Distribution Characteristics and Enrichment Mechanisms of Fluoride in Alluvial–Lacustrine Facies Clayey Sediments in the Land Subsidence Area of Cangzhou Plain, China

Abstract

Compression of clayey sediments not only causes land subsidence but also results in geogenic high fluoride groundwater. The distribution characteristics and enrichment mechanisms of fluoride in alluvial−lacustrine facies clayey sediments in the land subsidence area of Cangzhou Plain, China, were investigated using sample collection, mineralogical research, and hydrogeochemical and isotopic analysis. The results show that F⁻ concentration of groundwater samples ranged from 0.31 to 5.54 mg/L in aquifers. The total fluoride content of clayey sediments ranged from 440 to 792 mg/kg and porewater F⁻ concentration ranged from 0.77 to 4.18 mg/L. Clay minerals containing fine particles, such as muscovite, facilitate the enrichment of fluoride in clayey sediments, resulting in higher total fluoride levels than those in sandy sediments. The clay porewater F⁻ predominantly originated from the dissolution of water-soluble F and the desorption of exchangeable F from sediments. The F⁻ concentration in porewater was further influenced by ionic interactions such as cation exchange. The stable sedimentary environment and intense compression promoted the dissolution of F–bearing minerals and the desorption of adsorbed F in deep clayey sediments. The similar composition feature of δ²H−δ¹⁸O in deep groundwater and clay porewater samples suggests a significant mixing effect. These findings highlight the joint effects of hydrogeochemical and mineralogical processes on F behavior in clayey sediments.

1. Introduction

Groundwater is a vital resource that supplies drinking water for up to 70% of the global population [1]. Fluorine (F) is an essential trace element for human health, but chronic exposure to high F groundwater can result in waterborne fluorosis, including dental and skeletal fluorosis [2]. Geogenic high F groundwater has emerged as a global threat to water supply safety, according to reports from America, India, Bangladesh, Pakistan, and China [3,4,5,6]. High F groundwater affects around 70 million people across 29 provinces in China [7,8].

In addition to anthropogenic pollution, geogenic high F groundwater can be caused by burial−dissolution, evaporation−enrichment, and compression−release mechanisms. The burial−dissolution type high F groundwater mainly results from geological processes and water–rock interactions. Fluorite (CaF₂) dissolution is commonly acknowledged as the primary source of groundwater F [9]. Hydrogeochemical conditions have a considerable influence on groundwater F⁻ concentration. High F groundwater generally has low Ca²⁺ concentrations and a neutral to alkaline pH (7–9) [10,11]. Ionic interactions, such as cation exchange, can promote CaF₂ dissolution by decreasing Ca²⁺ concentration. Additionally, OH⁻ and HCO₃⁻ can displace F⁻ from the surfaces of amorphous Fe/Al hydroxides and clay minerals into groundwater due to similar anionic charge and radius [12,13]. The evaporation−enrichment type high F groundwater refers to the concentration and enrichment of F in groundwater during intensive evaporation in arid and semi-arid basins [14]. The compression−release type high F groundwater is formed by the mixture of high F porewater released from the clayey sediments during compression under overexploitation. The previous two categories have garnered substantial attention, the compression−release type high F groundwater is only becoming visible with increased overexploitation. For instance, Ye et al. indicated that the cumulative subsidence in the Cangzhou Plain (CZP) increased from 9 mm in 1975 to 2670 mm in 2013 due to overexploitation [15], with clay compression contributing up to 57.6% of the annual recharge to deep confined aquifers [16]. Meanwhile, researchers discovered the high F porewater in clayey sediments, based on the rise in regional groundwater F⁻ concentrations of the CZP [17,18]. Therefore, the release of high F porewater during clay compression may be a major cause of high F groundwater. However, the enrichment mechanisms of high F porewater in clayey sediments are frequently disregarded due to their low permeability and limited water supply, which hinder the understanding of the genesis of regional high F groundwater.

The North China Plain (NCP) has the most severe land subsidence in China, and the subsidence is currently progressing rapidly [19]. The decline in groundwater levels induced by overexploitation is not only the primary cause of land subsidence in the NCP, but it also deteriorates water quality. The CZP, located in the central eastern NCP, is a typical land subsidence zone with waterborne fluorosis. The presence of saline water in the eastern area of CZP because of seawater intrusion exacerbates groundwater overexploitation and the associated land subsidence in the western region [20]. The CZP was formed by alluvial−lacustrine deposition, resulting in large-scale clay layers and complicated sediment fluoride distribution under various sedimentary environments. Up to 57.6% of the groundwater in deep confined aquifers was replenished by the released porewater during the compression of clay layers, which was also the primary cause of land subsidence in the CZP [16]. The clayey sediments are mostly composed of illite, clinochlorite, fluorapatite, and biotite, which are predominant F-bearing clay minerals [21,22]. In addition to F-containing minerals, other sources of F exist in clayey sediments, such as Fe (III) complexed−F. Liu et al. discovered that the high levels of sediment F could transform into porewater through complex water−rock interactions in clay layers and then enter the adjacent aquifers during compression in the CZP, increasing the groundwater F⁻ concentration [23]. Identifying the enrichment mechanisms of high F porewater is an important basis for revealing the causes of compression−release type high F groundwater and effectively controlling waterborne fluorosis. However, existing research focuses mostly on the impact of clayey sediments on the groundwater quality in shallow aquifers [7,20,21]. The distribution characteristics and enrichment mechanisms of porewater F in clayey sediments under various sedimentary environments remain unknown.

In summary, taking the subsidence area in the western CZP as a typical research area, the representative sediment and groundwater samples from various depths were collected for mineralogical and hydrogeochemical studies by geological drilling and groundwater monitoring wells. The main objectives of this study are (1) to reveal the distribution characteristics of F in alluvial−lacustrine facies clayey sediments at varying depths in the land subsidence area, (2) to analyze the effect of the sedimentary environment on the occurrence of sediment F, and (3) to identify the formation mechanism of high F porewater in clayey sediments.

2. Materials and Methods

2.1. Study Area

The CZP is located in the eastern region of the NCP (115°6′–117°8′ N, 37°4′–38°9′ E; approximately 14,000 km²) and comprises Cangzhou City, Nanpi, Huanghua, Haixing, and Qingxian Counties, etc. (Figure 1b). This location has a warm temperate continental monsoon climate with distinct seasons and significant temperature differences. The multi-year average temperature is 12.2 °C, with a yearly average rainfall of 547.5 mm. The topography of CZP is modest and gently slopes from the southwest to the northeast. Cenozoic strata in CZP are dominated by alluvial and lacustrine deposits.

The Quaternary thickness in the CZP ranges from approximately 350 to 550 m. The Quaternary groundwater system is classified into four aquifers, from shallow to deep, based on lithology and hydrogeological conditions (Figure 1c,

Table 1. Geological information of Quaternary strata in CZP.

Aquifer	Lower Boundary Depth	Description
Aquifer I	−20~−60 m	Composed of fluvial–deltaic deposits that span the Holocene lower boundary and the Late Pleistocene upper boundary.
Aquifer II	−120~−180 m	Formed during the Late Pleistocene and is primarily composed of fluvial deposits interspersed with marine sediments.
Aquifer III	−250~−350 m	Formed during the Middle Pleistocene and is primarily composed of alluvial–lacustrine deposits.
Aquifer IV	−350~−550 m	Formed during the Early Pleistocene, consisting primarily of terrestrial sediments originated from river bends and floodplains.

).

The aquifers III and IV contain confined groundwater and are the primary distribution zones of high F groundwater [23]. Under natural conditions, groundwater flows from the southwest to the northeast along the topographic slope. However, over 40 years of overexploitation has caused considerable reductions in water levels in both shallow and deep aquifers, resulting in the formation of groundwater depression cones to varying degrees.

Land subsidence in the CZP began in the 1970s, initially localized in the central area of the urban depression cone, with a subsidence rate of only 9 mm/a. As urbanization accelerated and groundwater extraction increased, subsidence developed rapidly, beginning in 1980. The subsidence rate was highest in the 1990s, when numerous subsidence zones emerged. After 2005, the implementation of groundwater extraction limitations and the South-to-North Water Diversion Project significantly slowed the land subsidence rate. Previous research found that the accumulated subsidence in CZP increased from 9 mm in 1975 to 2670 mm in 2013 [15], with clay compression accounting for up to 57.6% of annual recharge to deep confined aquifers [16]. Based on the monitoring data through interferometric synthetic aperture radar (InSAR) and hydrostatic level, a contour map of accumulated subsidence from 1970 to 2023 was generated (Figure 1b). Although the accumulated subsidence in Cangzhou City is the largest, the annual average subsidence rate is relatively low due to limited exploitation at present, while the annual average subsidence rate in the western areas such as Xianxian County is relatively high. The overall subsidence rate steadily decreases from west to east. Therefore, this research mainly focuses on the western subsidence zone.

2.2. Sample Collection

A geological borehole (115°57′51.24″ E, 38°18′29.23″ N, surface elevation: 11.06 m) was set in the rapid subsidence zone in Xianxian County, southwestern CZP. The borehole depth is 400 m, and continuous core samples were collected from the entire depth. The underlying strata was primarily alluvial–lacustrine facies sedimentary, with alternating deposits of clay, silt, and sand. The total thicknesses of clay, silt, and sand layers were 214.22 m, 44.35 m, and 141.65 m, respectively. The borehole construction and samples collection were completed in October 2024.

Clayey sediment (total of 41) and sandy sediment samples (total of 10) were collected at various depths and immediately wrapped in preservative film and aluminum foil to reduce exposure to the atmosphere. The sediment samples were then placed in iron sheets to prevent extrusion. After transporting the samples to the laboratory, self-developed experimental equipment was utilized to collect porewater from the sediment samples by squeezing [24]. Due to human disturbance, it is hard to extract porewater from sandy sediment samples. A total of 20 groundwater samples around the borehole were collected from monitoring wells at depths ranging from 15 to 380 m in Xianxian in November 2024 (Figure 1). Samples were collected in HDPE containers using a low flow pump. The sampling bottles were rinsed at least three times with extracted water, and the water samples were filtered with the 0.45 μm membrane filters and filled with bottles with no bubbles. The samples for cation analysis were acidified using ultra-purified HNO₃ to pH < 2, and samples for anion analysis were collected directly after being filtered. Samples for oxygen and hydrogen isotope analysis were collected in 50 mL HDPE bottles without head space. All samples were stored and refrigerated at 4 °C, and the subsequent analysis was conducted within 7 days of sampling.

2.3. Chemical Analysis

2.3.1. Sediment Samples

The moisture content of sediment samples was determined through the stoving method. The porosity was measured using the cutting ring method, and the void ratio was calculated. Dry sediment samples were crushed for mineral composition measurement using X-ray diffraction (XRD) (Bruker D8, Walzbachtal, Germany). The total F (TF) content was determined after alkaline digestion with sodium hydroxide [25]. The water-soluble F (WF) was extracted using deionized water at about 25 °C, with continuous ultrasonic shaking for 30 min, followed by centrifugation at 4000 r/min for 5 min. The exchangeable F (EF) was extracted using 1 M MgCl₂ solution (pH = 7) and continuous ultrasonic shaking for 1 h, followed by centrifugation at 4000 r/min. The extraction solutions of TF, WF, and EF were quantified using a Fluorine–ion selective electrode (HZPL-T503) [10].

The particle size component of sediment samples was determined by a Mastersizer 2000 analyzer (Malvern Instruments, Malvern, UK), and the distribution curve of particle size frequency was generated using Sigma Plot 13.0 software. Test method: 1.0 g of each freeze-dried sediment sample was treated with hydrogen peroxide (15%, 15 mL), hydrochloric acid (10%, 10 mL), and distilled water (30~50 mL), followed by sodium metaphosphate (0.5 M, 10 mL) for testing. The particle sizes of sediment samples are classified as clay (<2 mm), fine silt (2~20 mm), coarse silt (20~50 mm), silty sand (50~100 mm), fine sand (100–250 mm), medium sand (250~500 mm), and coarse sand (500~2000 mm) particles, according to the US–made particle size classification standard.

2.3.2. Clay Porewater and Groundwater Samples

All 41 clay porewater and 20 groundwater samples were measured to reveal the hydrochemical characteristics. The pH, redox potential (Eh) of clay porewater and groundwater samples were measured using a multiparameter water quality monitor (HACH, HQ40d, Loveland, CO, USA). The alkalinity was evaluated by hydrochloric acid titration. The porewater F⁻ (PF) and other anions were analyzed using ion chromatography (Thermo Fisher Scientific, ICS-1100, Waltham, MA, USA), and cations were quantified using ICP–OES (PerkinElmer, avio200, Waltham, MA, USA) after filtration through 0.45 μm plastic filters. Analytical quality control was assured by the analysis of duplicate samples, blanks and standard reference materials according to national standards (HJ/T 164-2020). The relative standard deviations of the blanks and replicates were <5%. The chloride–alkali index (CAI) was used to assess the intensity of the cation exchange. The positive value of CAI indicates that (Ca²⁺ + Mg²⁺) in aqueous media is replaced by (Na⁺ + K⁺).

$$\mathrm{CAI}={\displaystyle \frac{\mathsf{\gamma }\left({\mathrm{Cl}}^{-}\right)-\mathsf{\gamma }\left({\mathrm{Na}}^{+}\right)-\mathsf{\gamma }\left({\mathrm{K}}^{+}\right)}{\mathsf{\gamma }\left({\mathrm{SO}}_4^{2-}\right)+\mathsf{\gamma }\left({\mathrm{HCO}}_3^{-}\right)}},$$

(1)

The hydrogen (²H) and oxygen (¹⁸O) isotopes of groundwater samples collected from monitoring wells were determined using a Finnigan MAT 253 mass spectrometer (Bremen, Germany). The δ²H and δ¹⁸O measurements were compared to internal standards calibrated using Vienna Standard Mean Ocean Water (V–SMOW), with analytical precisions of 0.6‰ and 0.2‰ for δ²H (V–SMOW) and δ¹⁸O (V–SMOW), respectively. Results were presented in standard δ notation as mille deviations from the V–SMOW standard. The measurement precision was ±0.1‰ for δ²H and ±1.0‰ for δ¹⁸O. The analyses described above were completed at the Analysis and Testing Research Center, Institute of Hydrogeology and Environmental Geology, Chinese Academy of Geological Sciences.

3. Results

3.1. Hydrochemical Characteristics of Clay Porewater and Groundwater Samples

According to the lithology and stratigraphic chronology analysis of borehole samples, the depth of −155 m represents the boundary between the Late Pleistocene and Middle Pleistocene, as well as the shallow and deep strata in this study.

The hydrochemical composition of 20 groundwater samples from monitoring wells is shown in

Table 2. Hydrochemical characteristics of groundwater samples from monitoring wells.

	Shallow Groundwater Samples			Deep Groundwater Samples
	Min	Max	Mean	Min	Max	Mean
K+ (mg/L)	0.19	2.6	0.77	0.15	12.2	1.5
Na+ (mg/L)	96.2	517	329.8	41.8	359	210.9
Ca2+ (mg/L)	3.13	149	62.02	6.51	202	37.92
Mg2+ (mg/L)	1.87	177	77.52	1.3	180	23.3
Cl− (mg/L)	79.3	512	295.7	20.8	506	151.8
SO42− (mg/L)	112	743	290.2	25	558	161.7
HCO3− (mg/L)	175	974	499.4	130	651	293.3
CO32− (mg/L)	0	30.4	4.6	0	18.2	5.9
F− (mg/L)	0.31	1.6	0.86	0.2	5.54	1.92
TDS (mg/L)	841	2220	1419	395	960	707
pH	7.18	9.36	7.87	6.93	8.57	8.04
Eh (mV)	−84.3	−11.5	−44.2	−82.8	2.4	−36.6
CAI	−0.69	0.08	−0.35	−0.35	−0.35	−0.35

, and the Piper diagram of them is shown in Figure S1 in the Supplementary Materials. The groundwater samples collected from monitoring wells in the subsidence zone covered aquifers I to IV at depths of −15~−380 m. Hydrochemical types included HCO₃·Cl–Ca·Na, HCO₃·Cl–Na, and SO₄·Cl–Na.

The results show that 50% of groundwater samples were above the Chinese government’s permitted limit of F in drinking water (1.0 mg/L), with a maximum F⁻ concentration of 5.54 mg/L. Gibbs diagram analysis (Figure 2) revealed that water−rock interactions mostly influenced the hydrochemical characteristics of groundwater samples. The groundwater samples collected are mostly from deep strata, which were less affected by evaporation. In addition, the rainfall infiltration in this region is hindered by clayey aquitards of different scales (Figure 1c). However, under the long-term of groundwater overexploitation, multiple groundwater subsidence funnels have formed in this area, resulting in slow lateral runoff and enhanced water−rock interactions [26]. The chemical substances and TDS levels of shallow groundwater were generally higher than those of deep groundwater (Table 2), which might be related to stronger water−rock interactions in shallow aquifers. However, the F⁻ concentration in deep groundwater was generally higher, suggesting that the deep groundwater F⁻ might have other recharge sources.

The hydrochemical composition of clay porewater samples collected from the borehole is shown in

Table 3. Hydrochemical characteristics of clay porewater samples from borehole.

	Shallow Clay Porewater Samples			Deep Clay Porewater Samples
	Min	Max	Mean	Min	Max	Mean
K+ (mg/L)	2.39	15.6	7.6	1.6	8.7	4.0
Na+ (mg/L)	44.6	89.4	63.9	34.8	102	66.2
Ca2+ (mg/L)	5.08	21.6	13.35	7.08	33.16	15.02
Mg2+ (mg/L)	6.12	20.4	12.45	8.4	30.4	16.9
Cl− (mg/L)	4.1	42	19.3	3.0	26.4	13.2
SO42− (mg/L)	7.6	23.9	15.5	4.6	20.9	14.3
HCO3− (mg/L)	120.6	521	283.8	256	540	367.8
CO32− (mg/L)	0	10.6	4.5	9	88	33.8
F− (mg/L)	0.77	2.51	1.48	1.73	4.18	2.77
TDS (mg/L)	554	914	700	587	1104	836.3
pH	7.63	9.01	8.42	8.11	9.19	8.69
Eh (mV)	−119.3	−51.7	−96.2	−143.7	−81.4	−120
CAI	−0.65	−0.30	−0.49	−0.91	−0.15	−0.42

. The hydrochemical type of clay porewater samples was uniformly HCO₃–Na type (Figure S1). Similarly to the groundwater in aquifers, the hydrochemical characteristics of clay porewater were mainly formed by water–rock interactions (Figure 2), and the F⁻ concentration of deep porewater was generally higher than that of shallow porewater. However, the chemical substance and TDS levels of deep clay porewater were higher than those of shallow porewater, indicating that the water–rock interactions in deep clayey sediments might be more sufficient.

Groundwater δ¹⁸O and δ²H values varied from −10.8‰ to −6.3‰ and from −79‰ to −49‰, respectively. The shallow groundwater samples had relatively higher values of δ¹⁸O (ranging from −10.6‰ to −6.3‰) and δ²H (ranging from −70‰ to −49‰) than those in the deep samples (ranging from −10.8‰ to −8‰ and from −79‰ to −67‰, respectively). The hydrogen and oxygen isotopic composition (δ¹⁸O and δ²H) of clay porewater samples varied between −7.8‰ and −5.4‰ and between −65‰ and −52‰, respectively [27].

3.2. Variation in Clayey Sediment Fluoride Levels from Borehole

The TF, WF, EF, and PF contents of clayey sediment samples were 440~792 mg/kg (average: 602 mg/kg), 1.7~20.7 mg/kg (average: 6.6 mg/kg), 1.2~14.0 mg/kg (average: 5.0 mg/kg), and 0.77~4.18 mg/L (average: 2.36 mg/L), respectively. The variations in the contents with increasing depth are shown in Figure 3. The sediment TF content varied significantly with depth, which was attributed to the soil parent material and sedimentary environment. The TF content of clayey sediments was generally higher than that of sandy sediments (286~524 mg/kg, average: 383 mg/kg). The WF and EF contents of shallow sediments were generally higher than those of deep sediments, while the trend of PF content with depth was the opposite. Except for WF and EF, other fluorides are non-transferable F present in sediment mineral structure and are different to release into porewater during water–rock interactions [28]; thus, they are not discussed in this study. The PF concentration initially increased with a fluctuation, peaking at −237 m, and then changed dramatically with depth.

3.3. Physical and Chemical Composition of Borehole Samples

The variations in sediment lithology, moisture content, porosity, and particle size composition with depth were shown in Figure 4. Under the influence of alluvial–lacustrine deposition, the Quaternary strata have formed a complex and diverse lithological structure, mainly divided into clayey sediments and sandy sediments. Clay, silty clay, and silt all belong to clayey sediments.

The moisture content and porosity of sediment samples were 16%~26.7% (average: 20.6%) and 22.6%~40.5% (average: 29.9%), respectively. Under self-weight stress and overexploitation, sediment consolidation gradually enhanced with depth, whereas moisture content and porosity gradually decreased. The proportions of clay (<2 mm), fine silt (2–20 mm), coarse silt (20–50 mm), and sand (>50 mm) particles were 8.75%~29.8% (average: 18.18%), 12.42%~39.77% (average: 26.71%), 16.54%~40.19% (average: 28.69%), and 4.81%~40.38% (average: 26.42%), respectively. The average contents of various particle size components in borehole sediments were comparable, which were mainly influenced by lithology and mineral composition, and showed no obvious variation patterns with depth (Figure 4).

Combined with the test results of XRD and particle sizes (Figure S2), the particles of clayey sediments are relatively fine, consisting primarily of clay and silt particles. The mineral types are mainly clay minerals including illite, chlorite, and muscovite, as well as quartz. The average proportion of clay minerals in a mineral composition exceeds 24%. The sandy sediments have relatively coarse particles, consisting primarily of coarse silt and sand, and the mineral types are mainly quartz and feldspar.

4. Discussion

4.1. Hydraulic Connection Between Clay Layer and Aquifer

The δ²H and δ¹⁸O values of clay porewater and groundwater samples were measured to identify the hydraulic connection between the clay layer and aquifer. The δ²H and δ¹⁸O values of clay porewater and deep groundwater samples (<−155 m) were fitted with two regression lines, as shown in Figure 5.

The δ²H and δ¹⁸O values of shallow groundwater samples (>−155 m) from the subsidence zone are dispersed near the Local Meteoric Water Line (δ²H = 7δ¹⁸O + 1.7) [29] and the Global Meteoric Water Line (δ²H = 8δ¹⁸O + 10). The d-excess values of δ²H and δ¹⁸O of shallow groundwater samples ranged from −8.6‰ to 7.8‰, suggesting that shallow groundwater is probably recharged by atmospheric precipitation. Most of the deep groundwater samples fell on the right side of the Local Meteoric Water Line (Figure 5), and the d-excess values of δ²H and δ¹⁸O of deep groundwater samples ranged from –14.4‰ to 0.1‰, suggesting the presence of recharge sources other than atmospheric precipitation. Given that the CZP is a coastal plain with severe land subsidence, two potential recharge sources could contribute: contemporaneous seawater intrusion and clay porewater compression–release. The depths of deep aquifers in western CZP are below the seawater-affected area [23], and the δ²H–δ¹⁸O isotope composition of deep groundwater differed significantly from seawater (δ²H: −20‰~−10‰, δ¹⁸O: −1‰~0, refer to [27]). The oxygen in water molecules undergoes isotope exchange with the oxygen in mineral lattice (such as calcite). The δ¹⁸O value in rocks is usually higher than that in water. After a long geological period, this exchange increases the proportion of δ¹⁸O in groundwater, leading to a deviation of the δ¹⁸O value to the right of the meteoric water lines [30]. Therefore, the modern seawater intrusion has little effect on the deep aquifers in this area. The slope of the Deep Groundwater Line (4.0) is similar to that of the Clay Porewater Line (3.4) (Figure 5), suggesting a mixing effect between clay porewater and deep groundwater.

4.2. Occurrence Forms of F in Sediments Under Different Sedimentary Environments

The parent material and sedimentary environment have a substantial impact on sediment fluoride content [31]. Particle size distribution is an important parameter for analyzing the sedimentary environment, and it is primarily influenced by transport medium, mode, and hydrodynamic conditions [32]. Furthermore, particle size parameters are closely related to the sedimentary environment, and specific sedimentary environments often exhibit corresponding characteristics of particle size parameters. The average kinetic energy of a transport medium is reflected in the average particle size (MG). The sorting and dispersion levels of particles are quantified by the standard deviation (σi). Kurtosis (KG) describes the sharpness or broadness of the particle size fraction curve, and an excessively high or low KG value implies significant sorting ability in a sedimentary environment. The grading standards and calculation methods of them are shown in Table S1. Furthermore, particle composition has a significant impact on sediment sorption capacity. The particle size of sediment directly affects its hydrodynamic properties, specific surface area, and sorption–desorption equilibrium, all of which then affect the fluoride content in sediments.

The σi and MG values of shallow sediment (>–155 m) were 7.6~18.3 and 11.75, respectively, indicating extremely poor sorting. The KG ranged from 1.06 to 2.44, indicating a kurtosis variation range from mesokurtic to very leptokurtic (Table S1). The hydrodynamic conditions of alluvial strata are unstable, and clay layers contain varying proportions of sand particles. There are four major types of particle size fraction curves of representative borehole samples from different depths (Figure 6). When the hydrodynamic condition of the sedimentary environment changes significantly, the sorting ability changes as well, so the sediments formed in diverse environments retain their original particle size characteristics, resulting in a clear bimodal or multimodal particle size fraction. The σi values of deep sediment (<−155 m) ranged from 1.7 to 5.2, with an MG of 3.28, indicating poor sorting but better than shallow sediment. The KG values varied from 0.84 to 1.32, indicating a distribution from mesokurtic to leptokurtic (Table S1). The particle size fraction curves of most samples are unimodal (Figure 6d), with a few being bimodal (Figure 6c), indicating that no significant environmental changes were discovered. Therefore, a more stable sedimentary environment in deep strata might lead to more sufficient water–rock interactions. The strong hydrodynamic conditions in shallow strata might be related to the intense alluviation that occurred after the last glacial period [33].

Correlation analyses between sediment F forms, mineral contents, and particle size components were performed to further analyze the occurrence mechanism of fluoride in sediments (Figure 7). The sediment TF content is significantly positively correlated with WF, indicating that the variation in WF and TF content were similar to depth. The WF and EF contents gradually decreased in depth, and both of them are significantly negatively correlated with the proportion of sand components in particle composition. The higher the content of sandy components, the lower the F content in sediments, which is consistent with the lower TF content in sandy sediments compared to clayey sediments. In addition, the WF content is significantly positively correlated with the proportions of fine silt in particle composition, clay minerals (illite, montmorillonite, chlorite, muscovite, etc.), and muscovite in mineral composition. The EF content is significantly positively correlated with the proportions of clay and fine silt in particle composition, as well as the clay minerals, chlorite, and muscovite in mineral composition. Therefore, the clay and fine silt components, and the clay minerals represented by muscovite, have a significant impact on the occurrence of transferable F in sediments.

Clayey sediments are composed of clinochlorite, fluorapatite, muscovite, biotite, etc., which are the most common F–bearing clay minerals (for example, muscovite, KAl₂(AlSi₃O₁₀)F₂) and can enrich groundwater F by dissolution [21,22]. This could account for the significant positive correlation between WF and clay mineral levels. Multiple layers of sand–silt–clay evolved in the CZP, and clay minerals with fine particles and a larger specific surface area than sandy minerals are more conducive to F enrichment in groundwater systems [28]. Clay minerals such as clinochlorite or muscovite can adsorb F via cation bond bridges, resulting in a significant positive correlation between their proportions and EF content (Figure 7). Except for clay minerals, high F groundwater is associated with the desorption and reductive dissolution of Fe (oxy)hydroxides in aquifers [18]. Li et al. [34] has confirmed the presence of abundant Fe (oxy)hydroxides in the clayey sediments of the CZP. Therefore, the physical and chemical properties of clayey sediments promote the formation of EF. Sediments TF, WF, and EF are significantly negatively correlated with the proportion of sand in particle composition. The main component of sandy sediments is quartz (SiO₂), which has relatively stable chemical properties and weak adsorption capacity compared to aluminosilicate and metal minerals in clayey sediments. Meanwhile, the content of F–bearing minerals in sandy sediments is relatively low, so the effect of sandy minerals on groundwater F enrichment is limited.

The F-bearing minerals in the CZP sediments are predominantly derived from F-rich rock masses widely distributed in the eastern foothills of the Taihang Mountains, west of the plain, and formed by weathering, transportation, and deposition. Although the hydrodynamic conditions in the shallow and deep layers were different, they shared the same sediment source, resulting in fluctuating TF content with depth but no significant difference between shallow and deep sediments. The proportion of transferable F (WF and EF) in TF was low, but they showed a significant decreasing trend with depth (Figure 3), indicating that they underwent a significant transformation under a relatively stable sedimentary environment and compression in deep strata, promoting the formation of high F porewater in deep clayey sediments.

4.3. Transformation Mechanisms of Clayey Sediment F into Porewater

The hydrochemical characteristics of all clay porewater samples were controlled by water–rock interactions based on the Gibbs diagram (Figure 2). The (HCO₃⁻/Na⁺) versus (Ca²⁺/Na⁺) diagrams were created to demonstrate the impact of three different lithologies weathering (carbonates, silicates, and evaporites) on hydrochemistry [35]. The clay porewater samples were mostly plotted in the silicate weathering zone (Figure 8), suggesting that silicate dissolution was the primary hydrogeochemical process. In contrast, most groundwater samples from monitoring wells fell within the evaporite dissolution zone, implying that the main minerals involved in water–rock interactions were different between clayey aquitards and sandy aquifers, which could also be one of the causes of high F porewater in clayey sediments.

The release of fluoride from clayey sediments and specific hydrochemical conditions are important factors affecting the enrichment of F in porewater [23]. The sources and influencing factors of porewater F⁻ were analyzed through principal component analysis (PCA) (Figure 9). The sediment TF, WF, EF, porewater pH, Ca²⁺, Na⁺, Cl⁻, HCO₃⁻, IS, and CAI levels were selected as important variables affecting porewater F⁻ concentrations. The KMO of all groundwater samples was higher than 0.5, indicating that the data were appropriate for PCA. Bartlett’s test showed a significance level of <0.05 (near to 0), suggesting a statistical relationship between these variables.

In shallow clayey sediments (>−155 m), the porewater pH, F⁻, Na⁺, HCO₃⁻, Cl⁻, and WF belong to the first principal component (total variance interpretation rate: 36.2%); the EF, CAI, and IS belong to the second principal component (total variance interpretation rate: 25.3%); and the TF and Ca²⁺ belong to the third principal component (total variance interpretation rate: 15%). The porewater F⁻ concentration is significantly positively correlated with WF, pH, and HCO₃⁻ levels, and significantly negatively correlated with Ca²⁺ concentration (Figure S3a). The first principal component represents factors that have an impact on porewater F⁻ concentration or have a similar variation with it. High F groundwater generally belongs to the Na–HCO₃ type, with relatively low Ca²⁺ concentrations and a neutral to alkaline pH (7–9) [11]. Under the water–rock interactions, high F porewater also exhibits the hydrochemical characteristics of high HCO₃⁻ and low Ca²⁺ concentrations. According to XRD results (Figure S2), fluorite (CaF₂) is the main component of WF, and it entered the porewater mostly through dissolution (Equation (2)). The saturation index of fluorite in clay porewater was −1.96~−1.13, indicating that all shallow porewater samples were undersaturated in terms of fluorite, which has a strong dissolution potential. Therefore, the dissolution of sediment WF has a significant impact on the enrichment of porewater F⁻. The dissolution of silicate minerals, such as albite, can release HCO₃⁻ into porewater, which accelerates the dissolution of fluorite by precipitating with Ca²⁺ (Equations (2)–(4)) [36]. This could account for the significant positive correlation between porewater F⁻ and HCO₃⁻ concentrations. Ionic interactions (desorption and cation exchange) are significant hydrogeochemical mechanisms influencing F mobilization in groundwater. The porewater F⁻ concentration and pH value belong to the same principal component, and they have a positive correlation (Figure S3a). Firstly, the clay porewater was weakly alkaline (Table 2), and the competitive adsorption between OH⁻ and F⁻ could facilitate the release of adsorbed F due to similar ion radius and charge numbers [37], which was summarized by Equations (5) and (6) [38]. Furthermore, under (weak) alkaline conditions, clay mineral surfaces exhibited a predominantly negative charge, promoting the desorption of adsorbed F into groundwater [39], and the OH⁻ may also restrict Ca²⁺ dissolution and enhance F⁻ enrichment in groundwater [28]. Therefore, the OH⁻ replacement and desorption might promote F release from minerals into porewater. The CAI value of shallow porewater samples varied from −0.65 to −0.30, suggesting that (Na⁺ + K⁺) in the aqueous medium was replaced by (Ca²⁺ + Mg²⁺), which was conductive to the enrichment of F⁻ in clay porewater.

CaF₂ → Ca²⁺ + 2F⁻

(2)

Ca²⁺ + 2HCO₃- → CaCO₃ + CO₂ + H₂O

(3)

Ca²⁺ + Mg²⁺ + 4HCO₃- → CaMg(CO₃)₂ + 2CO₂ + 2H₂O

(4)

KAl₂(AlSi₃O₁₀)F₂ + 2OH⁻ → KAl₂(AlSi₃O₁₀)(OH)₂ + 2F⁻

(5)

K(Mg, Fe)₃(AlSi₃O₁₀)F₂ + 2OH⁻ → KMg₃(AlSi₃O₁₀)(OH)₂ + 2F⁻

(6)

In deep clayey sediments (<−155 m), the porewater pH, F⁻, Ca²⁺, IS, WF and EF belong to the first principal component (total variance interpretation rate: 42.8%); the Na⁺, Cl⁻, and CAI belong to the second principal component (total variance interpretation rate: 20.8%); and the TF and HCO₃⁻ belong to the third principal component (total variance interpretation rate: 12.2%). The porewater F⁻ concentration is significantly positively correlated with WF, EF, pH, and HCO₃⁻ levels, and significantly negatively correlated with Ca²⁺ concentration (Figure S3b). Similarly to shallow porewater, the dissolution of WF and the hydrochemical properties of porewater, such as pH and HCO₃⁻ concentrations, have a significant impact on deep porewater F⁻ concentration. The saturation index of fluorite in porewater varied from −1.36 to −0.52, indicating that all deep porewater samples were under–saturated in terms of fluorite. The saturation index of fluorite in deep porewater was generally lower than that of shallow porewater, suggesting that the fluorite in deep porewater was more prone to precipitation. Meanwhile, the porewater F⁻ and IS belong to the same principal component. This could be due to the stronger water–rock interactions and higher concentrations of hydrochemical components in deep porewater under the stable sedimentary environment and clay layers compression. The CAI values of deep porewater samples ranged from −0.91 to −0.15, suggesting that (Ca²⁺ + Mg²⁺) substituted (Na⁺ + K⁺) in the aqueous medium, thereby facilitating CaF₂ dissolution. In addition, the EF and porewater F⁻ belong to the same principal component, and their positive correlation indicates a similar variation with depth. Therefore, the desorption of sediment EF in deep strata might be stronger. Deep clayey sediments had higher porewater F⁻ concentrations and lower sediment WF and EF contents than those of shallow clayey sediments (Figure 3), indicating that deep sediments might experience more sufficient water–rock interactions.

The deep aquifers (aquifers III and IV) in the subsidence zone were historically overexploited and experienced substantial clay layers compression. Clayey sediment compression not only releases a large amount of porewater into adjacent aquifers, but it also influences the water–rock interaction that occurs within them. The moisture content and porosity of deep samples were generally lower than those of shallow samples, and the moisture content of most samples was lower than the plastic limits. The fragmentation of sediment particles during compression enhanced the contact area between minerals and porewater, hence promoting mineral components’ dissolution. The release of sediment chemical substances might increase porewater IS [18,40], resulting in a positive correlation between F⁻ concentration and IS in deep porewater (Figure S3b). Therefore, long-term compression might have resulted in a lower WF content in deep samples than in shallow samples (Figure 3). The change from a loose to a compact arrangement of clay mineral particles lowered the specific surface area [41], and the adsorption sites on the mineral surface gradually decreased under compression, releasing the EF into porewater. In conclusion, overexploitation-induced compression enhanced the release of sediment F into porewater, increasing the porewater F⁻ concentration in deep clayey sediments (Figure 3).

This study investigates the distribution characteristics and enrichment mechanisms of fluoride in alluvial–lacustrine facies clayey sediments in the land subsidence area. The primary cause of high F porewater is the high background content of transferable F in sediments, which can enter the porewater via various water–rock interactions. Deep clayey sediments exhibit more complex F migration and transformation processes and higher porewater F⁻ concentrations under compression compared to shallow samples. Liu et al. discovered that the ¹⁴C ages of groundwater at varying depths from the CZP deep aquifer were close, suggesting an intense vertical mixing effect throughout the deep aquifer [23], so the released porewater from clayey sediments could affect the groundwater quality of the entire deep aquifer.

5. Conclusions

This study identified the distribution characteristics and enrichment mechanisms of fluoride in alluvial–lacustrine facies clayey sediments in the land subsidence area of CZP. Clayey sediments with fine particles have a higher F content and a greater potential for F enrichment in porewater than sandy sediments. Porewater F⁻ is mostly derived from the dissolution of sediment water-soluble F and the desorption of exchangeable F. The occurrence of released F in porewater was influenced by hydrochemical conditions such as pH. The stable sedimentary environment and clay layer compression accelerated the transformation of sediment transferable F into porewater by enhancing water–rock interactions. The hydraulic connection between deep aquifers and clay layers was detected using δ²H and δ¹⁸O isotopes. Therefore, the release of high F porewater from clayey sediments will increase the groundwater F⁻ concentration in adjacent aquifers.

Reasonable regulation of groundwater levels is crucial for scientifically preventing the formation of high F groundwater in land subsidence areas. Future research should investigate the distribution of high-quality water sources in this area, establish a mutual feedback relationship between groundwater level, clay compression, and groundwater F⁻ concentration, and consider the critical water level that affects the F release of clayey sediments in the evaluation system of groundwater mining output.