Profiles of geochemical and isotopic signatures from the Helan Mountains to the eastern Tengger Desert, northwestern China

Journal of Arid Environments 90 (2013) 77e87

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Profiles of geochemical and isotopic signatures from the Helan Mountains to the eastern Tengger Desert, northwestern China Z. Ding a, b, c, J. Ma b, *, W. Zhao b, Y. Jiang c, A. Love d a

Water Environment Institute, Chinese Academy for Environmental Planning, State Environmental Protection Administration, No. 8, Dayangfang Beiyuan Road, Beijing 100012, China Key Laboratory of Western China’s Environmental Systems (Ministry of Education), Lanzhou University, 222 South Tianshui Road, Lanzhou 730000, China c State Key Laboratory of Earth Surface Processes and Resource Ecology, College of Resources Science & Technology, Beijing Normal University, 19 Xinjiekouwai Street, Beijing 100875, China d School of the Environment, Flinders University, Adelaide 5001, South Australia, Australia b

a r t i c l e i n f o

a b s t r a c t

Article history: Received 24 November 2011 Received in revised form 10 July 2012 Accepted 5 October 2012 Available online 1 December 2012

We used environmental tracers to provide insights into groundwater evolution and flow processes on the eastern margin of China’s Tengger Desert and determine recharge sources and recharge timing along the flow paths. The chemical composition was generally Naþ enriched, with no dominant anions. The chemistry is strongly influenced by evaporation and subsequent dissolution of minerals during recharge in the rainy season. Other processes, including cation exchange and weathering, also contribute to the water composition. The median groundwater d18O value was around 9.1&, and most groundwater in the basin was depleted in heavy isotopes. d13C in deep groundwater ranged from 4.73 to 9.56&, indicating that carbonate mixing is common but that isolated carbonate still exists. Radiocarbon values in groundwater ranged from 5.4 to 63.5 pmc. We estimated a residence time of 17.9 and 19.6 kyr at two desert sites, revealing that some replenishment of desert aquifers occurred in the late Pleistocene and early Holocene, when some of the upper reaches of rivers were characterized as modern water. Ó 2012 Elsevier Ltd. All rights reserved.

Keywords: Environmental tracers Groundwater recharge Radiocarbon Residence time

1. Introduction In arid northwestern China, population dynamics and agricultural activities are putting heightened strain on the region’s limited water resources. Groundwater is the most important water supply in this area, so increased knowledge of the geochemical processes that control groundwater chemical composition would improve understanding of the region’s groundwater systems and guide management of this resource. The southern and western edges of the Tengger Desert in northwestern China have been widely studied to examine the stresses on the groundwater resources in this area (Chen et al., 2003; Ma et al., 2005, 2009a). However, few studies have addressed the desert’s eastern edge, bounded by the Helan Mountains, the Yellow River and China Loess Plateau begins (Fig. 1), which form a natural border for the desert (Wang et al., 2000). Any investigation of the eastern edge is not trivial due to the complex geology and the lack of prior data on the area. The hydrochemical models focusing on different chemical signatures have been applied to investigate groundwater recharge,

* Corresponding author. Tel.: þ86 931 8912436; fax: þ86 931 8912330. E-mail address: [email protected] (J. Ma). 0140-1963/$ e see front matter Ó 2012 Elsevier Ltd. All rights reserved. http://dx.doi.org/10.1016/j.jaridenv.2012.10.019

paleoclimate evolution, and anthropogenic effects in the deserts of northwestern China (Chen et al., 2003; Edmunds et al., 2006; Ma et al., 2009a, 2009b; Pachur et al., 1995; Zhang et al., 1998). Nevertheless, these models have not been applied to the eastern Tengger Desert due to a lack of sufficient high-quality data, so details of this region’s groundwater flow processes and the historical evolution of groundwater quality are poorly documented and understood. Thus, they require further investigation. In the present study, we used environmental tracers, including ions, stable isotopes, and radiocarbon, to provide insights into the evolution of groundwater flow processes at the eastern margin of the Tengger Desert in terms of the recharge sources and recharge times along the main flow paths. The goals of our analyses were to identify the groundwater system’s recharge sources, to understand the subsurface hydrogeochemical processes, to ascertain groundwater flow directions, and to estimate the groundwater residence time. By combining the hydrogeochemical evolution revealed by the tracer analysis with the recharge pathways derived from relationships among different parts of the Tengger Desert and discharge zones in the inner desert, it becomes possible to delineate the climate history and better understand the extent of anthropogenic effects. These results let us reconstruct the palaeohydrologic conditions. Such knowledge is essential for ensuring

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Z. Ding et al. / Journal of Arid Environments 90 (2013) 77e87

Fig. 1. Location of the study area in northwestern China and of the main geographic features of the eastern Tengger Desert. Sample locations are indicated by numbers with an HL prefix. AB represents the geological cross-section shown in Fig. 2.

the safety of the region’s supply of drinking water, and is especially important to support sustainable management of the region’s water resources and meet the needs of future generations, while accounting for the effects of climate change on these resources. 2. Study area and hydrogeology 2.1. General settings The Tengger Desert is the fourth largest desert in China, and covers an area of about 36 700 km2. The region’s dune fields are surrounded by goby (deserts with gravel and cobble surfaces) and are undergoing deflation or accumulation along the piedmont zones, fossil lake basins, and desert gorges. Lake records from the region provided evidence of intermittent, short-lived wet periods with rising lake levels during the earlier wet period before 10 000 yr BP (Zhang et al., 2004). The desert’s climate was extremely dry during the early Holocene, relatively wet from 7200 to 5200 cal yr BP, highly variable from 5200 to 3000 cal yr BP, and dry again since 3000 cal yr BP (Zhao et al., 2008). At present, there are no permanent lakes in the Tengger Desert (Zhang et al., 2004). Our study area included the lowland area along the western margin of the Helan Mountains that form the eastern boundary of the Tengger Desert. The Helan Mountains are approximately 150 km long and 30 km wide, and are oriented from northeast to

southwest. They rise abruptly from the desert, which lies at an elevation of 1000e1500 m asl, to reach a maximum height of 3556 m asl. The study area consists primarily of unconsolidated sand dunes, some smaller goby, and scattered oases, most of which tend to be oriented parallel to the Helan Mountains. This region contains numerous playas, lake basins, and marshes (e.g., Tonghu Lake), and the sediments were transported into the desert during intermittent periods of high aeolian activity with predominantly westerly winds during the Holocene. The climate is strongly continental. The annual precipitation decreases gradually from east to west, from 419.9 mm/yr in the Helan Mountains piedmont to 146.1 mm/yr in the western Tengger desert (AZWCB, 2002). However, the average potential evaporation increases from 1400 mm/yr in the piedmont to 2316.5 mm/yr in the desert area. The annual mean temperature at Alxa Zuoqi is 7.5  C, ranging from 23  C in January to 38  C in July. Areas of frozen ground can be found between late October and early April, with the maximum frozen depth reaching 2.07 m. Alxa Zuoqi is the capital of our study area (Fig. 1). It is a wellirrigated oasis that has contributed greatly to the local livestock and agriculture industries since its development in 1970 (Wang et al., 2000). Settlement is currently sparse and land use is limited to subsistence agriculture and grazing. There are no perennial streams, but a series of fan-edge depressions near modern alluvial fans have developed in the desert. The area has

Z. Ding et al. / Journal of Arid Environments 90 (2013) 77e87

a complex network of water channels for agricultural use, and water infiltration from irrigation represents a significant recharge source for the aquifer. 2.2. Hydrogeology The regional Quaternary aquifers are alluvial fans in the piedmont plains and fluvial deposits at lower elevations in the basin. Although there is no reliable geological cross-section available, data from a series of boreholes running from southwest to northeast (Fig. 1) provides clues to the regional geological background (Fig. 2) (AZWCB, 2002). The eastern Tengger Desert is part of the western marginal area of the North China Platform, with widespread absence of Early Ordovician (Tremadoc) strata (Zhou and Dean, 1996). The parallel active reverse faults and derivative Cenozoic uplift along the piedmont of the Helan Mountains resulted in complex groundwater hydraulic contacts between the alluvial sediments and the desert aquifer, with outcrops of Carboniferous, Permian, and Triassic systems. These systems may provide the lower confining unit for the region’s shallow aquifers. The first aquifer is buried at a depth of 10e40 m below the surface and is unconfined, with a thickness of 10e30 m, although there are 1- to 3-m-thick discontinuous aquiclude layers at the top. The second aquifer is deeper (approximately 120e260 m), and serves as the main developing layer, with a composition of coarse sand and gravel. The primary porosity of the aquifer is a mixture of framework porosity, diminished by a degree of carbonate cementation, and uncemented inter-particle porosity within the unconsolidated carbonate sands. To the western side, the surface topography is deeply incised, with middle- to late-Pleistocene alluvial fans stretching from small, narrow canyon mouths into the extensive dune fields of the Tengger Desert. The Lower Pleistocene sediments comprise semi-consolidated carbonateedolomite conglomerate and sandy conglomerate, which undergo a gradual transition towards finer grained and well-stratified fluviatile and lacustrine facies ranging in thickness from 50 m in the south to a thickness of less than 1 m in the west (Fig. 2). At the southern and northern edges of the study area, the Quaternary sediments comprise the desert’s major shallow aquifer, which can be found at shallow depths (20e70 m) in many locations and emerge to form seasonal lakes in many interdune areas. Confined or semi-confined conditions are apparent in the vicinity of some lakes, where travertine islands have formed as a result of fresh groundwater being forced upwards under pressure.

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Low-porosity lacustrine sediments or interbedded sandstone may provide the confining layers locally. Groundwater generally flows from the piedmont area towards the center of the basin. However, regional inputs to the shallow Quaternary aquifer and deeper groundwater systems are poorly understood. The measured permeability coefficient from borehole data varies from 2.5 to 45 m/d in this sub-basin, but the whole Quaternary sediment of the uplifted highlands (the piedmont of the Helan Mountains) is unsaturated. The depths to the water table range from 50 to 200 m in the upper alluvial fan to between 0.8 m (at Zini) and 6.24 m (at Luanjing) in the lowland areas west of the mountains. 3. Materials and methods Samples of water were obtained from surface water, shallow and deep wells, and springs in July 2008. Among the four surface water sites, one was at the Helan Reservoir, one was a seepage area in the piedmont of the Helan Mountains, one was at Zini Lake, and one was at a Yellow River irrigation channel. The major groundwater samples were obtained from farm wells along the eastern edge of the Tengger Desert and from the typical lacustrine facies, and were primarily from irrigation wells ranging in depth from 3 to 300 m (Table 1). The shallow wells were purged with a submersible pump for at least 20 min before sample collection, and all deep wells had been pumping continuously for 1 h prior to sampling. Each groundwater sample was divided into two portions. One portion was retained for the stable O, H, and C isotope analyses, and the other portion was filtered through 0.45-mm filter paper for chemical analysis. Samples for cation analysis were acidified to a pH <2 to preserve the cations using 1% HNO3. Unacidified samples were collected for the anion and stable isotope analysis. Unfiltered 1-L groundwater samples were collected for the radiocarbon (14C) analysis. On-site analyses included water temperature, specific electrical conductance (SEC), total alkalinity (as HCO 3 ) measured by titration, and pH. Major ion concentrations were analyzed by means of inductively coupled plasma optical-emission spectrometry (ARL3400C) at Lanzhou University. Samples for the stable isotope analysis (18O, 2H) were measured using a MAT-252 isotope-ratio mass spectrometer at the State Key Laboratory of Lithospheric Evolution, Institute of Geology and Geophysics, Chinese Academy of Sciences. d18O and d2H values were measured relative to internal standards that were calibrated using the IAEA standard mean ocean water

Fig. 2. Hydrogeological cross-section along transect AB in Fig. 1. Arrows represent the direction of subsurface water flow (AZWCB, 2002). Numbers represent borehole numbers from the research report of AZWCB, 2002.

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Table 1 Field data: ion concentrations and stable isotope parameters in groundwater and surface water. Station type

pH

Temp. ( C)

Depth (m)

HL1 HL7 HL9 HL26

Helan Reservoir Zini Lake Spring Yellow River

6.33 8.51 8.22 8.11

14.2 26.1 16.5 24.5

e e e e

HL2 HL4 HL6 HL14 HL24 HL25 HL27

Shallow Shallow Shallow Shallow Shallow Shallow Shallow

6.54 7.28 7.26 7.75 8.71 8.00 7.37

14.9 13.4 17.6 11.1 17.4 19.4 19.8

HL3 HL5 HL8 HL10 HL11 HL12 HL13 HL15 HL16 HL17 HL18 HL19 HL20 HL21 HL22 HL23

Deep well Deep well Deep well Deep well Deep well Deep well Deep well Deep well Deep well Deep well Deep well Deep well Deep well Deep well Deep well Deep well

7.65 7.67 8.75 8.07 7.81 7.41 7.81 8.35 8.22 7.48 7.36 7.19 7.67 7.43 8.74 7.98

16.8 16.1 18.6 19.1 15.5 13.7 14.1 13.6 14.7 17.4 18.6 18.3 27.4 12.7 22.5 16.1

well well well well well well well

Concentration (mg/L)

Saturation index

d18O (&)

d2H (&)

TDS

Na

K

Mg

Ca

Cl

SO4

HCO3

NO3

Calcite

Dolomite

755 1230 740 377

18.85 1653.27 123.21 32.20

3.25 106.92 6.39 2.21

12.67 338.45 29.59 16.46

107.76 100.49 40.25 64.89

31.09 2055.51 84.28 32.50

67.30 1975.83 161.79 68.74

242.65 76.94 198.26 210.10

12.42 0.00 11.61 6.43

0.75014 0.37719 0.56006 0.73857

2.0811 1.6217 1.3384 1.2312

8.64 1.50 e 9.65

71.77 34.44 e 72.29

9 4 4 20 30 6 3

564 1400 1550 555 480 1680 1870

114.01 341.48 435.91 60.01 52.07 385.84 332.58

3.48 14.95 12.70 3.84 8.19 25.64 36.16

15.17 36.16 28.18 29.10 21.44 51.56 63.57

61.35 49.74 42.15 69.34 31.31 65.81 100.37

79.02 348.30 319.81 39.77 62.64 316.38 329.30

127.03 288.12 356.21 81.04 87.00 392.08 394.34

236.73 239.69 313.67 355.09 142.04 337.34 366.93

24.31 61.67 0.00 13.18 26.74 29.77 83.47

0.8241 0.27362 0.27443 0.61219 0.81539 0.65929 0.25515

1.9056 0.33253 0.37376 1.1974 1.8217 1.5632 0.6609

9.06 e e e 6.89 e 7.09

73.26 e e e 63.69 e 64.14

100 100 100 200 120 100 100 100 100 300 260 100 100 100 129 100

1450 676 368 548 422 621 470 852 910 839 1260 1750 1330 990 392 1560

283.71 115.92 36.64 57.94 44.97 76.08 44.52 275.79 123.64 195.95 353.07 567.95 364.11 236.24 31.84 294.00

10.10 0.00 1.31 2.62 1.48 1.46 2.93 4.00 4.62 5.67 18.92 10.05 12.10 15.88 6.94 10.79

48.64 21.39 16.58 34.65 26.39 39.00 35.70 192.12 61.15 19.46 20.46 45.66 30.27 32.10 20.60 66.69

95.04 20.66 64.62 77.38 51.55 30.78 50.24 106.17 62.63 56.08 25.36 62.90 30.23 47.50 48.97 130.18

374.67 136.26 53.63 83.11 44.62 64.88 61.00 560.41 123.73 113.95 249.24 574.99 315.03 176.52 52.44 349.05

257.10 89.86 59.45 126.87 78.88 114.60 82.79 625.43 216.66 187.14 325.83 503.24 292.54 179.96 53.54 402.50

189.38 106.53 178.73 230.81 213.06 260.40 218.97 285.34 266.32 343.26 218.97 204.18 230.81 319.58 165.71 284.08

47.80 22.28 14.49 28.74 28.39 22.66 44.52 25.66 31.07 0.00 6.62 16.21 14.86 25.80 34.96 27.75

0.27243 0.4839 1.2512 0.76597 0.34845 0.20886 0.33804 1.0463 0.82827 0.18684 0.52385 0.40634 0.12423 0.026938 1.098 0.8572

0.60558 0.60321 2.2697 1.5338 0.75629 0.035007 0.87827 2.7095 2.0004 0.26364 0.79286 0.60177 0.092353 0.23535 2.1771 1.7774

11.00 10.72 10.38 9.81 10.12 10.26 e e 10.90 10.75 11.16 10.82 e 8.63 6.56 8.88

90.88 90.57 80.08 76.62 78.88 81.50 e e 85.54 84.13 87.07 84.73 e 77.47 61.02 77.95

Z. Ding et al. / Journal of Arid Environments 90 (2013) 77e87

Sample ID

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(SMOW), Greenland ice sheet project (GISP), and standard light Antarctic precipitation (SLAP) standards. Radiocarbon (13C, 14C) samples were prepared and analyzed by means of Atomic Mass Spectrometry (AMS) at the SSAMS Radiocarbon Dating Centre of the Australian National University. CO2 from dissolved inorganic carbon (DIC) was liberated by acidification under vacuum in sealed sidearm vessels. The percent modern carbon (pmc) values had 1s errors <0.5 in all cases. d13C values were expressed relative to the Vienna Pee Dee Belemnite (PDB) standard. 4. Results and discussion 4.1. Salinity and recharge environment In general, the samples had neutral to slightly alkaline pH (7.19e 8.75), with the exception of the Helan Mountains reservoir sample (HL1), with pH 6.33, and one shallow well sample (HL2), with pH 6.54. The total dissolved solids (TDS) concentration in groundwater samples increased gradually between the boreholes in the upper slopes and those at mid-slope to plains locations. TDS ranged from 368 to 1870 mg/L and a standard deviation of 446.9 mg/L. On this basis, the groundwater can be considered freshwater or moderately saline water. Higher TDS (>1400 mg/L) was found in shallow water (e.g., HL4, HL6, HL25, HL27) at the lowest elevations in the basin. Samples near the western foothills of the Helan Mountains (HL2, HL14) had lower TDS (<600 mg/L). The concentration difference of 1200 mg/L between HL24 and HL25 (1680 mg/L) is interesting because it suggests that the freshwater and saline water had different origins even though they were separated by only 2.9 km. A Quaternary clay sediment aquiclude may exist in this area and may prevent a hydraulic relationship between these two sites. HL24 is located in a scenic area that provides a good natural habitat with an abundant freshwater supply; in comparison, HL25 is located in a salt marsh that produces sodium products such as halite and mirabilite, thus its water may be derived from a mixture with water from the lower confined PermianeTriassic sandstone formation. The chemical composition of the water in our study area was generally enriched in Naþ, but had no dominant anion type (Fig. 3). An exception was HL14, which is a spring well in the Helan Mountains near Hongyuan Village and has a relatively low Cl/(SO2 4 þ

Fig. 3. Piper diagram comparing the cation and anion compositions of the surface and groundwater samples.

81

HCO 3 ) ratio. Deep wells in the southern Helan Mountains had ionic compositions similar to those of wells at the boundary between the mountains and the desert and similar to those of samples from the northern part of the Zini Lake area (e.g., HL7). Two pairs of wells were distinct from the other deep wells. HL8 and HL11, from the Helan piedmont area, had higher Ca2þ and HCO 3 concentrations than many other deep wells, and samples from HL10 (which is affected by the Fuda chemical industry) and HL22 (near Malian Lake) had higher Mg2þ, Ca2þ, and HCO 3 concentrations than other deep wells. The major trends in groundwater salinity can be traced back to the recharge source, recharge rate, and evaporation rate. In the Helan Mountains, the shallow groundwater sampled in this study was mainly recharged by precipitation (This will be discussed in Section 4.3). In different flow paths, the wells at higher elevations in the mountains (HL1 and HL2) had a lower potential evaporation rate, and therefore had low salinity. In contrast, the shallow wells in low-elevation areas of the desert experienced intense evaporation due to the warm and dry climate, with an average annual precipitation of only 150 mm and a groundwater depth of 10e30 m; thus, there seems little possibility for effective recharge of groundwater by precipitation, and the strong evaporation will lead to high salinity. In contrast, the groundwater in the deep wells is less influenced by evaporation, so its salinity is typically much lower than that in shallow wells. 4.2. Major ions as tracers Plots of dissolved ions against Cl (Fig. 4) showed correlations ranging from 0.15 to 0.99, and the shallow and deep groundwater did not differ significantly (Pearson Correlation Test, p < 0.05)  (Table 2). Naþ and SO2 4 were most strongly correlated with Cl . Two physical processes may explain these relationships. Mixing of two groundwater bodies with different end-member compositions (i.e., fresh and saline water) is one possible explanation. However, because of the large number of potential end members that can occur within the shallow aquifer’s flow regime, simple mixing of two discrete groundwater bodies seems unlikely. The similarity of the major ion compositions in groundwater from the Helan Mountains piedmont and from the eastern Tengger Desert is also consistent with a common hydrogeochemical history for water in the two areas. Concentration or dilution of a single water input due to variations in evaporation during recharge is a second possibility. In this scenario, groundwater with a low ionic concentration may derive from areas with high recharge rates, such as in areas with sand dunes. In contrast, brackish groundwater derives from areas with low recharge rates, where long residence times for the water within the unsaturated zone combined with higher evaporation rates concentrate ions before they reach the water table (Love et al., 1993). In both cases, evaporation is the primary control on Cl in the absence of geological sources (e.g., halite) in the study area. Significant evaporite salt deposits are present near Alxa Zuoqi Town (HL8 & HL9), where they are commercially extracted, but although leaching of these deposits into shallow groundwater is locally possible, there is no indication that this would have broader impacts on groundwater salinity. The strong linear relationships with Cl indicate that Naþ and SO2 4 variability is also strongly determined by evaporation rather than by mineral dissolution or 2þ against Cl show ion exchange. Plots of alkalinity (HCO 3 ) and Ca a wide scatter of the data (Fig. 4), but at the low end of the range of alkalinity and Ca2þ values, the concentrations of these ions fall close to the 1:1 line, which suggests that CaCO3 equilibrium controls the HCO 3 concentration. This process may occur during either the recharge process or as a result of CaCO3 precipitation within the aquifer system.

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Fig. 4. Scatterplots of the major ions as a function of the Cl concentration. The straight lines on the graph represent the correlation regression. The data for Na and Mg indicate evaporative concentration of the recharge water. In contrast, the wide scatter for the other ion data suggests a geochemical overprint on the recharge waters.

Table 2 Pearson correlation matrix for the heavy metal concentrations.

Na K Mg Ca Cl SO4 HCO3 NO3

Na

K

Mg

Ca

Cl

SO4

HCO3

NO3

1.00 0.93** 0.83** 0.27 0.98** 0.97** 0.25 0.19

1.00 0.81** 0.30 0.92** 0.92** 0.23 0.05

1.00 0.47* 0.92** 0.94** 0.25 0.10

1.00 0.37 0.39* 0.21 0.20

1.00 0.99** 0.32 0.15

1.00 0.26 0.17

1.00 0.19

1.00

*Correlation is significant at the 0.05 level (two-tailed); **Correlation is significant at the 0.01 level (two-tailed).

A comparison of the ratios of various pairs of ions to the Cl concentration provides further insights into the relative importance of mineral solution interactions (Fig. 5). We compared the trends for the chemistry of the major ions with the stoichiometry of known reactions to provide information on possible minerale solution interactions with the goal of developing a geochemical model for the region’s groundwater. We also compared the data with the marine aerosol ion ratios because differences between the marine and groundwater ratios can provide information on the addition or consumption of ions relative to their value in atmospheric fallout. Molar Na/Cl ratios ranged from 0.76 to 2.65 (Fig. 5), with two surface samples (HL1 in the Helan Mountains reservoir and HL26 in the Yellow River irrigation channel) having the lowest

Z. Ding et al. / Journal of Arid Environments 90 (2013) 77e87

83

Fig. 5. Relationships between major cations and anions. The first four scatterplots of data for various major ions, significantly normalized with respect to the conserved Cl anion, as   a function of the Cl concentration. The other graphs show the relationship between Ca2þ, Mg2þ and SO2 4 , HCO3 . Weak relationships with HCO3 indicate that calcite, dolomite, and gypsum dissolution does not control Ca or Mg abundance.

values, between 0.8 and 1.0. The high Na/Cl ratios in the freshest groundwater were probably controlled by watererock interactions, such as albite weathering:

NaAlSi3 O þ Hþ þ ð9=2ÞH2 O ¼ ð1=2ÞAl2 Si2 O3 ðOHÞ4 þ Naþ þ 2SiðOHÞ4 Deep groundwater sample HL15 (Yaoba) had a low Na/Cl ratio, probably as a result of Na loss or exchange reactions on soil particles and clays. The other cations may have been derived from the dissolution of silicate minerals (e.g., plagioclase feldspar, chlorite, biotite), carbonates (dolomite or calcite), and gypsum, of from cation exchange of Naþ for Ca2þ and Mg2þ on clay minerals (Fig. 5).

Thick carbonate deposits or carbonate cements may also be present in the field, since carbonate dissolution (calcite and dolomite) appears to be another major process that controls the Ca2þ and Mg2þ contents of the groundwater, especially in shallow groundwater. The saturation indices with respect to calcite (SI ¼ 1.25 to 0.82) and dolomite (SI ¼ 2.71 to 2.08) also showed that the most groundwater is saturated with respect to these minerals (Table 1). 4.3. Groundwater evolution based on isotopic data 4.3.1. Stable isotope ratios (d18O and d2H) Fig. 6 shows the relationship between the stable d18O and d2H values for 19 water samples from Table 1 and compares these ratios

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Fig. 6. Stable isotopic ratios for 19 of the water samples in Table 1. Average precipitation ratios for northwestern China (at nine stations) and the Local Meteoric Water Line (LMWL, diagonal line) were obtained from the IAEA GNIP database (www.iaea.or. at:80/programs/ri/gnip).

with the local meteoric water line (LMWL) calculated based on data from the Zhangye station. The results of isotopic measurements for groundwater from the eastern Tengger Desert show the typical pattern of variation for arid regions, with d18O values ranging between 11.2& and 1.5& (with an average of 9.1&), and d2H values ranging between 90.9& and 34.4& (with an average of 75.6&). The average ratio for 9 meteorological stations near the study area produces a LMWL for northwestern China with the following equation: d2H ¼ 7.56d18O þ 5.05 (r2 ¼ 0.97, p < 0.05) (Ma et al., 2009a). Because all of the values from our 19 samples fall well below the LMWL, it is obvious that strong evaporation has altered the original d2Hed18O relationship for precipitation to a slope less than that of the global value of 8 (Craig, 1961) as has been reported in many other arid regions (Darling, 2011; Gat, 1980; Ma et al., 2012). Isotope values in precipitation from the Yinchuan station had a d18O value of 3.9&. The weighted mean value from the Zhangye station, in a representative position in the Hexi Corridor (roughly the line of the ancient Silk Road) about 400 km from our study site, is plotted in Fig. 6 for seven years since 1986, and all but one of the results fall on or close to the LMWL, with a weighted mean composition of 6.5& for d18O and 44& for d2H. The values at this station, which represent a location with the heaviest rainfall in our study area, closely reflect the summer isotopic signatures because the majority of annual precipitation is provided by the East Asian summer monsoon, and these values are therefore most likely to be representative of the values in present-day local rainfall recharge in the Helan Mountains front. The open surface water sampled from the Helan Mountains reservoir is less isotopically depleted than the groundwater, with values of 8.64& for d18O and 73.3& for d2H, which lies near the local meteoric water line and therefore describes precipitation that occurred under cooler mountain conditions with limited evaporation. It is therefore representative of present-day infiltration caused by flash floods or subsurface flows through the alluvial fans flanking the western Helan Mountains. The Zini Lake sample is more enriched in heavy isotopes than the other samples as a result of evaporation from the lake surface, a process that is also reflected in salinities higher than most groundwater values at this site. Shallow groundwater showed enriched d18O values between 6.89 and 9.06& along an evaporation line that indicated kinetic fractionation effects of evaporation from pore moisture in unsaturated

parts of the soil (Clark and Fritz, 1997). HL24 (near Tonghu Lake) and HL27 (near the Mafuxiazi mineral field) are most likely to represent discharge zones for the shallow aquifer at the desert margins given the lower elevation of these samples compared with samples from the mountains, which indicates a strong difference in the hydraulic head between these sample locations and the location of sample HL2 (in the Helan Mountains piedmont); they may therefore be fed by the same recharge source or sources and may be hydraulically connected to the dune field areas. On average, deep groundwater tended to be more depleted (mean d18O of 10&) than shallow groundwater, except for deep water from HL21 (near Toudao Lake), HL22 (near Malian Lake), and HL23 (near an iron mineral industry facility) in the southern desert dune field and except for HL3 and HL5, which are located in the northern part of the Tengger Desert basin. The discrepancy for these five deep wells could result from differences in latitude, since rainfall becomes isotopically lighter with decreasing latitude (Gat, 1980). The samples closer to the mountains fall along an evaporation line with a lower degree of evaporative enrichment, but with a LMWL intercept of 10& for d18O, which is 1e2& heavier than the values from the desert dune field samples. This pattern is consistent with the hypothesis that the Helan Mountains are the primary source of recharge for the desert, and creates the possibility of detecting shifts in meteoric input signals associated with the effects of climatic variability in the mountains. Fig. 7 shows no statistically significant relationship between groundwater 18O depletion and Cl concentration for all of the data combined or for any subset of the data. As an independent indicator of ancient environments and salt constraints, the lower Cl concentration (an average of 260.5 mg/L) implies a long-term sustainability of a humid climate cycle, suggesting that the groundwater retains evidence of water supply during a cooler and wetter period of the ancient climate. 4.3.2. Carbonate chemistry d13C values of dissolved inorganic carbon (DIC) in deep groundwater samples ranged from 4.7 to 10.3& (Table 3). This agrees with the geochemistry data (Fig. 8), which indicate that carbonate dissolution, exchange, mixing, or a combination of the three played a significant role in determining the geochemistry of this water. Calcite saturation indices (Table 1) indicated that groundwater samples from the hypothesized recharge area are saturated with respect to calcite in some cases (e.g., HL10, HL15, HL22). However, these values are probably too enriched in DIC to be

Fig. 7. Relationship between Cl and

18

O.

Z. Ding et al. / Journal of Arid Environments 90 (2013) 77e87

85

Table 3 C isotope results and modeled radiocarbon ages for deep groundwater (100 m below the surface). Sample

HL3 HL5 HL8 HL10 HL11 HL12 HL17 HL18 HL19 HL21 HL22

Modeled age (years) Tamers

Ingerson and Pearson

Fontes and Garnier

IAEA

Evans et al.

Eichinger

963 4730 856 1810 1400 427 11 811 19 157 6320 1602 2339

5311 219 1791 3137 1332 1735 10 680 12 355 1724 350 4370

5447 368 1879 3179 1398 1775 10 650 12 056 2101 380 4456

1061 4081 2332 952 3012 2742 14 888 16 478 2420 4202 510

5752 674 2202 3454 1619 1919 10 444 11 598 2584 500 4858

3082 2150 45 1936 275 455 11 701 15 064 694 792 2670

derived from the dissolution of carbonate minerals alone because the stoichiometry of the reaction involving soil CO2 would result in a d13C of the DIC of around 11  2& (Mook, 1980). Some possible explanations for the relatively 13C enrichment of the groundwater compared with the surface water include incongruent dissolution and re-precipitation of carbonate minerals, methanogenesis, and isotope exchange with aquifer minerals (under open-system conditions with respect to atmospheric CO2 and therefore without much alteration of pH). The source of the DIC, which provides C for precipitation, was determined from the slope of the line [DIC/DIC0]  1 vs. {d13 CeDIC$[DIC/DIC0]  1} (Lojen et al., 2004). The d13C value of added DIC was 10.5&, which corresponds to the dissolution of CO2 from soil with a d13C value of 10.7&  0.15&. The low infiltration rates shown by the elevated Cl concentrations in shallow groundwater may allow for calcite saturation without the presence of significant solid-phase carbonates by allowing enough time for CO2eDIC equilibration during infiltration and by providing a source of Ca2þ from evaporative enrichment in the near-surface region, with possible additional contributions from silicate weathering. In addition, downgradient trends in the carbon isotopic contents of TDIC show heavier d13C values that correspond to diminishing 14 C activities in the deep wells of the study area. Groundwater with more than 50% of carbon after congruent dissolution must be of recent origin, and some of this water is likely to contain inputs of 14 C from thermonuclear weapon tests. The 14C specific activity of older water that has undergone a longer evolutionary process

Fig. 8. Estimation of the source of dissolved CO2 as added dissolved inorganic carbon (DIC).

d13C (PDB)

14

q

Corrected age (yr)

6.9 6.4 8.0 9.6 8.2 9.3 10.3 5.3 4.7 9.5 8.6

59.2 29.7 45.3 63.5 43.8 52.0 12.9 5.4 26.5 45.0 66.6

0.745 0.691 0.872 1.039 0.892 1.008 1.123 0.576 0.514 1.032 0.939

1896 6988 5416 4070 5875 5472 17 887 19 562 5483 6857 2837

C(A0) (pmc)

decreases toward insignificant values as d13C increases. Mixed signatures may occur where water with a high 14C content mixes directly with older water, possibly due to mixture with water following a rapid flow path from recharge into the aquifer (e.g., HL3), or as a result of a faulty borehole lining that gives rise to a bypass flow (e.g., HL8), or it may simply reflect the evolution of an open system (e.g., HL21 and HL22) (Table 3). 4.4. Groundwater residence time:

14

C age correction

Carbon-14 age dating of groundwater is complicated by inputs of dead carbon from reactions with carbonate minerals and organic matter, which can affect the DIC. Groundwater residence times can be estimated based on 14C activity if the dilution of the 14C by inactive C in the aquifers can be evaluated. Many correction models have been proposed based on chemical evolution or isotope dilution of 13C/12C derived from soil CO2 through interactions with aquifer carbonates. The groundwater ages calculated based on the models of Ingerson and Pearson (1964), Evans et al. (1979), Fontes and Garnier (1979), and Eichinger (1983) are to some extent in good agreement and provide similar solutions. Table 3 summarizes the results of age calculations with several models. For the corrections we used in these models, we assumed that the 14C values (which ranged from 5.4 to 66.6 pmc) and the d13C values (which ranged from 4.7 to 10.3) for the carbonate minerals were 0 pmc and 0&, respectively, whereas the corresponding corrections for the soil gas phase were 100 pmc and 23&, respectively (typical values for soil-zone CO2; Geyh, 2000). The groundwater ages provided by the Tamers and IAEA models were generally older than the Ingerson and Pearson, Fontes and Garnier, and Evans models, which gave similar ages. However, the Tamers model and the Ingerson and Pearson model both assume simple mixing between soil CO2 and carbonate minerals, and that isotope exchange does not occur, but the Ingerson and Pearson model uses d13C data to establish the amount of carbonate mineral weathering. The Fontes and Garnier model also uses d13C data to establish the amount of carbonate mineral weathering, and also accounts for isotope exchange. Given the potential errors, there appears to be little difference between the Ingerson and Pearson, Fontes and Garnier, and Evans models, which gave similar mean initial 14C activity values of 36.5, 36.1, and 35.1, respectively. Further work, including identifying other possible reactions that could have diluted the 14C activities, using different models for age correction, and geochemical modeling, would be required to better constrain the residence times. Although considering potential errors, it is more sophisticated for F&G Model to accurately reflects the groundwater ages for the samples of study areas. Based on the Fontes and Garnier groundwater ages, the oldest groundwater that

86

Z. Ding et al. / Journal of Arid Environments 90 (2013) 77e87

we sampled was at HL18 (5.4 pmc), with a corrected age of approximately 12 056 years. However, groundwater ages are expected to be greater farther along the flow path. Negative ages calculated for HL10 (3179 yr, 63.5 pmc) are most probably due to higher initial activities (A0) as a result of nuclear weapons testing since the 1950s, and indicates that “modern” water recharge has occurred since this time. HL10 is the first sampled well along the flow path, and lies closest to the recharge zone. In addition, HL3 (59.2 pmc) and HL22 (66.6 pmc) showed the properties of “modern” groundwater but are more likely to have originated from other sources. A modified version of the Pearson (1965) isotopic correction model developed by Clark and Fritz (1997) can be used to account for isotopic dilution as a result of incongruent dissolution of carbonates (Gates et al., 2008; Ma et al., 2010). The radiocarbon dilution factor q can be calculated based on the change in DIC d13C between the sample and the recharge zone:

q ¼

d13 CDIC  d13 Ccarb d13 Crech  d13 Ccarb

where the subscripts carb and rech represent solid-phase carbonate and recharge water, respectively. The mean of the calculated q values (0.76; Table 2) is intermediate between the ranges reported by Vogel (1970) for carbonate aquifers (0.65 < q < 0.75) and sedimentary basins with trace carbonates (0.75 < q < 0.9). However, the extreme q values for some of the water samples (more than 1 or less than 0.65) may have resulted from differences in climate, vegetation, and geographical factors, so the modern 13C-substrate value may not serve as a good model for the input value of the recharge source. We estimated residence times of 17.9 ka and 19.6 ka for samples HL17 and HL18 (in the desert areas) for the desert areas, suggesting that some replenishment of the desert aquifers occurred during the late Pleistocene and early Holocene, whereas some of the water at upstream positions in the flow path represents modern water. The modern (<50 yr) groundwater that mixed with older water in parts of the basin (e.g., HL22) probably resulted from irrigation returns. 5. Conclusions As an important part of the eastern Tengger Desert of northwestern China, the piedmont and plains areas west of the Helan Mountains are a representative area that can be used to examine stresses on the groundwater resources in this region. In response to the need for a sustainable regional groundwater management strategy, we performed the present study to investigate the groundwater quality, origins, and subsequent evolution in this region using hydrogeological and hydrogeochemical methods. The groundwater chemical compositions exhibited a transition moving along the flow path from the mountains into the desert basin. TDS increased along the flow paths, with an average value of 902.4 mg/ L. The groundwater was generally rich in Naþ and Kþ, but had no dominant anion type. The chemical composition of the groundwater and the relationships among the major ions indicate that direct infiltration of precipitation is not the major source of recharge for the plains and desert areas, which are strongly influenced by evaporation during the dry seasons and subsequent dissolution of minerals during the rainy season recharge. Other processes such as cation exchange and weathering contributed to the groundwater composition. The groundwater isotopic measurements showed a pattern of variation that is typical of arid regions, with d18O values ranging between 11.2& and 1.5&, and d2H values ranging between 90.9& and 34.4&. A comparison of the stable isotope

compositions of groundwater with local values in modern precipitation suggested that direct infiltration of precipitation is not an important source of groundwater recharge in the desert parts of the study area. The relationship between d18O and the Cl concentration suggests that the least isotopically depleted waters indicated wetter conditions in the past during recharge. d13C values of dissolved inorganic carbon (DIC) in deep groundwater samples ranged from 4.7 to 10.3&, which indicates that groundwater with more than 50 pmc after congruent dissolution must be of recent origin, and some of these samples may also contain inputs of 14C from thermonuclear weapon tests after the 1950s. The estimated residence times of 17.9 ka and 19.6 ka at sites HL17 and HL18 in the desert suggests that some replenishment of the desert aquifers occurred during the late Pleistocene and early Holocene, whereas some of the water at upstream positions in the flow path represents modern water. Acknowledgments This research was supported by the National Science Foundation of China (No. 41102145, 40872161) and the Keygrant Project of Chinese Ministry of Education (No. 310005). We thank Jianhua He, Hui Zhang, and Kunpeng Zhou for assistance with the fieldwork and laboratory analysis. We also thank the anonymous reviewers for their comments that have led to improvements in the paper. References AZWCB, 2002. Water-resources Investigations Report in Alxa Zuoqi. Alxa Zuoqi Water Conservancy Bureau, Alxa Zuoqi. Chen, F.H., Wu, W., Holmes, J., Madsen, D.B., Zhu, Y., Jin, M., Oviatt, J.G., 2003. A midHolocene drought interval as evidenced by lake desiccation in the Alashan Plateau, Inner Mongolia, China. Chinese Science Bulletin 48 (13), 1e10. Clark, I.D., Fritz, P., 1997. Environmental Isotopes in Hydrogeology. Lewis Publishers, Boca Raton. Craig, H., 1961. Isotopic variations in meteoric waters. Science 133 (3465), 1702e 1703. Darling, W.G., 2011. The isotope hydrology of quaternary climate change. Journal of Human Evolution 60 (4), 417e427. Edmunds, W.M., Ma, J., Aeschbach-Hertig, W., Kipfer, R., Darbyshire, F., 2006. Groundwater recharge history and hydrogeochemical evolution in the Minqin Basin, North West China. Applied Geochemistry 21 (12), 2148e2170. Eichinger, L., 1983. A contribution to the interpretation of 14C groundwater ages considering the example of a partially confined sandstone aquifer. Radiocarbon 25, 347e356. Evans, G.V., Otlet, R.L., Downing, R.A., Monkhouse, R.A., Rae, G., 1979. Some problems in the interpretation of isotope measurements in United Kingdom aquifers. Isotope Hydrology (Proc. Symp. Vienna, 1978) 2, 679e708. IAEA-SM228/34. Fontes, J.C., Garnier, J.M., 1979. Determination of the initial 14C activity of the total dissolved carbon: a review of the existing models and a new approach. Water Resources Research 15, 399e413. Gat, J.R., 1980. The isotopes of hydrogen and oxygen in precipitation. In: Fritz, P., Fontes, J.-Ch. (Eds.), Handbook of Environmental Isotope Geochemistry. The Terrestrial Environment, vol. 1. Elsevier, Amsterdam, pp. 21e47. Gates, J.B., Edmunds, W.M., Ma, J., Scanlon, B.R., 2008. Estimating groundwater recharge in a cold desert environment in northern China using chloride. Hydrogeology Journal 16, 893e910. Geyh, M.A., 2000. An overview of 14C analysis in the study of groundwater. Radiocarbon 42, 99e114. Ingerson, E., Pearson, F.J., 1964. Estimation of age and rate of movement of groundwater by the 14C method. In: Miyake, Y. (Ed.), Recent Researches in the Fields of Hydrosphere, Atmosphere and Nuclear Geochemistry. Maruzen, Tokyo, pp. 263e283. Lojen, S., Dolenec, T., Vokal, B., Cukrov, N., Mihel ci c, G., Papesch, W., 2004. C and O stable isotope variability in recent freshwater carbonates (River Krka, Croatia). Sedimentology 51, 361e375. Love, A.J., Herczeg, A.L., Armstrong, D., Stadter, M.F.F., Mazor, E., 1993. Groundwater flow regime within the Gambier Embayment of the Otway Basin, Australia: evidence from hydraulics and hydrochemistry. Journal of Hydrology 143, 297e 338. Ma, J.Z., Wang, X.S., Edmunds, W.M., 2005. The characteristics of ground-water resources and their changes under the impacts of human activity in the arid northwest China: a case study of the Shiyang River Basin. Journal of Arid Environments 61, 277e295.

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