geochemical characteristics and circulation of geothermal fluids in

International Journal of Applied Engineering Research ISSN 0973-4562 Volume 11, Number 18 (2016) pp 9364-9376 © Research India Publications. http://www.ripublication.com

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Geochemical characteristics and circulation of geothermal fluids in the

southern Xining basin, NE Tibetan Plateau, China

Huidi Lia,b, Yue Zhaob, Senqi Zhangc and Tianshui Yangd

aGeological Publishing House, Beijing, 100083, China.

bLaboratory of Crustal Deformation and Process, Institute of Geomechanics, Chinese Academy of Geological Sciences, Beijing, 100081, China.

cQinghai Institute of Geological Survey, Xining, 810012, China. dChina Geosciences University, Beijing, 100081, China.

Abstract

Chemical and isotopic investigation on thermal and cold waters has been carried out. Chemical analysis reveals the waters from the Yaoshuitan (YST) and Ping’an (PA) geothermal systems are of the HCO3 type, and those from the Xining (XN) geothermal system are of mixed character (Cl and SO4). The results from different geothermometers indicate that reservoir temperatures in the YST, XN, and PA geothermal systems are about 49~67°C, 61~102°C, and 45°C, respectively. An absence of tritium indicates the age of geothermal waters (except for sample 7) from the YST geothermal system is more than 50 years. The other geothermal waters contain tritium, implying mixing with shallow cold groundwater.

Keywords: Geothermal water, Geothermal system, Geochemistry, Geothermometry, Isotope

INTRODUCTION

Geothermal energy is the heat energy present inside the earth surface in the form of hot springs, fumaroles, volcanoes, and geysers [1, 2]. Geothermal source, clean energy source [3] and renewable enery [4, 5, 6], can be applied by the ground source heat pump [7]. High temperature geothermal resources in the southwest of China are mainly located in the Tibet-Yunnan geothermal zone extending from western Tibet to the Tengchong geothermal field in Yunnan province. The high temperature geothermal zone, with a heat flow of 61 ~ 140 mW/m2 [8], is located at the boundary of the Indian and Eurasian plates [9, 10], and has been studied in detail [9, 11, 12, 13]. In China, low-to-medium temperature geothermal resources are primarily used due to the limited distribution of high temperature geothermal resources. However, few studies have been carried out on such resources [14, 15]. The Xining basin with a heat flow of 32 ~ 75 mW/m2, located at the intra-plate, has one of the greatest potentials for being a low-to-medium temperature geothermal resource [8, 16]. Zhang et al. first elucidated the geothermal anomalies of the Xining basin, and proposed that they correlate with Tertiary folds and faults [17]. Wang et al. studied the formation mechanism, distribution and burying of geothermal anomalies of the Laji mountain front, and the results revealed that there is a close relationship between geothermal anomalies and

tectonics [18]. Tang and Qin further investigated geothermal resources of the Xining basin and divided them into four sub-regions, namely the Shuangshu depression, the Zongzhai-Dabuzi depression, the Caojiagou-Yunjiakou depression, and the Xining uplift. They proposed that the Zongzhai-Dabuzi depression is suitable for geothermal exploitation due to the great thickness, shallow depth and good penetrability of the thermal reservoir [19]. Based on the results provided by Tang and Qin[19], a new geothermal well (DR2005), was drilled at Dujiazhuang village in 2005[20]. Yu et al. assessed geothermal and mineral waters of the Yaoshuitan (YTS) geothermal system, and confirmed they are suitable for medical use and drinking. Although geothermal exploitation has a history of only two decades, geothermal water, currently used for bathing and spas, is badly wasted. Because the hydrothermal system is very sensitive to the conditions of exploitation, intensive exploitation has led to progressive reduction of flow in natural hot springs in the YST geothermal system[21]. It should be noted that most of the geothermal investigations in the Xining basin are mainly aimed at geothermal usage, whereas studies focusing on the nature and evolution of geothermal water (such as the age, origin, water and rock interaction or mixing between geothermal water and shallow groundwater) have been very limited. In addition, the Xining basin has a continental arid and semi-arid climate with a mean annual temperature of about 6 °C. The mean annual precipitation is only 386 mm (precipitation mainly occurs in July, August, and September, and accounts for 60% of annual rainfall), whereas the mean annual evaporation is 1763 mm. Due to this substantial water shortage, the sustainability of a hot water supply has become an important issue for the local government and residents. To better understand the nature and evolution of the geothermal systems, as well as to develop a sustainable management plan for the geothermal resources of the southern Xining basin, we have carried out chemical and isotopic analyses from well and spring waters in the YST, Xining (XN), and Ping’an (PA) geothermal systems.

GEOLOGICAL SETTING

The Xining basin, with an area of 9500 km2, is located at the northeastern margin of the Tibetan Plateau, which is the highest mountainous region induced by the India-Eurasia collision. This basin spreads over an area from the Daban

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mountain in the north to the Laji mountain in the south, and from the Riyue mountain in the west to a small upheaval in the east (Fig. 1a). Fault structures are presently the most conspicuous features in the Xining basin, and there are many active faults with strikes of NNE, NNW, NS, and NW. The Lajishan, Haiyan and Dabanshan faults are the main deep faults, with the others being shallower (Fig. 1b). Tectonic evolution of the Xining basin can be divided into three stages: (1) an intracontinental thrust and overthrust orogenic stage, which is characterized by crust shortened since the late Triassic; (2) a strike-slip orogenic stage, which

is characterized by a large-scale strike-slip fault and basin formed from the Jurassic to the Tertiary period; and (3) an uplift orogenic stage, which is characterized by mountain uplift and basin subsidence from the Pleistocene to the present. The Xining basin is covered mainly by Quaternary sediments; Tertiary and Proterozoic rocks; and Cretaceous, Jurassic and Triassic rocks (Fig. 1b). Rocks encountered in boreholes include sandstone, siltstone, limestone, dolomite, mudstone, granite, schist, quartzite and sandy conglomerate [16, 18].

Fig. 1: Location of the Xining basin in China; (a) Contour of water pressure of the Xining basin; (b) Geological and structural map of the Xining basin. 1. Quaternary; 2. Neogene; 3. Eogene; 4. Cretaceous; 5. Jurassic; 6. Triassic; 7. Proterozoic; 8. Sampling site; 9. Thrust faults; 10. Strike-slip fault; 11. Buried fault.

(a) (a)

(b)


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GEOTHERMAL SYSTEMS OF THE SOUTHERN

XINING BASIN

There are several geothermal systems in the Xining basin. Three of the systems, namely the YST, XN and PA geothermal systems, are located in the southern Xining basin and were selected for our study.

YST geothermal system

The YST geothermal system, with an area of about 800 m2, is located 1 km southwest of Huashan village and is tectonically situated at the southern edge of the Xining basin. Extensional faults (F1 and F2) with a strike of NNE and a compressive-torsional fault (F3) with a strike of almost SN are the dominant structures in this area (Fig. 2a). The outcrop of the YST geothermal system is sericite phyllite, sandstone, limestone, dolomite and slate, and borehole information revealed the main rock of the thermal reservoir is Proterozoic dolomitic limestone and dolomite in the YST geothermal system [18].

XN geothermal system

The XN geothermal system lies in the south part of the Xining basin. Samples were taken from the two thermal water wells of 8701 and DR2005. Well DR2005, with a depth of 1600 m, is located at Dujiazhuang village. Three faults exist near Well

DR2005 (Fig. 2b). The Zongzhai fault (F4) with a strike of NW is a deep fault certificated by geophysical survey. The active fault (F5) along Quanerwan and Xiejiazhai villages is well developed and substantiated by geophysical survey and remote sensing technology [20]. The F6 fault, located at Xinzhuang village, was identified by the difference in lithology. There are some ascending thermal springs which group along the F5 fault, indicating the emergence of geothermal fluid. Paleo-travertine deposits between the F4 and F5 faults indicate paleo-hydrothermal activity. Well DR2005 is located between the F4 and F5 faults, and its main aquifer is Cretaceous-Jurassic sandstone that has a thickness of 272 m. Wellhead and bottom-hole temperatures are 62.5 and 69.3C, respectively, and the temperature gradient is 3.9C/100 m [20]. Well 8701, with a depth of 758 m, is located at the Shengli hotel in Xining city. There are two faults including F7 with a strike of NNW and F8 with a strike of NNE in the south of the well (Fig. 2b). Well 8701 has a mean temperature gradient of 5.6C/100 m. Its main aquifer is Jurassic-Cretaceous sandstone and sandy conglomerate with a thickness of 180 m [16].

Fig. 2: Location and structural map of the sampling sites. (a) YST geothermal system; (b) XN geothermal system; (c) PA geothermal system. Figs. 3a, 3b and 3c are located about 1 km southwest of Huashan village, between Xining and Dujiazhuang, and between Binglingshan and Sanhe in Figure 1.


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PA geothermal system

The PA geothermal system lies about 18 km southwest of Ping’an county. Samples were taken from Well ZK10, Binglingshan spring and the river that are situated at the towns of Sanhe, Binglingshan and near Qijiachan, respectively (Fig. 2c). The fault (F9) with a strike of NNE has been well developed along the Qijiachuan River. According to the obvious difference in lithology, a horst and a graben are identified to the east and west of the Qijiachuan River, respectively. A compressive-torsional fault (F10) with a strike nearly EW has developed to the north of Sanhe, and F11 and F12 faults with a strike NNW exist near Binglingshan (Fig. 2c). Well ZK10 has a depth of 212 m, and the main rock of the

aquifer is Cretaceous sandstone of about 89 m thickness, having a permeability coefficient of 0.046 m/d. In summary, occurrence of thermal springs and well locations in the study area are always linked to faults. Precipitations, percolating to the reservoirs, are heated at depth and ascend to the surface through major faults (Fig. 2). The heat source is a high geothermal gradient. According to the contours of water pressure, water flows toward the center from the northern, western, and southern margins of the Xining basin (Fig. 1a). The conceptual model in Fig. 3 expresses the relationship between recharge, travel, and discharge of geothermal water in the Xining basin.

Fig. 3: Conceptual model of geothermal reservoir of the study area; 1. Quaternary; 2. Paleogene and Neogene; 3. Cretaceous; 4. Jurassic; 5. Lower Proterozoic; 6. Upper Proterozoic; 7. Heat flow; 8. Flow direction of groundwater; 9. Thrusting fault; 10. Spring; 11. Caprock; 12. Geothermal reservoir.

SAMPLING AND ANALYSES

Water samples from 16 sampling sites (four hot water wells, six hot springs, one river, and five cold springs) were collected from the three geothermal systems. During the sampling procedure, all sampling bottles were first washed three times using sample water. Filtered (0.45 µm) waters were collected and stored in 1 L polyethylene bottles for determination of cations and SiO2. Untreated waters were collected using polyethylene bottles for analyzing anion and isotopic compositions, and isotopic results were used to study the origin and characteristics of waters. Chemical analyses were carried out using: Inductively Coupled Plasma for Na, K, Ca, Mg, Li, SiO2 and B; Inductively Coupled Plasma - Mass Spectrometry for Sr, Br

and I; a titration method for Cl, F, HCO3, and SO4 at the National Research Center for Geo Analysis. Tritium analysis was performed at the State Key Laboratory of Earthquake Dynamics (SKLED), Institute of Geology of China Earthquake Administration. Stable isotope analyses were performed using mass spectrometry (MAT 253 EM) and were related to the Vienna Standard Mean Ocean Water (V-SMOW) standards. Oxygen isotopic ratios (18O/16O) of waters were determined by equilibrating the samples with CO2, and 2H compositions were determined from the H2 generated by the zinc-reduction method [22]. The precision for 2H and 18O are 2.0 and 0.2, respectively. Tritium was determined on electrolytically enriched water samples by low-level


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proportional counting, and the results are reported as 3H units. The locations of sampling sites are shown in Fig. 2.

WATER CHEMISTRY

Hydrogeochemical characteristics

The major goal of geochemical exploration is to obtain the subsurface composition of the fluids in a geothermal system and to use this to obtain information on temperature, origin, and flow direction [23, 24]. Chemical compositions of all samples are summarized in Table 1. Discharge temperatures of thermal springs (and wells) vary from 18 to 62.5°C, while those of cold springs range between 3.5 and 6.5°C and are close to the average annual air temperature at the discharge elevation. The total dissolved solids (TDS) of the cold and geothermal waters are 210 ~ 403 mg/L and 1041 ~ 34882 mg/L, respectively. The results of pH measurements show that all geothermal waters are close to neutral. Chemical analysis shows that the major cations are Ca, Mg, and Na, and the major anions are HCO3, SO4, and Cl. A Piper diagram plot (Fig. 4) indicates that Ca and Mg are the dominant cations and HCO3 is the main anion for samples from the YST and PA geothermal systems, and that Na is the dominant cation and Cl and SO4 are the main anions for samples from the XN geothermal system. Samples from the YST and PA geothermal systems are HCO3 waters (Ca-HCO3, 8; Ca-Mg-HCO3, 1, 2, 3, 4, 6, 7, and 11; Mg-Ca-HCO3, 5, 9, and 10; Ca-Na-Mg-HCO3, ZK10). Samples 8701-1 and DR2005 are Na-Cl-SO4 in composition, and sample 8701-2 is a Na-Cl-SO4-HCO3 type. All waters contain high Mg concentrations, implying they are

immature [25, 26]. In addition, Cl-rich geothermal waters have high Cl/Br ratios, suggesting they have likely interacted with evaporation [27]. Figure 5 shows a triangular diagram (SO4-Cl-HCO3). Based on this diagram, the waters can be classified into four types: (a) mature type (Cl rich); (b) steam heated water (SO4 rich); (c) peripheral water (HCO3 rich); and (d) volcanic water (Cl-SO4 rich) [28, 29, 30]. Samples with relatively low TDS from the YST geothermal system are located in the HCO3-corner as peripheral water, indicating short and shallow circulation [24, 29]. The relatively higher B in geothermal waters than in the cold spring waters suggests geothermal waters have a deeper circulation than cold waters [26, 29]. The chemical characteristics of the YST geothermal waters are consistent with the first stages of interaction between meteoric waters and rocks [31, 32, 33]. Waters with high TDS from wells 8701 and DR2005 are Cl-SO4 type, suggesting a mixture of Cl- and SO4-rich waters. Figure 6 illustrates the relationships of Na, Ca, Mg, HCO3, SiO2 and TDS versus Cl for the YST samples. Sodium, Ca, Mg, HCO3, SiO2 and TDS have a positive linear correlation with Cl. If the well sample is not considered (where geothermal water exploitation has led to a hydraulic change), the results for sample 7 always lie between those for cold and geothermal waters. This result indicates that sample 7 is likely a mixture of cold and geothermal waters, which is in agreement with the isotopic results discussed below.

Table 1: Chemical composition of all samples from the southern Xining basin (mg/L)

Sample

number Typea Locationb Date t(C) C pH Na K Ca Mg HCO3 SO4 Cl F B Sr Li Br I SiO2 TDS

ZK10 W SH 26/04/2003 18 6.84 325 76.8 784 156 2993 948.7 68.18 0.7 0.26 0.63 0.088 11.78 5302

8701-1 W SLH 21/9/1987 39.5 7.10 10920 64.5 391.7 171 1359 14110 7840.2 0.4 21.1 34882

8701-2 W SLH 22/7/2002 39.5 7.35 8300 50.0 271.4 112 1576 10791 5232.0 4.2 1.51 0.17 18.94 27662

DR2005 d W DJZ 28/09/2005 62.5 7.56 11100 124 270.1 78.3 1414 14200 6983 1.36 6.1 1.7 7.51 0.47 29.65 34200

1 W YST 06/06/2006 40 6.53 15.6 3.52 268 87.2 1178 38.2 12.1 0.3 0.41 0.51 0.029 0.04 0.007 21.15 1041

2 HS YST 07/06/2006 18.2 6.45 13.1 3.01 457 94.6 1774 43.4 12.9 0.4 0.35 0.49 0.026 0.04 0.008 19.08 1536

3 CS YST 07/06/2006 6.5 7.69 2.61 0.41 46.1 21.6 233 8.98 7.03 0.06 <0.02 0.07 0.001 0.01 0.001 4.98 210

4 HS YST 07/06/2006 18.1 6.28 11.8 2.64 407 98.9 1631 56.6 11.3 0.41 0.27 0.5 0.021 0.03 0.007 22.15 1433

5 CS YST 07/06/2006 5.5 7.1 4.39 1.05 74.9 55 461 23.6 6.25 0.02 <0.02 0.29 0.004 0.01 <0.001 5.94 403

6 HS YST 07/06/2006 22.5 6.31 15.7 3.56 491 94.3 1862 38.6 14.1 0.41 0.4 0.58 0.032 0.04 0.008 20.62 1615

7 HS YST 07/06/2006 21.5 6.12 9 1.86 309 88 1291 27.1 12.9 0.3 0.19 0.34 0.015 0.02 0.005 13.39 1111

8 HS YST 07/06/2006 27.3 6.55 15.6 3.73 509 105 1989 40.1 10.6 0.48 0.4 0.58 0.032 0.04 0.009 21.31 1707

9 CS YST 07/06/2006 5.3 7.67 5.4 0.9 43 49 323 42 7.03 0.05 <0.02 0.23 0.004 0.01 <0.001 5.20 316

10 CS YST 07/06/2006 3.5 7.54 4.26 0.86 43.9 49.4 332 36.5 6.64 0.04 <0.02 0.19 0.003 0.01 <0.001 4.38 313

11 CS YST 07/06/2006 6.5 7.3 3.64 0.79 81 23.3 334 13.6 8.2 0.05 <0.02 0.16 0.002 0.01 0.002 6.37 306 a W: well; HS: hot spring; CS: cold spring. b YST: Yaoshuitan; SH: Sanhe; SLH: Shengli hotel; DJZ: Dujiazhuang.


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Fig. 4: Piper triangular diagram for the major ions.

Fig. 5: Relative concentrations of Cl, SO4, and HCO3 for all samples.


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Fig. 6: Plots of Cl vs. Na, Ca, Mg, HCO3, SiO2 and TDS for YST samples.

Mineral saturation indices

Better estimation of reservoir temperature can be achieved by simultaneously considering the equilibrium state between a specific water and many hydrothermal minerals as a function of temperature [34]. The saturation index (SI) is used to describe the equilibrium state of certain minerals. Saturation indices are based on 10 logarithmic values with an SI of zero indicating exact saturation, negative values corresponding to undersaturation, and positive ones to oversaturaton [14]. SIs of hydrothermal minerals that are possibly present in the geothermal systems are calculated by the PHREEQC computer code [35], and results are given in Table 2. Anhydrite, gypsum, and halite are undersaturated in all waters, and quartz is oversaturated. Talc is undersaturated in most waters, whereas other mineral such as aragonite, calcite and dolomite are oversaturated.

Meteoric waters acquire saturation with calcite and Ca-HCO3 composition in the initial stages of interaction with rocks containing even small amounts of calcite [31]. The reason is that, at temperatures close to 25°C, the dissolution rate of calcite is two to six orders of magnitude higher than that of Al-silicates, depending on the pH [34]. Meteoric waters acquire Mg-HCO3 composition through interaction with rocks [36], and the main rocks in the study area containing MgO are dolomite and dolomitic limestone [18]. Ca and Mg concentrations are much higher in the geothermal waters than in the cold spring waters (Table 1), implying that water-rock interactions are more significant in geothermal waters than in cold waters. The prevalence of Mg and Ca among dissolved cations is consistent with the chemical and mineralogical characteristics of rocks, as well as with the high dissolution rates of the minerals involved. Low Ca and Mg content in the cold spring waters suggests that dissolution of Ca and Mg minerals has occurred to a minor extent [14].


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Table 2: Calculated PCO2 and saturation index (SI) values for waters from geothermal systems of the southern Xining basin using the PHREEQC program (1999)

Sample

number

Anhydrite Aragnite

onite

Calcite Dolomite Gypsum Halite Quartz Talc Log(PCO2)

ZK10 -0.50 1.06 1.21 1.98 -0.26 -6.35 0.37 -4.1 -0.51 8701-1 -0.18 0.55 0.68 1.42 -0.05 -2.96 0.37 -0.04 -1.04 8701-2 -0.37 0.73 0.87 1.76 -0.24 -3.24 0.30 0.69 -1.22

DR2005 -0.16 1.07 1.19 2.20 -0.21 -3.07 0.22 4.07 -1.37 1 -1.96 0.43 0.56 1.13 -1.81 -8.40 0.29 -2.70 -0.42 2 -1.83 0.40 0.55 0.67 -1.59 -8.41 0.57 -5.78 -0.31 3 -2.98 -0.16 0 -0.27 -2.73 -9.26 0.18 -3.45 -2.44 4 -1.74 0.15 0.30 0.25 -1.50 -8.51 0.64 -6.47 -0.18 5 -2.48 -0.32 -0.17 -0.43 -2.23 -9.10 0.28 -5.75 -1.57 6 -1.87 0.36 0.51 0.62 -1.64 -8.31 0.54 -6.00 -0.13 7 -2.11 -0.14 0.01 -0.22 -1.88 -8.57 0.36 -7.96 -0.09 8 -1.85 0.70 0.84 1.38 -1.64 -8.45 0.48 -3.85 -0.32 9 -2.42 -0.13 0.03 0.14 -2.16 -8.95 0.22 -2.67 -2.29 10 -2.46 -0.27 -0.11 -0.18 -2.21 -9.08 0.18 -3.99 -2.15 11 -2.63 -0.18 -0.02 -0.53 -2.37 -9.06 0.29 -5.34 -1.9

Chemical geothermometry

Cationic and silica geothermometers can be used to estimate the reservoir temperatures [37, 38, 39, 40, 41, 42, 43, 44, 45, 46, 47]. The temperatures obtained using geothermometers are given in Table 3. Previous studies showed that a geothermometer based on the simple Na/K ratio must be used with extreme caution [35]. When the

NaCa MM /)( value is larger than 1, the

temperature estimated by a Na/K geothermometer would be much higher than the actual temperature. For the YST and PA thermal waters, the

NaCa MM /)( values of 3.8~7.1 are larger

than 1, and the temperatures estimated by the Na/K geothermometer vary from 287 to 313°C. This range of values is clearly higher than the measured temperature (Table 3), suggesting that the Na/K temperatures for the YST and PA thermal waters are unreasonable. For the XN thermal waters, the

NaCa MM /)( values range from 0.005 to 0.3, and the

temperatures estimated by the Na/K geothermometer ranges from 62~102°C. These values generally agree with the measured temperatures, suggesting that the Na/K temperatures for the XN thermal waters may be reasonable [40]. Assuming equilibrium with quartz at depth, temperatures can be estimated from the silica content of waters discharged at the surface. The silica geothermometer, which was proposed by Fournier [38] based on the solubility of quartz, has been widely used to estimate subsurface temperatures of less than 250°C in hot spring systems (e.g., [32]). Minissale et al. [48] suggested that quartz is still the SiO2(aq)-controlling phase for deep solutions that do not re-equilibrate to the emergence temperature, and a SiO2 (quartz) geothermometer is generally applicable to geothermal systems at pH < 10 and without steam loss [49]. The temperatures estimated by the quartz

geothermometer for the YST, XN, and PA thermal waters are 49~67°C, 61~79°C, and 45°C, respectively. For the XN thermal waters, the quartz temperatures are well consistent with those estimated by the Na/K geothermometer, suggesting that the quartz geothermometer tends to be reliable in geothermal systems hosted in quartz bearing rocks, even at low temperatures. Considering that the mixing, as indicated by tritium results (see Section 6.3), has occurred in the samples 7, DR2005 and ZK10, and that it can affect the quartz temperatures, the actual temperatures for these three samples should be higher than the silica geothermometer temperatures.

Table 3: Temperature (in C) estimated by chemical geothermometers

Sample Measurement temperature(C) t(Na-K(G)) t (Na-K(V)) t (Quartz)

8701-1 39.5 76 62 65 8701-2 39.5 77 63 61

DR2005 62.5 102 88 79 ZK10 18 312 302 45

1 40 307 297 66 2 18.2 309 299 62 3 6.5 271 260 18 4 18.1 306 296 67 5 5.5 313 303 23 6 22.5 307 297 65 7 21.5 298 287 49 8 27.3 313 303 66 9 5.3 277 265 19 10 3.5 295 285 15 11 6.5 303 293 25

t (Quartz): Fournier (1977); t (Na-K(G)): Giggenbach (1988); t (Na-K(V)): Verma and Santoyo (1997).


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ISOTOPE RESULTS

Origin of geothermal waters

Stable isotope compositions are reported in Table 4. 2H and 18O values range from -58‰ to -47‰ and from -9.0‰ to -7.9‰ for the YST cold waters, from -66‰ to -56‰ and from -10.5‰ to -9.4‰ for the YST geothermal waters, -74‰ and -10.4‰ for the XN geothermal waters, and -73.49‰ ~ -75‰ and -10.55‰ ~ -10.6‰ for the PA geothermal waters. Isotopic data can adequately differentiate three possible origins of geothermal water: magmatic, oceanic and meteoric. The ranges of 18O and 2H for all the samples were -7.76‰ ~ -10.6‰ and -47‰ ~ -75‰, respectively. Our data do not show the presence of any significant amount of magmatic water, which generally has 18O of +6‰ ~ +9‰ and 2H of -40‰ ~ -80‰ [50]. The possibility of an oceanic origin is limited by the absence of highly enriched 18O, 2H and Cl values. Therefore, the origin of geothermal waters appears to be meteoric. 18O vs. 2H is plotted in Fig. 7, which shows stable isotope compositions along with the Global Meteoric Water Lines (GMWL) [51], further supporting meteoric water being the origin. Moreover, Fig. 7 reveals 2H and 18O are lighter in geothermal waters than in cold waters, seemingly the result of age and climate effects [52, 53]. Figure 7 indicates some cold waters clearly originate from modern precipitation, while geothermal waters are derived from the infiltration of precipitation. The relation between altitude and 18O and 2H values has also been demonstrated [54], and effects of climate and age are discussed below (Sections 6.2 and 6.3).

Table 4: Isotope data and ages for all samples from the geothermal systems of the southern Xining basin

Sample number t (C)

2H(‰) 18O(‰) 3H(TU)

ZK10 18 -75 -10.6 4.9±1.1

Binglingshan 19.5 -73.49 -10.55

DR2005 62.5 -74 -10.4 3.93±2.30

Qijiachuan river a -51.219 -7.76

1 40 -65 -10.4 ND

2 18.2 -62 -9.9 ND

3 6.5 -58 -9 28.69±3.17

4 18.1 -56 -9.4 ND

5 5.5 -47 -8.1 8.41±2.96

6 22.5 -65 -10.5 ND

7 21.5 -61 -9.8 5.51±2.89

8 27.3 -66 -10.3 ND

9 5.3 -49 -8.5 8.97±2.96

10 3.5 -48 -8.3 16.00±3.06

11 6.5 -53 -7.9 0.29±2.87 ND: not detected

Fig. 7: Plot of δ2H and δ18O for all samples.

Recharge altitude and temperature of geothermal waters

The 18O value and the altitude of precipitation infiltration are related by

G PH hK

(1)

Where H and h represent the altitudes of precipitation infiltration and the sampling site in meters above sea level (a. s. l.), respectively, G (‰) and P (‰) are 18O the values of geothermal water and precipitation, respectively, and K (‰/100 m) the altitude gradient of the isotope. The values of P and K in the study area are -6.9 [55] and -0.25 [13], respectively. Based on the above equation, the recharge altitude should be about 3480 ~ 3675 m a.s.l. for the geothermal waters from the YST geothermal system, 3200 ~ 3460 m a.s.l. for the cold waters from the YST geothermal system, about 3091 ~ 3108 m a.s.l. for the PA geothermal system, and 3132 m a.s.l. for the XN geothermal system. These results, combined with water flow directions (Fig. 1a), reveal that the YST and PA geothermal waters were recharged by precipitation from near the Laji Mountain, and that the XN geothermal waters might be mainly derived from near the Laji, Niuxin and Daban mountains. Based on the relation between 18O values and mean annual surface air temperature in Xining, 18O = 0.5579t - 13.393 (‰) from Zhang et al. [55], the temperature of the precipitation infiltration for the YST, XN and PA geothermal systems are about 5 ~ 7C, 5C, and 4 ~ 5C, respectively. The recharge temperature for geothermal waters is cooler by about 3C compared with the recharge temperature of 8 ~ 10C for cold waters from the YST geothermal system.


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Circulation of geothermal waters

Thermal waters and some cold waters from the YST geothermal system, recharged by precipitation infiltration, shift somewhat towards higher 2H or lower 18O values (Fig. 7), which may result from the influence of local precipitations. The differences in 2H and 18O can be related to age, climate and/or latitude of precipitation. Tritium results indicate that thermal waters are older than cold underground waters (discussed below). According to Liu et al.’s results [44], 2H values of Xining precipitation are about -60‰ for spring, -20‰ for summer, -60‰ ~ -80‰ for autumn, and -120‰ for winter, respectively, revealing the climate effect. Furthermore, Liu et al.’s results [56] also showed that 2H values decrease with latitude. However, plots of other geothermal waters (DR2005, ZK10, and Binglingshan) close to the global meteoric water line (GMWL) clearly indicate that geothermal waters are derived entirely from the deep circulation of meteoric water. Because all geothermal waters show no positive 18O shift, isotopic exchange with the country rocks did not occur significantly [47, 57]. Generally, a positive 18O shift, which is observed in high temperature geothermal systems [57], is not observed in the present low-temperature geothermal systems. Our result is also consistent with previous studies that showed no positive 18O shift in low-temperature geothermal systems [54, 58, 59]. Tritium has been widely used to identify modern water and as evidence of groundwater mixing [14, 60]. The tritium content for cold waters varies from about 0.29±2.87 to 28.69±3.17TU, whereas geothermal waters (with the exception of sample 7 from the YST geothermal systems) contain no tritium (Table 4). The lack of tritium suggests that the original recharging water remained uncontaminated from tritium produced by nuclear explosions since the 1950s, suggesting that geothermal waters have subsurface residence times of more than 50 years [58, 60]. In contrast, tritium content of cold waters from the YST geothermal system suggests that precipitation infiltration has taken place since the 1950s. It should be noted that the Xiaomajigou valley is the eastern margin of the YST geothermal system, and that the sample 11 is located at the exterior of the YST geothermal system. So the low tritium in the sample 11 may be attributed to local geological condition that could affect the precipitation recharge. Samples DR2005 and ZK10 contain tritium concentrations of 3.93±2.30 and 4.9±1.1TU, respectively. These results, combined with the carbon-14 ages of 22.28 kyr for DR2005 [20] and 22.50 kyr for ZK10 [61], suggest that meteoric waters in the XN and PA geothermal systems also have long subsurface residence time, and that shallow groundwater mixture occurred by upflow [62]. According to the equation of the isotopic admixture, the degree of mixing can be tentatively estimated.

( )cold water / warm water

( )

warm water admixture water

admixture water cold water

(2)

( )cold water (%)

( )

warm water admixture water

warm water cold water

(3)

warm water and cold water are average 18O values of the geothermal waters without tritium and the cold waters, respectively, and admixture water represents admixture water (sample 7). The ratio of cold water to geothermal water and the percentage of cold water in the YST geothermal water are about 0.21 and 17.24%, respectively.

SUMMARY AND CONCLUSIONS

Chemical compositions of geothermal and cold waters from the YST and PA geothermal systems indicate that all samples are bicarbonate waters, and that the main cations are Ca and Mg. This result suggests shallow circulation or the first stages of interaction between meteoric waters and rocks. For the XN geothermal system, Na is the main cation and SO4 and Cl are the main anions. Geothermometer results suggest that the temperatures of the YST, XN, and PA geothermal systems are 49~67°C, 61~102°C, and 45°C, respectively, and that quartz geothermometer is a reliable means of estimating the reservoir temperature in the southern Xining basin. A 2H and 18O plot (Fig. 7) shows the thermal and cold waters are of meteoric origin. The variability in the 2H and 18O ranges appears to result from the combined effects of age, climate and/or latitude. Geothermal waters and some cold waters in the YST geothermal system are shifted somewhat towards higher 2H or lower 18O values, which may be due to the influence of local precipitation. According to the relation between altitude and 18O, the range of recharge altitudes for geothermal waters is from 3091 to 3675 m a.s.l. The lack of tritium content implies that geothermal water from the YST geothermal system was recharged before the 1950s. Tritium content of geothermal waters shows a mixture of cold fresh groundwater by upflow. According to the admixture relation, the ratio of cold water to geothermal water is 0.21, and the percentage of cold water in the geothermal water is 17.24%. Although geothermal water has been recharged by precipitation, sustainable usage should be considered because geothermal waters have residence times of more than 50 years.

ACKNOWLEDGEMENTS

We are grateful to J.Y. Zhou and W.D. Shi for their help in the field. This work is supported by the China Geological Survey Project (1212010610108) and the National Natural Science Foundation of China (41172038). REFERENCES

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