Stable isotopic variations in west China: A consideration of moisture sources

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1 Click Here for Full Article JOURNAL OF GEOPHYSICAL RESEARCH, VOL. 112,, doi: /2006jd007718, 2007 Stable isotopic variations in west China: A consideration of moisture sources Lide Tian, 1,2,3 Tandong Yao, 1 K. MacClune, 2 J. W. C. White, 2 A. Schilla, 2 B. Vaughn, 2 R. Vachon, 2 and K. Ichiyanagi 4 Received 30 June 2006; revised 1 December 2006; accepted 19 December 2006; published 23 May [1] In this study, individual precipitation samples, collected over 2 years at stations in different climatic regions of west China (Tibetan Plateau region, Tianshan region, and Altay) were analyzed for the stable isotopes of precipitation to improve our understanding of how vapor transport impacts the modern stable isotopic distribution. Our results identify regional patterns in both d 18 O and deuterium excess (D excess, defined as dd 8d 18 O), and in particular we have identified the northward maximum extent of the southwest monsoon over the Tibetan Plateau. This demarcation is also the boundary for the fractionation effect of temperature on stable isotopes in precipitation. The patterns we have identified are as follows: (1) In the southern Tibetan Plateau, along the southern slope of the Himalayas, our results show a distinct seasonality for both d 18 O and D excess as a result of the shift of summer monsoon moisture and winter westerly moisture transport. The signals of d 18 O in the western Tibetan Plateau reveal that the region receives southwest monsoonal moisture. In the east of the plateau, stable isotopic variation shows alternation between monsoon intrusion and recycling of northern moisture. (2) In contrast, in Tianshan there is an apparent temperature effect in d 18 O, with enriched values occurring in summer and depleted values occurring in winter. Seasonal D excess values, opposite to those observed in the southern Tibetan Plateau, are controlled by differing seasonal evaporation conditions. (3) In Altay, the most northern mountain region, the seasonal d 18 O shows the same variation with that in Tianshan region. However, D excess shows no apparent seasonal variation. Citation: Tian, L., T. Yao, K. MacClune, J. W. C. White, A. Schilla, B. Vaughn, R. Vachon, and K. Ichiyanagi (2007), Stable isotopic variations in west China: A consideration of moisture sources, J. Geophys. Res., 112,, doi: /2006jd Introduction 1 Laboratory of Land Surface Processes Monitoring, Institute of Tibetan Plateau Research, Beijing, China. 2 Stable Isotope Laboratory, Institute of Arctic and Alpine Research, University of Colorado, Boulder, Colorado, USA. 3 Key Laboratory of Cryosphere and Environment, Cold and Arid Regions Environment and Engineering Research Institute, Chinese Academy of Science, Lanzhou, China. 4 Institute of Observational Research for Global Change, Yokosuka, Japan. Copyright 2007 by the American Geophysical Union /07/2006JD007718$09.00 [2] The seasonal and spatial patterns of meteoric stable isotopes ( 18 O and D) are well defined on the global scale as a result of work conducted by the IAEA/WMO Global Network of Isotopes in Precipitation (GNIP) and our understanding of the mechanisms that control stable isotopic fractionation in natural processes has shown great progress [Dansgaard, 1964; Majoube, 1971; Merlivat and Jouzel, 1979; Jouzel and Merlivat, 1984; Jouzel et al., 1997]. Both dd and d 18 O in precipitation have spatial and temporal variation that are dependent on conditions at vapor source, the degree of moisture recycling during water vapor transport, and the temperature at the point of final condensation. Deuterium excess (D excess, defined as dd 8d 18 O) [Craig, 1961] can provide complementary information to d 18 OordD. Though D excess has a global average value of 10, it varies spatially and temporally [Jouzel et al., 1997]. It is mainly controlled by the temperature and relative humidity of the air mass over the moisture source surface, although wind speed at the source also has a role [Merlivat and Jouzel, 1979]. At extremely low temperatures such as over the central Antarctic plateau, the supersaturation of vapor during snow formation can also influence D excess [Jouzel and Merlivat, 1984; Petit et al., 1991]. Reevaporation of raindrops below the cloud base in arid areas, especially in the case of light rain or virga, can give rise to low D excess values in subsequent precipitation [Gat and Tzur, 1967; Stewart, 1975]. Thus the seasonal variations of D excess can provide information about moisture sources and the hydrological cycle, and/or local climatic conditions. [3] The close relationship between isotopic depletion and air temperature [Rozanski et al., 1993; White et al., 1997] makes it possible to quantify climate well beyond recorded history. The use of stable isotopes of precipitation to interpret paleoclimate has been widely used in ice cores from polar regions [Lorius et al., 1985; Dansgaard et al., 1982, 1993; Greenland Ice-Core Project Members, 1993]; 1of12

2 however, similarly useful long-timescale ice cores have been extracted from midlatitude to low-latitude regions as well [Thompson et al., 1990, 1997, 2000]. In the past two decades many of these lower-latitude ice cores have been extracted from the extensive area of the Tibetan Plateau. In recent years, ice cores have also been drilled in other glaciers of Central Asia such as in Mongolia [Schotterer et al., 1997], the Altay Mountains [Henderson et al., 2006], and Pamirs plateau [Kreutz et al., 2003; Tian et al., 2006]. These records provide long-term climatic and environmental information on local and global scales [Thompson et al., 1990, 1997, 2000; Yao et al., 1997a, 1997b]. [4] Thompson et al. [2003] compared three ice core d 18 O profiles from different regions of the Tibetan Plateau and found large differences on both decadal and century scales. These disparities have been attributed to the diverse regional settings which receive moisture from varying moisture sources and experience different postdepositional processes. The Tibetan Plateau ice core programs highlight the need to understand the factors that drive the different behaviors of stable isotopes with modern precipitation studies. [5] Although sampling of meteoric stable isotopes in west China is relatively sparse compared to the vast area, the different climates, and the complicated topography, a few studies have been conducted. Araguas-Araguas et al. [1998] discussed the association of stable isotope patterns over Southeast Asia with different air mass trajectories. Similarly, Aizen et al. [1996] found that the isotopic components in precipitation of three different regions of central Asia (southeastern Tibet, northern Himalayas, and central Tianshan) are related to differing moisture contributions from either easterly or southerly (southwest monsoon) air mass trajectories. Tian et al. [2001] found that the spatial distributions of both d 18 O and D excess along a south-tonorth section in the eastern Tibetan Plateau revealed three distinct climatic-controlled regions, and showed that the southwest monsoon can reach the Tanggula Mountains. These studies illustrate the potential of using isotopic studies to explore the influence of different moisture sources on the stable isotope patterns in west China and ultimately to reconstruct the southwest monsoon s intrusion onto the Tibetan Plateau. In this paper, we use new d 18 O and D excess data to constrain the influence of moisture source and local effects on precipitation signals in west China. 2. Precipitation Sampling [6] There are three main mountain regions in west China (Figure 1). To the south, the Tibetan Plateau is the largest and highest plateau in the world. Large and long mountain ranges stretch in a general zonal configuration along this plateau. To the north are the Tianshan Mountains, the long mountain ranges that help form the Pamiris, and most northerly lie the Altay mountain ranges. [7] Within these three mountain regions there are 46,298 glaciers with a total area of 59,406 km 2 [Yao et al., 2004]. These high mountain ranges also impact moisture transport patterns to the interior of the plateau. In summer, the enormous southwest monsoon brings moisture from the Indian Ocean from the south and provides most of the moisture that precipitates in the Himalayas and on the southern Tibetan Plateau. For the northern part of the research area westerlies provide the moisture for summer precipitation [Numaguti, 1999]. In winter, the westerlies shift to the south, and bring little moisture to most of the Tibetan Plateau. Since each moisture source region brings with it a different isotopic signature, understanding how moisture transport affects modern precipitation isotopes is important for interpreting ice cores from different parts of west China. [8] In this study, we investigate seasonal isotopic compositions and D excess values for 7 precipitation collection sites in the three mountain regions. Figure 1 presents the locations of the precipitation sampling sites together with the GNIP sites that are located within the research area. On the Tibetan Plateau are Nyalam and Lhasa in the south, Shiquanhe and Gaize in the west, and Yushu to the east of the plateau. In the Tianshan Mountain region we use the data from Urumqi and Avalanche stations. In the Altay mountain region, we use the results from the Altay meteorological station. [9] Isotopic analyses were conducted on precipitation collected for over 2 years with the exception of Avalanche station, where the samples were collected only in Following each precipitation event, samples were collected at local meteorological or other permanent stations, and saved as one sample. Liquid samples were collected immediately after the end of rainfall and sealed in a plastic bottle; solid precipitation samples were collected and melted thoroughly at room temperature in a plastic bag before being sealed in a plastic bottle. Meteorological data (temperature and precipitation amount) were also recorded during the precipitation events. [10] Table 1 lists the number of precipitation samples from each station. The numbers of precipitation samples are dependent on the amount of local precipitation and the seasonal precipitation distribution. At Nyalam, precipitation occurs primarily during summer (June September) and spring (March May), while at Lhasa and Yushu, the summer period (June September) alone has the bulk of the precipitation. There is far less annual precipitation at Gaize and Shiquanhe, where what does fall occurs mainly in summer. At Altay, there is precipitation throughout the year, but the summer season still has the highest rainfall. Table 2 shows the measured precipitation-weighted monthly variations of stable isotopes of precipitation at different stations. [11] Most precipitation samples were analyzed for both 18 O and 2 H in Stable Isotope Laboratory, Institute of Arctic and Alpine Research, University of Colorado at Boulder. The measured precision is 0.2% for d 18 O and 1.0% for dd. The method for hydrogen isotopic measurement is introduced in detail by Vaughn et al. [1998]. Some samples (at Avalanche station) were measured in the Laboratoire des Sciences du Climat et de l Environnement, France [Tian et al., 2001]. Another part samples were measured for d 18 Oin the Key Laboratory of Cryosphere and Environment, Chinese Academy of Science, with a precision within 0.2%. The precipitation samples collected in 1999 at Nyalam were measured at Ecological Research Center of Kyoto University, Japan, with a precision of 0.1% and 1.5% for d 18 O and dd. Measured results at all labs were calibrated respect to VSMOW standards using the same normalization method. A comparison of samples measured both at Saclay in France and in Lanzhou, China showed the maximum 2of12

3 Figure 1. Map showing the locations of research area and main mountain ranges in west China. Solid circles are our precipitation sampling stations, and open circles are GNIP stations. The open triangle is the location of Dasuopu ice core in Xixiabangma. The thin line shows the general geographical extent of mountains over 3000 m in altitude. difference between the two labs is within 0.2% for individual precipitation samples and within 0.1% for standard deviation. The seasonal variations of stable isotopes at Nyalam are much the same for the 2 year period although measurements were performed at different labs showing that measurement results are consistent from between labs. 3. Seasonal Atmospheric Circulation Patterns in the Research Area [12] Most of the low-elevation areas of west China are arid because of the long distances from coastal water vapor source regions and the blocking of moisture by high mountains around the region. However, high-elevation mountain regions do intercept considerable moisture as a result of elevation and temperature changes. [13] Figure 2 illustrates the long-term mean wind vector field and specific humidity at different pressure levels around west China from the monthly mean NCAR/NCEP reanalysis data. During the summer period (June to September), the southwest monsoon brings moisture from the ocean to the south of the Tibetan Plateau. There are two trajectories of monsoon moisture: one is from the southern Indian Ocean across the Arabian Sea to the Tibetan Plateau, Table 1. Precipitation Sampling Information at Stations in West China Stations Observation Periods Altitude, m Number of Samples Annual Precipitation, m Nyalam Sep 1998 May Lhasa Sep 1998 May Gaize Sep 1999 Sep Shiquanhe Sep 1999 Sep Yushu Nov 2000 Nov Avalanche Mar 1996 Sep Altay Jun 2000 Dec of12

4 Table 2. Precipitation-Weighted Monthly Stable Isotopes of Precipitation at Stations in West China Station D Value Jan Feb Mar Apr May Jun Jul Aug Sep Oct Nov Dec Nyalam d 18 O dd Lhasa d 18 O dd Shiquanhe d 18 O dd Gaize d 18 O dd Yushu d 18 O dd Hetian (GNIP) d 18 O dd Urumqi (GNIP) d 18 O dd Altay d 18 O dd the other is through the Bay of Bengal northward to the inside of the Tibetan Plateau along the Yalongzangbo River valley (see 850 hpa level for the June to September period in Figure 2). In the northern part of west China, the midlatitude westerlies persist, providing moisture for this region from the west and northwest. The northern circulation system and southwest monsoonal system intersect in the middle of the Tibetan Plateau (500 hpa in Figure 2). Figure 2. Long-term mean divergent wind vector and specific humidity during (left) January May and (right) June September in 850 and 500 hpa pressure levels surrounding west China. The scale for the divergent wind vector (m s 1 ) is shown at the bottom right, and the color bar represents the specific humidity (g kg 1 ). The location of the Tibetan Plateau is shown by the 3000 and 4500 m contours. 4of12

5 Figure 3. Seasonal variations in observations of d 18 O, D excess, and precipitation of individual precipitation events over 2 years at Nyalam and Lhasa, south of the Tibetan Plateau. During the winter and spring period from January to May, westerly winds dominate over most of the research area and bring little moisture. [14] The different moisture origins and their seasonal variation strongly affect the stable isotopes in precipitation over west China. In the following sections, the variation in stable isotopes of precipitation in the different geographical regions (Tibetan Plateau, Tianshan mountain region, and Altay mountain region) will be discussed using results from several representative stations together with the GNIP stations. 4. Results 4.1. Tibetan Plateau [15] The summer monsoon provides most of the annual precipitation to this region. The precipitation amount decreases from east to west and from south to north along the moisture trajectory. Here we outline the results from Nyalam and Lhasa in the south, Shiquanhe and Gaize in the west and Yushu in the east of the Tibetan Plateau Nyalam and Lhasa (South of Tibetan Plateau) [16] Lhasa station ( E, N, 3658 m) is located in the Yalongzangbo River valley (the upper stream of the Brahmaputra River). Nyalam ( E, N, 3810 m) is located in the middle of the southern Himalayas, very close to Xixiabangma Peak. The area s summer precipitation is fed by the southwest monsoon. During winter, westerly winds return to the southern portion of the Tibetan Plateau and bring precipitation of a different origin. Previous work at Lhasa has shown that the seasonal distribution of stable oxygen isotopes is characterized by low d 18 O values during the summer monsoon season, and higher values throughout the reminder of the year [Tian et al., 2003]. [17] Here we present the temporal variation in the amount of precipitation, in d 18 O and in D excess of individual precipitation events at the Nyalam station from September 1998 to May The results from Lhasa, for the same period, are given as a comparison (Figure 3). The d 18 O values show distinct seasonality at the two stations as a result of moisture source changes between the summer monsoon season and winter westerlies. The low d 18 O observed in summer is the result of strong monsoonal activity. From June to September, vapor from the Indian Ocean moves through the Arabian Sea and the Bay of Bengal to the southern part of the Tibetan Plateau and is uplifted onto the southern Himalayas. This uplift brings intense precipitation, resulting in heavily depleted 18 Oin subsequent precipitation. One branch of this moisture moves along the Brahmaputra River Valley and reaches the deep interior of the plateau, which affects the d 18 Oof precipitation at Lhasa. The same monsoon process brings similar variations of d 18 O in summer precipitation to Nyalam. [18] Previous studies found extremely high D excess in both ice cores and glacial meltwater from the summits of the southern Himalayas [Tian et al., 2001]. The D excess in the Dasuopu ice core (in Xixiabangma) varies between 10 17% over the past 1000 years [Thompson et al., 2000]. High D excess was also found in the Qomolangma (Mount Everest) snow [Kang et al., 2001]. These values are much higher than those measured from precipitation from the southwest monsoon, which usually varies between 0 10% [Tian et al., 2001]. [19] The D excess values at Lhasa and Nyalam show distinct seasonal trends: low D excess values from June to September, during the monsoon season, and high D excess in the other months (Figure 3). We attribute the low summertime D excess values to the high humidity over the moisture source region. On the basis of meteorological data from the last 50 years, the summer monsoon season (May September) provides most of the annual precipitation at Lhasa (90%), while Nyalam receives only 47% of its annual total from the monsoon. For this reason, the amountweighted annual D excess value calculated for Nyalam is much higher than that calculated for Lhasa. This suggests that the high D excess values found in the Dasuopu ice core and glacial meltwater are partly caused by wintertime precipitation. [20] The seasonal precipitation pattern at Nyalam is different from both the other southern meteorological stations and any inland Tibetan Plateau meteorological station. There is not much winter precipitation to the south of the Himalayas such as in Bombay, Yangoon or New Delhi [Araguas-Araguas et al., 1998], indicating that the winter moisture at Nyalam (and at Dasuopu) is not likely to be of southern origin. Also, there is little precipitation during winter in the interior of the Tibetan Plateau. Analyzing winter wind vectors reveals that the wind on the southern side of the Himalayas is controlled by the prevailing 5of12

6 Table 3. High d Excess of Winter Precipitation (December May) at Kabul and Nyalam Jan Feb Mar Apr May Jun Jul Aug Sep Oct Nov Dec Kabul Nyalam westerlies (Figure 2). In winter, storms that pick up moisture from the Mediterranean bring precipitation to Iran and Afghanistan, and the high mountains intercepting most of the moisture. The westerly moisture enhanced by Mediterranean evaporation probably brings moisture to the southern site of Himalayas during winter. Back trajectories of each precipitation event during the observation period at Nyalam show strong westerly transport as well (figures are not shown here). [21] Relative to the D excess of moisture derived from the Indian Ocean, The D excess of moisture derived from the eastern Mediterranean is higher because of the very dry overlying air mass, the strong temperature contrast between the air and water and the isotopic disequilibrium between the moisture in the atmosphere and that of the water body [Gat and Carmi, 1970; Rindsberger et al., 1983; Gat et al., 2003]. The D excess values of the 50 vapor samples collected over the Mediterranean Sea during the winter of 1995, varied within 9 35 (with an average of 21%) [Gat et al., 2003]. These values are close to the D excess of 10 30% (average of 20%) in winter precipitation at Nyalam. Kabul is located along the same track of the winter westerlies as Nyalam and precipitation falls mainly in winter (December to May) [Araguas-Araguas et al., 1998]. We see that the periods of greatest winter precipitation at Kabul match well with the first rainfall maximum at Nyalam and that d 18 O and D excess show the same patterns in winter for the two sites (Table 3) Shiquanhe and Gaize (West Tibetan Plateau) [22] Shiquanhe ( E, N, 4280 m asl) and Gaize ( E, N, 4420 m asl) are located in the western region of the Tibetan Plateau. It is a semiarid region in part because of the long distance from the coast but primarily because of to the high mountains surrounding this area. The high Himalayas block moisture transport from the Indian Ocean to the south, and the Pamirs in the west block the moisture transport from the Atlantic Ocean. [23] Precipitation in this area occurs mainly in the summer season with a small number of storm events providing most of the annual precipitation. Figure 4 gives the temporal variation of d 18 O, D excess and precipitation amount of each precipitation event at Shiquanhe (August 1999 to September 2002) and Gaize (August 1999 to August 2001). Definitive seasonal trends in d 18 O and D excess values are not apparent for either location, likely due to the scarcity of winter precipitation. However, there are sharp depletions in the stable isotope values at both sites during summer precipitation events. In the summer of 2000, d 18 O values decreased from 3% to 22% at Gaize and 7% to 26% at Shiquanhe in the middle of July. These depletion events demark the intrusion of the southwest monsoon into western Tibet. The Himalaya Mountains and the distance from the sea ensure that only strong monsoon activity (revealed by heavy rainfall in Nyalam) can bring moisture to the west of the plateau. To verify a southwest monsoon source, we compared the temporal variations of d 18 O during the monsoon period in 2000 at Shiquanhe, and Gaize with that of Nyalam and find that they have similar temporal variations (Figure 5a). Since local temperatures can influence d 18 O, we compared the value of the sampled precipitation against local temperatures. In Figure 5b we present air temperature variations during the collected precipitation events at the three stations, and find little consistency among them. The regression results between air temperature and d 18 O at three stations are quite poor, suggesting air temperature is not the dominant influence on d 18 O. The similar temporal variations of d 18 Oatthe three stations reveals that the conditions that control precipitation along the southern Tibetan Plateau, homogenous moisture source and the similar atmospheric circulation, also control precipitation on the west of the plateau Yushu (East Tibetan Plateau) [24] This study uses observations from one station in eastern Tibet, Yushu ( N, E, 3681 m asl), which is located on the upper tributary of Chang Jiang River. To Figure 4. Seasonal variations in observations of d 18 O, D excess, and precipitation amount of individual precipitation events over 2 years at Shiquanhe and Gaize, western Tibetan Plateau. 6of12

7 Figure 5. (a) Temporal variations of d 18 O between west and south of the Tibetan Plateau, suggesting similar vapor origins during the 2000 summer monsoon season. (b) Regression results between d 18 O and air temperature. The poor correlation argues against temperature as the explanation for consistent temporal d 18 O variation at the three stations. the north is the high Baryan Har Shan Mountains, the highest peak of which is 5267 m. The seasonal air temperature and precipitation patterns are much the same as those at Lhasa. The annual precipitation is 482 mm, of which 93% is in summer (May to October). The mean annual air temperature is 3.0 C. [25] Although there is a distinctive seasonality in air temperature and precipitation, the seasonal variations of stable isotopes are not as obvious as at Lhasa. Figure 6 shows the temporal variation of d 18 O in individual precipitation events at Yushu from January 2000 to October The value for d 18 O is consistently low in winter, but shows strong fluctuation from 20 to 0% in summer. The strong depletion of d 18 O from the beginning of July in both years reflects the influence of the monsoon intrusion, which is accompanied by a decrease in D excess. However, the large fluctuation of stable isotopes throughout these 2 years reflects an apparent alternation between two kinds of moisture sources. One is the southwest monsoon, which results in very low d 18 O as found on the southern part of the plateau. The other is moisture from the arid north which has apparently undergone significant local moisture recycling and results in precipitation with much higher d 18 O and D excess [Tian et al., 2001]. From September on, with the quick retreat of southwest monsoon, an increase in D excess due to this recycling of moisture from the interior of central Asia is apparent in the precipitation. The temporal fluctuation of both d 18 O and D excess in precipitation at Yushu indicate that Yushu marks a boundary for the inland extent of the southwest monsoon Tianshan [26] Most of the vast northern area of the west China region is arid, particularly in the large, low-elevation desert basins, due to the long distance from the coast and the rain shadow effect of the surrounding large mountain ranges. Some mountains receive greater levels of precipitation because of their extreme elevation. We have observations in east Tianshan, at the GNIP station: Urumqi ( N, E, 919 m asl). We also have a short time series of observations from Avalanche station which is along the upper stream of the Kunes River, in the west Tianshan Mountains ( N, E, 1775 m). [27] The stable isotopic patterns in the northern Tianshan reveal a distinct seasonal variation of d 18 O: high d 18 Oin summer and low d 18 O in winter (Figure 7). This seasonality is opposite to that in the monsoon region of the southern plateau. The temporal variations of d 18 O are closely correlated with the temporal variations of air temperature during precipitation (Table 4). We found the regression slope (r) between d 18 O and temperature at Avalanche to be 0.81 d 18 O %/ C (r 2 = 0.67). [28] The D excess value at Urumqi is lower than the global average of 10 in summer, but higher (15 20%) in winter. The observations at Avalanche station took place over less than 1 year (March September), but hint at a similar seasonal trend in D excess: low D excess values during summer and high D excess in spring and autumn. A shallow ice core from the Tianshan Mountains of Kyrgyzstan, drilled at 5100 m, also displays a distinct D excess seasonality [Kreutz et al., 2003]. This seasonality of D excess is consistent with seasonal variations in moisture source evaporation conditions. Moisture for most of north- Figure 6. Seasonal variations in observations of d 18 O, D excess, and precipitation amount in individual precipitation process observed at Yushu, east of Tibetan Plateau. 7of12

8 Figure 7. Seasonal variations of d 18 O and D excess at Urumqi and Avalanche stations, Tianshan Mountains, plus precipitation amount at Urumqi. ern Eurasia (including our sites at Urumqi and Avalanche stations) is supplied from the northern Atlantic Ocean [Numaguti, 1999]. The high relative humidity over the Atlantic Ocean in summer gives rise to low D excess in the subsequent precipitation. However, in winter, the large deficit of moisture over the warm ocean surface leads to a high D excess in the resulting precipitation. This mechanism is also responsible for the D excess seasonality in west Europe [Rozanski et al., 1993]. Thus precipitation in Tianshan region is characterized by low D excess in summer and high D excess in winter. [29] Though having the same seasonality as that at Urumqi, the D excess from the Inilchek glacier, in Tianshan is 23 on average [Kreutz et al., 2003], which is significantly higher than that from the surrounding area of about 10 [Aizen et al., 2005]. Kreutz et al. [2003] postulated that this significantly higher D excess is related to the source water for the precipitation being derived from the Caspian/Aral Seas with subsequent influence by reevaporation from the surrounding desert regions. Yet very high D excess values in glacier ice (over 18) have also been found in summer precipitation at 7000 m of the Xixiabangma, in the central Himalayas [Tian et al., 2001]. This D excess value is much higher than that of the surrounding low-latitude summer precipitation found at Nyalam and Lhasa, though they receive the same monsoon moisture. [30] Except for the seasonal change of moisture sources, the vertical moisture stratification and the existence of mixing layer in the middle mountain region of high massif, can also affect the altitude variation of isotope, and there is a possibility to change the seasonal pattern of precipitation isotope as well. Quite a few studies on stable isotope in snow pits show discontinuity in the variation of d 18 O with altitude in high mountains as Mount Logan, Canada [Holdsworth et al., 1991; Goto-Azuma et al., 2006] and also in Cerro Aconcagua, Argentina [Holdsworth and Krouse, 2002]. This is explained by the existence of the mixing layer in the middle regions, which resulted in an iso-d layer in the middle region. This kind of work is still scarce in the Himalayas glaciers for the strong melting of the lower part of glaciers in summer. However, the strong uplift of southwest monsoon in the southern slope Himalayas probably prevent the preserving of vertical moisture stratification in this regions. In the Tianshan region, an investigation of the d 18 O altitudinal variation from 6200 m to 7500 m (free from melting in this elevation) of the snow pits on the Muztagata glacier, revealed a subsequent decrease of d 18 O with a gradient of 0.4%/100 m and no iso-d layer [Li et al., 2006]. Further lower regions are suffered from intensified summer snow melting. [31] After taking into consideration the different moisture sources, the fractionation in snow formation at extremely cold temperatures should also be considered when trying to understand the higher D excess from high-elevation ice cores [Jouzel and Merlivat, 1984; Petit et al., 1991] Altay [32] Altay ( N, E 735 m) station is located at the southwestern edge of the Altay Mountains. Moisture advects up the Ertis River Valley to reach this area. The annual precipitation at Altay station is about 194mm which is fairly evenly distributed throughout the year (summer months from May to October account for about 55% of annual precipitation) (Figure 8). The high mountains to the northeast receive from 350 to 600 mm of annual precipitation [Liu and You, 1982]. Table 4. Regression Between d 18 O and Air Temperature During Precipitation Events at All the Stations Stations Observation Periods Regression Result Coefficient Squared Nyalam Sep 1998 May 2001 d 18 O= 0.50T 6.69 (n = 285) Lhasa Sep 1998 May 2001 d 18 O= 0.27T (n = 153) Gaize Sep 1999 Sep 2001 d 18 O = 0.17T (n = 147) Shiquanhe Sep 1999 Sep 2002 d 18 O = 0.39 T (n = 65) Yushu Nov 2000 Nov 2002 d 18 O = 0.13 T (n = 339) Avalanche Mar 1996 Sep 1996 d 18 O = 0.81 T (n = 59) Altay Jun 2000 Dec 2002 d 18 O = 0.61 T (n = 233) of12

9 Figure 8. Multiyear average of seasonal pattern of precipitation and air temperature at Altay. The substantial winter precipitation is different from that in Tianshan region, where annual precipitation is mainly from summer. [33] The seasonal d 18 O follows the seasonal air temperature at Altay, which is the same as that at Tianshan (Figure 9). The D excess shows no obvious seasonality throughout the year, with values varying mainly between 0 20%, with a weighted average of 10 (Figure 9). Some extremely low D excess events seen in the data may be due to reevaporation of raindrops during light rainfall. The absence of seasonality in D excess at Altay is probably the result of a different moisture transport mechanism compared to that in the Tianshan Mountain region. During summer, westerlies transport moisture from the Atlantic Ocean. Liu and You [1982] have shown that winter moisture from the polar regions contributes a large portion (about 45 50%) of the annual precipitation. A polar influence is our leading hypothesis for the difference in seasonal D excess compared to that found at Tianshan, which receives moisture throughout the year exclusively from the Atlantic Ocean via the westerlies. [34] A nearby ice core taken from Belukha in the Siberian Altai, also reveals a weak seasonality in D excess [Aizen et al., 2005]. In this ice core, D excess from February precipitation is slightly higher than the annual average, and there is, in fact, less precipitation in the winter half year. The slightly higher D excess found in winter precipitation in Belukha ice is suspected to be due to inland sea evaporation [Aizen et al., 2005]. However, the lack of an apparent higher value in winter D excess at Altay excludes the possibility of inland sea evaporation as a water vapor source for precipitation in this area. The lack of winter and spring precipitation at Belukha indicates that the moisture source between Altay meteorological station and the Belukha ice core likely is different. 5. Conclusion and Discussion [35] In this paper, we present new results from our network of stable isotopes in precipitation together with GNIP data, aiming to an improved understanding of the effect of moisture transport on stable isotopes in west China. The new results have allowed us to develop a general pattern of how different moisture origins impact the spatial and seasonal behavior in isotopes in west China as shown in Figure 10. [36] In summer, the low D excess and low d 18 O of precipitation on the southern Tibetan Plateau (Nyalam, Lhasa, Gaize, Shiquanhe and Yushu) is influenced by the monsoon. While in the northern part of west China (Urumqi, Avalanche station), air masses from the Atlantic Ocean bring moisture to the Tianshan Mountains [Numaguti, 1999] with high d 18 O and low D excess of precipitation. This spatial contrast between the southern Tibetan Plateau and northern west China allow us to delineate the maximum northward extent of the southwest monsoon (Figure 10, top). [37] By winter, the monsoon dissipates and the westerly air currents resume control vast of west China (Figure 10, bottom). Although the low-altitude regions receives little moisture because of the blocking effect of these high mountains to their west, high elevations intercept substantial moisture from Atlantic Ocean via the westerlies. The d 18 O of winter precipitation has a low value because of the low air temperatures, while D excess is higher. The exception is in middle Himalayas, where both d 18 O and D excess of precipitation are higher than in the summer monsoon season. There is evidence that the farthest northern reaches of west China might be affected by moisture from polar air masses (such as at Altay). [38] Previous studies of the northern maximum extent of the summer monsoon have conflicted with one study setting this boundary along the Yalongzangbo River in the interior of the Tibetan Plateau [Araguas-Araguas et al., 1998] and another excluding almost the whole Tibetan Plateau [Johnson and Ingram, 2004]. Our isotope of precipitation data which are derived from a larger number of sites on the Tibetan Plateau allow for a more detailed exploration of the northern boundary of summer monsoon rains over the Tibetan Plateau. This study shows that the northern limit of the summer monsoon is north of the Yalongzangbo river into the middle of the Tibetan Plateau around N. [39] The isotopes in precipitation from the area to the north of the dividing line are dominated by the temperature effect, whereas the precipitation isotopes from the area to the south of this line are more strongly influenced by Figure 9. Seasonal variations in observations of d 18 O, D excess, and precipitation amount in individual precipitation events at Altay. 9of12

10 Figure 10. General patterns for winter and summer moisture transport which affect the seasonal isotopic patterns in west China. In summer, we draw a dividing line to separate the regions influenced by different air masses. This dividing line is located around N in the middle of the plateau. This information is vital for interpreting climate reconstructions from ice cores and lake sediments from the area, which contain information on the long-term variations in the position of this boundary, and because temporal changes in the north-south position of this boundary reflect changes in monsoon strength. 10 of 12

11 precipitation amount. In Table 4, we list the regression results of the d 18 O value of individual precipitation samples relative to temperature at the various stations. It should be noted that these T-d 18 O relationships are based on individual precipitation events over 2 years, which largely reflect the seasonal relationship between d 18 O and air temperature. This seasonal relationship may differ from that of the longterm interannual and century-scale temperature relationship, particularly in the monsoon dominated regions. [40] In addition we find that the D excess of precipitation at the low-elevation meteorological stations is much lower than that found in ice cores from nearby high-elevation sites. This is true between the Xixiabangma glacier and Nyalam station in the central Himalayas, and between the Inilchek glacier and Urumqi station in Tianshan. The large difference of D excess in precipitation between the cold, high-elevation glaciers and low-altitude meteorological stations may imply different moisture sources at different altitude levels and/or further fractionation during snow formation. [41] Despite these and other advances in our understanding of the stable isotopes of precipitation in this region, much work remains to be done, especially in northern west China. In the extensive arid area of northern west China, moisture recycling is intensive, and is not properly understood. This recycling of moisture is probably a strong influence on seasonal isotope patterns. Further work should include simultaneous intensive observations at more stations in the northern west China. [42] Acknowledgments. This work is supported by the National Natural Science Foundation of China (grant ), National Basic Research program of China (grant No.2005CB422002), National Natural Science Foundation of China (grant and grant ). and the Innovation Program of Chinese Academy of Sciences (KZCX3-SW-339). References Aizen, V., H. Aizen, J. Melack, and T. Martma (1996), Isotopic measurements of precipitation on central Asian glaciers (southeastern Tibet, northern Himalayas, central Tian Shan), J. Geophys. Res., 101, Aizen, V., E. Aizen, K. Fujita, S. Nikitin, K. Kreutz, and T. Nozomu (2005), Stable-isotope time series and precipitation origin from firn-core and snow samples, Altai glaciers, Siberia, J. Glaciol., 51, Araguas-Araguas, L., K. Froehlich, and K. Rozanski (1998), Stable isotope composition of precipitation over Southeast Asia, J. Geophys. Res., 103, 28,721 28,742. Craig, H. (1961), Isotopic variations in meteoric waters, Science, 133, Dansgaard, W. (1964), Stable isotope in precipitation, Tellus, 14(4), Dansgaard, W., H. B. Clausen, N. Gundestrup, C. U. Hammer, S. F. Johnsen, P. M. Kristinsdottir, and N. Reeh (1982), A new Greenland deep ice core, Science, 218, Dansgaard, W., et al. (1993), Evidence for general instability of past climate from a 250-kyr ice-core record, Nature, 364, Gat, J. R., and I. Carmi (1970), Evolution of the isotopic composition of atmospheric waters in the Mediterranean Sea area, J. Geophys. Res., 75, Gat, J. R., and Y. Tzur (1967), Modification of the isotopic composition of rainwater by processes which occur before groundwater recharge, in Isotopes in Hydrology, Proceedings of the Symposium in Vienna, 1966, pp , Int. At. Energy Agency, Vienna. Gat, J. R., B. Klein, Y. Kushnir, W. Roether, H. Wernli, R. Yam, and A. Shemesh (2003), Isotope composition of air moisture over the Mediterranean Sea: An index of the air-sea interaction pattern, Tellus, Ser. B, 55, Goto-Azuma, K., R. Koerner, M. Denuth, and O. Watanabe (2006), Seasonal and spatial variations of snow chemistry on Mount Logan, Yukon, Canada, Ann. Glaciol., 43, Greenland Ice-Core Project Members (1993), Climate instability during the last interglacial period recorded in the GRIP ice core, Nature, 364, Henderson, K., L. Andreas, G. Heinz, O. Susanne, P. Tatyana, and S. Margit (2006), Temporal variations of accumulation and temperature during the past two centuries from Belukha ice core, Siberian Altai, J. Geophys. Res., 111, D03104, doi: /2005jd Holdsworth, G., and H. R. Krouse (2002), Altitudinal variation of the stable isotopes of snow in regions of high relief, J. Glaciol., 48, Holdsworth, G., S. Fogarasi, and H. R. Krouse (1991), Variation of the stable isotopes of water with altitude in the St. Elias Mountains of Canada, J. Geophys. Res., 96, Johnson, K., and B. L. Ingram (2004), Spatial and temporal variability in the stable isotope systematics of modern precipitation in China: Implications for paleoclimate reconstructions, Earth Planet. Sci. Lett., 220, Jouzel, J., and L. Merlivat (1984), Deuterium and oxygen-18 in precipitation: Modeling of the isotope effects during snow formation, J. Geophys. Res., 89, 11,749 11,757. Jouzel, J., K. Froehlich, and U. Schotterer (1997), Deuterium and oxygen- 18 in present-day precipitation: Data and modeling, Hydrol. Sci., 42(5), Kang, S., K. Kreutz, P. Mayewski, D. Qin, and T. Yao (2001), Stable isotopic composition of precipitation over the northern slope of the central Himalaya, J. Glaciol., 48, Kreutz, K. J., C. P. Wake, V. B. Aizen, L. D. Cecil, and H. Synal (2003), Seasonal deuterium excess in a Tien Shan ice core: Influence of moisture transport and recycling in Central Asia, Geophys. Res. Lett., 30(18), 1922, doi: /2003gl Li, Z., T. Yao, L. Tian, B. Xu, and G. Wu (2006), Variations of d 18 Oin precipitation from the Muztagata Glacier, East Pamirs, Sci. China, Ser. D, 49(1), Liu, C., and G. You (1982), An introduction of the glacier inventory of Altay Mountains, in Glacier Inventory of China II Altay Mountains (in Chinese), edited by L. Chaohai, pp. 1 18, Sci. Press, Beijing. Lorius, C., J. Jouzel, C. Ritz, L. Merlivat, N. I. Barkov, Y. S. Korotkevich, and V. M. Kotlyakov (1985), A 150,000-year climatic record from Antarctic ice, Nature, 316, Majoube, M. (1971), Fractionnement en oxygène 18 et en deutérium entre l eau et sa vapeur, J. Chem. Phys., 197, Merlivat, L., and J. Jouzel (1979), Global climate interpretation of the deuterium-oxygen 18 relationship for precipitation, J. Geophys. Res., 84, Numaguti, A. (1999), Origin and recycling processes of precipitating water over the Eurasian continent: Experiments using an atmospheric general circulation model, J. Geophys. Res., 104, Petit, J. R., J. W. C. White, N. W. Young, J. Jouzel, and Y. S. Korotkevich (1991), Deuterium excess in recent Antarctic snow, J. Geophys. Res., 96, Rindsberger, M., M. Magaritz, I. Carmi, and D. Gilad (1983), The relation between air mass trajectories and the water isotope composition of rain in the Mediterranean Sea area, J. Geophys. Res., 88, Rozanski, K., L. Araguas-Araguas, and R. Gonfiantini (1993), Isotope patterns in modern global precipitation, in Climate Change in Continental Isotope Records, Geophys. Monogr. Ser., vol. 78, edited by P. K. Swart et al., pp. 1 36, AGU, Washington, D. C. Schotterer, U., K. Frohlich, H. W. Gaggeler, S. Sandjordj, and W. Stichler (1997), Isotope records from Mongolian and alpine ice cores as climate indicators, Clim. Change, 36, Stewart, M. K. (1975), Stable isotope fractionation due to evaporation and isotopic exchange of falling water drops: Application of atmospheric processes and evaporation of lakes, J. Geophys. Res., 80, Thompson, L. G., E. Mosley-Thompson, M. E. Davis, J. F. Bolzan, J. Dai, T. Yao, N. Gundestrup, X. Wu, L. H. Klein, and Z. Xie (1990), Holocene- Late Pleistocene climatic ice core records from Qinghai-Tibetan Plateau, Science, 246, Thompson, L. G., T. Yao, M. E. Davis, K. A. Henderson, E. Mosley- Thompson, P.-N. Lin, J. Beer, H.-A. Synal, J. Cole-Dai, and J. F. Bolzan (1997), Tropical climate instability: The last glacial cycle from a Qinghai- Tibetan ice core, Science, 276, Thompson, L. G., T. Yao, E. Mosley-Thompson, M. E. Davis, K. A. Henderson, and P.-N. Lin (2000), A high-resolution millennial record of the South Asian monsoon from Himalayan ice cores, Science, 289, Thompson, L. G., E. Mosley-Thompson, M. E. Davis, P.-N. Lin, K. Henderson, and T. A. Mashiotta (2003), Tropical glacier and ice core evidence of climate change on annual to millennial time scales, Clim. Change, 59, Tian, L., V. Masson-Delmotte, M. Stievenard, T. Yao, and J. Jouzel (2001), Tibetan Plateau summer monsoon northward extent revealed by measurements of water stable isotopes, J. Geophys. Res., 106, 28,081 28, of 12

12 Tian, L., T. Yao, P. F. Schuster, J. W. C. White, K. Ichiyanagi, E. Pendall, J. Pu, and W. Yu (2003), Oxygen-18 concentrations in recent precipitation and ice cores on the Tibetan Plateau, J. Geophys. Res., 108(D9), 4293, doi: /2002jd Tian, L., T. Yao, Z. Li, K. MacClune, G. Wu, B. Xu, Y. Li, A. Lu, and Y. Shen (2006), The recent rapid warming trend revealed from the isotopic record in Muztagata ice core, eastern Pamirs, J. Geophys. Res., 111, D13103, doi: /2005jd Vaughn, B. H., J. W. C. White, M. Delmotte, M. Trolier, O. Cattani, and M. Stievenard (1998), An automated system for hydrogen isotope analysis of water, Chem. Geol., 152(3), White, J. W. C., L. K. Barlow, D. Fisher, P. M. Grootes, J. Jouzel, S. J. Johnsen, M. Stuiver, and H. B. Clausen (1997), The climate signal in the stable isotopes of snow from Summit, Greenland: Results of comparisons with modern climate observations, J. Geophys. Res., 102, 26,425 26,439. Yao, T., L. G. Thompson, Y. Shi, D. Qin, and K. Jiao (1997a), Climatic changes since the Last Interglacial recorded in the Guliya ice core, Sci. China, Ser. D, 40, Yao, T., Y. Shi, and L. G. Thompson (1997b), High resolution record of paleo climate since the Little Ice Age from the Tibetan ice cores, Quat. Int., 37, Yao, T., S. Liu, J. Pu, Y. Shen, and A. Lu (2004), Glacial retreat high Asia and its influence on water resource in northwest China, Sci. China, Ser. D., 47, K. Ichiyanagi, Institute of Observational Research for Global Change, Yokosuka, Kanagawa , Japan. K. MacClune, A. Schilla, R. Vachon, B. Vaughn, and J. W. C. White, Stable Isotope Laboratory, Institute of Arctic and Alpine Research, University of Colorado, Boulder, CO 80309, USA. L. Tian and T. Yao, Laboratory of Land Surface Processes Monitoring, Institute of Tibetan Plateau Research, Beijing , China. 12 of 12