13 Late Pleistocene-Holocene Palaeohydrology of Monsoon Asia

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Content: 13 Late Pleistocene­Holocene Palaeohydrology of Monsoon Asia V.S. KALE,1 A. GUPTA2 AND A.K. SINGHVI3 1University of Pune, Pune, India 2University of Leeds, Leeds, UK 3Physical Research Laboratory, Ahmedabad, India 1 INTRODUCTION The monsoon system is a thermodynamic atmospheric circulation, characterised by strong seasonality of wind direction, temperature and precipitation (Ramage, 1971). The largest monsoon-dominated region in the world is in Asia (Figure 13.1). The Asian monsoon comprises the southwest (Indian) monsoon and the southeast (East Asian) monsoon. The former is the major source of precipitation over the Indian subcontinent and the western part of Southeast Asia and the latter is the dominant influence over the eastern part of southeast Asia and east Asia. The two systems, although largely independent, interact and play a significant role in the global hydrologic cycle (An et al., 2000; Kudrass et al., 2001). Multi-proxy records from China and the Indian and North Pacific Oceans denote that the Indian as well as the east-Asian monsoon systems were established about 8 million years ago (An et al., 2001). There is evidence of a stronger monsoon system at 3.5 million years and 2.6 million years (Qiang et al., 2001). Continental and marine palaeoclimatic records further indicate that since the onset of glaciation in the northern hemisphere around 2.5 million years ago (Shackleton et al., 1984), the strength of the Asian monsoon has varied on both long and short timescales (Overpeck et al., 1996; Liu et al., 1999; Lu et al., 1999; Qiang et al., 2001). These changes in the strength of the monsoons were linked to global processes (Schulz et al., 1998; Wang et al., 1999; Kudrass et al., 2001), and had a profound impact on the palaeohydrology, palaeogeography and geomorphology of monsoon Asia. This chapter attempts to reconstruct a framework for the late-Quaternary palaeoclimatic and palaeohydrological changes in monsoon Asia from the large number of investigations published so far (Table 13.1), from the Last Glacial Maximum (LGM) · Q1 to the present, that is covering the last 18 14C kyr BP· (radiocarbon) or 21.5 cal kyr BP (calibrated). 2 PALAEOHYDROLOGY DURING THE LAST GLACIAL MAXIMUM During the last glacial period, only the highly elevated areas of monsoon Asia, such as the Tibetan Plateau and the Himalayan Mountains, provided favourable conditions for the formation of large glaciers (Seltzer, 2001) and ice caps. Three ice caps on the Tibet Plateau, namely, Dunde, Guliya and Dasuopu (Figure 13.1) have provided evidence Palaeohydrology: Understanding Global Change. Edited by K.J. Gregory and G. Benito 2003 John Wiley & Sons, Ltd ISBN: 0-470-84739-5
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Indus Chang Jiang SOUTHEAST MONSOON
1
6
2
7
Taklamakan
3
8
Gobi
Rub al Khali
4
9
Du
5
Hindu Kush
SG
Q
Huanghe
H
TIBET
IM
L Thar
A
L Ganga
A
YA
Brahmaputra
Narmada Godavari
SOUTHWEST
Krishna
Arabian Sea
MONSOON
Ng
SOUTHWEST
Bay of Bengal
MONSOON
Mekong Irrawady
Yellow Sea
Xi
Nj
East China
Sea
Pearl Ty South China Sea DA SU N
km
0
000
INDIAN OCEAN
Figure 13.1 Map showing some important sites with climate proxy records mentioned in the text. The dashed lines with arrows show the wind patterns for the Asian summer monsoon (June to September). Key: 1 ­ extinct drainage of Sundaland; 2 ­ borehole sites; 3 ­ lake/peat sites; 4 ­ ice core sites; 5 ­ fluvial/flood sites; 6 ­ deep-sea core sites; 7 ­ Loess Plateau; 8 ­ deserts; 9 ­ Qinghai­Tibetan Plateau; Du ­ Dunde ice cap; G ­ Guliya ice cap; L ­ Lunkaransar Lake; Ng ­ Nilgiri; Nj ­ Nanjing flood site; Q ­ Qinghai Lake; S­Sumxi Co; Ty ­ Tianyang Lake; Xi ­ Xiaolangdi palaeofloods site (Thompson et al., 1989; 1997; 2000) of ­ (1) a pronounced climatic instability in the tropics; and (2) century-to-millennium scale fluctuations in the monsoon hydrologic cycle since the last glacial stage and even earlier. Ice-core records from Tibet suggest that the climate during the LGM was cool, dry and variable (Thompson et al., 1997). Evidence indicates the presence of large glaciers, depression of the glacier snowline by several hundred meters and atmospheric cooling of 5 to 7C during the Last glacial (Lehmkuhl et al., 1999; Seltzer, 2001). On the basis of pollen records from the Tianyang Lake (Figure 13.1), Zheng and Lei (1999) have inferred that the drop in temperature and precipitation during the LGM was much greater than during the three previous glacial periods (Oxygen Isotope Stages 6, 8, 10). Consequently, it is reasonable to suggest that during the LGM, the hydrological regime was notably different from that of the present. A coeval weakening of the summer monsoon over Asia has been indicated by­(1) past lake levels and lake sediment/peat in China (Fang, 1991; Gasse et al., 1991; Zheng
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Table 13.1 Proxy palaeoclimatic records from monsoon Asia
Proxy records
Area
References
Lakes/peat Lakes/peat Loess Boreholes Speleothems, calc tufa and groundwater · Q2 Ice cores Deep-sea cores
China Thailand India China Bangladesh China Nepal India Tibet ­ China Bay of Bengal/ Andaman Sea Arabian Sea South China Sea
Fang (1991; 1993), Gasse et al. (1991), Lister et al. (1991), Zheng and Lei (1999), An et al. (2000) Kealhofer and Penny (1998) Bhattacharyya (1989), Singh et al. (1990), Sukumar et al. (1993), Mazari et al. (1996), Kusumgar et al. (1995), Prasad et al. (1997), Enzel et al. (1999), Phadtare (2000) Kukla and An (1989), An et al. (1991), An et al. (1993), Maher et al. (1994), Porter and An (1995), Weijan et al. (1996), Chen et al. (1997) Goodbred and Kuehl (2000) Hori et al. (2001) Denninston· et al. (2000) Pawar et al. (1988), Yadava and Ramesh (1999), Sukhija et al. (1998) Thompson et al. (1989), Thompson et al. (1997), Thompson et al. (2000) Colin et al. (1998), Ahmed et al. (2000), Sangode et al. (2001) Van Compo et al. (1982), Duplessy (1982), Sirocko et al. (1993), Caratini et al. (1994), Overpeck et al. (1996), von Rad et al. (1999), Sarkar et al. (2000), Thamban et al. (2001) Huang et al. (1997), Wei et al. (1998), Wang et al. (1999), Pelejero et al. (1999)
and Lei, 1999; An et al., 2000) and India (Bryson and Swain, 1981; Singh et al., 1990; Sukumar et al., 1993; Phadtare, 2000); and (2) isotopic, microfaunal and geochemical data from the Arabian Sea (Van Compo et al., 1982; Sirocko et al., 1993) and the South China Sea (Huang et al., 1997; Wang et al., 1999). A contrasting strengthening of the winter monsoon during the Last glacial stage and evidence of aridity and humidity associated with global cooling and warming, respectively, up to the millennium scale have also been deduced from land and oceanic records (Duplessy, 1982; Singh et al., 1990; Porter and An, 1995; Colin et al., 1998; Schulz et al., 1998; Wang et al., 1999). The Asian Highlands (Himalaya, Tibet and Myanmar Ranges) are the source area of many large rivers including the Ganga, Brahmaputra, Indus, Mekong and Changjiang. The headwaters of these rivers are sustained at present by both glacial melt water and orographic monsoon rainfall. During the LGM, cold and dry conditions prevailed over
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much of the area, the southwest monsoon was weak and less moisture reached the headwaters of these large rivers (Duplessy, 1982). It is likely that even the springtime melt water was reduced (Emeis et al., 1995). Consequently, the rivers must have experienced a significant reduction in water and sediment discharge, such that many rivers experienced only highly seasonal or ephemeral flows. River discharges to the Bay of Bengal and the Arabian Sea were markedly reduced (Cullen, 1981; Duplessy, 1982; Cayre and Bard, 1999). Isotopic and geochemical analyses of deep confined groundwater in south India show that the period ca 18 14C kyr BP was marked by aridity (Sukhija et al., 1998). An arid phase between ca 20 and 16 14C kyr BP is also indicated by peat in Nilgiri, south India (Sukumar et al., 1993); and in southeast Asia a drop in rainfall by more than half has been suggested (Flenley, 1979; Verstappen, 1980). Planktonic foraminiferal assemblages from the South China Sea indicate a much higher temperature contrast between Summer and winter (ca 9C) during the Last glacial than that during the Holocene (Wei et al., 1998), suggesting a greater seasonality. Increased aridity and seasonality during the LGM corresponds to reduced streamflow and flow magnitudes in the rivers of south and southeast Asia (Kale and Rajaguru, 1987; Verstappen, 1997). At present, tropical cyclones forming over the adjoining seas are common components of the monsoon and are an important source of moisture. Pronounced continentality and cooler sea surfaces during the glacial times (Ruddiman, 1984) should have caused a diminution in the intensity and frequency of tropical storms. Oceanic reconstructions indicate fewer tropical cyclones in the South China Sea during the glacial period (Wang et al., 1999). A decrease in the incidence of typhoons over Japan has been inferred by Sugai (1993). By implication, similar conditions should have prevailed in the seas adjoining the Indian subcontinent. The reduction in the flow magnitude and the increased seasonality had a striking effect on the fluvial systems originating in the Asian highlands. Estimates of palaeomonsoon precipitation across the Chinese Loess Plateau indicate great decreases in rainfall in Central China during the glacial periods (Maher et al., 1994). In south China, the water discharge of the Changjiang River decreased remarkably, and the river occasionally became dry, exposing a 3- to 10-km-wide sandy channel bed. The dry sand bed provided the material for sand dunes on the southern bank of the river. Several phases of sand dune accretion indicate that the river was flowing only intermittently (Liu et al., 1997). Studies in Thailand by Loeffler et al. (1984) also suggest greatly reduced river flows and sand dune activity in the seasonal rivers. Apart from the decrease in precipitation and runoff, the cool and dry conditions during the LGM imply reduced weathering, affecting the sediment load of the Asian rivers. This is suggested by the clay mineralogical studies of sediments in the Bay of Bengal and the Andaman Sea deposited by the Ganga­Brahmaputra and the Irrawaddy rivers, respectively (Colin et al., 1999). Sangode et al. (2001) suggest that the sediment supply to the Bay of Bengal during the LGM increased proportionally from rivers of peninsular India as the southwest monsoon weakened over the Himalaya. The LGM period was also a period of eustatic low sea level (Fairbanks, 1989), when a large currently offshore area of southeast Asia and south Asia was exposed (Emmel and Curray, 1982; Pelejero et al., 1999). The exposed area on the Sunda shelf off southeast Asia was a coastal lowland drained by several major river systems (Gupta et al., 1987). Large rivers of monsoon Asia must have extended their courses across
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the exposed sea floor. Evidence of such courses is seen in submerged canyons (Liu et al., 1992; von Rad and Tahir, 1997; Goodbred and Kuehl, 2000) or deltas (Emmel and Curray, 1982; Wang et al., 1999). Evidence of increased continentality has been found in many parts of monsoon Asia (Kale and Rajaguru, 1987; Prins and Postma, 2000). Reduced discharges in the rivers of peninsular India have been related to both the weakening of the summer monsoon and the seaward shifting of the shoreline by more than 200 km along the West Coast of India (Kale and Rajaguru, 1987). Similar conditions should be expected for other rivers of monsoon Asia that were affected by a pronounced shoreline shift (ca 300­600 km). Sedimentary response to such a drastic decrease in precipitation, water discharge and a falling sea level during the LGM is covered by relatively few studies. Williams and Clarke (1984), Kale and Rajaguru (1987) and Joshi and Kale (1997) have noted periods of aggradation in the interior of the Indian Peninsula during the LGM, Kale and Rajaguru (1987) recorded deposition by non-meandering and bedload-dominant streams during the glacial period in northwest Deccan, and Tandon et al. (1997), Juyal et al. (2000) and Srivastava et al. (2001) found evidence of disruption of drainage systems in western India. In the Indian Desert, drainage was seriously affected and the fluvial processes were largely dormant because of increased aridity (Kar et al., 2001). Segmentation of streams by aeolian activity was a feature of this period (Kar, 1990). In western Nepal, Monecke et al. (2001) found deposits of highly mobile, braided rivers produced under glacial conditions. Gupta et al. (1987) are of the view that the Old Alluvium of Singapore was laid down by braided rivers that were characterised by seasonal flows and large floods. The presence of similar deposits in southeast Asia suggests the widespread occurrence of seasonal rivers (Loeffler et al., 1984). In central China, Porter et al. (1992) found a strong association between cold phases and stream aggradation. The upper-middle reaches of the Changjiang responded to the reduced precipitation by aggradation. In the lower reaches, incision by rivers, graded to a lower sea level of the glacial stage appears to be a common phenomenon. This is suggested by the presence of incised valleys of the Ganga­Brahmaputra and Changjiang Rivers below early Holocene sediments (Goodbred and Kuehl, 2000; Hori et al., 2001). All the climate proxies of this time, thus, indicate a cooler and drier period in monsoon Asia in contrast to warm and wet conditions of the succeeding Holocene. During the LGM a low sea level, increased seasonality, decreased rainfall and reduced frequency of tropical cyclones is indicated in a number of studies. Fluvial activity was highly variable and generally subdued.
3 THE DEGLACIAL­EARLY HOLOCENE PERIOD Sufficient oceanic evidence now indicates that the drier LGM climate ameliorated after about 13 to 12.5 14C kyr BP (15.3 to 14.7 cal kyr BP) in response to increased insolation and global warming (Sirocko et al., 1993). This conclusion has been derived mainly from Arabian Sea and South China Sea multi-proxy records (Sirocko et al., 1993; Overpeck et al., 1996; Huang et al., 1997; Wang et al., 1999); Chinese and Indian lake sediment (Singh et al., 1990; Fang, 1991; Lister et al., 1991; An et al., 2000) and Tibetan ice cores (Thompson et al., 1997). Computer modelling (CCM1) also shows that the summer monsoon strengthened significantly after 13.5 14C kyr BP (16 cal kyr BP) (Kutzbach et al., 1998). Reconstruction of Sea Surface Temperatures
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(SSTs) using oxygen isotopes of planktonic foraminifera from the eastern Arabian Sea indicates a 1.5 to 2.5C deglacial warming (Cayre and Bard, 1999). Marine records from the South China Sea also suggest increase in SSTs from the LGM to the Holocene (Wie et al., 1998; Steinke et al., 2001). Intensification of the southwest and also the southeast monsoon since the last deglaciation is currently perceived as stepwise and dramatic (Marcontonio et al., 2001). Whilst Sirocko et al. (1993) have inferred that the monsoon intensified episodically between 14.3 and 8.7 14C kyr BP in four steps, Overpeck et al. (1996) concluded that the monsoon strength increased suddenly in two steps, 13 to 12.5 14 C kyr and 10 to 9.5 14C kyr BP (ca 15.3­14.7 and 11.5­10.8 cal kyr BP). High-resolution SST records from the South China Sea show an abrupt warming (ca 1C <200 yr) at the end of the last glaciation, approximately synchronous with the Bшlling­Allerшd transition (Steinke et al., 2001). A major intensification of the southwest monsoon at 11.5 cal kyr BP, coinciding with a major climatic transition in the Greenland record, has been identified by Sirocko et al. (1996). Strengthening of the summer monsoon and related high precipitation occurred between 9.5 and 5.5 14C kyr BP (Sirocko et al., 1993; Overpeck et al., 1996). Glaciers on the Nanga Parbat expanded in the early Holocene, as a result of enhanced moisture levels (Phillips et al., 2000). In peninsular India, evidence of enhanced precipitation, discharge and groundwater levels is provided by the building of waterfall tufas between 9.8 and 8.1 ka (U/Th) in the rainshadow area of the Western Ghat (Pawar et al., 1988). Reconstruction of palaeorainfall over the Loess Plateau in China indicates much higher rainfall between 9 and 5 14C kyr BP (Maher et al., 1994). Fang (1991), on the basis of data from over 70 lakes distributed across China, found high lake levels and lake expansions between ca 9.5 and 3.5 cal kyr BP. However, a number of recent investigations indicate that the timing of the Holocene climatic optimum was not synchronous over monsoon Asia (Fang, 1991; Shi et al., 1993; Lehmkuhl, 1997; An et al., 2000). It reached a maximum in northeastern and central China ca 10,000 to 7,000 years ago, in the middle and lower Changjiang basin around 7,000 to 5,000 years ago, in southern China around 3,000 years ago (An et al., · Q3 2000), and in western China between 7,500 and 3,500 years BP (Zhou et al., 1991·). Pollen data from alpine peat suggests that in the central Himalaya, the highest monsoon intensity occurred between 6.0 and 4.5 cal kyr (Phadtare, 2000). The abrupt deglacial intensification of the monsoon had a catastrophic impact on the drainage systems of monsoon Asia. In many Asian rivers, the glacial­interglacial transition was marked by an abrupt and immense increase in the wet monsoon flow. Widespread evidence now exists to suggest shrinkage of deserts, revival of fluvial activity in the Indian subcontinent and Tibet, and increased weathering and fluvial erosion in many parts during early Holocene (Colin et al., 1999; Lehmkuhl et al., 1999; Goodbred and Kuehl, 2000). The presence of river-transported silts at ca 11.8 14C kyr BP in the Qinghai Lake record (Lister et al., 1991) indicates the revival of fluvial activity on the Tibetan Plateau. On the Deccan Plateau, evidence of widespread overbank flooding by suspended-load dominant meandering streams from ca 17 to 10 14C kyr BP has been reported by Kale and Rajaguru (1987). The increase in water discharge in many rivers at this time can also be attributed to an increase in the frequency of tropical cyclones (Huang et al., 1997; Wang et al., 1999) and reduced continentality as a result of high sea levels (Kale and Rajaguru, 1987). Appearance of storm-related early Holocene clay in the Taiwan lake sediments (Huang et al., 1997) suggests that the strengthening of the summer monsoon
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was accompanied by an increased number of tropical cyclones. Similarly, the early Holocene increase in the flood magnitude in the Ara River in Japan has been attributed to increased typhoon frequency (Grossman, 2001). Hydrological characteristics of the fluvial systems were also significantly affected in the early Holocene by the impact of a strong monsoon rainfall against the mountain chains, extensive slope failures and consequently enhanced sediment supply. Unusually high sediment output and water discharge under conditions of an intensified early Holocene monsoon in the Himalayan rivers is revealed by chronostratigraphic data from the deltaic deposits of Ganga­Brahmaputra (Weber et al., 1997; Goodbred and Kuehl, 2000). Borehole data and volume calculations indicate that an enormous amount of sediment was deposited between 11 and 7 cal kyr BP (Goodbred and Kuehl, 2000). It is most likely that the enormous deposition of sediment in the Ganga­Brahmaputra delta was preceded by, or coincided with, equally phenomenal erosion in the Himalayan Mountains and sediment accumulation at the mountain­plains interface. The deltaic sedimentation could have been penecontemporaneous with the building of mega-fans (of rivers such as Kosi, Gandak, etc.) at the foot of the Himalaya. Slope failure is a dominant denudational process in the Himalayan Ranges (Shroder, 1998), characterised by rapid uplift and high incision rates (Burbank et al., 1996). Landslides and debris flows that follow earthquakes and/or intense rainfall/snowmelt can block stream courses and create large lakes/dams. Later, sudden breaching of these dams generates catastrophic floods downstream. In addition, the advance of tributary glaciers into main valleys and the moraine dams created by retreating glaciers may also produce similar flood events. Numerous examples of floods, consequent to such dam failures have been reported from the Himalaya (Coxon et al., 1996; Hewitt, 1998; Shroder, 1998; Wohl and Cenderelli, 1998; Cornwell, 1998). The capacity and competence of such floods to transport sediment flux exceeds other denudational processes (Shroder, 1998). Evidence of postglacial enhanced temperatures and precipitation levels over the Himalaya suggests that landslide-dam, ice-dam and moraine-dam failure floods were a feature of the early Holocene humid period. Advance of glaciers, such as in the Nanga Parbat area, was favoured by increased moisture (Phillips et al., 2000). Sediment accumulation and later removal probably started to occur in the Himalaya early in the Holocene. In western Nepal, there is evidence of mobilisation and redeposition of morainic material by enormous debris flows. These debris flows were triggered by outburst floods from glacial lakes, or by strong monsoonal rains and earthquakes (Monecke et al., 2001). Most climatic proxies suggest that the stronger monsoon in the early Holocene was associated with high lake levels, increased flow discharges, floods and scouring. Evidence of huge floods and scouring in the Changjiang River is provided by a large number of uprooted trees near Ichung­Nanjing (Figure 13.1) that have yielded ages between 6 and 4.5 14C kyr BP (Yang, 1991b). Analysis of slackwater flood deposits indicates that about five large floods (>27,000 m3s-1) had occurred on the Huanghe River at Xiaolangdi (Figure 13.1) between 8 and 6 14C kyr BP (Yang · Q4 et al., 2000·). The largest known flood at 7,362 years BP on the river, with a magnitude of 42,900 m3s-1, was comparable to the catastrophic flood in 1843 AD (Yang et al., 2000). Similarly, the Pearl River, the second largest river in China in terms of discharge, also experienced a significant increase in flow (Wang et al., 1999). In Japan, Grossman (2001) deduced that large floods occurred on the Ara River in the early Holocene.
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Proxy records of the Holocene palaeohydrology in southeast Asia are somewhat sparse, but there are sufficient indications of increased precipitation and stream discharge in the early Holocene (Verstappen, 1997). Estimation of discharge by Bishop and Godley (1994) suggests greater bankfull discharges in the Yom River (Thailand) during the early mid-Holocene humid phase. By implication, other river valleys of the region would also have experienced similar changes. At present, high-magnitude floods, spaced over periods of a few years to decades, govern the channel morphology of many monsoonal rivers and the channel size increases with flood magnitude (Gupta, 1995). In the early Holocene, it is likely that the rivers responded by deepening and enlarging their channels to accommodate increased water discharges and to efficiently transport a large supply of sediments from upstream. Evidence from central China (Porter et al., 1992) and peninsular India is consistent with this notion. Widespread erosion and incision in the peninsular rivers was associated with enhanced monsoon precipitation. Rivers, such as the Son (Williams and Clarke, 1984), Sabarmati (Srivastava et al., 2001), Narmada (Gupta et al., 1999), Krishna and Godavari (Kale and Rajaguru, 1987) all responded to these changes by deepening and lowering their channels, giving rise to terraces. The early Holocene age of the soils of areally extensive interfluves in the upper Ganga Plains (Kumar et al., 1996) also implies that the Himalayan rivers were no longer aggrading. Similarly, the fluvio-sedimentary response of the Changjiang to the higher runoff in its upper-middle reaches was channel deepening (Yang, 1991a). By implication, other comparable rivers (Brahmaputra, Indus, Irrawaddy, Mekong, etc.) probably responded in a similar fashion, although neotectonic activities would have additionally accentuated incision and gorge formation (Hurtado et al., 2001). Deglaciation and a rapid rise in sea level during the early Holocene has been documented in a variety of locations, including the Indus delta (von Rad and Tahir, 1997; Prins and Postma, 2000); the Ganga­Brahmaputra delta (Goodbred and Kuehl, 2000; Banerjee, 2000); South China Sea (Emmel and Curray, 1982; Pelejero et al., 1999) and southeastern China (Chen, 1999; Zheng and Li, 2000; Hori et al., 2001). The rapid and sometimes accelerated rise (such as between 14.6 and 14.3 cal kyr; Hanebuth et al., 2000) in sea level had a significant impact on the drainage systems. The most striking evidence is from southeast Asia where the distal parts of large drainage systems, which were draining the subaerially exposed Sunda shelf during the late Pleistocene, were submerged (Gupta et al., 1987; Pelejero et al., 1999). Consequently, southeast Asia witnessed a large-scale disruption and disintegration of the drainage network. Elsewhere, the transgression was associated with rapid delta growth and delta progradation due to enormous sediment discharge from upstream (von Rad and Tahir, 1997; Goodbred and Kuehl, 2000; Hori et al., 2001; Saito et al., 2001). Evidence of cooling and strong seasonality associated with the Younger Dryas (YD) event have been found in the oxygen isotope records from the Arabian Sea (Cayre and Bard, 1999; Marcontonio et al., 2001), planktonic foraminiferal assemblages from the South China Sea (Wei et al., 1998; Steinke et al., 2001), central China loess (An et al., · Q5 1993), peat and aeolian-palaeosol sequence from east Asia (Weijian et al.,· 1996) and in lacustrine sediment from southeast Asia (Maloney, 1995). There is, however, very little direct evidence of fluvio-sedimentary response to this near-glacial event. This is not surprising as fluvial records on land lack the time-resolution and continuity required to understand the effects of shorter climatic events. A possibility of reduced discharge from the Irrawaddy and Salween Rivers into the Andaman Sea coinciding with the YD, has been suggested by Ahmed et al. (2000). There is also evidence of cessation
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of river-transported silt flux into the Qinghai Lake (China) in response to increased aridity around 10.8 14C kyr BP (Lister et al., 1991). However, there is no evidence of a YD age glacier advance from Tibet (Lehmkuhl, 1997), which, as suggested by Phillips et al. (2000), may be attributed to drier conditions during this period. After the early Holocene monsoon optimum, a progressively weakened monsoon and increasing aridity generally characterised the mid-Holocene (Steig, 1999). This change, after about 6 to 5 14C kyr BP, was much more gradual in comparison with the glacial­deglacial transition (Overpeck et al., 1996). However, because of variations in latitude, altitude and distance from the sea, the termination of the early Holocene humid phase in various basins was not synchronous. A trend towards aridity after 5.5 cal kyr BP (4.8 14C kyr BP) is indicated by marine records (Sirocko et al., 1993; Overpeck et al., 1996); lake levels in China and northwest India (Singh et al., 1990; Fang, 1991; Shi et al., 1993; Enzel et al., 1999); minor advance of glaciers in Lahul Himalaya (Owen et al., 1997) and alpine peat from the higher Himalaya (Phadtare, 2000). Borehole data from the Ganga­Brahmaputra delta also indicates a drop in sediment discharge in the Himalayan rivers after about 7 cal kyr BP (Goodbred and Kuehl, 2000). Enzel et al. (1999) have reported the desiccation of Lunkaransar Lake after about 4,800 14C kyr BP in Rajasthan. A short phase of aggradation in the Son River (Williams and Clarke, 1984) and in the upper Godavari and Krishna Basins (Kale and Rajaguru, 1987) also indicates reduction in rainfall and discharge over parts of the Indian Peninsula. The deposits associated with the T1 terrace in the Ganga Plains (Singh, 1996) were perhaps laid down during this phase. In Thailand, Bishop and Godley (1994) found evidence of lower discharges during the mid-late Holocene. Grossman (2001) noted a drop in flood magnitude on the Ara River, Japan around 5 14C kyr BP. Thus, the pattern of enhanced discharge, accelerated erosion and transported sediment of the early Holocene appears to have been replaced by a cooler, drier period (Steig, 1999), which was geomorphologically less active.
4 LATE HOLOCENE PALAEOHYDROLOGY During the late Holocene, temperature and precipitation conditions shifted in response to global climatic changes (Liu et al., 1998). Although the shift was neither abrupt nor synchronous, a conspicuous change around ca 3.5 14C kyr BP is indicated by marine, peat and lake records (Table 13.2). This was also the time of decrease in the input of fluvial mud into the Indus (von Rad and Tahir, 1997), stabilisation of the present configuration of the Ganga­Brahmaputra delta (Goodbred and Kuehl, 2000), and reduced discharge in the Yom River, Thailand (Bishop and Godley, 1994). Very few fine-resolution palaeoclimatic records exist to reconstruct millennial, century and decadal fluctuations in precipitation and runoff. Pollen from ice cores from the Dunde ice cap show evidence of relatively humid periods at 2.7 to 2.0, 1.5 to 0.8 and 0.6 to 0.8 cal kyr BP (Liu et al., 1998). The records also indicate prominent changes during the Medieval Warm (MW) period (790­620 yrs BP) and the Little Ice Age (LIA) (330­80 yrs BP) (Liu et al., 1998). This and other studies based on multiproxy data (Zhou et al., 1991; Sukumar et al., 1993; Mazari et al., 1996; Lehmkuhl, 1997; von Rad et al., 1999; Denniston et al., 2000) suggest ­ (1) relatively widespread drier periods from ca 2 to 1.5 14C kyr BP and around 1 14C kyr BP; (2) a significant warming during the MW period; and (3) a cooling during the LIA (Figure 13.3). Fluvial records lack comparable temporal resolution and the corresponding fluctuations in flow and sediment pattern have been poorly understood, although there are a
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Table 13.2 Timing of late Holocene change as indicated by palaeoclimate proxies
Climatic proxy/area Glacier ­ Tibet, China Lakes ­ China Palaeoceanographic record ­ South China Sea High Himalaya
Date 3 14C kyr BP 1000 BC 4 14C kyr BP 4­3.5 cal. yrs BP
Ice core and marine core ­ Tibet and Arabian Sea data Shallow water cores ­ Arabian Sea Upwelling indices ­ Arabian Sea Marine core ­ Arabian Sea Groundwater ­ south India Peat ­ Nilgiri, south India Loess and wind deposit ­ Korat Plateau, Thailand
3.4 14C kyr BP 3.5 14C kyr BP 3.5 14C kyr BP 3.5­4 14C kyr BP 4 14C kyr BP 3.5 14C kyr 3.5 ka
Remark Glacial advance Dramatic drop in lake levels Cooling event
Reference Lehmkuhl (1997) Fang (1993) Wei et al. (1998)
Sharp decrease in temperature and rainfall Significant increase in the biogenic and lithogenic components General weakening of the monsoon
Phadtare (2000) Rangarajan and Sant (2000) Nigam (1993)
Intensity of monsoon reduced
Naidu (1996)
Decreased precipitation Onset of unstable climate Onset of aridity Drier conditions and loess deposition
von Rad et al. (1999) Sukhija et al. (1998) Sukumar et al. (1993) Nutalaya et al. (1989)
few geological and historical records. Caratini et al. (1994) have inferred a reduction in the discharge of the Kali River into the Arabian Sea since about 2.2 14C kyr BP, due to reduction in rainfall over the Western Ghat. Slackwater palaeoflood hydrology indicates distinct periods of large and moderate floods in central and western India during the last 2 14C kyr BP (Ely et al., 1996, Kale et al., 2000). Sequences of extreme floods occurred twice: between ca 400 to 1000 AD and after the 1950s (Kale, 1999). A significantly reduced frequency of large floods between ca AD 1500 and the late 1800s may reflect the regional influence of the LIA (Kale, 1999). Nigam and Khare (1992) have identified large floods in 2000 and 1500 BC in peninsular India from oceanographic records. Rao et al. (1963) also reported archaeological evidence of these two floods from western India (Gujarat). A period of major flood activity ca 1.7 to 1.8 14C kyr BP has been inferred in northeast Thailand (Prinya et al., 1989; Bishop and Godley, 1994). Historical records are a valuable source of information on short-term climatic variability. In India, information on droughts is more extensive than that on floods. Information exists about major floods and cyclones in various river basins of India, but the records are mostly neither continuous nor uniform (Kale, 1998). In contrast, the Chinese historical documents provide a near-continuous record of thermal and
Late Pleistocene­Holocene Palaeohydrology of Monsoon Asia
223
LGM (Thompson et al., 1997)
Periods of monsoon
intensification
Monsoon peak Onset of aridity1
(Overpeck et al., 1996) (Overpeck et al., 1996) /drier condtions
Continentality Sea level
Cold­dry Low High High Low
Transitional
Warm-humid
YD Sumxi/Bangong Lakes­Tibet
Warm drier Present sea level Gasse et al. (1996)
Palaeorainfall­Loess Plateau, China Maher et al. (1994)
Loess/palaeosol­China Zhou et al. (1991)
Lakes Cn-Wn Tibet
Lake level­China
Fang (1991)
Sediment storage­Bangladesh Goodbred and Kuehl (2000)
Sea-level curve Pleistocene
Groundwater­Sn. India Lake level - India Holocene
Sukhija et al. (1998) Enzel et al. (1999)
20
18
16
14
12
10
8
6
4
2
0
Age (14C yr BP x 103)
· Q6
Figure 13.2· Periods of monsoon intensification during the last glacial/post-glacial
period indicated by proxy records. YD ­ Younger Dryas. 1 ­ After Singh, G., Wasson, R.J.
and Agrawal, D.P., 1990. Vegetation and seasonal climatic changes since the last full
glacial in the Thar Desert, Northwest India. Review of Palaeobotany and Palynology, 64,
351­358; Fang, J.Q., 1991. Lake evolution during the past 30,000 years in China and its
implications for environmental change. Quaternary Research, 36, 37­60; Overpeck, J.,
Anderson, D., Trumbore, S. and Prell, W., 1996. The Southwest Indian monsoon over the
last 18000 years. Climate Dynamics, 12, 213­225; Gasse, F., Fontes, J.Ch., Van Campo, E.
and Wei, K., 1996. Holocene environmental changes in Bangong Co basin (Western Tibet).
Part 4: discussion and conclusions. Palaeogeography, Palaeoclimatology, Palaeoecology,
120, 79­92; Lehmkuhl, F., 1997. Late Pleistocene, late-glacial and Holocene glacier
advances on the Tibetan plateau. Quaternary International, 38­39, 77­83; and others.
The generalised trend of the eustatic sea-level rise is after Fairbanks, R.G., 1989. A
17000 year glacio-eustatic sea level record: influence of glacial melting rates on the
Younger Dryas events and deep-ocean circulation. Nature, 342, 637­642. The beginning
and end of periods approximate
precipitation conditions and document floods, droughts, storms and lake levels (Jingtai et al., 1991; Zhang, 1991; Fang, 1993). Fang (1993) has identified three major periods of lake expansion ­ BC 500 to 0 AD, AD 650 to 950, and AD 1250 to 1650 in eastern China. The flood record on Huanghe extends back to the second century AD, and major floods on the Changjiang occurred in 1153, 1227, 1520, 1560, 1788, 1796, 1860 and 1870 AD, the last being the largest in the last millennium (cf. Baker et al., 1987). Apart from minor climatic and sea-level fluctuations, anthropogenic alterations in the drainage basins have been the dominant characteristics of the late Holocene in monsoon Asia (Singh, 1971; Kealhofer and Penny, 1998; Jiang and Piperno, 1999; Hori et al., 2001). Consequently, the palaeoclimatic and palaeohydrological signals are hard to decipher because the available signals are interwoven with human impacts.
224
Palaeohydrology: Understanding Global Change
BC AD Dunde ice cap pollen record, Tibet (Liu et al., 1998)
Great Drought
1790-1796 AD
MW (Thompson et al.
LIA
2000)
Offshore records NW Arabian Sea (Von Rad et al., 1999)
Lake expansions in China (Fang, 1993) Flood periods, Gansu Province, China (Jingtai et al., 1991) Speleothem, Nepal (Denniston et al., 2000)
Northwestern Himalaya (Mazari et al., 1996) Palaeoflood Record Cn-Wn India (Kale, 1999)
Himalayan lakes (Kusumgar et al., 1998)
1
3
2
4 5 Nilgiri peat (Sukumar et al., 1993)
0
2
4
6
8
10
12
14
16
18
20
Years Ч 102 AD
Figure 13.3 Palaeohydrology of the past 2,000 years. Key: 1 ­ Major floods on the Changjiang (cf. Baker et al., 1987); 2 ­ shift towards wetter climate; 3 ­ more or less same as present; 4 ­ drier than present; 5 ­ wetter than present. MW ­ Medieval Warm period and LIA ­ Little Ice Age (after Liu, K.B., Yao, Z. and Thompson, L.G., 1998. A pollen record of Holocene climatic changes from Dunde ice cap, Qinghai-Tibetan plateau. Geology, 26, 135­138.) Thin line represents no information. The beginning and end of periods approximate
5 DISCUSSION AND CONCLUSIONS The Asian monsoon system is possibly the major component of the global atmospheric circulation. It is clear from multi-proxy palaeoclimatic records that the strength of the monsoon is linked to glacial­interglacial cycles, the orbital forcing and the El · Q7 Nino Southern Oscillations (ENSO)· events (Duplessy, 1982; Thompson et al., 1989; Whetton et al., 1990; Sirocko et al., 1993; Schultz et al., 1998; Yang et al., 2000; Thompson et al., 2000). It has also been postulated that the summer monsoon initiates, amplifies and terminates climatic cycles in the northern hemisphere, by influencing the conditions that produce the greenhouse effect by injecting a large amount of water vapour into the atmosphere and by affecting snow accumulation rates (Kudrass et al., 2001). Understanding the linkage between monsoons and global climates is, therefore, an important aspect in reconstructing a framework for Quaternary climatic and hydrologic changes on the global scale. The physical, hydrological and biological environment of monsoon Asia is closely linked to the rhythm of the two monsoons. Changes in the intensity of monsoons connected with global cooling and warming during the Quaternary have strikingly affected the hydrological characteristics of the rivers of monsoon Asia. Although the poor resolution of terrestrial records precludes any firm derivation of hydrological parameters, virtually all the palaeoclimatic proxies suggest changes in temperature,
Late Pleistocene­Holocene Palaeohydrology of Monsoon Asia
225
precipitation, wind pattern, vegetation and sea level, which in turn caused large amplitude changes in the runoff, discharge, and sediment load across monsoon Asia. The LGM (21.5 cal kyr BP) was characterised by drier conditions with erratic and reduced fluvial activity. The summer monsoon intensified during the early Holocene (ca 10­5 14C kyr BP or 11.5­5.0 cal kyr BP), when maximum monsoon precipitation and an associated increase in fluvial activity appear to have occurred. The glacial­interglacial transition was abrupt and this had a catastrophic impact on the Asian rivers. About 5 to 3 14C kyr BP, a reduction in monsoon precipitation to a minimum occurred. Except in coastal areas and lower reaches of the rivers where sea-level changes intervened, cool, dry phases were associated with aggradation and loess deposition, and the warm, wet phases with large-scale erosion and enormous amounts of sediment transport. The response of fluvial systems to late-Quaternary climatic changes is broadly known but the reasons for lags and leads in the fluvial response, lake levels, sedimentation and soil formation are less well understood. The beginning and the end of increased or decreased monsoon episodes, recognised on the basis of ice- and deep-sea cores, do not exactly match those derived from evidence of increased or decreased precipitation on land (Denninston et al., 2000; Sarkar et al., 2000). It is intriguing to note that while the 3.5 14C kyr BP event is recognisable from both terrestrial and marine records, the YD event is less well represented in the land records. Further, although the monsoon intensity exerts a first-order control on erosional rates in the Himalaya and other highlands, there is little doubt that neotectonic activity has also played an equally important role throughout the Quaternary. However, there seems to be little information that specifically relates to the contribution of neotectonic activity. One of the important themes in future studies should be separating changes caused by tectonic activity from changes that would have occurred primarily because of climatic variations. Therefore, developing chronologies with improved time-resolution is critical to achieve a better understanding of the spatio-temporal differences in the character of the monsoons, palaeohydrological changes and the role of neotectonic activity in one of the most densely populated regions of the world.
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Please clarify the following queries: Q1 The notation "C14 kyr" has been changed to "14C kyr" to maintain consistency in the book. Please confirm if this is correct. Q2 Please clarify if `Denninston' is fine or should it read as `Denniston' as occuring in the references. Q3 Please clarify if Zhou et al., 1991 should be Zhous et al., 1991; because the latter has been provided in the reference list. Q4 Please clarify if this should be Yang et al., 2000a and 2000b. Q5 Please clarify if `Weijian' is fine or should it read as `Weijan' as occuring in Table 13.1 Q6 Figure 13.2 has not been cited in text. Please provide the place of citation. Q7 Please confirm that this is the correct expansion of "ENSO". Q8 We have expanded the initials to the full name. Please clarify if it can be retained as such. Q9 Please provide the page range for this reference. Q10 This author's name has not been cited in the text. Please clarify. Q11 This author's name has not been cited in the text. Please clarify. Q12 Please clarify if `Magnetogstratigraphic' is fine. Q13 Please clarify if `Desrairies' is fine. Q14 This author's name has not been cited in the text. Please clarify. Q15 This author's name has not been cited in the text. Please clarify. Q16 This author's name with the year 1987 has not been cited in the text. Please clarify. Q17 Please clarify if `Erlenkeuser' is fine.

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