# Hydrogeochemical Characteristics and Genesis of the Chazi Geothermal Field Area in Tibet

Authors: Mingzhong Song; Xujuan Lang; Zhennan Zhong; Fengxin Kang; Dawa Nan; Haoting Li; Haowen Yu; Shaoyun Liu; Sihang Han; Zhao Liu
Journal: Geofluids
 (2022)

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## Abstract

The Chazi geothermal field area is located in the large region of Shigatse in southern Tibet. The geothermal resources in this area are abundant, but their exploitation and utilization are low. By studying the water chemistry and isotope characteristics of geochemical fluids in the study area, information on water chemistry, heat storage temperature, recharge source, recharge elevation, and circulation depth was obtained. These results provide a scientific theoretical basis for improving the genetic mechanism of high-temperature geothermal systems in the study area. The type of geothermal fluid hydrochemicals in this area is mainly HCO3–Na. The isotopic geochemical method was used to determine that the recharge source of geothermal fluids was atmospheric precipitation, and the recharge elevation was 5200–6000 m. The geochemical thermometer, Na–K–Mg equilibrium diagram, and silica-enthalpy mixed model indicated the shallow and deep thermal storage temperatures of approximately 150 and 200°C, respectively, and the average circulation depth of 1163.38 m in the study area. Combined with the fracture structure and magmatic activity characteristics of the southwest Qinghai-Tibet Plateau, the source, storage, cover, and general situation of the area were preliminarily summarised, and the conceptual model of geothermal origin was established. The results can provide a scientific theoretical basis for the mechanism of high-temperature geothermal systems and subsequent drilling and resource development.


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## Body

## 1. Introduction

Under the current circumstances of energy conservation and emission reduction, a strategic action plan for energy development is vigorous establishment of renewable energy. As a clean, low-carbon, stable, and continuous noncarbon-based energy source, the use of geothermal resources is indispensable to achieve the goals of carbon peak and neutrality [1]. Tibet is the most geothermally active region and accounts for nearly a quarter of over 3,000 identified hydrothermally active areas in China [2, 3].Numerous studies have been conducted in the Yangbajing, Yangyi, and Dagejia geothermal fields. Ji and Ping et al. investigated the hydrochemical characteristics, genesis, and evolution of the Yangbajing geothermal field [4–6]. Qinghai and Chen studied the chemical anomaly of tungsten present in a high-temperature spring in the Targejia field [7]. Shijuan analysed the source of medium- and high-temperature geothermal fluids in Cuonna [8]. Qinghai et al. studied the hydrochemical characteristics of Yangyi geothermal field by conducting factor analyses [9]. Many scholars have studied the hydrochemical characteristics and genesis mechanism of geothermal fluids in a high-temperature geothermal display area of Tibet [10, 11]. A complete model of the geothermal system in Yangbajing field was established to improve the genetic model of the high-temperature geothermal system in this area. However, studies have shown that the distribution of geothermal water in the geotropic region of Tibet is highly nonuniform, and the geological background and reservoir characteristics of various geothermal areas are different. Therefore, for this area, the genetic model of a high-temperature geothermal system is not suitable, and this model cannot be directly applied to other geothermal display areas [12]. Located in Xigaze of the Tibet Autonomous Region, the Chazi geothermal area is undergoing rapid urbanisation. The geothermal resources of the region are undeveloped, and hot water is largely discharged in the form of springs, which cause serious waste of resources. Studies in this area have focused on plate tectonics [13–15], sedimentary evolution [16–18], geochemistry [19, 20], geological hazards [21, 22], geochronology [23, 24], and deposit geology [25, 26], with few studies focusing on the hydrogeochemical characteristics, evolution of geothermal fluids, models of geothermal genesis, and genetic conceptual models in the geothermal area. Shaoqiang et al. and Haoting et al. have estimated the temperature of geothermal reservoirs in the Chazi geothermal field by using the silica geothermometer and the chalcedony geothermometer, respectively, and concluded that the geothermal fluid is mixed with cold water in this geothermal field, but temperature correction of the thermal storage has not been conducted [27, 28]. In 1975, the Tibetan Plateau Scientific Expedition and Research Team of the Chinese Academy of Sciences conducted a field investigation of geothermal resources in the Xigaze area and recorded basic information, such as physical geography, eruption, and hydrochemical characteristics of exposed hot springs, in the Journal springs in Tibet [29]. In summary, further studies are required as limited literature is available on the Chazi geothermal area.This study investigated the hydrochemical characteristics, material sources, and characteristics of geothermal reservoirs in the Chazi geothermal area by using hydrogeochemistry, isotope geochemistry, and mixture models [30]. The investigation of the geothermal genetic mechanism of the Chazi geothermal area can contribute to a good conceptual model of the geothermal system, which can provide proof for the rational development and utilization of the geothermal resources in the area.


## 2. Overview of the Study Area

### 2.1. Geological Conditions of the Study Area

In the southwest of the Qinghai-Tibet plateau and the eastern segment of the Alpine-Himalayan structural belt (Figure1(a)), the Chazi geothermal area is situated along the Gangdese-Tengchong landmass. Its magma arc zone extends westward, eastward, and northward, and its south borders extend on the fore-arc basin by a deep fault [31].Figure 1
Geodetic structure location map and geological sketch of the study area.The exposed strata in this area feature volcanic rocks, sedimentary rocks of the Paleogene Linzizong group and the Quaternary strata. From the base to top, the Linzizong group includes Dianzhong, Nianbo, and Pana Formation. Only the Dianzhong Formation (E1-2d) and Pana Formation (E2p) appear in the survey area. The Dianzhong Formation (E1-2d) is mainly present in Duozebu Mountain, the south-western corner of the survey area, and is distributed in clusters around ridges with an exposed area of 3.5 km2. The lithology comprises neutral–medium acidic volcanic lava with acid volcaniclastic rock, in the form of light grey to off-white massive rhyolite and rhyolite clastic lava, without a rocky bottom in the survey area. The formation is unconformably overlain by Pana Formation, with a thickness of >1000 m. The Pana Formation (E2p) is developed around the periphery of the Dianzhong Formation, with an exposed area of approximately 9 km2. The lithology of the Pana Formation comprises dark grey blocky dacite, trachyte andesite, purple-red blocky andesite, and andesitic and porphyroclastic lava, without a rocky peak in the survey area. Eruption is covered by the Dianzhong Formation with angular unconformity, with a thickness of >2400 m [32, 33].


### 2.2. Tectonic Conditions of the Study Area

The fault structure is the major structure of heat control and conductivity in the study area. Fault development is mainly in NE and NW directions. The NW trending tectonic is a normal fault (Figure1(b)) with water-blocking features. Boiling hot springs are distributed within a few kilometres along the NW fault zone, where large areas of calcifications and calcified mounds exist. The left region of the NE fault is a water-conducting structure, located in the western thermal area. At its intersection with the NW trending fault, the fracture development leads to a crush zone, which provides a satisfactory channel for the storage of deep hot water and migration of underground hot water. The underground hot water is obstructed by the NW fault and the normal fault during migration, resulting in rising springs [34]. With good hydrogeological conditions, the Quaternary pore groundwater is composed of Holocene alluvium, impingement sand, sand gravel, and Holocene marsh deposits and has high water abundance. The Quaternary pore groundwater, rock weathering fractures, and structural phreatic aquifers flow through the geothermal area from the northwest to southeast.


### 2.3. Hydrogeology of the Study Area

In the study, groundwater area can be classified into pore and fissure water available in loose rock and bedrock, respectively, according to the water-bearing medium. In loose rocks, the pore water is mainly distributed in the intermountain valley and piedmont; this water belongs to the quaternary alluvial type and is mainly supplied by atmospheric precipitation and lateral runoff of groundwater in the bedrock mountain area. The runoff conditions of pore water are strictly controlled by the topography, and this water moves from the front of the mountain to lowland under the action of gravity. There are two main drainage methods: (1) to supply the water to a river through runoff and (2) to discharge the water in the form of evaporation. In the bedrock, groundwater is mainly supplied by atmospheric precipitation, and its runoff moves from a high to low ground. This water is drained in the form of springs and through evaporation and lateral runoff that supplies the pore groundwater of quaternary loose rocks [35].


## 2.1. Geological Conditions of the Study Area

In the southwest of the Qinghai-Tibet plateau and the eastern segment of the Alpine-Himalayan structural belt (Figure1(a)), the Chazi geothermal area is situated along the Gangdese-Tengchong landmass. Its magma arc zone extends westward, eastward, and northward, and its south borders extend on the fore-arc basin by a deep fault [31].Figure 1
Geodetic structure location map and geological sketch of the study area.The exposed strata in this area feature volcanic rocks, sedimentary rocks of the Paleogene Linzizong group and the Quaternary strata. From the base to top, the Linzizong group includes Dianzhong, Nianbo, and Pana Formation. Only the Dianzhong Formation (E1-2d) and Pana Formation (E2p) appear in the survey area. The Dianzhong Formation (E1-2d) is mainly present in Duozebu Mountain, the south-western corner of the survey area, and is distributed in clusters around ridges with an exposed area of 3.5 km2. The lithology comprises neutral–medium acidic volcanic lava with acid volcaniclastic rock, in the form of light grey to off-white massive rhyolite and rhyolite clastic lava, without a rocky bottom in the survey area. The formation is unconformably overlain by Pana Formation, with a thickness of >1000 m. The Pana Formation (E2p) is developed around the periphery of the Dianzhong Formation, with an exposed area of approximately 9 km2. The lithology of the Pana Formation comprises dark grey blocky dacite, trachyte andesite, purple-red blocky andesite, and andesitic and porphyroclastic lava, without a rocky peak in the survey area. Eruption is covered by the Dianzhong Formation with angular unconformity, with a thickness of >2400 m [32, 33].


## 2.2. Tectonic Conditions of the Study Area

The fault structure is the major structure of heat control and conductivity in the study area. Fault development is mainly in NE and NW directions. The NW trending tectonic is a normal fault (Figure1(b)) with water-blocking features. Boiling hot springs are distributed within a few kilometres along the NW fault zone, where large areas of calcifications and calcified mounds exist. The left region of the NE fault is a water-conducting structure, located in the western thermal area. At its intersection with the NW trending fault, the fracture development leads to a crush zone, which provides a satisfactory channel for the storage of deep hot water and migration of underground hot water. The underground hot water is obstructed by the NW fault and the normal fault during migration, resulting in rising springs [34]. With good hydrogeological conditions, the Quaternary pore groundwater is composed of Holocene alluvium, impingement sand, sand gravel, and Holocene marsh deposits and has high water abundance. The Quaternary pore groundwater, rock weathering fractures, and structural phreatic aquifers flow through the geothermal area from the northwest to southeast.


## 2.3. Hydrogeology of the Study Area

In the study, groundwater area can be classified into pore and fissure water available in loose rock and bedrock, respectively, according to the water-bearing medium. In loose rocks, the pore water is mainly distributed in the intermountain valley and piedmont; this water belongs to the quaternary alluvial type and is mainly supplied by atmospheric precipitation and lateral runoff of groundwater in the bedrock mountain area. The runoff conditions of pore water are strictly controlled by the topography, and this water moves from the front of the mountain to lowland under the action of gravity. There are two main drainage methods: (1) to supply the water to a river through runoff and (2) to discharge the water in the form of evaporation. In the bedrock, groundwater is mainly supplied by atmospheric precipitation, and its runoff moves from a high to low ground. This water is drained in the form of springs and through evaporation and lateral runoff that supplies the pore groundwater of quaternary loose rocks [35].


## 3. Water Sampling and Tests

In this study, field investigation and systematic sampling were conducted in the study area in 2019, and 15 typical hot spring water samples, 3 river water samples, and 1 cold water well sample were collected. The collected samples are numbered as follows: Taegejia Hot Springs (DR28), Semi Hot Springs (DR32), Charong Qurani Hot Springs (DR36), Dingle Hot Springs (DR37), Polah Hot Springs (DR40), Qugu Hot Springs (DR42), Luo Hot Spring (DR44), Rub Hot Spring (DR45), Qugulongbu Hot Spring (DR47), Shalei Hot Spring (DR48), Tamar Quzhen Hot Spring (DR9), Yinjian Quzhen Hot Spring (DR11), Gabu Nidan Hot Spring (DR15), Chazi Hot Springs (CKLSQ1, CKLSQ2), Chaqu River Water (CQSH, CQXH), Darong Tsangbu River Water (DLZBHX), and Lengshui Well (J05). The sampling position distribution is shown in Figure1(b). The hydrochemical data of 10 hot springs in the “Tibet Hot Springs” are as follows: Tagejia Hot Springs (DR28-1, DR28-2, DR28-3), Semi Hot Springs (DR32-1), Qugu Hot Springs (DR42-1), Seluo Hot Spring (DR44-1), Rub Hot Spring (DR45-1), Qugulongbu Hot Spring (DR47-1), Shalei Hot Spring (DR48-1), Chazi Hot Spring (CKLSQ-1), and five sets of hydrogen and oxygen isotope data of the Gejia Hot Springs.The pH, water temperature, geographical location, and altitude of the region were measured during the study. The samples were collected and stored in PTEF bottles, and the sample bottle and sample plug were rinsed thrice before sampling. The samples were placed in a portable refrigerator. The chemical analysis of the water samples was performed at the Central Laboratory of Tibet Geological Survey Bureau. Hydrogen and oxygen isotopes were detected in the Institute of Hydrogeology and Environmental Geology, CAGS. The pH value was measured using the glass electrode method by using a 3-STARpH meter, and the reading accuracy was ≤0.02. TDS was tested by employing the dry-weight method; the main instruments used were a type 101 electric heat drum air drying box (oven temperature control accuracy ±1°C) and an AL104-IC electronic analysis balance (sense 0.1 mg). K+, Na+, Ca2+, and H2SiO4 were detected through inductively coupled plasma optical emission spectrometry (ICP-OES) (RSD<3%) (Optima 5300DV inductively coupled plasma emission spectrometer, Pergin Elmer, USA). Mg2+, Li, W, Rb, Fe, Mn, Cs, Ge, Sr, and As were detected through inductively coupled plasma mass spectrometry (RSD<8%) (Optima 5300DV inductively coupled plasma emission spectrometer, Perkin Elmer, USA). The total arsenic was measured through hydride generation atomic fluorescence spectrometry (RSD<5%); the HCO3− content was determined through acid–base titration (RSD<2.5%); the SO42− content was measured through disodium edetate titration (RSD<1.5%); the Cl− concentration was estimated using the nitrate volumetric method (RSD<3%); the F and Br contents were estimated through ion chromatography (RSD<2%); the sulphur content was measured through sulphide by using iodometry (RSD<4%); the HBO2 content was determined by H-acid–methylimine spectrophotometry with the minimum determination of 1 μg. The sample monitoring method was in line with the method used for drinking natural mineral water (GB8538-2016) and the test method employed for groundwater quality determination (DZ/T0064-93). The δ2H-δ18O isotope detection instrument used was a liquid water isotope laser spectrometer (model number L2130i, made in China). The test principle is based on the cavity axis integrated cavity output spectroscopy. Degree of accuracy was δ2Hv−SMOW‰≤±1‰ and δ18Ov−SMOW‰≤±0.2‰. The high detection accuracy and standard acceptable errors indicated the reliable results.


## 4. Results and Analysis

### 4.1. Hydrogeochemical Characteristics of Geothermal Fluids

According to the hydrochemical analysis data (Tables1 and 2), the geothermal fluid present in the study area is alkaline brackish water, with the pH of >7 and of TDS>1g/L. The hydrochemical type was HCO3–Na, with trace components of HBO2, Li, Sr, and F−. The Li content exceeded the limit of medical mineral water (minimum limit value is 1 mg/L), and HBO2 and F− contents were considerably higher than the limit of medically permitted mineral water (minimum limit values are 50 and 2 mg/L, respectively), indicating the physiotherapy value and utilization prospects. For trace elements, As and F− concentrations increased slightly, whereas the B concentration decreased because of three reasons: (1) the content of fluoride ions in alkalinity is higher than in acidity as the fluorine complex is easy to be hydrolysed in an alkaline environment; (2) alkaline water is conducive to the dissolution of fluorine-containing minerals [36]; (3) magma and late enrichment lead to an increase in the As content. The hydrochemical composition of the geothermal fluid in the Chazi geothermal area has not changed considerably in recent years.Table 1
Major constituents of the water samples in the study area.
Field numberGeothermal display areaSample typeWater temperature/°CpHρB/(mg/L)Hydrochemical typeNa+K+Ca2+Mg2+HCO3-CO32-Cl-SO42-TDSDR28Lagejia springHot spring858.3386.235.12.90.0449.2106.915889.91732HCO3-NaDR32Semi springHot spring898.7790.5125.50.70.00571.3722.810.14350Cl·CO3-NaDR36Charongqurene springHot spring657.1564.758.253.36.7807.50262.94362333.7HCO3·SO4- NaDR37Dengle springHot spring436.6837.272.366.814.61201.10919.337.53442.1Cl·HCO3-NaDR40Pola springHot spring306.6107.15.18.50.299.8090.543.6421.4Cl ·HCO3-NaDR42Qugu springHot spring478.9117.92.40.7<0.1268.920.94.212.2469.2HCO3-NaDR44Seluo springHot spring446.7195.040.2195.837.71105.20226.62.91893.1HCO3·cl-ca·NaDR45Cha springHot spring487.1476.426.670.99.21246.50200.528.32242.3HCO3-NaDR47QugulongbuHot spring666.9454.725.350.37.6616.90280.6250.81903.5HCO3·cl-NaDR48Shalei springHot spring757.7115.410.428.87378.507.259.2710.3HCO3-NaXZDR9TamaquzhenHot spring637.2407.431.64012.6856.30190.578.21919HCO3-NaXZDR11YinjianquzhenHot spring396.559.45.421020.9811.908.934.81113HCO3-CaXZDR15GabudanniHot spring677.0451.429.719.2<0.1584.40259.4169.91922HCO3·Cl-NaCKLSQ1Chazi VillageHot spring877.8481.939.112.8<0.1654.987.1165.1121.91806.9HCO3-NaCKLSQ2Chazi VillageHot spring747.9476.941.19.5<0.1707.358.1167.8117.51844.1HCO3-NaCQSHChazi VillageChaqu River water7.57.513.11.440.96.6115.901.4147.3279HCO3·SO4-CaCQXHChazi VillageChaqu River water25.58.6363.730.119.21.82424.1108.2118.1106.11386.1HCO3-NaDLZBXHChazi VillageDalong Zangbo River water6.58.38.12.544.54.978.41.13.175.3233.2SO4·HCO3-CaJ05Chazi VillageWell water47.55.72.3252.178.702.716.6115.6HCO3-CaDR28-1Lagejia springHot spring85.59.1365417.21.92612341651321570CO3·Cl -HCO3DR28-2Lagejia springHot spring788.4388.9350.60.2590.636112.11021712HCO3-NaDR28-3Lagejia springHot spring858.73503900447.2100.8151.777.81600HCO3-NaDR32-1Semi springHot spring868.77001043.44.249030667465.84340Cl·CO3·HCO3-NaDR42-1Qugu springHot spring458.91182.53.41.322745.713.928.8227HCO3·CO3·SO4·Cl-NaDR44-1Seluo springHot spring407.317533.320222.91040—157741380HCO3·Cl-Ca·NaDR45-1Cha springHot spring44.57.344326.573.812.51130—17232.91560HCO3·Cl-Na·CaDR47-1QugulongbuHot spring65.27442.526.553.29.4493—2722801600HCO3·Cl-Na·CaDR48-1Shalei springHot spring767.512512.534.38.338.1—9.626.8510HCO3-Na·CaCKLSQ-1Chazi VillageHot spring877.946845124.2801—1631441560HCO3·Cl·SO4-Na∗Note: DR28-1, DR28-2, DR28-3, DR32-1, DR42-1, DR44-1, DR45-1, DR47-1, DR48-1, and CKLSQ-1 are quoted from the report of Wei et al. [29].Table 2
Trace components of the water samples in the study area.
Field numberρB/(mg/L)HBO2SiO2Fe2+Fe3+F-NO3-NO2-PO43-LiSrAsDR28421.6286.9<0.05<0.0525.80.3<0.00330.11960.28.6DR322130536.5<0.05<0.0522.4<0.05<0.00330.11933.90.212.2DR3650.783.9<0.03<0.038.70.8<0.0010.01193.62.70.2DR37425.687.20.1<0.038.65.2<0.0010.0119142.84.9DR4020.455.1<0.05<0.053.81.5<0.00330.1190.60.20.5DR426.250.2<0.03<0.033.20.02<0.0010.01190.20.30.0006DR44181.634.60.2<0.030.90.1<0.0010.01196.21.80.0006DR45165.169.7<0.03<0.030.30.1<0.0010.011933.90.0006DR47200.368.40.3<0.036.40.1<0.0010.01192.826.2DR4826.676.7<0.03<0.030.20.1<0.0010.01190.31.80.0006XZDR9360.9170.8<0.050.199.40.06<0.00330.01194.61.25.6XZDR117.640<0.05<0.050.71.6<0.00330.01190.32.20.0002XZDR15416.7138.3<0.050.3910.5<0.05<0.00330.01196.61.16CKLSQ1116.3129.9<0.03<0.0315.50.1<0.0010.01192.20.70.7CKLSQ2120.7144<0.03<0.0316.90.1<0.0010.01192.20.70.9CQSH0.535.10.98<0.032.40.02<0.0010.01190.010.50.0049CQXH93.6104.30.26<0.0312.50.1<0.0010.01191.60.60.6DLZBXH1.815.7<0.03<0.030.30.03<0.0010.01190.040.30.0008J050.2———0.32.5<0.001———0.0014DR28-144022.7——25———7.3—5.4DR28-2207.887.5——30.4———5.4——DR28-3165.6120——27.5———12.3—5.5DR32-11750116.2——13.3———35—30.8DR42-1836——4.9———<0.1——DR44-116032.2——0.9———4.9——DR45-117357.8——0.8———2.8——DR47-118778.8——3.6———2.7—4DR48-12574.7——3.5———0.3——CKLSQ1-1170145——9.7———2.2—0.4The hydrochemistry of surface water in CQSH, CQXH, and DLZBXH was HCO3·SO4–Ca, HCO3–Na, and SO4·HCO3–Ca (Figure 2), respectively, with the main cations of Ca2+ and Na+, main anions of SO42− and HCO3−, and lower TDS than that of hot spring water. Table 1 presents the differences in the hydrochemical composition between hot and surface water. However, the hydrochemical composition of the water samples collected on the downstream of Chaqu was similar to that of hot water samples, which could explain the downstream water flowing through the hot water area to mix with the geothermal water due to short runoff distance.Figure 2
Chemical piper diagram of the water samples.


### 4.2. Source of Materials

#### 4.2.1. Ionic Component

The Na+/Cl− ratio is considerable higher than 1 (Table 3), indicating strong dissolution and filtration in groundwater [37, 38]. Hot water exhibits strong dissolution for the surrounding rocks during upwelling, enabling Na+, K+, and Ca2+ in the potash feldspar and plagioclase to migrate into the hot water [10, 39]. The pH of the geothermal fluid was 7.83–7.93, a range in which the carbonic acid balance favoured HCO3−. The HCO3− content in the fluid samples collected in downstream Chaqu (CQXH) was drastically higher than that in upstream Chaqu presumably because the lateral runoff occurred during the vertical upwelling of geothermal water and was subsequently mixed into the river [40].Table 3
Ratios of Na and Cl in the study area.
Field numberGeothermal display areaSample typeρB/(mg/L)Na+/Cl-Na+Cl-DR28Lagejia springHot spring386.21582.4DR32Semi springHot spring790.5722.81.1DR36Charongqurene springHot spring564.7262.92.1DR37Dengle springHot spring837.2919.30.9DR40Pola springHot spring107.190.51.2DR42Qugu springHot spring117.94.228.3DR44Seluo springHot spring195226.60.9DR45Cha springHot spring476.4200.52.4DR47QugulongbuHot spring454.7280.61.6DR48Shalei springHot spring115.47.216.1XZDR9TamaquzhenHot spring407.4190.52.1XZDR11YinjianquzhenHot spring59.48.96.7XZDR15GabudanniHot spring451.4259.41.7CKLSQ1Chazi VillageHot spring481.9165.12.9CKLSQ2Chazi VillageHot spring476.9167.82.8CQSHChazi VillageChaqu River water13.11.49.3CQXHChazi VillageChaqu River water363.7118.13.1DLZBXHChazi VillageDalong Zangbo River water8.13.12.6J05Chazi VillageWell water5.72.82.1DR28-1Lagejia springHot spring3651652.2DR28-2Lagejia springHot spring388.9112.13.5DR28-3Lagejia springHot spring350151.72.3DR32-1Semi springHot spring7006741.0DR42-1Qugu springHot spring11813.98.5DR44-1Seluo springHot spring1751571.1DR45-1Cha springHot spring4431722.6DR47-1QugulongbuHot spring442.52721.6DR48-1Shalei springHot spring1259.613.1CKLSQ-1Chazi VillageHot spring4681632.9Almost all the thermal spring samples in the study area are observed below the 1 : 1 line (Figure3(a)), indicating that considerable ion replacement occurs during runoff. The low concentrations of Ca2+ and Mg2+ (Table 1) suggest that in the water, Ca2+ and Mg2+ are replaced with absorbed Na+ in the surrounding rock. The γNa/γ (Na+Ca+Mg) content of hot water is higher than that of cold water, and the value is close to 1 (Figure 3(c)), which indicates considerable ion exchange between Na+–Ca2+ and Na+–Mg2+ in geothermal fluids in this area. Ca2+ and Mg2+ available in the geothermal fluid are replaced with Na+ present on the rock and soil surfaces, which confirmed the aforementioned speculation. Almost all the fluids are distributed above the 2 : 1 line (Figure 3(b)), which indicates that the main mineral dissolved by carbonate in groundwater is calcite.Figure 3
Ratio diagrams of the ions of water samples in the study area.
(a)(b)(c)The high concentration of SO42− is a major result of the strong dissolution of geothermal fluids that occurs in the bottom strata of the Paleogene Pana formation (E2p) dominated by dacite, andesite, and andesitic and porphyroclastic lava according to the aforementioned discussion [41]. In the study area, the variation range of δ34S is small, and all fall within the range of atmospheric sulphate (Figure 4). The δ34S of DR32 and XZDR9 deviated from the sulphur of meteorites, suggesting that the sulphur of DR32 and XZDR9 is mainly leached from the surrounding rocks of the sedimentary cover. The δ34S of XZDR10 is close to that of meteorite sulphur, indicating that the sulphur source is located deep within the earth’s crust and may be regenerated magma produced by remelting in the crust. The mixing of deep sulphur is uniform, and no significant sulphur isotope differentiation can be observed in the geothermal fluid ascent. Some geothermal water samples located in the atmospheric sulphate range are distributed in the volcanic sulphur (SO2 and SO4) range. The geological conditions and the fact that no volcanic activity occurs in the study area indicate that the sulphur in geothermal fluids is caused by the dissolution of the surrounding rocks.Figure 4
δ34S-ρSO42− plot of geothermal water in the study area.The main sources of Cl− are human activities, deep magma, and leaching effects on rocks. Because this area is vast and not populous, with few Cl− minerals in the stratum, it is speculated that Cl− in geothermal fluids may come from deep magma. The Cl− concentration in the hot springs of the study area is approximately 200mg/L; however, that in the Semi spring and Dengle spring is higher (Table 1). The deep matter in these two hot springs is speculated to mix to a large extent. The low Cl− content of the Shalei spring distributed near the Yarlung Zangbo River may be caused by mixing with river water.


#### 4.2.2. Trace Elements

In general, the F− concentration in the hot water in the study area is high, which is related to the presence of fluorine-containing minerals feldspar and biotite in the strata (Table 2) [42]. Under natural conditions, B and Li in the groundwater majorly originate from the dissolution of related minerals in rocks. The B and Li concentrations in geothermal fluids are positively correlated with the Cl− concentration (Figure 5), indicating that B, Li, and Cl− have the same material source. In addition to the dissolution of rock minerals, other sources are responsible for high B and Li concentrations [43–46]. Studies have revealed that B, Li, Rb, and As in the geothermal water in Tibet originate from the residual magma in the late-stage magmatism [47]. Considering the geothermal and geological conditions of this area, the rich content of Li and Sr in the hot water might be caused by the long-term leaching effect of deep geothermal fluids on volcanic rock formation (Paleogene Dianzhong Formation (E1-2d)).Figure 5
Correlation diagram between trace components and Cl− in the study area.


### 4.3. Isotope Characteristics of the Geothermal Fluids

#### 4.3.1. Source of Recharge

Hydrogen and oxygen isotopes are typically used to trace the recharge source for the groundwater, determine the intensity of groundwater runoff, and identify the water-rock interaction between the groundwater runoff and surrounding rocks [48–50]. Craig analysed the δ2H and δ18O values of more than 400 natural water samples on a global scale and obtained the global meteoric water line: δ2H=8δ18O+10 [51]. Scholars have often used the positional relation between the hydrogen and oxygen stable isotope data in the study area to determine the replenishment source of underground fluids [51].In this study, the hydrogen and oxygen isotope data measured for 6 samples and that collected for 5 samples (Table4) were used to analyse the recharge source and recharge elevation of the geothermal fluid in the area. Figure 6 reveals that all the hydrogen and oxygen isotopes were distributed near the global atmospheric rainfall equation line [51], indicating that the main recharge source for the underground hot water system was atmospheric rainfall. Moreover, the δ18O and δ2H of the geothermal fluid and surface water did not differ considerably, which indicated that the recharge source of the underground hot and surface water comes from the surrounding atmospheric precipitation infiltration and not through long-distance migration; that is, the water–rock interaction is not strong [52, 53]. Another possible reason is that the geothermal fluids in the study area were mixed with surface cold water during the rising process.Table 4
Data ofδ2H–δ18O of the geothermal fluid in the study area.
Field numberGeothermal display areaSample typeδ2HVSMOW(‰)δ18OVSMOW (‰)Data sourcesDR28-1Lagejia springHot spring-149-18.6Quoted from literature [47]DR28-2Lagejia springHot spring-154-18.6DR28-3Lagejia springHot spring-153-18.6DR28-4Xiaolupang springHot spring-155-18.6CMQChangmaqu River waterRiver water-130-18.3CKLSQ1Chazi VillageHot spring-163-20.5MeasuredCKLSQ2Chazi VillageHot spring-163-20.6CQSHChazi VillageRiver water-136-17.6CQXHChazi VillageRiver water-148-18.7DLZBXHChazi VillageRiver water-164-20.9Figure 6
δ2H–δ18O isotope relationship of various water samples.


#### 4.3.2. Recharge Elevation

Theδ2H and δ18O of atmospheric precipitation often exhibit the characteristics of temperatural, continental, seasonal fluctuation, etc. The following method can be used to estimate the replenishment elevation of each hot spring in the study area based on the elevation effect of the stable isotope of oxygen.
(1)H=δG−δPK+h,where H is the recharge elevation, m; h is the elevation of the sampling point, m; δG is the δ18O of the hot spring, ‰; δP is the δ18O of the atmospheric precipitation, ‰; K is the elevation gradient of the δ18O of the atmospheric precipitation, δ/100 m. In this paper, the CQSH data (δ18O is −17.6‰) at the upstream point of Chaqu is used as the δ18O value of atmospheric precipitation in Chazi Hot Spring, and the CMQ (δ18O is −18.29‰) at the Changmaqu River water point is used as the δ18O value of atmospheric precipitation in the Gejia Hot Spring. The elevation gradient of δ18O in Tibet (−0.26‰/100 m) is used as the K value [54]. The calculation results are presented in Table 5.Table 5
Calculated values of the geothermal fluid replenishment elevation in the study area.
Field numberCKLSQ1CKLSQ2DR28-1DR28-2DR28-3DR28-4Sampling elevation484148375054505850475050Replenishment elevation595659905165518851775184The geothermal fluid replenishment elevation range of the study area is 5200–6000 m (Table5). This estimated value is close to the elevation data of the sampling point. It is inferred that the replenishment of geothermal water mainly comes from atmospheric rainfall infiltration replenishment received by the mountains near the hot spring outcrop area and a small amount of alpine snowmelt replenishment.


### 4.4. Temperature of Geothermal Reservoirs and Cycle Depth Estimation

#### 4.4.1. Geochemical Geothermometer for the Temperature Estimation of Geothermal Reservoirs

The geothermometer is an effective method for estimating the temperature of thermal storage, and each temperature scale has certain assumptions. Therefore, when using the geothermometers to study the temperature of the thermal reservoir, the applicable scope and conditions of each temperature scale should be considered to select a suitable geothermometer to estimate the thermal storage temperature. The commonly used geothermometers mainly include cationic geochemistry, SiO2 geothermometer, isotope geothermometer, and gas geothermometer, among which the most commonly used are cationic geochemistry and SiO2 geothermometers. According to the hydrochemical characteristics of the geothermal fluid in the study area, the geothermal geology, and the applicable conditions of the geothermal geothermometers (Table 6), this study used the SiO2 geothermometer (quartz, chalcedony) and Na-K geothermometer to estimate thermal storage temperature for the study area [55–57]. Because the study region is located in a high-altitude area in Tibet, the boiling point of water is low. In this study, we designated the hot springs with exposure temperatures of >80°C as boiling springs. That is, when the exposure temperature of the sampling point was ≥80°C, the maximum steam loss formula was used when the SiO2 geothermometer was used to estimate thermal storage temperature. When the exposure temperature of the sampling point did not reach 80°C, the no-steam loss formula was used. The calculation results are shown in Table 7.Table 6
Geothermometer function of geothermal water in the study area.
GeothermometersCalculation formulaScope of applicationQuartz1T=13095.19−logSiO2−273.15No loss of steamQuartz2T=15225.75−logSiO2−273.15Maximum steam loss 0-250°CChalcedony1T=10324.69−logSiO2−273.15High temperature thermal storageChalcedony2T=12645.31−logSiO2−273.15Maximum steam lossNa-KT=1217logNa/K+1.483−273.15t>150°CTable 7
Estimated results of the reservoir temperature in the study area.
Field numberSampling temperature/°CSiO2 geothCation geothNa-K/°CQuartz geoth (minimum steam loss)/°CQuartz geoth (maximum steam loss)/°CChalcedony geoth (minimum steam loss)/°CChalcedony geoth (maximum steam loss)/°CDR2885—189.2—170224.8DR3289—230.8—216.7251.8DR3665127.6—99.9—234.7DR3743129.7—102.1—220.8DR4030106.4—76.8—179.8DR4247102—72.1—131.6DR444485.4—54.4—297.4DR4548117.9—89.3—189.6DR4766117—88.3—189.3DR4875122.9—94.7—223.9XZDR963169.4—146.8—212.7XZDR113991.7—61.1—224.9XZDR1567156.2—131.7—201CKLSQ187—145.4—122.3216.2CKLSQ274158.7—134.5—220.6The reservoir temperature calculated using chalcedony (minimum steam loss) is not the true reflection of the hydrothermal system because the estimated values are not substantially different from the reservoir temperature of the thermal springs (Table7). The reservoir temperatures estimated using quartz and chalcedony geothermometer (maximum steam loss) of DR28, DR32, and CKLSQ1 are the same, indicating that the temperatures acquired using this formula are highly reliable (Table 7). For other sample points, the reservoir temperature calculated using quartz geothermometer (no steam loss) is mostly distributed in 120–160°C, with an average of 126°C. The estimated values are slightly low combined with the geothermal background. The estimated value of Na–K geothermometer is higher than that of SiO2 geothermometer. The low value of SiO2 geothermometer may be caused by the following reasons: (1) the hot spring fluid was mixed with varying degrees of cold water during the ascent. (2) Saturated SiO2 dissolved in hot water above 180°C precipitated [48]. From the Na–K–Mg triangle diagram (Figure 7), the groundwater in the study area had not reached a water–rock equilibrium state, and cold water was mixed in the geothermal fluid. This is the probable reason for the low reservoir temperature estimation by using SiO2 geothermometer. Most thermal springs are located near the partial equilibrium zone, which shows that the method of reservoir temperature estimation by using the Na–K geothermometer presents certain reference significance. The estimation results acquired using the Na-K geothermometer are considered to reflect deep reservoir characteristics, and those obtained by employing SiO2 geothermometer are considered to reflect shallow reservoir characteristics combined with the applicable conditions of Na–K geothermometer.Figure 7
Na–K–Mg triangle diagram of the geothermal fluid in the study area.


#### 4.4.2. Silica-Enthalpy Mixing Model to Estimate the Temperature of Geothermal Reservoirs

According to the aforementioned analysis, the hot water in the study area is likely to mix with the shallow cold water during the ascent. A strong linear relationship exists between Cl and B in the geothermal fluids (Figure8). CQSH is the upstream sample, and CQXH is the downstream sample. Most geothermal fluid samples are located at the cold water end (Figure 8). No significant differences can be observed in the compositions of δ2H and δ18O of the geothermal fluids (Figure 6). Therefore, the geothermal fluid in this area can be considered a mixture of deep geothermal fluids and cold water [58, 59].Figure 8
Relationship between Cl and B contents in the study area.In 1974, Fournier and Truesdell developed a silica-enthalpy equation method to estimate the true thermal storage temperature before the cold water is mixed. The principle of this method is that the mixing of hot and cold water inevitably causes the initial enthalpy and SiO2 content of deep water to decrease to the final enthalpy and SiO2 content of spring water. According to this principle, the model equation is derived:
(2)ScX1+Sh1−X1=Ss,(3)ρcSilicaX2+ρhSilicaX21−X2=ρsSilica,where Sc is the enthalpy of cold water (J/g); Ss is the final enthalpy of spring water (J/g), below 100°C, and the enthalpy of saturated water is equal to the water Celsius degree. For >100°C, the relationship between temperature and enthalpy of saturated water is presented in Table 8. Sh is the initial enthalpy of hot water (J/g); ρcSilica is the silica mass concentration of cold water (mg/L); ρsSilica is the silica mass concentration of the spring water (mg/L); ρhSilica is the initial silica mass concentration of hot water (mg/L); and X is the mixing proportion of cold water.Table 8
Relationship between water temperature, enthalpy, and mass concentration of SiO2.
t/°CenthalPy (∗4.1868J/g)ρcSiO2/(mg/L)505013.5757526.6100100.148125125.180150151125175177185200203.6265225230.9365250259.2486275289614300321692In this study, the enthalpy of cold water was determined to be the average temperature of the study area at 5°C in the year, and the final enthalpy of the hot spring was the enthalpy value corresponding to the measured temperature value. The silica mass concentration of cold water (upstream of Chaqu) was 35.05 mg/L, and the mass concentration of SiO2 in the hot spring takes the actual measured value this time (Table 2). Substitute the concentration into the model equation to obtain the values of X1 and X2 for each hot spring point (Table 9), and then, project the calculated values of X1 and X2 and their corresponding temperatures into the scatter diagram, two curves can be drawn. The abscissa corresponding to the intersection of the two curves is the estimated value of heat storage temperature before mixing, and the ordinate corresponding to the intersection of the two curves is the mixing ratio of cold water (Figure 9).Table 9
Calculated values ofX1 and X2 hot springs in the study area.
Temperature/°C5075100125150175200225250275300DR28X1-0.78-0.140.160.330.450.530.60.650.690.720.75X212.6930.81-18.45-4.6-1.8-0.68-0.10.240.440.560.62DR32X1-0.87-0.20.120.30.420.510.580.630.670.70.73X224.2760.35-37.73-10.16-4.58-2.34-1.18-0.52-0.110.130.24DR36X1-0.330.140.370.50.590.650.70.730.760.790.81X23.276.78-2.77-0.090.460.670.790.850.890.920.93DR37X10.160.460.60.680.740.780.810.830.850.870.88X23.427.17-3.02-0.160.420.650.770.840.880.910.92DR40X10.440.640.740.790.830.850.870.890.90.910.92X21.933.37-0.540.560.780.870.910.940.960.970.97DR42X10.070.40.560.650.710.760.790.810.830.850.87X21.72.79-0.170.660.830.90.930.950.970.970.98DR44X10.130.440.590.680.730.770.80.830.850.860.88X20.980.951.031.011111111DR45X10.040.390.550.640.710.750.780.810.830.850.86X22.615.1-1.670.230.620.770.850.90.920.940.95DR47X1-0.360.130.360.490.580.650.690.730.760.790.81X22.554.95-1.580.260.630.780.850.90.930.940.95DR48X1-0.5600.260.420.520.590.650.690.720.750.78X22.935.93-2.220.070.540.720.820.870.910.930.94XZDR9X1-0.290.170.390.520.60.660.710.740.770.80.82X27.317.06-9.48-2.02-0.510.090.410.590.70.770.79XZDR11X10.240.510.640.720.770.80.830.850.870.880.89X21.231.590.620.890.940.970.980.980.990.990.99XZDR15X1-0.380.110.350.480.580.640.690.730.760.780.8X25.7913.22-6.97-1.3-0.150.310.550.690.770.820.84CKLSQ1X1-0.82-0.170.140.320.440.520.590.640.680.710.74X25.412.22-6.32-1.11-0.050.370.590.710.790.840.86CKLSQ2X1-0.530.010.270.430.530.60.650.690.730.760.78X26.0613.9-7.42-1.42-0.210.270.530.670.760.810.83Figure 9
Si-enthalpy model of the study area.The mixing ratio of cold water in the hot springs in the study area is relatively large, and the heat storage temperature before the cold water mixed is generally above 150 °C, which indicates the characteristics of high-temperature heat storage in the study area (Figure9). The abnormalities are the five hot springs: DR28, DR32, DR42, DR44, and XZDR9. They present no intersection in Figure 5, indicating that these five hot springs are affected by the absence of cold water mixing. The thermal storage temperatures of DR36, DR37, DR40, DR45, DR47, DR48, XZDR11, XZDR15, CKLSQ1, and CKLSQ2 before the mixing of cold water are 170°C, 217°C, 169°C, 169°C, 145°C, 149°C, 104°C, 241°C, 200°C, and 234°C, respectively, and their mixing ratios of cold water are 64%, 82%, 85%, 74%, 57%, 52%, 65%, 74%, 59%, and 71%, respectively.In summary, the average mixing ratio of hot spring and cold water in the study area is 75%, and the average value of the heat storage temperature before the cold water mixed is 200°C, which is basically the same as the heat storage temperature obtained in the Na-K-Mg balance diagram and the result of the Na-K geothermometer. This reflects the real storage temperature before the cold water is mixed in the study area, that is, the deep heat is stored in the study area. The calculation results of the SiO2 (maximum steam loss) geothermometer are distributed around 150°C, indicating shallow heat storage in the study area.


#### 4.4.3. Circulation Depth

The foregoing hydrogen and oxygen isotope research results showed that the geothermal fluid in the study area is mainly of atmospheric rainfall replenishment, and the temperature of hot water chiefly depends on deep-circulation geothermal heating. Therefore, the calculation of the temperature of underground hot water storage can be used to estimate the cycle depth of the hot water with the following formula:(4)Z=GTz−T0+Z0,where Z is the calculated cycle depth, m; G is the geothermal heating gradient, m/°C, with a value of 6.67 m/°C [60]; TZ is the heat storage temperature, °C, taking the heat storage temperature determined using an existing mixed model; T0 is the multiyear average temperature of the replenishment area, °C, its value is 5 °C; and Z0 is the depth of the normal temperature zone, m, the value is 20 m. The calculation results of the cycle depth are shown in Table 10.Table 10
Calculation of underground hot water circulation depth in in the study area.
Field numberSampling point elevation/mReservoir temperature/°CCycle depth/mDR285054202.11335DR324262237.71572DR3648861701121DR3752592121401DR4049231601054DR424382109.4716.4DR444279183.61212DR4541571631074DR474173144947.1DR484170148973.8XZDR95030185.41223XZDR114377102667XZDR1547332361561CKLSQ4839216.51431The circulation depths of DR42, DR47, DR48, and XZDR11 hot springs are shallow, ranging from 667 to 974 m (Table10). The circulation depths of all the other hot springs are >1000 m. DR32 and XZDR15 hot springs show the maximum circulation depth of 1500 m. Overall, the average circulation depth of the study area is 1163.38 m, and the geothermal conditions are satisfactory.


## 4.1. Hydrogeochemical Characteristics of Geothermal Fluids

According to the hydrochemical analysis data (Tables1 and 2), the geothermal fluid present in the study area is alkaline brackish water, with the pH of >7 and of TDS>1g/L. The hydrochemical type was HCO3–Na, with trace components of HBO2, Li, Sr, and F−. The Li content exceeded the limit of medical mineral water (minimum limit value is 1 mg/L), and HBO2 and F− contents were considerably higher than the limit of medically permitted mineral water (minimum limit values are 50 and 2 mg/L, respectively), indicating the physiotherapy value and utilization prospects. For trace elements, As and F− concentrations increased slightly, whereas the B concentration decreased because of three reasons: (1) the content of fluoride ions in alkalinity is higher than in acidity as the fluorine complex is easy to be hydrolysed in an alkaline environment; (2) alkaline water is conducive to the dissolution of fluorine-containing minerals [36]; (3) magma and late enrichment lead to an increase in the As content. The hydrochemical composition of the geothermal fluid in the Chazi geothermal area has not changed considerably in recent years.Table 1
Major constituents of the water samples in the study area.
Field numberGeothermal display areaSample typeWater temperature/°CpHρB/(mg/L)Hydrochemical typeNa+K+Ca2+Mg2+HCO3-CO32-Cl-SO42-TDSDR28Lagejia springHot spring858.3386.235.12.90.0449.2106.915889.91732HCO3-NaDR32Semi springHot spring898.7790.5125.50.70.00571.3722.810.14350Cl·CO3-NaDR36Charongqurene springHot spring657.1564.758.253.36.7807.50262.94362333.7HCO3·SO4- NaDR37Dengle springHot spring436.6837.272.366.814.61201.10919.337.53442.1Cl·HCO3-NaDR40Pola springHot spring306.6107.15.18.50.299.8090.543.6421.4Cl ·HCO3-NaDR42Qugu springHot spring478.9117.92.40.7<0.1268.920.94.212.2469.2HCO3-NaDR44Seluo springHot spring446.7195.040.2195.837.71105.20226.62.91893.1HCO3·cl-ca·NaDR45Cha springHot spring487.1476.426.670.99.21246.50200.528.32242.3HCO3-NaDR47QugulongbuHot spring666.9454.725.350.37.6616.90280.6250.81903.5HCO3·cl-NaDR48Shalei springHot spring757.7115.410.428.87378.507.259.2710.3HCO3-NaXZDR9TamaquzhenHot spring637.2407.431.64012.6856.30190.578.21919HCO3-NaXZDR11YinjianquzhenHot spring396.559.45.421020.9811.908.934.81113HCO3-CaXZDR15GabudanniHot spring677.0451.429.719.2<0.1584.40259.4169.91922HCO3·Cl-NaCKLSQ1Chazi VillageHot spring877.8481.939.112.8<0.1654.987.1165.1121.91806.9HCO3-NaCKLSQ2Chazi VillageHot spring747.9476.941.19.5<0.1707.358.1167.8117.51844.1HCO3-NaCQSHChazi VillageChaqu River water7.57.513.11.440.96.6115.901.4147.3279HCO3·SO4-CaCQXHChazi VillageChaqu River water25.58.6363.730.119.21.82424.1108.2118.1106.11386.1HCO3-NaDLZBXHChazi VillageDalong Zangbo River water6.58.38.12.544.54.978.41.13.175.3233.2SO4·HCO3-CaJ05Chazi VillageWell water47.55.72.3252.178.702.716.6115.6HCO3-CaDR28-1Lagejia springHot spring85.59.1365417.21.92612341651321570CO3·Cl -HCO3DR28-2Lagejia springHot spring788.4388.9350.60.2590.636112.11021712HCO3-NaDR28-3Lagejia springHot spring858.73503900447.2100.8151.777.81600HCO3-NaDR32-1Semi springHot spring868.77001043.44.249030667465.84340Cl·CO3·HCO3-NaDR42-1Qugu springHot spring458.91182.53.41.322745.713.928.8227HCO3·CO3·SO4·Cl-NaDR44-1Seluo springHot spring407.317533.320222.91040—157741380HCO3·Cl-Ca·NaDR45-1Cha springHot spring44.57.344326.573.812.51130—17232.91560HCO3·Cl-Na·CaDR47-1QugulongbuHot spring65.27442.526.553.29.4493—2722801600HCO3·Cl-Na·CaDR48-1Shalei springHot spring767.512512.534.38.338.1—9.626.8510HCO3-Na·CaCKLSQ-1Chazi VillageHot spring877.946845124.2801—1631441560HCO3·Cl·SO4-Na∗Note: DR28-1, DR28-2, DR28-3, DR32-1, DR42-1, DR44-1, DR45-1, DR47-1, DR48-1, and CKLSQ-1 are quoted from the report of Wei et al. [29].Table 2
Trace components of the water samples in the study area.
Field numberρB/(mg/L)HBO2SiO2Fe2+Fe3+F-NO3-NO2-PO43-LiSrAsDR28421.6286.9<0.05<0.0525.80.3<0.00330.11960.28.6DR322130536.5<0.05<0.0522.4<0.05<0.00330.11933.90.212.2DR3650.783.9<0.03<0.038.70.8<0.0010.01193.62.70.2DR37425.687.20.1<0.038.65.2<0.0010.0119142.84.9DR4020.455.1<0.05<0.053.81.5<0.00330.1190.60.20.5DR426.250.2<0.03<0.033.20.02<0.0010.01190.20.30.0006DR44181.634.60.2<0.030.90.1<0.0010.01196.21.80.0006DR45165.169.7<0.03<0.030.30.1<0.0010.011933.90.0006DR47200.368.40.3<0.036.40.1<0.0010.01192.826.2DR4826.676.7<0.03<0.030.20.1<0.0010.01190.31.80.0006XZDR9360.9170.8<0.050.199.40.06<0.00330.01194.61.25.6XZDR117.640<0.05<0.050.71.6<0.00330.01190.32.20.0002XZDR15416.7138.3<0.050.3910.5<0.05<0.00330.01196.61.16CKLSQ1116.3129.9<0.03<0.0315.50.1<0.0010.01192.20.70.7CKLSQ2120.7144<0.03<0.0316.90.1<0.0010.01192.20.70.9CQSH0.535.10.98<0.032.40.02<0.0010.01190.010.50.0049CQXH93.6104.30.26<0.0312.50.1<0.0010.01191.60.60.6DLZBXH1.815.7<0.03<0.030.30.03<0.0010.01190.040.30.0008J050.2———0.32.5<0.001———0.0014DR28-144022.7——25———7.3—5.4DR28-2207.887.5——30.4———5.4——DR28-3165.6120——27.5———12.3—5.5DR32-11750116.2——13.3———35—30.8DR42-1836——4.9———<0.1——DR44-116032.2——0.9———4.9——DR45-117357.8——0.8———2.8——DR47-118778.8——3.6———2.7—4DR48-12574.7——3.5———0.3——CKLSQ1-1170145——9.7———2.2—0.4The hydrochemistry of surface water in CQSH, CQXH, and DLZBXH was HCO3·SO4–Ca, HCO3–Na, and SO4·HCO3–Ca (Figure 2), respectively, with the main cations of Ca2+ and Na+, main anions of SO42− and HCO3−, and lower TDS than that of hot spring water. Table 1 presents the differences in the hydrochemical composition between hot and surface water. However, the hydrochemical composition of the water samples collected on the downstream of Chaqu was similar to that of hot water samples, which could explain the downstream water flowing through the hot water area to mix with the geothermal water due to short runoff distance.Figure 2
Chemical piper diagram of the water samples.


## 4.2. Source of Materials

### 4.2.1. Ionic Component

The Na+/Cl− ratio is considerable higher than 1 (Table 3), indicating strong dissolution and filtration in groundwater [37, 38]. Hot water exhibits strong dissolution for the surrounding rocks during upwelling, enabling Na+, K+, and Ca2+ in the potash feldspar and plagioclase to migrate into the hot water [10, 39]. The pH of the geothermal fluid was 7.83–7.93, a range in which the carbonic acid balance favoured HCO3−. The HCO3− content in the fluid samples collected in downstream Chaqu (CQXH) was drastically higher than that in upstream Chaqu presumably because the lateral runoff occurred during the vertical upwelling of geothermal water and was subsequently mixed into the river [40].Table 3
Ratios of Na and Cl in the study area.
Field numberGeothermal display areaSample typeρB/(mg/L)Na+/Cl-Na+Cl-DR28Lagejia springHot spring386.21582.4DR32Semi springHot spring790.5722.81.1DR36Charongqurene springHot spring564.7262.92.1DR37Dengle springHot spring837.2919.30.9DR40Pola springHot spring107.190.51.2DR42Qugu springHot spring117.94.228.3DR44Seluo springHot spring195226.60.9DR45Cha springHot spring476.4200.52.4DR47QugulongbuHot spring454.7280.61.6DR48Shalei springHot spring115.47.216.1XZDR9TamaquzhenHot spring407.4190.52.1XZDR11YinjianquzhenHot spring59.48.96.7XZDR15GabudanniHot spring451.4259.41.7CKLSQ1Chazi VillageHot spring481.9165.12.9CKLSQ2Chazi VillageHot spring476.9167.82.8CQSHChazi VillageChaqu River water13.11.49.3CQXHChazi VillageChaqu River water363.7118.13.1DLZBXHChazi VillageDalong Zangbo River water8.13.12.6J05Chazi VillageWell water5.72.82.1DR28-1Lagejia springHot spring3651652.2DR28-2Lagejia springHot spring388.9112.13.5DR28-3Lagejia springHot spring350151.72.3DR32-1Semi springHot spring7006741.0DR42-1Qugu springHot spring11813.98.5DR44-1Seluo springHot spring1751571.1DR45-1Cha springHot spring4431722.6DR47-1QugulongbuHot spring442.52721.6DR48-1Shalei springHot spring1259.613.1CKLSQ-1Chazi VillageHot spring4681632.9Almost all the thermal spring samples in the study area are observed below the 1 : 1 line (Figure3(a)), indicating that considerable ion replacement occurs during runoff. The low concentrations of Ca2+ and Mg2+ (Table 1) suggest that in the water, Ca2+ and Mg2+ are replaced with absorbed Na+ in the surrounding rock. The γNa/γ (Na+Ca+Mg) content of hot water is higher than that of cold water, and the value is close to 1 (Figure 3(c)), which indicates considerable ion exchange between Na+–Ca2+ and Na+–Mg2+ in geothermal fluids in this area. Ca2+ and Mg2+ available in the geothermal fluid are replaced with Na+ present on the rock and soil surfaces, which confirmed the aforementioned speculation. Almost all the fluids are distributed above the 2 : 1 line (Figure 3(b)), which indicates that the main mineral dissolved by carbonate in groundwater is calcite.Figure 3
Ratio diagrams of the ions of water samples in the study area.
(a)(b)(c)The high concentration of SO42− is a major result of the strong dissolution of geothermal fluids that occurs in the bottom strata of the Paleogene Pana formation (E2p) dominated by dacite, andesite, and andesitic and porphyroclastic lava according to the aforementioned discussion [41]. In the study area, the variation range of δ34S is small, and all fall within the range of atmospheric sulphate (Figure 4). The δ34S of DR32 and XZDR9 deviated from the sulphur of meteorites, suggesting that the sulphur of DR32 and XZDR9 is mainly leached from the surrounding rocks of the sedimentary cover. The δ34S of XZDR10 is close to that of meteorite sulphur, indicating that the sulphur source is located deep within the earth’s crust and may be regenerated magma produced by remelting in the crust. The mixing of deep sulphur is uniform, and no significant sulphur isotope differentiation can be observed in the geothermal fluid ascent. Some geothermal water samples located in the atmospheric sulphate range are distributed in the volcanic sulphur (SO2 and SO4) range. The geological conditions and the fact that no volcanic activity occurs in the study area indicate that the sulphur in geothermal fluids is caused by the dissolution of the surrounding rocks.Figure 4
δ34S-ρSO42− plot of geothermal water in the study area.The main sources of Cl− are human activities, deep magma, and leaching effects on rocks. Because this area is vast and not populous, with few Cl− minerals in the stratum, it is speculated that Cl− in geothermal fluids may come from deep magma. The Cl− concentration in the hot springs of the study area is approximately 200mg/L; however, that in the Semi spring and Dengle spring is higher (Table 1). The deep matter in these two hot springs is speculated to mix to a large extent. The low Cl− content of the Shalei spring distributed near the Yarlung Zangbo River may be caused by mixing with river water.


### 4.2.2. Trace Elements

In general, the F− concentration in the hot water in the study area is high, which is related to the presence of fluorine-containing minerals feldspar and biotite in the strata (Table 2) [42]. Under natural conditions, B and Li in the groundwater majorly originate from the dissolution of related minerals in rocks. The B and Li concentrations in geothermal fluids are positively correlated with the Cl− concentration (Figure 5), indicating that B, Li, and Cl− have the same material source. In addition to the dissolution of rock minerals, other sources are responsible for high B and Li concentrations [43–46]. Studies have revealed that B, Li, Rb, and As in the geothermal water in Tibet originate from the residual magma in the late-stage magmatism [47]. Considering the geothermal and geological conditions of this area, the rich content of Li and Sr in the hot water might be caused by the long-term leaching effect of deep geothermal fluids on volcanic rock formation (Paleogene Dianzhong Formation (E1-2d)).Figure 5
Correlation diagram between trace components and Cl− in the study area.


## 4.2.1. Ionic Component

The Na+/Cl− ratio is considerable higher than 1 (Table 3), indicating strong dissolution and filtration in groundwater [37, 38]. Hot water exhibits strong dissolution for the surrounding rocks during upwelling, enabling Na+, K+, and Ca2+ in the potash feldspar and plagioclase to migrate into the hot water [10, 39]. The pH of the geothermal fluid was 7.83–7.93, a range in which the carbonic acid balance favoured HCO3−. The HCO3− content in the fluid samples collected in downstream Chaqu (CQXH) was drastically higher than that in upstream Chaqu presumably because the lateral runoff occurred during the vertical upwelling of geothermal water and was subsequently mixed into the river [40].Table 3
Ratios of Na and Cl in the study area.
Field numberGeothermal display areaSample typeρB/(mg/L)Na+/Cl-Na+Cl-DR28Lagejia springHot spring386.21582.4DR32Semi springHot spring790.5722.81.1DR36Charongqurene springHot spring564.7262.92.1DR37Dengle springHot spring837.2919.30.9DR40Pola springHot spring107.190.51.2DR42Qugu springHot spring117.94.228.3DR44Seluo springHot spring195226.60.9DR45Cha springHot spring476.4200.52.4DR47QugulongbuHot spring454.7280.61.6DR48Shalei springHot spring115.47.216.1XZDR9TamaquzhenHot spring407.4190.52.1XZDR11YinjianquzhenHot spring59.48.96.7XZDR15GabudanniHot spring451.4259.41.7CKLSQ1Chazi VillageHot spring481.9165.12.9CKLSQ2Chazi VillageHot spring476.9167.82.8CQSHChazi VillageChaqu River water13.11.49.3CQXHChazi VillageChaqu River water363.7118.13.1DLZBXHChazi VillageDalong Zangbo River water8.13.12.6J05Chazi VillageWell water5.72.82.1DR28-1Lagejia springHot spring3651652.2DR28-2Lagejia springHot spring388.9112.13.5DR28-3Lagejia springHot spring350151.72.3DR32-1Semi springHot spring7006741.0DR42-1Qugu springHot spring11813.98.5DR44-1Seluo springHot spring1751571.1DR45-1Cha springHot spring4431722.6DR47-1QugulongbuHot spring442.52721.6DR48-1Shalei springHot spring1259.613.1CKLSQ-1Chazi VillageHot spring4681632.9Almost all the thermal spring samples in the study area are observed below the 1 : 1 line (Figure3(a)), indicating that considerable ion replacement occurs during runoff. The low concentrations of Ca2+ and Mg2+ (Table 1) suggest that in the water, Ca2+ and Mg2+ are replaced with absorbed Na+ in the surrounding rock. The γNa/γ (Na+Ca+Mg) content of hot water is higher than that of cold water, and the value is close to 1 (Figure 3(c)), which indicates considerable ion exchange between Na+–Ca2+ and Na+–Mg2+ in geothermal fluids in this area. Ca2+ and Mg2+ available in the geothermal fluid are replaced with Na+ present on the rock and soil surfaces, which confirmed the aforementioned speculation. Almost all the fluids are distributed above the 2 : 1 line (Figure 3(b)), which indicates that the main mineral dissolved by carbonate in groundwater is calcite.Figure 3
Ratio diagrams of the ions of water samples in the study area.
(a)(b)(c)The high concentration of SO42− is a major result of the strong dissolution of geothermal fluids that occurs in the bottom strata of the Paleogene Pana formation (E2p) dominated by dacite, andesite, and andesitic and porphyroclastic lava according to the aforementioned discussion [41]. In the study area, the variation range of δ34S is small, and all fall within the range of atmospheric sulphate (Figure 4). The δ34S of DR32 and XZDR9 deviated from the sulphur of meteorites, suggesting that the sulphur of DR32 and XZDR9 is mainly leached from the surrounding rocks of the sedimentary cover. The δ34S of XZDR10 is close to that of meteorite sulphur, indicating that the sulphur source is located deep within the earth’s crust and may be regenerated magma produced by remelting in the crust. The mixing of deep sulphur is uniform, and no significant sulphur isotope differentiation can be observed in the geothermal fluid ascent. Some geothermal water samples located in the atmospheric sulphate range are distributed in the volcanic sulphur (SO2 and SO4) range. The geological conditions and the fact that no volcanic activity occurs in the study area indicate that the sulphur in geothermal fluids is caused by the dissolution of the surrounding rocks.Figure 4
δ34S-ρSO42− plot of geothermal water in the study area.The main sources of Cl− are human activities, deep magma, and leaching effects on rocks. Because this area is vast and not populous, with few Cl− minerals in the stratum, it is speculated that Cl− in geothermal fluids may come from deep magma. The Cl− concentration in the hot springs of the study area is approximately 200mg/L; however, that in the Semi spring and Dengle spring is higher (Table 1). The deep matter in these two hot springs is speculated to mix to a large extent. The low Cl− content of the Shalei spring distributed near the Yarlung Zangbo River may be caused by mixing with river water.


## 4.2.2. Trace Elements

In general, the F− concentration in the hot water in the study area is high, which is related to the presence of fluorine-containing minerals feldspar and biotite in the strata (Table 2) [42]. Under natural conditions, B and Li in the groundwater majorly originate from the dissolution of related minerals in rocks. The B and Li concentrations in geothermal fluids are positively correlated with the Cl− concentration (Figure 5), indicating that B, Li, and Cl− have the same material source. In addition to the dissolution of rock minerals, other sources are responsible for high B and Li concentrations [43–46]. Studies have revealed that B, Li, Rb, and As in the geothermal water in Tibet originate from the residual magma in the late-stage magmatism [47]. Considering the geothermal and geological conditions of this area, the rich content of Li and Sr in the hot water might be caused by the long-term leaching effect of deep geothermal fluids on volcanic rock formation (Paleogene Dianzhong Formation (E1-2d)).Figure 5
Correlation diagram between trace components and Cl− in the study area.


## 4.3. Isotope Characteristics of the Geothermal Fluids

### 4.3.1. Source of Recharge

Hydrogen and oxygen isotopes are typically used to trace the recharge source for the groundwater, determine the intensity of groundwater runoff, and identify the water-rock interaction between the groundwater runoff and surrounding rocks [48–50]. Craig analysed the δ2H and δ18O values of more than 400 natural water samples on a global scale and obtained the global meteoric water line: δ2H=8δ18O+10 [51]. Scholars have often used the positional relation between the hydrogen and oxygen stable isotope data in the study area to determine the replenishment source of underground fluids [51].In this study, the hydrogen and oxygen isotope data measured for 6 samples and that collected for 5 samples (Table4) were used to analyse the recharge source and recharge elevation of the geothermal fluid in the area. Figure 6 reveals that all the hydrogen and oxygen isotopes were distributed near the global atmospheric rainfall equation line [51], indicating that the main recharge source for the underground hot water system was atmospheric rainfall. Moreover, the δ18O and δ2H of the geothermal fluid and surface water did not differ considerably, which indicated that the recharge source of the underground hot and surface water comes from the surrounding atmospheric precipitation infiltration and not through long-distance migration; that is, the water–rock interaction is not strong [52, 53]. Another possible reason is that the geothermal fluids in the study area were mixed with surface cold water during the rising process.Table 4
Data ofδ2H–δ18O of the geothermal fluid in the study area.
Field numberGeothermal display areaSample typeδ2HVSMOW(‰)δ18OVSMOW (‰)Data sourcesDR28-1Lagejia springHot spring-149-18.6Quoted from literature [47]DR28-2Lagejia springHot spring-154-18.6DR28-3Lagejia springHot spring-153-18.6DR28-4Xiaolupang springHot spring-155-18.6CMQChangmaqu River waterRiver water-130-18.3CKLSQ1Chazi VillageHot spring-163-20.5MeasuredCKLSQ2Chazi VillageHot spring-163-20.6CQSHChazi VillageRiver water-136-17.6CQXHChazi VillageRiver water-148-18.7DLZBXHChazi VillageRiver water-164-20.9Figure 6
δ2H–δ18O isotope relationship of various water samples.


### 4.3.2. Recharge Elevation

Theδ2H and δ18O of atmospheric precipitation often exhibit the characteristics of temperatural, continental, seasonal fluctuation, etc. The following method can be used to estimate the replenishment elevation of each hot spring in the study area based on the elevation effect of the stable isotope of oxygen.
(1)H=δG−δPK+h,where H is the recharge elevation, m; h is the elevation of the sampling point, m; δG is the δ18O of the hot spring, ‰; δP is the δ18O of the atmospheric precipitation, ‰; K is the elevation gradient of the δ18O of the atmospheric precipitation, δ/100 m. In this paper, the CQSH data (δ18O is −17.6‰) at the upstream point of Chaqu is used as the δ18O value of atmospheric precipitation in Chazi Hot Spring, and the CMQ (δ18O is −18.29‰) at the Changmaqu River water point is used as the δ18O value of atmospheric precipitation in the Gejia Hot Spring. The elevation gradient of δ18O in Tibet (−0.26‰/100 m) is used as the K value [54]. The calculation results are presented in Table 5.Table 5
Calculated values of the geothermal fluid replenishment elevation in the study area.
Field numberCKLSQ1CKLSQ2DR28-1DR28-2DR28-3DR28-4Sampling elevation484148375054505850475050Replenishment elevation595659905165518851775184The geothermal fluid replenishment elevation range of the study area is 5200–6000 m (Table5). This estimated value is close to the elevation data of the sampling point. It is inferred that the replenishment of geothermal water mainly comes from atmospheric rainfall infiltration replenishment received by the mountains near the hot spring outcrop area and a small amount of alpine snowmelt replenishment.


## 4.3.1. Source of Recharge

Hydrogen and oxygen isotopes are typically used to trace the recharge source for the groundwater, determine the intensity of groundwater runoff, and identify the water-rock interaction between the groundwater runoff and surrounding rocks [48–50]. Craig analysed the δ2H and δ18O values of more than 400 natural water samples on a global scale and obtained the global meteoric water line: δ2H=8δ18O+10 [51]. Scholars have often used the positional relation between the hydrogen and oxygen stable isotope data in the study area to determine the replenishment source of underground fluids [51].In this study, the hydrogen and oxygen isotope data measured for 6 samples and that collected for 5 samples (Table4) were used to analyse the recharge source and recharge elevation of the geothermal fluid in the area. Figure 6 reveals that all the hydrogen and oxygen isotopes were distributed near the global atmospheric rainfall equation line [51], indicating that the main recharge source for the underground hot water system was atmospheric rainfall. Moreover, the δ18O and δ2H of the geothermal fluid and surface water did not differ considerably, which indicated that the recharge source of the underground hot and surface water comes from the surrounding atmospheric precipitation infiltration and not through long-distance migration; that is, the water–rock interaction is not strong [52, 53]. Another possible reason is that the geothermal fluids in the study area were mixed with surface cold water during the rising process.Table 4
Data ofδ2H–δ18O of the geothermal fluid in the study area.
Field numberGeothermal display areaSample typeδ2HVSMOW(‰)δ18OVSMOW (‰)Data sourcesDR28-1Lagejia springHot spring-149-18.6Quoted from literature [47]DR28-2Lagejia springHot spring-154-18.6DR28-3Lagejia springHot spring-153-18.6DR28-4Xiaolupang springHot spring-155-18.6CMQChangmaqu River waterRiver water-130-18.3CKLSQ1Chazi VillageHot spring-163-20.5MeasuredCKLSQ2Chazi VillageHot spring-163-20.6CQSHChazi VillageRiver water-136-17.6CQXHChazi VillageRiver water-148-18.7DLZBXHChazi VillageRiver water-164-20.9Figure 6
δ2H–δ18O isotope relationship of various water samples.


## 4.3.2. Recharge Elevation

Theδ2H and δ18O of atmospheric precipitation often exhibit the characteristics of temperatural, continental, seasonal fluctuation, etc. The following method can be used to estimate the replenishment elevation of each hot spring in the study area based on the elevation effect of the stable isotope of oxygen.
(1)H=δG−δPK+h,where H is the recharge elevation, m; h is the elevation of the sampling point, m; δG is the δ18O of the hot spring, ‰; δP is the δ18O of the atmospheric precipitation, ‰; K is the elevation gradient of the δ18O of the atmospheric precipitation, δ/100 m. In this paper, the CQSH data (δ18O is −17.6‰) at the upstream point of Chaqu is used as the δ18O value of atmospheric precipitation in Chazi Hot Spring, and the CMQ (δ18O is −18.29‰) at the Changmaqu River water point is used as the δ18O value of atmospheric precipitation in the Gejia Hot Spring. The elevation gradient of δ18O in Tibet (−0.26‰/100 m) is used as the K value [54]. The calculation results are presented in Table 5.Table 5
Calculated values of the geothermal fluid replenishment elevation in the study area.
Field numberCKLSQ1CKLSQ2DR28-1DR28-2DR28-3DR28-4Sampling elevation484148375054505850475050Replenishment elevation595659905165518851775184The geothermal fluid replenishment elevation range of the study area is 5200–6000 m (Table5). This estimated value is close to the elevation data of the sampling point. It is inferred that the replenishment of geothermal water mainly comes from atmospheric rainfall infiltration replenishment received by the mountains near the hot spring outcrop area and a small amount of alpine snowmelt replenishment.


## 4.4. Temperature of Geothermal Reservoirs and Cycle Depth Estimation

### 4.4.1. Geochemical Geothermometer for the Temperature Estimation of Geothermal Reservoirs

The geothermometer is an effective method for estimating the temperature of thermal storage, and each temperature scale has certain assumptions. Therefore, when using the geothermometers to study the temperature of the thermal reservoir, the applicable scope and conditions of each temperature scale should be considered to select a suitable geothermometer to estimate the thermal storage temperature. The commonly used geothermometers mainly include cationic geochemistry, SiO2 geothermometer, isotope geothermometer, and gas geothermometer, among which the most commonly used are cationic geochemistry and SiO2 geothermometers. According to the hydrochemical characteristics of the geothermal fluid in the study area, the geothermal geology, and the applicable conditions of the geothermal geothermometers (Table 6), this study used the SiO2 geothermometer (quartz, chalcedony) and Na-K geothermometer to estimate thermal storage temperature for the study area [55–57]. Because the study region is located in a high-altitude area in Tibet, the boiling point of water is low. In this study, we designated the hot springs with exposure temperatures of >80°C as boiling springs. That is, when the exposure temperature of the sampling point was ≥80°C, the maximum steam loss formula was used when the SiO2 geothermometer was used to estimate thermal storage temperature. When the exposure temperature of the sampling point did not reach 80°C, the no-steam loss formula was used. The calculation results are shown in Table 7.Table 6
Geothermometer function of geothermal water in the study area.
GeothermometersCalculation formulaScope of applicationQuartz1T=13095.19−logSiO2−273.15No loss of steamQuartz2T=15225.75−logSiO2−273.15Maximum steam loss 0-250°CChalcedony1T=10324.69−logSiO2−273.15High temperature thermal storageChalcedony2T=12645.31−logSiO2−273.15Maximum steam lossNa-KT=1217logNa/K+1.483−273.15t>150°CTable 7
Estimated results of the reservoir temperature in the study area.
Field numberSampling temperature/°CSiO2 geothCation geothNa-K/°CQuartz geoth (minimum steam loss)/°CQuartz geoth (maximum steam loss)/°CChalcedony geoth (minimum steam loss)/°CChalcedony geoth (maximum steam loss)/°CDR2885—189.2—170224.8DR3289—230.8—216.7251.8DR3665127.6—99.9—234.7DR3743129.7—102.1—220.8DR4030106.4—76.8—179.8DR4247102—72.1—131.6DR444485.4—54.4—297.4DR4548117.9—89.3—189.6DR4766117—88.3—189.3DR4875122.9—94.7—223.9XZDR963169.4—146.8—212.7XZDR113991.7—61.1—224.9XZDR1567156.2—131.7—201CKLSQ187—145.4—122.3216.2CKLSQ274158.7—134.5—220.6The reservoir temperature calculated using chalcedony (minimum steam loss) is not the true reflection of the hydrothermal system because the estimated values are not substantially different from the reservoir temperature of the thermal springs (Table7). The reservoir temperatures estimated using quartz and chalcedony geothermometer (maximum steam loss) of DR28, DR32, and CKLSQ1 are the same, indicating that the temperatures acquired using this formula are highly reliable (Table 7). For other sample points, the reservoir temperature calculated using quartz geothermometer (no steam loss) is mostly distributed in 120–160°C, with an average of 126°C. The estimated values are slightly low combined with the geothermal background. The estimated value of Na–K geothermometer is higher than that of SiO2 geothermometer. The low value of SiO2 geothermometer may be caused by the following reasons: (1) the hot spring fluid was mixed with varying degrees of cold water during the ascent. (2) Saturated SiO2 dissolved in hot water above 180°C precipitated [48]. From the Na–K–Mg triangle diagram (Figure 7), the groundwater in the study area had not reached a water–rock equilibrium state, and cold water was mixed in the geothermal fluid. This is the probable reason for the low reservoir temperature estimation by using SiO2 geothermometer. Most thermal springs are located near the partial equilibrium zone, which shows that the method of reservoir temperature estimation by using the Na–K geothermometer presents certain reference significance. The estimation results acquired using the Na-K geothermometer are considered to reflect deep reservoir characteristics, and those obtained by employing SiO2 geothermometer are considered to reflect shallow reservoir characteristics combined with the applicable conditions of Na–K geothermometer.Figure 7
Na–K–Mg triangle diagram of the geothermal fluid in the study area.


### 4.4.2. Silica-Enthalpy Mixing Model to Estimate the Temperature of Geothermal Reservoirs

According to the aforementioned analysis, the hot water in the study area is likely to mix with the shallow cold water during the ascent. A strong linear relationship exists between Cl and B in the geothermal fluids (Figure8). CQSH is the upstream sample, and CQXH is the downstream sample. Most geothermal fluid samples are located at the cold water end (Figure 8). No significant differences can be observed in the compositions of δ2H and δ18O of the geothermal fluids (Figure 6). Therefore, the geothermal fluid in this area can be considered a mixture of deep geothermal fluids and cold water [58, 59].Figure 8
Relationship between Cl and B contents in the study area.In 1974, Fournier and Truesdell developed a silica-enthalpy equation method to estimate the true thermal storage temperature before the cold water is mixed. The principle of this method is that the mixing of hot and cold water inevitably causes the initial enthalpy and SiO2 content of deep water to decrease to the final enthalpy and SiO2 content of spring water. According to this principle, the model equation is derived:
(2)ScX1+Sh1−X1=Ss,(3)ρcSilicaX2+ρhSilicaX21−X2=ρsSilica,where Sc is the enthalpy of cold water (J/g); Ss is the final enthalpy of spring water (J/g), below 100°C, and the enthalpy of saturated water is equal to the water Celsius degree. For >100°C, the relationship between temperature and enthalpy of saturated water is presented in Table 8. Sh is the initial enthalpy of hot water (J/g); ρcSilica is the silica mass concentration of cold water (mg/L); ρsSilica is the silica mass concentration of the spring water (mg/L); ρhSilica is the initial silica mass concentration of hot water (mg/L); and X is the mixing proportion of cold water.Table 8
Relationship between water temperature, enthalpy, and mass concentration of SiO2.
t/°CenthalPy (∗4.1868J/g)ρcSiO2/(mg/L)505013.5757526.6100100.148125125.180150151125175177185200203.6265225230.9365250259.2486275289614300321692In this study, the enthalpy of cold water was determined to be the average temperature of the study area at 5°C in the year, and the final enthalpy of the hot spring was the enthalpy value corresponding to the measured temperature value. The silica mass concentration of cold water (upstream of Chaqu) was 35.05 mg/L, and the mass concentration of SiO2 in the hot spring takes the actual measured value this time (Table 2). Substitute the concentration into the model equation to obtain the values of X1 and X2 for each hot spring point (Table 9), and then, project the calculated values of X1 and X2 and their corresponding temperatures into the scatter diagram, two curves can be drawn. The abscissa corresponding to the intersection of the two curves is the estimated value of heat storage temperature before mixing, and the ordinate corresponding to the intersection of the two curves is the mixing ratio of cold water (Figure 9).Table 9
Calculated values ofX1 and X2 hot springs in the study area.
Temperature/°C5075100125150175200225250275300DR28X1-0.78-0.140.160.330.450.530.60.650.690.720.75X212.6930.81-18.45-4.6-1.8-0.68-0.10.240.440.560.62DR32X1-0.87-0.20.120.30.420.510.580.630.670.70.73X224.2760.35-37.73-10.16-4.58-2.34-1.18-0.52-0.110.130.24DR36X1-0.330.140.370.50.590.650.70.730.760.790.81X23.276.78-2.77-0.090.460.670.790.850.890.920.93DR37X10.160.460.60.680.740.780.810.830.850.870.88X23.427.17-3.02-0.160.420.650.770.840.880.910.92DR40X10.440.640.740.790.830.850.870.890.90.910.92X21.933.37-0.540.560.780.870.910.940.960.970.97DR42X10.070.40.560.650.710.760.790.810.830.850.87X21.72.79-0.170.660.830.90.930.950.970.970.98DR44X10.130.440.590.680.730.770.80.830.850.860.88X20.980.951.031.011111111DR45X10.040.390.550.640.710.750.780.810.830.850.86X22.615.1-1.670.230.620.770.850.90.920.940.95DR47X1-0.360.130.360.490.580.650.690.730.760.790.81X22.554.95-1.580.260.630.780.850.90.930.940.95DR48X1-0.5600.260.420.520.590.650.690.720.750.78X22.935.93-2.220.070.540.720.820.870.910.930.94XZDR9X1-0.290.170.390.520.60.660.710.740.770.80.82X27.317.06-9.48-2.02-0.510.090.410.590.70.770.79XZDR11X10.240.510.640.720.770.80.830.850.870.880.89X21.231.590.620.890.940.970.980.980.990.990.99XZDR15X1-0.380.110.350.480.580.640.690.730.760.780.8X25.7913.22-6.97-1.3-0.150.310.550.690.770.820.84CKLSQ1X1-0.82-0.170.140.320.440.520.590.640.680.710.74X25.412.22-6.32-1.11-0.050.370.590.710.790.840.86CKLSQ2X1-0.530.010.270.430.530.60.650.690.730.760.78X26.0613.9-7.42-1.42-0.210.270.530.670.760.810.83Figure 9
Si-enthalpy model of the study area.The mixing ratio of cold water in the hot springs in the study area is relatively large, and the heat storage temperature before the cold water mixed is generally above 150 °C, which indicates the characteristics of high-temperature heat storage in the study area (Figure9). The abnormalities are the five hot springs: DR28, DR32, DR42, DR44, and XZDR9. They present no intersection in Figure 5, indicating that these five hot springs are affected by the absence of cold water mixing. The thermal storage temperatures of DR36, DR37, DR40, DR45, DR47, DR48, XZDR11, XZDR15, CKLSQ1, and CKLSQ2 before the mixing of cold water are 170°C, 217°C, 169°C, 169°C, 145°C, 149°C, 104°C, 241°C, 200°C, and 234°C, respectively, and their mixing ratios of cold water are 64%, 82%, 85%, 74%, 57%, 52%, 65%, 74%, 59%, and 71%, respectively.In summary, the average mixing ratio of hot spring and cold water in the study area is 75%, and the average value of the heat storage temperature before the cold water mixed is 200°C, which is basically the same as the heat storage temperature obtained in the Na-K-Mg balance diagram and the result of the Na-K geothermometer. This reflects the real storage temperature before the cold water is mixed in the study area, that is, the deep heat is stored in the study area. The calculation results of the SiO2 (maximum steam loss) geothermometer are distributed around 150°C, indicating shallow heat storage in the study area.


### 4.4.3. Circulation Depth

The foregoing hydrogen and oxygen isotope research results showed that the geothermal fluid in the study area is mainly of atmospheric rainfall replenishment, and the temperature of hot water chiefly depends on deep-circulation geothermal heating. Therefore, the calculation of the temperature of underground hot water storage can be used to estimate the cycle depth of the hot water with the following formula:(4)Z=GTz−T0+Z0,where Z is the calculated cycle depth, m; G is the geothermal heating gradient, m/°C, with a value of 6.67 m/°C [60]; TZ is the heat storage temperature, °C, taking the heat storage temperature determined using an existing mixed model; T0 is the multiyear average temperature of the replenishment area, °C, its value is 5 °C; and Z0 is the depth of the normal temperature zone, m, the value is 20 m. The calculation results of the cycle depth are shown in Table 10.Table 10
Calculation of underground hot water circulation depth in in the study area.
Field numberSampling point elevation/mReservoir temperature/°CCycle depth/mDR285054202.11335DR324262237.71572DR3648861701121DR3752592121401DR4049231601054DR424382109.4716.4DR444279183.61212DR4541571631074DR474173144947.1DR484170148973.8XZDR95030185.41223XZDR114377102667XZDR1547332361561CKLSQ4839216.51431The circulation depths of DR42, DR47, DR48, and XZDR11 hot springs are shallow, ranging from 667 to 974 m (Table10). The circulation depths of all the other hot springs are >1000 m. DR32 and XZDR15 hot springs show the maximum circulation depth of 1500 m. Overall, the average circulation depth of the study area is 1163.38 m, and the geothermal conditions are satisfactory.


## 4.4.1. Geochemical Geothermometer for the Temperature Estimation of Geothermal Reservoirs

The geothermometer is an effective method for estimating the temperature of thermal storage, and each temperature scale has certain assumptions. Therefore, when using the geothermometers to study the temperature of the thermal reservoir, the applicable scope and conditions of each temperature scale should be considered to select a suitable geothermometer to estimate the thermal storage temperature. The commonly used geothermometers mainly include cationic geochemistry, SiO2 geothermometer, isotope geothermometer, and gas geothermometer, among which the most commonly used are cationic geochemistry and SiO2 geothermometers. According to the hydrochemical characteristics of the geothermal fluid in the study area, the geothermal geology, and the applicable conditions of the geothermal geothermometers (Table 6), this study used the SiO2 geothermometer (quartz, chalcedony) and Na-K geothermometer to estimate thermal storage temperature for the study area [55–57]. Because the study region is located in a high-altitude area in Tibet, the boiling point of water is low. In this study, we designated the hot springs with exposure temperatures of >80°C as boiling springs. That is, when the exposure temperature of the sampling point was ≥80°C, the maximum steam loss formula was used when the SiO2 geothermometer was used to estimate thermal storage temperature. When the exposure temperature of the sampling point did not reach 80°C, the no-steam loss formula was used. The calculation results are shown in Table 7.Table 6
Geothermometer function of geothermal water in the study area.
GeothermometersCalculation formulaScope of applicationQuartz1T=13095.19−logSiO2−273.15No loss of steamQuartz2T=15225.75−logSiO2−273.15Maximum steam loss 0-250°CChalcedony1T=10324.69−logSiO2−273.15High temperature thermal storageChalcedony2T=12645.31−logSiO2−273.15Maximum steam lossNa-KT=1217logNa/K+1.483−273.15t>150°CTable 7
Estimated results of the reservoir temperature in the study area.
Field numberSampling temperature/°CSiO2 geothCation geothNa-K/°CQuartz geoth (minimum steam loss)/°CQuartz geoth (maximum steam loss)/°CChalcedony geoth (minimum steam loss)/°CChalcedony geoth (maximum steam loss)/°CDR2885—189.2—170224.8DR3289—230.8—216.7251.8DR3665127.6—99.9—234.7DR3743129.7—102.1—220.8DR4030106.4—76.8—179.8DR4247102—72.1—131.6DR444485.4—54.4—297.4DR4548117.9—89.3—189.6DR4766117—88.3—189.3DR4875122.9—94.7—223.9XZDR963169.4—146.8—212.7XZDR113991.7—61.1—224.9XZDR1567156.2—131.7—201CKLSQ187—145.4—122.3216.2CKLSQ274158.7—134.5—220.6The reservoir temperature calculated using chalcedony (minimum steam loss) is not the true reflection of the hydrothermal system because the estimated values are not substantially different from the reservoir temperature of the thermal springs (Table7). The reservoir temperatures estimated using quartz and chalcedony geothermometer (maximum steam loss) of DR28, DR32, and CKLSQ1 are the same, indicating that the temperatures acquired using this formula are highly reliable (Table 7). For other sample points, the reservoir temperature calculated using quartz geothermometer (no steam loss) is mostly distributed in 120–160°C, with an average of 126°C. The estimated values are slightly low combined with the geothermal background. The estimated value of Na–K geothermometer is higher than that of SiO2 geothermometer. The low value of SiO2 geothermometer may be caused by the following reasons: (1) the hot spring fluid was mixed with varying degrees of cold water during the ascent. (2) Saturated SiO2 dissolved in hot water above 180°C precipitated [48]. From the Na–K–Mg triangle diagram (Figure 7), the groundwater in the study area had not reached a water–rock equilibrium state, and cold water was mixed in the geothermal fluid. This is the probable reason for the low reservoir temperature estimation by using SiO2 geothermometer. Most thermal springs are located near the partial equilibrium zone, which shows that the method of reservoir temperature estimation by using the Na–K geothermometer presents certain reference significance. The estimation results acquired using the Na-K geothermometer are considered to reflect deep reservoir characteristics, and those obtained by employing SiO2 geothermometer are considered to reflect shallow reservoir characteristics combined with the applicable conditions of Na–K geothermometer.Figure 7
Na–K–Mg triangle diagram of the geothermal fluid in the study area.


## 4.4.2. Silica-Enthalpy Mixing Model to Estimate the Temperature of Geothermal Reservoirs

According to the aforementioned analysis, the hot water in the study area is likely to mix with the shallow cold water during the ascent. A strong linear relationship exists between Cl and B in the geothermal fluids (Figure8). CQSH is the upstream sample, and CQXH is the downstream sample. Most geothermal fluid samples are located at the cold water end (Figure 8). No significant differences can be observed in the compositions of δ2H and δ18O of the geothermal fluids (Figure 6). Therefore, the geothermal fluid in this area can be considered a mixture of deep geothermal fluids and cold water [58, 59].Figure 8
Relationship between Cl and B contents in the study area.In 1974, Fournier and Truesdell developed a silica-enthalpy equation method to estimate the true thermal storage temperature before the cold water is mixed. The principle of this method is that the mixing of hot and cold water inevitably causes the initial enthalpy and SiO2 content of deep water to decrease to the final enthalpy and SiO2 content of spring water. According to this principle, the model equation is derived:
(2)ScX1+Sh1−X1=Ss,(3)ρcSilicaX2+ρhSilicaX21−X2=ρsSilica,where Sc is the enthalpy of cold water (J/g); Ss is the final enthalpy of spring water (J/g), below 100°C, and the enthalpy of saturated water is equal to the water Celsius degree. For >100°C, the relationship between temperature and enthalpy of saturated water is presented in Table 8. Sh is the initial enthalpy of hot water (J/g); ρcSilica is the silica mass concentration of cold water (mg/L); ρsSilica is the silica mass concentration of the spring water (mg/L); ρhSilica is the initial silica mass concentration of hot water (mg/L); and X is the mixing proportion of cold water.Table 8
Relationship between water temperature, enthalpy, and mass concentration of SiO2.
t/°CenthalPy (∗4.1868J/g)ρcSiO2/(mg/L)505013.5757526.6100100.148125125.180150151125175177185200203.6265225230.9365250259.2486275289614300321692In this study, the enthalpy of cold water was determined to be the average temperature of the study area at 5°C in the year, and the final enthalpy of the hot spring was the enthalpy value corresponding to the measured temperature value. The silica mass concentration of cold water (upstream of Chaqu) was 35.05 mg/L, and the mass concentration of SiO2 in the hot spring takes the actual measured value this time (Table 2). Substitute the concentration into the model equation to obtain the values of X1 and X2 for each hot spring point (Table 9), and then, project the calculated values of X1 and X2 and their corresponding temperatures into the scatter diagram, two curves can be drawn. The abscissa corresponding to the intersection of the two curves is the estimated value of heat storage temperature before mixing, and the ordinate corresponding to the intersection of the two curves is the mixing ratio of cold water (Figure 9).Table 9
Calculated values ofX1 and X2 hot springs in the study area.
Temperature/°C5075100125150175200225250275300DR28X1-0.78-0.140.160.330.450.530.60.650.690.720.75X212.6930.81-18.45-4.6-1.8-0.68-0.10.240.440.560.62DR32X1-0.87-0.20.120.30.420.510.580.630.670.70.73X224.2760.35-37.73-10.16-4.58-2.34-1.18-0.52-0.110.130.24DR36X1-0.330.140.370.50.590.650.70.730.760.790.81X23.276.78-2.77-0.090.460.670.790.850.890.920.93DR37X10.160.460.60.680.740.780.810.830.850.870.88X23.427.17-3.02-0.160.420.650.770.840.880.910.92DR40X10.440.640.740.790.830.850.870.890.90.910.92X21.933.37-0.540.560.780.870.910.940.960.970.97DR42X10.070.40.560.650.710.760.790.810.830.850.87X21.72.79-0.170.660.830.90.930.950.970.970.98DR44X10.130.440.590.680.730.770.80.830.850.860.88X20.980.951.031.011111111DR45X10.040.390.550.640.710.750.780.810.830.850.86X22.615.1-1.670.230.620.770.850.90.920.940.95DR47X1-0.360.130.360.490.580.650.690.730.760.790.81X22.554.95-1.580.260.630.780.850.90.930.940.95DR48X1-0.5600.260.420.520.590.650.690.720.750.78X22.935.93-2.220.070.540.720.820.870.910.930.94XZDR9X1-0.290.170.390.520.60.660.710.740.770.80.82X27.317.06-9.48-2.02-0.510.090.410.590.70.770.79XZDR11X10.240.510.640.720.770.80.830.850.870.880.89X21.231.590.620.890.940.970.980.980.990.990.99XZDR15X1-0.380.110.350.480.580.640.690.730.760.780.8X25.7913.22-6.97-1.3-0.150.310.550.690.770.820.84CKLSQ1X1-0.82-0.170.140.320.440.520.590.640.680.710.74X25.412.22-6.32-1.11-0.050.370.590.710.790.840.86CKLSQ2X1-0.530.010.270.430.530.60.650.690.730.760.78X26.0613.9-7.42-1.42-0.210.270.530.670.760.810.83Figure 9
Si-enthalpy model of the study area.The mixing ratio of cold water in the hot springs in the study area is relatively large, and the heat storage temperature before the cold water mixed is generally above 150 °C, which indicates the characteristics of high-temperature heat storage in the study area (Figure9). The abnormalities are the five hot springs: DR28, DR32, DR42, DR44, and XZDR9. They present no intersection in Figure 5, indicating that these five hot springs are affected by the absence of cold water mixing. The thermal storage temperatures of DR36, DR37, DR40, DR45, DR47, DR48, XZDR11, XZDR15, CKLSQ1, and CKLSQ2 before the mixing of cold water are 170°C, 217°C, 169°C, 169°C, 145°C, 149°C, 104°C, 241°C, 200°C, and 234°C, respectively, and their mixing ratios of cold water are 64%, 82%, 85%, 74%, 57%, 52%, 65%, 74%, 59%, and 71%, respectively.In summary, the average mixing ratio of hot spring and cold water in the study area is 75%, and the average value of the heat storage temperature before the cold water mixed is 200°C, which is basically the same as the heat storage temperature obtained in the Na-K-Mg balance diagram and the result of the Na-K geothermometer. This reflects the real storage temperature before the cold water is mixed in the study area, that is, the deep heat is stored in the study area. The calculation results of the SiO2 (maximum steam loss) geothermometer are distributed around 150°C, indicating shallow heat storage in the study area.


## 4.4.3. Circulation Depth

The foregoing hydrogen and oxygen isotope research results showed that the geothermal fluid in the study area is mainly of atmospheric rainfall replenishment, and the temperature of hot water chiefly depends on deep-circulation geothermal heating. Therefore, the calculation of the temperature of underground hot water storage can be used to estimate the cycle depth of the hot water with the following formula:(4)Z=GTz−T0+Z0,where Z is the calculated cycle depth, m; G is the geothermal heating gradient, m/°C, with a value of 6.67 m/°C [60]; TZ is the heat storage temperature, °C, taking the heat storage temperature determined using an existing mixed model; T0 is the multiyear average temperature of the replenishment area, °C, its value is 5 °C; and Z0 is the depth of the normal temperature zone, m, the value is 20 m. The calculation results of the cycle depth are shown in Table 10.Table 10
Calculation of underground hot water circulation depth in in the study area.
Field numberSampling point elevation/mReservoir temperature/°CCycle depth/mDR285054202.11335DR324262237.71572DR3648861701121DR3752592121401DR4049231601054DR424382109.4716.4DR444279183.61212DR4541571631074DR474173144947.1DR484170148973.8XZDR95030185.41223XZDR114377102667XZDR1547332361561CKLSQ4839216.51431The circulation depths of DR42, DR47, DR48, and XZDR11 hot springs are shallow, ranging from 667 to 974 m (Table10). The circulation depths of all the other hot springs are >1000 m. DR32 and XZDR15 hot springs show the maximum circulation depth of 1500 m. Overall, the average circulation depth of the study area is 1163.38 m, and the geothermal conditions are satisfactory.


## 5. Conceptual Model of the Heat Reservoir in the Research Area

### 5.1. Heat Source in the Study Area

Geophysical surveys showed AMT measurement anomalies, gravity anomalies, and magnetic anomalies present in the study area. The gravity field in the area is high and low in the south and north, respectively. The overall appearance is low gravity anomalies, and high gravity anomalies are in some areas. The magnetic anomalies are generally distributed in NE–SW orientation. The Cenozoic, neotectonic activities, mainly composed of active faults, fault basins (fault-sag belts), fault-block mountains (uplift belts), and accompanying hydrothermal phenomena (hot springs), and seismic activities, are considerably frequent in the study area. Most faults are NW-trending faults dominated by normal and translational faults. The area is mainly the Ranyongcuo-Xurucu fault zone; the faults in the area are mainly distributed in the Quaternary graben of the Xurucuo-Chazi area. The Xurucuo-Chazi area is composed of a stepped normal fault in the NW direction and cuts the early NWW direction into various structural units. The Quaternary, the NW-trending fault zone, has been active for many times, thus controlling the evolution of graben basins. Basin margin faults form linear fault triangles. The NNE, NNW, and SN directions are connected end to end, showing the characteristics of tracking faults. A series of hot springs or cold water springs are linearly distributed along the active faults on both the sides of the graben. The SN-trending graben was formed in the area because the collision of the Paleogene Indian plate with the Gangdise plate, and the Indian plate continued to push northward, thus forming an intracontinental orogenic uplift on both the sides of the fault basin. The crust uplifts to a certain height, forming two sets of rupture surfaces in NE and NW directions. The uplifted high mountains form fault-block mountains (uplift belts), resulting in an altitude of 6000 m on both the sides of the fault depression, such as the Duo Zebu Mountain. The neotectonic belts in the study area are mostly closely associated with uplifts and fault depressions. Numerous fault depression zones often constitute several bead-shaped fault basins and lakes. These fault basins are often suitable places for hydrothermal activities. For example, the grid framework active structural belt and the Xu Rucuo active structural belt constitute several bead-shaped fault basins, boundary faults, and uplift mountain bodies. Combined with the great contribution of the local melt in the shell to the heat source, that is, the heat source background in the high-temperature geothermal display area of southern Tibet, the author suggests that the good heat source conditions in this study area are created under the complementary effects of local melts, tectonic movement, hydrothermal activity, and seismic activity.


### 5.2. Migration Channel of Underground Hot Water in the Study Area

Structural fractures are often favourable places for underground hot water storage and migration. The hydrothermal activity in the study area is mainly controlled by the Dajeling-Angren-Renbu-Langxian-Metuo fault, which extends in the EW direction in the northern boundary of the Yarlung Zangbo River junction and the Dangre Yongcuo-Xu Rucuo fracture that spreads in the SN direction. In the study area, Semi Hot Springs, Chazi Hot Springs, and other high-temperature hydrothermal display areas are exposed in the fault basin controlled by these faults. The Chazi fault and Duozebu fault spread in the near SN direction, the Zharinanmu Co South-Cuomei fault spreads in the near EW direction, and other secondary faults intersect with the main fault in the northwest and northeast directions. A huge underground water transport network is formed, which provides good conditions for the upwelling and transport of underground hot water from the deep.


### 5.3. Thermal Storage Analysis

The reservoirs in the study area are divided into shallow and deep reservoirs. The shallow thermal storage is chiefly Quaternary loose rock pore water-bearing rock group, and the lithology is alluvial-diluvial sandy gravel layer. The deep thermal storage is mainly bedrock fractured water-bearing rock group and mainly constitutes magmatic rocks with well-developed fractures and high permeability [60]. The fractures in the rock mass are developed, and the main faults are the intersection fracture zones. These fault fracture zones are underground hot water reservoirs.The cap rocks in the study area chiefly constitute the rock masses of the Pana and Dianzhong Formations. The lithology is mainly shale, sandstone, limestone, andesite, etc. They have low thermal conductivity, large thickness, high water resistance, and heat retention, providing suitable heat retention conditions for the underground hot water in the study area.Based on the aforementioned analysis, conclusions, and previous research results, the conceptual model of geothermal origin for the study area is summarised. The geothermal fluid in the study area mainly receives the infiltration and replenishment of atmospheric precipitation and snowmelt water, and the replenishment area is roughly located in the Duozebu mountainous area from the southwest. The cold water forms the infiltration of faults along the strong tectonic activity. As the distance from the deep magma melt gets closer, the temperature gradually increases. When the appropriate conditions are reached, it migrates upwards and continuously dissolves with the surrounding rock during migration. It exists in magmatic rocks with well-developed fissures and high permeability, forming deep thermal reservoirs. Because the bedrock fissure thermal reservoir is not completely enclosed, the underground hot water continues to migrate upward along the basement fissure and enters the Quaternary aquifer, creating a shallow thermal reservoir. Near the surface, the geothermal fluid is mixed with the cold water on the surface and is exposed at the topographical cut to form hot springs (Figure10).Figure 10
Concept model of the geothermal origin in the study area.


## 5.1. Heat Source in the Study Area

Geophysical surveys showed AMT measurement anomalies, gravity anomalies, and magnetic anomalies present in the study area. The gravity field in the area is high and low in the south and north, respectively. The overall appearance is low gravity anomalies, and high gravity anomalies are in some areas. The magnetic anomalies are generally distributed in NE–SW orientation. The Cenozoic, neotectonic activities, mainly composed of active faults, fault basins (fault-sag belts), fault-block mountains (uplift belts), and accompanying hydrothermal phenomena (hot springs), and seismic activities, are considerably frequent in the study area. Most faults are NW-trending faults dominated by normal and translational faults. The area is mainly the Ranyongcuo-Xurucu fault zone; the faults in the area are mainly distributed in the Quaternary graben of the Xurucuo-Chazi area. The Xurucuo-Chazi area is composed of a stepped normal fault in the NW direction and cuts the early NWW direction into various structural units. The Quaternary, the NW-trending fault zone, has been active for many times, thus controlling the evolution of graben basins. Basin margin faults form linear fault triangles. The NNE, NNW, and SN directions are connected end to end, showing the characteristics of tracking faults. A series of hot springs or cold water springs are linearly distributed along the active faults on both the sides of the graben. The SN-trending graben was formed in the area because the collision of the Paleogene Indian plate with the Gangdise plate, and the Indian plate continued to push northward, thus forming an intracontinental orogenic uplift on both the sides of the fault basin. The crust uplifts to a certain height, forming two sets of rupture surfaces in NE and NW directions. The uplifted high mountains form fault-block mountains (uplift belts), resulting in an altitude of 6000 m on both the sides of the fault depression, such as the Duo Zebu Mountain. The neotectonic belts in the study area are mostly closely associated with uplifts and fault depressions. Numerous fault depression zones often constitute several bead-shaped fault basins and lakes. These fault basins are often suitable places for hydrothermal activities. For example, the grid framework active structural belt and the Xu Rucuo active structural belt constitute several bead-shaped fault basins, boundary faults, and uplift mountain bodies. Combined with the great contribution of the local melt in the shell to the heat source, that is, the heat source background in the high-temperature geothermal display area of southern Tibet, the author suggests that the good heat source conditions in this study area are created under the complementary effects of local melts, tectonic movement, hydrothermal activity, and seismic activity.


## 5.2. Migration Channel of Underground Hot Water in the Study Area

Structural fractures are often favourable places for underground hot water storage and migration. The hydrothermal activity in the study area is mainly controlled by the Dajeling-Angren-Renbu-Langxian-Metuo fault, which extends in the EW direction in the northern boundary of the Yarlung Zangbo River junction and the Dangre Yongcuo-Xu Rucuo fracture that spreads in the SN direction. In the study area, Semi Hot Springs, Chazi Hot Springs, and other high-temperature hydrothermal display areas are exposed in the fault basin controlled by these faults. The Chazi fault and Duozebu fault spread in the near SN direction, the Zharinanmu Co South-Cuomei fault spreads in the near EW direction, and other secondary faults intersect with the main fault in the northwest and northeast directions. A huge underground water transport network is formed, which provides good conditions for the upwelling and transport of underground hot water from the deep.


## 5.3. Thermal Storage Analysis

The reservoirs in the study area are divided into shallow and deep reservoirs. The shallow thermal storage is chiefly Quaternary loose rock pore water-bearing rock group, and the lithology is alluvial-diluvial sandy gravel layer. The deep thermal storage is mainly bedrock fractured water-bearing rock group and mainly constitutes magmatic rocks with well-developed fractures and high permeability [60]. The fractures in the rock mass are developed, and the main faults are the intersection fracture zones. These fault fracture zones are underground hot water reservoirs.The cap rocks in the study area chiefly constitute the rock masses of the Pana and Dianzhong Formations. The lithology is mainly shale, sandstone, limestone, andesite, etc. They have low thermal conductivity, large thickness, high water resistance, and heat retention, providing suitable heat retention conditions for the underground hot water in the study area.Based on the aforementioned analysis, conclusions, and previous research results, the conceptual model of geothermal origin for the study area is summarised. The geothermal fluid in the study area mainly receives the infiltration and replenishment of atmospheric precipitation and snowmelt water, and the replenishment area is roughly located in the Duozebu mountainous area from the southwest. The cold water forms the infiltration of faults along the strong tectonic activity. As the distance from the deep magma melt gets closer, the temperature gradually increases. When the appropriate conditions are reached, it migrates upwards and continuously dissolves with the surrounding rock during migration. It exists in magmatic rocks with well-developed fissures and high permeability, forming deep thermal reservoirs. Because the bedrock fissure thermal reservoir is not completely enclosed, the underground hot water continues to migrate upward along the basement fissure and enters the Quaternary aquifer, creating a shallow thermal reservoir. Near the surface, the geothermal fluid is mixed with the cold water on the surface and is exposed at the topographical cut to form hot springs (Figure10).Figure 10
Concept model of the geothermal origin in the study area.


## 6. Conclusions

(1)
The geothermal area of Chazi, Tibet, features high-temperature geothermal reservoirs, and its hot water is moderately alkaline with high salinity and the hydrochemical type of HCO3–Na. The ratio of Na+/Cl− is considerably >1, indicating that the groundwater undergoes dramatic dissolution that enables Na+, K+, and Ca2+ available in potash feldspar and plagioclase to migrate into hot water. The HCO3− content in the geothermal water is considerably higher than that in the river water, which might indicate the dilution of hot water where runoff occurs after upwelling with the mixed surface cold water, enriching cold water with HCO3−.(2)
Theδ2H-δ18O isotope analysis revealed that the main recharge source in the Chazi geothermal area is atmospheric precipitation, with no obvious δ18O drift. According to the height effect of δ2H and δ18O on precipitation, the recharge elevation is 5200–6000 m.(3)
The temperature of geothermal reservoirs was estimated and corrected using geochemical thermometers, silica enthalpy, and the mixture model of chlorine enthalpy. The temperature of geothermal reservoirs in shallow areas is 143.7–150.6 °C, and that in deep areas is 200~ 233°C; the average mixing ratio of cold water is 75%, and the average circulation depth is 1163 m.(4)
The comprehensive analysis indicated that deep hot water mixes with cold water during upwelling and reaches the deep thermal reservoir layer. Subsequently, the mixed water rises, again mixing with shallow cold water and reaching the shallow thermal reservoir layer. The NW normal fault provides a channel for deep geothermal excretion, and the sand gravel layer of the quaternary alluvium creates a cover for thermal reservoirs. The prominent recharge direction is from alpine to valley, and the main source of recharge is ice and snow meltwater, followed by atmospheric precipitation.



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*Source: 5815996-2022-09-23.xml*