Contribution of deep-earth fluids to the geothermal system: A case study in the Arxan volcanic region, northeastern China

Abstract

Investigations of the hot spring water and gas in the volcanic region are involved in assessing geothermal resources and understanding groundwater circulation, volcano, and earthquake activities. The origins of water and gas of the hot springs, lakes, rivers, and rain in the Arxan volcanic region (AVR), northeastern (NE) China, were investigated by conducting a field survey and geochemical analysis. The low electrical conductivity (40–835 μS/cm) and low total dissolved solids (TDS, 23.83–540.00 mg/L) of the water samples indicate that they are fresh water. δ¹⁸O and δD values of the water samples range from −4.1% to −16.0% and from −61.3% to −119.9%, respectively. Enrichment of heavy isotopes in the rainwater and the crater lake waters was caused by evaporation. The component H₂O of the water samples predominantly originated from the meteoric water, with less than 1 vol% contributed by deep-earth fluids. Ions in the rain sample were predominantly derived from sea salt and continental aerosol. Ions in the surface water samples had multiple origins (mineral dissolution, atmospheric, and anthropogenic sources). While the ions in the hot spring water were predominantly derived from both the dissolution of rocks and deep-earth fluids, the latter contributed 73%–87% of Cl⁻ and 86%–99% of Na⁺ to the hot spring waters. Gases from the hot springs were composed of more than 95% N₂ and less than 5% O₂ and Ar, with ³He/⁴He ratios of 0.14–1.17 R_A (R_A=1.4×10⁻⁶). Excess N₂, Ar, He, and CO₂ of the hot springs were mainly derived from both the crust and upper mantle. About 3%–23% of the total He in the bubbling gases from the crater lake waters and hot springs is derived from the mantle, implying a supplement of heat energy from the mantle to the geothermal systems. Significantly, about 12% of the He dissolved in the Budonghe water is derived from the mantle, indicating that plenty of mantle-derived heat transported by deep-earth fluids keeps the river water from freezing. Our results indicate that Cl and Na ions and ³He/⁴He ratio are the feasible geochemical indicators for source partitioning of geothermal fluids.

1 Introduction

The origins of geothermal fluids in the volcanic areas are involved in investigations of natural resources, environment, and hydrological cycle. Ion concentrations and isotope ratios of groundwater are the efficient indicators of water origin, circulation process of groundwater, geothermal potential, and volcanic activity (Giggenbach et al., 1993; Shaw et al., 2003; Taran, 2009, 2011; Benavente et al., 2016; Tardani et al., 2016; Rizzo et al., 2019). Therefore, geochemical observations of fluids can reveal the fluid origins, magmatic and seismic processes, and dynamic processes in the deep earth.

It is a major challenge in geoscience to decipher the complexity of fluids and minerals in the earth’s interior. Sources of dissolved solids in groundwater can be divided into three categories: dissolution of minerals in circulation, deep-earth fluids, and atmospheric and anthropogenic sources. The deep-earth fluids, meaning waters, gases, and supercritical fluids that are derived from the deep crust and mantle, may be composited with different genetic types of fluids such as initial, magmatic, metamorphic, deep hydrothermal, and diagenetic fluids that usually mix into geothermal fluids. The origins of ions and gases can be traced by using concentrations and isotopic ratios of the chemical species and techniques of statistical analysis (Han et al., 2019; Li, 2020; Chen et al., 2021). A variable non-sea-salt contribution usually causes the slope to differ from the seawater ratio and the intercept to differ from zero (Keene et al., 1986). Cl⁻ and Na⁺ are the most reliable indicators for discriminating and partitioning sources of ions in the spring waters because they tend to exist in aqueous solution during water circulation, while other ions are partially fixed in secondary minerals.

Ion concentrations, δ¹⁸O and δD, of water have been widely used for investigating the origins of geothermal fluids. For instance, the high values of δ¹⁸O, δD, δ¹³C, and Cl⁻ concentrations of spring waters in Oita Plain, Japan, implied an origin of andesitic-magmatic steam or metamorphic water, while the low values of δD and Li and B concentrations of groundwater indicated meteoric water (Amita et al., 2005). Li et al. (2012) investigated the water vapor sources of meteoric water in NE China using δ¹⁸O and δD data. The hydrochemical data indicated that the main stream water of the Kherlen River in eastern Mongolia, neighboring to the west of the Arxan volcanic region (AVR), were recharged by meteoric water from the headwater region of more than 1,650 m altitude (Tsujimura et al., 2007). δ¹⁸O and δD values of the lake, river, and well waters in the Lake Hulun Basin neighboring to the north of the AVR indicated that the waters originated from meteoric water and suffered evident evaporation (Gao et al., 2018).

The concentrations of gaseous components, He/Ne, ⁴He/²⁰Ne, N₂/Ar, ³He/⁴He, CO₂/³He, and δ¹³C, provide insight on the sources of the gaseous species in the geothermal and volcanic systems (Giggenbach et al., 1993; Shaw et al., 2003; Du et al., 2005; Taran, 2009; 2011; Xu et al., 2013; Benavente et al., 2016; Rizzo et al., 2019; Bini et al., 2022), secondary processes of geothermal fluids (Zhang et al., 2016; Barry et al., 2020), volcanic and seismic activities (Chen et al., 2014; Zhou et al., 2015; Tardani et al., 2016), and contribution of volcanic gas to environment (Sun et al., 2020; Zhao et al., 2021). The regional variations in ³He/⁴He, δ¹³C of CO₂, δ¹⁵N, and ⁸⁷Sr/⁸⁶Sr values of hydrothermal fluids along an intra-arc fault system in the southern Volcanic Zone of the Chilean Andes indicated the geothermal gas was a mixture of mantle-derived gas and radiogenic gas in the crust (Tardani et al., 2016). The molecular and isotopic data of hot spring fluids and fumarolic gases in the Tateyama volcanic hydrothermal system in Japan (Seki et al., 2019), the Tengchong volcanic area (Du et al., 2005) and western Sichuan Province (Du et al., 2006) in southwestern China, and the Wudalianchi volcanic field (Du et al., 1999; Xu et al., 2013), Changbaishan volcano (CBV) (Wei et al., 2016), and the Songliao continental rift system (Zhao et al., 2019) in NE China indicated the gases were predominantly derived from the mantle and crustal sources.

The mantle worldwide and regionally appears heterogeneous. The crust and upper mantle in NE China are characterized by fluctuation of Moho depth, mantle uplift in the rift valley and geophysical and geochemical heterogeneities (Jia and Zhang, 2020). The crust thickness in the Greater Khingan Range orogenic belt ranges from 34.5 to 43.5 km, while it becomes 32.4–36.2 km in the Songliao Basin in the east (Li et al., 2014; Jia and Zhang, 2020). The statistic histogram of 279 ³He/⁴He ratios of mantle xenoliths in eastern China shows multiple peaks with a wider range from 0.1 to 12 R_A (excluding eight individual data from 12.1 to 33 R_A; R_A is atmospheric ³He/⁴He =1.4×10⁻⁶) (Cui et al., 2022). Isotopic ratios of noble gases in mantle xenoliths in the orogenic belt and rift valley indicate that the mantle in NE China is heterogeneous. Obviously, the geochemical characteristics of the mantle in NE China differ from those in the middle ocean ridge and volcanic arc (Sano and Marty, 1995; Shaw et al., 2003; Roulleau et al., 2015).

The previous investigations in AVR mainly involved geology, volcanology, geothermics and hydrology (Tang, 1984; Sun, 1999; Zhang, 2017; Gu et al., 2017; Chen et al., 2021). The hydrogeochemical investigations for 36 hot springs (defined as their temperatures being 5°C higher than the local annual atmospheric temperature) around the Hot Spring Museum, a natural museum established for tourism and spa in the Arxan city, suggested that the hot spring waters were chemically classified into Ca·Na-HCO₃·SO₄, Na·Ca-HCO₃, and Na-HCO₃ (Gu et al., 2017; Zhang, 2017), and the residence time of high-temperature hot spring waters ranges from 70 a to 90 a, and that of low-temperature hot springs/shallow groundwater is about 10 a (Gu et al., 2017). Recently, the groundwater in the basalts of Arxan was considered exogenous water from the Tibetan Plateau based on the data of the water balance relationship and δ¹⁸O and δD values (Chen et al., 2021). Zhao et al. (2021) reported that gases from two hot springs in the AVR were mainly composed of N₂ with low ³He/⁴He ratios (c. 0.1 R_A), high ⁴He/²⁰Ne ratios (150–380), and δ¹³C values of CO₂ (-6.2% to -13.6%). They concluded that the spring gases were mainly derived from the crust. So far, the origins of hot spring fluids in the AVR remain debated. This paper aims at revealing the contribution of deep-earth fluids to the geothermal system in the AVR based on the molecular and ion concentrations and isotope compositions of the hot spring fluids.

2 Geological and hydrothermal setting

The AVR is famous for a lot of hot springs and Quaternary volcanoes. It is located at the west margin of northeastern (NE) China, where volcanic eruptions were very violent in the Cenozoic Era and produced more than 590 volcanoes and about 50,000 km² of exposed basalt (Liu et al., 2001). The altitudes in the AVR range from 900 m to 1,700 m (Figure 1A). The annual mean temperature is −2.5°C, monthly averages of temperature range from 10.5°C to −31.2°C with the lowest temperature of -45.7°C recorded on 1^st January 2001. Annual precipitation was 450 mm during 1951–2015, of which 80%–90% fell during June–September, whereas annual evaporation capacity is up to 1,116 mm (Gu et al., 2017; Chen et al., 2021).

FIGURE 1

Fan et al. (2011)).

The AVR tectonically belongs to the Greater Khingan Range orogenic belt that connects with the northern part of the continental rift valley on the east. There are mainly five groups of faults in the study area: NEE, NNE, NW, NNW, and EW trending faults. The NEE-trending Halaha River fault cuts the lithosphere with a length of ca. 500 km. The deep-cut faults provide two channels for Quaternary magma migration from the mantle: one displays a high temperature and fluid-enriched body approaching the depth of 10–12 km, and another has become cold above a depth of 30 km (Tang et al., 2005). The series of NNW-trending basement rifts deeply cut the crust, favoring the deep cycle of groundwater and the formation of hot springs.

Mesozoic igneous rocks are widely distributed in the AVR with a small area of Paleozoic clastic strata (Figure 1C). The Paleozoic strata are composed of Late Ordovician metamorphic clastic rocks and volcanic-sedimentary clastic rocks, Late Silurian metamorphic clastic rocks, and Early–Middle Devonian clastic rocks sandwiched with biogenic limestone. The Mesozoic strata are mainly Late Jurassic rhyolitic lava and volcanic clastic rocks. The Cenozoic strata are Neogene clastic rocks and black basalts and Quaternary sediments (Fan et al., 2011; Xie et al., 2011).

Magmatic rocks in the study area are mainly the Early Indosinian and the Late-Middle Yanshanian porphyroid potassic granite, monzonitic granite, and moyite and Neogene and Quaternary basalts (Figures 1C, D). Alkaline basalts in the Greater Khingan Range are characterized by high MgO and Ni concentrations, high CaO/Al₂O₃ ratios, enrichment of large lithophile elements and positive Nb-Ta anomalies, and moderately depleted Sr–Nb–Hf isotopic ratios. The basalts could be derived from the deep primary mantle, possibly from ancient primordial peridotite (Fan et al., 2003; Xue et al., 2019). Quaternary basalts in the AVR cover an area of ca. 100 km² and erupted from more than 54 craters in the Middle-Late Pleistocene and Holocene (Fan et al., 2011). The Yanshan Volcano is the youngest volcanic crater and has a radiocarbon age of 2,000 years (Bai et al., 2005).

Groundwaters in the study area can be classified into three types: (1) pore water in Quaternary fluvial sediments, discharging through river beds; (2) bedrock fissure water in Late Jurassic volcanic rocks, recharged by meteoric water and discharged through underflow to the river valley; and (3) vein-type fissure water in the fault zones (Chen et al., 2021). Surface waters include river water (the Halaha River and the Arshangole River), barrier lakes, and crater lakes. There are 45 hot springs around the Hot Spring Museum in Arxan town (Figure 1B), whose water temperatures range from 2 to 40°C. A unique phenomenon is that the distance between the low-temperature hot spring (2°C) and the high-temperature hot spring (40°C) is just 0.3 m, which could be caused by different amounts of deep-earth fluid recharge.

The heat flow in the Greater Khingan Range is 40.2 mW/m². The estimated Moho heat flow is 33.2 mW/m² (Sun, 1999). Though the regional heat flow is low, three geothermal fields have been found in the AVR (Qi et al., 2012). One is the Jinjiangguo (JJG) Valley geothermal field, which is composed of two geothermal zones controlled by the NE and NNW trending faults in granite and characterized by lower resistance. The geothermal water age was dated as 8,150 ± 90 a (Qi et al., 2012). The second one is the Yinjianggou Valley (YJG) geothermal field, which is controlled by a NNW-trending fault. The third one is the Wuliqiao geothermal field, which is controlled by EW- and NNW-trending faults (Qi et al., 2012). The thermal energy is derived from the mantle (magma) and radiogenic heat in the crust (Qi et al., 2012). The heat energy of the hot springs around the Hot Spring Museum could be mainly derived from radiogenic heat in the crust (Sun, 1999).

3 Sample and method

The water and gas samples were collected from the AVR in August 2010 (Figure 1). The field measurements of water temperature, electrical conductivity (EC), and pH were conducted with the portable instruments. The samples for hydrochemical analysis were collected with the 250-ml plastic bottles, and for stable isotopic analysis of H and O with the 2-ml plastic bottles. The gas samples were collected by the gas drainage method in 1,000-ml glass bottles and sealed with rubber caps (Du et al., 2006). Totally, 17 water samples were collected from the springs, crater lakes, river, and rain, and eight gas samples were collected including seven bubbling gases and one dissolved gas in the water of the Budonghe River (at the sampling site Ⅴ in Figure 1). The rain sample is the thundershower water dropped down from the building rooftop after filtering dust particles.

Concentrations of Li⁺, Na⁺, NH₄⁺, K⁺, Ca²⁺, Mg²⁺, F⁻, Cl⁻, NO₃⁻, and SO₄^2- were measured with a Dionex ICS-900 ion chromatography system with the standard configuration (reproducibility within ±2%). The HCO₃⁻ concentrations were measured by the standard titration procedures with a ZDJ-100 potentiometric titrator (reproducibility within ±2%) (Chen et al., 2014). Ion charge balance errors (I.B.) of the hydrochemical data are less than ±5%. The total dissolved solid (TDS) value was calculated by, , subtracting total ion concentration minus the half of HCO₃⁻ concentration (Table 1). Stable isotope ratios of H and O were measured with a MAT 253 mass spectrometer at Hehai University, and the δ¹⁸O and δD values were reported referring to the V-SMOW standard with errors of 0.1% and 2‰, respectively (Chen et al., 2021).

The N₂, O₂, Ar, CO₂, CH₄, and He concentrations of the gas sample were analyzed with a Finnigan MAT-271 mass spectrometer, with a precision of ±0.1%. Helium and neon isotope compositions of the gas samples were measured with an MM5400 mass spectrometer at the Laboratory of Gas Geochemistry, the Institute of Geology and Geophysics, the Chinese Academy of Sciences. δ¹³C values of CH₄ and CO₂ were measured with the GC-IRMS analytical system, a gas chromatography (Agilent 6890)–stable isotope ratio mass spectrometer, and the values of C¹³C are reported relative to PDB in per mill with an error of ±0.5‰ (Li et al., 2007; Zhou et al., 2015). It is worthy of mention that the He concentration and ⁴He/²⁰Ne ratios of dissolved gas in the Budonghe water are double of atmospheric values. The ³He/⁴He ratio was less than the atmospheric ratio, which indicates the sample was not contaminated by air during sampling and analysis.

4 Results

The temperatures of the hot spring waters are in a range of 2.3°C–37.5°C. The water temperatures of the springs no. 5 at the Hot Spring Museum and Wuliquan (at 2.5 km northwest Arxan town) are lower than those of Budonghe water (7.5°C). The water samples have low electric conductivity (40–835 μS/cm) and low TDS (23.83–540.00 mg/L). TDS values of the surface (river and lake) waters are lower than those of the hot spring waters. Abundance of the majority cations and anions is generally in the order of Na⁺>Ca²⁺>Mg²⁺>K⁺>NH₄⁺ and HCO₃⁻>SO₄²⁻>Cl⁻, NO₃⁻>F⁻, respectively. δ¹⁸O and δD values of the water samples range from -4.1% to −16.0% and -61.3% to -119.9%, respectively. The lake waters are more enriched in heavy isotopes (¹⁸O and D) than the hot spring waters (Table 1).

The molecular and isotope compositions of the gas samples are listed in Table 2 including some published data (Zhao et al., 2021) in the AVR for comparison. The N₂ concentrations of the gas samples from the AVR are more than 95%, and others together are less than 5%. ³He/⁴He ratios are in a range of 0.20×10⁻⁷ ∼ 1.64×10⁻⁶. ⁴He/²⁰Ne ratios of the surface waters are approximately equal to the atmospheric value (0.32), but those of hot spring waters are much higher than the atmospheric value. CO₂ concentrations are lower (0.13%–0.41%) and δ¹³C of CO₂ from -22.3% to -6.2%. δ¹³C of methane (C₁) from Tianchi is -51.5%, indicating the biogenic origin of C₁, but δ¹³C of C₁ from the geothermal well in JJG is 1.4‰, hinting at an abiogenic origin. ²¹Ne/²²Ne and ²⁰Ne/²²Ne ratios are approximated by the atmospheric values (0.029 and 9.78, respectively).

5 Discussion

5.1 Origins of the water components

5.1.1 Isotopic compositions of hydrogen and oxygen

Most δ¹⁸O and δD values of the water samples are scattered nearby the global meteoric water line (GMWL, δD = 8.17δ¹⁸O + 10.35, Rozanski et al., 1993) and the local meteoritic water line (LMWL) in NE China (δD = 7.20δ¹⁸O − 2.39, Li et al., 2012), but some shift far to the right side of the lines (Figure 2). The coordinates (−13.1% and −97.0%) of the intersection between the GMWL and LMWL can be regarded as the mean value of the initial source water (Han et al., 2019). The slop and intercept of the LMWL are smaller than those of the GMWL, indicating evaporation effect and multiple origins of waters (Tsujimura et al., 2007; Li et al., 2012; Han et al., 2019). The isotopic data of the spring and river waters in the AVR, the spring waters in Wudalianchi (Du et al., 1999) and CBV (Shangguan et al., 1996), and the river waters in Oita Plain, Japan (Amita et al., 2005), are scattered along the GMWL, indicating that those waters are mainly derived from the Pacific Ocean. This conclusion is supported by the fact that the meteoric water in northeastern China is mainly derived from the Pacific Ocean during the summer monsoon season. The wide ranges of δ¹⁸O and δD values (Table 1) indicated that the different kinds of waters experienced different processes of isotopic fractionation.

FIGURE 2

Rozanski et al., 1993); sold line (LMWL) is the local meteoric water line (Li et al., 2012); red circle is the intersection of the GMWL and LMWL; red solid circle is the rain sample; for comparison, our data and the reported data of spring and river waters in the AVR (Gu et al., 2017; Li, 2020; Chen et al., 2021), Oita HSW and Oita RW - hot spring water and river water in Oita Plain and Arita - deep geothermal fluid in Japan (Amita et al., 2005), WDLC - spring water in the Wudalianchi volcanic area (Du et al., 1999), CBS - data of hot spring water in the Tianchi volcanic area in Changbai Mountains (Shangguan et al., 1996), and Hulun Lake - lake water data (Liang et al., 2017) were plotted.

δ¹⁸O and δD values of the rain sample (AW-0) are scattered far from the LMWL (Figure 2). The heavier isotope compositions of the rain sample are concordant with those of rain in the Hulun Lake Basin (Liang et al., 2017), NE China (Li et al., 2012), and eastern Mongolia in summer (Tsujimura et al., 2007). Therefore, the heavier isotope compositions of the rain sample can be attributed to the seasonal isotopic variations of meteoric water. In the east-west extended climate zone from the Greater Khingan Range to the Mongolia Plateau, the isotopic compositions of H and O in meteoric water are characterized by obviously seasonal variation of more negative values in winter than those in summer (Tsujimura et al., 2007; Gao et al., 2018; Li, 2020).

The surface water samples (AW-1–AW-5) were relatively enriched in heavy isotopes but depleted in the TDS (<62 mg/L) (Table 1; Figure 2). δ¹⁸O and δD values of water samples from the Dichi (a crater lake of the maar volcano) and Budonghe are similar. The water samples from the Tianchi and Tuofengling (TFL) Tianchi, two crater lakes on the volcanic cones, are more enriched in heavy isotopes. The Dujuehu is a barrier lake, whose water is more enriched in D (Table 1). The d-excess (δD-8.17δ¹⁸O) of the lake water samples is much lower than that of the LMWL (Figure 2). Such isotopic shifts can be mainly attributed to the evaporation of raindrops and lake water caused by higher evaporation capacities.

δD and δ¹⁸O values of seven water samples collected at the Hot Spring Museum and three water samples from the hot springs in JJG and previous results (Sun, 1999; Gu et al., 2017; Li, 2020; Chen et al., 2021) scattered nearby the LMWL, which indicates that the hot spring waters in the AVR mainly originated from meteoric water (Figure 2). The depletion of heavy isotopes in the hot spring waters can be explained by the air temperature effect because the isotopic fractionations of O and H in meteoric water are mainly related to the thermal dynamics of water vapor condensation. Both δD and δ¹⁸O values of meteoric water and surface air temperature in north China (latitude higher than 36°N) have positive correlation. δ¹⁸O/T gradient was 0.52‰/°C, and δD/T gradient was 4.27‰/°C (Zhang et al., 2008). Let us take the Hailun station in the Songliao Basin (47.45°N, 126.93°E, an altitude of 236 m, an annual mean of 5.5°C, and 480 km to NE Arxan) of the Chinese Network of Isotopes in Precipitation as a reference, the mean values of δ¹⁸O and δD of meteoric water are −12.5% and −92.2% at the Hailun station (Liu et al., 2014). The temperature difference of 8.0°C between Hailun and Arxan results in the δ¹⁸O and δD differences of 4.2% and 34.2%, respectively. Consequently, the δ¹⁸O and δD values of meteoric water in Arxan are estimated to be, respectively, -16.7% and -126.4%. In reverse, the isotopic coordinates of intersection between the GMWL and LMWL, regarding the mean value of recharge source water, minus the O and H isotope fractionation caused by decreasing temperature, are equal to −17.3% and 121.2%, respectively, which are concordant with the isotopic ratios of the spring waters in the AVR (Table 1).

5.1.2 Ions in the waters

The TDS values of all water samples range from 23.83 to 540.00 mg/L, indicating the waters are fresh. The waters can be chemically classified into four groups, namely, Na-HCO₃, Ca-HCO₃, Ca·Mg-HCO₃, and Na·Ca-HCO₃ by Sokolov’s method (Sokolov, 1966). All the spring water samples are Na-HCO₃ type, four samples from the crater lakes are Ca·Mg-HCO₃, Na·Ca-HCO₃, Ca-HCO₃, and Na-HCO₃, one from the Budonghe River is Ca·Mg-HCO₃, and a rain sample is Ca-HCO₃ type (Table 1).

The results of the correlation analysis for the hydrochemical parameters show that there is a strong positive correlation (r > 0.8, Table 3) between TDS and Li⁺, Na⁺, K⁺, NH₄⁺, HCO₃⁻, Cl⁻, F⁻, and SO₄^2-, indicating that the TDS of the water samples is mainly controlled by deep-earth fluids and the dissolution of rocks. This is supported by the higher Na⁺ and Cl⁻ concentrations, TDS, and temperatures of the hot spring waters and the residence time (Gu et al., 2017). Cui et al. (2022) reported that 85% of the TDS of continental meteoric water was composed of SO₄^2-, Ca²⁺, NO₃⁻, NH₄⁺, and Cl⁻, and NO₃⁻ was mainly anthropogenic origin. NO₃⁻ has shown an obviously positive correlation with NH₄⁺, Cl⁻, Ca²⁺, and K (r> 0.5) and a weak positive correlation with TDS (Table 3), indicating the contribution of atmospheric and anthropogenic sources, While the weak negative correlation between the couple ions of Mg-Li, Mg-Na, SO₄-Mg, HCO₃-Mg, and TDS-Mg (Table 3) may reflect water–rock interaction (dissolution, ion exchange and deposition).

In the diagram of Na–Mg–Cl concentrations (Figure 3), the data of water samples from the AVR scatter far from the marine precipitation lines (Keene et al., 1986) and show evidently different slops, indicating the contribution of those ions from marine origin is negligible. The lines of Cl-Na and Cl-TDS in the water samples and solutions leaching granodiorite at ambient conditions (Du et al., 2010) have similar slopes and intersections, while those of Na-Mg are different. TDS, Cl, and Na in the samples are much higher than those in the leaching solutions (Figure 3), which indicate the ions mainly originate from Na-Cl-enriched deep geothermal fluids and the dissolution of Na-bearing minerals.

FIGURE 3

Keene et al. (1986); blue-dashed and brown-dashed point lines for solutions of soaking granodiorite grains for 168 h, which were extended for comparison, and two points for data of soaking 2,000-μm grains of granodiorite (Gran) and trachyandesite (Tran) under the ambient condition for 12 h (Du et al., 2010); point lines for data of this study. (B) Diagram of Na-TDS-Cl; the inset is at the enlarged left bottom corner; the short extended dashed line for data of solutions of soaking granodiorite grains of different sizes under the ambient condition for 168 h (Du et al., 2010).

Plagioclase is a main component of the granite, granodiorite, and andesitic basalts that are widely distributed in the AVR (Figure 1). Na and Ca cations can be dissolved into solution through the reactions between plagioclase and CO₂-bearing water.

The decomposition reactions of plagioclase can be promoted by an increase in temperature, CO₂ partial pressure, and a mixture of Cl-S-enriched deep-earth fluids.

The Mg/Na and K/Na ratios of the rain and surface water samples from the AVR are concordant with those of marine precipitation (Keene et al., 1986) and water-soluble ions (Mg/Na=0.6 and K/Na=0.4) in aerosol samples in NE China (Shen et al., 2007). This indicates that the ions in the rain sample and the surface waters are mainly derived from sources of continental aerosol, sea salt, and rock dissolution.

The Cl⁻ concentrations of the rain and surface water samples are much lower than those of the hot spring waters, except for the sample AW-14 (Table 1; Figure 3), indicating the contribution of atmospheric and anthropogenic Cl to the hot spring waters is neglectable. Additionally, the values of Cl⁻ and Na⁺ concentrations and Mg/Na, Cl/Na, and Cl/Mg ratios of rock solutions differ from those of the hot spring water samples (Figure 3). The Na⁺ concentration in the solution is about 2 magnitudes higher than Cl⁻ concentration if granitic and andesitic rocks are dissolved in equal proportion. The experiments of soaking basalt and trachyandesite grains display that concentrations of Cl⁻, SO₄^2-, and Na⁺ approach the highest values in short time, and then, the highest value of Cl⁻ of about 3 mg/L show no obvious variation with increasing soaking time, while others varied with soaking time (Du et al., 2010). Similarly, the observed and experimental results indicated that the meteoric waters picked up chloride more rapidly than they would be congruent dissolution of basalts, resulting in Cl⁻ of 1–5 mg/L in the spring water (Gislason and Hans, 1987; Gislason et al., 1993). Such data indicate that the amount of Cl⁻ in groundwater contributed by meteoric water dissolving igneous rocks ranges from 2 to 4 mg/L. Therefore, the ions in the hot spring water could be a mixture of rock solution and deep-earth fluids.

Cl⁻ percentages of dissolution of rock and deep-earth fluids in the hot spring waters can be estimated by the two-member linear mixing model (Figure 4). Assuming Cl⁻ concentration of the end member of dissolution of rock be 3 mg/L (Gislason and Hans, 1987; Du et al., 2010) and that of the end member of deep-earth fluids be 20,000 mg/L (Oita deep geothermal water Cl: 18,649–23,787 mg/L and Na: 12,213–15,813 mg/L (Amita et al., 2005)), the volume percentages of deep-earth fluids in the AVR hot spring waters were estimated in the range from 0.041 to 0.103; in other words, 73%–87% of Cl⁻ in the hot spring waters were derived from deep-earth fluids. Moreover, taking the Na⁺ concentration of 3 mg/L in the 168-h soaking solution (Du et al., 2010) as the end member of dissolution of igneous rocks and that of 13,000 mg/L in Oita deep-earth fluids as another, the volume percentages of deep-earth fluids in the spring waters were estimated in the range of 0.14–1.66; that is, 86%–99% of Na⁺ in the hot spring waters were derived from deep-earth fluids (Table 1). Obviously, less than 1 vol% of deep-earth fluids mixed in the hot spring waters can be ignored when assessing groundwater volume, but the contributions of ions and heat energy from the deep-earth fluids to the geothermal system are significant.

FIGURE 4

Amita et al. (2005); Cl AVR and Mg AVR are in this study; others (solutions of soaking granodiorite and trachyandesite) are after Du et al. (2010); LPR is the local precipitation region (rectangle); WRR is the water–rock reaction region (oval); DTR is the deep-earth fluid region; AMW displays approximately andesitic-magmatic water.

5.2 Origins of gases

5.2.1 Molecular compositions of the hot spring gases

The measured N–He–Ar abundance system in AVR obviously differs from the gaseous components in air-saturated water (ASW) (Table 2). The ratios of N₂/O₂ and Ar/O₂ of the samples are 44–67 and 0.07–1.11, respectively, except for 137.6 and 3.1 of the gas samples from the hot springs in JJG, which obviously differ from those of air (3.7 and 0.04) and those in ASW (1.81 and 0.03 (Weiss, 1970)). The N₂/He ratios of the gas samples range from 116.4 to 663.0, obviously differing from the atmospheric (15,000) and ASW values (11,400 (Weiss, 1970; Taran, 2011)). He concentrations in the gas samples are approximately 4 magnitudes larger than the atmospheric value, but the N₂/He ratios are about 2 magnitudes lower than ASW’s value, and the ⁴He/²⁰Ne ratios of 257.8–385.3 of the hot spring gases are much higher than the atmospheric value (0.318 (Porcelli et al., 2002)) and ASW’s value (0.222 at 3°C (Weiss, 1971)). The disparities of molecular parameters between the samples and air or ASW likely indicate that amounts of atmospheric N₂ and He carried by meteoric water to the hot spring gases are not significant.

Sources of gases can be illustrated by the He–Ar–N₂ ternary diagram (Giggenbach et al., 1993; Rizzo et al., 2019; Bini et al., 2022). O₂ and Ar in hot springs are often considered to be derived from air based on the fact that O₂ and Ar concentrations in the atmosphere are much higher than those in the lithosphere, and hot spring water is mainly derived from meteoric water. In the case where all gaseous components in underground water originate from air, a part or all of the oxygen consumed by microbe respiration and the oxidation of organic and inorganic matter in the water circuit may result in N₂-rich gas. The low CO₂ concentrations in the gas samples indicate that the O₂ consumption of microbe respiration and the oxidation of organic matter in the hydrological circuit are neglectable. Data of our two gas samples are plotted nearby a magmatic source in the He–Ar–N₂ ternary diagram (Figure 5, Taran, 2011), indicating the gaseous components mainly originate from deep-earth fluids (Benavente et al., 2016). The data of the same hot springs obtained at different times (Zhao et al., 2021) shift to the air source in Figure 5, probably hinting at temporal variation or air contamination. The excess N₂ (N₂^*), Ar (Ar^*), and amount of non-atmospheric components can be estimated using the concentration of N₂/O₂ and Ar/O₂ ratios of ASW (Taran, 2009). Assuming the measured O₂ is derived from air and without consideration of O₂ consume in the hydrological circuit, Ar* can be estimated by Ar*= (Ar-O₂/20.5)/Ar, and excess N₂ can be estimated by N* =(N₂-1.814 O₂)/N₂ (Table 2). The calculated results indicated that more than 95% of N₂ and more than 91% of Ar in the gas samples were of no-atmospheric origin (Table 2). N₂^*/Ar^* ratios ranged from 52.21 to 88.04, less than those of geothermal gases in the volcanic arc (107–388 (Roulleau et al., 2015)) and mantle (350 (Taran, 2011)). The low N₂^*/Ar^* ratios can be attributed to the contribution of radiogenic Ar produced by ⁴⁰K that is relatively enriched in the granodiorite and alkaline basaltic rocks (Figure 1). Higher concentration He can be attributed to mantle-derived He and radiogenic He produced by the U-Th series enriched in the granodiorite and acid-intermediate volcanic rock in the crust.

FIGURE 5

5.2.2 Isotopic ratios of helium, carbon, and neon

5.2.2.1 He

In most cases, helium in geothermal fluids originates from atmospheric, crustal, and mantle sources (Sano et al., 1985; Ballentine et al., 2002; Zhou et al., 2015; Bini et al., 2022). Several methods were proposed for source partitioning of He in geothermal gases. For example, Sano et al. (1985) used an equation set to calculate the He percentages of atmospheric, crust, and mantle sources on the basis of the given ³He/⁴He and ²⁰Ne/⁴He ratios of those sources. Using a plot of CO₂/³He vs. δ¹³C with the given values of the same parameters for the atmospheric, crust, and mantle sources, the source partitioning of He and CO₂ in the geothermal gases can be illustrated (Sano and Marty, 1995; Hilton, 1996). Using the corrected ³He/⁴He ratios deducted for atmospheric He by the ⁴He/²⁰Ne ratio of air or ASW, the no-atmospheric He can be estimated by the two-end member model of crust and mantle sources (Ballentine et al., 2002; Zhou et al., 2015).

The samples of gases in the hot springs in AVR are characterized by low ³He/⁴He and CO₂/³He ratios (Table 2), which obviously differ from those of hydrothermal fluids in the Wudalianchi volcanic area (low ³He/⁴He and high CO₂/³He) in the rift valley (Du et al., 1999; Xu et al., 2013) and Tianchi volcanic area (high ³He/⁴He and low CO₂/³He) neighboring the rift valley to the west in NE China (Shangguan et al., 1996), and much less than those of volcanic and geothermal gases in the volcanic arc and middle ocean ridge (Sano and Marty, 1995; Shaw et al., 2003). The ⁴He/²⁰Ne ratios of the gas samples are in the range of 0.34–385.3, and ⁴He concentrations and ³He/⁴He ratios of the spring gases are independent (Table 2). The CO₂/³He ratios of the gas samples are 3 magnitudes less than those of the mantle (MORB-type) (1.5×10⁹ (Sano and Marty, 1995)), 3–5 magnitudes less than the values of fumaroles and hot springs in the volcanic areas in Japan (7.74×10⁹–1.18×10¹¹ (Sano and Marty, 1995)) and in the Nicaraguan volcanic front (Shaw et al., 2003), and 8 magnitudes less than the values of sediment and limestone (1×10¹³ (Sano and Marty, 1995; Xu et al., 2013)). Therefore, it can be concluded that that He in the gas samples has multiple origins.

The ²⁰Ne/²²Ne ratios of the gas samples are close to the atmospheric value, and the ²¹Ne/²²Ne ratios are slightly less than the atmospheric value (Table 2), indicating that Ne is mainly derived from air. Assuming all the ²⁰Ne in the geothermal gases is derived from air, the measured ³He/⁴He ratios of geothermal gases can be corrected by ⁴He/²⁰Ne ratios of air or air-saturated water (Hilton, 1996; Ballentine et al., 2002; Xu et al., 2013; Zhou et al., 2015; Zhao et al., 2021). For eliminating the air contribution resulted from air dissolved in meteoric water into groundwater, the measured ³He/⁴He ratios were corrected by the ASW’s ⁴He/²⁰Ne ratio using the equation, (³He/⁴He)^* = {(³He/⁴He)_sample—r}/(1—r), where r = (⁴He/²⁰Ne)_{(air or ASW)}/(⁴He/²⁰Ne)_sample (Xu et al., 2013). ASW’s ⁴He/²⁰Ne ratio was calculated using the data of He and Ne solubility in water at 3 °C (4.814×10⁻⁵ ml/kg and 21.71×10⁻⁵ ml/kg (Weiss, 1971)), and ²⁰Ne content is 90.48% of total Ne (Porcelli et al., 2002), regardless of isotopic fractionations of He and Ne caused by air dissolution. The corrected ³He/⁴He ratios of the gas samples from the Tianchi, JJG Tianchi crater lakes, the Budonghe water, and spring no. 0 differ evidently from the measured ratios, but those of the hot spring gas samples with high ⁴He/²⁰Ne ratios and high He concentrations show no change (Table 2). The percentages of atmospheric He and no-atmospheric He in the gas samples can be calculated by the following equation:where is the measured value; Atm is the percentage of atmospheric He, is the atmospheric ratio; and is the ASW-corrected ratio or no-atmospheric He. The calculation results show the percentages of atmospheric He (Atm) in the hot spring gases are less than 73, and those of no-atmospheric He range from 27 to 100 (Figure 6; Table 2).

FIGURE 6

Helium derived from the crustal and mantle origins can be quantitatively estimated by the two-member mixing model using the ASW (or air)-corrected ³He/⁴He ratios (Ballentine et al., 2002; Zhou et al., 2015; Seki et al., 2019; Zhao et al., 2021). Traditionally, assuming the mantle be homogenous, the ³He/⁴He ratios of the mantle and crust sources were proposed to be 8 ± 1 R_A and 0.02 R_A, respectively (Sano and Marty, 1995). However, the chemical heterogeneity of the upper mantle must be emphasized to estimate the amount of mantle-derived He in the hot spring gas, which is involved in assessing heat flux and potential heat resources in the geothermal fields. The ³He/⁴He ratios of basalts and mantle xenoliths in the different tectonic regions such as the mid-ocean ridge, volcanic arc, subcontinent, and orogenic zone evidently differ from each other and vary in a wide region within the same tectonic unit (Graham, 2002; Porcelli et al., 2002; Xu and Liu, 2002; Zhao et al., 2021; Randazzo et al., 2022). The ³He/⁴He ratio for the European subcontinental lithospheric mantle source was proposed as 6.1 ± 0.9 R_A (Randazzo et al., 2022). The isotopic compositions of noble gases in mantle peridotite xenoliths in eastern China vary from place to place and are obviously lower than MORB values (Cui et al., 2022). The lithospheric mantle and the asthenosphere in the continent area are important sources of gases in the upper mantle, which differs from the orogenic ranges to the rift valley in NE China (Xu et al., 2013). The mantle peridotite xenoliths in the Arxan–Chahe region (Liu et al., 2001; Sui et al., 2012) are geochemically similar to those in the Hannuoba region (Song and Frey, 1989; E and Zhao, 1987). Both the Arxan and Hannuoba volcanic regions are located in the orogenic belt of the Greater Khingan Range–Taihang Mountains, neighboring the rift valley on the east. The alkali basalts in the two regions formed in Miocene–Holocene, of which the REE pattens and spiderweb diagrams are similar (Zhao, 2010). The mantle xenoliths found in the two regions are mainly Al-rich lherzolites (E and Zhao, 1987). Therefore, the upper mantle from the Hannuoba region to the south part of the Greater Khingan Range can be considered the same, implying that there is no obvious difference between the mantle sources of He in both the AVR and Hannuoba region in the Cenozoic Era. Therefore, ³He/⁴He of the mantle source in the AVR can be represented by the average (2.1×10⁻⁶) of ³He/⁴He (n=31) of mantle peridotite xenoliths in the Hannuoba region (Cui et al., 2022). On the assumption that the ³He/⁴He ratios of the crustal and upper mantle sources are 2×10⁻⁸ and 2.1×10⁻⁶, respectively, contributions of mantle-derived He to the total He of the hot spring gases were estimated in a range of 3%–23% using the ASW-corrected ³He/⁴He ratios and percentages of no-atmospheric He in the total He and percentages of crustal He range from 9 to 94 (Table 2). Source partitioning can be illustrated by the diagram of ³He/⁴He vs. ⁴He/²⁰Ne (Figure 6). The percentages of mantle-derived He in the high-temperature spring gases in Figure 6 are slightly higher than the calculated values (Table 2). In consideration of the uncertainty of the different methods, however, it is clear that the calculated results are in general comparable with the diagram of ³He/⁴He vs. ⁴He/²⁰Ne (Figure 6). He percentages of the primary mantle source in the gas samples range from 1% to 6%, estimated with the MORB value (Table 2). The bubble gas in the Tuofenglin crater lake has the highest percentage of mantle He and is composed of 68% of atmospheric He, 23% of mantle He, and 9% of crustal He. The high percentages of mantle He are found in the crater lakes (Figure 1), indicating that the mantle He migrates upwards to the surface through fractures in the channel of magma migration and mixes with atmospheric and crustal He during migration. The contribution of atmospheric He to the gases in the high-temperature hot springs in the Jinjianggou area is negligible. Gas in the low-temperature spring of no. 0 around the Hot Spring Museum (Figure 1B) contains about 3% of atmospheric He, but gas in the high-temperature hot spring on no. 34 lack of atmospheric He (Table 2; Figure 6). Such a small amount of atmospheric He may be carried into the spring by cool water recharge containing dissolved atmospheric He. Specially, the amount of mantle-derived He in the Budonghe water was estimated as high as 12%, which indicated that plenty of heat energy was transported by deep-earth fluids from the upper mantle to the Budonghe area in the Wuliqiao geothermal field. He in the hot spring gases with high concentrations is predominantly of crust origin, with less than 10% of mantle-derived He, indicating He accumulation in the geothermal reservoirs. The mantle-derived He emits upwards to the surface through the deep-cut faults and transports plenty of heat energy to the geothermal systems, which is supported by the data that about 80% of total heat flow is derived from the mantle in the study area (Sun, 1999).

5.2.2.2 CO₂

Concentrations of CO₂ in the hot spring gases in the AVR are less than 1%, and δ¹³C values are in the range of −15.9–−22.3% (Table 2). There are three scenarios for the origin of CO₂ in the spring gases in the AVR. The first one is that CO₂ in the hot springs is likely derived from the mantle based on the mean value of −22.6% (n = 105) of CO₂ in mantle xenoliths and minerals enclosed in Cenozoic basalts in eastern China (Cui et al., 2022). The statistical result of the δ¹³C values of CO₂ in the mantle xenoliths worldwide also shows a bimodal distribution with peak values of −5% and -25% (Deines, 2002). The second one is that the CO₂ is likely to have originated from biogenic and metamorphic gases in the crust. The δ¹³C values of CO₂ originating from Jurassic coal seams are in a range of −11%–−28% (Du and Liu, 1991). Biogenic CO₂ of concentrations less than 6% in the gas reservoirs and gas seepages in China has δ¹³C values between −10% and −22% (Dai et al., 1996). The third one is that the CO₂ is most likely a mixture of biogenic and abiogenic CO₂. δ¹³C data (−4.7%–-6.4%) of CO₂ in the fumarolic gases at the Longonot Volcano, Kenya, indicated a magmatic origin with minor contributions from biogenic CO₂ (Robertson et al., 2016).

5.2.2.3 N₂

Nitrogen, in some instances, is the main component in hot spring, fumarole, and volcanic gases, which were identified as a mixture of atmospheric, mantle, and crustal N₂ (Roulleau et al., 2015; Tardani et al., 2016; Zhao et al., 2021). N₂ is the predominant component in the hot spring gases in the AVR, and δ¹⁵N values are in a range of +1.3–+1.9% (Table 2). The excess N₂ content is larger than 92%. Combining with ⁴He/²⁰Ne and ³He/⁴He ratios and the lack of organic matter to produce lot of metamorphic nitrogen in the study area, it can be considered that the excess N₂ mainly originated from the mantle with mixing crust-derived N₂, as reported by Zhao et al. (2021).

6 Conclusion

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