Groundwater recharge history and hydrogeochemical evolution in the Minqin Basin, North West China

Applied Geochemistry Applied Geochemistry 21 (2006) 2148–2170 www.elsevier.com/locate/apgeochem

Groundwater recharge history and hydrogeochemical evolution in the Minqin Basin, North West China W.M. Edmunds a,*, Jinzhu Ma b, W. Aeschbach-Hertig c, R. Kipfer d, D.P.F. Darbyshire e b

a Oxford Centre for Water Research, Oxford University Centre for the Environment, Oxford University, Oxford OX1 3QY, UK CAEP, Key Laboratory of West China’s Environmental System (Ministry of Education), Lanzhou University, Lanzhou 730000, China c Institut fu¨r Umweltphysik, Universita¨t Heidelberg, D-69120 Heidelberg, Germany d Isotope Geology, ETH, CH-8092, Zurich, Switzerland e NERC Isotope Geoscience Laboratory, Keyworth, Nottingham, NG12 5GG, UK

Received 17 November 2005; accepted 31 July 2006 Editorial handling by R. Fuge Available online 30 October 2006

Abstract The Minqin Basin is a type area for examining stress on groundwater resources in the Gobi Desert, and has been investigated here using a combination of isotopic, noble gas and chemical indicators. The basin is composed of clastic sediments of widely differing grain size and during the past half century over 10 000 boreholes have been drilled with a groundwater decline of around 1 m a 1. Modern diffuse recharge is unlikely to exceed 3 mm a 1, as determined using unsaturated zone profiles and Cl mass balance. A small component of modern (<50 a) groundwater is identified in parts of the basin from 3 H–3He data, probably from irrigation returns. A clear distinction is found between modern waters with median d18O values of 6.5 ± 0.5& and most groundwaters in the basin with more depleted isotopic signatures. Radiocarbon values as pmc range from 0.6% to 85% modern, but it is difficult to assign absolute ages to these, although a value of 20% modern C probably represents the late Pleistocene to Holocene transition. The d13C compositions remain near-constant throughout the basin (median value of 8.1& d13C) and indicate that carbonate reactions are unimportant and also that little reaction takes place. There is a smooth decrease in 14C activity accompanied by a parallel increase in 4He accumulations from S–N across the basin, which define the occurrence of a regional flow system. Noble gas temperatures indicate recharge temperatures of about 5.6 C for late Pleistocene samples, which is some 2–3 C cooler than the modern mean annual air temperature and the recharge temperature obtained from several Holocene samples. Groundwaters in the Minqin Basin have salinities generally below 1 g/L and are aerobic, containing low Fe but elevated concentrations of U, Cr and Se (mean values of 27.5, 5.8 and 5.3 lg L 1, respectively). Nitrate is present at baseline concentrations of around 2 mg L 1 but there is little evidence of impact of high NO3 from irrigation returns. Strontium isotope and major ion ratios suggest that silicate reactions predominate in the aquifer. The results have important implications for groundwater management in the Minqin and other water-stressed basins in NW China – a region so far destined for rapid development. The large proportion of the water being used at present is in effect being mined and significant changes are urgently needed in water use strategy.  2006 Elsevier Ltd. All rights reserved.

*

Corresponding author. E-mail address: [email protected] (W.M. Edmunds).

0883-2927/$ - see front matter  2006 Elsevier Ltd. All rights reserved. doi:10.1016/j.apgeochem.2006.07.016

W.M. Edmunds et al. / Applied Geochemistry 21 (2006) 2148–2170

1. Introduction The Minqin Basin has recently been the focus of attention due to increasing stress on its water resources and environmental degradation and the Prime Minister has indicated a national priority to rehabilitate this important and historic area. The basin may be considered as a type area for much of northern China where in modern times an arid or semi-arid belt lies at the northern limit of the area reached by the SE monsoon. To the east this region is fed by the Yellow River, originating in the heart of the Tibetan Plateau and which then passes through the arid region of the Gobi Desert. Water from the river has been used traditionally to alleviate water shortage, although withdrawals have increased significantly and grossly reduced its flow in recent years such that it no longer reaches the coast in most years (Liu and Jun, 2004). Minqin is one of a series of basins in NW China which are fed directly from surface runoff from the Tibetan Plateau. The Shiyang River (Fig. 1) like others along the Hexi Corridor area (roughly the line of the ancient Silk Road) feeds (or rather fed) a terminal lake (Lake Zhuyeze) in the arid Gobi Desert but this lake progressively diminished in size splitting into two smaller lakes around the first century AD due to upstream abstraction. This process continued over the last two millennia with the complete disappearance of the lake by the mid 1950s (Chen et al., 2003). The natural emplacement and withdrawal of groundwater in the terminal lake basin has to be explained against this setting. In a palaeohydrological context it is possible that wetter periods could have contributed to local direct recharge, although at the present day with around 100 mm rainfall this is likely to be low. The Minqin and similar basins may form part of a regional groundwater flow regime with subsurface inflows and outflows towards the Yellow River. Short term climatic fluctuations could have had an impact historically on recharge to the basin aquifer, but for 2 ka human impacts are more likely to have been felt on the hydrological regime as a whole (Ma et al., 2005). For many years there has been significant abstraction of upstream water from the Shiyang River using traditional multi-channel irrigation (Chen and Qu, 1992), which in turn has led to over-exploitation of groundwater to augment supplies. The aquifers of the mountain alluvial fans in

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the vicinity of Wuwei are highly productive, being replenished by the river system and therefore attractive for development at the expense of the downstream areas. This situation is exacerbated by rapid population growth and socio-economic development along the Hexi Corridor. Over the last 50 a, man-made oases, including Minqin, have developed rapidly in the terminal regions of various river basins and irrigation areas have expanded to 3.8 million ha (Wang and Cheng, 1999, 2000). Degradation of the hydrogeological regime is therefore a feature of this region with falling water levels a widespread occurrence (Ma et al., 2000; Feng et al., 2000). Sometimes the reasons for this decline are obvious from water balance studies but the system as a whole is less well understood. Past work on the Minqin Basin area has dealt largely with the chemical properties of the water. However, Shi et al. (2001) reported stable isotope data showing values of 6.6 to 9.7 d18O which they claim to represent modern recharge, modified by evaporation and that the groundwater is sourced from the Shiyang River. These authors also report He isotope data which show some enrichment and that the 3He/4He ratios denote a component of mantle He. Radiocarbon results are also available from the neighbouring Ejina (Heihe) Basin which give Holocene ages for confined groundwater (Wu et al., 2002). Further evidence for Holocene as well as recent recharge in the same basin is also given by Chen et al. (2004). The objectives of the present paper are to consider a wide selection of geochemical indicators to provide an understanding of the evolution of the groundwater in the Minqin Basin in relation to its recharge sources, timing of recharge and recharge history especially since it is likely that much of the water being exploited may have been recharged under different climatic conditions to the present day. Here the authors present the first evidence from central Asia on noble gas palaeo-recharge temperatures and provide the first application of multi-tracer approaches to resolving some of the key hydrogeological questions in the region. The controls on the natural baseline water quality and subsequent impacts of the significant water resource development of the past few decades on the groundwater quality are considered against knowledge of the recharge history. Comprehensive approaches to groundwater understanding combining chemistry and isotopic

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Fig. 1. Location map of the Shiyang River Basin showing its relationship to the Minqin Basin in NW China. The main hydrogeological regions linked to geological structure are shown (I–III). The site of the unsaturated zone drilling profile at Beitu is shown.

indicators are relatively few in China with the exception of (Zongyu et al., 2003) but have been applied elsewhere (Edmunds et al., 2003). Improved understanding of the fragile nature of the resource therefore is fundamental to many important decisions that will need to be made in China (as in other arid regions), in relation to demographic change and the reform of agricultural practices (Ma et al., 2005).

2. Regional geology and hydrogeology The Minqin Basin, fed by the Shiyang River is one of a number of terminal lake basins in Chinese Inner Mongolia with an origin in the Tibetan Plateau. From its source the Shiyang River transects three different geological units: (I) the Qilian Palaeozoic geosynclinal fold zone, (II) the Hexi Corridor depression (Wuwei Basin) with its associated struc-

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tures and (III), the northern Minqin Basin, limited to the north by the Alasha Arc structural zone (Fig. 1). From the end of the Palaeozoic and continuing through the Mesozoic era, coinciding with the uplift of the Qilian geosyncline, the embryonic Hexi Corridor was formed. This was followed by a complex tectonic episode dissecting the Hexi Corridor into two parts: the southern Wuwei Basin and the northern Minqin Basin. From the late Tertiary, especially from the end of the Pliocene and at the beginning of the early Pleistocene, the Tibetan mountains underwent a rapid uplift (Li et al., 1979) and the basins subsided further. At the same time intensive denudation and erosion from the mountains led to significant transfer of clastic material to the basin depressions. This deposition mainly during the Quaternary, of thick diluvial and alluvial sediments, and some aeolian and lacustrine deposits, led to the formation of the main aquifer system (Fig. 2). In the southern part of the Shiyang River basin, the diluvial aquifer is formed of highly permeable cobble and gravel deposits some 200–300 m thick in which the present day water levels range from 50 m to 200 m below surface. The specific capacity is about 3–30 L s 1 m 1, and the coefficient of permeability is up to 100–600 m d 1 (Qu, 1991). This allows a large amount of surface runoff in the piedmont fan to seep down and recharge the aquifer. From the northern edge of this diluvial fan, the aquifer comprising inter-bedded cobble gravel, fine sand and clay becomes confined or semi-confined with piezometric levels at less than 5 m depth. In many places at the onset of confined conditions, the groundwater then overflows as springs and re-emerges as streams.

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Although the Hongya and Alagu mountains form a natural geological boundary between the Wuwei and Minqin Basins, a hydraulic connection exists between them (Fig. 1). However, at the present day the groundwater barely crosses the tectonic zone to recharge the aquifer of Minqin Basin. The Hongya Mts and some small outcrops in the Minqin Basin consist of Palaeozoic schist, gneiss, metamorphic sandstone, and inter-bedded Cretaceous and Tertiary argillaceous rocks and conglomerates. The Pliocene and early Pleistocene sediments in the Basin are inter-bedded, poorly consolidated fluvio-lacustrine sandstones with a thickness averaging 200–300 m (Fig. 2). The overlying Middle and Late Pleistocene sediments are comprised of fluvial sands in the south of the basin but inter-bedded sandyclay and clay extend northward between 80 and 130 m thick. The specific capacity of these ranges from 10 L s 1 m 1 to 1 L s 1 m 1 and the coefficient of permeability is about 70–10 m d 1 northward (Qu, 1991). The region is arid with mean annual precipitation in the Minqin Basin of 100 mm or less with rainfall occurring in the summer months at the northern limit of the SE monsoon. The mean annual temperature is 8.2 C but with an annual range from 40 C to 27 C. Groundwater development was very limited until the middle 20th century. Intensive development and irrigated agriculture using both surface and groundwater was facilitated by the building of the Hongyashan Reservoir in the1950s. In 1965 there were only 165 shallow wells but by 1976 there were 10 000 wells with over 14 000 by the end of the century to a maximum depth of 320 m.

Fig. 2. Schematic hydrogeological cross section of the Minqin Basin.

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Previous studies have identified 3 water quality zones mainly in a south to north direction: (i) a fresh water zone with 0.3–0.48 g/L total mineralisation; (ii) a locally developed brackish zone some 60 m thick overlying the fresh water; (iii) a deeper brackish or saline aquifer some 150–240 m thick with a total mineralisation of 3–17 g/L (Chen et al., 1982; Li, 1991). 3. Methods Samples were taken for isotope and chemical analysis during two field seasons in 1999. Following preliminary interpretation of these data a further sampling campaign was conducted in 2000 during which samples for noble gases and additional water samples were collected; surface water samples from the Hongya Reservoir were also taken as well as groundwaters from the new South Lake development area. Locations of all samples are shown in Fig. 3 where they have been subdivided according to well depth. Most samples were collected from boreholes penetrating the shallow or deep aquifer and these fell into two distinct groups (<100 m and 200–300 m) corresponding to the customary drilling practice. The borehole head works consisted of discharge mains without sample valves with estimated flow rates of 30 L s 1. Samples were taken anaerobically by inserting a sampling tube against the pressure of the water discharge main. All samples for chemical analysis were filtered (0.45 lm membrane filters) and an aliquot acidified with 1% HNO3 for analysis of trace elements. One litre samples for radiocarbon were collected. Samples for noble gases and 3H (about 23 cm3) were collected in crimped Cu tubes. On-site analysis included temperature, specific electrical conductance (SEC), total alkalinity (as HCO3 ) by titration, pH (on most samples) and redox potential (eH). Chloride, NO3–N, Br, F and I were analysed by automated colorimetry. Filtered and acidified samples were analysed for major cations, SO4, and a wide range of trace elements either by ICP-OES (inductively coupled plasma optical emission spectroscopy) or ICP-MS (inductively coupled plasma mass spectrometry). Calibrations for cation analyses were performed using appropriately diluted standards and both laboratory and international reference materials were used as checks for accuracy. Instrumental drift during ICP-MS analysis was corrected using In and Pt internal standards.

Samples for stable isotope analysis (18O, 2H, 13C) were measured by isotope ratio mass spectrometry. An internal check on the quality of the data was made by determining the ionic balance; the balance lay below ±6% except for four samples. Precision of measurement for stable isotope and radioactive analysis was ±0.1& for d18O, d13C and ±2& for d2H. Graphite samples were prepared from HCO3 in the untreated water at the NERC Radiocarbon facility East Kilbride, Scotland and sent for AMS measurement of radiocarbon at the University of Arizona. Strontium was separated from aliquots of unacidified water samples by conventional ion exchange procedures. Isotope ratios were measured on a Finnegan MAT 262 mass spectrometer operated in dynamic mode. The average 87Sr/86Sr ratio obtained for the NBS 987 international isotope standard during the two periods of the study were 0.710251 ± 0.000015 (n = 18) and 0.710258 ± 0.000026 (n = 18) and therefore all measured ratios have been normalised to a value of 0.710248. Replicate analyses (n = 18) of North Atlantic seawater yielded a mean ratio of 0.709170 ± 0.000026 (2r). Noble gases were analysed by mass spectrometry in the noble gas laboratory of ETH Zurich, Switzerland, according to the methods discussed by Beyerle et al. (2000). Measurement precision ranged between 0.5% and 1% for He, Ne and Ar, and between 1% and 2% for Kr and Xe. After gas extraction for the noble gas analysis the water was resealed in the Cu tubes and analysed on 10 samples approximately one year later for regrown 3He to determine the 3H concentration. Due to the long decay time and the use of a highly sensitive compressor ion source, the precision and detection limit for 3H was better than 0.1 TU. To examine the age and climatic relationships further, samples for 3H, He isotopes, and noble gases were collected from 16 sites where radiocarbon and stable isotope data were also available. Helium and noble gas data for 15 wells were obtained, (one sample was lost during analysis). For two wells, results from duplicate samples are available, which reproduced within the experimental errors (see above), except for Xe, where a deviation of about 4% occurred in both cases. A profile of unsaturated sand at Beitu to the SE of the Minqin Basin (Fig. 3) was taken by auguring for moisture analysis (Edmunds et al., 1988) to determine modern recharge rates using

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Fig. 3. Site map of the Minqin Basin showing the sampling locations.

the Cl mass balance technique and to measure isotopic inputs to the shallow aquifer. This work was an extension of other profile studies in the Badain Jaran Desert NW of the area (Ma and Edmunds, 2006). Water was extracted by elutriation of 50 g sand samples with 30 mL demineralised water for Cl analysis and directly by centrifugation for isotope samples.

4. Results and discussion The geochemical results are shown in Table 1 (site information field data, major and minor ions), Table 2 (trace elements, stable isotopes of water and Sr isotope ratios) Table 3 (C isotopes) and Table 4 (3H, and noble gases); the sites from which samples were taken are shown in Fig. 3. The results are first

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eH (mv)

T (C)

11.1 14.3 17.5 17.5 19.6 21.1 23.2

224 220 173 235 210 240 214

47 36 41 40 31 44

23.2 23.6 23.6 24.3 27.5 27.9

45 35 46 34 32 33 30 26 29 27 28 51 21 22 25 50 53 3 24 23 4

28.2 28.9 29.6 32.5 33.9 36.1 36.8 39.3 41.1 41.4 45.0 50.0 52.5 56.8 56.8 58.9 62.5 62.5 62.9 63.6 65.7

Location

Well depth

Map no

Distance from Hongya

Gengmin Bashe Shangxing Shiyishe Yigan Shuiguansuo Ergan Shuiguansuo Puyang Yishe Zhongguan Yishe Desert Part(water Co.4) Chuanxing Liushe Zhonggou Yishe Tianbing Liushe Qinfen Fenchang Zhongping Ershe Xialei Datan Fenchang Dencao Ershe Longyi Sanshe Xiadong Yishe Haozihu Wushe Huang’an Sanshe Beigu Qishe Shangquan Sishe Sanjie Dongsanshe Daxi Qishe Hongdong Ershe Beixing Yishe Fengqin Yishe Quanshan Zhengfu Heping Sishe Gaoleiwang Yishe Liuyu Yishe Mingqi Ershe Jianggui Village Huayin Ershe Huazai Sishe Alttong village

100 100 100 100 96 100 220

43 42 38 48 37 39 49

100 100 100 80 100 80 100 100 100 100 100 100 100 100 100 100 100 100 300 90 80 90 250 26 280 300 220

pH (field)

7.46

7.41 7.11 7.02

7.08 7.87 6.82 7.58

Tot Min (mg L 1)

Cl (mg L 1)

9.1 10.1 13 10.7 9.6 10.5 11.2

674 1085 766 1733 1137 1468 911

51.7 90 63.6 211 120 157 115

190 403 238 687 417 562 324

244 300 262 345 275 336 213

199 217 274 140 191 247

9.8 9.03 10.3 10.8 9.9 10

1348 888 1563 1134 567 1701

151 72.9 185 125 45.9 247

511 289 633 442 130 659

247 250 91 269 1 196 311 333 169 246 325 197 255 331 250 70 143 250 195 279 250

10.1

1080 1738 2188 1828 994 539 883 1056 611 696 676 1129 800 888 2237 1949 1242 1629 668 3388 1023

120 226 270 223 207 49.1 102 137 62.2 84 61.5 134 129 156 325 509 369 271 194 630 208

404 718 935 758 283 124 308 434 214 244 246 283 293 265 938 675 311 537 156 1400 306

9.6 10.3 11.6 11.1 11.2 11.7 11.1 12.2 10.6 12.2 13.4 11.6 10.2 12.8 11.3 14 11.5 13.8

SI calcite

0.09

0.45 0.47 0.07

0.19 0.09 0.44 0.21

SO4 (mg L 1)

HCO3 (mg L 1)

NO3–N (mg L 1)

Br (mg L 1)

F (mg L 1)

Si Fe (mg (mg L 1) L 1)

1.1 1.58 1.28 2.86 1.8 2.19 1.46

0.037 0.045 0.043 0.135 0.057 0.077 0.072

0.23 0.23 0.34 0.25 0.45 0.19 0.21

6.0 7.2 7.3 7.3 7.2 7.3 7.3

<.02 0.02 <.02 0.03 0.02 0.02 0.02

87 53.4 92 52.8 33 97

1.91 1.39 2.31 1.49 0.899 2.19

0.110 0.040 0.125 0.070 0.028 0.122

0.33 0.39 0.16 0.31 0.39 0.31

7.5 7.9 7.5 6.9 6.6 7.1

0.02 <.02 0.03 0.02 <.02 0.03

54.7 108 149 109 12.7 16.5 43 40.6 22.4 24.7 29.7 15.3 38.2 39.3 135 99 59.3 118 26.6 238 36.9

1.5 3.15 3.35 2.37 0.436 0.501 1.28 1.43 0.763 0.835 0.886 0.466 1.47 1.35 2.5 2.78 2.5 3.49 2.08 4.69 1.99

0.070 0.101 0.180 0.129 0.038 0.020 0.057 0.073 0.032 0.046 0.038 0.029 0.031 0.042 0.269 0.075 0.082 0.080 0.145 0.331 0.100

0.24 0.34 0.25 0.30 1.00 0.47 0.29 0.30 0.36 0.35 0.32 0.56 0.46 0.29 0.33 0.21 0.64 1.35 0.41 0.22 0.73

7.2 7.4 7.4 7.1 5.9 6.6 7.5 6.9 7.4 7.1 7.2 7.2 7.0 7.3 7.2 7.8 6.2 7.5 5.1 6.5 5.8

0.02 0.03 0.03 0.03 <.02 <.02 <.02 0.02 <.02 <.02 <.02 <.02 0.03 <.02 0.39 0.03 <.02 0.36 <.02 0.07 0.24

Na (mg L 1)

K (mg L 1)

Ca (mg L 1)

Mg (mg L 1)

Sr (mg L 1)

1.0 2.1 0.8 6.7 2.1 4.2 0.6

63 95.8 70.4 180 147 175 93.7

3.6 4.4 4.6 6.6 7 6.2 5.5

87.3 115 85.5 185 103 145 108

33.9 75 41.3 112 66 83 51.4

298 290 290 226 251 314

2.3 0.5 3.0 0.8 0.3 1.5

171 89.2 201 164 51.9 243

8.6 8.1 7.1 5.1 5.6 7.3

119 84.8 152 118 49.8 132

252 302 367 350 181 211 229 192 169 179 192 378 150 240 382 190 156 420 123 465 202

0.4 4.0 5.0 1.6 0.4 0.2 0.2 0.5 0.9 0.4 0.2 1.6 0.2 0.8 4.0 0.4 0.7 4.0 1.6 6.0 1.16

130 191 275 247 286 103 99.6 148 82.9 102 80 287 135 117 319 301 276 172 120 514 216

5.8 8.2 6.7 9.1 3.9 2.7 4.7 3.6 2.8 2.9 3.2 4.1 2.7 3.3 5.4 5.3 2.8 7.2 1.7 5.3 4.3

113 181 180 130 20.4 32.8 96.6 100 56.7 58.6 63.4 26 52.3 67 129 169 67.4 100 45.1 130 48.4

W.M. Edmunds et al. / Applied Geochemistry 21 (2006) 2148–2170

Table 1 Locations of all groundwater samples from the Minqin Basin referred to in the study and locations of which are shown in Fig. 3

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61 53.4 164 25.2 136 37.5 82 47.8 471 78 123 477 625

3.4 1.41 7.99 1.82 3 3.04 5.41 2.84 8.28 4.73 8.94 5.37 11.96

0.105 0.038 0.181 0.110 0.202 0.060 0.300 0.100 0.290 0.139 0.211 0.391 4.103

0.46 0.51 0.42 0.33 0.37 0.45 0.41 0.70 0.83 0.44 0.36 0.70 0.90

6.6 7.2 5.9 6.1 7.2 6.9 5.9 6.2 7.0 6.6 6.1 9.6 8.1

0.08 0.02 0.05 0.06 0.03 0.03 0.34 0.17 0.03 0.03 0.75 1.12 0.14

discussed below in relation to the present day inputs to the aquifer from rainfall and the river. Then the question of modern direct recharge to the basin is considered using results from unsaturated zone moisture. The origin and age of the groundwater are discussed using mainly stable isotope and noble gas data, and also chemical data including Sr isotope ratios. The natural chemical characteristics of the groundwater are also considered as tracers, as well as their baseline characteristics against which the extent of modern recharge induced from irrigation or urban activities may be established, as well as quality criteria for use.

1.05 0.33 0.02 0.01 0.37 0.28 0.01 0.06 Results for field data, major ions and key trace elements.

6.8 7.48 7.71 7.01 7.34 7.66 6.74 6.68

280 80 300 290 70 270 315 255 60 260 280 65 100 Banhu Sanshe Dongqu Shuiguansuo Chuxian Sishe Xianshe Village Yinke Sishe Shouhao Yishe Wansheng village nr Zhongqu Xisui Ershe Huokan Ershe Shuisheng Qishe Yanghe Bashe Zhengxing Yishe

14 19 11 6 52 15 5 2 18 12 13 17 16

Well Map depth no

66.1 67.5 68.6 68.9 68.9 69.3 70.7 71.8 71.8 73.9 75.7 77.1 77.9

4.38 7.11 7.51 7.48

293 250 287 150 230 290 314 148 145 267 202 98 152

13.6 10.6 11.8 15.2 10.5 9.3 14.8 14.2 10.7 12.4 11.2 11.1 10.6

3.41 0.32 0.01 0.46

924 944 2907 665 4242 887 1371 1218 5448 1438 2512 5952 5598

265 81 860 172 1100 236 421 274 2090 339 420 1060 1670

249 285 992 156 1400 271 439 412 >2500 512 1130 2330 >2500

129 310 172 139 390 116 99.4 157 406 159 163 948 770

0.5 1.5 0.7 0.52 1.9 0.8 1.1 1.12 2.9 0.7 0.2 0.2 0.6

145 127 533 131 1060 143 223 240 2060 257 413 822 2120

2.8 7.5 3.8 1.9 8.6 2.2 3 3.8 46.9 2.5 3.7 28.6 31.2

71.3 78.2 182 39.4 145 80.5 102 82.5 371 89.9 260 287 381

4.1. Recharge to the aquifer

Location

Distance from Hongya

pH eH (field) (mv)

T (C)

SI calcite

Tot Min Cl SO4 (mg (mg L 1) (mg L 1) L 1)

HCO3 NO3–N (mg L 1) (mg L 1)

Na K Ca Mg Sr (mg L 1) (mg L 1) (mg L 1) (mg L 1) (mg L 1)

Br F (mg L 1) (mg L 1)

Si (mg L 1)

Fe (mg L 1)

W.M. Edmunds et al. / Applied Geochemistry 21 (2006) 2148–2170

Five categories of recharge to the Minqin aquifer are possible – modern local rainfall, past rainfall (wetter climates), modern river recharge, past river recharge and lateral groundwater flow. The terminal lake may have had an influence on groundwater composition as a result of discharge. Modern end members can be explored using O and H isotopic compositions and available data are plotted in Fig. 4. Modern rainfall from the region is represented by Zhangye station some 250 km NW of Minqin, but in a representative position in the Hexi Corridor. The range of compositions for Zhangye over 7 a, taken from the IAEA network (http:// isohis.iaea.org) lies between 0 and 24.7& d18O, but for consideration of the most likely composition of surface runoff and groundwater recharge the volume-weighted mean values are considered the most relevant. Weighted mean values are plotted for seven non-consecutive years since 1986 and all except one set of data for 1990 (d18O 4.4& and d2H 25.8&) plot on or close to the meteoric line with an overall mean composition of d18O 6.5& and d2H 44&. Modern values for surface runoff are represented by the Hongya Reservoir which aggregates the annual composition of the flow of the Shiyang River and also for the Tangjia saline lake. The composition of the reservoir outflow is d18O 3.1& and d2H 27.9& and lies below the GMWL (Fig. 4) indicating some evaporation. The isotopic composition of the two open water bodies plot on a line with slope of 4.9, comparable to the theoretical slope of 4.5 for direct evaporation from an open water body under conditions of low humidity (Clark and Fritz, 1997) and produces an intercept near to the mean modern rainfall values for Zhangye.

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Table 2 Results for representative trace elements discussed in the study Li (lg L 1)

Al (lg L 1)

Cr (lg L 1)

Mn (lg L 1)

Co (lg L 1)

Ni (lg L 1)

Cu (lg L 1)

Zn (lg L 1)

As (lg L 1)

Se (lg L 1)

Rb (lg L 1)

Mo (lg L 1)

Cd (lg L 1)

Sb (lg L 1)

Ba (lg L 1)

U (lg L 1)

43 42 38 48 37 39 49 47 36 41 40 31 44 45 35 46 34 32 33 30 26 29 27 28 51 21 22 25 50 53 3 24 23 4 14 19 11 6 52 15 5 2 18 12 13 17 16

14 29 18 45 32 30 20 43 29 36 23 18 39 22 49 63 46 27 15 18 19 12 13 14 27 24 22 51 49 46 70 35 116 52 44 40 106 34 104 47 68 69 327 65 132 274 462

6 7 11 2 2 2 7 10 1 4 4 4 1 5 2 7 3 53 7 8 13 6 2 3 4 14 7 5 1 1 4 4 4 15 16 2 4 7 3 43 9 12 3 8 29 2 5

2.3 2.1 2.4 1.3 2.3 1.7 2.8 2.7 1.8 3 3.8 3.8 1.9 3.2 <.5 3 4.1 16.8 7.6 3.2 7 5.3 7.2 8.2 35.7 14.7 8.6 <.5 18.7 15.3 1.4 4.8 0.6 30.7 20.1 0.8 22.1 25.7 4.8 1.6 38.1 13.8 1.9 19.5 2.9 1 1.2

0.36 0.33 1.16 1.37 0.57 1.4 1.89 78.77 0.14 0.14 0.44 1.44 2.21 2.13 10.39 0.93 0.73 6.44 0.76 0.57 0.49 0.29 0.33 0.92 0.6 1.4 1.44 12.86 2.23 2.68 0.47 17.07 145.03 2.75 2.79 1.03 3.71 4.63 1.33 9.56 1.8 9.07 11.51 2.42 37.05 3419 852.11

0.03 0.03 0.05 0.23 0.07 0.14 0.07 0.46 <.02 0.15 0.05 0.03 0.16 0.06 0.35 0.23 0.11 0.06 0.03 0.02 0.06 0.02 0.02 0.02 <.02 <.02 0.03 0.22 <.02 0.02 0.1 0.02 0.88 0.04 <.02 <.02 <.02 0.06 0.13 <.02 0.04 0.12 <.02 <.02 <.02 3.14 0.98

<.2 <.2 <.2 <.2 <.2 <.2 <.2 0.7 <.2 <.2 <.2 <.2 0.2 <.2 <.2 <.2 <.2 <.2 <.2 <.2 <.2 <.2 <.2 <.2 0.2 <.2 <.2 0.5 <.2 0.3 <.2 <.2 0.3 0.3 <.2 <.2 <.2 <.2 0.5 <.2 <.2 <.2 <.2 <.2 <.2 2.7 3.9

0.3 1.1 0.4 1 0.7 1 0.5 1.3 0.5 1.1 0.7 0.6 1.2 0.7 1 1.4 1.1 1.4 0.6 0.7 1.2 0.8 0.7 0.5 2.8 1.5 1 1.7 1.6 1.5 1.2 0.9 2.9 2 1.1 1.5 2.8 0.8 5 1 1.4 1.5 11.7 2 2.4 4.9 12.3

2 3.4 2.4 2.9 2.8 4.5 1.6 19.3 2 2.4 3.2 2.6 3 4.4 3.3 3.6 2.7 3.1 1.2 2.3 2.7 5.4 2 2.9 5.7 3.2 1.7 6.7 4.2 5.3 2 2.1 5.7 3.5 2.9 4.2 4 7.3 7.8 3.8 2.6 12.5 17.6 44.4 5.4 8.4 20.1

<1 <1 <1 <1 <1 <1 <1 <1 <1 <1 <1 <1 <1 <1 <1 <1 <1 5 1 <1 1 1 1 1 2 2 1 <1 1 2 <1 1 1 2 2 1 2 2 4 1 1 1 5 2 <1 5 2

1.9 12.6 1.9 15.2 4 6 1.6 3.8 3.9 4.9 2.1 1 5 1.6 7.6 10.5 6 1.2 <.5 1.2 4.5 1.8 2 0.5 1.4 1.6 1.8 13 1.5 1.9 12.1 1.6 22.6 1.7 1.5 4.1 2.7 1.4 7.3 1.5 2 2.9 34.5 2 2.3 3 45.7

0.82 0.67 1.06 0.81 0.99 1.2 1.02 1.87 1.05 1.16 0.92 0.69 1.01 0.91 0.97 0.89 0.89 1.16 0.27 0.38 0.22 0.2 0.1 0.37 0.23 0.5 0.2 0.18 0.47 0.6 0.13 0.39 0.71 0.89 0.5 0.32 0.93 0.36 0.12 0.57 0.5 0.74 3.6 0.63 0.9 0.63 1.33

2.2 1.5 2.1 0.9 3.1 0.9 1.2 2.9 1.8 0.9 1.5 2.8 1.4 1.3 1.1 1 1.6 18.9 5.1 1.7 2.5 3.5 3.2 2.6 3.6 2.1 1.1 1.4 1 4.1 10.7 2.3 2.6 4.1 2 4.4 1.7 2.3 1.3 2.9 1.6 4.8 9.3 2 1.5 2.1 18

<.05 <.05 <.05 <.05 <.05 <.05 <.05 <.05 <.05 <.05 <.05 <.05 <.05 <.05 <.05 <.05 <.05 <.05 <.05 <.05 <.05 <.05 <.05 <.05 <.05 0.11 <.05 <.05 <.05 0.06 0.09 <.05 <.05 <.05 0.06 <.05 0.25 <.05 0.08 <.05 0.06 0.1 <.05 0.83 0.06 <.05 <.05

0.07 <.05 0.07 0.06 <.05 <.05 <.05 0.08 <.05 <.05 <.05 <.05 <.05 <.05 <.05 <.05 <.05 0.07 <.05 <.05 <.05 <.05 <.05 <.05 <.05 <.05 <.05 <.05 <.05 <.05 <.05 <.05 <.05 <.05 <.05 <.05 0.13 <.05 <.05 <.05 <.05 <.05 <.05 0.06 <.05 <.05 <.05

29.5 40.7 33.2 34.1 30.3 29.6 27.0 34.5 37.5 29.8 21.9 29.9 24.1 18.2 38.5 29.5 15.1 19.6 22.1 32.9 41.2 30.4 20.6 11.9 7.9 17.5 18.9 11.7 20.3 19.0 38.1 24.1 7.6 13.4 26.4 15.7 13.6 23.3 9.6 17.8 19.9 16.1 9.2 15.0 12.8 20.2 9.5

12.4 18.1 15.9 30.0 21.5 26.1 18.4 24.5 17.7 26.8 12.7 11.9 27.4 18.3 23.2 37.1 25.5 8.5 8.3 16.2 13.0 8.9 8.7 10.8 17.7 7.7 13.4 38.6 15.3 9.9 47.9 5.3 47.9 11.3 7.9 26.8 15.7 5.7 33.3 5.6 5.6 10.0 97.9 13.0 9.5 65.4 354.1

Isotopic data (O, H, C, Sr) are also given.

d18O (&) 8.13 7.51 7.64 7.49 7.53 7.38 8.38 8.83 9.01 8.18 8.40 8.76 7.94 8.03 8.78 7.14 9.06 10.17 9.83 9.49 10.01 11.54 10.34 10.03 10.11 10.35 10.10 8.99 9.68 10.45 7.6 10.65 8.85 9.5 10.49 7.29 10.13 9.7 8.89 10.61 10.3 9.5 8.49 10.39 10.29 2.95 4.99

d2H (&) 56.6 53.7 56.2 51.0 53.5 55.0 57.7 55.1 57.4 60.6 61.2 60.9 57.8 58.1 57.8 51.4 57.4 71.9 65.9 60.5 65.9 70.8 63.4 65.7 73.8 69.3 68.2 59.9 60.3 67.3 70 71.6 61.8 65 68.5 47.9 67.9 61 59.7 71.5 69 68 58.1 68.6 68.5 31.6 38.8

87/86

Sr (pmc)

0.712836 0.713010 0.712952 0.712693 0.712566 0.712953 0.712787

14 C (pmc)

d13C (&)

78.6

7.8

73.7

9.3

80.2

8.0

69.3

7.3

72.5

8.4

85.5

10.1

42.9 51.6 26.8 15.1 76.9

8.2 9.0 5.9 6.1 10.8

8.2 45.1

6.8 9.1

10.7

6.8

5.3

8.0

5.2

7.6

0.7

8.4

2.8 6.3

8.2 8.9

0.712801

0.712779 0.712557 0.712795 0.712808 0.712756

0.712544 0.712719 0.712114

0.712242 0.711924

0.712662

0.711779 0.711504

W.M. Edmunds et al. / Applied Geochemistry 21 (2006) 2148–2170

Map no

W.M. Edmunds et al. / Applied Geochemistry 21 (2006) 2148–2170

2157

Table 3 Results for noble gases and tritium for representative samples from the Minqin Basin Sample

He (10 8 cm3 STP g 1)

Ne (10 7 cm3 STP g 1)

Ar (10 4 cm3 STP g 1)

Kr (10 8 cm3 STP g 1)

Xe (10 8 cm3 STP g 1)

3

He/4He (10 6)

M2 M4 M12a M12b M13 M15 M21 M22 M26 M28 M29 M30 M34 M36a M36b M43 M49

259.10 288.02 168.56 168.06 208.44 175.92 19.070 19.763 6.393 5.298 7.365 6.882 5.264 5.748 5.689 8.100 5.215

2.150 2.263 2.221 2.237 2.187 2.421 2.104 2.132 1.960 2.117 1.851 1.911 2.039 2.222 2.206 3.055 2.069

3.868 3.990 3.990 3.992 3.924 4.131 3.805 3.776 3.587 3.791 3.565 3.601 3.564 4.087 4.101 3.888 3.560

9.079 9.149 9.232 9.289 9.060 9.528 8.936 8.827 8.376 8.667 8.373 8.423 8.189 9.530 9.721 8.323 8.020

1.310 1.367 1.320 1.379 1.365 1.375 1.277 1.248 1.219 1.252 1.237 1.260 1.129 1.386 1.451 1.140 1.151

0.478 0.568 0.431 0.414 0.470 0.412 0.611 0.552 1.123 1.914 1.004 1.057 2.171 1.322 1.386 1.491 1.406

Errors (%)

±0.6

±0.8

±0.5

±1.6

±1.2

±1.9

3

H (TU)

0.00 ± 0.07

0.05 ± 0.03 1.52 ± 0.05 7.40 ± 0.14 0.48 ± 0.08 0.39 ± 0.05 7.38 ± 0.12 1.71 ± 0.07 0.25 ± 0.07 0.01 ± 0.07

Table 4 Results of fitting equation 1 to the measured concentrations of Ne, Ar, Kr, and Xe Sample

v2

P (%)

T (C)

DNe (%)

Radiogenic He (10

M2 M4 M12a M12b M13 M15 M21 M22 M26 M28 M29 M30 M34 M36a M36b M43 M49

0.16 1.45 0.09 0.34 1.82 0.07 0.77 0.64 0.03 0.19 0.34 1.39 0.53 0.01 0.03 0.14 0.54

69 23 76 56 18 80 38 42 85 66 56 24 47 94 86 71 46

6.2 ± 0.6 5.3 ± 0.4 6.5 ± 0.7 4.9 ± 0.3 5.1 ± 0.3 5.4 ± 0.3 7.1 ± 0.3 7.6 ± 0.6 8.0 ± 0.3 7.9 ± 0.4 7.6 ± 0.3 7.1 ± 0.3 10.8 ± 1.0 5.1 ± 0.7 3.5 ± 0.3 11.3 ± 0.4 10.1 ± 0.4

20.2 25.3 24.6 23.4 21.0 34.1 18.6 20.9 11.6 20.2 4.7 7.7 19.5 22.7 19.7 79.4 20.3

254.2 ± 1.5 282.7 ± 1.4 163.5 ± 1.0 162.9 ± 0.8 203.6 ± 1.1 170.3 ± 0.7 14.2 ± 0.1 14.8 ± 0.1 1.85 ± 0.07 0.40 ± 0.06 3.13 ± 0.06 2.48 ± 0.09 0.50 ± 0.09 0.67 ± 0.08 0.62 ± 0.08 0.28 ± 0.12 0.35 ± 0.06

8

cm3 STP g 1)

3

H–3He age (yr)

15.6 ± 3.0 18.6 ± 0.5

23.0 ± 1.1 14.1 ± 5.1 19.3 ± 2.0

v2 is the sum of the weighted squared deviations between modeled and measured concentrations. Its expected value is the number of degrees of freedom, which is one, since there are four constraints (Ne, Ar, Kr, Xe) and three free parameters (T, A, F) for each sample. p is the probability for v2 to be equal to or larger than the actual value due to random errors, although the model is correct. Fits with p < 1% are usually rejected (not occurring here). T is the model parameter for the equilibration temperature (noble gas temperature, NGT). DNe (%) = (Nemeas/Neeq 1) Æ 100% is the relative Ne excess above solubility equilibrium. radiogenic He is the He excess above the atmospheric He components predicted by the model. 3 H–3He age is the age calculated from 3H and 3He data where applicable (3H > 1 TU).

Isotopic data were obtained from four samples of moisture extracted from the 16 m moisture profile from unconsolidated sands at Beitu, including one from near the water table. These results are also

shown in Fig. 4 and plot along a line with a slope of 2.6 with an intercept on the meteoric water line of 6.5&, which is identical to the global weighted mean value for the rainfall from Zhangye. Chloride

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Fig. 4. Stable isotope plot of the ground and surface waters of the Minqin Basin and surrounding area. Mean annual rainfall for Zhangye station is shown. Sample numbers of outliers relate to those in Table 2.

was also measured on these profile samples and gave a mean Cl in moisture of 193.7 mg L 1 corresponding to a recharge rate of 2.8 mm a 1 based on an annual rainfall of 150 mm a 1 and Cl in rainfall of 2.8 mg L 1. This result is comparable to that from the Badain Jaran Desert some 150 km to the NW (Ma and Edmunds, 2006) where mean recharge rates in unconsolidated sands were between 0.95 and 1.33 mm a 1, in an area with 89 mm a 1 rainfall. Isotopic compositions for runoff from the Shiyang River entering the Minqin Basin (Hongyashan Reservoir) are also comparable with the modern regional rainfall and the direct local recharge through sands. It is unlikely therefore that modern recharge is significantly influenced by high altitude precipitation or snowmelt from the higher Qilian mountain range since this would be characterised by a lighter isotopic composition (although no field data are available to confirm). The groundwaters therefore are significantly lighter than modern recharge and must represent palaeowaters or their mixtures with modern water. 4.2. Origin and age of the groundwater The stable isotope data for all groundwaters (Fig. 4) show a range from 7.1& to 11.5& d18O. There are two exceptions (16 and 17), both

of intermediate depth, located in the extreme north of the basin in the area of the former lake, which lie along an evaporation line and are slightly saline. The deeper groundwaters (with depths ranging from 200 to 320 m) found mainly in the north of the basin generally have the lightest isotopic compositions. One shallow well (20) has a strongly depleted composition and another (3) shows evaporation from an isotopically light source. The interpretation of these data suggests that the parent groundwater is palaeowater with an isotopic composition of circa 10.5& d18O and unrelated to modern day recharge. There is some disturbance of the stratification and mixing induced by pumping, as well as possibly recycled irrigation water. Some of the heavier intermediate depth groundwaters ( 7& to 8& d18O) lie close to the main bifurcation of the Shiyang River around Minqin implying that the river source may also have contributed to recharge in the past. Radiocarbon analyses for 20 groundwaters are shown in Table 3 as percent modern carbon (pmc) where the range is from 0.6% to 85% modern. The d13C values lie between 6& and 9&. The plot of 14C along the flow line (section A) shows a progressive decrease in activity with lowest values of 0.6% modern found in the deepest waters in the north of the basin (Fig. 5). The d13C values however show no trend across the basin and have an average

2159

L

L

W.M. Edmunds et al. / Applied Geochemistry 21 (2006) 2148–2170

Fig. 5. Plot of

14

C, d13C, Cl and HCO3 in groundwaters along flow line A (see Fig. 2).

value of 8.1& d13C. Whilst the depth relationships indicate an aging trend from south to north, derivation of an absolute age of the groundwaters requires a knowledge of the source and initial activity of CO2 in the soil at the point and time of recharge, the activity of 14C in the carbonate minerals that may react to produce the total dissolved inorganic C as HCO3 in the groundwater, as well as of the stoichiometry of congruent and incongruent reactions along the flow path. The aquifer and recharge areas are almost exclusively composed of siliceous clastic, metamorphic or granitic rocks with an absence of carbonate minerals. The groundwaters are generally undersaturated with respect to calcite, a further indication that carbonate minerals are rare or absent in the formation. It is probable that the TDIC has been derived from soil CO2 reactions

with silicate minerals, in which case no dilution of the 14C activity would have taken place and that the values observed are due to closed system decay with no further evolution of TDIC within the formation. The vegetation in the Gobi Desert is predominantly C3 species (Wang et al., 2005) and this would give rise to a soil CO2 d13C of around 23&. The heaviest values that would be expected for TDIC would then be around 14& d13C, after allowing for isotopic fractionation at the near neutral pH (Table 2). Thus there remains a discrepancy in the model that could lead to correction of the radiocarbon data to give absolute ages based on available data. Until improved environmental data becomes available the 14C values as pmc are used to show relative ages and likely age bands to separate modern, Holocene and late Pleistocene groundwaters.

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A rough estimate of the initial 14C activity of the recharging groundwater might be provided by using the values obtained from samples that are identified as modern based on their 3H and 3H–3He signatures. From such an analysis (see below), an upper limit of the initial 14C activity of about 80 pmc is obtained. Using this value and assuming no incongruent reaction with carbonate minerals, in accordance with geochemical and isotopic indications, one can place the boundary between late Pleistocene and Holocene samples at 14C activities of roughly 20 pmc, corresponding to an estimated age of 11– 12 ka. Because 80 pmc is considered an upper limit of the initial 14C activity, such ages are maximum estimates, although probably correct within a few ka. The radiocarbon data are plotted vs d18O in Fig. 6 and there is a clear separation into isotopically light waters with late Pleistocene signatures and modern waters (>60% pmc) with heavier compositions. There is good evidence that isotopically light waters persisted into the Holocene. A distinct anomaly is found in site 29 in the centre west of the basin where very light water of Holocene age is found; this feature is considered further below. The 3H concentrations of about 7.4 TU in two wells (No. 28 and 34), clearly identify these as having a young (post-bomb-peak) component derived most probably from irrigation water. In combina-

tion with the He isotope data, these samples yield reliable 3H–3He ages of about 19 and 23 a. These samples also contain high 14C activities but are depleted in d18O – an anomaly which is discussed below. Two further samples (wells 26 and 36) have lower, but clearly significant 3H concentrations (1.5 and 1.7 TU respectively), indicating a young component that can be dated to an age of approximately 15 a. For the samples beyond the dating range of 3 H–3He (older than about 50 a), 4He can be used as a proxy for residence time in addition to 14C. Radiogenic 4He is produced in crustal rocks by the a-emitting nuclides in the U/Th decay series and thus accumulates in groundwater over time. Quantitative 4He-dating is however often impractical because unknown accumulation rates and the possibility of external sources contributing to the non-atmospheric excess He in groundwater. In this study, the terrigenic He excess is calculated by subtracting the atmosphere derived He (solubility equilibrium plus excess air components) obtained from the modelling of the heavier noble gases (see below) from the total measured He. Large excesses on the order of 2 · 10 6 cm3 STP g 1 ( 50 times higher than the solubility equilibrium concentration) were found in wells 2, 4, 12, 13 and 15. An analysis of the isotopic composition of these excesses shows that they have 3He/4He

Fig. 6. Plot of d18O against radiocarbon (expressed as percent modern carbon (pmc)) for shallow and deep groundwaters from the Minqin Basin.

W.M. Edmunds et al. / Applied Geochemistry 21 (2006) 2148–2170

ratios between (4–6) · 10 7, clearly lower than the atmospheric ratio Ra of 1.384 · 10 6, but also much higher than the typical value of 2 · 10 8 for crustal radiogenic He. These elevated 3He/4He ratios might indicate admixing of a small mantle-derived He component, as indicated also by Shi et al. (2001). Despite this finding, the non-atmospheric 4He excess is referred to as radiogenic in the following. In Fig. 7 it can be seen that there is a distinct order of magnitude difference in radiogenic 4He between waters of Modern/Holocene and those of late Pleistocene age based on 14C activities. Radiogenic 4He also increases systematically with distance from the Hongyashan Reservoir (inset in Fig. 7). Thus radiogenic 4He clearly serves as a useful qualitative age indicator in the study area and the regular age/ distance relationship from the two main isotopic residence time indicators (14C and 4He) suggests a homogeneous flow system across the basin. Noble gas recharge temperatures (NGTs) were calculated from the concentrations of dissolved atmospheric noble gases (Ne, Ar, Kr, Xe) according to the procedures and models outlined by Aeschbach-Hertig et al. (1999, 2000) and Kipfer et al. (2002). A uniform recharge altitude of 1250 m was assumed for all samples in these calculations; it is considered that the main recharge area is the coarse-grained aquifer in the piedmont of the Qilian Mts. Different models describing the composition of

2161

the excess air component were fitted to the measured concentrations to derive the best model parameters, including NGT. The closed-system equilibration model of Aeschbach-Hertig et al. (2000) was found to provide the best (and very good) fit for all samples and was thus applied in all calculations of NGT as well as of radiogenic 4 He and tritiogenic 3He for dating (see above). The fit results are summarised in Table 4, and in Fig. 8 the NGTs are plotted against radiocarbon as pmc. Waters from the late Pleistocene (wells 2, 4, 12, 13 and 15, having the lowest 14C activities and highest radiogenic He concentrations) form a distinct group (Fig. 8) with recharge temperatures ranging from 4.9 to 6.5 C with a mean value of 5.6. Except for well 36, the younger wells have higher recharge temperatures, although they seem to fall into two different groups. The first group (wells 34, 43, and 49) has much higher NGTs between 10.1 and 11.3 C (averaging at 10.7 C). These wells all have high 14C activities (above 74 pmc) and low radiogenic 4He contents, indicating young ages. The second group (wells 22, 26, 28, 29, and 30) has intermediate NGTs averaging at 7.6 C. With the exception of well 28, these wells have lower 14C activities (27–52 pmc) and higher radiogenic 4He contents, indicating older, but still Holocene ages. These two groups are also well distinguished based on the distance from the

Fig. 7. Radiogenic helium contents (4He) of Minqin Basin groundwaters plotted against radiocarbon expressed as pmc and also (inset) against distance from Hongya Reservoir.

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W.M. Edmunds et al. / Applied Geochemistry 21 (2006) 2148–2170

Fig. 8. Plot of noble gas recharge temperature against radiocarbon as pmc for shallow and deep groundwaters from the Minqin Basin.

Hongya Reservoir (see Fig. 3). Although well 28 would seem to belong to the intermediate group based on its location and NGT, it is clearly modern as shown by its 3H–3He age of 19 a. The modern mean annual air temperature (MAAT) in Minqin is 8.2 C, fitting better with the mean NGT of the intermediate group than that of the young group. Hence, there is some uncertainty about the NGT representative for Holocene conditions. The difference in NGT between the two groups of Holocene samples could be related to different recharge altitudes or recharge processes influencing the two areas. NGTs are usually assumed to closely represent the ground temperature in the recharge area, independent of the recharge process, because of gas exchange and equilibration near the water table (e.g. Stute and Schlosser, 1999). Although it is conceivable that under extreme conditions, such as rapid recharge along preferential pathways during flood events, such a final equilibration might not be reached and NGTs might differ from the soil temperature, such effects have never been demonstrated. If NGTs track ground temperatures well, an explanation for the NGT differences in the study area could be temporal and spatial variations in the relationship between ground and air temperature. Beyerle et al. (2003) showed that, in a tropical, semi-arid setting, ground temperatures can be about 3 C higher than MAAT, and hypothesized that this

difference was much smaller during past wet periods with increased vegetation. Since the two groups of Holocene samples in the Minqin Basin represent different periods and possibly different areas of recharge, such effects may play a role, although there are no independent indications that could be used to corroborate this. One major anomaly is found in sample 36, which was measured in duplicate with NGT of 5.1 and 3.5 C. These consistent data raise the possibility of a cold and wet phase during the late Holocene (at least as cold as the Last Glacial Maximum) with flooding possibly allowing local but not regional recharge. 4.3. Geochemical indicators of origin and groundwater evolution 4.3.1. Chloride and major ions In Table 1 the results are arranged in order of distance from Hongya Reservoir. The most important element is Cl, which is an inert tracer denoting inputs and providing information on physical processes especially and evaporation occurring during recharge as well as any additions along flow paths. The groundwaters are generally fresh with Cl ranging from 46 to 2090 mg L 1 (average value 321 mg L 1 and total mineralisation with an average of 763 mg L 1). The higher salinities show the influence of mixing with evaporated waters

W.M. Edmunds et al. / Applied Geochemistry 21 (2006) 2148–2170

In contrast to the main Minqin Basin, a new development project at South Lake taps the phreatic aquifer beneath the desert. This water with a Pleistocene isotopic signature has very low Cl (18.5 mg L 1) and total mineralisation (365 mg L 1). This differs from the water of the main Minqin Basin and because of its low salinity is considered to be indicative of the regional flow system

L

from the original lake basin (Fig. 5). Higher than normal salinities in the central and upper parts of the basin are considered to be the result of some irrigation return waters. The results overall suggest that an internal Cl source is required to produce the higher Cl values, since these are too high to be derived directly only from rainfall recharge.

2163

L

L

L Fig. 9. Plot of Br /Cl and Na/Cl for Minqin Basin groundwaters.

2164

W.M. Edmunds et al. / Applied Geochemistry 21 (2006) 2148–2170

and of the more abundant recharge during the late Pleistocene. It also contrasts markedly with the mean Cl of the Beitu unsaturated zone (194 mg L 1), which is representative of the low rates of modern recharge. The Br/Cl ratio may be used as an indicator of the origin of salinity, with values equal to or exceeding the marine ratio likely to originate from marine water or marine formation waters; values much lower than the marine ratio indicate an evaporite source (Edmunds, 1996). The plot of Br/Cl for the waters of the Minqin Basin (Fig. 9) indicates clearly a non-marine origin for the groundwaters with a strongly depleted ratio (average 0.0005 as compared with sea water of 0.0035). The very low ratio of Br/ Cl in Tangjia Lake is representative of an evaporite end member. Modern waters coming from the Hongyashan Reservoir are slightly enriched compared with the main body of groundwater but these may have been subject to human influence. The Br/Cl distributions in the groundwaters show that Cl may have originated partly by evaporative enrichment from rainfall but with some addition from within the continental sediments. High Cl waters display similar chemistry to the freshwaters and indicate evolution by evaporation. One exception is the brackish water (16) from the terminal lake area which has enrichment in Br/Cl indicating evolution involving organic rich lake sediments.

The salinity relationships are explored further using the plot of Cl vs d18O (Fig. 10). From the above discussion of d18O and 14C it is clear that the least isotopically-depleted waters are also the younger ones. However, there is no obvious relationship between salinity and age (or depth), which suggests that the flux of solutes into the basin has been variable with time. Any significant evaporative enrichment due to irrigation return water can be discounted in the waters so far investigated. Most groundwaters are neutral to slightly alkaline with pH mean value of 6.6–7.9. The waters have SO4/Cl around 2.5 and HCO3 in the range 100–770 mg L 1 with the highest values occurring in some shallow groundwaters. Saturation indices for calcite (not shown) calculated using PHREEQC (Parkhurst and Appelo, 1999) are just below saturation (mean values 0.24 and a range 0.66 to 0.28). This underlines the mainly continental character of the aquifer and suggests the virtual absence of carbonate minerals and silicate weathering being the dominant process. The groundwaters are also close to dolomite saturation (mean value SI dolomite 0.38, range 2.32 to 0.40) and this reflects the mMg/Ca ratios close to or exceeding 1. This may be explained by the weathering of reactive mafic minerals in the sediments derived from diorites in the Qilian Shan source rocks.

L Fig. 10. Plot of Cl vs d18O in Minqin Basin groundwaters.

W.M. Edmunds et al. / Applied Geochemistry 21 (2006) 2148–2170

A further significant characteristic of the groundwaters is the very low mK/Na ratio (mean 0.017), which would denote the dominance of albite over K-feldspars in the catchment area. The plot of Na vs Cl can be used as a first-order indicator of albite reaction along the flow lines (Fig. 9). The ratio mNa/Cl averages 1.5 and is similar to that in the Hongyashan Reservoir and the absence of any progressive increase across the basin indicates that little water rock interaction takes place over time. 4.3.2. Strontium isotopes and groundwater provenance Strontium isotopes were analysed from selected sites within the basin to constrain the origins of groundwater and solutes of different generations and depths. Samples of rainfall were also obtained in the surrounding region. The sediments of the Shiyang River Basin are predominantly clastic material but it was beyond the scope of the present study to conduct analyses of the various rock assemblages across the basin. The results are given in Table 2 and plotted as 87Sr/86Sr vs 1000/Sr in Fig. 11. The Minqin shallow groundwaters display a narrow range of Sr isotope signatures (0.712544– 0.713010) with variable Sr concentrations, whereas the deep well samples are less radiogenic and have a wider range of 87Sr/86Sr values (0.711504–

Fig. 11.

87

2165

0.712566). The most radiogenic signature is observed in the South Lake groundwater and there is a general trend towards less radiogenic values of 87 Sr/86Sr ratio towards the terminal lake (Table 2). The Hongyashan Reservoir ratio (0.712557) is similar to that of shallow groundwater. Local rain sampled at Baddam and Yabulai yielded 87Sr/86Sr ratios of 0.71097 and 0.711682 and significantly lower Sr concentrations (0.011 and 0.079 lg L 1, respectively) than observed in the groundwaters and the reservoir. Four samples of unsaturated sand from Beitu yielded a narrow range of whole rock 87Sr/86Sr ratios (0.720341–0.721258) which are comparable with that obtained by Yokee et al. (2004) for a sample of sand from the Tengger Desert (0.72045). These values are significantly higher than those reported for desert sands and river deposits from the Tarim Basin (Liu et al., 1994; Asahara et al., 1999) and for fine fractions (<20 lm) of loess from the Jungar and Qaidam Basins (Sun, 2002). Acetic acid leaching of the Beitu samples resulted in a mean Sr isotope signature for the secondary soil carbonate minerals of 0.712506 ± 0.000156. Yokee et al. (2004) using sequential extraction obtained a Sr isotope signature of 0.71269 for the evaporite minerals and 0.71281 for the carbonate minerals in the Tengger sand. They also studied the fine grained

Sr/86Sr vs 1000/Sr plot of waters from the Minqin Basin showing the likely relationship with source compositions and rainfall.

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fraction (mean particle size 15.7 lm) and found that the Sr isotope ratios of the extracted and residual fractions were consistently lower than that of the host rock reflecting differences in mineralogical composition with grain size. The Sr isotope signature for the water and acetic acid soluble minerals 0.7111 ± 0.0004 (mainly halite and calcite) in the fine sand fractions appears to be relatively uniform over a wide area. It is also very comparable to the 87 Sr/86Sr ratio obtained by Palmer and Edmond (1989) for the high stage of the Yellow River (0.7111 ± 0.0001). Although there is an overlap between the range of 87Sr/86Sr ratios displayed by the Minqin shallow groundwaters and the Sr isotope signature obtained for carbonate minerals from the Beitu sand samples, other indicators such as the major ion ratios suggest silicate weathering is the dominant process. Waters from the shallow wells extend the trend established between local rain and waters of the Hongya Reservoir towards the more radiogenic Sr isotope compositions displayed by silicate minerals (Fig. 11). However, the deep Minqin waters define a separate mixing line between groundwater compositions and less radiogenic signatures exhibited by fine grained evaporite and freshwater carbonate minerals, possibly influenced by changing lithology with depth. Elevated Cl and salinity levels in shallow waters at the northern end of the basin have been attributed to the influence of mixing with evaporated waters from the original lake basin. The more radiogenic signature displayed by the South Lake palaeowater could be explained by a greater contribution of labile Sr from the surface of silicate minerals, the result of time-dependent water rock interaction. 4.3.3. Redox controls and trace element evolution Redox relationships have a significant influence on the hydrochemistry of the Minqin Basin groundwaters. The groundwaters have near neutral pH and positive eH values (mean 216 mv) (Table 1). Most groundwaters also contain NO3, and Fe is generally present below detectable concentrations (0.03 mg L 1). Taken together these indicators suggest that, despite no measurements of dissolved O2, the groundwaters are oxidising and, bearing in mind the age indicators, that many have remained aerobic for several millennia; this phenomenon is similar to continental sediments found in other sedimentary basins (Winograd and Robertson, 1982; Edmunds et al., 2003). Plotted on an eH–pH diagram (not

shown here) the groundwaters fall inside the field of Fe(OH)3 stability and the low measured concentrations for dissolved Fe are consistent with theoretical considerations (Hem, 1985). The concentrations of nitrate as NO3–N are low but detectable with a mean value of 1.6 mg L 1 NO3–N (Fig. 12). Only a few samples have NO3 concentrations above 2 mg L 1 and the impact of irrigation return flow at least at the depths sampled cannot be significant. In the main agricultural area at the immediate water table the effects are predicted to be higher. There is no correlation between the ‘‘modern’’ waters as indicated by age indicators and the samples having higher than background NO3 and therefore the scatter in NO3 may be entirely natural. The redox related parameters are plotted for the Minqin Basin in Fig. 12, including key trace elements – U, Cr, Mo and Se against distance and depth. The aerobic conditions favour the stability of some trace metals and high concentrations have been described for certain elements from other continental sedimentary basins from the UK and Algeria for example (Smedley and Edmunds, 2001; Edmunds et al., 2003). The concentrations of U are particularly high (mean value 27.5 lg L 1), indicating a mineralogical source in the continental deposits and also that U is stabilised under the near neutral and aerobic conditions in the basin as the complex (UO2)(CO3)H2O. The aerobic and mildly alkaline conditions also favour the stability of Se as SeO24 (Smedley and Edmunds, 2001; Fordyce, 2005) and the mean concentrations in the Minqin groundwaters are 5.8 lg L 1 with maxima occurring in the most saline waters where values exceed 20 lg L 1. The mean values found here however are below the World Heath Organisation guide value of 10 lg L 1 and are not anomalous, unlike in certain areas of NE and central China (Tan, 1989). Chromium is also a significant trace element being stabilised as CrO3 and is the only trace metal to show a progressive increase, most likely due to reaction with mafic minerals along the profile and higher concentrations are also prevalent at depth. Shallow groundwaters have an average of 5.3 lg L 1 Cr, whilst the deeper waters have an average of 15.6 lg L 1. Molybdenum, stabilised as the MoO24 complex, is detectable in all samples and has an average concentration of 3.2 lg L 1 but with no obvious spatial relationships. Manganese in the groundwaters remains at very low concentrations (0.23 lg L 1 except for three outliers) under the oxidising conditions.

L

L

L

L

L

L

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Fig. 12. Downgradient plot of redox related parameters in groundwaters from the Minqin Basin.

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Barium concentrations (median 21 lg L 1) are rather similar and calculations using PHREEQC indicate that almost all waters are at barite saturation (mean SI 0.17) controlled by SO4 concentrations. Fluoride is also low (mean 0.41 lg L F) and all waters are undersaturated with respect to fluorite (mean SI 1.79). The distribution and concentration ranges are consistent with a non-carbonate aquifer and with a relatively uniform pattern probably having rainfall as a major source, concentrated by evaporation during recharge. The concentrations of other minor and trace elements, which are significant for environment and/or health, are shown in Table 2 in order of distance from the reservoir. Whilst some trace elements forming oxy-anions (U, Cr, Se), are of high abundance, others such as V, As, Sb, Li and Mo are low in concentration and indicate low geochemical abundance of these elements in this lithology. Other trace metals Al, Zn, Cu, Ni, Co, Cd also are very low abundance with mean concentrations at or below 1 lg L 1 representing their natural baseline values. 5. Conclusions A multi-tracer approach involving chemical isotopic and noble gas indicators has enabled the groundwater evolution in the Minqin Basin to be constructed. This has been hampered to some extent by lack of detailed knowledge of the hydrogeology, yet the geochemical patterns also provide insight into the flow regime. It appears that in this stratified basin a continuous flow regime exists from the Qilian Mts to discharge areas in the terminal lake area, and possibly beyond. There is a regular increase in groundwater residence time from south to north, indicated by radiocarbon and also by indicators such as the accumulation of 4He and by stable isotopes. Although the assignment of an absolute age is problematic, the value of around 20% modern carbon can be taken as the approximate boundary between waters of late Pleistocene and Holocene age. It may be possible to provide more accurate age estimates in the future, when more data are available on C isotope geochemistry in the region. The characterisation and extent of modern recharge is determined using chemical and isotopic tracers. The isotopic signature of modern rainfall coincides with that of evaporated unsaturated zone moisture as well as that of the Hongyashan Reservoir. The Cl mass balance confirms the small

amount of modern recharge beneath dune areas. Stable isotopes (d18O and d2H) provide a distinction between modern and ancient waters. The signature of groundwater from the late Pleistocene differs markedly from that of the Holocene, as shown by the stable isotopic compositions (around 10.5& d18O as compared with values of 7& at the present day). This primary difference is similar to that observed by Zongyu et al. (2003) for the North China Plain near Beijing. These authors also show a progressive enrichment in d18O values from about 7 ka BP and an optimum around 6–4 ka BP equivalent to the global warming and a strengthened monsoon. The late Pleistocene mean recharge temperature was about 5.6 C, which is some 2–3 cooler than the modern mean annual air temperature and the recharge temperature shown by several Holocene samples. However, a group of modern samples indicate even higher recharge temperatures of about 10.7 C, which complicates the assessment of the exact magnitude of the temperature change between the late Pleistocene and the Holocene. Thus it is clear that the late Pleistocene climatic conditions were significantly cooler than modern, but it is difficult to assign an unambiguous value for the cooling. A significant finding from the present study is the distinct cooler climatic anomaly, shown by NGT results for water of Holocene age. A few other Holocene waters from shallow depths also have very light stable isotope signatures (but no NGTs). This suggests the likelihood of unstable Holocene climates with one or more cold and wet phases (at least as cold as the LGM) with flooding allowing local but not regional recharge. It is clear therefore that the groundwaters in the Minqin Basin (and adjacent areas such as South Lake) represent continuous recharge mainly during the cold and wet phases of the late Pleistocene and also during the early Holocene. With a flow path of some 50 km from the recharge area the radiocarbon data suggest a minimum groundwater velocity of around 2 m a 1. Recharge to the aquifer from the Shiyang River is considered to have been of minor importance compared with regional groundwater flow. Until the modern era the higher groundwater levels were probably also feeding the terminal lake. The relative fresh water quality of the regional groundwater could also indicate that this is part of a much large flow system controlled ultimately by the base level of the Yellow River.

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Diffuse recharge in the area is very low (>3 mm a 1) and the groundwater resource is essentially non-renewable, as shown also by the falling water table (Ma et al., 2005). Until the mid-20th century it is likely that the upstream aquifer was locally fed along its course by recharge from the Shiyang River but the upstream abstraction and dam construction has all but eliminated this source. The presence of a minor component of modern recharge is confirmed for several shallow wells from the 3H (3He–3H) evidence. This is likely to be the result of irrigation return waters, since some of these samples are from areas away from the river channels and due to the fact that the natural river flows were unlikely to have occurred much since the 1960s. The quality of the deeper groundwater in the basin is generally good, around 1 g/L, and human impacts (indicated by low NO3 concentrations) are so far minimal, although pollution would be expected to be retained in the unsaturated zone or at the immediate water table where few samples were taken. The hydrogeochemistry is very typical of a non-carbonate continental aquifer (sands and inter-bedded clays) and the groundwaters are aerobic, allowing NO3 to remain stable. A further consequence of the oxidising conditions is to allow build up of some trace elements (U, Cr and Se). The concentrations of Cr in all groundwaters are still below WHO limits (50 lg L 1) as is the case for Se (10 lg L 1); no guidelines for upper U concentrations in water are currently given. Fluoride concentrations are low and may create dental problems. Arsenic concentrations are very low and pose no health problem. The scientific results have important implications for groundwater management in the Minqin Basin, probably the most water-stressed of the basins in NW China and in a region destined for rapid development under the West Development strategy. Since the bulk of the water resource is derived from recharge during a past wetter climate any renewal of groundwater under the modern rainfall regime and its arid geographical location is going to very limited. The rapidly falling water table is the combined result of exploiting palaeowater, high pumping rates and the intensive agricultural development using water inefficient practices. An improved hydrogeological model is required to quantify the available resources, but for development purposes it is prudent to regard the resource as non-renewable and therefore exhaustion of recoverable groundwater would take place within decades without drastic

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measures to curtail abstraction rates. If this historically important area is not to become a wasteland, then groundwater must be more highly valued and be considered a strategic reserve to be available for the next and subsequent generations. Acknowledgments We would like to acknowledge receipt of a NERC award (1141.1005) for analysis of radiocarbon. The research is supported by the National Science Foundation of China (No. 40302031). This work also forms part of a wider UK–China collaboration and we acknowledge the support of the Royal Society through its link scheme (PEK/0992/ 306). We wish to thank George Darling and Janice Trafford for respectively coordinating the stable isotope and chemical analytical work and Jo Green for technical support in the Sr isotope laboratory. This paper has benefited considerable from reviews by Andrew Herczeg and Luc Aquilina and we warmly acknowledge their advice and suggestions for improvement. References Aeschbach-Hertig, W., Peeters, F., Beyerle, U., Kipfer, R., 1999. Interpretation of dissolved atmospheric noble gases in natural waters. Water Resour. Res. 35, 2779–2792. Aeschbach-Hertig, W., Peeters, F., Beyerle, U., Kipfer, R., 2000. Palaeotemperature reconstruction from noble gases in ground water taking into account equilibration with entrapped air. Nature 405, 1040–1044. Asahara, Y., Tanaka, T., Kamioka, H., Nishimura, A., Yamazaki, T., 1999. Provenance of the north Pacific sediments and process of source material transport as derived from Rb–Sr isotopic systematics. Chem. Geol. 158, 271–291. Beyerle, U., Aeschbach-Hertig, W., Imboden, D.M., Baur, H., Graf, T., Kipfer, R., 2000. A mass spectrometric system for the analysis of noble gases and tritium from water samples. Environ. Sci. Technol. 34, 2042–2050. Beyerle, U., Ru¨edi, J., Leuenberger, M., Aeschbach-Hertig, W., Peeters, F., Kipfer, R., Dodo, A., 2003. Evidence for periods of wetter and cooler climate in the Sahel between 6 and 40 kyr BP derived from groundwater. Geophys. Res. Lett. 30, 1173. Chen, L., Qu, Y., 1992. Water and Land Resources and their Rational Utilisation and Development in the Hexi Region. Science Press, Beijing, 143-176. Chen, L., Qu, Y., Chen, H., 1982. The land and water resources rational utilization in the Shiyang river basin. J. Desert Res. 2, 6–13. Chen, F., Wu, W., Holmes, J.A., Madsen, D.B., Zhu, Y., Jin, M., Oviatt, C.G., 2003. A mid-Holocene drought interval as evidenced by lake dessication in the Alashan plateau, Inner Mongolia, China. Chin. Sci. Bull. 48, 1401–1410. Chen, Z.Y., Nie, Z., Zhang, H.S., He, M.L., 2004. Groundwater renewability based on groundwater ages in the Heihe Valley

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