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Article

Apatite as an Indicator of Sedimentary Environment and Diagenesis for the Hengyang Neoproterozoic Iron Formation, South China

by
Chuangye Zhang
1,2,
Lei Liu
1,2,*,
Kuanxin Huang
1,2,3 and
Tianyang Hu
4,5
1
School of Geosciences and Info-Physics, Central South University, Changsha 410083, China
2
State Key Laboratory of Critical Mineral Research and Exploration, Central South University, Changsha 410083, China
3
Urban Geological Survey and Monitor Institute of Hunan Province, Changsha 410007, China
4
Henan Institute of Ultrapure Mineral Materials, Zhengzhou 450016, China
5
Henan Academy of Geology, Zhengzhou 450016, China
*
Author to whom correspondence should be addressed.
Minerals 2026, 16(4), 392; https://doi.org/10.3390/min16040392
Submission received: 6 February 2026 / Revised: 2 April 2026 / Accepted: 7 April 2026 / Published: 10 April 2026
(This article belongs to the Section Mineral Geochemistry and Geochronology)
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Abstract

Neoproterozoic iron formations (NIFs), deposited during Cryogenian glaciation events, are critical for understanding early Earth oxidation events and the evolution of glacial–interglacial environments. Apatite, a common accessory mineral in iron formations, holds significant implications for sedimentary environments and diagenetic processes, but these aspects remain underexplored. This study focuses on the Hengyang NIF in the Nanhua Basin, South China. Using whole-rock geochemistry and major and trace element analysis of apatite, we investigate the environmental significance of apatite and associated diagenetic processes. Our results show that the Hengyang NIF are formed through the mixing of low-temperature hydrothermal fluids, seawater, and terrigenous detrital materials, with hydrothermal contributions increasing progressively from the bottom to the top of the iron formation layers. Whole-rock geochemical proxies indicate that the depositional water column evolved from relatively oxidizing to weakly oxidizing conditions. The study further demonstrates that the rare earth element patterns in apatite, characterized by middle rare earth element enrichment, are primarily controlled by porewater chemistry during diagenesis. In contrast, Ce anomalies and the V/Cr and V/(V + Ni) ratios in apatite, which are strongly influenced by fluid–rock interactions and magnetite recrystallization, no longer reliably reflect the primary depositional environment. The Th/U ratio in apatite, due to its geochemical stability, aligns with whole-rock trends and serves as a more reliable redox proxy. Based on these findings, we propose a three-stage depositional-diagenetic model: the early and late stages are characterized by high-energy, rapid sedimentation with minimal diagenetic modification, while the middle stage is dominated by low-energy, stagnant conditions with slow sedimentation rates, leading to prolonged diagenesis and significant decoupling of mineral geochemical signatures. This study emphasizes the need to distinguish between sedimentary and diagenetic signals when using mineral geochemical proxies to reconstruct paleoenvironments and provides new insights into the genesis of Neoproterozoic iron formations.

1. Introduction

Neoproterozoic iron formations (NIFs) represent archives of early Earth surface system evolution, particularly documenting the interactions among atmospheric-oceanic redox conditions, biological activity, and global glaciation events [1,2,3,4]. Unlike Archean banded iron formations, which formed under low atmospheric oxygen levels and pervasive deep-sea anoxia [5,6], NIFs were deposited during a period of elevated atmospheric oxygen concentrations [5]. Globally, NIFs are closely associated with the Sturtian and Marinoan glaciations. Their formation has been attributed to processes such as post-glacial sea level rise, the development of chemically stratified marine basins, and the upwelling of iron-rich fluids [7,8,9,10,11,12,13,14]. The Nanhua Basin in South China preserves continuous Neoproterozoic volcanic-sedimentary sequences, within which NIFs provide a critical window into paleo-marine environments and metallogenic processes during this period [15,16,17,18,19,20,21]. However, significant controversies persist regarding the material sources of these iron formations (IFs), the redox state of the paleo-ocean during deposition, and the extent to which diagenetic processes have modified the primary geochemical signals [20].
Apatite, a common accessory mineral in IFs, has gained attention as a potential proxy for reconstructing paleo-seawater chemistry and diagenetic fluid properties due to its capacity to host abundant rare earth elements (REE) and yttrium (Y), as well as its sensitivity to fluid compositions [22,23,24,25,26,27]. Traditionally, trace element compositions of bulk rocks and minerals have been widely used to reconstruct paleoenvironments. However, growing evidence suggests that diagenetic processes, particularly interactions between pore waters and mineral phases, may significantly modify or even obscure primary sedimentary geochemical signals, leading to potential misinterpretations of paleoenvironmental conditions [28,29,30]. Therefore, distinguishing sedimentary signals from diagenetic overprints is essential for accurately interpreting the environmental records of IFs and remains a key challenge in current research.
NIFs in South China’s Nanhua Basin were deposited across diverse environments from continental shelf (e.g., Fulu NIF) to deep-water settings (e.g., Xinyu NIF) [15,16,17,18,19,20]. These NIFs are considered to have formed during the Sturtian deglaciation (694 ± 17 and 691 ± 12 Ma) [20,21], when glacial influence was limited, allowing open-water conditions to develop [19,20]. This facilitated direct O2 exchange between the atmosphere and ocean, promoting the oxidation and deposition of Fe2+ to form IFs [15,16,17,18,19,20,21] (Figure 1a). The Hengyang region, situated in the Nanhua Basin’s interior, hosts extensive NIFs with abundant apatite occurrences, and the genesis of the Hengyang iron formation has been insufficiently studied.
Here, we focus on the Hengyang NIFs of the Nanhua Basin and apatite within NIFs. We present whole-rock geochemical data of the Hengyang NIFs and major and trace element analyses of apatite. The aim is to characterize the depositional environment. Additionally, we assess the reliability of apatite as a proxy for primary sedimentary signals, considering the potential influence of diagenetic processes.

2. Geological Setting and Samples

The Nanhua Basin is situated within the South China Block, which formed through the amalgamation of the Yangtze Block and Cathaysia Block during the early Neoproterozoic [20]. The South China Block preserves a complete volcanic and sedimentary rock sequence spanning the Tonian–Cryogenian periods [19,31].
The Neoproterozoic sedimentary succession of the Nanhua Basin (850–540 Ma) can be divided into three stages [20,32,33]: Stage I comprises the Sibao Group (also known as the Lengjiaxi Group or the Fenjingshan Group, 850–820 Ma), characterized by continental alluvial and littoral–shallow marine deposits, followed by the Banxi Group (814–715 Ma), which is dominated by volcanic and pyroclastic facies [34]; Stage II (715–635 Ma) records glacial–interglacial sedimentary sequences, including the Chang’an, Fulu, Datangpo, and Nantuo formations [35]; Stage III (635–540 Ma) is represented by platform carbonate deposits formed under transgressive marine conditions. Within Stage II, the Chang’an and Fulu Formations represent major glacial deposits correlated with the Sturtian glaciation [36] (>659 Ma), whereas the Nantuo Formation corresponds to the Marinoan glaciation [37,38]. The Datangpo Formation records interglacial sedimentation between these two ice ages [39,40].
The NIFs of the Nanhua Basin were deposited during the Sturtian glaciation and occur at the base of the lower Fulu Formation. These deposits are distributed across a range of depositional environments from the east–west–trending continental shelf to deeper-water settings of the basin and have locally undergone metamorphism to form slate and schist (Figure 1a) [19]. The Fulu IFs formed in a shelf environment and the Xinyu IFs in a basin setting, whereas the Jiangkou and Hengyang IFs were deposited in relatively deeper sub-basins within the basin [39,41,42]. Lan et al. (2015) utilized detrital zircon U-Pb geochronological data from the IF-hosting Xifang formation, which is composed of iron-bearing rocks, as well as the underlying Shenshan Formation and Shangshi Formation, to determine the depositional age of the Nanhua IFs between 694 ± 17 and 691 ± 12 Ma [35]. Meanwhile, Wang et al. (2025) used the weighted average age of the youngest detrital zircons in the host rocks of the intermediate conglomerates, which is approximately 694 Ma [21]. This suggests that the NIF in the Nanhua Basin was deposited during the deglacial stage around 694 Ma ago.
In the Hengyang area, the Fulu Formation consists predominantly of littoral–shallow marine flysch deposits, primarily composed of clastic rocks interbedded with sandstones and carbonates (Figure 1b and Figure 2). The iron formation is hosted within the Fulu Formation, with pebbly sandstone observed 30–80 m beneath the iron formation stratum and carbonate rocks appearing 30–50 m above it. The sedimentary sequence follows a rhythmic pattern: pebbly sandstone → clastic rocks → iron formation → carbonate rocks, indicative of a transgressive sequence. The overlying strata consist of Nantuo Formation tillite, which unconformably overlies the iron formation, while the underlying strata are composed of sandstone (Figure 2). The northeastern part of the study area is occupied by the Guandimiao Pluton, emplaced during the Late Triassic (ca. 226–220 Ma), and mainly composed of amphibole–muscovite monzogranite [43]. Thermal metamorphism has variably altered both the wall rocks and iron formation, with the metamorphic grade increasing with proximity to intrusive bodies. Two southeast-trending dykes (~2 m wide) occur in the area and locally intrude the IFs and associated clastic metasedimentary rocks (Figure 1b). The sampling sites are located ~500 m from the Guandimiao granite pluton and ~100 m from the nearest dyke. Petrographic observations reveal no occurrence of typical contact metamorphic minerals or recrystallization of quartz, suggesting that the samples were not significantly affected by contact metamorphism and largely preserve their primary characteristics.
Samples were collected from the Hengyang Qidong mining area (112°00′36″ E–112°01′19″ E; 26°55′38″ N–26°57′03″ N). This study focuses on eight magnetite quartzite samples, primarily composed of magnetite and quartz. The sampling locations are far from the granite veins and away from the granite body. The collected IF samples have a low degree of metamorphism, avoiding the influence of igneous rock. The iron formation in the study area is classified into distinct stratigraphic units based on TFe2O3 content. The upper unit exhibits high TFe2O3 content (TFe2O3 > 60%). The middle unit contains moderate TFe2O3 content (60% ≥ TFe2O3 > 30%). In contrast, the lower unit is characterized by notably low TFe2O3 content (TFe2O3 ≤ 30%). Field sampling was conducted across these stratigraphic intervals. Samples 24-MC-4, 5, 6, 7, 8, and 9 were collected along a vertical profile from an exposed excavation face in the northern mining area, representing a continuous section from the lower to the intermediate stratigraphic units of the iron formation (Figure 2 and Figure 3b). Additionally, samples ZK3701-1 and ZK3701-2 were obtained from the upper unit of the iron formation. ZK3701 is a borehole. Although the borehole is located in the Quaternary cover area, the samples ZK3701-1 and ZK3701-2 were collected from the subsurface Neoproterozoic iron formation lithologies at depths of 480 to 520 m. Accordingly, the collected iron formation samples were categorized into three groups corresponding to these units: bottom stratum unit (24-MC-7, 8, 9), intermediate stratum unit (24-MC-4, 5, 6), and upper stratum unit (ZK3701-1, 2).
The iron formation samples are dark gray in color (Figure 3c). Magnetite in the iron formation occurs as disseminated grains, serving as the primary iron-bearing mineral (Figure 3d–f). It typically exhibits equigranular textures and occasionally appears as tabular or prismatic crystals. Magnetite grains range from euhedral to anhedral, with diameters of 0.1–0.5 mm (finer-grained in some samples). Quartz inclusions within magnetite, likely formed during post-depositional alteration, are also observed. Quartz constitutes 30%–50% of the iron formation and is uniformly distributed as fine-grained aggregates or cryptocrystalline phases.

3. Analytical Methods

3.1. Whole-Rock Geochemical Analyses

In this study, five iron formation samples were analyzed for whole-rock major and trace elements analyses. The five samples (24-MC-5, 24-MC-6, 24-MC-8, ZK3701-1, ZK3701-2) represent new test data for this study. All analytical work was conducted at the Key Laboratory of Magmatic Ore Formation and Prospecting, Ministry of Natural Resources, Xi’an Geological Survey Center, China Geological Survey.
Whole-rock major elements were determined using an X-ray fluorescence (XRF) spectrometer. The samples were pulverized and ground to a 200-mesh powder. Three aliquots were prepared for analysis. The first aliquot underwent acid digestion, evaporation to dryness, dissolution, and volumetric analysis using an Agilent 5100 ICP-OES (Agilent Technologies Inc., Santa Clara, CA, USA), which allowed for preliminary estimation of S, Ca, Fe, Mn, and Cr concentrations. The second aliquot was dried, weighed, and mixed with a flux (lithium tetraborate, potassium metaborate, and lithium nitrate) in a platinum crucible. The mixture was fused at 1050 °C using a RYL-4S high-precision fusion machine (Luoyang Kepuxin Experimental Equipment Co., Ltd., Luoyang, China). Major elements were subsequently measured using a PANalytical PW 2424 XRF spectrometer (Malvern Panalytical B.V., Almelo, The Netherlands), achieving detection limits of 0.01% and analytical errors < 2%.
Trace elements were analyzed using Agilent 7900 inductively coupled plasma mass spectrometry (ICP-MS, Agilent Technologies Inc., Santa Clara, CA, USA) with a combined internal–external standard method, using Rh as the internal standard. Each element underwent five replicate scans to ensure precision, with relative standard deviations (RSD) < 5%. Certified reference materials (BCR-2 basalt, BHVO-2 basalt, AGV-2 andesite) were analyzed in each batch to monitor procedural accuracy, ensuring trace element deviations <10% from reference values. All trace element analyses were maintained within a 10% error margin.

3.2. Backscattered Electron Imaging and Apatite Element Analysis

Apatite single-mineral separation and analytical sample preparation were carried out at Langfang Yuneng Testing Technology Co., Ltd. in Langfang, China. Apatite grains were extracted from these IF samples: 24-MC-(4~9), ZK3701-(1, 2). The rock samples collected in the field were first crushed and sieved to 80–120 mesh. Heavy mineral fractions were then separated using electromagnetic separation and heavy-liquid techniques, and transparent, well-formed apatite grains were handpicked under a binocular microscope. Representative apatite grains were subsequently mounted in epoxy resin and, after curing, the mounts were ground and polished to obtain a flat surface. Backscattered electron (BSE) imaging of apatite grains on the single-mineral mounts was then performed. Electron probe microanalysis (EPMA) was employed to characterize mineral morphology, internal textures, and determine mineral compositions.
BSE images were obtained at the Key Laboratory of Metallogenic Prediction and Monitoring of the Ministry of Education at the School of Geosciences and Info-Physics, Central South University in Changsha, China. Analysis was performed using a SHIMADZU EPMA-1720H electron probe microanalyzer (Shimadzu Corporation, Kyoto, Japan). The samples consisted of apatite grains mounted in epoxy resin and coated with a 15–20 nm carbon layer. Analyses were conducted at an accelerating voltage of 10 keV with a beam current of 0.5–5 nA.
Major element analyses of apatite from IF samples 24-MC-5, 24-MC-8, and ZK3701-1 were conducted at the Key Laboratory for the Study of Magmatism and Mineralization, the Ministry of Natural Resources (Xi’an Center of Geological Survey), China Geological Survey, Xi’an, China. The analyses were performed using a JEOL JXA-8230 electron probe micro-analyzer (EPMA, JEOL Ltd., Tokyo, Japan)under the following operating conditions: an accelerating voltage of 20 kV, a beam current of 20 nA, and a beam diameter of 3 μm. The measured elements included CaO, P2O5, MgO, FeO, and SO3, as well as the halogens F and Cl. Detection limits for major elements were 0.01%.
Apatite trace elements were analyzed using a Teledyne Cetac HE 193 nm laser ablation system (Teledyne CETAC Technologies, Omaha, NE, USA)coupled with an Analytik Jena PlasmaQuant MS Elite ICP-MS (Analytik Jena, Jena, Germany) at the Key Laboratory of Metallogenic Prediction and Monitoring of the Ministry of Education at the School of Geosciences and Info-Physics, Central South University in Changsha, China. Analysis parameters included Ar/He carrier gas, 1.5–2 J/cm2 energy density, 40 μm spot size, 5 Hz frequency, and 70 s ablation time (20 s background, 30 s signal, 20 s washout). NIST612 was used as the external standard, with 43Ca as the internal standard. The analytical precision was better than 5%. Data were processed using GLITTER 4.4.4 (developed by GEMOC, Macquarie University, Sydney, Australia), and were carried out according to the standard method [44].

4. Results

4.1. Whole-Rock Major and Trace Elements

The whole-rock chemical compositions of all samples are presented in Table 1. The major and trace element data for samples 24-MC-4, 24-MC-7, and 24-MC-9 are from another article co-authored by the authors of this paper, Huang et al., 2025 [45]. The IF samples display high TFe2O3 + SiO2 contents (78.4–94.9 wt.%), which is consistent with the geochemical characteristics typically observed in Neoproterozoic iron formation [46]. They also show elevated clastic components, such as Al2O3 (2.12–9.55 wt.%) and Zr (49–276 ppm). Total REE contents (ΣREE) range from 73.4 to 260 ppm, with low Y/Ho ratios of 25.2–27.9. Post-Archean Australian Shale (PAAS)-normalized REY (rare earth elements and yttrium) patterns show a relative enrichment in heavy REE (HREE) and a depletion in light REE (LREE; Figure 4). The overall REY patterns profiles are relatively flat, with no significant Eu anomalies (Eu/Eu* = 0.96–1.10; Eu/Eu* = 2(Eu)N/[(Sm)N + (Gd)N]) or Ce anomalies (Ce/Ce* = 0.95–1.00; Ce/Ce* = 2(Ce)N/[(La)N + (Pr)N]).

4.2. Textures and Element Compositions of Apatite

A total of 21 major element analyses of apatite from samples 24-MC-5, 24-MC-8, and ZK3701-1 are summarized in Table S1. The apatite from the upper stratum unit (ZK3701) exhibits consistent concentrations of P2O5 (40.45–41.82 wt%) and CaO (53.88–54.72 wt%), while showing considerable variations in F (1.95–4.94 wt%) and Cl (0.05–0.37 wt%) contents. In contrast, apatite from the intermediate (24-MC-5) and bottom (24-MC-8) stratum units displays relatively uniform P2O5 (40.87–42.56 wt%) and CaO (52.31–54.88 wt%) concentrations, with moderate F contents (3.05–4.23 wt%) and very low Cl levels (≤0.04 wt%).
To further clarify the mineral species and their chemical evolution, an F-Cl-OH ternary diagram was constructed (Figure 5). The results show that apatite grains from the bottom and intermediate units plot tightly at the pure fluorapatite end-member. This chemical purity, combined with the presence of abundant alteration voids (Figure 6b,c), indicates that these grains represent authigenic fluorapatite that stabilized through extensive diagenetic recrystallization and decarbonation of metastable precursors. Conversely, the apatite from the upper stratum unit exhibits a more scattered distribution with higher OH and Cl components, consistent with its interpreted terrigenous detrital origin and subsequent weathering overprints.
The apatite grains in the upper stratum unit (ZK3701-1, ZK3701-2) exhibit relatively smooth surfaces, are almost free of irregular impurities or fluid inclusions, and display distinct rounding. Their widths range from approximately 0.05 to 0.3 mm (Figure 6a). The apatite grains in the intermediate stratum unit (24-MC-4, 24-MC-5, 24-MC-6) show rough surfaces with abundant irregular impurities and fluid inclusions. Some grains contain small, irregularly shaped cores at their centers. The grain widths are about 0.03 to 0.15 mm (Figure 6b). The apatite grains from the bottom stratum unit (24-MC-7, 24-MC-8, 24-MC-9) typically present rough surfaces containing numerous irregular impurities and fluid inclusions, with small, irregularly shaped cores located centrally within the grains. These grains are predominantly subhedral in shape, with widths ranging from approximately 0.03 to 0.15 mm (Figure 6c).
A total of 146 trace element analyses of apatite from eight samples are summarized in Table S2. The apatite compositions exhibit considerable variation across different stratigraphic units (Figure 6a–c):
Apatite from the upper stratum unit (ZK3701-1, ZK3701-2): These samples show high ΣREE content (average: 2469 ppm; “average” refers to the arithmetic mean), with LREE-enriched patterns (average LREE/HREE ratio: 3.0; Figure 6a). They have low Ti (average: 6.6 ppm) and Zr (average: 0.56 ppm) concentrations. Ce anomalies are negligible (average Ce/Ce*PAAS = 0.94), while Eu anomalies are heterogeneous (Eu/Eu*PAAS = 0.28–2.04, average: 1.03).
Apatite from the intermediate stratum unit (24-MC-4, 24-MC-5, 24-MC-6): These apatite samples have lower ΣREE content (average: 1228 ppm) and exhibit pronounced MREE (middle REE) enrichment (MREE/MREE* average: 2.49; Figure 6b; here MREE is the average of Eu, Gd, Tb, and Dy (normalized to PAAS), MREE* is the average of the LREE and HREE values). Ti (average: 39 ppm) and Zr (average: 16 ppm) concentrations are moderately elevated. Weak negative Ce anomalies (average Ce/Ce*PAAS = 0.81) and weak positive Eu anomalies (average Eu/Eu*PAAS = 1.12) are present.
Apatite from bottom stratum unit (24-MC-7, 24-MC-8, 24-MC-9): These samples show high ΣREE contents (average: 2983 ppm), with mildly MREE-enriched patterns (MREE/MREE* average: 1.54; Figure 6c). Elevated Ti (average: 235 ppm) and Zr (average: 55 ppm) concentrations are observed. Both samples lack Ce anomalies (average Ce/Ce* = 0.97) and Eu anomalies (average Eu/Eu*PAAS = 0.97).

5. Discussion

5.1. Source of the Hengyang NIF

NIFs are typical chemical sedimentary rocks, although not all are purely chemical precipitates [2,52,53]. It is widely accepted that seawater and hydrothermal fluids serve as the primary sources of ferrous and siliceous components in IFs, with most IFs incorporating clastic materials during deposition [54]. The Hengyang NIF exhibits elevated high-field-strength elements (HFSEs, e.g., Zr > 49 ppm) and higher ∑REE concentrations, which can be attributed to the presence of mafic and ultramafic components in the clastic fraction [55]. Strong positive correlations between Al2O3 and TiO2 (R2 = 0.64) in whole-rock samples, along with negative correlations between TFe2O3 and Al2O3 (R2 = −0.76; Figure 7), suggest significant clastic contamination during deposition [19,56].
The degree of mixing between hydrothermal metal deposits and detrital material can be evaluated using the Al/(Al + Mn + Fe + Na + K + Ca) ratio and Fe/Ti ratio [46] (Figure 8a). The Hengyang NIF plot along a mixing trend between the end-members of pure hydrothermal precipitates (submarine hydrothermal sediments, BHS) and the upper continental crust (UCC) [57,58]. This indicates that hydrothermal contributions increase upward from the lower to upper stratigraphic units, while clastic input diminishes, with the uppermost layers showing the strongest hydrothermal influence. The Fulu NIF is interpreted to have been more strongly influenced by detrital material input and was likely deposited in a shelf environment [19]. In contrast, the Xinyu NIF and Hengyang NIF exhibit geochemical signatures indicative of a greater hydrothermal contribution, suggesting deposition in a deeper basin setting [20,21].
Whole-rock REE compositions of the IFs are generally considered relatively robust against diagenetic modification, thus reflecting the characteristics of the source-region REY distribution [3,52,59]. High-temperature hydrothermal fluids (>350 °C) typically exhibit pronounced Eu positive anomalies [47,60], whereas low-temperature hydrothermal fluids (T < 200 °C) lack significant Eu anomalies [47]. The absence of Eu anomalies in Hengyang NIF (Eu/Eu* = 0.95–1.11) indicates that high-temperature hydrothermal fluids were not the primary source of the material, which aligns with characteristics observed in NIFs from other regions globally [20,61,62].
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Figure 8. Source discrimination diagram for sedimentary materials (using whole-rock data). (a) Plot of Fe/Ti vs. Al/(Al + Fe + Mn + Na + K + Ca) for the Hengyang NIF samples (modified from Cox et al., 2013 [46]); (b) Zr-Y/Ho discrimination diagram, from Bolhar et al., (2004) [62]; (c) The component mixing lines of Y/Ho vs. Sm/Yb; (d) The component mixing lines of Sm/Yb vs. Eu/Sm (after Alexander et al., 2008) [63]. SW: modern seawater from Bolhar et al., (2004) [62]; HTH: high-T Hydrothermal from fluids Bau and Dulski, (1999) [47]; LTH: low-T Hydrothermal fluids from Michard et al., (1993) [48]; BHS: Submarine hydrothermal sediments from Marchig and Gundlach, (1982) [57]; E-MORB: enriched mid-ocean ridge basalt, OIB: ocean island basalt, N-MORB: normal MORB from Taylor and Mclennan, (2009) [64]; UCC: upper continental crust from Cribb and Barton, (1996) [58]; S-type granite from Whalen et al., (1987) [65].
Figure 8. Source discrimination diagram for sedimentary materials (using whole-rock data). (a) Plot of Fe/Ti vs. Al/(Al + Fe + Mn + Na + K + Ca) for the Hengyang NIF samples (modified from Cox et al., 2013 [46]); (b) Zr-Y/Ho discrimination diagram, from Bolhar et al., (2004) [62]; (c) The component mixing lines of Y/Ho vs. Sm/Yb; (d) The component mixing lines of Sm/Yb vs. Eu/Sm (after Alexander et al., 2008) [63]. SW: modern seawater from Bolhar et al., (2004) [62]; HTH: high-T Hydrothermal from fluids Bau and Dulski, (1999) [47]; LTH: low-T Hydrothermal fluids from Michard et al., (1993) [48]; BHS: Submarine hydrothermal sediments from Marchig and Gundlach, (1982) [57]; E-MORB: enriched mid-ocean ridge basalt, OIB: ocean island basalt, N-MORB: normal MORB from Taylor and Mclennan, (2009) [64]; UCC: upper continental crust from Cribb and Barton, (1996) [58]; S-type granite from Whalen et al., (1987) [65].
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Y3+ and Ho3+, which have similar ionic radii, are expected to exhibit similar geochemical behavior. However, differences in surface complexation capacities result in Ho precipitating from seawater approximately twice as rapidly as Y, making the Y/Ho ratio a useful diagnostic tool for distinguishing marine from non-marine depositional environments [66]. Modern seawater exhibits Y/Ho ratios ranging from 44 to 74 [51], while chondritic meteorites show Y/Ho ratios of 26–28. Upper crustal rocks and terrigenous sediments display Y/Ho ratios similar to those of chondritic values [67]. The Hengyang NIF has an average whole-rock Y/Ho ratio of 27.8 (Figure 8b). Zr is commonly enriched in detrital minerals such as zircon and therefore serves as a proxy for terrigenous input. In the Zr–Y/Ho diagram (Figure 8b), samples with higher Zr concentrations tend to reflect stronger detrital contributions, which may overprint the primary seawater Y/Ho signal. The Hengyang NIF shows an average whole-rock Y/Ho ratio of 27.8, close to chondritic values, suggesting that detrital input has partially overprinted the original marine signature.
The component mixing model allows for a more precise evaluation of the relative contributions of seawater and hydrothermal fluids in the NIFs. In the Sm/Yb vs. Y/Ho diagram (Figure 8c), the samples plot along a mixing line between low-temperature hydrothermal fluids and the upper continental crust. In the Sm/Yb vs. Eu/Sm diagram (Figure 8d), the samples align with the mixing trend between low-temperature hydrothermal fluids and seawater. These results highlight that low-temperature hydrothermal fluids and seawater were key contributors to the Hengyang NIF, with substantial detrital input masking primary marine geochemical signals.

5.2. The Genesis and REY Feature of Apatite

Before explaining the geochemical indicators of apatite, it is necessary to assess the potential influences of detrital contamination, diagenesis, and metamorphism. For bulk-rock samples, the lack of correlation between Al2O3 and P2O5 contents indicates a limited contribution from detrital phosphorus overall (Figure 9), despite the presence of aluminosilicate detrital components in the studied IFs [68]. Apatite exhibits strong resistance to weathering and transport, making it a robust indicator for provenance analysis, particularly in settings with complex source evolution [68,69]. However, micro-scale analyses reveal distinct origins and alteration histories for apatite across different stratigraphic units.
In the upper stratum unit, apatite grains are relatively scarce and exhibit a smooth, subrounded morphology indicative of transport abrasion [70,71]. Geochemically, these grains display LREE-enriched patterns with diverse shapes and variable Eu anomalies (Figure 6a), which are characteristic of terrigenous detrital input [72]. Furthermore, weathering processes are known to elevate (Y/Y*)N and (La/Nd)N ratios, resulting in a positive correlation between these parameters [73]. The apatite in the upper stratum exhibits exactly this correlation (Figure 10), confirming that its REY distribution has been significantly affected by weathering. Therefore, the upper-layer apatite is predominantly of detrital origin and unsuitable for marine paleoenvironmental reconstruction.
In stark contrast, apatite in the bottom and intermediate stratum units is much more abundant and displays fundamentally different characteristics. These grains lack any evidence of transport-related abrasion; instead, they typically contain abundant inclusions and pervasive alteration voids. Based on our EPMA quantitative analyses and the F-Cl-OH ternary diagram (Figure 5), the apatite from these lower and intermediate units is classified as highly pure fluorapatite. The combination of these textural features (alteration voids) and chemical purity provides direct petrographic evidence for intense fluid–rock interactions during diagenesis. It is interpreted that metastable primary marine phosphate precipitates underwent extensive diagenetic recrystallization and decarbonation driven by pore fluids, ultimately stabilizing as authigenic fluorapatite [74,75]. Given the low regional metamorphic grade (lower greenschist facies) of the Hengyang region, the compositions and textures of these apatite grains are considered to preserve the geochemical signals of diagenetic pore waters rather than subsequent metamorphic overprints.
Apatite from the bottom stratum unit shows HREE enrichment, whereas apatite from the intermediate stratum unit exhibits pronounced MREE enrichment. This MREE enrichment is strongly indicative of diagenetic processes controlled by pore water chemistry, rather than high-grade metamorphism. Similar MREE-enriched patterns are well-documented in modern deep-sea pelagic sediments, where primary rare earth-bearing phases include biogenic apatite (e.g., fish bones and teeth), Fe–Mn (oxyhydr)oxides, and sediment pore water [76,77]. Phosphorites with anomalous rare earth enrichment commonly display MREE-enriched distribution patterns, while Fe–Mn (oxyhydr)oxides often exhibit a characteristic MREE bulge pattern [78]. Current research indicates that the adsorption of rare earth elements from seawater by Fe–Mn (oxyhydr)oxides, followed by their subsequent release into pore water during reductive dissolution in diagenesis, is a key process controlling REY enrichment in authigenic apatite within marine sediments [59,79,80,81]. Therefore, the MREE enrichment feature of the middle-layer apatite in the Hengyang NIF is attributed to diagenetic pore water interactions (the “Fe-Mn pump” mechanism) rather than primary seawater precipitation. In contrast, the HREE enrichment observed in the bottom stratum unit is likely due to a relatively shorter diagenetic duration or weaker fluid–rock interactions, allowing it to preserve more of the residual characteristics of primary seawater or early sediments.
The extent of these diagenetic modifications can be quantitatively assessed using (La/Yb)N and (La/Sm)N ratios (Figure 11a). The data from the bottom and intermediate stratum units plot along a distinct trajectory indicative of apatite recrystallization during late diagenesis [82,83]. Furthermore, Morad and Felitsyn (2001) [84] demonstrated that during diagenesis-induced MREE (particularly Nd) enrichment, Ce anomalies often exhibit a positive correlation with (La/Sm)N values. Our samples display this exact strong positive correlation (Figure 11b), confirming that the REY redistribution occurred during diagenetic recrystallization, thereby rendering the Ce/Ce* unreliable for paleo-marine environmental reconstruction. Crucially, this chemical evidence perfectly aligns with our petrographic observations (Figure 6); the abundant alteration voids and irregular surfaces of the intermediate and bottom stratum apatite serve as direct physical evidence of this fluid-related diagenetic dissolution and recrystallization. This explains why the apatite in these specific layers differs morphologically from typical unaltered sedimentary authigenic apatite, which generally exhibits smoother surfaces and better-developed crystal forms [68].

5.3. Response to Diagenetic Processes from Apatite and Whole-Rock Geochemistry

From the lower to the upper stratum units of the Hengyang NIF, whole-rock analyses reveal a continuous decrease in the Th/U ratio. Under reducing conditions, soluble U6+ is reduced to immobile U4+ and becomes enriched in sediments, whereas Th remains relatively immobile. Therefore, decreasing Th/U ratios reflect increasingly reducing depositional conditions [85,86]. Concurrent upward increases in V/(V + Ni) and V/Cr ratios further support a transition toward more reducing conditions, as V is preferentially enriched relative to Ni and Cr as conditions become progressively less oxygenated (or more restricted) [87,88] (Figure 12).
The Th/U ratios in apatite separated from different stratigraphic layers also display a decreasing trend from lower to upper strata, which is fully consistent with the whole-rock data. However, the V/(V + Ni) and V/Cr ratios in apatite deviate from the whole-rock trends (Figure 13). During diagenesis, V, Cr, and Ni exhibit markedly different partition coefficients for magnetite, with V strongly favoring adsorption (V: 6.6 ± 3.8, Cr: 1.2 ± 1.1, Ni: 3.0 ± 1.5) [89]. Magnetite preferentially incorporates V4+ via inner-sphere complexation and lattice substitution, significantly reducing pore water V concentrations. The resulting relative increases in V/Cr and V/(V + Ni) are inherited by crystallizing apatite, preventing these ratios from reflecting whole-rock redox trends. Magnetite recrystallization during diagenesis alters pore water trace element compositions relative to the original depositional environment (Figure 13).
The Th/U ratio in apatite can serve as a reliable indicator of sedimentary redox conditions [86,90]. Its consistency between whole-rock and apatite samples arises from the geochemical “inertness” of Th and U during diagenesis. Both whole-rock and apatite samples show initial U enrichment during the depositional stage—under anoxic bottom-water conditions, soluble U6+ is reduced to insoluble U4+ and incorporated into sediments. This process establishes the system’s characteristic low initial Th/U ratio. Consequently, even during apatite crystallization in diagenetic environments, the Th/U ratio in surrounding pore fluids remains predominantly controlled by the inherent Th/U ratio of the sediments themselves. Thus, the Th/U ratio fundamentally reflects the macro-scale redox state at the water–sediment interface during deposition (Figure 14). Once incorporated into sediments, and subsequently into the apatite crystal lattice, both Th and U exhibit low mobility under the low-temperature diagenetic to very low-grade metamorphic conditions evidenced in our samples [91].
The Ce anomaly (Ce/Ce*) in IFs is widely used to infer the redox conditions of the precipitating water column. Oxygenated water columns typically show a negative Ce anomaly, as the oxidation of Ce3+ to Ce4+ in aqueous solutions decreases Ce solubility, leading to the preferential association of Ce4+ with iron–manganese oxyhydroxides and organic matter for transport [92,93]. Bau and Dulski (1996) [59] proposed calculating the praseodymium anomaly (Pr/Pr*) = 2Pr/(Ce + Nd) (normalized to post-Archean Australian shale, PAAS). A true negative Ce anomaly resulting from Ce’s chemical behavior leads to (Pr/Pr*) > 1, while a positive Ce anomaly results in (Pr/Pr*) < 1. In this study, (Pr/Pr*) is slightly greater than 1, and the whole-rock Ce anomaly in the Hengyang NIF is slightly less than 1 (Figure 15). Since the oxidation potential of Ce3+/Ce4+ is higher than that of Fe2+/Fe3+ [94], the oxygen content in the water column where the iron formation at the base of the Fulu Formation was deposited could oxidize Fe2+ to Fe3+, but not Ce3+ to Ce4+. This suggests that the Hengyang NIF were deposited in a weakly oxidizing environment.
Most of the apatite from the bottom and intermediate stratum units of the Hengyang NIFs shows MREE enrichment, influenced by both differential REE partitioning in hydroxyl apatite and diagenetic processes [78]. The apatite samples from the Hengyang NIF show an even stronger lg(Ce/Ce*)-(La/Sm)N positive correlation, confirming that the Ce anomalies in these samples reflect diagenetic overprinting rather than original depositional conditions (Figure 11b and Figure 15).
Trace elements in apatite are predominantly controlled by diagenetic pore water interactions [77,78,95]. The Th/U ratio’s unique properties allow it to co-vary between whole-rock and apatite, reliably recording depositional redox trends. Whole-rock V/Cr and V/(V + Ni) ratios can also indicate redox conditions, but their corresponding ratios in apatite are distorted (Figure 13). Ce anomalies, particularly in apatite, require further scrutiny, as magnetite recrystallization during diagenesis renders them unsuitable for paleo-environmental reconstruction. In summary, the Hengyang NIF records a depositional evolution from oxidizing to weakly oxidizing conditions from lower to upper strata. However, diagenetic processes varied across layers, with increasing diagenetic duration significantly altering apatite trace element compositions. This makes apatite Ce anomalies, V/Cr, and V/(V + Ni) unreliable for environmental reconstruction, whereas the Th/U ratio—due to its inherent stability—faithfully preserves redox signatures in both whole-rock and apatite (Figure 14).
During the diagenetic process of IFs, significant differences in partition coefficients exist between various elements in pore water and solid phases, such as magnetite and apatite. Prolonged diagenesis exacerbates these differences, leading to systematic variations in elemental ratios. At the sediment–water interface, FeOOH adsorbs HPO42− and REE3+ from seawater. As sediment burial depth increases, oxidation weakens progressively, and FeOOH undergoes reductive dissolution, releasing Fe2+, HPO42−, and REE3+. This increases the concentrations of these ions in pore water, promoting the precipitation of authigenic apatite. In the suboxic zone, Fe2+ produced by FeOOH reduction combines with residual FeOOH to form magnetite; deeper in the anoxic zone, it co-precipitates with sulfides as FeS; in the oxic zone, it is reoxidized to FeOOH. Apatite formed in the suboxic zone inherits the chemical signatures of the pore water (Figure 16).
As observed in apatite, the Ce anomaly, V/Cr, V/(V + Ni), and (La/Sm)N ratios exhibit synchronous changes, with samples 24-MC-5 and 24-MC-6 showing a distinct decreasing trend (Figure 13). This suggests that these samples experienced a prolonged diagenetic process, leading to more pronounced modifications to their elemental ratios (Figure 16). Microscopic examination further reveals that magnetite crystals in samples 24-MC-5 and 24-MC-6 are larger and better crystallized, whereas those in samples 24-MC-7, 8, and 9 are smaller. Over time, the sedimentary environment transitioned from moderately oxidizing to weakly oxidizing conditions, with the intermediate diagenetic phase being the most prolonged, corresponding to a lower-energy hydrodynamic setting. The decoupling of whole-rock and apatite geochemical signals reaches its maximum in the intermediate stratum unit (e.g., samples 24-MC-5, 6). This decoupling, alongside the distinct MREE enrichment and modified Ce anomalies, robustly supports a depositional model where this middle stage corresponds to a stagnant, low-energy sub-basin environment with slow sedimentation, allowing for prolonged diagenetic fluid–mineral interactions.

6. Conclusions

(1)
The absence of Eu anomalies, relative enrichment in HREE, and near-chondritic Y/Ho ratios (~27.8) in the Hengyang NIF collectively constrain low-temperature hydrothermal fluids and seawater as a critical contributor to the Hengyang NIF. However, significant detrital input likely masked the original seawater-derived Y/Ho signature.
(2)
Apatite from the bottom and intermediate stratigraphic units of the Hengyang NIF is interpreted to be primarily syndepositional. Its variable REE patterns, including HREE enrichment and pronounced MREE enrichment, were mainly controlled by pore water chemistry during diagenesis. These patterns likely reflect processes such as the dissolution of Fe–Mn hydroxides, releasing MREEs, and the degradation of organic matter. The positive correlation between Ce anomalies and (La/Sm)N in apatite further indicates strong diagenetic modification.
(3)
Trace element characteristics of both whole-rock and apatite collectively record an evolution from relatively oxic to weakly reducing conditions from the base to the top of the deposit. Systematic decreases in Th/U ratios in both whole-rock and apatite reliably indicate a reduction in bottom-water oxygenation. However, Ce anomalies, V/Cr, and V/(V + Ni) ratios in apatite, influenced by magnetite recrystallization and pore water interactions during diagenesis, cannot be directly used for paleoenvironmental reconstruction.

Supplementary Materials

The following supporting information can be downloaded at https://www.mdpi.com/article/10.3390/min16040392/s1, Table S1: Major elements compositions (wt%) of the apatite in the Hengyang NIF; Table S2: Trace elements compositions (ppm) of the apatite in the Hengyang NIF.

Author Contributions

Conceptualization, C.Z. and L.L.; formal analysis, L.L.; investigation, C.Z., K.H. and L.L.; resources, K.H. and L.L.; data curation, C.Z.; writing—original draft preparation, C.Z. and L.L.; writing—review and editing, C.Z., T.H. and K.H.; visualization, C.Z. and L.L.; funding acquisition, L.L. and K.H. All authors have read and agreed to the published version of the manuscript.

Funding

This research was supported by the National Natural Science Foundation of China (No. 41972198), the National Natural Science Foundation of Hunan (Nos. 2022JJ30702, 2025JJ80062, 2026JJ50447), the Exploration Program (Exploration of the Miaochong Iron Deposit in Hengyang City, Hunan Province) and the Open Research Fund Program of Hunan Mine Carbon Sequestration and Sink Enhancement Engineering Technology Research Center (No. 2025KSGTZH05).

Data Availability Statement

All data supporting the findings of this study are provided in the tables within the main text of the article and Supplementary Materials.

Acknowledgments

We appreciate anonymous reviewers for their critical and constructive comments and suggestions.

Conflicts of Interest

The authors declare no conflicts of interest.

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Figure 1. (a) Distribution of NIFs in the Nanhua Basin of South China [19]; (b) Simplified geological map for Hengyang NIF in Hunan Province. Samples ZK3701-1 and ZK3701-2 were obtained from the upper unit of the iron formation via borehole ZK3701.
Figure 1. (a) Distribution of NIFs in the Nanhua Basin of South China [19]; (b) Simplified geological map for Hengyang NIF in Hunan Province. Samples ZK3701-1 and ZK3701-2 were obtained from the upper unit of the iron formation via borehole ZK3701.
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Figure 2. Correlation of representative Neoproterozoic strata in the middle and southeast portions of the Yangtze Block (modified from Wang et al. (2025) [21]; Ye et al (2024) [20]). Abbreviations: NT—Nantuo Formation, DTP—Datangpo Formation, GC—Gucheng Formation, Fulu—Fulu Formation, CA—Changan Formation.
Figure 2. Correlation of representative Neoproterozoic strata in the middle and southeast portions of the Yangtze Block (modified from Wang et al. (2025) [21]; Ye et al (2024) [20]). Abbreviations: NT—Nantuo Formation, DTP—Datangpo Formation, GC—Gucheng Formation, Fulu—Fulu Formation, CA—Changan Formation.
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Figure 3. Photographs of typical iron formation from the Hengyang NIF. (a) Contact relationship between granite dyke and Hengyang NIF. (b) Field photograph showing the iron formation. (c) Hand specimen photograph. (d–f) Photomicrographs of representative samples. Mag: magnetite.
Figure 3. Photographs of typical iron formation from the Hengyang NIF. (a) Contact relationship between granite dyke and Hengyang NIF. (b) Field photograph showing the iron formation. (c) Hand specimen photograph. (d–f) Photomicrographs of representative samples. Mag: magnetite.
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Figure 4. PAAS-normalized REY pattern of the Hengyang NIF samples compared with the average compositions of hydrothermal fluids, modern seawater, and other well-studied NIFs. The vertical gray shaded areas highlight the Eu and Y anomalies, respectively. Data sources: modern seawater and high-T Hydrothermal fluids from Bau and Dulski (1999) [47]; low-T Hydrothermal fluids from Michard et al. (1993) [48]; Rapitan BIF form Klein and Beukes (1993) [49]; Fulu IF from Wang et al. (2025) and Li et al. (2014) [21,50]; Xinyu IF from Wu et al. (2022) [19].
Figure 4. PAAS-normalized REY pattern of the Hengyang NIF samples compared with the average compositions of hydrothermal fluids, modern seawater, and other well-studied NIFs. The vertical gray shaded areas highlight the Eu and Y anomalies, respectively. Data sources: modern seawater and high-T Hydrothermal fluids from Bau and Dulski (1999) [47]; low-T Hydrothermal fluids from Michard et al. (1993) [48]; Rapitan BIF form Klein and Beukes (1993) [49]; Fulu IF from Wang et al. (2025) and Li et al. (2014) [21,50]; Xinyu IF from Wu et al. (2022) [19].
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Figure 5. Ternary plot of halogen and hydroxyl mole fractions (F-Cl-OH) for apatite from the Hengyang NIF. The diagram on the right provides a magnified view of the F-rich corner. Apatite from the bottom and intermediate units converges toward the pure fluorapatite end-member. In contrast, apatite in the upper stratigraphic units shows a more scattered distribution.
Figure 5. Ternary plot of halogen and hydroxyl mole fractions (F-Cl-OH) for apatite from the Hengyang NIF. The diagram on the right provides a magnified view of the F-rich corner. Apatite from the bottom and intermediate units converges toward the pure fluorapatite end-member. In contrast, apatite in the upper stratigraphic units shows a more scattered distribution.
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Figure 6. Backscattered scanning electron microscope images and PAAS-normalized REY patterns of apatite in the Hengyang NIF. (a) BSE images of representative grains from samples ZK3701-2 and ZK3701-1, and corresponding REY patterns of apatite in the upper Hengyang NIFs; (b) BSE images of representative grains from samples 24-MC-4, 24-MC-5, and 24-MC-6, and corresponding REY patterns of apatite in the intermediate Hengyang NIFs; (c) BSE images of representative grains from samples 24-MC-7, 24-MC-8, and 24-MC-9, and corresponding REY patterns of apatite in the bottom Hengyang NIFs. Ap: apatite. Data sources: PAAS-normalization refers to McLennan (1989) [51].
Figure 6. Backscattered scanning electron microscope images and PAAS-normalized REY patterns of apatite in the Hengyang NIF. (a) BSE images of representative grains from samples ZK3701-2 and ZK3701-1, and corresponding REY patterns of apatite in the upper Hengyang NIFs; (b) BSE images of representative grains from samples 24-MC-4, 24-MC-5, and 24-MC-6, and corresponding REY patterns of apatite in the intermediate Hengyang NIFs; (c) BSE images of representative grains from samples 24-MC-7, 24-MC-8, and 24-MC-9, and corresponding REY patterns of apatite in the bottom Hengyang NIFs. Ap: apatite. Data sources: PAAS-normalization refers to McLennan (1989) [51].
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Figure 7. Correlations of Al2O3 with TFe2O3 (a) and TiO2 (b) for Hengyang NIFs compared with other NIFs of the Nahua Basin. Data for Xinyu NIFs are from Wang et al. (2025) and Li et al. (2014) [21,50], whereas data for Fulu NIFs are from Wu et al. (2022) [19].
Figure 7. Correlations of Al2O3 with TFe2O3 (a) and TiO2 (b) for Hengyang NIFs compared with other NIFs of the Nahua Basin. Data for Xinyu NIFs are from Wang et al. (2025) and Li et al. (2014) [21,50], whereas data for Fulu NIFs are from Wu et al. (2022) [19].
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Figure 9. Relationships between bulk-rock Al2O3 (wt.%) contents and bulk-rock P2O5 contents (wt.%).
Figure 9. Relationships between bulk-rock Al2O3 (wt.%) contents and bulk-rock P2O5 contents (wt.%).
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Figure 10. Relationships between (Y/Y*)N and (La/Nd)N in apatite in Hengyang NIF. (a) Plot for samples ZK3701-2 and ZK3701-1, showing a positive correlation (R2 = 0.49); (b) Plot for samples 24-MC-4, 24-MC-5, and 24-MC-6, displaying no linear correlation (R2 = 0.001); (c) Plot for samples 24-MC-7, 24-MC-8, and 24-MC-9, indicating no linear correlation (R2 = 0.02).
Figure 10. Relationships between (Y/Y*)N and (La/Nd)N in apatite in Hengyang NIF. (a) Plot for samples ZK3701-2 and ZK3701-1, showing a positive correlation (R2 = 0.49); (b) Plot for samples 24-MC-4, 24-MC-5, and 24-MC-6, displaying no linear correlation (R2 = 0.001); (c) Plot for samples 24-MC-7, 24-MC-8, and 24-MC-9, indicating no linear correlation (R2 = 0.02).
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Figure 11. Correlations of PAAS-normalized (La/Sm)N with (La/Yb)N (a) and (Ce/Ce*)N (b) for apatite in Hengyang NIF. (a) after Gadd et al., 2016 [82]. The modern seawater field and trajectories for early diagenesis (adsorption), late diagenesis (recrystallization), and substitution are from Reynard et al. 1999 [83]; (b) after Morad and Felitsyn, 2001 [84].
Figure 11. Correlations of PAAS-normalized (La/Sm)N with (La/Yb)N (a) and (Ce/Ce*)N (b) for apatite in Hengyang NIF. (a) after Gadd et al., 2016 [82]. The modern seawater field and trajectories for early diagenesis (adsorption), late diagenesis (recrystallization), and substitution are from Reynard et al. 1999 [83]; (b) after Morad and Felitsyn, 2001 [84].
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Figure 12. Profiles of Th/U, V/(V + Ni), and V/Cr values of the Hengyang NIF through iron formation sections. Profiles of Th/U (left, gray), V/(V+Ni) (middle, gold), and V/Cr (right, pink).
Figure 12. Profiles of Th/U, V/(V + Ni), and V/Cr values of the Hengyang NIF through iron formation sections. Profiles of Th/U (left, gray), V/(V+Ni) (middle, gold), and V/Cr (right, pink).
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Figure 13. Box plots of geochemical parameters in sedimentary-origin apatite from different samples of the Hengyang NIF: (a) (La/Sm)N; (b) V/(V + Ni); (c) Ce/Ce*; and (d) V/Cr. Boxes denote the interquartile range (IQR = Q3 − Q1), encompassing the central 50% of the data. Outliers are represented by diamonds.
Figure 13. Box plots of geochemical parameters in sedimentary-origin apatite from different samples of the Hengyang NIF: (a) (La/Sm)N; (b) V/(V + Ni); (c) Ce/Ce*; and (d) V/Cr. Boxes denote the interquartile range (IQR = Q3 − Q1), encompassing the central 50% of the data. Outliers are represented by diamonds.
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Figure 14. Box plots of geochemical parameters in sedimentary-origin apatite from different samples of the Hengyang NIF: (a) Th/U; (b) Y/Ho. Boxes denote the interquartile range (IQR = Q3 − Q1), encompassing the central 50% of the data. Outliers are represented by diamonds.
Figure 14. Box plots of geochemical parameters in sedimentary-origin apatite from different samples of the Hengyang NIF: (a) Th/U; (b) Y/Ho. Boxes denote the interquartile range (IQR = Q3 − Q1), encompassing the central 50% of the data. Outliers are represented by diamonds.
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Figure 15. Composition of the Hengyang NIF and apatite plotted on PAAS-normalized (Pr/Pr*)N vs. (Ce/Ce*)N diagram (after Bau and Dulski, 1996 [59]).
Figure 15. Composition of the Hengyang NIF and apatite plotted on PAAS-normalized (Pr/Pr*)N vs. (Ce/Ce*)N diagram (after Bau and Dulski, 1996 [59]).
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Figure 16. The formation processes of apatite and magnetite (modified from Haley, 2004 and Pufahl (2017) [77,96]). The boxes on the left illustrate the REE patterns in the water column under different redox conditions (Oxic, Suboxic, and Anoxic). Solid arrows represent direct precipitation or chemical transformation pathways, while dashed arrows indicate migration, adsorption, or reductive dissolution processes. In the suboxic zone, the reductive dissolution of FeOOH-rich particles releases Fe2+, HPO42−, and REE3+, leading to the formation of authigenic apatite (Ap) and magnetite (Mag).
Figure 16. The formation processes of apatite and magnetite (modified from Haley, 2004 and Pufahl (2017) [77,96]). The boxes on the left illustrate the REE patterns in the water column under different redox conditions (Oxic, Suboxic, and Anoxic). Solid arrows represent direct precipitation or chemical transformation pathways, while dashed arrows indicate migration, adsorption, or reductive dissolution processes. In the suboxic zone, the reductive dissolution of FeOOH-rich particles releases Fe2+, HPO42−, and REE3+, leading to the formation of authigenic apatite (Ap) and magnetite (Mag).
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Table 1. Major (wt%) and trace (ppm) element compositions of the Hengyang NIF.
Table 1. Major (wt%) and trace (ppm) element compositions of the Hengyang NIF.
Upper Stratum UnitIntermediate Stratum UnitBottom Stratum Unit
ElementsZK3701-1ZK3701-224-MC-424-MC-524-MC-624-MC-724-MC-824-MC-9
SiO229.9329.2351.6343.0243.8366.4472.6756.35
TiO20.190.300.270.240.400.660.360.50
Al2O33.882.124.845.736.819.557.026.86
TFe2O360.9265.6836.1744.7541.7111.9813.4228.56
MnO0.030.040.090.070.140.140.090.09
MgO2.061.613.502.473.212.462.122.10
CaO0.880.331.110.911.272.361.470.90
Na2O0.070.020.170.430.340.020.040.04
K2O0.570.281.321.281.792.041.631.64
P2O50.600.120.300.290.160.680.150.59
LOI0.720.510.96−0.18−0.013.351.391.54
Li60.2172112.743.448.72927.6
Be0.970.380.791.421.292.361.621.52
Sc65.910.46.17.316.89.413.6
V92104848487645085
Cr406060509016090170
Co4.93.76.26.17.95.53.13.2
Ni94.710.817.822.713.612.29.8
Cu0.40.41.41.40.82.20.91.1
Zn5463665385523852
Ga64.96.98.49.313.499.8
Ge0.280.270.230.280.350.220.160.25
As0.80.3<0.2<0.2<0.20.40.21.1
Rb20.911.946.1476110359.160.6
Sr57.921.663.372.5191.573.689.8124
Y35.325.921.319.824.95116.842.7
Zr52701024972276100218
Nb4.67.16.45.36.118.67.810.6
Mo0.450.480.610.720.630.390.490.52
Ag<0.010.010.060.030.020.020.01<0.01
Cd<0.020.020.030.020.030.02<0.020.06
In0.0360.030.030.0350.0470.0620.0330.054
Sn11.41.32.11.720.91.3
Sb12.57.730.480.633.791.292.42.27
Cs9.422.517.036.165.287.694.983.11
Ba38661.243835535011607371825
La16.811.312.61717.436.61234.6
Ce36.126.627.737.936.58223.777.5
Pr4.463.473.54.674.469.772.799.19
Nd18.614.51418.517.839.110.636.5
Sm4.153.633.113.874.167.942.037.76
Eu1.050.870.680.760.871.650.521.57
Gd4.784.093.33.594.297.672.027.62
Tb0.80.730.540.590.751.230.341.25
Dy5.174.43.373.464.487.612.237.31
Ho1.231.010.780.740.991.830.571.65
Er3.642.942.342.032.785.691.954.68
Tm0.560.460.360.290.40.910.330.72
Yb3.462.782.361.632.375.782.284.35
Lu0.580.460.390.260.371.010.410.72
Hf1.322.21.32.25.93.13.8
Ta0.250.310.340.320.460.790.430.42
W2.924.21.90.720.60.9
Tl0.080.050.180.220.220.430.210.21
Pb99.724.312.415.41415.76.1
Bi0.160.160.250.080.130.090.120.16
Th2.423.033.493.315.489.794.896.64
U0.420.530.380.420.650.610.470.57
Note: data for samples 24-MC-4, 24-MC-7, and 24-MC-9 are derived from Huang et al., 2025 [45].
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Zhang, C.; Liu, L.; Huang, K.; Hu, T. Apatite as an Indicator of Sedimentary Environment and Diagenesis for the Hengyang Neoproterozoic Iron Formation, South China. Minerals 2026, 16, 392. https://doi.org/10.3390/min16040392

AMA Style

Zhang C, Liu L, Huang K, Hu T. Apatite as an Indicator of Sedimentary Environment and Diagenesis for the Hengyang Neoproterozoic Iron Formation, South China. Minerals. 2026; 16(4):392. https://doi.org/10.3390/min16040392

Chicago/Turabian Style

Zhang, Chuangye, Lei Liu, Kuanxin Huang, and Tianyang Hu. 2026. "Apatite as an Indicator of Sedimentary Environment and Diagenesis for the Hengyang Neoproterozoic Iron Formation, South China" Minerals 16, no. 4: 392. https://doi.org/10.3390/min16040392

APA Style

Zhang, C., Liu, L., Huang, K., & Hu, T. (2026). Apatite as an Indicator of Sedimentary Environment and Diagenesis for the Hengyang Neoproterozoic Iron Formation, South China. Minerals, 16(4), 392. https://doi.org/10.3390/min16040392

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