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Article

Geochemical Evidence Constraining Genesis and Mineral Scaling of the Yangbajing Geothermal Field, Southwestern China

by
Hu Yang
1,
Xingcheng Yuan
2,*,
Yongling Chen
1,
Jiawei Liu
2,
Chun Zhan
1,
Guosen Lv
2,
Junfeng Hu
1,
Minglu Sun
3 and
Yunhui Zhang
2
1
Research Center of Applied Geology, China Geological Survey, Chengdu 610036, China
2
Yibin Research Institute, Southwest Jiaotong University, Yibin 644000, China
3
College of Engineering, Tibet University, Lasa 850000, China
*
Author to whom correspondence should be addressed.
Water 2024, 16(1), 24; https://doi.org/10.3390/w16010024
Submission received: 2 November 2023 / Revised: 27 November 2023 / Accepted: 30 November 2023 / Published: 20 December 2023
(This article belongs to the Topic Groundwater Pollution Control and Groundwater Management)
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Abstract

:
The Yangbajing geothermal field, a renowned high-temperature geothermal resource in Tibet of southwestern China, has been utilized for power generation for several decades. To improve geothermal exploitation in the Yangbajing, genesis and mineral scaling have yet to be further revealed. In this study, hydrochemistry and D-O-Sr isotopy were employed for analyzing genesis and mineral scaling in the Yangbajing geothermal field. The geothermal waters were weakly alkaline and had a high TDS content (1400–2900 mg/L) with the Cl-Na, Cl·HCO3-Na, and HCO3·Cl-Na types. The dissolution of silicate minerals (sodium and potassium feldspars) and positive cation exchange controlled the hydrogeochemical process. The geothermal water was recharged from snow-melted water and meteoric water originating from the Nyainqentanglh Mountains and Tangshan Mountains. The geothermal waters possessed the highest reservoir temperature of 299 °C and the largest circulation depth of 2010 m according to various geothermometers. The geothermal waters can produce CaCO3 and SiO2 scaling during vertical and horizontal transport. These achievements can provide a scientific basis for the sustainable development and conservation of the high-temperature geothermal resources in Yangbajing and elsewhere.

1. Introduction

The exploration of clean and renewable energy has become a significant research topic due to the energy crisis and environmental issues [1,2]. The tremendous potential of high-temperature geothermal resources for power generation has attracted substantial attention and research [3,4]. Revealing the genesis is the key for the development of a high-temperature geothermal system due to complicated physical–chemical processes. Recharge source, water–rock interaction, and reservoir temperature are the significant constituents of genesis analysis [5,6,7,8]. In addition, the scaling issue in geothermal wells has posed a serious limitation in the utilization of geothermal energy [9,10].
The application of hydrogeochemical approaches has played a significant role in tracing fluid sources, revealing water–rock interactions, estimating reservoir temperatures and hot/cold water mixing ratios, and predicting fluid fouling trends [10,11,12]. Stable isotopes (δD and δ18O) are widely used to trace the recharge source [13,14]. The correlation of ions and strontium isotopes (87Sr/86Sr) has been used to reveal the water–rock interaction characteristics of geothermal fluids during subsurface runoff and mixing with the parent geothermal fluids [15,16,17]. Reactive elements, such as Na, K, Mg, Ca, and SiO2, have been used to assess the mineral equilibrium and estimate the reservoir temperature [18,19,20,21]. In addition, the mineral saturation index (SI) can determine the dissolution equilibrium state of minerals and scaling tendency in geothermal fluids [22].
The Yangbajing area is a representative high-temperature geothermal field with great exploration potential in Tibet, southwestern China [23]. Numerous investigations have been conducted into the hydrochemical evolution of the Yangbajing geothermal field. These studies indicated that the geothermal water in this field was derived from the recharge of meteoric water and heated by subsurface magma [23,24,25,26,27]. Moreover, the characterization and sources of geothermal gases (He and CO2) in the geothermal field were analyzed, and it was concluded that supercritical geothermal fluids did not exist in the geothermal field of the Yangbajing [2]. However, the reservoir temperature achieved in previous studies remained controversial with a wide range. And scarce knowledge of mineral scaling hampered geothermal exploitation.
The objective of this research is to constrain the genesis of the geothermal system in Yangbajing using hydrochemistry and D-O-Sr isotopy. (1) A correlation of hydrochemical parameters and Sr isotopy were used to reveal the factors dominating hydrochemistry. (2) Various geothermometers were applied to estimate the reservoir temperatures. (3) δD and δ18O were applied to determine the recharge area. (4) Mineral scaling was quantitatively measured by Phreeqc modeling. The achievements of this study are expected to provide valuable knowledge for geothermal exploitation in the Yangbajing and similar high-temperature geothermal fields.

2. Geological and Geothermal Setting

The Yangbajing high-temperature geothermal field is located in the hinterland of the Qinghai-Tibetan Plateau, about 90 km northwest of Lhasa (Figure 1a). The geography exhibits a characteristic distribution of higher terrain in the northwest and lower terrain in the southeast, with the highest elevation area being Mount Nyainqentanglh (7162 m) [8]. The climate belongs to a semi-arid monsoon type with an average annual temperature of 2.5 °C and average annual rainfall of 382.8 mm. The Zangbuqu is the main river traversing the study area [24].
The collision of the Indian and Eurasian plates during the Cenozoic led to the uplift of the Tibetan Plateau. Subduction and thrusting of the Indian plate beneath the Tibetan Plateau resulted in the melting of the crustal contact zone, providing beneficial geological conditions for the formation of geothermal resources [28]. The Tibetan Plateau is divided from north to south into the Songpan-Ganzi Massif, the Qiangtang Massif, the Lhasa Massif, and the Himalayan Massif. These massifs are separated by the east–west trending Kunlun suture zone, the Jinshajiang suture zone, the Bangong-Nujiang suture zone, and the Indus-Yarluzangbo suture zone (Figure 1a) [28,29]. The Yangbajing geothermal field is located in the central part of the Yadong-Gulu rift system within the Lhasa Massif, sandwiched between the Nyainqentanglh Mountains in the northwest and the Tangshan Mountains in the southeast (Figure 1b) [2]. NNE–SSW and NE–SW extensional faults are distributed in front of the Nyainqentanglh Mountains and Tangshan Mountains, providing favorable space for the circulation of geothermal fluids (Figure 1b,c) [24].
The strata exposed in the geothermal field consist of Paleozoic gneisses and mélange in the northwestern part and Carboniferous-Permian slates and Upper Cretaceous breccias in the southeastern part (Figure 1b). The basement rocks of the geothermal field are composed of Late Yanshan-Himalayan granite which outcrops locally in the north [24,25]. The granite was formed by multistage magmatism from the Cretaceous to Neogene periods (K-Ar radiometric ages of 87.4 Ma and 8.1 Ma) [30]. The Yangbajing geothermal field has a substantial coverage of Quaternary sediments on its surface, including Holocene alluvial deposits, flood deposits, and glacial accumulations; Upper Pleistocene alluvial deposits, sinister, and glacial accumulations; and Middle Pleistocene glacial deposits (Figure 1d) [25].
Water 16 00024 g001
Figure 1. (a) Structural map of the Himalayan-Tibet orogenic belt (modified from refs. [31,32,33]). AKMS: Ayimaqin–Kunlun–Mutztagh suture; JS: Jinshajiang Suture; BNS: Bangong-Nujiang suture; YZS: YanglungZangbo suture; MBT: Main Boundary Thrust. (b) Regional geological map of Yangbajing geothermal field (reference [31]). (c) The sampling location of geothermal water and the phenomenon of typical geothermal water exposed to the surface. (d) Yangbajing geothermal field cross-section A-A’.
Figure 1. (a) Structural map of the Himalayan-Tibet orogenic belt (modified from refs. [31,32,33]). AKMS: Ayimaqin–Kunlun–Mutztagh suture; JS: Jinshajiang Suture; BNS: Bangong-Nujiang suture; YZS: YanglungZangbo suture; MBT: Main Boundary Thrust. (b) Regional geological map of Yangbajing geothermal field (reference [31]). (c) The sampling location of geothermal water and the phenomenon of typical geothermal water exposed to the surface. (d) Yangbajing geothermal field cross-section A-A’.
Water 16 00024 g001
The findings of the INDEPTH (International Depth Profiling of Tibet and the Himalayas) project indicate that the magmatic melts present at a depth of 15–20 km provide a heat source for the geothermal system in the Yangbajing field [34,35,36,37,38,39]. During the field sampling conducted in June 2023, the temperature of the outcrop of the geothermal spring water was recorded to reach a maximum of 80 °C. Additionally, certain geothermal springs exhibited noticeable boiling phenomena, accompanied by the release of hot gas. The flow rate of these geothermal springs varied between approximately 1.5 L/s and 3.5 L/s (Figure 2). The geothermal field consists of two underlying geothermal reservoirs: (1) The shallow thermal reservoir is buried at a depth of 180–280 m. The northern region is primarily composed of eroded granite and Quaternary sandstone, while the southern region is predominantly comprised of Quaternary sediments. The reservoir temperature is 140–160 °C, and the geothermal fluids exhibit a characteristic of lateral replenishment in the NW–SE direction. (2) The deep thermal reservoir is buried at a depth of 950–1850 m. This reservoir is comprised of cranberries and fissures, while the surface is composed of very weathered cranes and dark clouds. The reservoir temperature is 250–278 °C [23,34].

3. Methods

3.1. Sampling and Analysis

In June 2023, five sets of geothermal water samples (YBJS1–YBJS4, YBJT1) and one set of river water samples (YBJR1) were collected in the Yangbajing area of Tibet. Before sampling, temperature, pH, electrical conductivity (EC), oxidation reduction potential (ORP), and total dissolved solids (TDS) were measured using a German Multi3630IDS portable multiparameter water quality analyzer. HCO3− was determined via Gram titration. Prior to bottling, all water samples were filtered through 0.2 μm filter membranes and filled into 550 mL high-density polyethylene bottles. These bottles were washed and rinsed at least three times before sampling. In order to prevent precipitation of SiO2 in the water, the geothermal water samples were diluted fivefold using deionized water prior to analysis for SiO2 [6].
The collected samples were sent to Kehui Testing (Tianjin) Technology Co., Ltd. for hydrochemical and D-O-Sr isotope analyses. Major cations (K+, Na+, Ca2+, and Mg2+) were analyzed via inductively coupled plasma emission spectrometry (ICP-OES), while anions (Cl−, SO42−, and F−) were analyzed via ion chromatography (Diona ICS-1100). Silicon and other trace elements were detected by ICP-MS (7000C, Agilent). The ion charge balance was calculated to be less than ±10% for most samples (Table 1). Hydrogen and oxygen isotopes (δD and δ18O) were measured using the water balance hydroxide isotope analysis method (MAT251EM). The obtained data were subsequently standardized to Vienna Standard Mean Ocean Water (VSMOW) using the conventional δ (‰) expression. The accuracy of the detection of δD and δ18O was ±0.5% and ±0.1%, respectively. Strontium isotopes (87Sr/86Sr) were tested via MC-ICP-MS using NIST SRM 987 Sr isotope standards as standards. In addition, this study also cited the sample data (YBJT2 and ZK4001) studied by references [24,25].

3.2. Geochemical Calculations

The recharge sources of geothermal water were assessed using stable hydrogen (δD) and oxygen (δ18O) isotopes, including recharge elevation, recharge temperature, and the proportion of magma water contribution. The recharge sources were calculated based on correlation equations and isotope bivariate mixing models [40,41,42]. Reservoir temperatures have been assessed using a variety of geothermometers, including silica geothermometers (Equations (1) and (2)) [18], multimineral equilibria [43], and chloride–enthalpy diagrams [19]. In addition, the SI of various minerals in the geothermal water was calculated using the PHREEQC3.0 software to assess the mineral equilibrium state and scaling trends [44].
Quartz geothermometer: T (°C) = 1309/(5.19 − log(SiO2)) − 273.15
Chalcedony geothermometer: T (°C) = 1032/(4.69 − log(SiO2)) − 273.15
where SiO2 is the chemical component in the geothermal water in mg/L.
A geothermometer is used to estimate reservoir temperature using the relationship between the content of specific chemical components in geothermal water and temperature. The basic basis is that the mineral and aqueous solutions in the deep thermal reservoir reach equilibrium. The temperature of geothermal water decreases during the rising process, but the chemical composition remains relatively stable, so as to estimate the reservoir temperature [45]. However, because most of the geothermal water exposed on the surface has not reached the complete water–rock equilibrium state, the reservoir temperature estimated by the geothermometer is generally low and needs to meet the corresponding conditions [12]. The multimineral equilibrium method is to determine the equilibrium temperature by constructing the relationship between the log (Q/K) of the mineral component and the temperature. The temperature at the convergence of the mineral curve is the estimated reservoir temperature in the state of mineral equilibrium. The principle of this method is to correct the CO2 degassing during the rise in geothermal water by adding equimolar components of HCO3− and H+ to the file. In this study, all calculations were performed using the SOLVEQ-XPT program to calculate the SI change trend of a group of water-bearing minerals at different reservoir temperatures. The minerals selected were those present in the geothermal system. This method allows an accurate assessment of the temperature at which minerals in geothermal water reach equilibrium by correcting for the effects of boiling decompression cooling [6]. Since chlorine–enthalpy diagrams can indicate the cooling modes experienced by geothermal water during the ascent process, such as conduction cooling, adiabatic cooling, and mixing with cold groundwater, the estimated temperature is close to the highest geothermal reservoir temperature [14].

4. Results

4.1. Hydrochemical Characteristics

The characteristics of hydrochemical parameters are presented in Table 1. The pH value of the geothermal water in Yangbajing was weakly alkaline (8.29–9.19). The exposed temperature of the geothermal springs and borehole water were 67–80 °C and 76–159 °C. The boreholes had a higher exposed temperature than the geothermal springs. The TDS contents of the geothermal springs and borehole water were 1400–1613 mg/L and 1599–2900 mg/L (Figure 3). The main anions were Cl− and HCO3−, and the main cation was Na+ in geothermal springs. The main anion was Cl− and the main cation was Na+ in geothermal borehole waters. All geothermal waters showed higher concentrations of major elements (except Ca2+ and Mg2+) and trace elements (SiO2, F, Sr, Li, B) compared to river water (Table 1 and Figure 3). The hydrochemical characteristics of geothermal waters in the Yangbajing field were comparable with those in the Kangding [3] and Rehai [46] high-temperature geothermal systems in southwestern China. Hence, the Yangbajing field is possibly related to magma existence and fault activity as well.
The Piper trilinear diagram showed that the hydrochemical types of geothermal water samples were predominantly of the Cl-Na, Cl·HCO3-Na, and HCO3·Cl-Na types (Table 1 and Figure 4). The hydrochemical type might be attributed to the reservoir lithology of the geothermal field being granite and sandstone (rich in silicate minerals) [34]. Interestingly, the process of mixing hot and cold water was indicated: Cl-Na-type water (deep reservoir) and HCO3-Ca·Na-type water (shallow cold water) were mixed to form Cl-HCO3·Na and HCO3-Cl·Na-type water (shallow reservoir) (Figure 4). The shallow reservoir in Yangbajing was formed due to the upward flow of geothermal fluids from the deep reservoir into the Quaternary aquifer [24].

4.2. D-O-Sr Isotopes

The hydrogen (δD) and oxygen (δ18O) isotope ratios of geothermal water ranged from −150.7 ‰ to −139.0 ‰ and −17.50 ‰ to −16.00 ‰, respectively. River water was more enriched in hydrogen (δD) and oxygen (δ18O) than geothermal water (−121.8 ‰ and −15.72 ‰, respectively) (Table 1). The strontium isotope (87Sr/86Sr) ratios of geothermal water varied from 0.71203 to 0.71259; the strontium isotope (87Sr/86Sr) ratio of river water was lower than that of geothermal water, at 0.70890 (Table 1).

5. Discussion

5.1. Recharge Source by δD and δ18O

The stable hydrogen (δD) and oxygen (δ18O) isotopes of water samples followed the global meteoric water line (GMWL) [13] and the local meteoric water line (LMWL) [25], revealing the origin mainly from meteoric water (Figure 5a). The stable oxygen (δ18O) isotopic compositions of geothermal water (from −17.50 ‰ to −16.00 ‰) exhibited a notable positive isotopic shift compared to GMWL and LMWL, i.e., the “δ18O” drifting. This phenomenon could be attributed to oxygen isotopic exchange during the water–rock interaction [6].
The recharge source of geothermal water can be included in three end members: magmatic water (rectangle A, [48]), snow-melted water (rectangle B, [25]), and meteoric water (rectangle C, [3]) (Figure 5a). The river water was closer to end element C, suggesting its atmospheric recharge was dominant. Almost all geothermal waters fell in the straight line between snow-melted and magmatic water. The contribution of magmatic water in the geothermal water of Yangbajing was estimated to be between 15% and 23% using a bivariate isotopic mixing model.
Due to the “δ18O” drift phenomenon, δD was more suitable for estimating the recharge elevation of geothermal water. According to the elevation gradient of δD on the Tibetan Plateau (–2.6%/100 m, [49]), the recharge elevations estimated by Equation (3) were between 5008 m and 5458 m (Table 1). In the study area, the Nyainqentanglh Mountains in the northwest and the Tangshan Mountains in the southeast were the recharge areas of geothermal water (Figure 5b). Based on the temperature effect of stable hydrogen isotopes (Equation (4)) [40], the recharge temperature of the geothermal water ranged from −20 °C to −16 °C.
H = (δR − δP)/K + h
where H (m) is the recharge elevation, δR (‰) is the δD of the geothermal water sample, δP (‰) is the δD of the river water sample, and h (m) is the elevation of the river.
δD =3 × T − 92
where T (°C) is recharge temperature of the geothermal water.

5.2. Process Controlling Hydrochemical Compositions

5.2.1. Correlation of Hydrochemical Compositions

In this study, Pearson correlation analysis was performed to understand the correlation characteristics among the hydrochemical variables. As shown in Figure 6, the correlation between Cl−, Na+, K+, SiO2, Li, and B was very positive, with the correlation coefficients higher than 0.85, indicating that these variables might have the same source. However, Mg2+ and Ca2+ were negatively correlated with other ions, suggesting that there were other factors influencing the changes in Mg2+ and Ca2+ concentrations.
The relationships between Cl and other major ions are effective in tracing hydrogeochemical processes during the formation of geothermal water [12,50]. Among them, Cl correlated well with Na and K (R2 = 0.87 and 0.94) (Figure 7a,b), and the ratio of Na/Cl was low, up to a maximum of only 1.0, indicating that the geothermal water circulation depth was shallow [51]. The negative correlation (R2 = 0.80 and 0.57) between Cl and Ca and Mg in water samples (Figure 7c,d), along with their low concentrations of Ca and Mg, showed the absence of carbonate rocks in reservoir rocks. This conclusion aligned with the regional geological characteristics of the study area. The reason for the decrease in Ca2+ concentration might be the CaCO3 scaling in the geothermal well [10,52]. Cl in the water samples correlated well with SO4 (R2 = 0.91), but there was little correlation with deep borehole water (Figure 7e). This occurred because the deep geothermal fluid is laden with H2S gas. As it rises and experiences heat loss, a portion of it transforms into condensate water, which is subsequently leached out by the surrounding rocks and oxidized into SO4 [27]. Consequently, the SO4 concentration in the geothermal fluid increases during ascent due to the oxidization of H2S deep gas. A strong correlation of Cl between Li and B was observed for all water samples (R2 = 0.92 and 0.99; Figure 7f,g), demonstrating that geothermal waters have a common primary geothermal fluid and have undergone mixing [53]. The favorable connection between Cl and SiO2 in the geothermal water from Yangbajing (R2 = 0.92; Figure 7h) was attributed to the strong association between the concentration of SiO2 and temperature [18]. The poor correlation between Cl and F (R2 = 0.48; Figure 7i) indicates that there were other geochemical processes affecting the concentration of F in shallow geothermal fluids.
In the study area, Cl-bearing minerals were not found in the strata. Consequently, the elevated concentration of Cl− in geothermal water is attributed to magma, as magma has the capacity to supply a significant quantity of Cl− to geothermal water [31,54]. Excess concentrations of Na+ and K+ relative to Cl− might be derived from the dissolution of silicate minerals. Chemical activity diagrams of the Na2O-Al2O3-SiO2-H2O and K2O-Al2O3-SiO2-H2O systems can determine the dissolution process of silicate minerals [55]. The phase boundaries were plotted at 100 °C (black) and 200 °C (red) (Figure 8a,b). The dominant factors influencing the presence of Na+ and K+ in geothermal water appeared to be the dissolution of sodium feldspar and potassium feldspar, as seen in Equations (5) and (6) [6].
2NaAlSi3O8 + 3H2O + 2CO2→Al2(SiO5)(OH)4 + 4SiO2 + 2Na+ + HCO3−
2KAlSi3O8+3H2O+2CO2→Al2(SiO5)(OH)4+4SiO2+2K++HCO3−
Moreover, ratio plots of (Na+ + K+ − Cl−) and (Ca2+ + Mg2+) − (SO42− + HCO3−) can illustrate the effect of ion exchange on groundwater [56]. Schoeller indices CAI-I (=(Cl− − (Na+ + K+))/Cl−) and CAI-II (=(Cl− − (Na+ + K+))/(HCO3− + SO42− + CO32− + NO3−)) can further confirm the type of cation exchange [57]. The geothermal samples were distributed near the ion exchange line, and the CAI-I and CAI-II values were below zero (Figure 8c,d). This signifies that positive ion exchange occurred, i.e., Ca2+ in the geothermal water was replaced by Na+ in the granite.

5.2.2. Water–Rock Interactions Based on 87Sr/86Sr Ratio

The initial abundance of strontium isotopes (87Sr/86Sr) varies widely between rock types and can serve as a conservative tracer of water–rock interactions [58]. Previous studies showed that the 87Sr/86Sr ratios of silicate rocks in the Yangtze River Basin of China are typically >0.715 [59]. The 87Sr/86Sr ratios in marine carbonate rocks range from 0.707 to 0.709 [60]. In addition, the 87Sr/86Sr ratios in granites from the Yangbajing geothermal field in the study area were 0.710 to 0.718 [30]. The 87Sr/86Sr ratios of geothermal water consistently aligned with those of granite (Figure 9a), suggesting that the predominant source of strontium was the granite reservoir. The river water was located in the range of carbonate rocks, which could be related to the dissolution of carbonate rocks (Figure 1b). In addition, the negative correlation between Sr2+ and Ca2+ in geothermal water also confirmed that Sr2+ did not come from the dissolution of carbonate minerals (Figure 9b). Geothermal water leaches Sr2+ from silicate minerals during water–rock interactions, leading to an increase in Sr2+ concentration in geothermal water [61].
In conclusion, the geothermal water all came from the same primary geothermal fluid (stored in granite), which mixed with the cold water as it rose. The dissolution of silicate minerals (sodium and potassium feldspars) and positive cation exchange controlled the hydrochemical evolution, resulting in the formation of Cl-Na, Cl·HCO3-Na, and HCO3·Cl-Na types of geothermal water.

5.3. Geothermometry

5.3.1. Equilibrium State of Geothermal Waters

Geothermometers can estimate reservoir temperatures in geothermal systems, and the temperature-related water–rock equilibrium needs to be analyzed first [15]. The Na-K-Mg triangular diagram can classify waters as fully equilibrated waters, partially equilibrated or mixed waters, and immature waters [21]. It can be seen that the geothermal waters of Yangbajing were all located in partially equilibrated waters, with individual water samples close to the full equilibrium point (Figure 10). It indicates that the geothermal water of the Yangbajing geothermal field had reached a partial equilibrium state with the surrounding rocks, characteristic of partially equilibrated water. This occurrence was closely associated with the mixture of shallow cold water [6]. By comparing the data of Laoyulin and Rehai, we found that the samples in these three regions were not randomly distributed in the Na–K–Mg ternary diagram but linearly distributed along the isolines that tend to Na-K (Figure 10) [21]. Using Na-K geothermal thermometers, it was possible to estimate that deep reservoirs with temperatures as high as 240 °C, 282 °C, and 260 °C might exist beneath Yangbajing, Laoyulin, and Rehai. It was an interesting phenomenon that geothermal water from magmatic geothermal systems with higher Na+ and K+ concentrations did not change Na/K ratios during mixing with cold water, and thus, reservoir temperatures could be characterized intuitively in the Na–K–Mg ternary diagram. The deep borehole water ZK4001 in the Yangbajing geothermal system was located on the temperature line of 290 °C. The sample was collected at the depth of the borehole [24], indicating that the deep reservoir temperature in Yangbajing was actually higher.

5.3.2. Temperature and Depth of Geothermal Reservoir

The reservoir temperature of the geothermal water in Yangbajing was measured via the solute geothermometers, quartz and chalcedony [18,21], and the results are shown in Table 2. For geothermal springs and borehole water, the chalcedony geothermometer estimated temperatures at 132–152 °C; the quartz geothermometer estimated temperatures at 157–174 °C. The SI of quartz and chalcedony calculated by PHREEQC3.0 software was 0.03–0.44 and −0.20–0.15, respectively. It can be seen that the saturation degree of quartz in geothermal water is higher, so the temperature estimated by the quartz geothermometer was selected to represent the reservoir temperature. Most of the geothermal water in this study belonged to partially equilibrated water, and there was the influence of cold water mixing. Therefore, the solute thermometer could only reflect the shallow reservoir temperature after mixing with cold water. Additionally, research is necessary to investigate the influence of various cooling mechanisms, including boiling depressurization cooling, adiabatic cooling, mixing with cold groundwater cooling, and conduction cooling [14].
During the process of boiling and depressurization, geothermal water undergoes the inevitable loss of some dissolved gases (e.g., CO2). A corrective process for CO2 degassing can be constructed using the multimineral equilibrium method so as to estimate the reservoir temperature under the influence of boiling decompression cooling. [43]. The SOLVEQ-XPT program was employed in this study to establish thermodynamic equilibrium between the minerals (Figure 11). Calculations were performed for two sample points, YBJT1 (from shallow reservoirs) and ZK4001 (from deep reservoirs), because in this study, these two points were the samples with the highest Cl concentration in the two reservoirs. The selected minerals began to converge in the temperature intervals of 200–220 °C and 268–276 °C, respectively, when 0.1 mol/L and 0.05 mol/L component species (H+ and HCO3−) were added to these two places in the file to correct degassing (Figure 11). The multimineral equilibrium method estimated higher reservoir temperatures than the quartz geothermometer due to the correction for the effect of boiling decompression on temperature (Table 2). This suggests that the temperature estimated by the multimineral equilibrium method might be the reservoir temperature prior to cold water mixing. However, only the deep borehole sample ZK4001 had temperatures (268–276 °C) close to those estimated by the Na-K geothermometer in the Na–K–Mg ternary diagram (240 °C and 290 °C), whereas reservoir temperatures (200–220 °C) estimated by the shallow borehole sample YBJT1 were low. Geothermal water from the same parent geothermal fluid should have the same reservoir temperature. Therefore, it is necessary to continue to explore the difference in reservoir temperatures estimated for different types of geothermal water, that is, to consider the influence of other cooling mechanisms on the deep reservoir temperature estimation of shallow geothermal water.
There was a significant correlation (R2 = 0.95) between Cl concentration and enthalpy for all water samples (Figure 12a). Therefore, in this study, a chlorine–enthalpy diagram [19] was used to analyze the effect of adiabatic cooling, conduction cooling, and cold water mixing on the temperature change in the geothermal water ascent process (Figure 12b). Drill hole ZK4001 was sampled at 1495 m below ground at temperatures up to 159.3 °C, with high Na+ and Cl− concentrations and low Ca2+ and Mg2+ concentrations (Table 1). Hence, adiabatic cooling was the main cooling method used to form the ZK4001 [14]. Most of the samples taken near the surface were below the local boiling temperature (85 °C), indicating that conduction cooling was experienced. We used the quartz-calculated temperatures to represent the transition point B between conduction cooling and cold water mixing; the line connecting the cold water point A and the transition point B as the cold water mixing line; and the intersection of the extension of A and B with the adiabatic cooling line, C, to represent the reservoir temperature (Figure 12b). The calculated deep reservoir temperatures were 257–299 °C for shallow geothermal water, and the cold water mixing ratios were 37–47% (Table 2). It was inferred that reservoir one (299 °C) was the parent geothermal fluid in the deep reservoir of the Yangbajing geothermal field. Reservoir temperatures estimated from the geothermal water collected at shallow depths or at the surface were essentially the same as those estimated from the geothermal water collected deeper in the subsurface (ZK4001: 1459 m) due to consideration of the effects of a range of cooling mechanisms (Table 2).
In conclusion, the temperatures of the geothermal water in Yangbajing continued to decrease due to a series of effects during the cyclic rise. We synthesized the temperatures estimated by various methods and used the reservoir temperatures estimated by the chlorine–enthalpy diagram (257–299 °C) as the deep reservoir temperature; the reservoir temperatures estimated by the quartz geothermometer (157–174 °C) as the temperature of the shallow reservoir mixed with cold water; and the temperatures estimated by the multimineral equilibrium (YBJT1: 200–220 °C) as the temperatures between the deep and the shallow reservoirs, because the effects of a series of cooling mechanisms resulted in the differential temperature variations between the reservoirs. The shallow reservoir temperatures (157–174 °C) and deep reservoir temperatures (257–299 °C) estimated in the present study were basically consistent with the temperatures of geothermal water in underground boreholes measured by previous researchers (150–175 °C and 250–278 °C for shallow and deep boreholes, respectively) [23,34].
The geothermal gradient of Yangbajing is 6.67 m/°C [31]. According to the reservoir temperatures estimated in this study, the shallow reservoir depth was 1059–1172 m and the deep thermal reservoir depth was 1724–2010 m.

5.3.3. Conceptual Model of Geothermal System Genesis in Yangbajing

Based on the geological and topographic features (Figure 1b,d), this study proposed the genesis conceptual model of the geothermal system in Yangbajing (Figure 13). The geothermal water was recharged from snow-melted water and meteoric water originating from the Nyainqentanglh Mountains in the northwest and the Tangshan Mountains in the southeast, respectively. Further, the geothermal water was partially mixed with magmatic water. Recharged water rapidly infiltrated through fracture zones and fissures and circulated along geothermal water channels formed by faults and fracture systems. The minerals involved in water–rock interactions during circulation were mainly silicate minerals (sodium feldspar and potassium feldspar), while positive cation exchange occurred, resulting in the formation of weakly alkaline water of the Cl-Na, Cl·HCO3-Na, and HCO3·Cl-Na types. Cl, Li, and B in geothermal water all came from magmatic water at depth. During the ascent, the CO2 and H2S that accompany geothermal fluids brought out from the depths became condensate due to condensation and participated in water–rock interactions to form HCO3– and SO42– [2,26]. Due to the high-temperature molten magma body beneath the Yangbajing [23], the heated high-temperature geothermal fluid rose along the interior of the pipe formed by the two sets of stretching normal faults, NNE–SSW and NE–SW. The temperature started to drop due to adiabatic cooling. A deep geothermal reservoir with high temperatures (257–299 °C) consisting of granitic mylonite and fissured granite was formed near a depth of 1724 m to 2010 m from the surface. Geothermal fluids continued to rise along the fracture pipe, and the temperature continued to fall due to conduction cooling. After mixing occurred with shallow cold water near the surface (1059–1172 m), a shallow reservoir with medium temperatures (157–174 °C) consisting of Quaternary sediments and sandstone was formed. The primary factor contributing to the presence of binary mixing relationships in geothermal water was the mixing with shallow cold water. The geothermal fluid in the shallow reservoir in the northwest continued to flow towards the southeast of the geothermal field through joint fissure channels, eventually overflowing to the surface to form hot springs.

5.4. Assessment of Scaling Trends in Geothermal Fluids

Scaling is one of the major problems in the utilization and production of geothermal fluids [62]. The change in hydrodynamic characteristics, gas partial pressure, and pH value will occur during the rising process of geothermal water, which will affect the change in scaling characteristics. In this research, the PHREEQC3.0 program was employed to model the alteration in equilibrium conditions of diverse minerals across a temperature range spanning from the wellhead temperature to 200 °C and the pH range of 4 to 10, respectively. Two samples, YBJS1 and YBJT1, were selected for simulation (Figure 14), because these two points had higher Ca2+ concentrations relative to the other samples (Table 1). Both samples exhibited supersaturation of aragonite and calcite across the entire temperature range (Figure 14a,b), indicating that intensive degassing of the geothermal water occurred during its ascent, resulting in a tendency for carbonate minerals to supersaturate and precipitate [6,10]. The carbonate minerals in the two samples gradually changed from an undersaturated state to a supersaturated state with the increase in the pH value (Figure 14c,d). This is because of the geothermal fluid flowing from the deep (high temperature and high pressure) to the shallow (low temperature and low pressure). During this transition, CO2 degassing and the boiling of geothermal water occur, resulting in an increase in pH. Consequently, this promotes the precipitation of carbonate minerals [10,63]. It might be inferred that the occurrence of CaCO3 scaling took place at a certain stage during the upward movement of geothermal water from the reservoir to the surface in the Yangbajing. Lei et al. analyzed the XRD of the scaling material from the wellheads of Yangbajing and concluded that the volume share of CaCO3 was more than 99% [52]. And by constructing a wellbore calcium carbonate scaling evaluation model, the maximum scaling rate of one well was simulated to be 0.067 mm/d, which occurred at a burial depth of 97.5 m.
It was noteworthy that the silicate minerals in the geothermal water of Yangbajing basically reached a state of supersaturation as the temperature decreased and the pH did not exceed 9 (Figure 14). During the ascent, as the geothermal water mixed with the shallow cold water, there was a rapid decrease in temperature and pressure. The SiO2 concentration in the shallow fluid surpassed the saturation point of quartz and chalcedony, triggering precipitation. Thus, the geothermal fluid slowly precipitated SiO2 in the form of siliceous cementation in the transport channel, forming a self-enclosed cemented cap layer [8,27]. Sulfate minerals were undersaturated throughout the temperature and pH ranges (Figure 14), indicating that sulfate scaling did not occur in the geothermal waters of Yangbajing.

6. Conclusions

In this study, the following conclusions were drawn by investigating the hydrogeochemical characteristics, isotopic characteristics, and scaling characteristics of geothermal water in the Yangbajing geothermal field:
(1) The geothermal water was recharged from snow-melted water and local meteoric water located at higher altitudes on the Nyainqentanglha Mountains in the northwestern part and on the Tangshan Mountains in the southeastern part, and some of the magmatic water was mixed in the deep geothermal water.
(2) Weakly alkaline waters of the Cl-Na, Cl·HCO3-Na, and HCO3·Cl-Na types were identified in the Yangbajing geothermal field. The geothermal water all came from the same primary geothermal fluid in the granitic reservoir, which mixed with the cold water as it rose. The dissolution and positive cation exchange of silicate minerals (sodium and potassium feldspars) controlled the hydrochemical evolution.
(3) The temperature range of the shallow reservoir was between 157 °C and 174 °C, and the burial depth was 1059–1172 m; the temperature range of the deep reservoir was between 257 °C and 299 °C, and the burial depth of the reservoir was 1724–2010 m. The mixing ratio of geothermal water with cold water in the process of geothermal water ascension was about 37–47%. The geothermal water produced CaCO3 and SiO2 scaling during vertical and horizontal transport.

Author Contributions

Conceptualization, X.Y.; data curation, Y.C. and G.L.; formal analysis, J.L.; investigation, H.Y., X.Y., Y.C., C.Z., G.L. and M.S.; methodology, H.Y. and J.L.; software, Y.C., C.Z. and J.H.; supervision, Y.Z.; validation, J.L. and Y.Z.; writing—original draft, H.Y.; writing—review and editing, X.Y. All authors have read and agreed to the published version of the manuscript.

Funding

The Scientific Key R&D project of the Tibet Autonomous Region, grant number [XZ202201ZY0021G], the National Natural Science Foundation of China, grant number [42072313, 42102334], the Sichuan Provincial Department of Science and Technology Projects, grant number [2022NSFSC0413, 2023YFS0356], the Yibin Government Foundation, grant number [SWJTU2021020007, SWJTU2021020008].

Data Availability Statement

All of the research data have been provided in the paper.

Acknowledgments

We would like to thank the anonymous reviewers and editor for their constructive comments on this manuscript.

Conflicts of Interest

The authors declare no conflict of interest.

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Figure 2. Photographs of geothermal water and river on site. (a) YBJS1, (b) YBJS2, (c) YBJS3, (d) YBJS4, (e) YBJT1, and (f) YBJR1. The temperatures marked in the figure were the exposure temperature of the site.
Figure 2. Photographs of geothermal water and river on site. (a) YBJS1, (b) YBJS2, (c) YBJS3, (d) YBJS4, (e) YBJT1, and (f) YBJR1. The temperatures marked in the figure were the exposure temperature of the site.
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Figure 3. Schöoller diagram of all water samples in Yangbajing.
Figure 3. Schöoller diagram of all water samples in Yangbajing.
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Figure 4. Piper trilinear diagram of all water samples from Yangbajing [47].
Figure 4. Piper trilinear diagram of all water samples from Yangbajing [47].
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Figure 5. (a) The relationship between δD value and δ18O value of all water samples in Yangbajing; (b) Yangbajing geothermal water recharge source area.
Figure 5. (a) The relationship between δD value and δ18O value of all water samples in Yangbajing; (b) Yangbajing geothermal water recharge source area.
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Figure 6. Pearson correlation of Yangbajing geothermal water. Red indicates positive correlation and blue indicates negative correlation.
Figure 6. Pearson correlation of Yangbajing geothermal water. Red indicates positive correlation and blue indicates negative correlation.
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Figure 7. Relationship between Cl and hydrochemical parameters. (a) Na, (b) K, (c) Mg, (d) Ca, (e) SO4, (f) B, (g) Li, (h) SiO2, and (i) F. The black dotted line is the fitting line of all samples; the purple double-dotted line is the fitting line to remove the deep borehole water (ZK4001). The method used for fitting is divided into the least squares method.
Figure 7. Relationship between Cl and hydrochemical parameters. (a) Na, (b) K, (c) Mg, (d) Ca, (e) SO4, (f) B, (g) Li, (h) SiO2, and (i) F. The black dotted line is the fitting line of all samples; the purple double-dotted line is the fitting line to remove the deep borehole water (ZK4001). The method used for fitting is divided into the least squares method.
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Figure 8. Hydrochemical analysis diagrams. (a) log(SiO2) vs. log(Na+/H+), (b) log(SiO2) vs. log(K+/H+), (c) Na+ + K+ − Cl− vs. Ca2+ + Mg2+ − (HCO3− + SO42−), and (d) CAI-I vs. CAI-II.
Figure 8. Hydrochemical analysis diagrams. (a) log(SiO2) vs. log(Na+/H+), (b) log(SiO2) vs. log(K+/H+), (c) Na+ + K+ − Cl− vs. Ca2+ + Mg2+ − (HCO3− + SO42−), and (d) CAI-I vs. CAI-II.
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Figure 9. (a) Sr2+ vs. 87Sr/86Sr, (b) Sr2+ vs. Ca2+.
Figure 9. (a) Sr2+ vs. 87Sr/86Sr, (b) Sr2+ vs. Ca2+.
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Figure 10. Na–K–Mg triangle diagram [21]. Laoyulin in Kangding and Rehai in Tengchong were from references [3,46].
Figure 10. Na–K–Mg triangle diagram [21]. Laoyulin in Kangding and Rehai in Tengchong were from references [3,46].
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Figure 11. Results of estimation of shallow (a) and deep (b) geothermal reservoir temperatures via multivariate mineral–water equilibrium thermometers. The red star in the figure represents the position and corresponding temperature when the mineral group reaches equilibrium.
Figure 11. Results of estimation of shallow (a) and deep (b) geothermal reservoir temperatures via multivariate mineral–water equilibrium thermometers. The red star in the figure represents the position and corresponding temperature when the mineral group reaches equilibrium.
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Figure 12. (a) The fitting relationship between chloride and enthalpy of Yangbajing water sample. (b) Chlorine–enthalpy diagram of Yangbajing water sample. The relationship between enthalpy and temperature is as follows: enthalpy = temperature × 4.1868; the unit is J/g.
Figure 12. (a) The fitting relationship between chloride and enthalpy of Yangbajing water sample. (b) Chlorine–enthalpy diagram of Yangbajing water sample. The relationship between enthalpy and temperature is as follows: enthalpy = temperature × 4.1868; the unit is J/g.
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Figure 13. Conceptual model of genesis of Yangbajing geothermal system. The numbers in the figure are as follows: 1. Pleistocene glacial deposits in the Quaternary; 2. Paleogene Eocene Pana Formation trachyte and ignimbrite.; 3. Miocene Jieli medium–fine-grained biotite monzonitic granite; 4. fine-grained biotite monzonitic granite in the Eocene Yangbajingbingzhan; 5. granitic mylonite belt; 6. granitic gneiss; 7. gneiss; 8. granoamphibolite; 9. maritime carbonate rocks; 10. migmatite; 11. rainfall and snow melt infiltration recharge; 12. heating groundwater; 13. geothermal fluid rise; 14. geothermal fluid lateral flow; 15. deep-layer volatiles; 16. rivers; 17. normal fault; 18. joints and fissures; 19. steaming ground and geothermal spring; 20. shallow reservoirs; 21. deep reservoirs. This diagram is based on [2,23,24,31].
Figure 13. Conceptual model of genesis of Yangbajing geothermal system. The numbers in the figure are as follows: 1. Pleistocene glacial deposits in the Quaternary; 2. Paleogene Eocene Pana Formation trachyte and ignimbrite.; 3. Miocene Jieli medium–fine-grained biotite monzonitic granite; 4. fine-grained biotite monzonitic granite in the Eocene Yangbajingbingzhan; 5. granitic mylonite belt; 6. granitic gneiss; 7. gneiss; 8. granoamphibolite; 9. maritime carbonate rocks; 10. migmatite; 11. rainfall and snow melt infiltration recharge; 12. heating groundwater; 13. geothermal fluid rise; 14. geothermal fluid lateral flow; 15. deep-layer volatiles; 16. rivers; 17. normal fault; 18. joints and fissures; 19. steaming ground and geothermal spring; 20. shallow reservoirs; 21. deep reservoirs. This diagram is based on [2,23,24,31].
Water 16 00024 g013
Water 16 00024 g014
Figure 14. The saturation index of typical minerals in geothermal water changes at different temperatures and pH.
Figure 14. The saturation index of typical minerals in geothermal water changes at different temperatures and pH.
Water 16 00024 g014
Table 1. Hydrochemical and isotopes results of geothermal water and river water samples.
Table 1. Hydrochemical and isotopes results of geothermal water and river water samples.
NO.TypeElevation (m)T(°C)pHTDSNa+K+Mg2+Ca2+Cl−SO42−
mg/L
YBJS1HS4291678.34 1613408.3445.700.248.99408.6459.33
YBJS24280689.14 1400374.8436.220.041.23389.6058.03
YBJS34280809.121471390.2443.300.041.22397.6855.43
YBJS44280769.19 1467392.1440.440.011.41393.5250.50
YBJT1BW4285768.29 1599410.0450.870.5110.59435.9253.19
YBJT243251108.34 1700300.2038.600.234.70514.0058.40
ZK4001DBW44161598.40 2900709.00135.000.132.101020.0027.00
YBJR1RW4346138.26 9810.461.052.2219.652.9413.55
HCO3−SiO2FSrLiB87Sr/86SrδD (‰)δ18O (‰)C.B. (%)H.T.R.E.(m)
mg/L
496.66139.600.040.477.6040.460.71219−144.8−16.883.71Cl·HCO3-Na5229
90.61157.600.150.606.9140.010.71204−148.5−17.39−11.67Cl-Na5372
160.58165.000.020.467.3640.200.71203−150.7−17.46−9.51Cl-Na5458
761.62170.600.020.467.3238.930.71207−149.2−17.3115.11HCO3·Cl-Na5400
503.54182.200.010.367.5840.940.71205−150.1−17.334.69Cl·HCO3-Na5435
163.60165.8019.600.216.4058.700.71235−142.1−17.5015.17Cl-Na5127
363.00581.5018.000.3325.00119.000.71259−139.0−16.002.54Cl-Na5008
80.2912.800.020.100.110.320.70890−121.8−15.720.99HCO3-Ca·Na−
Note: “−” means no calculation. (“HS” represents hot spring; “BW” represents borehole water; “DBW” represents deep borehole water; “RW” represents river water; “C.B.” represents charge balance; “H.T.” represents hydrochemical type; “R.E.” represents recharge elevation; YBJT2 and ZK4001 (sampling depth was 1495 m) referring to [24,25].
Table 2. Calculations of reservoir temperature and cold water mixing ratio (a and b refer to [18]; c refers to [43]; d refers to [19]).
Table 2. Calculations of reservoir temperature and cold water mixing ratio (a and b refer to [18]; c refers to [43]; d refers to [19]).
NO.Quartz (a)Chalcedony (b)Multimineral Equilibria (c) Chlorine–Enthalpy (d)
TTMix
(°C)(°C)(°C)(%)
YBJS1157132−28146
YBJS2164141−29647
YBJS3167144−29645
YBJS4169147−29945
YBJT1174152200–22028842
YBJT2168145−25737
ZK4001−−268–276−−
Note: “−” means no calculation.
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Yang, H.; Yuan, X.; Chen, Y.; Liu, J.; Zhan, C.; Lv, G.; Hu, J.; Sun, M.; Zhang, Y. Geochemical Evidence Constraining Genesis and Mineral Scaling of the Yangbajing Geothermal Field, Southwestern China. Water 2024, 16, 24. https://doi.org/10.3390/w16010024

AMA Style

Yang H, Yuan X, Chen Y, Liu J, Zhan C, Lv G, Hu J, Sun M, Zhang Y. Geochemical Evidence Constraining Genesis and Mineral Scaling of the Yangbajing Geothermal Field, Southwestern China. Water. 2024; 16(1):24. https://doi.org/10.3390/w16010024

Chicago/Turabian Style

Yang, Hu, Xingcheng Yuan, Yongling Chen, Jiawei Liu, Chun Zhan, Guosen Lv, Junfeng Hu, Minglu Sun, and Yunhui Zhang. 2024. "Geochemical Evidence Constraining Genesis and Mineral Scaling of the Yangbajing Geothermal Field, Southwestern China" Water 16, no. 1: 24. https://doi.org/10.3390/w16010024

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