to west. The study site’s streams are ephemeral, bearing runoff only after rainfall events. The receiving stream is also ephemeral and flows into a tributary of the Huang River (Huang He) (Fig. 1). The receiving stream’s catchment, which contains the study site, covers 2.32 km2 (Fig. 1). Loess is the parent material in the whole catchment. The summit altitudes of the divides reach 1800–2000 m asl. Total drainage basin length is 600 m, with a stream length of 520 m and a most recent channel incision of about 120 m. Correspondingly, slope gradients widely exceed 25 and the relative relief (ratio of maximum difference in altitude and drainage basin length) ranges between 0.15 and 0.35. Depending on slope gradient and slope aspect, the vegetation cover density varies between 0 and 50%, with an average of 28.5%; the lowest vegetation cover is on the south-facing slopes. Along the lower course of the receiving stream, fluvial accumulation terraces are visible at three different levels along the main drainage way (Fig. 1). At the confluence of the study site’s channel and the receiving stream, the latter’s channel is boxshaped, and its floor is located approximately 25.5 m below the valley bottom of the tributary channel draining the study site. The lowest accumulation terrace is located 15 m above the valley floor of the receiving stream, followed by a second terrace 8 m upslope (23 m above the receiving stream’s valley floor), and a third accumulation terrace after another 2.5 m (25.5 m above the receiving stream’s valley floor). The study site’s drainage pattern is dendritic, with three secondorder channels (after Horton, 1945) draining the headwater area and feeding the main channel. In the upper part of the study site, channels indicate the temporary occurrence of concentrated runoff, and the high frequency of riffle-pool sequences indicates turbulent runoff (Clifford, 1993) as an expression of high flow velocity (Leopold et al., 1992); both factors are also reflected in the occurrence of V-shaped valleys (Fig. 2). The lower part of the study site is characterised by a gully-like, box-shaped valley. This valley section has a total length of 95 m; the valley bottom is infilled, and the channel is up to 65 m wide (Figs. 2 and 3). The longitudinal profile of this valley section is step-like, and its floor is adjusted to the uppermost terrace level of the receiving stream: in four steps the channel of the box-shaped valley overcomes 30 m difference in altitude. Sediments exposed along the steps are layered and composed of silt. These sediments are termed ‘loess-like’ as the origin of these
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valley infills is loess that has been removed by hill wash and fluvial erosion (Wenske et al., 2011). Valley cross-profiles (Fig. 2) are measured for the headwater area (A), the confluence of the three major receiving gullies (B) and the lower course before it drains into the main valley (C). Crossprofile A shows two V-shaped valleys with slopes up to 45 . Typical steppe vegetation such as Artemisia and Stipa is present all along the cross-profiles. Canopy cover varies from 5% on steep slopes to more than 50% on artificial terraces, where slope amounts to less than 10 . The north-facing slope is shaped by up to 4 m wide terraces, created for soil conservation. However, at present the site is not used for agriculture. These terraced slopes lack linear erosion features. Instead, several piping forms indicate subsurface runoff; they occur predominantly at the terrace edges. Correspondingly, piping occurs throughout the drainage basin of the receiving stream, wherever terraces have been built as soil conservation measures. A pipe inventory shows that 30% of the pipes recorded appear in artificial terraces, which cover about 15% of the total catchment area. Cross-profile B is located immediately downstream of the confluence of the three channels draining the headwater area. The maximum slope is 45 in this transition zone between the upper headwater area and its lower part. The north-facing slopes have been shaped by artificial terraces built for soil conservation purposes. In the downslope parts, terrace width increases up to 8–9 m, corresponding to the slope angle of 25 . The flat topography of the channel indicates an infilled valley. The vegetation cover resembles that of cross-profile A: Artemisia and Stipa dominate, with the densest canopy cover (more than 50%) along the flat valley floor. The flat topography in the central part of cross-profile C again marks the valley infill. Maximum channel width is 30 m; adjacent slopes incline by as much as 70 . Small mass movements occur all along the steep slopes. Valley infills are riddled by piping.
4.2. Chronology In the downstream section of the study site the main drainage way (Figs. 3 and 4) is divided into four levels (L1–L4) by a step-pool sequence with three steps. All levels are almost planar, owing to the fluvially infilled valley floor and separated from each other by steep, cascade-like steps. The valley fills are exposed at the steps and sediments show a clear layering with occasional cross beds, confirming deposition by fluvial processes. To reconstruct valley evolution, four samples were extracted from the valley fills exposed in the steps (Figs. 3 and 4). OSL0807-1 was sampled from the valley fill of section L2 at 1.1 m below the present-day valley bottom; IRSL dating yielded an age of 22.5 1.5 ka BP. OSL0807-2 and OSL0807-3 were extracted from section L3, with OSL0807-2 sampled from the valley fill at 5.4 m, and OSL0807-3 at 2.0 m below the present-day valley floor. IRSL dating yielded an age of 20.7 1.4 ka BP for OSL0807-2 and 17.8 1.2 ka BP for OSL0807-3. OSL0807-4 was sampled from the valley fills of section L4, extracted 2.0 m below the valley floor; IRSL sediment ages are 17.7 0.8 ka BP (Table 1). In addition, sediments from the accumulation terraces of the study site’s receiving stream were sampled for dating immediately downstream of the confluence (Fig. 2). Four samples were taken from the three recorded accumulation terraces of the valley fill. OSL0807-7 was taken from the terrace level L2, 7.0 m below surface and about 16 m above the present-day valley floor. The sediment structure indicates that accumulation processes were primarily fluvial; the silty texture underlines that the deposits derive from fluvially transported loess in the headwater area. IRSL dating of OSL0807-7 yielded an age of 19.1 1.5 ka BP (Table 1).
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Fig. 1. Location of the study site. The overview map (left) shows the location in the southwestern Chinese Loess Plateau (a) with regard to the Yellow River as the major regional drainage system (b; map source: Fu et al., 2006). The detailed map (c) gives an overview of the geomorphology of the receiving stream’s drainage basin; the location of the study site at the right lower course is marked by the black frame (elevation data are based on Aster DEM, 2001).
5. Discussion 5.1. Present-day processes According to Derbyshire (2001) the study site is located in the highest collapsibility zone of the Loess Plateau. It displays severe erosional damage such as small mass movements, gully erosion and piping (Figs. 2 and 4). This is mainly due to the texture of loess, which causes high soil erodibility (Planchon et al., 1987). Throughout the headwater area, erosion features dominate along the slopes and control present-day morphodynamics, whereas accumulation features are absent. In their studies of the Chinese Loess Plateau, Valentin et al. (2005) and Guanglu et al. (2004) observe that slope steepness and absence of vegetation cover are the major factors contributing to rill and gully erosion. These results compare to some degree with the study site: rill erosion is initiated at all slopes with poor vegetation cover and a slope gradient of more than 5 , progressing to gully erosion with increasing slope gradient. The construction of artificial terraces increases slope stability at the headwater area of the study site and explains the absence of linear erosional forms. Most artificial terraces are bounded by low
earth benches to retain runoff. Consequently, the water infiltrates in small desiccation cracks at the edge of the terraces, where the hydraulic gradient is highest, and enters the subsurface drainage system. Piping develops especially in these positions. Thus, piping forms at the study site are frequently caused by the lack of drainage at terraces created as soil conservation measures (Hudson, 1995). Accordingly, the slope surface is not exposed to erosion processes, whereas resulting subsurface runoff causes severe erosion damage, superficially visible where pipe systems collapse and at worst build gully heads (Derbyshire, 2001). With regard to gully formation, Bocco (1991) observes that 60% of the gullies in semi-arid areas were formed as a result of piping; because the subterranean runoff systems lead to instability and subsidence of the soil, the resulting small depressions are starting points for further gully formation. 5.2. Late Quaternary landscape history The lower part of the study site is dominated by a gully-like valley. The sediment character of the outcrops along the cascadelike steps clearly indicates that present-day valley shaping results from the very recent dissection of the fluvially deposited infill of an
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Fig. 2. Geomorphological map of the study site. Black squares show sampling locations for luminescence dating with samples 1–4 extracted from the valley infills of the study site and samples 5–7 extracted from accumulation terraces of the receiving stream. Diagrams on the right show valley cross-profiles A–C.
older, more strongly incised valley. The valley fill consists of layered silts, originating from erosion of the loess of the tributary slopes and headwater divides. The dated samples scatter between 23 and 17 ka and are in stratigraphic order with the oldest sample in the deposits exposed downstream and the youngest in the deposits exposed upstream. Ages of the valley fill document that the infill process occurred during MIS 2, around the Last Glacial Maximum. Investigation of palaeoenvironmental proxies (by, among others, Frenzel, 1994; Schlu¨tz, 1999; Yan et al., 1999; Zhou et al., 2001; Hong et al., 2003) indicate alternating, but overall cold and semi-
arid conditions changing to dry-subhumid conditions between 18.5 and 17.0 ka BP. Feng et al. (2007) verified this statement for the investigated area. For the Chinese Loess Plateau, Chen et al. (1997) found that the post-glacial climate not far from the study site was slightly drier between 17.2 and 15.4 ka BP. The shift to the modern atmospheric circulation pattern with an intensified Asian monsoon circulation with increased summer rainfall set in around the Bølling-Allerød interstadial (e.g., Wang et al., 2001). It is assumed that the trigger for this change in monsoon circulation was the warming of the North Atlantic Ocean (Overpeck et al., 1996).
Fig. 3. Longitudinal profile of the main thalweg. Luminescence samples derive from different levels. Further details in Table 1.
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Fig. 4. Sample location. The photo shows the tributary catchment with 4 sample sites. (a) marks sample OSL0807-1, (b) OSL0807-2, (c) OSL0807-3 and (d) OSL0807-4.
Given these overall climatic conditions, dominant fluvial processes during MIS 2 corresponded to the processes known today from drylands: periodic to erratic rainstorm events generate concentrated surface runoff, whereas events with low rainfall intensity are not runoff-effective because of the high infiltration capacity of loess (Le Bissonais, 1990). The ubiquitous fines of the parent material in the Chinese Loess Plateau are easily eroded and cause high suspended load in the concentrated runoff. Downstream infiltration loss, which is typical for dryland rivers, causes a sediment overload and the deposition of valley fills (Schwanghart and Schu¨tt, 2008). This process is facilitated by the downstream decreasing slope of the channel. The same process occurres at different drainage basin scales. A similar formation process is likely for the accumulation terrace systems of the receiving stream. In view of the high erosional sensitivity of loess sediments in this area, accumulation rates of up to 6 m are consistent with short runoff periods. The time frame of the valley fills and terrace accumulations preclude a derivation from soil erosion processes resulting from human activity. As the Asian monsoon circulation intensified during the Holocene, humidity increased and the vegetation cover became denser (Zhou et al., 2001). Thus, it may be concluded that intensive rainfall events characterised the rainfall-runoff regime both during MIS 2 and during the subsequent Holocene (MIS 1). Such storm events were erratic during MIS 2 but became more periodic with the change to MIS 1. During the Holocene, increased annual precipitation and more reliable rainfall occurrence facilitated vegetation growth, which in turn confined erosion processes (Hudson, 1995). These relationships are controlled by the decreasing ratio of annual mean precipitation to effective annual mean precipitation (runoffeffective precipitation) with increasing vegetation cover (Langbein and Schumm, 1958; Douglas, 1967). Hence, the area’s relief was stable as long as a natural vegetation cover was present. Landscape destabilisation and soil erosion started after clearing and cultivation. With reference to the piedmont of the Qinling Mountains,
Huang et al. (2006) showed that significant and sustained soil erosion processes – as documented in the deposition of colluvial deposits – started approximately during the Bronze Age (Xia and Shang Dynasties). By contrast, early Neolithic cultures destabilised the landscape system only temporarily (Huang et al., 2006). Accordingly, it is assumed that the LGM valley fills were dissected during the early Holocene. However, correlative sediments are lacking. The data set obtained for the fluvial terrace (OSL0807-7) is not sufficient for further chronological interpretation. With no independent age control, IRSL age estimates from single samples must be regarded with caution (Lehmkuhl et al., 2007). However, the present IRSL ages are considered to be maximum deposition ages and so provide a rough chronological frame for the sediments under study. Data validation is limited by the lack of any data on soils, palaeosols or organic material. Nevertheless, the dated age of the second fluvial terrace (19.1 1.5 ka BP) is in accordance with results from the tributary catchment. 6. Conclusion This study of a small drainage basin shows that the strong relief, characterised by deeply incised V-shaped valleys, originated during MIS 2 at the latest. During this time the level of the receiving stream channel was approximately 25 m higher in altitude than today (w1700 m asl), documented by alluvial deposits at this elevation. The study site’s drainage was aligned to this level. Prior to MIS 2, the tributary valley was strongly V-shaped. This may have been caused by climate-triggered incision processes as well as by tectonic activity (Dijkstra, 2001); both are high-probability factors likely to have affected the shaping of the valley. During MIS 2, incision processes stopped, and hill wash and fluvial accumulation processes created valley fills and accumulation terraces. These processes are typical for cold, dry climates in areas with erratic rainfall, high sediment supply, and downstream decreasing runoff
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due to infiltration loss. However, since the study site is located in an area of high tectonic activity, this factor cannot be excluded as an additional influence affecting relief decline. During the early Holocene, enhanced humidity and a denser vegetation cover stabilised the landscape, coinciding with limited erosion. It is assumed that dissection of the valley fills is due to human impact and, thus constitutes a soil erosion process. The onset of soil erosion processes coincides with intensified land use during the late Holocene; prior to this, the natural vegetation has been disturbed only occasionally. At the same time, the receiving stream channel was deepened by at least 25 m close to the confluence of the study site’s channels. Here too, the possible erosion-triggering impact of tectonic activity has to be taken into account as a controlling factor, in addition to human-induced changes in runoff behaviour. Today, the land surface is again stable: artificial terraces reduce the local relief and promote infiltration. Bush planting offers additional protection against soil erosion. However, insufficient drainage of the artificial terrace systems entails severe piping processes which locally have even developed into gully erosion. Dense vegetation cover on the slopes mitigates soil erosion, so rill erosion is rare. Acknowledgments The authors would like to thank especially Silke Wilhelm from Freie Universita¨t Berlin for her help during field work and all the Chinese, Mongolian and German students of the Eurasia Network for their assistance during the Summer School on Landscape Sensitivity in 2007. We thank Anne Beck for English-language editing. References Aitken, M.J., 1998. An Introduction to Optical Dating. Oxford University Press, Oxford, pp. 280. An, Z., Kukla, G.J., Porter, S.C., Xiao, J., 1991. Magnetic susceptibility evidence of monsoon variations on the Loess Plateau of central China during the last 130,000 years. Quaternary Research 36, 29–36. Bariss, N., 1977. Gullying in a semi-arid loess region of North America. Great PlainsRocky Mountains Geographical Journal 6 (2), 125–132. Bocco, G., 1991. Gully erosion: processes and models. Progress in Physical Geography 15, 392–406. Bork, H.-R., Li, Y., Zhao, Y., Zhang, J., Shiquan, Y., 2001. Land use changes and gully development in the Upper Yangtze River Basin, SW-China. Journal of Mountain Research 19 (2), 97–103. Science Press, Chengdu. Bryan, R.B., Jones, J.A.A., 2000. The significance of soil piping processes: inventory and prospect. Geomophology 20, 209–218. Casali, J., Lopez, J.J., Giraldez, J.V., 1999. Ephemeral gully erosion in southern Navarra (Spain). Catena 36, 65–84. Chen, F., Bloemendal, J., Wang, J., Oldfield, F., 1997. High-resolution multi-proxy climate records from Chinese loess: evidence for rapid climatic changes over the last 75 kyr. Palaeogeography, Palaeoclimatology, Palaeoecology 130, 323–335. Cheng, H., Zou, X., Wu, Y., Zhang, C., Zheng, Q., Jiang, Z., 2007. Morphology parameters of ephemeral gully in characteristics hillslopes on the Loess Plateau of China. Soil and Tillage Research 94, 4–14. Clifford, N.J., 1993. Formation of riffle-pool sequences: field evidence for an autogenetic process. Sedimentary Geology 85, 39–51. Derbyshire, E., 2001. Geological hazards in loess terrain, with particular reference to the loess regions of China. Earth-Science Reviews 54, 231–260. Derbyshire, E., Meng, X., 2000. Loess as a geological material. In: Derbyshire, E., Meng, X., Dijkstra, T.A. (Eds.), Landslides in the thick loess terrain of North-West China, Chichester, pp. 47–90. Derbyshire, E., Van Asch, T., Billard, A., Meng, X., 1995. Modelling the erosional susceptibility of landslide catchments in thick loess: Chinese variations on a theme by Jan de Ploey. Catena 25, 315–331. Dijkstra, T.A., 2001. Geotechnical thresholds in the Lanzhou loess of China. Quaternary International 76/77, 21–28. Ding, Z., Rutter, N.W., Liu, T., 1993. Pedostratigraphy of Chinese loess deposits and climatic cycles in the last 2.5 Myr. Catena 20, 73–91. Ding, Z., Liu, T., Rutter, N.W., Yu, Z., Guo, Z., Zhu, R., 1995. Icevolume forcing of Asian winter monsoon variations in the past 800,000 years. Quaternary Research 44, 149–159. Domroes, M., 2001. Ra¨umliche und zeitliche Variabilita¨t der Sommerniederschla¨ge in China. Geographische Rundschau 10, 36–41. Douglas, I., 1967. Man, vegetation and sediment yield of rivers. Nature 215, 925–928.
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