Groundwater Geochemistry in the Karst-Fissure Aquifer System of the Qinglian River Basin, China

Abstract

The Qinglian River plays a significant role in China’s national water conservation security patterns. To clarify the relationship between hydrogeochemical properties and groundwater quality in this karst-fissure aquifer system, drilling data, hydrochemical parameters, and δ²H and δ¹⁸O values of groundwater were analyzed. Multiple indications (Piper diagram, Gibbs diagram, Na⁺-normalized molar ratio diagram, chloro-alkaline index 1, mineral saturation index, and principal component analysis) were used to identify the primary sources of chemicals in the groundwater. Silicate weathering, oxidation of pyrite and chlorite, cation exchange reactions, and precipitation are the primary sources of dissolved chemicals in the igneous-fissure water. The most relevant parameters in the karst water are possibly from anthropogenic activities, and other chemicals are mostly derived from the dissolution of calcite and dolomite and cation exchange reactions. Notably, the chemical composition of the deep karst water from the karst basin is mainly influenced by the weathering of carbonate and cation exchange reactions and is less affected by human activities. The hydrogeochemical properties of groundwater in the karst hyporheic zone are influenced by the dissolution of carbonates and silicates, evaporation, and the promotion effect of dissolution of anorthite or Ca-containing minerals. Moreover, the smallest slope of the groundwater line from the karst hyporheic zone among all groundwater groups revealed that the mixing effects of evaporation, isotope exchange in water–rock interaction or deep groundwater recharge in the karst hyporheic zone are the strongest. The methods used in this study contribute to an improved understanding of the hydrogeochemical processes that occur in karst-fissure water systems and can be useful in zoning management and decision-making for groundwater resources.

1. Introduction

The Qinglian River originates from Shikengkong (also called Mengkengshi), which is the highest mountain in Guangdong Province. As a principal source of the Beijiang River, a tributary of the Pearl River, the Qinglian River Basin is important for maintaining the ecological security of the Guangdong–Hong Kong–Macao Greater Bay Area and establishing a strong ecological security barrier in the southern hill and mountain belt of China. Groundwater is a crucial freshwater resource worldwide. However, some water points in the Qinglian River Basin have been abandoned due to water quality and quantity problems. Some water points meant for use as drinking water sources are only used for domestic washing, and most of the groundwater used as agricultural irrigation water is in a natural runoff state. Moreover, the forest land is mainly located in the northern granite distribution area, while the cultivated land and construction zones are mainly distributed in karst basins, depressions, and other karst areas, where groundwater is easily polluted [1]. The karst hyporheic zone, the lithologic contact area of granite and limestone (Figure 1), is an area where material and energy are actively exchanged between karst water and groundwater from the surrounding matrices (including sediment and bedrock) [2]. It is significant for microbial action, subsurface ecosystem evolution, and pollutant transport, adsorption, and conversion.

Groundwater availability is affected by many factors, including hydrogeochemical processes [3], urbanization [4,5], industrial [6] and agricultural production [7], and other human influences [1,8]. The combined application of hydrogeochemical and isotopic methods is an essential requirement for investigations into water cycle processes in groundwater systems, notably in karst hyporheic zones [2]. The hydrogeochemical characteristics of groundwater reflect the influence of various external factors on the groundwater system, help reveal the evolution of the groundwater system [4,9,10,11,12,13,14], and provide insights into cation exchange processes [15,16]. Isotope hydrology has a wide range of applications in the water cycle [17]. The difference in δ²H and δ¹⁸O values in the water cycle can be used to identify recharge sources [18] and water–rock interactions in the groundwater environment [19]. Numerous scientists have used hydrogeochemical approaches and isotope hydrology to examine groundwater systems [20,21,22], achieving significant success. For instance, Gibbs diagrams serve as a valuable tool for qualitatively analyzing the mechanisms underlying the origins of various ions in groundwater [23,24]. The Na⁺-normalized molar ratio diagram of groundwater encompasses representative lithological end members of carbonate, silicate, and evaporite minerals, facilitating the tracing of the solute sources and hydrogeochemical processes [25]. The principal component analysis (PCA) is perhaps the most common dimensionality reduction method, and parameters in the same principal component (PC) analysis with positive loadings indicate the same origin or similar geochemical behavior [26,27]. Additionally, the chloro-alkaline index 1 (CAI-1, [28]) and the mineral saturation index (SI) [29] have been extensively utilized in groundwater hydrochemistry studies alongside the δ¹⁸O values [30].

Groundwater quality in the Qinglian River Basin is intimately tied to groundwater geochemistry. Studies on the groundwater geochemical characteristics in this area have not yet been carried out. Hence, a detailed investigation of groundwater geochemistry is needed to evaluate the groundwater resources. In the present study, the primary hydrogeochemical properties controlling the groundwater quality in different landforms were analyzed based on hydrochemical parameters and δ²H and δ¹⁸O values of groundwater. The study of the karst-fissure aquifer system can provide a scientific foundation for managing the groundwater resources of the Qinglian River Basin. For example, the slight degradation of shallow karst water quality may be primarily associated with local agricultural activities. The water resource managers and practitioners in Yangshan County, Qingyuan City can gain significant insight into the physical processes and monitoring techniques concerning natural water quality.

2. Study Area

The study area is located in Yangshan County, Qingyuan City (latitude and longitude: 24°24′06″~24°55′38″ N and 112°33′22″~113°01′18″ E). The climate is characterized by subtropical monsoons, with an average temperature of 21.8 °C and an average annual rainfall of 1946 mm. The natural surface water system in the granite and part of the karst area is relatively developed, and the main stream is the Qinglian River. Surface runoff flows from northeast to southwest along the valley, traversing through the basin and into the Lianjiang River, with major tributaries including the Henglongqiao, Jiakeng, Kengzai, Huangben, Leicai, and Litou rivers (Figure 2).

2.1. Geological Structure

The study area has undergone several tectonic events, including the Indosinian, Caledonian, and Yanshanian movements [31,32,33]. These events resulted in a complicated geological structure with many giant folds and faults. In addition, some small fractures are distributed in the N−W and E−W directions, with most formed through brittle deformation (Figure 2). Most of the study area is underlain by limestone strata from the Triassic, Permian, Carboniferous, and Devonian periods. Karst strata are divided into two types: bare and covered. The area of exposed limestone is extensive, primarily located in the towns of Lingbei and Qinglian. Covered limestone strata occur primarily in the karst basin, with Quaternary alluvium overlying them. Few clastic strata were discovered, primarily Permian strata dispersed in a band in Lingbei Town and Huangben Town. The northern portion of the study area is composed of Jurassic and Cretaceous granites, which were formed during the Yanshanian Period [32]. According to core observations, slightly or moderately weathered granites (core Z01) and limestones (core Z03 and Z08 in the karst areas) are frequently found (Figure S1).

2.2. Landform

The Qinglian karst-fissure groundwater system covers approximately 1270 km² and is oriented northeast. In northern areas where granite occurs, the groundwater boundary coincides with the surface watershed. The ridgelines of Shikengkong, Shiqikengding, and Dashan mountains form the eastern boundary of the groundwatershed in these locations, while the western watershed is confined by the ridgelines of Yankengshan and Dongyue. However, the watershed for karst water is shared with the Yangcheng karst water system in the southwest and the Qingkeng karst water system in the southeast.

The geomorphology is classified into two types: karst erosion and dissolution geomorphology, and tectonic erosion (denudation) geomorphology. According to the morphological characteristics, it is further separated into six terrains: peak cluster valley, peak cluster depression, karst basin, eroded and denuded high and low mountains, and hills (Figure 1 and Figure 3a). The peak cluster valley is a landform composed of isolated peaks and cliffs, sinkholes, karst funnels, and stone buds developed between adjacent isolated peaks. The elevations typically range from 500 m to 800 m, with relative heights exceeding 300 m. The peak cluster depression is a landform composed of interconnected upward-pointing rock peaks that surround depressions [34]. These peaks are mainly conical with slopes ranging from 40° to 50°, and the bottoms of the depressions are mostly covered with 1–5 m of soil. Its elevation is less than 500 m, with relative heights varying from 50 m to 150 m. The center of the system is a karst basin bordered by large areas of clastic and granite bedrock and bare limestone. It covers 33.94 km² and accounts for 2.69% of the study area.

Figure 2. Study area in the southern hill and mountain belt, China [35] (a). Geological map of the Qinglian River Basin, Guangdong Province, China, (b), and the even gross geological section (c).

Figure 2. Study area in the southern hill and mountain belt, China [35] (a). Geological map of the Qinglian River Basin, Guangdong Province, China, (b), and the even gross geological section (c).

Figure 3. Location of samples for hydrochemical parameter analysis (a) and isotopic analysis (b).

Figure 3. Location of samples for hydrochemical parameter analysis (a) and isotopic analysis (b).

The elevation of eroded and denuded high mountain ranges from 800 m to 1902 m, and the relative elevation is greater than 400 m. The elevation of tectonic eroded low mountain is from 500 to ~800 m, and the elevation of eroded and denuded hills is less than 500 m.

3. Materials and Methods

3.1. Sample Collection and Analysis

A total of 53 groundwater samples, including 3 drill boreholes, 14 underground river outlets, 1 underground river inlet, 30 springs, and 5 wells (Figure 3a), were collected between July 2023 to January 2024 for hydrochemical parameter analysis. Notably, the sampling equipment for samples taken from boreholes mainly includes centrifugal and submersible pumps. The borehole was cleaned by pumping water, and the water discharged was more than 3 times the water stored in the borehole before sampling to ensure that the samples collected were representative. Two bottle samples were taken from each sampling site. First, natural water was extracted using a syringe and then filtered with 0.45 µm membrane filters (Milipores). The filtered water was poured into a 250 mL Teflon bottle with 2.5 mL nitric acid (pure nitric acid–ultra-pure water, 1:1; v:v) added in advance, and finally shaken to ensure pH < 2. The other 2000 mL Teflon bottle was filled with natural water directly after washing it with natural water 3 times. All water samples were collected and stored in a cool place before being sent to the laboratory.

The collected water samples were delivered to the Guangdong Geological Experimental Test Center for testing. The test parameters were sodium (Na⁺), potassium (K⁺), calcium (Ca²⁺), magnesium (Mg²⁺), chloride (Cl⁻), sulfate (SO₄²⁻), bicarbonate (HCO₃⁻), total dissolved solids (TDS), pH, total hardness (TH), and metasilicic acid (H₂SiO₃). The cations were tested using an inductively coupled plasma–optical emission spectrometer (Optima 8300DV, Perkin Elmer, Waltham, MA, USA) with the WinLab32 software. The quantified concentrations of calcium and magnesium were converted into their respective calcium carbonate equivalents, which are subsequently summed to determine the total hardness. The contents of HCO₃⁻ and Cl⁻ were determined using titration with a glass burette in the lab. SO₄²⁻ was determined with turbidimetry using ultraviolet–visible spectrophotometers (Uvmini-1240, Shimazu, Kyoto, Japan). The pH values of sample were measured with the glass electrode method using a pH meter (PHS-3C, Shanghai REX Instrument Factory, Shanghai, China) in the lab. TDS was measured using the gravimetry with a drying box (Shanghai Qixin Scientific Instrument Co., LTD., Shanghai, China) and a balance (METTLER TOLEDO, Zurich, Switzerland). The H₂SiO₃ contents were analyzed using spectrophotometry (S22PC, Shanghai REX Instrument Factory, Shanghai, China). All chemical reagents used in the test were analytical reagent-grade (AR, Table S1).

All parameter tests of groundwater referred to DZ/T 0064-2021 “Groundwater Quality Analysis Method” (Ministry of Natural Resources of the People’s Republic of China, 2021). The charge balance error (CBE) was used to check the accuracy of the data set [36,37,38]. The unit of cations and anions in the formula is meq/L. Equation (1) is as follows:

$$\mathrm{C}\mathrm{B}\mathrm{E}=\left({\displaystyle \frac{\sum Cations-\sum Anions}{\sum Cations+\sum Anions}}\right)\times 100$$

(1)

The accuracy of the results was generally within ±10% [13], and 89% of the samples had an accuracy between −5% and +5%, indicating that the analysis results were reasonable.

A total of 44 groups of atmospheric precipitation, surface water, and groundwater samples were collected for the analysis of hydrogen and oxygen isotopic values and tritium concentration, which were used to study the characteristics of the groundwater system in the Qinglian River Basin. The sample collection information is shown in Figure 3b and Figure S2. The first 22 samples in Table S2 were tested at the Institute of Karst Geology of Chinese Academy of Geological Sciences (IKG, CAGS), using an American L2130-i high-precision δ²H and δ¹⁸O water isotopic analyzer (Picarro, Waltham, MA, USA) with an autosampler controller, coordinator, and ChemCorrect software. The remaining 22 samples listed in Table S2 were tested at the China University of Geosciences (CUG), using the same isotopic analyzer (L2130-i, Picarro, Waltham, MA, USA) as the IKG and CAGS used. The standard deviations of hydrogen and oxygen isotope measurements were ±1.0‰ and ±0.20‰, respectively. The hydrogen and oxygen isotopic compositions were reported in the standard δ notation as per mil (‰) relative to the Vienna standard mean ocean water (V-SMOW) standard. Equation (2) is as follows:

δ_sample = (R_sample − R_VSMOW)/R_V-SMOW × 1000 (‰)

(2)

where R is the ratio of isotope abundance (²H/H and ¹⁸O/¹⁶O) for the collected water sample and the Vienna standard mean marine water standard.

In addition, Geographic Information System (GIS) software, namely MapGIS 6.7 (Wuhan Zhongdi Cyber S&T. Co., LTD, Wuhan, China), was initially used to view, edit, manage, and analyze geographic data. CorelDRAW Graphics Suite X8 (Corel Corp., Ottawa, ON, Canada) and Golden Software Grapher 20.0 (Golden Software LLC, Golden, CO, USA) were used for the post-production of the figures.

3.2. Analytical Methods

According to the Shukarev classification [39], the hydrochemical types of the water samples were determined. The Piper diagram [40] was drawn using Golden Software Grapher20.0.

The main sources of soluble ions in natural water, weathering and dissolution of rock or soil, atmospheric input (or sedimentation), evaporation and concentration, and anthropogenic effects, can be qualitatively analyzed by using a Gibbs diagram [23]. The ordinate of a Gibbs diagram refers to the content of TDS, and the abscissa refers to the content ratio of γCl⁻/(γCl⁻ + γHCO₃⁻) or γNa⁺/(γNa⁺ + γCa²⁺), where γ is the ion equivalent concentration (meq/L). Such a plot can distinguish the control factors of natural water chemical components, such as evaporation and concentration, rock weathering, or atmospheric precipitation [11,24]. When the dominant process is atmospheric precipitation, the TDS value is low and the γCl⁻/(γCl⁻ + γHCO₃⁻) or γNa⁺/(γNa⁺ + γCa²⁺) ratio is high, resulting in the sample data in the lower right corner of the plot. When the dominant effect is rock weathering, TDS values are moderate, the γCl⁻/(γCl⁻ + γHCO₃⁻) or γNa⁺/(γNa⁺ + γCa²⁺) ratio is low, and sample data are located in the middle region of the plot. When the main process is evaporation and crystallization, the TDS value and γCl⁻/(γCl⁻ + γHCO₃⁻) or γNa⁺/(γNa⁺ + γCa²⁺) ratio are high, and sample data are located in the upper right of the plot [23,41].

Generally, Ca²⁺ and Mg²⁺ are mainly derived from the weathering of carbonate and silicate minerals [42,43]. Na⁺ and K⁺ come from the dissolution of evaporite and weathering of silicate minerals [43], and HCO₃⁻ comes from carbonate and silicate minerals, and the soil and atmosphere in shallow conditions [12,44]. The representative lithologic end members of carbonate, silicate, and evaporite minerals can be shown in a Na⁺-normalized molar ratio diagram. Generally, the chemical composition of the carbonate end members is found to be as follows: Ca/Na ratio close to 50, Mg/Na ratio close to 10, and HCO₃/Na ratio close to 120. Therefore, the carbonate end members fall in the upper right corner of the diagram [45]. However, the molar ratios of Ca/Na and HCO₃/Na in the silicate end members are lower than those in the carbonate end members. The values of the silicate end members are Ca/Na = 0.35 ± 0.15, Mg/Na = 0.24 ± 0.12, and HCO₃/Na = 2 ± 1, which fall in the middle of the figure [43].

In order to understand the cation exchange between groundwater and the aquifer bedrock, the CAI-1 [46] of the groundwater samples was calculated. The index is usually calculated using the following formula:

$$\mathrm{C}\mathrm{A}\mathrm{I}-1={\displaystyle \frac{\gamma {\mathrm{C}\mathrm{l}}^{-}-\left(\gamma {\mathrm{N}\mathrm{a}}^{+}+\gamma {\mathrm{K}}^{+}\right)}{\gamma {\mathrm{C}\mathrm{l}}^{-}}}$$

(3)

The CAI-1 is commonly used to determine the direction and intensity of cation exchange [28,46]. When CAI-1 is negative, it suggests a groundwater environment with low salinity and chloralkali imbalance. Mg²⁺ and/or Ca²⁺ in the groundwater exchange ions with Na⁺ and/or K⁺ on the aquifer materials. When CAI-1 is positive, cation exchange takes place in the reverse sequence [19].

The SI is one of the most widely used indexes in groundwater hydrochemistry [29]. The formula is as follows:

$$\mathrm{S}\mathrm{I}=\mathrm{l}\mathrm{o}\mathrm{g}{\displaystyle \frac{\mathrm{I}\mathrm{A}\mathrm{P}}{\mathrm{K}\mathrm{T}}}$$

(4)

where IAP is the ion activity product of ionic species in solution and Ksp refers to the solubility product of a mineral. When SI = 0, the mineral is at equilibrium in an aqueous solution, also known as saturation. Undersaturation (SI < 0) causes mineral phases in groundwater to dissolve. SI > 0 indicates supersaturation, where mineral phases in groundwater tend to precipitate [24]. PHREEQC software was used to compute the SI of the groundwater samples.

In water–rock interactions, ¹⁸O in minerals and ¹⁶O in groundwater often exchange isotopes at low temperatures as follows:

CaC¹⁶O₂¹⁸O (calcite) + H₂¹⁶O⟺CaC¹⁶O₃ + H₂¹⁸O

(5)

$${\mathrm{Si}}^{16}{\mathrm{O}}^{18}\mathrm{O} \left(\mathrm{quartz}\right)+2{{\mathrm{H}}_2}^{16}\mathrm{O}\iff {\mathrm{Si}}^{16}{\mathrm{O}}_2+{{\mathrm{H}}_2}^{18}\mathrm{O}$$

(6)

$${\mathrm{CaAl}}_2{{\mathrm{Si}}_2}^{16}{{\mathrm{O}}_7}^{18}\mathrm{O} \left(\mathrm{anorthite}\right)+{{\mathrm{H}}_2}^{16}\mathrm{O}\iff {\mathrm{CaAl}}_2{{\mathrm{Si}}_2}^{16}{\mathrm{O}}_8+{{\mathrm{H}}_2}^{18}\mathrm{O}$$

(7)

The δ¹⁸O value in the groundwater becomes increasingly positive following isotope exchange reactions [30].

The SPSS 19 (SPSS Inc, Chicago, IL, USA) was performed on the PCA of parameters of Na⁺, K⁺, Ca²⁺, Mg²⁺, Cl⁻, SO₄²⁻, HCO₃⁻, pH, and H₂SiO₃ to explore the sources and influencing factors of groundwater chemistry [26]. TH and TDS were excluded since the former is the sum of dissolved Ca²⁺ and Mg²⁺, and the latter contains all soluble solids (including cations and anions) in the analysis.

4. Results

4.1. Hydrochemical Characteristics of Groundwater

The hydrochemical composition of the groundwater in the Qinglian River Basin is listed in

Table 1. Statistical summary of the geochemical composition of groundwater samples in the Qinglian River Basin.

Landforms	Statistics	pH	H2SiO3	TH	TDS	K+	Na+	Ca2+	Mg2+	HCO3−	Cl−	SO42−
		−	mg/L	mg/L	mg/L	mg/L	mg/L	mg/L	mg/L	mg/L	mg/L	mg/L
Eroded tectonic mountains (N = 9)	Max 1	7.57	36.80	34.00	110.00	2.64	15.60	12.60	0.60	56.00	8.00	6.80
	Min 2	5.86	17.90	3.30	25.00	0.88	2.32	1.06	0.05	11.00	1.00	0.10
	Ave 3	6.80	26.09	8.63	46.89	1.79	4.92	2.87	0.20	21.56	2.50	1.68
Karst hyporheic zone (N = 6)	Max	8.09	19.30	241.00	272.00	3.24	7.11	84.10	12.60	273.00	7.00	31.80
	Min	6.39	7.11	30.00	66.00	0.28	0.61	10.80	0.78	30.00	2.00	3.90
	Ave	7.29	12.10	135.27	166.00	1.32	2.40	47.40	4.09	154.50	3.00	9.88
Peak cluster valley (N = 20)	Max	7.86	10.20	295.00	322.00	2.78	3.72	104.00	18.30	332.00	9.00	30.20
	Min	6.96	4.90	156.00	175.00	0.24	0.27	58.50	1.56	190.00	1.00	0.80
	Ave	7.37	7.57	227.35	249.60	0.72	1.06	81.99	5.48	254.65	3.30	6.99
Peak cluster depression (N = 14)	Max	7.93	18.80	315.00	319.00	4.03	2.33	118.00	20.10	360.00	8.00	13.26
	Min	7.17	4.85	128.00	157.00	0.10	0.31	47.60	2.04	149.00	1.00	3.10
	Ave	7.50	8.74	225.27	238.85	1.18	1.05	80.45	4.56	242.37	3.13	7.63
Karst basin (N = 4)	Max	7.48	32.90	327.00	411.00	28.60	13.40	122.00	17.10	292.00	17.00	81.30
	Min	7.07	9.40	201.00	237.00	0.46	1.19	59.50	5.03	218.00	5.00	6.90
	Ave	7.22	16.73	246.75	325.75	8.30	6.29	84.88	8.45	240.25	11.50	33.83

¹ Max: maximum, ² Min: minimum, ³ Ave: arithmetic average.

and Figure S3. The pH of igneous-fissure water in the eroded tectonic mountains is low, ranging from 5.86 to 7.57, with an average value of 6.80. The groundwater in the karst hyporheic zone, peak cluster valley, peak cluster depression, and karst basin are mildly alkaline, with average pH values of 7.29, 7.37, 7.50, and 7.22, respectively. The metasilicic acid content in fissure water is the highest, with an average value of 26.09 mg/L. The shallow karst water in the karst basin has the second highest metasilicic acid content, with an average of 16.73 mg/L. The average metasilicic acid level in the karst hyporheic zone samples is 12.10 mg/L, falling between the eroded tectonic mountains and the karst area. However, the distribution of TDS and total hardness (measured as CaCO₃) is opposite to that of metasilicic acid. Groundwater samples from the eroded tectonic mountains have the lowest TDS and total hardness content, while karst basins had the highest levels. The chemical contents in the groundwater from the karst hyporheic zone fell somewhere between the first two. The levels of most parameters steadily rise from the recharge area (eroded tectonic mountains, karst hyporheic zone, peak cluster valley, and peak cluster depression) to the confluence area (karst basin).

From the recharge area to the confluence area, the Ca²⁺ and Mg²⁺ levels gradually increase. The Ca²⁺ and Mg²⁺ contents of groundwater in the karst hyporheic zone are between the first two. The content of Na⁺ is higher in the groundwater samples from the eroded tectonic mountains and karst basins, with average values of 4.92 mg/L and 6.29 mg/L, respectively. However, Na⁺ is very low in the karst hyporheic zone, peak cluster valley, and peak cluster depression, with average values of 2.40 mg/L, 1.06 mg/L, and 1.05 mg/L, respectively. The average K⁺ content of the karst water sampled in the karst basin is the highest, at 8.30 mg/L, which may be attributed to agricultural activities [47]. Other places exhibit low K⁺ content.

HCO₃⁻ is the main anion in the groundwater of the peak cluster valley, peak cluster depression, and karst basin, with average contents of 254.65 mg/L, 242.37 mg/L, and 240.25 mg/L. Igneous-fissure water has the lowest HCO₃⁻ content (21.56 mg/L). The content of HCO₃⁻ in the karst hyporheic zone is between that found in the karst water and igneous-fissure water, indicating that the igneous-fissure water circulates faster than karst water and is greatly influenced by precipitation. The karst basin has the highest average Cl⁻ level in groundwater, at 11.50 mg/L. This may be due to the gentle topography and significant evaporation within the karst basin [48]. Other locations contain moderate levels of Cl⁻. The average SO₄²⁻ level in the sampled igneous-fissure water is 1.68 mg/L, whereas the karst basin has the highest average (33.83 mg/L, Figure S3h). It is worth noting that the largest SO₄²⁻ content of igneous-fissure water is 6.81 mg/L at borehole Z01, and the rest are extremely low, ranging from 0.1 to 0.6 mg/L. The contents of SO₄²⁻ in the karst hyporheic zone, peak cluster valley, and peak cluster depression are similar, averaging 9.88 mg/L, 6.99 mg/L, and 7.63 mg/L, respectively.

The results of the PCA are summarized in

Table 2. Principal component (PC) loadings for chemical parameters in the Qinglian River Basin. Bold values are > of 0.60, bold and italics values are <−0.60.

Chemical Parameters	PCs—Igneous-Fissure Water		PCs—Karst Water			Pcs—Groundwater In The Karst Hyporheic Zone
	PC1	PC2	PC1	PC2	PC3	PC1	PC2
pH	0.21	–0.91	–0.26	–0.75	0.14	–0.65	0.08
Ca2+	0.98	0.01	0.09	0.96	0.12	–0.76	0.57
H2SiO3	0.72	0.01	0.85	–0.09	0.07	0.79	–0.48
K+	0.34	0.82	0.89	–0.09	–0.05	0.98	–0.19
SO42−	0.95	0.07	0.94	0.13	0.08	–0.07	0.97
Cl−	0.92	0.05	0.81	0.31	0.06	0.91	–0.11
Mg2+	0.95	0.23	0.08	–0.02	0.99	–0.27	0.94
Na+	0.99	0.02	0.94	0.21	–0.05	0.92	–0.25
HCO3−	0.99	0.00	–0.13	0.88	0.30	–0.72	0.59
Eigenvalues	6.26	1.55	2.44	4.02	1.12	4.86	2.84
Explained variance (%)	69.60	17.23	44.69	27.08	12.43	54.02	31.52
Cumulative % of variance	69.60	86.83	44.69	71.77	84.20	54.02	85.53

. The cumulative variance of PCA for the groundwater sample in the eroded tectonic mountains, karst areas, and karst hyporheic zone were 86.83%, 84.20%, and 85.53%, respectively, indicating that the dataset is acceptable for PCA.

The two-factor model controlled the groundwater quality parameters for the groundwater in the eroded tectonic mountains and karst hyporheic zone. In the igneous-fissure water, the most relevant parameters of PC1 are HCO₃⁻ (0.99), Na⁺ (0.99), Ca²⁺ (0.98), SO₄²⁻ (0.95), Mg²⁺ (0.95), Cl⁻ (0.92), and H₂SiO₃ (0.72), and PC2 is dominated by the parameter of K⁺ (0.82). The pH value has negative loading in PC2 of the igneous-fissure water. In the karst hyporheic zone, the PC1 shows an agreement among K⁺ (0.98), Na⁺ (0.92), H₂SiO₃ (0.79), while Ca²⁺ (–0.76), HCO₃⁻ (–0.72), and pH (–0.65) have negative loadings. Moreover, the PC2 is mainly controlled by SO₄²⁻ (0.97) and Mg²⁺ (0.94), with a small amount affected by Ca²⁺ (0.57) and HCO₃⁻ (0.59).

The karst water, including the groundwater in peak cluster valley, peak cluster depression, karst basin, is controlled by three factors. In PC1 of karst water, positive correlation exists among SO₄²⁻ (0.94), Na⁺ (0.94), K⁺ (0.89), H₂SiO₃ (0.85), and Cl⁻ (0.81). The relevant parameters of PC2 are Ca²⁺ (0.96) and HCO₃⁻ (0.88), with a reasonable negative correlation of pH (–0.75). The PC3 is categorized by a strong loading of Mg²⁺ and minor loading of HCO₃⁻.

4.2. Results of Isotopic Values of ²H and ¹⁸O

Hydrogen and oxygen isotopic values of collected samples are listed in Table S2. The δ²H and δ¹⁸O values of the igneous-fissure water samples collected during the wet season varied from −46.48‰ to −35.99‰ and from −7.64‰ to −6.19‰, with an average of −39.93‰ and −6.77‰, respectively. During the dry season in the same area, the δ²H and δ¹⁸O values of groundwater were marginally positive, ranging from −42.16‰ to −30.34‰ and from −6.95‰ to −5.22‰, with average values of −37.39‰ and −6.41‰, respectively. The δ²H and δ¹⁸O values of the karst water from the peak cluster valley, peak cluster depression, and karst basin varied from −47.00‰ to −35.10‰ and from −7.52‰ to −6.04‰, respectively. The δ²H and δ¹⁸O values of the groundwater from the karst hyporheic zone ranged from −42.20‰ to −38.00‰ and from −7.05‰ to −6.42‰, respectively.

5. Discussion

5.1. Hydrogeochemical Properties of the Igneous-Fissure Water

As shown in Table 1, the dominant cations and anions in the igneous-fissure water sampled in the eroded tectonic mountains are Na⁺ > Ca²⁺ > K⁺ > Mg²⁺ and HCO₃⁻ > Cl⁻ > SO₄²⁻. The hydrochemistry in distinct lithologic distribution areas varies greatly, as shown in the Piper diagram. According to Figure 4, the igneous-fissure water (red triangular symbols) is dominated by Na–Ca–HCO₃- and Na–HCO₃-type water.

The chemical weathering of rocks is a significant factor in the formation of dissolved substances in groundwater [43]. As shown in Figure 5a, half of the igneous-fissure water plots are within the rock-weathering control area and half within the precipitation control area, suggesting that the chemical components of igneous-fissure water (except borehole Z01) in the eroded tectonic mountains are influenced by both rock weathering and precipitation. This aligns with the topography and climate conditions in the area. Furthermore, to some extent, the rough topography, abundant rainfall, rapid subsurface runoff, and dense vegetation cover in the eroded tectonic mountains inhibited evaporation.

Mixing diagrams of Na-normalized molar ratios are used to further trace the solute sources and hydrogeochemical processes [25,49]. In Figure 5b, the igneous-fissure water sampled in the eroded tectonic mountains is plotted in the center of the diagram, showing that the primary driver of hydrogeochemical composition is silicate weathering. The primary minerals discovered in the core Z01 were K-feldspar (KAlSi₃O₈), plagioclase (NaAlSi₃O₈−CaAl₂Si₂O₈), and quartz (Si₂O), followed by biotite K(Mg, Fe)₃(AlSi₃O₁₀)(F, OH)₂) and amphibole (Ca, Na)₂(Mg, Fe, Al)₅(Al, Si)₈O₂₂ (OH)₂) (Figure S1), all of which potentially offer basic ingredients for metasilicic acid, Na⁺, and K⁺ in the igneous-fissure water [15]. The observation of typical plagioclase-series feldspar that occurs in the granites of the study area is consistent with the published sources [31,32]. Plagioclase is a solid solution series composed of albite and anorthite in varying proportions [20]. Therefore, the possible mineral dissolution reactions contributing to the chemical compositions of the igneous-fissure water are as follows:

$$\begin{gathered} 9{\mathrm{NaAlSi}}_3{\mathrm{O}}_8 \left(\mathrm{Albite}\right)+8{\mathrm{CO}}_2+23{\mathrm{H}}_2\mathrm{O}={\mathrm{NaAl}}_7{\mathrm{Si}}_{11}{\mathrm{O}}_{30}(\mathrm{OH}{)}_6 (\mathrm{Na}\text{-}\mathrm{Montmorillonite})+ \\ {\mathrm{Al}}_2{\mathrm{Si}}_2{\mathrm{O}}_5(\mathrm{OH}{)}_4 (\mathrm{Kaolinite})+8{\mathrm{Na}}^{+}+8{{\mathrm{HCO}}_3}^{-}+14{\mathrm{H}}_2{\mathrm{SiO}}_3 \end{gathered}$$

(8)

CaAl₂Si₂O₈ (Anorthite) + 2CO₂ + 6H₂O = Al₂O₃·3H₂O + Ca²⁺ + 2HCO₃⁻ + 2H₂SiO₃

(9)

2KAlSi₃O₈ (K-feldspar) + 2CO₂ + 7H₂O = Al₂Si₂O₅(OH)₄ (Kaolinite) + 2K⁺ + 2HCO₃⁻ + 4H₂SiO₃

(10)

SiO₂ (Quartz) + H₂O = H₂SiO₃

(11)

This study found that the metasilicic acid content of most igneous-fissure water met the “Drinking natural mineral water (GB8537-2018)” [50] requirement of ≥25 mg/L. The content of metasilicic acid ranges from 17.90 to 36.80 mg/L, with an average of 26.09 mg/L, indicating a prospective exploration region for metasilicate mineral water. Neither ion exchange nor biomass absorption affect metasilicic acid concentrations in groundwater [51]. Accordingly, it can be used as one of the indicators to determine the weathering degree of silicate minerals [52]. Moreover, the Na⁺ and Cl⁻ components do not indicate a (molar) ratio of Cl⁻/Na⁺ of 1 (Table S3) and their SI values of halite range from −10.42 to −8.42 (Figure 6 and Figure 7, Table S4), confirming that these components were produced from the weathering of silicate minerals rather than halite [53].

In the PCA, the most relevant parameters of PC1 in the igneous-fissure water are HCO₃⁻, Na⁺, Ca²⁺, and H₂SiO₃, indicating that dissolutions of albite and anorthite are the most dominant hydrogeochemical processes in the igneous-fissure water. The loadings of SO₄²⁻, Mg²⁺ in the PC1 are as significant as the HCO₃⁻, Na⁺, Ca²⁺, and H₂SiO₃, indicating an originate of mineral dissolution for them. The SI values of anhydrite and gypsum for igneous-fissure water are extremely low (−6.73~−3.26), suggesting that the SO₄²⁻ are not derived from the dissolution of anhydrite and gypsum. As we all know, SO₄²⁻ can be derived from aqueous pyrite oxidation [55]. The SO₄²⁻ content of the shallow igneous-fissure water are extremely low. However, small amounts of pyrite and galena were found in the Z01 core (Figure S1), which account for the largest SO₄²⁻ content in the deep igneous-fissure water (6.80 mg/L). Similarly, the chlorite found in the Z01 core (Figure S1) contributes to the existence of Mg²⁺ [56]. Therefore, the chemicals in the deep igneous-fissure water were not primarily derived from the precipitation (Figure 5a), and the weathering dissolution of minerals, such as aragonite, calcite, or dolomite, but rather from silicate minerals, and the oxidation of pyrite and chlorite. PC2 is dominated by the parameter of K⁺ (0.82), with a negative loading of pH value, which represents an anthropogenic source of the chemicals in the igneous-fissure water. However, the average content of K⁺ is extremely low (1.68 mg/L), indicating little human influence on the igneous-fissure water.

Cation exchange plays a significant role in the origin of chemical components in igneous-fissure water. Figure S3l shows that the igneous-fissure water has a chloralkali imbalance, with CAI values ranging from −18.43 to −1.49. The Mg²⁺ and/or Ca²⁺ in the fissure water exchanges ions with Na⁺ and/or K⁺ in the aquifer materials [46,57].

5.2. Hydrogeochemical Properties of the Karst Water

In Figure S3, the contents of SO₄²⁻ and Cl⁻ are relatively high in the karst water, and the shallow karst water from the karst basin has the highest concentrations of SO₄²⁻ and Cl⁻. In general, sources of SO₄²⁻ in groundwater include precipitation [58], mineral dissolution (e.g., anhydrite and gypsum) [53] and oxidation [55], and agricultural fertilizer application [7]. In the same precipitation-covered area, the SO₄²⁻ content in the northern igneous-fissure water is very low, indicating that SO₄²⁻ is not derived from precipitation. The SI values of anhydrite and gypsum for karst water are all unsaturated (−3.9~−1.61, Table S4). Other sulfur-containing minerals (e.g., pyrite) were not found in the cores (Z03 and Z08, Figure S1) from the karst areas, indicating that the SO₄²⁻ did not originate from the weathering of sulfur-containing minerals. Although the deep igneous-fissure water is affected by the aqueous oxidation of pyrite, its SO₄²⁻ content (6.81 mg/L) is much lower than that of most karst water (Table 1). Moreover, human activities (agricultural activities) occur frequently in karst areas. Therefore, the primary source of SO₄²⁻ in the karst water is agricultural activity rather than the dissolution of sulfur-containing minerals and precipitation. The SI values of halite for karst water are significantly less than 0 (−11.12~−8.21, Table S4). Chemicals in groundwater generated from halite dissolution should have equal concentrations of Na⁺ and Cl⁻ [53]. The Cl⁻/Na⁺ ratio of the karst water in the underground river entrance (A49, 1.32) is a little bit smaller than that in the outlet (A50, 1.48), which may be influenced by the ion exchange process along the groundwater flow path. Nevertheless, the Cl⁻/Na⁺ ratios of most samples in the karst areas (including the karst basin) are, remarkably, greater than 1 (Table S3), indicating that Cl⁻ in these samples may originate from significant evaporation or agricultural activities to varying degrees. In the karst water, SO₄²⁻, Na⁺, K⁺, and Cl⁻ all belong to PC1 (Table 2), which also indirectly provides proof for this analysis.

The contents of TDS, HCO₃⁻, Ca²⁺, Na⁺, and K⁺ are relatively high in the karst water, especially in the karst basin. Generally, the karst basin represents the final stage of karst landform evolution [59]. The karst basin is a confluence area. According to the mixing diagrams using Na-normalized molar ratios (Figure 5b), the chemical composition of the deep karst water (borehole Z08) from the karst basin is influenced by carbonate weathering, while the shallow karst water (such as A52) is affected by the weathering of carbonate and silicate minerals. Moreover, the metasilicic acid content in the shallow karst water of the karst basin is comparatively high, with a maximum content of 32.90 mg/L. The metasilicic acid content in deep karst water (sampled taken from borehole Z08) is 14.90 mg/L, demonstrating that shallow karst water is more susceptible to silicate weathering than deep karst water. Human activities affect the chemical composition of shallow karst water in the karst basin, resulting in Ca− HCO₃−SO₄-type water (Figure 4). However, the deep karst water (borehole Z08) from the karst basin was Ca−Mg−HCO₃-type water, which is predominantly affected by mineral weathering of its dolomitic limestone [11] and less affected by human activities.

The relevant parameters of PC2 are primarily Ca²⁺ (0.96) and HCO₃⁻ (0.88), which may represent the hydrogeochemical process of calcite dissolution. The chemical types of groundwater in the area of the peak cluster valley, peak cluster depression, and parts of the karst basin and karst hyporheic zone are primarily of the Ca–HCO₃ type (Figure 4). According to the Gibbs diagram (Figure 5a), groundwater samples in the peak cluster valley, peak cluster depression, and karst basin plot in the rock-weathering control area indicate significant water–rock interactions in these regions. However, the hydrogeochemical properties of groundwater in the peak cluster valley, peak cluster depression, and karst basin are somewhat distinct in the present study.

Geology, landform, and soil must all be considered when classifying hydrochemical components in groundwater [2,60]. The karst water in the peak cluster valley and peak cluster depression is controlled by the weathering of carbonate minerals, as shown by the Gibbs and mixing diagrams of Na-normalized molar ratios (Figure 5b). The SI values of calcite in the karst water of the peak valley and peak depression are similar (Figure 6 and Figure 7). However, the SI values for dolomite in the karst water from the peak cluster depression are greater than those from the peak cluster valley (Figure 6). The karst water from the peak cluster depression contains high levels of Mg²⁺, HCO₃⁻, and TDS, while the peak cluster valley has high levels of Ca²⁺, HCO₃⁻, and TDS. The Mg²⁺ level in the karst water from the western peak cluster depression is slightly greater than that from the peak cluster valley (Figure S3e), which may be attributed to their lithological differences. Dolomite (CaMg(CO₃)₂) is a mineral with higher Mg²⁺ than calcite or aragonite. The dissolved products of dolomite are Mg²⁺, Ca²⁺, and HCO₃⁻, resulting in an increase in Mg²⁺ content in groundwater [24]. The main mineral dissolution reactions in the karst water are as follows:

CaCO₃ (Calcite) + CO₂ + H₂O = Ca²⁺ + 2HCO₃⁻

(12)

CaMg(CO₃)₂ (Dolomite) + 2CO₂ + 2H₂O = Ca²⁺ + Mg²⁺ + 4HCO₃⁻

(13)

Some karst water from the western karst depressions is Ca−Mg−HCO₃-type water, which also supports this analysis. The related parameters of PC3 in the karst water are Mg²⁺ and HCO₃⁻, which may represent the dissolution of dolomite. Owing to the low SiO₂ content and limited presence of soluble free quartz in carbonate minerals, the H₂SiO₃ concentrations in the karst water from the peak cluster valley and peak cluster depression are low [20,61].

5.3. Hydrogeochemical Properties of Groundwater in the Karst Hyporheic Zone

Karst is relatively well developed in the karst hyporheic zone as a result of fracture and fragmentation, and underground rivers are developed as well. The outlet of an underground river is typically found along the edge of a karst valley and at the front of a platform. As a result, the hydrogeochemical properties of groundwater in these places differ from those found elsewhere. In Figure 5, some springs (A10 and A20) plot near the end member of the silicate rock, while others (such as A21 and A24) plot near the end member of the carbonates. The groundwater in these places is a mixture of karst-fissure and other waters. The chemical parameter contents, SI, and CAI values of most samples from the karst hyporheic zone are intermediate between those of groundwater in the granite and karst areas (Table 1, Figure 6, Figure 7 and Figure S3). These findings indicate that the hydrogeochemical properties of groundwater in the karst hyporheic zone are influenced by the weathering of carbonates and silicates [43].

According to the PCA, the most relevant parameters of PC1 are K⁺, Na⁺, and H₂SiO₃, which represent the products originated from the dissolution of albite and K-feldspar, respectively [15]. The parameters of Ca²⁺, HCO₃⁻ and pH have negative loadings in PC1, indicating that the chemicals derived from the dissolution of albite and K-feldspar are negatively correlated with that of the dissolution of Ca-containing minerals. Some groundwater samples had slightly elevated Ca²⁺ and metasilicic acid levels. Owing to the involvement of calcite and other minerals in the hydrogeochemical reactions, the Ca²⁺ activity in the aqueous solution is constrained by the dissolution equilibrium of calcite, making it difficult for the activity product of [Ca²⁺]/[H⁺]² in the solution to reach the saturation equilibrium level of anorthite [20]. Therefore, the dissolution of anorthite or Ca-containing minerals is promoted, resulting in an increase in the concentrations of Ca²⁺ and metasilicic acid in groundwater in the karst hyporheic zone. Moreover, the PC2 is influenced by SO₄²⁻ and Mg²⁺, with a small amount of Ca2⁺ and HCO₃⁻, suggesting that the groundwater chemicals in some areas are possibly derived from the aqueous oxidation of pyrite [55] and the dissolution of dolomite [24].

Spring A20 plots close to borehole Z01 in the granite area on the Gibbs diagram (Figure 5) and shows a similar relationship between SI and TDS (Figure 7), indicating that the two hydrogeochemical processes are also similar. In particular, spring A10 is located in the karst area, with an extremely low TDS content together with high levels of metasilicic acid and Na⁺. It is Ca−Na−HCO₃-type water. It deviates from the karst water samples and is closer to the igneous-fissure water samples shown in Figure 4, Figure 5 and Figure 7, indicating a similar hydrogeochemical process to that of igneous-fissure water. Therefore, the scale should be extended to the location of spring A10.

5.4. Isotopic Values of ²H and ¹⁸O

Using precipitation data from Baojinggong (BJG) Karst Cave in Yingde City from 2011 to 2014 [62], the Qingyuan hilly area in 2022 [63], and the samples analyzed in 2023 (Table S2), the meteoric water line for the Beijiang River Basin, namely the local meteoric water line (LMWL), was fitted as δD = 8.60δ¹⁸O + 14.74 (R² = 0.98). The δ²H and δ¹⁸O values of karst- and igneous-fissure water are distributed near to the global meteoric water line (GMWL; δD = 8 δ¹⁸O + 10 [64]; Figure 8), suggesting that the sources of groundwater are similar, and precipitation is the dominant recharge source. In this study, the samples plot to the left side of the LMWL due to moisture recycling [65]. Precipitation first accesses the saturated zone or aquifer through the soil and unsaturated zone and then enters the saturated zone or aquifer following partial evaporation [18].

The values of δ²H and δ¹⁸O not only help to explore the sources of groundwater recharge but also effectively track the water–rock interactions in groundwater based on changes in oxygen isotopic values [66,67]. The whole rock δ¹⁸O values of carbonate and silicate rocks are reported to be generally high [68,69]. In general, the δ¹⁸O values of igneous minerals increase gradually due to the continuous fractionation and crystallization of magma [70]. Mafic minerals with ¹⁸O-depletion crystallize early, while silica minerals (e.g., quartz) crystallize later and can have δ¹⁸O values of up to 12.9‰ [71]. Isotope exchange during the water–rock interaction leads to the enrichment of ¹⁸O isotopes in groundwater, which is called “oxygen positive drift” [72]. In addition to evaporation and isotope exchange, deep groundwater recharge also contributes to the ¹⁸O-enrichment. In the karst hyporheic zone, ¹⁸O-depletion in the groundwater from underground river (A21) and surface water (A22) after rainfall (Figure 9) correlate with the recharge of precipitation. However, slight enrichment of ¹⁸O isotope observed in the spring (A20) is possibly related to the recharge of deep groundwater with ¹⁸O-inrichment. The multiple effects make the groundwater line slope in the Qinglian River Basin less than LMWL.

For the wet season samples, the fitting equation for the igneous-fissure water line is δD = 7.82 δ¹⁸O + 13.07 (R² = 0.97, Figure 8), while for the dry season samples, it is δD = 6.73 δ¹⁸O + 5.74 (R² = 0.94). The slopes of the igneous-fissure water in wet and dry seasons are 7.82 and 6.73, respectively, indicating that the evaporation or isotope exchange in the igneous-fissure water is stronger in the dry season than in the wet season, which can be attributed to the higher precipitation and rapid groundwater runoff during the wet season.

The fitting equation for the karst water line during the dry season is δD = 8.00 δ¹⁸O + 13.06 (R² = 0.98). The fitting equation for the groundwater from the karst hyporheic zone during the dry season is δD = 6.61 δ¹⁸O + 5.77 (R² = 0.99). The slope of the fitting equation for groundwater from the karst hyporheic zone is the smallest among all groundwater sample groups, indicating the strongest mixing effects of evaporation, isotope exchange in water–rock interaction and deep groundwater recharge in the karst hyporheic zone.

The sources of the main chemicals in groundwater have been accurately identified by the analyzed drilling data, hydrochemical parameters, and δ²H and δ¹⁸O values. However, the contribution ratio of factors affecting water quality cannot be quantified, and further assessment is recommended using models that can identify sources and quantify contributions based on the physical and chemical properties of substances. Currently, receptor models that can be used to calculate the contribution ratio of various factors include positive matrix factorization (PMF) [73], absolute principal component score–multiple linear regression (APCS/MLR) [74], potential source contribution function (PSCF) [75], and PCA−APCS−MLR [76].

On the other hand, there are some of the difficulties remain when using this sampling approach to characterize groundwater geochemistry. For example, deep groundwater analysis requires drilling to collect water samples, resulting in the high cost of collecting deep groundwater. Additionally, the pH values in the study cannot be tested in the field, and it is recommended that pH values should be tested in situ in future similar studies.

6. Conclusions

This study explored the hydrogeochemical properties of groundwater in different landforms by analyzing drilling data, hydrochemical parameters, and δ²H and δ¹⁸O values of groundwater. The major points of the study are summarized as follows:

• It is the first specific analysis of the groundwater geochemistry in the karst-fissure aquifer system of the Qinglian River Basin in northern Guangdong Province. Further quantified assessment on the contribution ratio of factors affecting water quality is recommended.
• Weathering of silicate minerals, oxidation of pyrite and chlorite, cation exchange reactions, and precipitation are the primary sources of dissolved chemicals in the igneous-fissure water from the eroded tectonic mountains.
• The natural chemical substances in the karst water are mostly derived from the weathering of carbonate minerals (calcite and dolomite) and cation exchange reactions. Moreover, the PCA suggests that the most relevant parameters in the karst water are possibly from anthropogenic activities, which are also closely related to groundwater quality in karst areas. Furthermore, the chemical composition of the deep karst water from the karst basin is mainly influenced by the weathering of carbonate and cation exchange reactions and is less affected by human activities.
• The hydrogeochemical properties of groundwater in the karst hyporheic zone are influenced by the weathering of silicates and carbonates, and the promotion effect of dissolution of anorthite or Ca-containing minerals. In addition, the mixing effects of evaporation, isotope exchange in water–rock interaction, and deep groundwater recharge in the karst hyporheic zone are the strongest in the Qinglian River Basin, as indicated by δ²H and δ¹⁸O analyses. The hydrogeochemical characteristics of the karst interaction zone are so unique that further studies are recommended.

The analysis of drilling data, hydrochemical parameters, and δ²H and δ¹⁸O values were conducted in a significantly different lithology area (Qinglian River Basin) and show hydrogeochemical processes. The exploration gives a good overview of the groundwater quality controlling factors within the region. Understanding the hydrogeochemical processes in the Qinglian River Basin is critical for scientific management and decision-making regarding karst and fissure groundwater.