Identification of Magmatic Fluid Inputs and Geochemical Evidence of the Mantle-Derived Components in Magma-Heated Geothermal Systems

Abstract

Magma-heated geothermal systems have garnered significant attention in academia due to their unique formation mechanisms and vast potential. This paper focuses on the Rehai, Ruidian, and Banglazhang geothermal fields in the Tengchong area. We present the element geochemistry and isotope compositions of hot springs, cold springs, and surface water to explore magmatic fluid input into geothermal systems and investigate the release of deep mantle-derived components. Based on our findings, we propose a conceptual model and theoretical framework for geothermal system genesis constrained by magmatic heat source influences. Results indicate that magma-heated geothermal systems coexist with three types of geothermal water: neutral chloride-rich water, acidic sulfate-rich water, and alkaline bicarbonate-rich water. The infusion of magmatic fluids into geothermal systems. The enrichment of trace elements in hot springs is jointly controlled by magmatic fluid input and host rock leaching. The magma chamber is the primary factor influencing the reservoir temperature. The parent geothermal fluid can be identified within the geothermal system. During circulation, the parent geothermal fluid undergoes three cooling processes: adiabatic cooling, conductive cooling, and mixing with cold water. We propose that the release of mantle-derived materials is a key factor in element enrichment within magma-heated geothermal systems, and mantle-derived components are more enriched in areas with active magma chambers. The findings of this study provide insights into magmatic fluid input into geothermal systems and highlight the critical role of the release of mantle-derived components in the formation of high-temperature geothermal resources.

1. Introduction

As a vital component of clean energy systems, geothermal energy holds an irreplaceable position in the global energy revolution and carbon reduction strategies [1,2,3]. Among the various types of geothermal resources, high-temperature geothermal resources are particularly noteworthy due to their immense development potential and unique research value [4,5,6]. In China, high-temperature geothermal resources are primarily distributed in the Yunnan–Tibet geothermal belt, encompassing regions such as southern Tibet, western Sichuan, and western Yunnan [7]. The high-temperature geothermal activities in these areas are closely linked to deep magmatic processes triggered by the collision between the Indian Plate and the Eurasian Plate [8,9,10].

Magma-heated geothermal systems are high-temperature geothermal systems that are influenced by both heat conducted from shallow magma chambers and the direct replenishment of magmatic fluids [11,12,13,14]. Due to their exceptionally high reservoir temperatures and distinctive geochemical characteristics, magma-heated geothermal systems have attracted significant attention from the academic community [15,16,17,18,19,20]. These geothermal systems are predominantly located at plate boundaries and in mantle plume activity regions [21,22], such as the Pacific Ring of Fire geothermal belt [23], the Mediterranean–Himalayan geothermal belt [24,25,26], and the Mid-Atlantic Ridge geothermal belt [27,28]. The presence of magmatic heat sources not only maintains the high-temperature state of the system but also significantly alters the hydrochemical and isotope composition of the geothermal fluids through the input of magmatic fluids [5,29,30,31,32,33,34,35,36,37,38].

Geothermal systems are near-surface manifestations of heat transfer from the Earth’s core to the surface, and fluids play a pivotal role in the transfer of heat and mass [39,40]. Recent studies have demonstrated that fluid geochemical data possess unique advantages in deciphering the genesis of geothermal systems. Such data not only accurately trace fluid sources but also reflect key geochemical processes occurring within reservoirs [10,34,41,42,43,44,45,46,47,48,49,50,51,52]. Geochemical investigations have revealed that magmatic heat-source geothermal systems exhibit the following distinctive characteristics: (1) enrichment in mantle-derived helium and volatile gases from magmatic degassing, indicating the presence of magmatic heat sources [53,54,55]; (2) the occurrence of acidic springs (pH < 7), closely associated with magmatic volatiles, particularly H₂S [56,57,58]; (3) high concentrations of major elements, such as Cl and trace elements, including B, As, Li, Rb, and Cs, in the geothermal water, reflecting contributions from magmatic fluids [5,10,19,24,59]; and (4) significant positive δ¹⁸O shifts, indicative of intense water–rock interactions [60,61]. These features collectively constitute critical hydrogeochemical markers for identifying magmatic heat-source geothermal systems. However, previous research on magmatic heat-source geothermal systems has primarily focused on genesis models and mechanisms of element enrichment. The identification and quantification of magmatic fluid input, as well as the mechanisms of the release of mantle-derived components within geothermal systems, require further investigation.

Similar to most magmatic heat-source geothermal systems worldwide, the formation of the Tengchong geothermal system is closely related to magmatic activity. Geophysical surveys indicate the presence of a shallow magma chamber beneath the Tengchong region [62,63,64], providing a stable and continuous heat source for the geothermal system. This unique magma-hydrothermal interaction mechanism makes the Tengchong geothermal area an ideal natural laboratory for studying the formation and evolution of magmatic heat-source geothermal systems. Most previous studies have focused on the heat sources, reservoir temperatures, and genesis of hot springs in the Tengchong region, while research on the quantitative analysis of hot spring material sources and the impact of mantle-derived material on geothermal water has been relatively limited. In this study, three groups of high-temperature geothermal water from the Tengchong region were selected as research subjects: the Rehai (RH), Ruidian (RD), and Banglazhang (BLZ) geothermal water. By analyzing the fluid geochemical data, the input of deep magmatic fluids was determined, and the cooling processes experienced by the parent geothermal fluid beneath the geothermal fields were investigated. The proportions of magmatic water mixed into the geothermal water were calculated using deuterium and oxygen isotopes. Helium isotope analysis (³He/⁴He (R/Ra) data are from [65,66]) was employed to assess the impact of the release of mantle-derived material on the composition of the geothermal water.

The major objectives of this study were (1) to employ geochemical methods to identify the presence of magmatic heat sources within the studied geothermal systems; (2) to quantitatively analyze the contributions of magmatic fluids to the geothermal systems and identify the chemical components derived from the magmatic sources; (3) to investigate the impact of the magmatic heat sources on the deep reservoir temperatures, identify parent geothermal fluids, and analyze their circulation processes; and (4) to explore the processes of the release of mantle-derived components in magma-heated geothermal systems.

2. Geologic Setting

The Tengchong Block is situated at the southeastern escape margin of the India–Eurasia suture zone (Figure 1a), in the western segment of the Sanjiang Tethys Orogenic Belt. It is part of the southeastern deformation zone of the Qinghai–Tibet Plateau (Figure 1c) and has been shaped by the subduction, escape, rotation, and strike-slip movements following the closure of the Nujiang Ocean (170–100 Ma) and the Myitkyina Ocean (150–65 Ma) [67]. The internal tectonic deformation of the Tengchong Block is highly distinct: the eastern boundary is defined by the nearly N-S-trending Gaoligong strike-slip fault zone connecting it to the Baoshan–Menglian Block, and the western boundary is delineated by the N-S-trending large strike-slip fault zone. Numerous strike-slip faults are developed between these fault zones, and they predominantly trend NE and N-S [68]. The formation of the volcanic rocks, the seismic activity, and the distribution of the geothermal areas in Tengchong are closely associated with the extensive fault development in the region.

The Tengchong Block experienced frequent magmatic activity during the Mesozoic and Cenozoic, including two major magmatic events during the Yanshanian and Himalayan periods. The stratigraphic development within the region is incomplete, with some strata missing [69]. The lithology of the Tengchong Block includes high-grade metamorphic rocks, Devonian–Permian sedimentary rocks, and Mesozoic–Cenozoic volcanic rocks (Figure 1d). The Proterozoic Gaoligongshan Group metamorphic basement is considered to be the oldest stratum in the Tengchong Block and is primarily composed of gneiss, schist, marble, slate, and amphibolite [70,71].

Tengchong is part of the Himalayan geothermal region (Figure 1b), characterized by abundant hot springs associated with volcanic activity, large-scale hydrothermal activity, and intense hydrothermal alteration. Sinter deposits are commonly present near hot spring outcrops. The hydrothermal activity within the region manifests in diverse forms, including typical hot springs, high-temperature boiling springs, natural gas vents, steam escaping from the ground, and hydrothermal explosions. The Rehai geothermal field, as the largest and most active hydrothermal area in Tengchong [24], features hot springs and boiling springs as its primary manifestations, along with numerous gas vents, heat-dissipating grounds, and extensive sinter deposits. The geothermal reservoir in the region consists of metamorphic rocks of the Gaoligongshan Group and the Yanshanian granites, while the caprock is composed of Tertiary Nanlin Formation sandstone and conglomerates. N-NE-trending faults serve as critical conduits for the upward migration of geothermal fluids. Studies have confirmed the presence of an uncooled magma chamber beneath the Rehai geothermal field, which serves as the heat source for this high-temperature geothermal system [7,62,72]. Based on MT survey results, a high-conductivity body approximately 7 km deep and 20 km thick has been inferred beneath the Rehai area, and it is to be a cooling magma chamber [73]. Based on carbon isotope fractionation calculations, ref. [74] estimated that the current temperature of the Tengchong–Heshun–Rehai magma chamber is 438–773 °C, with an average of 566 °C.

Figure 1. (a) Geotectonic map of the major Cenozoic fracture zones in Asia modified from [75]. (b) Map showing the location of the Himalayan geothermal region. (c) Tectonic map of the Sanjiang Tethys orogenic belt modified from [76]. (d) Geologic map of Tengchong and location of the study area modified from [77]. (e) Location of the Rehai samples modified from [5]. (f) Location of the Ruidian samples modified from [78]. (g) Location of the Banglazhang samples modified from [5].

Figure 1. (a) Geotectonic map of the major Cenozoic fracture zones in Asia modified from [75]. (b) Map showing the location of the Himalayan geothermal region. (c) Tectonic map of the Sanjiang Tethys orogenic belt modified from [76]. (d) Geologic map of Tengchong and location of the study area modified from [77]. (e) Location of the Rehai samples modified from [5]. (f) Location of the Ruidian samples modified from [78]. (g) Location of the Banglazhang samples modified from [5].

3. Materials and Methods

3.1. Sampling and Analysis

The sampling was conducted between April and May 2023, during which 14 water samples were collected, including 10 geothermal water samples, three surface water samples, and one cold spring water sample. To obtain more comprehensive hydrochemical data, 22 additional water sample datasets were obtained from published studies [24,57,78,79]. The sampling locations are shown in Figure 1e–g. When sampling, all water samples were filtered through 0.45 μm membranes on-site. In the field, a multiparameter water quality (HI991301, HANNA, Woonsocket, RI, USA) was used to measure the temperature, pH, electrical conductivity (EC), redox potential (Eh), and total dissolved solids (TDSs) of each water sample. The water samples were collected in 500-mL high-density polyethylene (HDPE) bottles, which were rinsed three times prior to sampling. For the samples intended for cation and trace element analysis, 2 mL of high-purity nitric acid was added to acidify the samples to a pH of <2. After sampling, the bottles were immediately sealed with parafilm.

The water sample analyses were conducted by the Groundwater and Environmental Monitoring Center of the Institute of Hydrogeology and Environmental Geology, Chinese Academy of Geological Sciences, Ministry of Natural Resources. The water chemistry and hydrogen-oxygen isotope analyses were conducted according to the Methods for Groundwater Quality Analysis. The major cations and trace elements were analyzed using an inductively coupled plasma optical emission spectrometer (Avio 550 Max, Waltham, MA, USA), and the anions were analyzed using an ion chromatograph (Metrohm 930, Herisau, Switzerland). The anion and cation balance error control was within 3%. The hydrogen and oxygen isotopes were analyzed using a water isotope analyzer (L2030i, Santa Clara, CA, USA) employing wavelength scanning-cavity ring-down spectroscopy (WS-CRDS, Santa Clara, CA, USA). The D/H and ¹⁸O/¹⁶O isotope ratios were calibrated against the Vienna Standard Mean Ocean Water (V-SMOW) international standard, and the analytical precisions for δD and δ¹⁸O were 1‰ and 0.1‰, respectively.

3.2. Analysis Methods

The magma-heat-source type geothermal system is a geothermal system strongly influenced by underground magma [24]. The magmatic fluids released from the magma chamber beneath it are a more significant factor in the geochemistry of geothermal water than the rock-water reactions within the heat reservoir. Magmatic fluids are high-temperature gases and superheated solutions that separate from magma as it rises and cools underground; they transport heat and dissolved elements. To determine whether a geothermal system is influenced by magmatic fluids, it is necessary to assess whether magmatic water has been mixed into the geothermal water; this requires first identifying the parent geothermal fluid.

3.2.1. Proportion of Magmatic Water Mixing

The proportion of the magmatic water mixed in the geothermal fluids was calculated based on the hydrogen and oxygen isotope compositions of the geothermal water, magmatic water, and atmospheric precipitation. The specific calculation formula is as follows [60]:

$$X_a={\displaystyle \frac{\left({\delta }_d-{\delta }_m\right)}{\left({\delta }_a-{\delta }_m\right)}}$$

(1)

where X_a is the mixing ratio of magmatic water; and δ_d, δ_a, and δ_m are the δD or δ¹⁸O composition (‰ V-SMOW) of the geothermal water, magmatic water, and atmospheric precipitation, respectively.

3.2.2. Identification Methods for Parent Geothermal Fluids

The Cl-enthalpy model and Cl-deuterium model have been widely used to identify the presence of parent geothermal fluids and their associated processes [80,81]. These models have also been applied to estimate the temperatures of parent geothermal fluids and analyze their cooling and mixing mechanisms during ascent [19,24,82,83].

To identify parent geothermal fluids using the Cl-enthalpy model, it is first necessary to estimate the reservoir temperature. The Na-K geothermometer effectively reflects deep reservoir temperatures [24,84,85,86]. Hence, it was selected for use in this study. On the Cl-enthalpy diagram, cooling and mixing lines were plotted to interpret the cooling and mixing mechanisms of the geothermal fluids. The adiabatic cooling line was drawn by connecting the saturation vapor point (enthalpy value of 2779.4 J/g; chloride ion concentration of 0 mg/L) to the sample points, while the mixing line connected the cold water points to the sample points. The intersection of these two lines represents the parent geothermal fluid.

The Cl-deuterium model can be employed to study the mixing and cooling processes affecting the composition of the geothermal fluid [81]. Using this model, the isotope compositions of the parent geothermal fluids must first be calculated. For a single water–vapor separation process at a given temperature, the isotope compositions of the parent geothermal fluids can be determined [19,81]. The proportion of the vapor phase (f_v) after water–vapor separation can be calculated using the following formula:

$$f_v={\displaystyle \frac{\left(H_0-H_l\right)}{(H_v-H_l)}}$$

(2)

where H₀ is the initial enthalpy value of the parent geothermal fluid; and H_l and H_v are the enthalpy values of the liquid and vapor phases after water–vapor separation at a specific temperature.

Assuming rapid isotopic fractionation between the water and vapor phases during ascent after water–vapor separation, the fluid in the reservoir remains in isotopic equilibrium. Thus, the isotopic mass balance equation is as follows:

$${\delta }_v{f}_v+{\delta }_l\left(1-f_v\right)={\delta }_0$$

(3)

where δ_v and δ_l are the isotope δ values of the vapor and liquid phases, respectively; and δ₀ is the isotope δ value of the parent geothermal fluid. The isotopic difference between the vapor and liquid phases can be approximated using the isotopic fractionation coefficient α at a specific temperature [87]:

$${\delta }_l-{\delta }_v={10}^{3}ln{\alpha }_{l-v}$$

(4)

$${\delta }_0={\delta }_l-f_v\ast {10}^{3}ln{\alpha }_{l-v}$$

(5)

The isotopic fractionation coefficients for hydrogen and oxygen isotopes between 0 and 374.1 °C can be calculated using the following formulas [88]:

$$\begin{gathered} {10}^{3}ln{\alpha }_{l-v}\left(\mathsf{\delta }\mathrm{D}\right)=1158.8\left({\displaystyle \frac{T^3}{{10}^9}}\right)-1620.1\left({\displaystyle \frac{T^2}{{10}^6}}\right)+794.84\left({\displaystyle \frac{T}{{10}^3}}\right) \\ -161.04+2.9992\left({\displaystyle \frac{{10}^9}{T^3}}\right) \end{gathered}$$

(6)

$$\begin{gathered} {10}^{3}ln{\alpha }_{l-v}\left({\mathsf{\delta }}^{18}\mathrm{O}\right)=-7.685+6.7123\left({\displaystyle \frac{{10}^3}{T}}\right)-1.6664\left({\displaystyle \frac{{10}^6}{T^2}}\right) \\ +0.35041\left({\displaystyle \frac{{10}^9}{T^3}}\right) \end{gathered}$$

(7)

where T is the temperature at which isotopic fractionation occurs (K).

4. Results

4.1. Hydrochemical Characteristics

The physical properties, stable hydrogen–oxygen isotopes, and major and trace element analysis results for the water samples are presented in

Table 1. Physical properties and stable isotope compositions of water samples from the Tengchong region.

Sites	No.	Type	T (°C)	pH	TDS (mg/L)	δD (‰)	δ18O (‰)	Hydrochemical Type
Rehai	RH01	Hot spring	92	8.75	2330	−63.7	−8.3	Na-Cl·HCO3
	RH02	Hot spring	96	8.74	2228	−62.8	−8.8	Na-Cl·HCO3
	RH03	Hot spring	92.5	7.9	2104	−62.6	−8.1	Na-Cl·HCO3
	RH04	Hot spring	50	4.85	110	−62	−8.5	Ca·Na-SO4·HCO3
	RH05	Hot spring	96	7.24	2502	−65	−7.1	Na-Cl·HCO3
	RH06	Hot spring	72	1.81	2420	−55	−8.1	Ca-SO4
	RH07	Hot spring	90	8.08	1474	−69	−9	Na-HCO3·Cl
	RH08	Hot spring	86	8.3	1516	−68.8	−9.3	Na-Cl·HCO3
	RH09 a	Hot spring	97.5	9.33	1677	-	-	Na-Cl
	RH10 a	Hot spring	89.5	9.15	1778	-	-	Na-Cl
	RH11 a	Hot spring	79	8.43	1153	-	-	Na-Cl·HCO3
	RH12 b	Hot spring	91.8	4.5	461	−64	−7.2	Na-SO4·Cl
	RH13 b	Hot spring	70	8.32	1460	−65	−8.4	Na-HCO3·Cl
	RH14 c	Hot spring	91	3.1	617	−58.6	−5.06	Na-SO4
Ruidian	RD01	Hot spring	83	7.43	1278	−83	−10.7	Na-HCO3
	RD02 d	Hot spring	90	7.6	1314	-	-	Na-HCO3
	RD03 d	Hot spring	86	7	1307	-	-	Na-HCO3
	RD04 d	Hot spring	75	7.47	1055	-	-	Na-HCO3
	RD05 d	Hot spring	85	6.91	1403	-	-	Na-HCO3
	RD06 c	Hot spring	77	7.2	1113	−74.5	−11.24	Na-HCO3
	RD07 c	Hot spring	87	8.1	1270	−82.4	−11.39	Na-HCO3
	RD08 c	Hot spring	86.5	7.7	1273	−85.9	−10.76	Na-HCO3
	RD09 c	Hot spring	56	7.9	783	−71.1	−10.24	Na-HCO3
	RD10 c	Hot spring	64	8.12	1143	−78.4	−10.73	Na-HCO3
	RD11 c	Hot spring	55	7.5	1223	−81.4	−8.54	Na-HCO3
Banglazhang	BLZ01	Hot spring	90.5	7.76	767	−72	−9.2	Na-HCO3
	BLZ02 a	Hot spring	75	7.97	517	−76	−10.7	Na-HCO3
	BLZ03 a	Hot spring	87	8.45	513	−77	−10.6	Na-HCO3
	BLZ04 a	Hot spring	88	9.13	527	−78	−11.4	Na-HCO3·SO4
	BLZ05 a	Hot spring	63	8.25	490	−75	−10.4	Na-HCO3·SO4
	BLZ06 a	Hot spring	53	7.14	458	−80	−11.8	Na-HCO3·SO4
	BLZ07 a	Hot spring	91	8.36	530	−78	−11.4	Na-HCO3
Entire Tengchong	Q01	Cold spring	19.2	6.68	168	−64	−9	Mg·Ca-HCO3
	W01	Surface water	20.1	8.67	140	−66	−9.1	Ca-HCO3·SO4
	W02	Surface water	25.3	8.45	99	−62	−8.8	Ca·Na-HCO3
	W03	Surface water	24.5	7.63	204	−61	−8.4	Na-HCO3·Cl

Note: ^a = data from [78]; ^b = data from [57]; ^c = data from [89]; and ^d = data from [24]. - = not analyzed.

and

Table 2. Hydrogeochemistry of water samples from the Tengchong region (in mg/L).

No.	Na	K	Ca	Mg	HCO3	CO3	Cl	SO4	F	As	B	Sr	Li	Rb	Cs	SiO2
RH01	745.0	103.5	0.50	n.d.	509.5	261	612.7	20.85	15.68	0.338	10.81	0.084	7.988	1.653	0.818	218.2
RH02	716.6	99.5	0.55	n.d.	519.3	246	591.7	18.55	15.4	0.768	10.42	0.073	7.156	1.543	0.778	196.8
RH03	681.0	93.36	0.49	n.d.	677.3	138	569.0	17.48	15.48	0.775	9.99	0.059	7.268	1.528	0.781	155.8
RH04	6.4	3.45	9.54	2.18	25.02	0	1.8	25.09	0.19	0.003	n.d.	0.071	0.006	0.012	0.002	48.3
RH05	832.4	115.8	0.56	n.d.	1199	0	705.5	11.65	18.17	0.965	12.07	0.095	9.182	1.843	0.932	304.6
RH06	18.2	26.00	77.42	8.52	0	0	61.3	1855	2.32	0.074	n.d.	0.108	1.388	0.376	0.039	229.5
RH07	460.5	67.13	1.83	0.14	567.5	72.01	348.4	22.58	8.12	0.259	6.09	0.056	4.526	0.943	0.469	158.1
RH08	472.3	67.02	1.26	0.05	456.4	127.2	360.6	27.68	9	0.277	6.25	0.025	4.318	0.937	0.481	184.5
RH09 a	511.4	66	2.9	0	87.3	21.7	621.6	23.8	15.9	0.841	-	-	8.425	1.405	0.724	174.9
RH10 a	573.3	83.8	3.2	0	147.1	25.1	642.3	0	15.9	1.042	-	-	9.381	1.631	0.757	181.7
RH11 a	471.9	68.1	5.7	0.1	285.8	8.9	368	31.1	9.1	0.402	-	-	5.607	0.945	0.435	199.3
RH12 b	58.5	21.7	2.16	0.61	12.8	0	64.3	162	1.38	0.070	1.03	0.01	0.520	0.250	0.050	140.0
RH13 b	375	68.9	4.88	0.54	708	39.1	362	34	8.97	0.310	6.12	0.06	4.310	0.970	0.500	199.0
RH14 c	58	23	9.43	3.21	0	0	16	303	1.2	0.040	0.75	-	0.600	0.200	n.d.	195.0
RD01	428.3	40.87	3.81	3.48	701.7	102	160.0	25.86	8.02	0.880	3.82	0.460	2.131	0.370	0.553	130.6
RD02 d	326	30.8	15.5	2.03	928	16.9	130.9	27.3	-	0.742	5.44	0.49	1.509	0.276	0.433	251.8
RD03 d	290.1	29.2	20.4	1.96	906.2	4.1	117.6	25.2	-	0.682	4.89	0.52	1.400	0.258	0.407	222.6
RD04 d	250.2	26.6	17.7	1.74	752.7	7.1	105	25.5	-	0.811	6.12	0.45	1.242	0.236	0.340	197.1
RD05 d	320.2	33.9	21.6	2.12	902.2	3.4	119.1	27.1	-	0.696	5.10	0.52	1.438	0.284	0.422	254.0
RD06 c	350	34	22.2	3.26	775	0	139	32	7	0.640	3.25	-	1.800	0.300	0.440	123.0
RD07 c	410	45	4.45	4.2	895	36.3	140	31.5	7.04	0.640	2.90	-	2.200	0.470	0.710	129.0
RD08 c	400	42	23.6	3.57	912	0	157	16	8.5	0.800	4.00	-	2.100	0.330	0.440	148.0
RD09 c	245	28.5	7.45	3.25	520	24.6	76.9	14.6	4.04	0.090	1.70	-	1.290	0.250	0.560	104.0
RD10 c	378	36.6	7.52	2.52	768	24.6	133	25	6.85	0.560	3.78	-	1.850	0.360	0.720	126.0
RD11 c	400	36	23.6	3.66	837	0	164	24	8.5	0.800	3.25	-	2.100	0.330	0.520	140.0
BLZ01	245.2	22.39	1.94	0.08	427.1	24	19.6	39.86	20.27	0.081	3.60	0.021	2.138	0.385	0.755	146.0
BLZ02 a	160.7	15.7	7.6	0	260.1	2.4	20	47.2	22.2	0.184	-	-	2.874	0.361	0.729	223.5
BLZ03 a	148.5	15.5	9.6	0.2	163.2	4.5	19.4	46.6	20.7	0.175	-	-	2.598	0.334	0.656	231.2
BLZ04 a	164.6	14.8	4.5	0	47.2	6.3	19.4	45.2	20.4	0.191	-	-	2.915	0.364	0.720	270.9
BLZ05 a	146.5	15.4	10.1	0.2	206.7	3.5	19	68.5	18.7	0.139	-	-	2.478	0.316	0.614	207.2
BLZ06 a	140.5	12	15	0.4	220.5	0.3	18.2	96.5	20.3	0.109	-	-	2.509	0.290	0.570	214.7
BLZ07 a	155.9	14.1	6.5	0	196.4	4.4	24.1	34.9	22.4	0.169	-	-	2.783	0.341	0.688	232.9
Q01	10.8	4.31	12.00	9.79	103.7	0	5.3	6.88	0.35	n.d.	n.d.	0.097	0.007	0.017	0.000	56.8
W01	9.7	2.14	25.10	5.15	73.22	0	4.6	37.53	0.95	0.005	n.d.	0.047	0.007	0.009	0.001	11.4
W02	8.7	1.66	14.59	3.23	76.27	0	1.8	4.91	0.35	0.001	n.d.	0.041	0.011	0.006	0.003	21.9
W03	43.0	7.55	14.47	2.99	122.6	0	29.8	7.18	1.24	0.013	0.36	0.091	0.234	0.060	0.025	39.0

Note: ^a = data from [78]; ^b = data from [57]; ^c = data from [89]; and ^d = data from [24]. The hydrogen and oxygen stable isotope data for RH01, RH02, RH03, and RH08 are from [79], and those for RH06 data are from [57]. - = not analyzed; and n.d.= not detected (Mg < 0.013 mg/L, As < 0.001 mg/L, B < 0.05 mg/L).

. Based on the sampling locations and pH, the geothermal water samples were categorized into four groups: (I) acidic geothermal water from Rehai (RH04, RH06, RH12, and RH14); (II) neutral to alkaline geothermal water from Rehai (RH01-RH03, RH05, RH07-RH11, and RH13); (III) alkaline geothermal water from Ruidian (RD01-RD11); and (IV) alkaline geothermal water from Banglazhang (BLZ01-BLZ07). The pH range of the acidic geothermal water from RH is 1.81–4.85, and the TDS values range from 110 to 2420 mg/L (average = 902 mg/L). The neutral to alkaline geothermal water sample from Rehai has a pH range of 7.24–9.33 and TDS values ranging from 1153 to 2502 mg/L (average = 1822.2 mg/L), and they are characterized by higher Cl⁻ and TDS levels overall. The TDS values of the neutral to alkaline geothermal water from RD and BLZ are 783–1403 mg/L and 458–767 mg/L, respectively, which are lower than those of the neutral to alkaline geothermal water from RH. The cold spring and surface water samples have a pH range of 6.68–8.67 and TDS values ranging from 99 to 204 mg/L.

4.2. Stable Isotope Composition

The δD and δ¹⁸O values of the acidic geothermal water from RH range from −64‰ to −55‰ and from −8.5‰ to −5.06‰, respectively. The δD and δ¹⁸O values of the neutral-alkaline geothermal water from RH range from −69‰ to −62.6‰ and from −9.3‰ to −7.1‰, respectively. The δD and δ¹⁸O values of the geothermal water from RD range from −85.9‰ to −71.1‰ and from −11.39‰ to −8.54‰, respectively. The δD and δ¹⁸O values of the geothermal water from BLZ range from −80‰ to −72‰ and from −11.8‰ to −9.2‰, respectively. The δD and δ¹⁸O values of the non-geothermal water range from −66‰ to −61‰ and from −9.1‰ to −8.4‰, respectively. The δD and δ¹⁸O values of the geothermal water from RH dominate among all of the types of geothermal water.

5. Discussion

5.1. Identification of Magmatic Heat Sources

Hydrochemical analysis is an effective method for identifying the heat sources of geothermal systems. Magma-heated geothermal systems typically include three types of hot springs: (I) neutral chloride-rich thermal water, (II) acidic sulfate-rich thermal water, and (III) alkaline bicarbonate-rich thermal water [24,34,90]. Type I thermal water is considered to be derived from parent geothermal fluids that have undergone adiabatic processes. Type II water originates from acidic geothermal steam separated from the parent geothermal fluid, which ascends to shallow cold water layers. Type III water is formed during the migration of type I thermal water toward the periphery of the hydrothermal zone, where the temperature and pressure are lower and dissolved CO₂ undergoes further hydrolysis, enhancing water–rock interactions. Different types of hot springs experience distinct hydrogeochemical processes. According to the Piper diagram (Figure 2), the neutral-alkaline geothermal water from RH is chloride-rich water, including Na-Cl·HCO₃, Na-Cl, and Na-HCO₃·Cl types. The acidic geothermal water from RH is sulfate-rich water, including Na-SO₄·Cl, Na-SO₄, Ca-SO₄, and Ca·Na-SO₄·HCO₃ types. The geothermal water from RD and BLZ is bicarbonate-rich water, primarily Na-HCO₃ type, as well as some Na-HCO₃·SO₄ type water in the BLZ geothermal field. The non-geothermal water types include Mg·Ca-HCO₃, Ca-HCO₃·SO₄, Ca·Na-HCO₃, and Na-HCO₃·Cl. Notably, surface water sample W03 is Na-HCO₃·Cl type water, and it has higher Na⁺ and Cl⁻ concentrations than the other surface water samples, indicating mixing of Na⁺ and Cl⁻-rich thermal water in W03. In the three typical high-temperature geothermal fields selected in the study area, all three types of hot springs were observed, suggesting the presence of a magmatic heat source deep underground.

5.2. Identification of Magmatic Fluid Input in Geothermal Systems

5.2.1. Recharge Source Identified by δD and δ¹⁸O

Stable hydrogen and oxygen isotopes have commonly been used as effective tracers for identifying groundwater sources, water–rock interactions, and the groundwater distribution. By determining the primary recharge sources of deep geothermal fluids, the circulation processes of the geothermal fluids can be explored [19,91,92,93]. On the δD-δ¹⁸O diagram (Figure 3), the hydrogen and oxygen isotope distributions of the water samples from the study area plot close to the global meteoric water line (GMWL: δD = 8δ¹⁸O + 10) [94] and the local meteoric water line (LMWL: δD = 7.8δ¹⁸O + 14.9) [89], indicating that precipitation infiltration is the primary recharge source of the geothermal water in the study area.

In geothermal systems with homogeneous rock composition and water-to-rock ratios, oxygen isotope shifts in deep hydrothermal fluids are primarily driven by vaporization and mixing with cold water during ascent [95]. Consequently, these shift values serve as crucial indicators for tracing hydrothermal evolution [89]. Notably, the RH samples exhibit a significant rightward δ¹⁸O shift. A linear fit of the RH geothermal water yields a slope of 1.68, characteristic of a magmatic fluid mixing line, suggesting the occurrence of deep magmatic fluid recharge [60,96]. The δD and δ¹⁸O values of the geothermal water samples from RD and BLZ are relatively low, and the degrees of the δ¹⁸O shift are less pronounced than those in the RH geothermal water, indicating a smaller contribution from magmatic fluids enriched in δ¹⁸O. The rightward δ¹⁸O shift is primarily associated with water–rock interactions under high-temperature conditions.

Analysis confirms the presence of magmatic fluid recharge in the RH geothermal water. Using Equation (1), the proportion of magmatic water mixing was calculated, and the stable hydrogen and oxygen isotope compositions of the magmatic water were determined to be δD = −20‰ and δ¹⁸O = 10‰ [60]. A local precipitation sample was collected from the Tengchong area, and its stable isotope composition was determined to be δD = −102‰ and δ¹⁸O = −14.3‰. Our calculations indicate that magmatic water contributes 21–38% to the RH geothermal water. Except for RH14 (with a mixing ratio of 38%), the contributions of the magmatic water to the other hot springs range from 21% to 30%, which are comparable to the values for high-temperature geothermal fields in southern Tibet that are influenced by magmatic fluid input (mixing ratios of 14–31%) [20].

5.2.2. Identification of Magma-Sourced Components

Based on the concentrations of the major ions in the water samples, a Schoeller diagram was plotted (Figure 4). For the RH neutral to alkaline geothermal waters, the predominant cation is Na⁺, ranging from 375 to 832.4 mg/L, while Cl⁻ has the anion with the highest concentration, ranging from 348.4 to 705.5 mg/L. The RH acidic geothermal water is enriched in Ca²⁺ and Mg²⁺, which is attributed to the significant mixing of shallow cold Ca²⁺-rich and Mg²⁺-rich water. The dominant anion is SO₄²⁻, and sample RH06 has the highest SO₄²⁻ concentration, reaching 1855 mg/L. Overall, the SO₄²⁻ concentrations of the acidic geothermal water are higher than those of the other types of geothermal water. The geothermal water samples from RD and BLZ exhibit similar ionic compositions, with Na⁺ as the dominant cation and HCO₃⁻ as the dominant anion. Figure 4 illustrates that the magmatic source-related Na⁺ and Cl⁻ [97] are most enriched in the RH neutral-alkaline geothermal water, while the SO₄²⁻, associated with acidic H₂S geothermal steam heating [57], is most enriched in the RH acidic geothermal water. HCO₃⁻, linked to silicate weathering and dissolution, is most enriched in the RD geothermal water.

Diagnostic trace elements such as B, Li, Rb, and Cs are considered to be the typical geothermal suite, and the concentrations of these elements are primarily influenced by magmatic degassing, dissolution of host rock minerals, and mixing with cold groundwater or surface water [20,98,99]. The relationships among the B, Li, Rb, and Cs concentrations are shown in Figure 5. The RH neutral-alkaline geothermal water exhibits high concentrations of B (6.09–12.07 mg/L), Li (4.31–9.38 mg/L), Rb (0.937–1.843 mg/L), and Cs (0.435–0.932 mg/L). The elevated trace element concentrations of thermal water samples are not only related to release from host rocks but also to magmatic volatiles [19]. Due to the relatively shallow depth of the magmatic heat source in the RH geothermal field, the influence of magmatic fluid input cannot be neglected. The high B, Li, Rb, and Cs concentrations are jointly contributed by magmatic fluids and leaching of the host rocks.

The Cl-SO₄-HCO₃ ternary diagram is primarily used to interpret the major geochemical processes of geothermal fluids [86]. Geothermal water is classified into deep Cl water (near the Cl vertex), peripheral circulation water (near the HCO₃ vertex), and steam-heated water (near the SO₄ vertex). As shown in Figure 6a, the acidic geothermal water from RH is a typical steam-heated type, characterized by very low pH values and high SO₄²⁻ concentrations, which results from the dissolution of acidic gases such as H₂S and SO₂ released from magma. The neutral to alkaline geothermal water from RH mainly plots in the deep Cl water area and its vicinity, indicating that it is influenced by recharge from Cl-rich magmatic fluids. In contrast to the first two types, the geothermal water samples from RD and BLZ are classified as peripheral circulation water (steam-condensed water) due to their high HCO₃⁻ contents, indicating that the elements in these water samples primarily originate from leaching of the host rocks. The Cl-Li-B ternary diagram illustrates the relative Cl, Li, and B concentrations, with varying degrees of conservative behavior (Figure 6b), showing the distinct distribution patterns among the different types of geothermal water. The geothermal water from RH and RD deviates from the rock region due to the absorption of B or Cl geothermal steam. The geothermal water from RH trends toward the absorption of low B/Cl steam, suggesting that RH represents a relatively older geothermal system, whereas the geothermal water from RD corresponds to a relatively younger geothermal system. The geothermal water from BLZ and some of the acidic geothermal water samples from RH, due to limited Cl supplementation, plot near the rock region. The Cl-Li-B ternary diagram indicates that the Cl, B, and Li concentrations in the geothermal water samples are influenced by magmatic activity within the geothermal area.

5.3. Constraints of Magmatic Heat Source on Reservoir Temperature

The processes of dilution, mixing, and re-equilibration of geothermal water can affect the accuracy of geothermal thermometers [101]. The Na-K-Mg ternary diagram, first introduced by [86], is a commonly used and efficient method for determining geothermal equilibrium states and the extent of water–rock interactions. As shown in Figure 7, the neutral-alkaline geothermal water from RH primarily plots within the fully equilibrated and partially equilibrated or mixed water zones, indicating a high degree of water–rock interaction. The hot springs exposed along the N-S-trending faults (RH01-RH03, RH05, RH09, and RH10) exhibit higher water–rock equilibrium levels than those exposed along the E-W-trending Zaotang River faults (RH07, RH08, RH11, and RH13), suggesting that the fault zones not only serve as water conduits but also play a crucial role in controlling the water–rock equilibrium. The acidic geothermal water from RH and the geothermal water from RD plot entirely within the immature water zone, which is associated with the mixing of shallow cold Ca-rich and Mg-rich waters, shallow circulation depths, and fast flow rates. Immature water may yield significantly overestimated or underestimated temperatures when using cation geothermometers [20].

Significant differences exist in the reservoir temperature estimates based on various geothermometers (Table 3 and Figure 8a). Due to the influence of cold water mixing and SiO₂ re-equilibration, the temperatures calculated using the K-Mg, Li-Mg, and SiO2 geothermometers tend to be underestimated, often reflecting the temperature of the geothermal water near the spring outlets [102]. Na+ is considered to reflect the equilibrium between alkaline feldspar and deep fluids. However, geothermal water rarely achieves re-equilibration during ascent, making the Na ion geothermometer relatively accurate for calculating the reservoir temperatures of high-temperature geothermal systems [86,103]. Since the Na-Li geothermometer provides a wide temperature range, and the Na-K geothermometer yields more accurate temperature estimates, the Na-K geothermometer is deemed to be more reliable for predicting deep reservoir temperatures. Using the Na-K geothermometer, the deep reservoir temperatures in RH (excluding the acidic springs classified as immature water), RD, and BLZ were calculated to be 254–286 °C, 224–245 °C, and 220–236 °C, respectively. According to the analysis results presented in Figure 7, the samples from RD were identified as immature water, indicating that the water in RD is formed by the mixing of deep reservoir water and infiltrating cold water. The temperature of the shallow reservoir was calculated using the quartz geothermometer [24], yielding temperatures of 139–197 °C.

Geochemical thermodynamic models have been proven to be effective in predicting reservoir temperatures [104,105,106]. Using the PHREEQC geochemical program and the phreeqc.dat thermodynamic database, multi-component mineral equilibrium temperatures were inferred for the geothermal water samples from the study area. Sodium feldspar was selected as the appropriate aluminum mineral phase to enforce dissolution-precipitation equilibrium, and the temperature calculations were set within the range of 50–300 °C, which was divided into 60 steps. As shown in Figure 8b, for water sample RH03, K-feldspar, quartz, chalcedony, aragonite, and kaolinite converge well near logQ/K = 0 at a temperature of 204 °C. The simulation results indicate that the reservoir temperatures are within the ranges of 142–226 °C (average = 199 °C) for RH, 168–185 °C (average = 175 °C) for RD, and 147–163 °C (average = 157 °C) for BLZ.

The RH geothermal field is located above the central Tengchong–Heshun–Rehai magma chamber, approximately 7 km from the deep magma chamber [73], and has the highest reservoir temperatures. The BLZ and RD geothermal fields have lower reservoir temperatures as they are situated above the edge of the magma chamber and connected to the magma chamber via faults, respectively [65], which indicates that the magma chamber beneath the geothermal fields serves as the primary heat source of these magma-heated geothermal systems, and both the distance from the magma chamber and the temperature of the magma chamber are influencing the deep reservoir temperatures.

Table 3. Equilibrium temperatures estimated using the selected chemical geothermometers.

No.	Measured Temperature	Na-K	Na-K-Ca	K-Mg	Na-Li	Mg-Li	Quartz (No Loss)	Quartz (Max Loss)	Chalcedony	SI Method
		(1)	(2)	(3)	(4)	(5)	(6)	(7)	(8)	(9)
RH01	92	260	275	-	305	-	186	173	166	207
RH02	96	260	272	-	299	-	179	167	158	207
RH03	92.5	259	272	-	305	-	164	155	140	204
RH04	50	-	223	61	146	7	100	101	70	142
RH05	96	260	276	-	308	-	211	193	195	224
RH06	72	-	296	93	565	107	190	176	170	156
RH07	90	264	250	194	297	228	164	155	141	214
RH08	86	262	253	220	291	256	175	164	153	207
RH09 a	97.5	254	238	-	347	-	171	161	149	195
RH10 a	89.5	265	249	-	347	-	174	163	151	206
RH11 a	79	263	236	202	315	248	180	168	159	226
RH12 b	91.8	-	266	125	289	116	157	149	133	198
RH13 b	70	286	250	168	312	195	180	168	158	200
RH14 c	91	-	253	103	302	97	178	167	157	198
RD01	83	229	212	118	243	133	153	146	128	175
RD02 d	90	228	192	118	238	131	196	181	178	168
RD03 d	86	233	191	117	241	129	187	174	167	169
RD04 d	75	237	192	116	243	127	179	167	158	171
RD05 d	85	237	195	120	235	129	197	182	178	173
RD06 c	77	230	191	114	245	129	149	143	124	173
RD07 c	87	240	218	118	248	131	152	145	127	185
RD08 c	86.5	236	198	119	247	132	160	152	136	181
RD09 c	56	245	206	109	247	119	139	135	113	180
RD10 c	64	230	203	120	242	134	151	144	125	176
RD11 c	55	224	189	114	247	132	157	149	133	171
BLZ01	90.5	225	205	160	287	205	159	151	135	163
BLZ02 a	75	230	187	-	356	-	188	174	168	158
BLZ03 a	87	236	187	133	354	194	190	176	171	160
BLZ04 a	88	224	188	-	355	-	202	186	184	155
BLZ05 a	63	236	187	132	350	192	182	170	162	160
BLZ06 a	53	220	171	114	356	178	185	172	164	147
BLZ07 a	91	224	184	-	356	-	191	177	171	153

^a = data from [78]; ^b = data from [57]; ^c = data from [89]; and ^d = data from [24]. “-” denotes negative values. (1) T = 1390/[1.750 + lg(Na/K)] − 273.15 [86]; (2) T = 1647/[lg(Na/K) + (lg(Ca_0.5/Na) + 2.06)/3 + 2.47] − 273.15 [84]; (3) T = 4410/[13.95 − lg(K₂/Mg)] − 273.15 [86]; (4) T = 1590/[0.779 + lg(Na/Li)] − 273.15 [107]; (5) T = 2200/[5.47 + lg(Mg_0.5/Li)] − 273.15 [108]; (6) T = 1309/[5.19 − lg(SiO₂)] − 273.15 [80]; (7) T = 1522/[5.75 − lg(SiO₂)] − 273.15 [80]; (8) T = 1032/[4.69 − lg(SiO₂)] − 273.15 [80]; (9) SI method [104].

Table 3. Equilibrium temperatures estimated using the selected chemical geothermometers.

No.	Measured Temperature	Na-K	Na-K-Ca	K-Mg	Na-Li	Mg-Li	Quartz (No Loss)	Quartz (Max Loss)	Chalcedony	SI Method
		(1)	(2)	(3)	(4)	(5)	(6)	(7)	(8)	(9)
RH01	92	260	275	-	305	-	186	173	166	207
RH02	96	260	272	-	299	-	179	167	158	207
RH03	92.5	259	272	-	305	-	164	155	140	204
RH04	50	-	223	61	146	7	100	101	70	142
RH05	96	260	276	-	308	-	211	193	195	224
RH06	72	-	296	93	565	107	190	176	170	156
RH07	90	264	250	194	297	228	164	155	141	214
RH08	86	262	253	220	291	256	175	164	153	207
RH09 a	97.5	254	238	-	347	-	171	161	149	195
RH10 a	89.5	265	249	-	347	-	174	163	151	206
RH11 a	79	263	236	202	315	248	180	168	159	226
RH12 b	91.8	-	266	125	289	116	157	149	133	198
RH13 b	70	286	250	168	312	195	180	168	158	200
RH14 c	91	-	253	103	302	97	178	167	157	198
RD01	83	229	212	118	243	133	153	146	128	175
RD02 d	90	228	192	118	238	131	196	181	178	168
RD03 d	86	233	191	117	241	129	187	174	167	169
RD04 d	75	237	192	116	243	127	179	167	158	171
RD05 d	85	237	195	120	235	129	197	182	178	173
RD06 c	77	230	191	114	245	129	149	143	124	173
RD07 c	87	240	218	118	248	131	152	145	127	185
RD08 c	86.5	236	198	119	247	132	160	152	136	181
RD09 c	56	245	206	109	247	119	139	135	113	180
RD10 c	64	230	203	120	242	134	151	144	125	176
RD11 c	55	224	189	114	247	132	157	149	133	171
BLZ01	90.5	225	205	160	287	205	159	151	135	163
BLZ02 a	75	230	187	-	356	-	188	174	168	158
BLZ03 a	87	236	187	133	354	194	190	176	171	160
BLZ04 a	88	224	188	-	355	-	202	186	184	155
BLZ05 a	63	236	187	132	350	192	182	170	162	160
BLZ06 a	53	220	171	114	356	178	185	172	164	147
BLZ07 a	91	224	184	-	356	-	191	177	171	153

^a = data from [78]; ^b = data from [57]; ^c = data from [89]; and ^d = data from [24]. “-” denotes negative values. (1) T = 1390/[1.750 + lg(Na/K)] − 273.15 [86]; (2) T = 1647/[lg(Na/K) + (lg(Ca_0.5/Na) + 2.06)/3 + 2.47] − 273.15 [84]; (3) T = 4410/[13.95 − lg(K₂/Mg)] − 273.15 [86]; (4) T = 1590/[0.779 + lg(Na/Li)] − 273.15 [107]; (5) T = 2200/[5.47 + lg(Mg_0.5/Li)] − 273.15 [108]; (6) T = 1309/[5.19 − lg(SiO₂)] − 273.15 [80]; (7) T = 1522/[5.75 − lg(SiO₂)] − 273.15 [80]; (8) T = 1032/[4.69 − lg(SiO₂)] − 273.15 [80]; (9) SI method [104].

5.4. Identification of Parent Geothermal Fluids and Analysis of Cooling Processes

5.4.1. Chloride-Enthalpy Model

The analysis described in Section 5.2 revealed the presence of magmatic water recharge of the RH geothermal water, suggesting the likely existence of parent geothermal fluids beneath the RH geothermal field. Based on the chloride-enthalpy diagram (Figure 9), we successfully identified the presence of parent geothermal fluids, with an enthalpy value of 1575.8 J/g (337 °C, derived from the International Steam Tables [109]) and a Cl− concentration of 357 mg/L. Thus, the neutral-alkaline geothermal water in RH originates from parent geothermal fluids, while the acidic geothermal water is not derived from parent geothermal fluids and likely originates from shallow cold water that is unaffected by parent geothermal fluids.

Based on the various formation processes of hot springs, they can be categorized into three types. (1) Hot springs formed by direct adiabatic cooling of parent geothermal fluids, which then ascend to the surface through reservoir 1 (RH05). These hot springs exhibit the highest chloride ion concentrations among all three types, and their hydrochemical characteristics closely resemble those of deep geothermal fluids. (2) Hot springs formed by the mixing of parent geothermal fluids with infiltrating cold water, with a temperature of around 313 °C. The mixed geothermal fluids reach reservoir 2 (RH01, RH02, RH03, RH09, and RH10) via adiabatic cooling and subsequently ascend to the surface, and they undergo boiling during their ascent. (3) Hot springs formed by the direct mixing of parent geothermal fluids with infiltrating cold water within reservoir 3 (RH07, RH08, RH11, and RH13), which is located along the mixing line. These fluids ascend through adiabatic cooling at approximately 252 °C, and once the temperature drops to the boiling point, no further boiling occurs. They then ascend to the surface through conductive cooling. These hot springs have temperatures below the local boiling point. The hydrochemical characteristics of the hot springs are influenced by the fault activity in various directions. Reservoirs 1, 2, and 3 correspond to hot springs exposed along faults with different orientations, and while these reservoirs are relatively independent, they all originate from the same parent geothermal fluid. Notably, the acidic samples from the RH region plot above the mixing boundary. Some of the acidic hot springs approach the local boiling point but are characterized by low chloride ion concentrations and low TDS contents. These springs are formed by shallow underground cold water heated by high-temperature flash steam separated from deep geothermal fluids.

5.4.2. Chlorine-Deuterium Model

The application of the chlorine-deuterium model [81] constrains the existence of parent geothermal fluids and validates the results obtained from the chlorine-enthalpy model. Assuming that steam separation occurs in a single process at a given temperature, the hydrogen and oxygen isotope compositions of the parent geothermal fluids were calculated using Equations (2)–(7). The δD value and Cl⁻ content of the parent geothermal fluids were determined to be −80.3‰ and 357 mg/L, respectively. Figure 10 presents the chlorine-deuterium diagram, showing that the neutral-alkaline geothermal water from RH originates from the parent geothermal fluids, whereas the acidic geothermal water does not. The geothermal water samples in Groups 1, 2, and 3 in Figure 10 correspond to reservoirs 1, 2, and 3 in Figure 9. The geothermal fluid mixing and cooling processes indicated by both models are consistent. For Groups 1 and 2, the adiabatic cooling process maintains the temperature of the geothermal water, with boiling occurring and the water–rock interactions reaching complete equilibrium (Figure 7). The geothermal water samples in Group 3 primarily undergo conductive cooling, with faster temperature loss than in Groups 1 and 2, and they exhibit lower water–rock interaction levels, achieving partial equilibrium. The geothermal water samples in Group 4 are mainly influenced by cold water mixing, and the water–rock reactions do not reach equilibrium.

The geothermal water samples from the RD and BLZ plots below the cooling line have relatively low Cl concentrations, indicating minimal influence from the parent geothermal fluid. According to existing research, three magma chambers are present beneath the Tengchong volcanic area. The central magma chamber is situated only 7 km away from the RH geothermal reservoir and exhibits the highest activity [65]. RD is situated 10 km north of the northern magma chamber and is connected to it via fractures [110]. The BLZ geothermal field is located at the edge of the southern magma chamber [57]. Therefore, the geothermal water in RD and BLZ is less affected by the parent geothermal fluid. The RH geothermal field, which is located close to the magma chamber, is significantly influenced by the parent geothermal fluid, and the chemical characteristics spring water are the most similar to those of the deep parent geothermal fluid.

5.5. Release of Deep Mantle-Derived Components

In Section 5.4, the parent geothermal fluid was successfully identified. However, whether the chemical composition of the geothermal water is influenced solely by the parent fluid or also by the release of mantle-derived components requires further discussion. Since ³He primarily originates from the mantle and ⁴He predominantly originates from the crust, the helium isotope ratio ³He/⁴He is an effective tracer for identifying the origin of mantle-derived magmatic fluids, assessing their dilution by crustal materials, and monitoring short-term variations caused by volcanic activity [50,53].

By mapping the spatial distribution of the ³He/⁴He ratios of the water samples from Tengchong, the present-day activity in the volcanic region, and the intensity of the release of deep mantle-derived material were investigated. As shown in Figure 11, three mantle-derived helium enrichment zones were identified in the Tengchong, primarily located in the central and southern regions. These zones are characterized by recent intense volcanic activity and significant release of mantle-derived materials. The first zone has the highest ³He/⁴He ratios, with a maximum value of 5.92Ra and an average value of greater than 4.5Ra. The second zone has an average ratio of 3.75Ra, and the third one has an average of 2Ra. The other regions have an average ³He/⁴He ratio of 0.8Ra. It is evident that the first zone has the highest proportion of mantle-derived helium and the most intense recent volcanic activity. This zone also corresponds to the distribution of the RH hot springs, where the geothermal reservoir temperature is the highest in Tengchong. The RD geothermal field is located north of the second zone and is connected to the deep magma chambers via faults, while the BLZ geothermal field lies on the periphery of the third zone and has reservoir temperatures lower than those of the RH geothermal field.

The relatively high R/Ra values in Tengchong suggest that a greater proportion of mantle-derived volatile components ascend to the surface along faults that extend to great depths. Additionally, in the geothermal systems constrained by magmatic heat sources, the high concentrations of trace elements are linked to magmatic volatiles. However, it remains uncertain whether the element compositions of the geothermal water originate from the upper mantle or entirely from the magma chambers within the crust. By analyzing the relationship between the helium isotopes and the Cl, B, As, Li, Rb, and Cs concentrations of the geothermal water (Figure 12), we found that helium isotopes and trace element concentrations are positively correlated. Among these, the correlations between the ³He/⁴He ratio and the Cl, B, Li, and Rb concentrations are particularly strong, suggesting that these trace elements may originate from the release of mantle-derived components.

To further investigate the factors influencing the release intensity of mantle-derived components, we constructed a plot of ³He/⁴He versus ⁴He/²⁰Ne for the gas samples (Figure 13). Given the high correlations between the ³He/⁴He ratios and the B and Li concentrations (Figure 12), as well as the fact that the B and Li characteristics of the geothermal water can be used to effectively identify whether the residual magmatic fluid components enter the geothermal water through magmatic degassing [5,111], the bubble size in Figure 13 represents the Li and B concentrations. As shown in Figure 13, the samples in Group 1 not only have high ³He/⁴He ratios but also elevated concentrations of Li and B, exhibiting positive correlations between the ³He/⁴He ratio and the trace element concentrations. Thus, the release of mantle-derived components is a crucial factor in the enrichment of elements in the magmatic heat-source geothermal systems. The RH geothermal field is the region with the highest trace element enrichment in the study area, likely due to the strong activity of the magma chamber associated with RH, which releases the most mantle-derived volatile components. Along the path from springs 1 to 5 (Figure 13), it is evident that, as the distance from the magma chamber increases, the concentration of mantle-derived components in the emitted gases decreases, and the Li and B concentrations also decrease. The fractures extending beneath the springs serve as conduits for the release of mantle-derived substances, transporting gases and chemical components from the mantle upward. Concurrently, the deep convection of geothermal fluids, accompanied by intense thermal convection, provides heat sources for the formation of high-temperature geothermal systems. The uplift of the upper mantle beneath the Tengchong Volcano has created a tectonic pathway for magma supply from deep sources and for magma eruptions from magma chambers. The locally ascending asthenosphere is highly molten and relatively rich in volatile elements such as chlorine, boron, and lithium, providing a source of compositions for geothermal water.

6. Patterns of Magmatic Fluid Input to Hot Springs

Based on the findings of this study regarding the input of magmatic fluids and the release mechanisms of mantle-derived volatiles, a genetic model for magmatic heat-source geothermal systems was established using the high-temperature geothermal fields in Tengchong as an example (Figure 14).

The Tengchong volcano originates from the deep mantle above the mantle transition zone. As illustrated in Figure 14a, the subduction of the Indian Plate penetrates the discontinuity at a depth of 410 km and extends to the mantle transition zone at a depth of 660 km. The dehydration associated with plate subduction accelerates the upwelling of high-temperature materials, altering the mantle substances that ascend along tensile fractures and erupt at the surface to form volcanoes [112]. The upwelling of hot asthenosphere materials has led to magma eruptions and the formation of three crustal magma chambers at depths of 5–25 km [65], which provide localized high-temperature heat sources for the geothermal systems and continuously release substances such as Cl, B, and Li into the geothermal systems (Figure 14b). Influenced by the magmatic heat sources, several typical high-temperature geothermal fields have formed in the Tengchong area, including the Rehai geothermal field (Figure 14c), the Ruidian geothermal field (Figure 14d), and the Banglazhang geothermal field (Figure 14e). Although these high-temperature geothermal fields shared certain intrinsic connections during their formation, their geochemical characteristics exhibit some differences.

The primary recharge sources of the RH geothermal water are precipitation and magmatic fluids. Precipitation from the surrounding mountains (metamorphic rock zones of the Gaoligong Mountain range) infiltrates deeply through a series of N-S-, N-W-, and N-E-trending faults and circulates along the major faults. During deep circulation, the water primarily absorbs heat from the magma chambers within the crust. The magmatic fluid component in the geothermal water originates from parent geothermal fluids, which have temperatures of up to 337 °C. As the parent geothermal fluids ascend, they undergo adiabatic cooling due to the decreasing temperature and pressure of the surrounding environment and are further cooled by the infiltration of shallow cold water, forming geothermal reservoirs with temperatures ranging from 254 to 286 °C. The mid-alkaline geothermal water in RH, influenced by deeper magmatic fluids, exhibits high concentrations of ions such as Cl, Li, and B ions. Under high-temperature conditions, intense water–rock interactions occur between the geothermal water and surrounding rocks, leading to rapid increases in the concentrations of Na, K, SiO₂, and certain trace elements such as B, Li, Rb, and Cs. Ultimately forming Cl-rich geothermal water, which is dominated by Na-Cl·HCO₃ and Na-Cl type water. The acidic geothermal water in RH is replenished by significant amounts of shallow cold water, and its heat source primarily consists of deep geothermal steam rich in H₂S. The dissolved H₂S in the geothermal water provides SO₄, and the ionization of the H₂S in the water releases protons, rendering the geothermal water acidic. This ultimately results in sulfate-type geothermal water, such as the Ca-SO₄ type.

The RD geothermal water is primarily recharged by precipitation infiltrating from the mountainous regions on the eastern and western sides. Under the influence of gravity, the precipitation circulates along geothermal water channels composed of faults. During circulation, the weathered granite zone absorbs volatile components and high-temperature heat released from the magma chambers, forming geothermal reservoirs. Due to the greater distance of the RD geothermal activity area from the magma chambers, the heat supply to this area is lower than that to the RH and BLZ areas, resulting in deep reservoir temperatures ranging from 224 to 245 °C, which are lower than those of the RH geothermal field. Influenced by magmatic volatile components and water–rock interactions, the geothermal water becomes enriched in Li and B. Driven by density differences, the deep geothermal water migrates upward, bypasses the low-permeability layers, and recharges the shallow gravel aquifers, forming shallow geothermal reservoirs. These shallow reservoirs are further influenced by lateral cold water recharge and conductive cooling processes, ultimately forming Na-HCO₃ type geothermal water.

The geothermal water in BLZ is replenished by precipitation from the mountainous region on the eastern side. The infiltrating cold water undergoes dissolution and filtration as it circulates through the Yanshanian granite and the Cambrian metamorphic rocks. Due to the area’s history of multiple large-scale fractures, these fractures cut deeply, providing excellent conduits for the downward infiltration of shallow cold water and the upward migration of hot water. Cold water enters the aquifer and accumulates in large-scale water-conducting fracture zones, absorbing heat from the magma chambers to form thermal reservoirs with temperatures of 220–236 °C. During its circulation, the geothermal water undergoes intense water–rock interactions with the surrounding rocks, eventually ascending to the surface and forming hot springs, which are primarily characterized by Na-HCO₃ and Na-HCO₃·SO₄ type water.

Figure 14. (a) Conceptual model of the subduction of the Indian Plate modified from [113]. TVC: Tengchong volcano. (b) Schematic diagram of the current distribution of the magma chambers and their formation mechanisms in the Tengchong volcanic region [66]. (c) Conceptual model of the genesis of the geothermal system in the Rehai geothermal field, modified from [114]. (d) Conceptual model of the genesis of the geothermal system in the Ruidian geothermal field, modified from [110]. (e) Conceptual model diagram of the genesis of the geothermal system in the Banglazhang geothermal field, modified from [115]. (f) Theoretical framework for the genesis mechanism of the Tengchong geothermal system, constrained by magmatic heat sources.

Figure 14. (a) Conceptual model of the subduction of the Indian Plate modified from [113]. TVC: Tengchong volcano. (b) Schematic diagram of the current distribution of the magma chambers and their formation mechanisms in the Tengchong volcanic region [66]. (c) Conceptual model of the genesis of the geothermal system in the Rehai geothermal field, modified from [114]. (d) Conceptual model of the genesis of the geothermal system in the Ruidian geothermal field, modified from [110]. (e) Conceptual model diagram of the genesis of the geothermal system in the Banglazhang geothermal field, modified from [115]. (f) Theoretical framework for the genesis mechanism of the Tengchong geothermal system, constrained by magmatic heat sources.

In summary, the presence of magmatic heat sources constrains the formation of the high-temperature geothermal systems in Tengchong (Figure 14f). The magma chambers not only act as heat sources but also transport chemical substances into the geothermal systems. The subduction of the Indian Plate triggered magmatism, during which magmatic water replenishes the geothermal water, and mantle-derived volatile substances are released into the geothermal systems, contributing to the chemical composition of the geothermal water. Additionally, under the high-temperature conditions, the geothermal water undergoes intense water–rock interactions with the surrounding rocks, which is another source of the chemical composition of the geothermal water. The three key factors, namely, magmatic water replenishment, the release of mantle-derived component, and water–rock interactions, are crucial contributors to the composition of the geothermal water. The geothermal water originates from the mixing of precipitation and magmatic fluids, and it undergoes various cooling processes during its ascent before reaching the surface. The Tengchong area contains three types of geothermal water: neutral chloride-rich geothermal water, acidic sulfate-rich geothermal water, and alkaline bicarbonate-rich geothermal water. The coexistence of these three types of geothermal water further demonstrates that the geothermal system is constrained by magmatic heat sources.

7. Conclusions and Outlooks

7.1. Conclusions

In this study, we conducted a systematic hydrogeochemical analysis of geothermal fluids in a magma-heated geothermal system. Our focus was on examining the differences among various hot springs, identifying the input processes of magmatic fluids, and determining the presence of parent geothermal fluids. From the perspective of mantle sources, we analyzed the processes of element enrichment of the geothermal water. Based on the geochemical analysis results, we propose that the hydrochemical types of magma-heated geothermal systems can be classified into three categories: neutral chloride-rich thermal water, acidic sulfate-rich thermal water, and alkaline bicarbonate-rich thermal water. Neutral chloride-rich thermal water is primarily influenced by the mixing of Cl-rich magmatic water, acidic sulfate-rich thermal water is mainly formed via heating by deep geothermal steam, and alkaline bicarbonate-rich thermal water results from water–rock interactions combined with the mixing of shallow cold water. The primary recharge source for the hot springs is precipitation, accompanied by the input of magmatic fluids, and the input proportion ranges from 21% to 38%. The trace element enrichment (B, Li, Rb, and Cs) of the hot springs is jointly controlled by the input of magmatic fluids and leaching of the host rocks. The magma chamber beneath the geothermal fields serves as the main heat source for the geothermal systems, and both the distance of the heat reservoir from the magma chamber and the temperature of the magma chamber influence the temperatures of the heat reservoirs. A Cl-enthalpy model was used to identify the parent geothermal fluid (337 °C), which ascends to the surface through adiabatic, mixing, and conductive cooling processes, confirming the critical role of magmatic fluids in the deep circulation. The release of mantle-derived materials is a key factor in the element enrichment of the magma-heated geothermal systems, and the mantle-derived components are more enriched in regions with high magma chamber activity. In this study, we elucidated the pivotal role of magmatic fluids in deep geothermal circulation, quantitatively analyzed the mixing proportion of magmatic water in magmatic heat-source geothermal systems, and thoroughly investigated the sources of the components. Our findings have significant scientific implications for the development and utilization of magmatic heat-source geothermal resources.

7.2. Outlooks

This study investigates the genesis model of the Tengchong magmatic heat source geothermal system from the perspectives of hydrochemistry and isotopics and summarizes the genesis model of this type of geothermal system based on three aspects: magmatic water recharge, high-temperature water-rock interactions, and the release of mantle-derived components. However, research on the quantitative contributions of these various sources to the multiple elements in geothermal fluids remains insufficient, which represents a direction for future research.