Baseline Groundwater Quality before Shale Gas Development in Xishui, Southwest China: Analyses of Hydrochemistry and Multiple Environmental Isotopes (²H, ¹⁸O, ¹³C, ⁸⁷Sr/⁸⁶Sr, ¹¹B, and Noble Gas Isotopes)

Abstract

The baseline quality of pre-drilling shallow groundwater is essential for the evaluation of potential environmental impacts of shale gas development. The Xishui region in the northern Guizhou Province of Southwest China has the potential for shale gas development but there is a lack of commercial production. As for the future environmental concerns in this undeveloped area, this study presented the hydrochemical and isotopic characteristics of shallow groundwater and its dissolved gas before shale gas development and determined the sensitive monitoring indicators. Results showed that shallow groundwater with an average pH of 7.73 had low total dissolved solids (TDS) ranging between 102 and 397 mg/L, with the main water chemistry types of HCO₃-Ca and HCO₃-Ca·Mg. The quality of most groundwater samples satisfied the drinking water standards of China. The mass concentration of dissolved methane in groundwater was below the detection limit (<0.01 mg/L), suggesting the low baseline value of hydrocarbon. The shallow groundwater was mainly recharged by local precipitation based on water isotopes. Water chemistry was modified by the dominant dissolution of carbonate rocks and partial dissolution of clastic rocks, as indicated by δ¹³C-DIC, ⁸⁷Sr/⁸⁶Sr, and δ¹¹B. Evidence from carbon isotopes of dissolved methane and CO₂ (δ¹³C-CH₄ and δ¹³C-CO₂) and noble gas isotopes (³He/⁴He and ⁴He/²⁰Ne) demonstrated that the biogenic methane mainly originated from acetate fermentation and the dissolved noble gas was a result of the dissolution of air. Based on the geochemical and isotopic differences between shallow groundwater and flowback and produced water (including shale gas) from the Weiyuan and Fuling shale gas fields as well as shale gas from Xishui, this study has provided the sensitive monitoring indicators and methods for identifying potential pollution of regional shallow groundwater related to shale gas development in the future.

1. Introduction

Due to advances in horizontal drilling and hydraulic fracturing, the rapid increase of shale gas production in North America has affected the global energy pattern [1,2]. China has developed shale gas exploration since 2012 and has the largest technically recoverable shale gas resources worldwide, which are estimated to include 21.8 trillion m³ of natural gas [3,4]. The black shale from the Lower Silurian Longmaxi Formation is the main shale gas reservoir in South China, accounting for 20% of national shale gas resources [5]. In the Sichuan Basin, Southwest China, several large gas fields have been developed with the Longmaxi Formation as reservoir formation, including Fuling, Weiyuan, and Changning [6]. In the Guizhou Province of Southwest China, the Longmaxi Formation is mainly located in Northern Guizhou and its shale gas reserves can be evaluated to 1.83 × 10¹² m³, which has the potential for large commercial shale gas development [7,8,9].

However, large-scale hydraulic fracturing (HF) may cause potential risks to the geological environment [10,11,12]. One of the widely concerned environmental issues related to shale gas development is the groundwater pollution problems [13,14,15]. Groundwater pollution caused by shale gas production is mainly divided into stray gas (natural gas) pollution and wastewater pollution through different pollution pathways, such as wastewater leakage from impounding reservoirs, fluid leakage associated with casing and cementing defects during hydrocarbon extraction, and migration of deep reservoir fluids to shallow groundwater through induced faults [16,17,18]. Some earlier studies indicated that dissolved CH₄ of shallow groundwater in Pennsylvania increased as it moved closer to shale gas wells [13,19]. This is most likely due to the small sample size, which was from a planned sampling campaign in a known contamination area [20]. A few follow-up studies using a larger dataset and a geospatial analysis tool suggested that such an impact was negligible except for a few hotspots [20,21]. In Texas, specifically in Parker-Hood counties, a detailed noble gas study indicated that the stray gas in shallow groundwater was most likely from the Strawn Group rather than the Barnett Shale [22,23]. Furthermore, lacking pre-drilling baseline data of shallow groundwater quality makes these findings of stray gas contamination controversial [24,25]. Recent research has compared the post-drilling and pre-drilling data and found that high concentrations of methane and dissolved solids of shallow groundwater were not related to the shale gas development, but to natural processes in most areas [25,26,27]. Thus, establishing geochemical baseline datasets of shallow groundwater systems before shale gas production is necessary, which can not only provide direct evidence when contamination occurs, but can also reduce controversies of shallow groundwater pollution after hydraulic fracturing [28,29,30]. Currently, a few sites, such as Pennsylvania, USA [31], Colorado, USA [32], Quebec, Canada [33], North Yorkshire, UK [34], and Chongqing, China [35,36], have developed baseline investigation of groundwater before drilling shale gas wells.

This study aims to present a case for the comprehensive evaluation of shallow groundwater quality before shale gas production. The study area is located in the Xishui region of Northern Guizhou, where there are plans for shale gas development, but where there is not yet commercial production. Previous studies in the USA and in the Qaidam Basin in North China show that hydrocarbon isotopes of methane can be used to identify the stray gas contamination of shallow groundwater [13,19,25], while hydrochemistry and isotopes of boron, lithium, and strontium are accessible and sensitive to recognizing the shallow groundwater pollution from flowback and produced water (FPW) [37,38,39,40,41]. Thus, this study investigated the systematic geochemical baseline values of groundwater and dissolved gas from 21 springs supplying local drinking water, by using multiple hydrochemical and isotopic compositions, including major elements, trace elements, and environmental isotopes (²H, ¹⁸O, ¹³C, ⁸⁷Sr/⁸⁶Sr, ¹¹B, ³He). Comparing shallow groundwater (including dissolved gas) of Xishui with the FPW (including shale gas) from the Weiyuan gas field (FPW-W) [42], the Fuling gas field (FPW-F) in Sichuan Basin [43], and a shale gas sample from Xishui, this study also concludes that sensitive geochemical indicators and integrated methods of groundwater pollution identification should be included in a routine monitoring program.

2. Study Area and Sampling

The study area is located in the Xishui area (25°06’35″–28°50′15″ N and 105°50′20″–106°44′30″ E), a northern region of Guizhou Province, Southwest China (Figure 1). It is typical of a subtropical humid monsoon climate with an average annual temperature of 13.1 °C. The annual average precipitation is 1138 mm (max 1160 mm; min 777 mm), of which approximately 79% occurs in May–October. Multiple rivers developing in the area belong to the Wujiang river system, including the Chishui River and the Xishui River. The study area is a mountainous landform with rare arable land. The land use/land cover (LULC) in Xishui is mainly forest (62%) and farmland (31%) [44].

Tectonically, the study area belongs to the transitional zone between the northern part of the Central Guizhou Uplift and the southeastern region of the Sichuan Basin [45]. The sampling area is located in the northwestern limb of the Sangmuchang anticline, while the axial direction is 40–50° north by east [46]. The oldest stratum exposed at the core of the anticline is the Proterozoic dolomite rock. Those strata from the core to northwestern (from old to new) comprise the Cambrian dolomite, Ordovician shale and limestone, Lower Silurian shale, Permian and Triassic limestone, and Jurassic and Upper Cretaceous mudstone and sandstone [47]. As the major shale gas reservoir in the study area, the black shale of the Lower Silurian Longmaxi Formation is deposited in the deep water continental sedimentary environment, the same as those in the Sichuan Basin [45,48]. It is characterized by large thickness (50–200 m), sapropelic kerogen, and a moderate burial depth (generally less than 1–2 km) [49]. As for its organic chemical characteristics, it is high in organic matter content (w(TOC) of 0.5–3.92%), gas content (saturated adsorption gas content of 1.52–2.77 m³/t), and thermal maturity, while it is low in porosity and permeability [50,51,52], indicating the advantages of shale gas accumulation and hydraulic fracturing.

Shallow groundwater aquifers in the study area can be divided into six types according to the lithology, water-bearing media, water yield property based on lithology, and water abundance (Figure 1) [53]: (I) Fracture-cave water mainly occurring in the limestone and dolomite of Lower Cambrian to Lower Ordovician (Є₁–O₁), Lower Permian (P₁), and Lower Triassic (T₁), with the discharge of springs of 10–100 L/s. (II) Fracture-cave water commonly presenting in the limestone with little mudstone of Upper Permian to Lower Triassic (P₂–T₁) and Middle Triassic (T₂), with the discharge of springs of 1–10 L/s. (III) Cave-fracture water generally occurring in the shale with little limestone of Lower to Upper Ordovician (O₁–O₃) and Lower Silurian (S₁) (including Longmaxi Formation), with the discharge of springs of 0.1–10 L/s. (IV) Pore-fracture water only flowing in the sandstone of Upper Triassic (T₃), with a relatively rich water yield of 1000–5000 m³/d for a well with the drawdown value of 40 m. (V) One kind of fracture water existing in the shale of Proterozoic to Lower Cambrian (Pt–Є₁) and the sandstone of Upper Jurassic to Upper Cretaceous (J₃–K₂), with a groundwater runoff modulus of 1.0–2.0 L/(s·km²). (VI) Other fracture water distributing in the mudstone of Lower to Middle Jurassic (J₁–J₂), with an extremely low groundwater runoff modulus of less than 1.0 L/(s·km²). Because of the various lithology and the strong anisotropy of shallow aquifers, shallow groundwater in the study area may occur heterogeneously with mixing sources of geochemical components [54]. In Figure 1, the major direction of groundwater flow in the study area is from the watershed boundary to the Xishui River (the largest river in study area). Fracture groundwater in the study area is usually recharged by local rainfall through infiltration and represents the average chemical and isotopic compositions of rainfall in the past several months to several years based on the local hydrogeological conditions [53]. Therefore, the seasonal variation has been homogenized. Shallow groundwater drainages mainly occur in springs and recharges into river water [54].

As the Xishui region is one of the most favorable exploration zones of shale gas in northern Guizhou [9,49], several exploratory wells have been developed for evaluating potential shale gas resources of the Longmaxi Formation, including XK-1 (shown in Figure 1). However, the study area is mainly an underdeveloped area without conventional and unconventional oil and gas development as well as other industrial activities. Therefore, it is an ideal site to access the baseline quality of local shallow groundwater.

3. Experiment and Analyses

Because there was almost no well in the study area, the 21 shallow groundwater samples were all collected from springs (outcrops of groundwater) in June 2018, some of which were used to provide drinking water for local people. The sites of the springs were distributed in different types of aquifers with various lithology (like silicate and carbonate rocks) and diverse ages (Cambrian to Cretaceous) (Figure 1). Despite complex lithology and different ages of aquifers in the study area, we concluded all samples to represent the holistic shallow groundwater and provide general baseline quality of the local shallow groundwater system. A shale gas sample (SG-1) was collected from the scientific shallow shale gas well (XK-3) in March 2019.

Field parameters of water samples, including temperature, pH, oxidation-reduction potential (ORP) and electrical conductivity (EC), were measured in situ by a portable multi-parameter meter (HQ40D, Hach). The titration of HCO₃/CO₃ was accomplished in the field using a 16900 Digital Titrator (Hach) with 0.8 mol/L sulfuric acid for titration and phenolphthalein and methyl orange as indicators. After the filtration through 0.45 μm membrane in the field, all samples were sealed in different capacities of polyethylene bottles according to the sample volumes required by different chemical tests.

Major ions and water stable isotopes (δ¹⁸O and δ²H) were analyzed in the Institute of Geology and Geophysics, Chinese Academy of Sciences (IGG-CAS). Major cations (K⁺, Na⁺, Ca²⁺, and Mg²⁺) and anions (F⁻, Cl⁻, NO₃⁻ and SO₄²⁻) were measured by an Ion Chromatography (Dionex ICS1100) with a detection limit of 0.05 mg/L. After water samples were filtered by 0.22 μm membrane as sample pretreatment in the laboratory, δ¹⁸O and δ²H of samples were measured by a laser absorption water isotope analyzer (Picarro L1102–i) with detection precisions of ±0.1 and ±0.5‰, respectively. Trace elements were determined by Inductively Coupled Plasma Mass Spectrometry (ICP-MS) of NexION 300D in the Beijing Research Institute of Uranium Geology (BRIUG). For the concentrations of Hg and As, an atomic fluorescence spectrometer (ASF2202) of BRIUG was used for the measurement, with analytical limits of 0.6 and 0.1 μg/L, respectively. The content of Br^- was tested in BRIUG using a Gas Chromatography Analyzer (Panna A90), with a detection limit of 0.02 mg/L. For the measurement of the content of CH₄ in groundwater, each water sample was sealed in two 40-mL brown glass bottles added with hydrochloric acid as an antimicrobial agent and the samples were analyzed by Gas Chromatograph Mass Spectrometer (GC/MS) in Shanghai SEP Analytical Service Co., Ltd., with a detection limit of 0.01 mg/L.

The carbon isotope of dissolved inorganic carbon (δ¹³C-DIC) in groundwater was determined by Isotope Ratio Mass Spectrometers (IRMS) of MAT253 in the Laboratory for Stable Isotope Geochemistry, IGG-CAS. The detection accuracy is ±0.1‰ (standard value of VPDB). For the measurement of boron isotope (δ¹¹B), water samples were condensed to the required testing concentration and they were measured by Thermal Ionization Mass Spectrometry (TIMS) in the Qinghai Institute of Salt Lakes, CAS. The measurement precision is better than 0.4‰. The detection limit of δ¹¹B is generally 10–20 μg/L. The ratio of strontium isotope (⁸⁷Sr/⁸⁶Sr) was analyzed by TIMS (Phoenix) in BRIUG. The ⁸⁷Sr/⁸⁶Sr of the standard sample (NBS 987) from the National Institute of Standards and Technology (NIST) is 0.710248 ± 0.000008. The uncertainty of the measurement of ⁸⁷Sr/⁸⁶Sr is ±0.000008. ${\mathsf{\epsilon }}_{\mathrm{Sr}}^{\mathrm{SW}}$ represents the normalization of the ⁸⁷Sr/⁸⁶Sr ratio of a water sample, by using the strontium isotope value of modern seawater (0.709169) [38]:

$${\mathsf{\epsilon }}_{\mathrm{Sr}}^{\mathrm{SW}}={10}^4\left(\right({}^{87}\mathrm{Sr}/{}^{86}\mathrm{Sr}{)}_{\mathrm{sample}}/({}^{87}\mathrm{Sr}/{}^{86}\mathrm{Sr}{)}_{\mathrm{seawater}}-1)$$

(1)

The samples of dissolved gas were collected from 12 selected springs through a conventional negative pressure method. A dissolved gas sample was injected into a 50 mL glass bottle with pristine water using the method of dissolved gas collection [55] and the bottle was further stored in a 500 mL brown plastic jar full of the pristine water. The tests for gas components (volumetric percentage) and the isotopes of carbon and noble gas were accomplished in the Key Laboratory of Petroleum Resources Research, CAS. The gas compositions of gas samples were tested by IRMS (MAT 271). The relative standard deviation was ±5% and the detection limit was 0.0001%. The carbon isotopes of methane (δ¹³C-CH₄) and carbon dioxide (δ¹³C-CO₂) were measured by IRMS (Thermo-Fisher Scientific Delta Plus XP) with an error range of ±0.2‰. The isotope values of noble gas (³He/⁴He and ⁴He/²⁰Ne) as well as the volumetric percentage of He were analyzed by a noble gas isotope mass spectrometer (Noblesse, Nu instruments). The measurement precision is ±3%. The standard gas of AIRLZ2007, representing Ra, is from the air on Gaolan Mountain, West China, of which the value of ³He/⁴He is 1.39 × 10⁻⁶. The results of ³He/⁴He can be expressed as R/Ra:

R/Ra = (³He/⁴He)_sample/(³He/⁴He)_{standard gas}

(2)

The full results are provided in Tables 1 and 2, including samples of groundwater, dissolved gas, and shale gas. The geochemical data of FPW-W and FPW-F come from Ni et al. (2018) [42] and Huang et al. (2020) [43], respectively. Besides, the isotopic and geochemical data of shale gas in the Weiyuan gas field are from Cao (2017) [56]. Though the Longmaxi Formation in the study area is undeveloped for shale gas, it is reasonable to compare natural shallow groundwater (including dissolved gas) with the FPW (including shale gas) from two gas fields extracting shale gas from the Longmaxi Formation, for concluding the major differences and establishing methods for identifying pollution from FPW and shale gas.

4. Results and Discussion

4.1. Baseline Values of Shallow Groundwater

4.1.1. Hydrochemistry

TDS values varied between 102 and 397 mg/L, with an average value of was 237 mg/L. The hydrochemical types of shallow groundwater were mainly HCO₃-Ca and HCO₃-Ca·Mg with low contents of Cl and Na (averages of 2.0 and 2.6 mg/L, respectively), except that only two samples (XS-17, XS-18) belonged to HCO₃·SO₄-Ca due to the high percentages of milligram equivalent of SO₄ (26.6 and 25.9 meq%) (Figure 2). The cations of shallow groundwater were dominated by Ca and Mg, with the average percentages of milligram equivalent of 76 and 21 meq%, respectively while the anions primarily consisted of HCO₃ (average of 80 meq%) with a minor amount of SO₄ (average of 16 meq%). The concentrations of both potassium and fluorine were low in groundwater, with ranges of 0.1–2.8 mg/L and 0.1–0.3 mg/L, respectively. In the karst area of Guizhou province, shallow aquifers affected by the dissolution of carbonate rock usually develop the HCO₃-Ca and HCO₃-Ca·Mg water types and low TDS of groundwater, indicating that the dissolution of dolomite and limestone may contribute to the compositions of major ions of local shallow groundwater [57,58,59].

Compared with the undetected content of sulfate in FPW [42,43], local shallow groundwater samples had a relatively high content of SO₄, ranging between 13 and 77 mg/L, with an average value of 34 mg/L, but far lower than the drinking water standard of China (GB5749-2006) of 250 mg/L. The main sources of sulfate in groundwater can be from human activity, precipitation, oxidation of sulfides and dissolution of sulfate-rich evaporite minerals (i.e., gypsum, anhydrite) [60,61]. The residents live dispersedly in the study area and the sampling spring outcrops were far away from the residence, suggesting that it had little possibility of groundwater pollution from the sewage water. The saturation indexes (SI) of gypsums and anhydrite of shallow groundwater, calculated by PHREEQC 3.0 with the database of llnl.dat [62], were all lower than −1.5. The low values of SI may be related to a quick circulation rate, strong renewability, and a relatively low content of evaporite minerals in the local karst area, implying that the dissolution of sulfate minerals may have little impact on local groundwater. The average milligram equivalent ratio of (SO₄)/(SO₄ + HCO₃ + CO₃ + Cl + NO₃) of groundwater was 16%, which was lower than values of the area influenced by oxidation of sulfides in Southwest China (24–29%) [63], indicating that most local groundwater may be neglectfully affected by oxidation of sulfides. The sulfate content of rain in nine cities evenly distributed in Southwest China had an average value of 10.6 mg/L [64], while the enrichment factor (evapotranspiration from rain to spring) had the range of 2–5 based on some sites with similar hydrogeological conditions in Southwest China [65,66], suggesting that the spring water with no loss or other source during recharge should have the range of SO₄ concentrations between 21 and 53 mg/L. The theoretical range covered the SO₄ values of most local springs, but two samples of XS-17 (74 mg/L) and XS–19 (77 mg/L) significantly exceeded the range. These two samples also had high ratios of (SO₄)/(SO₄ + HCO₃ + CO₃ + Cl + NO₃) of 22% and 23%, respectively, close to the range of 24–29%, suggesting that the oxidation of sulfides may affect two samples. Thus, the sulfate of most samples was mainly derived from local precipitation, while two samples with high sulfate concentrations may be additionally affected by the oxidation of sulfide deposits.

The concentrations of NO₃ in groundwater varied between <0.05 and 10.5 mg/L, with an average value of 4.0 mg/L, except two samples (XS-15, XS-17) with high values of 63.9 and 28.4 mg/L, respectively. The NO₃ concentrations of most water samples were below the threshold limit of 44 mg/L for the drinking water standard in China, except for one sample (XS-15). A few corn fields were in the vicinity of the sampling sites of XS-15, and XS-17, and the high content of nitrate may be related to agricultural pollution such as nitrogen fertilizer.

Bromide concentrations of most water samples were below the detection limit of 0.02 mg/L and only five samples had quite low contents, ranging between 0.02 and 1.00 mg/L, with an average value of 0.25 mg/L. In China, there is no standard limit of methane for both drinking water (GB5749-2006) and groundwater quality (GB/T 14848-2017). In America, the defined action level of methane for hazard mitigation (10–28 mg/L), recommended by the US Office of the Interior, can be considered as a standard limit of methane for natural groundwater [67]. The methane contents of all groundwater samples were less than 0.01 mg/L and were far lower than the action level, implying that the CH₄ contents of shallow groundwater in the region were low background values. If the stray gas from shale gas wells leaks into shallow groundwater, it may cause the increasing mass concentrations of CH₄ of the polluted groundwater.

The concentrations of all measured trace elements in each sample were below the threshold limits of the drinking water standard (GB5749-2006). The contents of Hg and As were undetected, while other typical toxicity indexes like Cd, Cr, Pb, and Sb were all lower than the threshold limits. Compared with other trace elements, Sr, Ba, Zn, and Cu had much higher averages and wider ranges of variation. For example, the concentrations of Sr, Ba, Zn, and Cu were 18–1949, 5.1–236, 0.36–163, and 0.05–88 μg/L, with averages of 444, 53, 13, and 6.9 μg/L.

According to the data of FPW-W and FPW-F, the hydrochemical characteristics were distinctly different from those of shallow groundwater, shown in Figure 2 and Figure 3. The hydrochemical types of FPW-W and FPW-F were Cl-Na, with high TDS (averages of 18,568 and 62,075 mg/L) and salinity (averages of Cl contents of 10,998 and 38,756 mg/L) (Figure 2). Unlike the extremely low content of Na in shallow groundwater, the averages of concentrations of Na in FPW-W and FPW-F were 3648 and 11,286 times higher than those of shallow groundwater, as the result of a mixing of high-saline formation water entrapped in Longmaxi formation (Figure 3a). Besides, when HCO₃ is only derived from the dissolution of carbonate minerals (calcite and dolomite) and/or anorthite reacting with the aqueous soil CO₂ in groundwater, the milligram equivalent ratio of (Ca + Mg)/(HCO₃) should be 1:1 [68,69]. In Figure 3b, shallow groundwater samples mostly fell along the 1:1 line with the average equivalent ratio of 1.2, suggesting that the concentrations of dissolved inorganic carbon (DIC) were related to the carbonate dissolution. In contrast with the high correlation coefficient of groundwater (r = 0.91), both FPW-W and FPW-F showed no correlation between (Ca + Mg) and (HCO₃). Compared with shallow groundwater (Br/Cl: 0.009 to 0.18), FPW-W and FPW-F had the higher concentrations of Br (21–179 mg/L, 139–153 mg/L) with lower molar ratios of Br/Cl (1.7 × 10⁻³ to 3.9 × 10⁻³, 6.5 × 10⁻⁴ to 6.9 × 10⁻⁴). FPW-W and FPW-F from the Sichuan Basin were also abundant in some trace elements due to the origin of evaporated seawater [42,43]. Ba, Sr, Rb, B, and Li in both FPW-W and FPW-F were significantly higher than that of shallow groundwater by at least two orders of magnitude, suggesting that Ba, Sr, Rb, B, and Li may rise in groundwater mixed with FPW (Figure 3d–f).

4.1.2. Isotopes of Oxygen, Hydrogen, Carbon, Strontium, and Boron

δ¹⁸O and δ²H

The stable isotopes of oxygen (¹⁸O) and hydrogen (²H) can effectively trace the origin and hydrological processes of groundwater [70]. The values of δ¹⁸O-H₂O and δ²H-H₂O of shallow groundwater in the study area had ranges of −8.2 to −6.4‰, and −47.7 to −36.8‰, respectively (

Table 1. Chemical and isotopic composition of shallow groundwater samples of the study area.

Sample ID	Water Type	T (°C)	pH	EC (μs/cm)	ORP (mV)	TDS (mg/L)	Major Ions (mg/L)									Trace Elements (μg/L)
							Ca	Mg	Na	K	HCO3	SO4	NO3	F	Cl	Br	Ba	Sr	Rb
XS-1	HCO3-Ca	23.1	8.10	277	−46	170	56	2.2	0.1	0.3	173	18	0.2	0.8	0.1	<20	37.6	94	0.24
XS-2	HCO3-Ca·Mg	15.5	8.07	350	−14.9	218	44	22.3	<0.05	0.7	210	37	2.8	1.0	0.1	20	17.7	760	0.56
XS-3	HCO3-Ca·Mg	16.6	7.76	398	−26.3	240	50	24.7	<0.05	0.5	215	48	2.4	0.8	0.1	<20	11.6	913	0.58
XS-4	HCO3-Ca·Mg	19.0	8.13	276	−45.1	168	32	18.5	<0.05	0.7	178	15	4.6	1.0	0.1	<20	25.8	246	0.54
XS-5	HCO3-Ca·Mg	15.8	7.74	370	−21.7	241	50	24.9	<0.05	0.5	210	51	1.7	0.9	0.1	<20	12.9	1038	0.61
XS-6	HCO3-Ca·Mg	16.3	7.92	247	−31.5	159	32	18.0	<0.05	0.6	168	16	2.9	0.9	0.1	100	26.1	101	0.49
XS-7	HCO3-Ca	17.8	7.35	328	−1.5	230	75	2.0	2.0	1.2	222	13	8.4	2.5	0.1	1000	52.7	119	0.29
XS-8	HCO3-Ca	17.6	7.59	343	−14.1	235	76	3.3	1.1	0.4	234	26	1.8	1.3	0.1	<20	38.3	177	0.22
XS-9	HCO3-Ca	22.5	8.01	366	−38.3	235	72	5.4	1.8	0.6	198	39	6.1	2.4	0.1	<20	53.1	239	0.46
XS-10	HCO3-Ca	17.8	7.44	429	−5.7	286	82	9.6	2.6	0.9	249	50	5.5	1.4	0.1	<20	56.2	419	1.00
XS-11	HCO3-Ca	17.6	7.20	426	+5.5	287	99	3.3	0.7	0.2	305	22	<0.05	1.1	0.1	<20	171	159	0.14
XS-12	HCO3-Ca	18.1	7.12	514	+10	370	108	11.5	3.2	0.7	390	36	3.2	1.4	0.1	<20	123	637	0.76
XS-13	HCO3-Ca	21.6	8.15	397	−50.4	235	72	5.3	0.6	0.1	207	38	4.5	2.9	0.3	<20	9.4	621	0.21
XS-14	HCO3-Ca·Mg	20.4	8.29	338	−57.6	210	43	24.9	<0.05	0.3	231	19	2.9	1.1	0.1	<20	5.1	18	0.35
XS-15	HCO3-Ca	18.1	7.66	426	−22.2	397	76	6.9	5.0	0.8	146	54	63.9	7.6	0.2	<20	51.4	256	0.39
XS-16	HCO3-Ca	17.8	7.67	415	−22.9	268	76	13.8	<0.05	0.9	288	19	3.4	1.6	0.1	<20	50.0	119	0.87
XS-17	HCO3-SO4-Ca	18.1	7.27	540	+1.6	259	101	8.5	4.1	2.8	247	74	28.4	7.9	0.1	<20	20.2	1091	1.28
XS-18	HCO3-SO4-Ca	21.4	8.19	210.1	−50.4	143	34	5.0	5.2	0.6	108	30	<0.05	0.9	0.2	20	67.7	151	0.21
XS-19	HCO3-Ca	18.7	7.39	549	−6	379	103	13.4	6.0	0.7	310	77	7.5	2.9	0.2	<20	236	1969	0.38
XS-20	HCO3-Ca	15.2	7.93	235	−36.7	155	48	1.5	1.6	2.3	129	19	10.5	1.4	0.1	<20	23.5	79	1.59
XS-21	HCO3-Ca	18.0	7.25	152.3	+0.2	102	26	1.6	2.1	1.5	78	16	0.2	0.8	0.1	100	32.2	127	2.30
Sample ID	Trace Elements (μg/L)				CH4 (mg/L)		SI-Calcite		PCO2		δ18O (‰)	δ2H (‰)	δ13C-DIC (‰)		87Sr/86Sr		${\mathsf{\epsilon }}_{\mathbf{Sr}}^{\mathbf{SW}}$		δ11B (‰)
	Li	B	Cu	Zn
XS-1	0.24	2.92	0.50	6.64	<0.01		0.70		10−2.80		−6.8	−39.5	−12.6		0.71058		19.9
XS-2	0.53	3.89	0.08	4.15	<0.01		0.51		10−2.74		−8.1	−47.3	−12.0		0.70917		−0.01
XS-3	0.63	3.01	0.32	9.45	<0.01		0.28		10−2.41		−8.2	−47.3
XS-4	0.05	3.13	0.07	2.57	<0.01		0.45		10−2.84		−7.7	−44.4
XS-5	0.72	2.79	0.05	1.02	<0.01		0.24		10−2.40		−8.2	−47.7
XS-6	0.06	1.82	0.35	2.49	<0.01		0.18		10−2.67		−7.6	−44.3	−11.2		0.71236		45.0
XS-7	1.53	6.44	0.15	1.99	<0.01		0.11		10−1.97		−7.1	−43.8	−11.9		0.71222		43.0
XS-8	0.58	4.87	0.15	4.01	<0.01		0.36		10−2.19		−7.0	−40.6	−13.4		0.71202		40.2
XS-9	0.98	5.26	0.28	2.07	<0.01		0.74		10−2.66		−6.8	−38.9
XS-10	4.76	9.84	0.17	1.00	<0.01		0.25		10−2.02		−7.6	−45.0	−10.3		0.71071		21.8
XS-11	0.06	1.39	0.07	0.83	<0.01		0.18		10−1.69		−6.4	−37.4	−13.8		0.71159		34.2
XS-12	5.88	12.7	52.5	41.7	<0.01		0.23		10−1.51		−7.1	−42.1	−11.6		0.71037		16.9		6.44
XS-13	0.07	1.21	0.13	0.80	<0.01		0.88		10−2.79		−6.5	−36.8
XS-14	0.06	1.05	0.45	163	<0.01		0.82		10−2.90		−7.6	−43.9	−12.5		0.71168		35.4
XS-15	0.74	9.06	0.30	1.69	<0.01		0.23		10−2.46		−6.8	−41.4	−11.7		0.70797		−16.9
XS-16	1.11	7.08	0.17	0.36	<0.01		0.52		10−2.19		−6.9	−40.5			0.70965		6.8
XS-17	0.40	7.63	0.32	4.30	<0.01		0.15		10−1.85		−6.4	−36.9	−7.8		0.70801		−16.4
XS-18	0.41	8.69	0.56	1.08	<0.01		0.36		10−3.11		−7.3	−42.2	−12.7		0.71270		49.8
XS-19	10.6	39.3	0.86	2.52	<0.01		0.38		10−1.87		−7.2	−43.9	−13.6		0.71111		27.4		−7.79
XS-20	0.21	4.20	88.1	10.1	<0.01		0.25		10−2.80		−7.2	−39.4	−13.3		0.71283		51.6
XS-21	4.10	3.25	0.16	3.01	<0.01		-0.82		10−2.31		−7.9	−45.5	−12.6		0.71034		16.5

). The global meteoric water line (GMWL) [71] and the local meteoric water line (LMWL) are usually combined to reveal the origin of groundwater. The LMWL is composed of isotopic data of δ¹⁸O and δ²H obtained from the Zunyi station of the GNIP database [72]. The data from GNIP consist of monthly average values of stable isotopes from 1986 to 1991, so the LMWL is based on five-year data reflecting the overall situation of the area.

For the shallow groundwater, most samples fell along the LMWL, suggesting that shallow groundwater mainly originated from the local precipitation (Figure 4). However, most groundwater samples were located slightly above the LWML with a relatively high weighted average d-excess of 15.8‰ compared with an average annual d-excess of 11.8‰, implying that the source of precipitation may have more fractions of terrestrial recycling vapor when recharging the groundwater [73]. A similar phenomenon also occurred in Chishui City, which is adjacent to Zunyi City [74]. Besides, the δ¹⁸O and δ²H of precipitation will become depleted with the rise of altitude and the decline of temperature, such as the depletion of δ¹⁸O will be between −0.15 and −0.5‰ per 100 m rise of altitude [70]. The samples from the outcrops of springs above the altitude of 1000 m had more depleted values of δ¹⁸O (an average value of −7.4‰) than those below 1000 m (an average value of −6.8‰). For example, the samples of XS-2, 3, and 5 with most depleted isotopic values (average δ¹⁸O value of −8.2‰) were collected from outcrops located in Cambrian strata with higher altitudes, while the samples of XS-11, 13, and 17 with most enriched values (an average δ¹⁸O value of −6.4‰) were from the outcrops of springs with lower altitudes (Figure 4). Thus, the altitude effect may cause the differences of δ¹⁸O and δ²H between various groundwater samples.

Distinguished from shallow groundwater, the stable isotopes of FPW-W were much heavier, with δ¹⁸O and δ²H ranging from −1.3 to +2.5‰, and from −27.4 to −15.6‰, respectively (Figure 4). It was the result of the blending consisting of injected fracturing water with low values of δ¹⁸O and δ²H, and formation water derived from evaporated seawater with high δ¹⁸O and δ²H [42]. The values of δ¹⁸O (−1.6 to −1.3‰) and δ²H (−29.1 to −27.4‰) of FPW-F were slightly below the range of FPW-W, but they were significantly different from shallow groundwater [43].

δ¹³C-DIC

The content of dissolved inorganic carbon (DIC) and the variation of δ¹³C reflect the specific geochemical processes and the biogeochemical cycle of carbon [75,76]. The sources of DIC in groundwater mainly consist of dissolved CO₂ from soil and atmosphere, oxidation of organic matter, and water–rock interaction. In this study, the pH of groundwater was from 7.12 to 8.29. When the values of pH are between 6.4 and 8.3, and the DIC in water is mainly composed of bicarbonate (HCO₃⁻) [77,78]. Thus, HCO₃⁻ was the major species of DIC in groundwater.

The concentrations of HCO₃⁻ were in the range of 78 and 390 mg/L, with an average value of 214 mg/L. As for the partial pressure of carbon dioxide (P_CO2), the P_CO2 in groundwater were all higher than the P_CO2 of the atmosphere (10^−3.5), with a range from 10^−3.11 to 10^−1.51, indicating that other sources of DIC had important impacts on groundwater. Except for sample XS-21 that had the negative value of SI-calcite (−0.82), most values of SI-calcite in groundwater were larger than 0, ranging from 0.11 to 0.88, suggesting that shallow groundwater mostly had the over-saturated level of calcite. Meanwhile, the values of δ¹³C of DIC in groundwater varied between −13.8 and −7.8‰ with an average value of −12.0‰ without an obvious correlation with concentrations of HCO₃ (r = 0.33) (Figure 5). The δ¹³C-DIC values of groundwater samples overlapped the results of groundwater in Zunyi City (−14.4 to −6.3‰) [79]. According to the positive values of SI-calcite, the ratios of (Ca + Mg)/(HCO₃) close to 1:1, high P_CO2 within the range of soil CO₂ (10⁻³–10⁻¹), and high ⁸⁷Sr/⁸⁶Sr of some samples shown in Figure 6b, the values of δ¹³C-DIC in local groundwater may be affected by the dissolution of carbonate rocks, the exchange of carbon isotopes between DIC (HCO₃) in water, and soil CO₂ in an open groundwater system, and the input of silicate weathering (e.g., albite, anorthite).

Compared with the depleted δ¹³C-DIC of groundwater, the FPW from Fuling gas field had more enriched values of δ¹³C-DIC varying between +13.6 and +16.2‰ with HCO₃ contents of 300–338 mg/L (Figure 5). The positive values of FPW may be related to the methanogenesis process causing the large isotopic fractionation between CH₄ and CO₂ (and hence DIC), but not the normal processes of both rainfall-infiltration and mineral dissolution generally controlling the δ¹³C values of shallow groundwater in karst zones [43].

⁸⁷Sr/⁸⁶Sr

Due to negligible isotope mass fractionation and various compositions in rocks and minerals, strontium isotopes are usually used for tracing hydrochemical processes in different geological settings, such as in the sourcing and transporting of materials, migration and transformation of pollutants, mixing patterns of water systems, weathering, and water–rock interaction [57,80,81]. In most cases, carbonate rocks and evaporates have higher concentrations of strontium and lower ⁸⁷Sr/⁸⁶Sr compared with silicate rocks with lower contents of strontium and higher ⁸⁷Sr/⁸⁶Sr [82]. The carbonate rocks in Xishui are mainly marine carbonates. The ⁸⁷Sr/⁸⁶Sr ratios of marine carbonates throughout the Phanerozoic were approximately within a range between 0.7068 and 0.7091 [83]. Meanwhile, the results of some studies show that Sr derived from the weathering of silicate rocks commonly had a ratio of ⁸⁷Sr/⁸⁶Sr higher than 0.7150 [84,85,86]. The concentrations of Sr in shallow groundwater ranged between 0.02 and 1.97 mg/L, with an average value of 0.44 mg/L. The ratios of ⁸⁷Sr/⁸⁶Sr of groundwater varied from 0.70797 to 0.71283 and had an average value of 0.71083, while ${\mathsf{\epsilon }}_{\mathrm{Sr}}^{\mathrm{SW}}$ had a range of −16.9 to 51.6 with an average value of 23.5. In Figure 6a, most of the water samples were in the transition between carbonate and silicate with three samples in the scope of carbonate. Besides, the groundwater affected by silicate weathering commonly had higher ⁸⁷Sr/⁸⁶Sr ratios and lower δ¹³C-DIC values than the groundwater controlled by carbonate weathering [87]. In Figure 6b, the springs in the range of transition had a δ¹³C-DIC average of −12.4‰, while the others in the range of carbonate had a δ¹³C-DIC average value of −10.5‰. Thus, silicate weathering partially affected the compositions of carbon and strontium isotopes for shallow groundwater. The relatively high ratios of ⁸⁷Sr/⁸⁶Sr also existed in the groundwater of urban areas in Zunyi, with a range of 0.70729 to 0.71727, reflecting the input of silicate weathering [88].

Compared with local groundwater, the FPW-W and FPW-F had higher contents of Sr (34–109 mg/L, 447–542 mg/L) and larger ratios of ⁸⁷Sr/⁸⁶Sr (0.71926–0.72001, 0.72397–0.72413), related to the mineral dissolution of bulk host shale (0.740246–0.740248) [42,43].

δ¹¹B

The relatively large mass difference between boron isotopes (¹⁰B and ¹¹B) leads to various isotopic values of δ¹¹B in natural water (e.g., seawater, river, groundwater), ranging between −16 and +59‰ [89]. Among different waters, groundwater has a wide range of δ¹¹B, between −15.9 and +32.4‰ [90]. The boron in unpolluted groundwater may be derived from water–rock interaction, atmospheric input (e.g., meteoric cyclic salts), or mixing with water of different origin [91,92]. The variability of δ¹¹B in natural groundwater is the result of the large differences among the natural B sources and partitioning reactions during adsorption/desorption [93]. According to data from different sites in different countries [43,94,95,96,97], the δ¹¹B of fresh groundwater (TDS < 1 g/L without anthropogenic pollution) had a wide range of δ¹¹B, between −14 and +27‰ with B contents of 7.8–670 μg/L (Figure 7). Apart from one sample (XS-19) with a high concentration of B (39.3 μg/L), the other samples had relatively low concentrations of boron varying between 1.1 and 12.7 μg/L, with an average value of 5.0 μg/L, overlapping the range of fresh groundwater. Due to the excessively low contents of B in shallow groundwater, the δ¹¹B values of most samples were undetectable, except for the samples of XS-12 (+6.4‰) and XS-19 (−7.8‰). In Figure 7, both XS-12 and XS-19 were in the range of global fresh groundwater, implying that the B contents and δ¹¹B of local groundwater belonged to the natural background values of fresh groundwater. The depleted δ¹¹B value and relatively high B concentration of XS-19 was likely related to water–rock interaction, such as adsorption/desorption between water and aquifer sediment [98], leaching of host rock [93], and dissolution of clay minerals [94].

Although the δ¹¹B data of groundwater in this study were limited and groundwater had a relatively wide range of δ¹¹B, shallow groundwater had significantly lower B concentrations and δ¹¹B, in contrast with FPW from Weiyuan (B: 2.4–56.2 mg/L; δ¹¹B: +23.8 to +27.6‰) [42], Fuling (B: 7.3–8.4 mg/L; δ¹¹B: +23.9 to +26.1‰) [43], the Marcellus shale (B: 5.2–62.6 mg/L; δ¹¹B: +25.5 to +32.3‰) [39], and the Fayetteville shale (B: 2.4–21.1 mg/L; δ¹¹B: +26.4 to +33.2‰) [99]. But the molar ratios of B/Cl for FPW-W (6.5 × 10⁻³ to 21 × 10⁻³) and FPW-F (6.5 × 10⁻⁴ to 6.9 × 10⁻⁴) were within or below the range of local shallow groundwater (1.3 × 10⁻³ to 44 × 10⁻³). Due to the significant differences of δ¹¹B and B contents, it is possible to use boron concentrations and isotopes to identify the groundwater pollution related to FPW.

4.2. Origin of Dissolved Gas

4.2.1. Compositions of Dissolved Gas

The compositions of dissolved gas samples are shown in

Table 2. Volume fraction, isotopic composition of dissolved gas samples of shallow groundwater and a shale gas sample (SG-1).

Sample ID	N2 (%.vol)	O2 (%.vol)	Ar (%.vol)	CH4 (%.vol)	CO2 (%.vol)	H2S (%.vol)	He (%.vol)	H2 (%.vol)	δ13C-CH4 (‰)	δ13C-CO2 (‰)	R/Ra	4He/20Ne
XS-1	81.05	16.96	1.34	0.008	0.636	UD		UD		−20.4
XS-2	76.37	21.81	1.32	0.020	0.484	UD		UD	−53.1	−20.0
XS-3	81.62	16.33	1.38	0.023	0.646	UD		UD
XS-4	78.60	19.76	1.15	0.034	0.440	UD		UD		−19.2
XS-8	76.51	21.66	1.16	0.081	0.582	UD	0.0006	UD	−50.6	−20.3	1.68	0.46
XS-9	75.70	22.13	1.40	0.032	0.727	0.0009		UD	−51.5	−20.5
XS-11	74.63	22.02	1.41	0.006	1.924	UD		UD
XS-16	81.42	15.35	1.52	0.375	1.336	UD	0.0009	UD		−21.1	1.19	0.61
XS-17	85.32	10.69	1.58	0.149	2.261	UD	0.0007	UD		−19.2	1.59	0.50
XS-18	75.11	22.93	1.44	0.007	0.505	UD		UD		−21.6
XS-19	72.57	25.04	1.43	0.003	0.958	UD		UD		−21.9
XS-20	72.87	25.45	1.43	0.003	0.252	UD		UD		−20.6
SG-1	1.88	0.15	0.02	97.40	0.015	UD	0.0538	UD	−31.9		0.015	383

. In 12 gas samples, the volumetric percentages of N₂ had a narrow range between 72.57 and 85.32% with an average value of 77.65% while the percentages of O₂ varied between 10.69 and 25.45% with an average value of 20.01%. The composition of gas samples were similar to those of natural air with the percentages of N₂ and O₂ of around 78% and 21%, respectively, implying that the dissolved gas in spring water may be related to the air. The contents of CO₂ and Ar were lower than 2.5%, with averages of 0.90 and 1.38%, respectively. The percentages of CH₄ were in a range between 0.0025% and 0.3754%, showing the low content of hydrocarbons. The proportions of H₂S were lower than the detection limit, except for the sample XS-9 with a value of 0.0009%. Only three samples (XS-8, 16, 17) had the percentages of He (0.0006, 0.0009, and 0.0007%). The values of H₂ were below the measurement limit. The acidic gases (e.g., HCl and SO₂) were absent in the gas samples. As for the shale gas sample (SG-1), the volumetric percentages of CH₄, N₂, O₂, CO₂, Ar, and He were 97.40%, 1.88%, 0.15%, 0.015%, 0.02%, and 0.0538%, respectively. The percentages of H₂S and H₂ were lower than the measurement limit.

4.2.2. Carbon and Helium Isotopes

The geofluid system may have different sources of dissolved methane, including biogenic, thermogenic, and abiotic methane. Understanding the origin of methane in shallow groundwater is important, which can be used to identify the sources of potential methane pollution in groundwater, like stray methane from shale gas wells [11,25,100]. The carbon isotope is a useful tracer for determining the origin of dissolved methane in the water system and for tracing the stray gas pollution [13,19]. The impact of the degassing sampling method on the isotopic composition of methane can be negligible because the low isotope fractionation between CH₄ (dissolved) and CH₄ (gas) (1.00033–1.0005) leads to little difference between the δ¹³C-CH₄ (dissolved) and δ¹³C-CH₄ (gas) (0.3–0.5‰) [101,102].

The δ¹³C values of biogenic methane are generally lower than −50‰, while the δ¹³C-CH₄ of thermogenic and abiotic origin can be commonly higher than −50‰ [70]. Besides, the abiotic methane derived from the modern seafloor hydrothermal vents has enriched δ¹³C values, ranging between −26 and −9‰ [103,104]. Based on the unique isotopic compositions of different kinds of methane, the origin of methane in geological systems can be determined. The δ¹³C-CH₄ value of a shale gas sample from Xishui was −31.9‰, suggesting the thermogenic origin of methane. The shale gas from the Weiyuan gas field had high values of δ¹³C-CH₄ varying between −37.3 and −34.3‰, indicating the typical thermogenic origin of methane without abiotic methane (Figure 8) [56]. Unlike the methane from shale gas, the δ¹³C-CH₄ values of shallow groundwater were lower, with a narrow range between −53.1 and −50.5‰, indicating the biogenic origin of dissolved methane.

The pathways of biological methanogenesis normally includes acetate fermentation and CO₂ reduction. Two pathways can occur solely or jointly in freshwater systems with the effect of mixed bacteria decomposing organic materials and producing CH₄ [70]. The fractionation of ¹³C between coexisting CO₂ and CH₄ (α¹³C_CH4-CO2) is an effective tool for identifying two pathways:

α¹³C_CH4-CO2 = (δ¹³C_CH4 + 1000)/(δ¹³C_CO2 + 1000)

(3)

The fractionation factor of CO₂ reduction producing CH₄ is generally less than 0.935, while that of acetate fermentation is usually higher than about 0.950 [105,106]. The α¹³C_CH4-CO2 values of shallow groundwater were all higher than 0.950 in a narrow range from 0.966 to 0.969 (Figure 8), demonstrating that acetate fermentation was the main origin of dissolved methane for local groundwater. In the reaction of acetate fermentation (2CH₂O → CH₃COOH → CH₄+CO₂, CH₂O stands for complex organic materials), the produced CH₄ and CO₂ will have δ¹³C values of −61 and +9‰, assuming δ¹³C for organic matter is about −26‰ and the enrichment factor between CH₄ and CO₂ is 70‰. However, the values of δ¹³C-CO₂ of shallow groundwater were more depleted, ranging between −21.9 and −20.0‰, close to soil CO₂ (−23‰ for C3 vegetation). The P_CO2 of groundwater (10^−3.11 to 10^−1.51) was close to the P_CO2 of soil (10⁻³ to 10⁻¹), showing that the local groundwater may be in an open groundwater system. Therefore, the depleted values of δ¹³C-CO₂ were probably related to both finite acetate fermentation with limited CH₄ and the open groundwater system dominated by soil CO₂.

Due to the inert chemical feature, noble gases are widely used to trace the origins of gas and fluid. As the different isotopic compositions of helium in atmosphere, mantle, and crust with little content of ²⁰Ne in mantle and crust, the ratios of ³He/⁴He and ⁴He/²⁰Ne can effectively identify the sources of gases (including fluids) and evaluate the influence of atmosphere in gases or fluids, such as thermal fluid and natural gas [107,108,109].

The ³He/⁴He ratio in the atmosphere (Ra) is around 1.39 × 10⁻⁶ [109,110], while ⁴He/²⁰Ne of the atmosphere can be 0.318 [107]. The ³He/⁴He ratios in other geological systems are usually expressed by R. The R/Ra ratio in the mantle, commonly represented by mid-oceanic ridge basalts (MORB), has an average value of 8.75 [111]. The typical R/Ra ratio of the crust is approximately 0.025 [112]. According to the different isotopic values of noble gas in the air, mantle and crust, the plot of R/Ra versus ⁴He/²⁰Ne can be applied for tracing the origin of gas, with the three end-members including air (R/Ra = 1, ⁴He/²⁰Ne = 0.318), mantle (R/Ra = 8.75, ⁴He/²⁰Ne = 100,000), and crust (R/Ra = 0.025, ⁴He/²⁰Ne = 100,000) (Figure 9). Although the ⁴He/²⁰Ne ratio of 100,000 is the assumed value for mantle and crust, the extremely high ⁴He/²⁰Ne can occur in the natural gas from several gas fields, such as the Hugoton-Panhandle gas field in the USA (up to 46,000) [113] and gas wells in Polish and Ukrainian Carpathians (up to 41,200) [114].

In Figure 9, the dissolved gas of shallow groundwater was close to the range of air, with the R/Ra ratios ranging from 1.19 to 1.68 and the ⁴He/²⁰Ne ratios varying between 0.46 and 0.61, indicating that the dissolved noble gas was mainly derived from dissolution of air. For the shale gas from the Weiyuan gas field, the R/Ra ratios varied between 0.019 and 0.025, with high ⁴He/²⁰Ne ratios of 22,570 to 44,505, conforming to the typical characteristic of crustal natural gas enriched in radiogenic ⁴He [56]. Compared with the shale gas from Weiyuan, the shale gas from Xishui had lower ratios of R/Ra (0.015) and ⁴He/²⁰Ne (383). Because of the large differences between shale gas and dissolved gas of groundwater in terms of isotopic compositions of both noble gas and CH₄, the isotopes of noble gas and CH₄ may help to recognize potential shale gas pollution occurring in local shallow groundwater during the shale gas exploitation.

4.3. Sensitive Indicators for Monitoring Groundwater

4.3.1. Sensitive Indicators

Due to the complex dissolved organic and inorganic components, wastewater generated from shale gas wells may cause potential or controversial environmental impacts on surface water, groundwater, and stream sediment [11]. Thus, the sensitive monitoring indicators are essential to monitor post-drilling groundwater quality and trace potential groundwater pollution caused by shale gas development. Most of the (semi) volatile organic compounds, pesticides and polychlorinated biphenyls (PCBs) are detected sporadically with low concentrations in FPW and therefore, the analysis of them may be unnecessary [115]. Based on the details of Section 4.1, FPW can be used as an end-member for identifying sensitive indicators. When FPW has significantly different geochemical and isotopic compositions compared with shallow groundwater, sensitive indicators can be gained through end-member comparison analysis [38,39,43]. Due to the lack of the shale gas wells in the Xishui region where the Longmaxi Formation is the future target stratum for shale gas, FPW-W and FPW-F are considered as hypothetic end-members for identifying indicators, assuming the compositions of ions and isotopes of the future FPW in Xishui will be similar to either FPW-W or FPW-F.

For the natural groundwater without contamination, the value of a specific ion and/or its isotopic composition has a random variation in a range of (average value ± 3σ) with a probability of 99.7% (Pauta criterion), where σ is the standard deviation. According to the determination method from Huang et al. (2020) [43], if shallow groundwater (SG) is mixed with X% of FPW, the specific ion and/or its isotopic composition in the polluted groundwater will change, where Δ represents the variation. For ions,

Δ_ion = ((X/100)C_FPW + (1 − X/100)C_SG) − C_SG)

(4)

For isotopes (e.g., δ¹³C, δ²H, δ¹⁸O, δ¹¹B, ${\mathsf{\epsilon }}_{\mathrm{Sr}}^{\mathrm{SW}}$),

Δ_isotope = ((X/100)C_FPWR_FPW + (1 − X/100)C_SGR_SG)/((X/100)C_FPW + (1 − X/100)C_SG) − R_SG

(5)

where C is the average concentration for specific ions of FPW-W or FPW-F and shallow groundwater (SG) and R is the isotopic value. Under the condition of |Δ| = 3σ, X% can be regarded as a threshold value of mixing rate from FPW-W and FPW-F causing statistical changes in ions or/and isotopes. The threshold values are put in order from smallest to largest and the first several indexes can be chosen as sensitive indicators, although sensitive indicators may vary from different sites with variable hydrogeological conditions.

Although the concentrations of most inorganic components (e.g., Na, Cl) in FPW increase with flowback time, we choose the averages of these components to represent major geochemical characteristics of the FPW. The results (Figure 10,

Table 3. The threshold values of specific ions and their isotopes during the mixing process between shallow groundwater from Xishui and FPW-W or FPW-F. X₁ represents the threshold value for identifying the mixing of FPW-W, while X₂ stands for threshold values for recognizing the mixing of FPW-F. GB5479-2006 stands for the drinking water standard of China.

Indicators	Units	GB5749-2006	FPW-W	FPW-F	Shallow Groundwater				X1 (%, |Δ| = 3σ)	X2 (%, |Δ| = 3σ)
			Average	Average	Max	Min	Average	3σ
EC	μs/cm	n/a	31,881.1	87,175	549	152	361	308	0.98	0.35
TDS	mg/L	1000	18,628	62,075	397	102	237	227	1.2	0.37
Na	mg/L	200	6937	21,459	6.0	<0.05	2.6	5.4	0.08	0.02
Ca	mg/L	n/a	228	1462	108	26.4	64.5	74.6	46	5.3
Mg	mg/L	n/a	32	172	24.9	1.5	10.8	24.5	113	15
HCO3	mg/L	n/a	548	318	390	78	214	211	63	203
Cl	mg/L	250	11,575	38,756	7.9	0.8	2.0	5.9	0.05	0.02
Br	mg/L	n/a	66	146	1.0	<0.02	0.23	1.1	1.7	0.78
B	μg/L	500	39,856	7840	39	1.1	6.6	23.8	0.06	0.30
Li	μg/L	n/a	16,810	38,777	10.6	0.05	1.6	7.7	0.05	0.02
Sr	μg/L	n/a	59,014	497,750	1969	18	444	1420	2.4	0.29
Ba	μg/L	700	14,2305	810,474	236	5.1	53.4	168	0.12	0.02
Rb	μg/L	n/a	529	737	2.3	0.1	0.6	1.6	0.29	0.21
δ18O	‰	n/a	0.8	−1.5	−6.4	−8.2	−7.3	1.6	20	28
δ2H	‰	n/a	−23.5	−28.3	−36.8	−47.7	−42.3	10	53	71
δ11B	‰	n/a	26.3	24.6	6.4	−7.8	−0.7	21.3	0.06	0.46
${\mathsf{\epsilon }}_{\mathrm{Sr}}^{\mathrm{SW}}$	mg/L	n/a	150	209	51.6	−16.9	23.5	62.9	0.74	0.05

) show that less than 0.1% of FPW-W would lead to significant changes of Li (0.05%), Cl (0.05%), δ¹¹B (0.06%), B (0.06%), Na (0.08%) for shallow groundwater while less than 0.1% of FPW-F would cause statistical rises of Li (0.02%), Cl (0.02%), Na (0.02%), Ba (0.02%), and ${\mathsf{\epsilon }}_{\mathrm{Sr}}^{\mathrm{SW}}$ (0.05%) for shallow groundwater. Besides, a mineral may precipitate from supersaturated conditions during the mixing of different waters. By using a mix of the geochemical model PHREEQC [62], if shallow groundwater blends with either 0.12% of FPW-W or 0.02% of FPW-F (threshold values of Ba), the saturation indexes (SI) of barite (BaSO₄) in both mixtures are 0.44, suggesting possible precipitation of Ba. However, the saturation indexes merely represent the tendency of hydrochemical processes (subsaturation means dissolution and supersaturation suggests precipitation) and only a few cases demonstrate the control of solutes by equilibrium with minerals [78,116]. The other ions (e.g., Ca, Mg, Br) and stable isotopes (δ¹⁸O and δ²H) are less sensitive to identify FPW with threshold values higher than 0.5%, such as at least 53% of FPW-W mixed with shallow groundwater causing statistical increases of δ¹⁸O and δ²H (shown in Table 3).

For collecting more indicators with enough sensibility to monitor the quality of shallow groundwater, those tracers of both FPW-W and FPW-F with threshold values below 0.1% can be considered as sensitive monitoring indicators. This means that a small portion of FPW releasing to shallow groundwater in Xishui can be likely recognized by using indicators of Na, Li, Cl, Ba, B, δ¹¹B, and ${\mathsf{\epsilon }}_{\mathrm{Sr}}^{\mathrm{SW}}$, which ought to be analyzed as a priority for tracing potential contamination of the future FPW in Xishui. For example, assuming B and Sr behave conservatively, a mixture consisting of 0.1% FPW-F and 99.9% shallow groundwater would lead to obvious increases of 13.7‰ for δ¹¹B and 98.3 for ${\mathsf{\epsilon }}_{\mathrm{Sr}}^{\mathrm{SW}}$, as well as significant rises in concentrations of Sr (112%), B (118%), Rb (115%), Na (828%), Ba (1463%), Cl (1910%), and Li (2416%) (using Mix of geochemical model PHREEQC).

Although the spring water samples were collected in the wet season (June 2018), the little temporary variability of geochemical and isotopic compositions of springs as groundwater in the study area has limited impact on the changes of threshold values of specific indicators. Springs in the study area represents the average chemical and isotopic compositions of rainfall in the past several months to several years [53] and the seasonal variation has been homogenized. According to the geochemical and isotopic data of spring water collected in Jiaoshiba, Chongqing with similar hydrogeological conditions [43], the major ions (Na, Cl), trace elements (B, Li,) and isotopes (δ¹³C-DIC) of springs had no significant variation between the dry season and the wet season. Besides, the geochemical (Na, Cl, B, Li, Ba, Sr) and isotopic compositions (⁸⁷Sr/⁸⁶Sr, δ¹³C-DIC) of a spring (unpublished data), collected in Xishui region in March, 2019 (dry season), were in the range of spring water samples (wet season) in this study. Assuming that the contents of major ions and trace elements of 21 spring water samples in the dry season would be abnormally three times higher than that in wet season (⁸⁷Sr/⁸⁶Sr and δ¹¹B are assumed to be stable), the threshold values of sensitive indicators used for identifying the mixing of FPW-W and shallow groundwater (including the dry and the wet season) would remain low, such as Na (0.21%), Cl (0.13%), B (0.14%), Li (0.11%), and δ¹¹B (0.13%). Although the δ¹¹B of most spring water samples are undetectable due to extremely low boron concentrations, the contaminated groundwater mixed with FPW would be likely qualified to be measured for δ¹¹B. For example, when 0.06% of FPW-W (threshold value) mingles with 99.94% of shallow groundwater, the boron concentration of polluted groundwater would be 30.6 μg/L, which can be totally qualified to the δ¹¹B measurement in the laboratory. Meanwhile, a spring, polluted by FPW in Jiaoshiba, had a high boron concentration of 60.3–69.0 μg/L and could be measured for δ¹¹B [43]. Therefore, if the pollution of FPW occurs, the polluted spring will be likely eligible to the δ¹¹B detection and the δ¹¹B can be used to verify the contamination of FPW.

The sensitive indictors should be included in the routine monitoring program and used in priority for tracing potential contamination of FPW during the shale gas development. As for the future shale gas development in the Xishui region, the routine monitoring parameters of shallow groundwater should contain the concentrations of Na, Li, Cl, Ba, and B, while ${\mathsf{\epsilon }}_{\mathrm{Sr}}^{\mathrm{SW}}$ and δ¹¹B can be further used to confirm potential contamination of PFW, assuming that the pollution trend and/or abnormal values occur. Meanwhile, when a small portion of FPW is mixed with shallow groundwater, some ions can reach or exceed the limit values of the drinking water standard, causing groundwater quality to deteriorate. For example, when the portion of FPW-F is 0.08%, the concentration of Ba in contaminated groundwater can reach the drinking water standard in China (700 μg/L), which means that 1 m³ of FPW-F can pollute 1250 m³ of natural shallow groundwater. Therefore, the hazard of FPW should be highly concerning.

4.3.2. Method for Identifying Pollution

Compared with the complex species of degradable organic components, it may be relatively accessible to trace the groundwater pollution related to shale gas development using stray gas and dissolved inorganic constituents. In this study, the methods could be composed of two parts for tracing pollution of inorganic components and stray gas.

For identifying the dissolved inorganic components of FPW, the pre-drilling shallow groundwater and FPW after drilling could be chosen as two end-members. Based on sensitive monitoring indicators discussed in Section 4.3.1, the concentrations of Na, Li, Cl, B, and Ba can be used to identify potential contamination trends of post-drilling shallow groundwater in the Xishui region, while isotopes of δ¹¹B and ${\mathsf{\epsilon }}_{\mathrm{Sr}}^{\mathrm{SW}}$ can further demonstrate whether FPW is the pollution source of groundwater. Elemental ratios of FPW are also different from that of shallow groundwater. For example, FPW-W and FPW-F had lower concentration ratios of Br/Cl (3.8 × 10⁻³–8.7 × 10⁻³, 3.4 × 10⁻³–4.1 × 10⁻³) and higher values of Sr/Ca (0.18–0.45, 0.32–0.35), compared with shallow groundwater with high Br/Cl (0.02–0.40) and low Sr/Ca (4.2 × 10⁻⁴–0.02). When 0.1% of FPW-F is mixed with 99.9% of groundwater, Br/Cl of contaminated groundwater would sharply decrease to 9.6 × 10⁻³, while Sr/Ca would significantly increase to 1.4 × 10⁻², in contrast with the high Br/Cl (average of 0.14) and low Sr/Ca (average of 6.8 × 10⁻³) of shallow groundwater. Therefore, specific elemental ratios should also be considered as sensitive indexes for tracing contamination of FPW. When the pollution of shallow groundwater occurs, the mixing ratio of FPW in the polluted groundwater can be appropriately quantified, assuming Cl, B, and Sr behave conservatively. According to the method from Huang et al. (2020) [43], assuming the polluted shallow groundwater (PSG) is a simple mixture of natural shallow groundwater (SG) and FPW without water–rock reactions, precipitation, or other hydrochemical processes, the mixing ratio (X%) of FPW in polluted groundwater could be approximately obtained based on the solute and isotopes mass balances. For specific ions (i.e., Cl), the X is given by:

X_ion = 100(C_PSG − C_SG)/(C_FPW − C_SG)

(6)

For isotopes of conservative ions (i.e., δ¹¹B and ${\mathsf{\epsilon }}_{\mathrm{Sr}}^{\mathrm{SW}}$), the X is given by:

X_isotope = 100(C_PSGR_PSG − C_SGR_SG)/(C_FPWR_FPW − C_SGR_SG)

(7)

where C is the average concentration for a specific ion of PSG, SG, and FPW, and R is the isotopic value of a specific ion. Although the contents of some ions (e.g., Na, Cl) in FPW increase with flowback time, the concentrations of these tracers for FPW at the early flowback stage are still distinctive to the shallow groundwater. For example, the Cl content of FPW-W can be up to 11,167 mg/L in the first day after hydraulic fracturing. Therefore, the averages of tracers are chosen to roughly represent the overall FPW. With the calculation of different sensitive indicators, an appropriate range of mixing ratios of FPW can be obtained, which can quickly and evidently assess shallow groundwater contamination caused by FPW.

As methane is the main component in shale gas (over 90% in Weiyuan shale gas), the stray methane is the major stray gas contamination. The δ¹³C of biogenic methane generally developing in shallow groundwater was significantly different to that of thermogenic methane in shale gas (Figure 8). It is, therefore, a logical deduction that if thermogenic methane of shale gas leaks into the shallow groundwater with limited biogenic methane, the δ¹³C of mixing methane in polluted groundwater is likely to shift, which may be affected by oxidation or diffusion. Besides, the ratios of R/Ra and ⁴He/²⁰Ne of dissolved gas in Xishui were significantly different from that of shale gas (Figure 9). Thus, with the dissolved gas of pre-drilling shallow groundwater and shale gas as two end-members, a deterministic model containing δ¹³C-CH₄, R/Ra, and ⁴He/²⁰Ne can be used to trace potential stray shale gas leaking into shallow groundwater in the Xishui region. In the zone of shallow groundwater in the Xishui region, δ¹³C-CH₄ of biogenic origin is from −50 to −80‰) [70], while R/Ra and ⁴He/²⁰Ne vary between the air (R/Ra = 1, ⁴He/²⁰Ne = 0.318) and the highest values of dissolved noble gas in Xishui (R/Ra = 1.68, ⁴He/²⁰Ne = 0.61). According to the data of shale gas from three gas fields in the Sichuan Basin (Fuling, Weiyuan, Changning) [56,117] and Xishui, the zone of shale gas can be restricted to the ranges of δ¹³C-CH₄ (−37.3 to −26.7‰), R/Ra (0.007 to 0.028), and ⁴He/²⁰Ne (383 to 44,505). If shale gas strays into a shallow groundwater system in Xishui during the shale gas development, the dissolved methane and noble gas in contaminated groundwater may be in the transition range between two end-members (Figure 11). Though this method needs to be further verified, it may provide a new tool to trace the stray gas contamination related to shale gas development. Besides, other gas indexes, such as δ¹³C-C₂H₆, δ²H-CH₄, and ³⁶Ar/⁴⁰Ar, should be added to the future baseline investigation of shallow groundwater and further provide more robust results about shale gas pollution, due to their different compositions between natural dissolved gas and shale gas [19,100].

5. Conclusions

This study provides baseline water quality before shale gas exploitation in the Xishui shale gas play. The geochemical and isotopic compositions for FPW are significantly different from shallow groundwater. Therefore, sensitive tracers can be obtained from end-member analysis. Taking shallow groundwater of Xishui and FPW from Weiyuan and Fuling as two end-members, some sensitive tracers (Na, B, Cl, Li, Ba, δ¹¹B, and ${\mathsf{\epsilon }}_{\mathrm{Sr}}^{\mathrm{SW}}$) should be analyzed as a priority during monitoring of groundwater quality in the study area. Besides, the indexes of δ¹³C-CH₄, R/Ra, and ⁴He/²⁰Ne can be used to identify the mixing of stray shale gas in the polluted groundwater. With these sensitive inorganic indicators and gas indexes, we can not only effectively monitor the change of the post-drilling groundwater quality, but also confirm and quantify the potential or possible groundwater pollution related to FPW and shale gas in future shale gas development.