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Article

Mineral Chemistry and Iron Isotope Characteristics of Magnetites in Pertek Fe-Skarn Deposit (Türkiye)

by
Hatice Kara
1,*,
Cihan Yalçın
2,
Mehmet Ali Ertürk
1 and
Leyla Kalender
1
1
Department of Geological Engineering, Fırat University, 23119 Elazığ, Türkiye
2
SRG Engineering and Consultancy Ltd. Sti, 23100 Denizli, Türkiye
*
Author to whom correspondence should be addressed.
Minerals 2025, 15(4), 369; https://doi.org/10.3390/min15040369
Submission received: 18 February 2025 / Revised: 26 March 2025 / Accepted: 28 March 2025 / Published: 1 April 2025
(This article belongs to the Section Mineral Deposits)
Browse Figures
Versions Notes

Abstract

:
This study investigates the mineral chemistry and iron isotope composition of the Pertek Fe-skarn deposit in the Eastern Taurides, Turkey, to elucidate skarn formation and ore genesis through chemical and isotopic parameters. The deposit consists of substantial and dispersed magnetite ores formed by the intrusion of a dioritic suite into marbles. Mineral assemblages, including hematite, goethite, andradite garnet, hedenbergite pyroxene, calcite, and quartz, exhibit compositional variations at different depths within the ore body. Magnetite is commonly associated with hematite, goethite, garnet, pyroxene, calcite, and quartz. Extensive LA–ICP–MS analysis of magnetite chemistry reveals elevated trace element concentrations of titanium (Ti), aluminum (Al), vanadium (V), and magnesium (Mg), distinguishing Pertek magnetite from low-temperature hydrothermal deposits. The enrichment of Ti (>300 ppm) and V (>200 ppm), along with the presence of Al and Mg, suggests formation from high-temperature hydrothermal fluids exceeding 300 °C. Discriminant diagrams, such as Al+Mn versus Ti+V, classify Pertek magnetite within the skarn deposit domain, affirming its medium- to high-temperature hydrothermal origin (200–500 °C), characteristic of skarn-type deposits. Magnetite thermometry calculations yield an average formation temperature of 414.53 °C. Geochemical classification diagrams, including Ni/(Cr+Mn) versus Ti+V and TiO2-Al2O3-MgO+MnO, further support the skarn-type genesis of the deposit, distinguishing Pertek magnetite from other iron oxide deposits. The Fe-skarn ore samples display low total REE concentrations, variable Eu anomalies, enrichment in LREEs, and depletion in HREEs, consistent with fluid–rock interactions in a magmatic–hydrothermal system. The δ56Fe values of magnetite range from 0.272‰ to 0.361‰, while the calculated δ56Fe_aq values (0.479‰ to 0.568‰) suggest a magmatic–hydrothermal origin. The δ57Fe values (0.419‰ to 0.530‰) and the calculated 103lnβ value of 0.006397 indicate re-equilibration of the magmatic–hydrothermal fluid during ore formation.

1. Introduction

Skarn deposits are one of the most abundant ore types in the Earth’s crust and form in rocks of nearly any age. These deposits, which can contain a variety of metals, including Fe, Cu, Zn, W, Sn, Mo, and Au, and subordinate rare earth elements (REEs), U, F, B, Bi, Be, and Co, form when magma intrudes and interacts with carbonate rocks. Magnetite is one of the primary ore minerals found in many skarn deposit types [1]. The iron skarn deposits are mined because of their high magnetite content, which primarily consists of calcsilicate contact metasomatic rocks [1]. Calcic magnetite skarns are associated to comparatively mafic intrusions, and in the prograde stage, they are represented by garnet (grossular-andradite) and pyroxene (ferrosilite), while at the retrograde stage, they are represented by amphibole, chlorite, and epidote. However, amphibole, serpentine, phlogopite, talc, and chlorite are found in the retrograde stage, and forsterite, spinel, and diopside are found in the prograde stage of magnesian magnetite skarns, which are linked to felsic plutons in dolomitic host rocks [2,3].
The Tethyan metallogenic belt, spanning from the Western Mediterranean to the Chinese Peninsula, is among the world’s most important mineral belts, with various deposits associated with post-collisional magmatism [4]. Pertek Fe-skarn deposit located at East Türkiye is part of the Tethyan Metallogenic Belt (Figure 1a) or the western extension of the Tethyan Eurasian Orogenic Belt [5,6]. The Pertek Fe-skarn represents a notable iron ore deposit situated within the Keban Metamorphic marbles [7]. The study area comprises substantial and dispersed magnetite ores resulting from the intrusion of a dioritic suite in contact with marbles. Previous mineralogical studies of the Pertek Fe-skarn deposit have identified a variety of skarn assemblages, including garnet, pyroxene, epidote, and amphibole. However, no detailed investigations using mineral chemistry and stable isotope studies have been conducted to better constrain the origin and source of the ore fluids. Magnetite (Fe2+Fe3+2O4) offers a reliable means of examining Fe-skarn mineralization [7]. This study focuses on the mineral chemistry and iron isotope composition of magnetite in the Pertek Fe-skarn deposit, highlighting the development of skarn and ore formation using chemical and isotopic parameters.

2. Geological Background

The geological history of Türkiye is shaped by the convergence and interaction of different tectonic units, such as the Pontides, Anatolides, Taurides, and Arabian Platform [8]. The Taurid block comprises Cambrian basement rocks, overlain by a succession of Paleozoic to Early Tertiary deposits. The Southeast Anatolian Orogenic Belt (SAOB) is a geologically intricate component of the larger Alpine-Himalayan system [9]. It is in Southeastern Anatolia, between the Taurides and the Arabian Platform.
The opening of the southern branch of the Neo-Tethys Ocean, situated between the Arabian Plate and the Tauride Anatolide Platform, commenced during the Late Triassic period [10,11,12]. This process persisted and was accompanied by the forming of mid-ocean ridge basalts (MORB) along a prominently developed oceanic ridge, spanning from the Late Triassic to the Early Cretaceous era [13,14]. The initiation of northward subduction within an intra-oceanic setting commenced during the Late Cretaceous epoch. This geological process led to ophiolites and arc magmatic rocks forming near the subduction zone. The formation of the Elazığ magmatic complex can be attributed to the subduction of the Neo-Tethyan oceanic slab and the subsequent collision between the Arabian Plate and the Tauride Anatolide Platform [15,16,17,18,19,20,21,22,23].
The study area is located west at the Pertek region within the Tauride block, between Elazığ and Tunceli (Figure 1b). It is worth noting that a substantial portion of the study area is currently submerged due to the presence of the Keban Dam Lake. The units observed in the study area are Keban Metamorphic rocks, Elazığ magmatic complex, Alibonca Formation and Karabakır Formation. Keban Metamorphic rocks mainly consist of recrystallized limestone, calc-schist, marble, metaconglomerate and calc-phyllites [15]. The Keban metamorphic rocks, which form the basement surrounding the study area, are primarily represented by marble and greenschists (Figure 1b). These rocks underwent metamorphism during the Late Cretaceous period and were deposited between the Permian and Cretaceous periods. The geological formation known as the Elazığ magmatic complex (Figure 1b) was initially designated as the Yüksekova complex but subsequently changed to the Elazığ magmatic complex. The Elazığ magmatic complex comprises volcanic, volcano-sedimentary, and plutonic rocks with intermediate to felsic compositions. These rocks include gabbro, diorite, tonalite, quartz diorite, granite, granodiorite, quartz monzodiorite, quartz monzonite, and monzodiorite, as documented in several studies [15,16,17]. Alibonca Formation overlies Keban Metamorphic rocks and Elazığ magmatic complex with an angular unconformity. The Alibonca Formation is represented in the study area by conglomerate, sandy limestone, limestone and marl. Based on the fossils found in the limestones of this unit, it was dated as Lower Miocene [24]. The Karabakır Formation overlies the Albonca Formation, Elazığ magmatic complex and Keban Metamorphic rocks with an angular unconformity. The Karabakır Formation is entirely represented by terrestrial environment sediments and terrestrial volcanic products [25].
Two major skarn zones have been identified in the Pertek Fe-skarn deposit: endoskarn and exoskarn. The endoskarn zone is defined by the presence of minerals such as garnet, and pyroxene, which originated during the final phases of metasomatism in proximity to dioritic rocks. The exoskarn zone is linked to carbonate rocks and comprises goethite, chlorite, magnetite, and quartz generated during retrograde metasomatism.
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Figure 1. (a) Location map of the study area and major tectonic setting of Anatolia (modified from [5]). (b) Geology map of the study area (modified from [26]). (c) Simplified geologic sketch map of the Pertek Fe-skarn mineralization.
Figure 1. (a) Location map of the study area and major tectonic setting of Anatolia (modified from [5]). (b) Geology map of the study area (modified from [26]). (c) Simplified geologic sketch map of the Pertek Fe-skarn mineralization.
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3. Materials and Methods

3.1. Sample Preparation

A total of 18 samples were collected from ore zones with Fe-skarn mineralization. To avoid metal contamination, the samples were ground into fine powder using Tungsten Carbide milling equipment. For geochemical analyses, powdered samples were subjected to a two-step acid digestion procedure using a Berghoff microwave digestion system. In the first step, approximately 0.2 g of each sample was weighed into Teflon vessels. A mixture of 6 mL HF (40%), 3 mL HNO3 (65%), and 1 mL HCl (37%) was added to each vessel. The samples were then digested at 135 °C for 45 min under microwave conditions to ensure the complete breakdown of silicate minerals and resistant matrices. In the second step, to prevent the precipitation of fluoride complexes (e.g., CaF2 and MgF2) that may form following HF digestion, a 5% boric acid solution was added to each sample. This step was carried out at 130 °C for 15 min under microwave heating. After digestion, the solutions were filtered and diluted appropriately for subsequent ICP–MS analysis.

3.2. Major and Trace Element Analysis

A comprehensive whole-rock geochemical investigation, encompassing major and rare earth elements (REEs), was performed at the Geochemistry Research Laboratories of Istanbul Technical University (ITU-JAL). The composition of major elements was analyzed using X-ray fluorescence (XRF) spectrometry (BRUKER S8 TIGER), whereas trace elements and rare earth elements (REEs) were quantified through Inductively Coupled Plasma Mass Spectrometry (ICP–MS) (ELAN DRC-e Perkin Elmer). The detection limits for REEs were less than 1 ppb, with uncertainty calculated at less than 5% for main elements and less than 10% for REEs.

3.3. X-Ray Diffraction (XRD) Examination

Mineral phases were determined using X-ray diffraction (XRD) at Istanbul Technical University (ITU-JAL). The analysis utilized a D8 ADVANCE BRUKER diffractometer with a Cu-sealed tube X-ray source producing Cu Kα radiation (1.5406 Å) at 40 kV and 40 mA. Mineral compositions were analyzed using X’Pert HighScore Plus version 3.0 software.

3.4. Iron Isotope Analysis (MC–ICP–MS)

The iron isotope analysis was conducted at ALS Scandinavia AB in Luleå, Sweden, utilizing Multicollector-Inductively Coupled Plasma Mass Spectrometry (MC–ICP–MS) with a Thermo Scientific Neptune Plus apparatus. Sample Preparation and Matrix Correction: Using a binocular microscope, magnetite grains were isolated from pulverized rock samples to achieve 99% purity. Samples were subjected to aqua regia leaching to eliminate contamination before isotopic analysis. Matrix effects were mitigated by employing column separation of iron with AG1-X8 resin before analysis. Calibration and Uncertainty: Internal standardization was executed utilizing standard-sample bracketing (SSB) with IRMM-014 and BHVO-2. Measured δ56Fe values were expressed concerning IRMM-014. The long-term precision exceeded ±0.05‰ (2SD), and the accuracy was within ±0.02‰. Reproducibility: Repeated measurements of reference materials revealed a long-term reproducibility within ±0.04‰.

3.5. Mineralogical Chemistry Assessment (LA–ICP–MS)

Mineral chemistry analyses were performed using a Perkin Elmer NexION 2000 mass spectrometer combined with an ESI NWR-213 solid-state laser ablation system in the Geochronology and Geochemistry Laboratory at Istanbul University–Cerrahpaşa (IUC-GGL). Laser spot size of 80 μm was selected for mineral chemistry analyses. For the analysis of each spot, ~20 s of background, 30 s of ablation, and 50 s of washout time was selected. Laser repetition rate was 5 Hz, and energy density was 3–5 J/cm2. Helium was used as a carrier gas with 0.6 l/s flow rate. The instrument was calibrated prior to the mineral chemistry analyses with ThO/Th ratio, which was always <0.05%. The quadrupole ion deflector (QID) settings were adjusted to provide maximum signal for light, medium, and heavy masses. The ICP–MS device has two detectors (analog and pulse detectors) with a wide dynamic range (up to 109) that can measure major and trace elements. The accuracy of our LA–ICP–MS measurements was ensured by separately measuring the 43Ca and 29Si content of BCR-2g glass standard with secondary electron microscopy-energy dispersive X-ray spectroscopy (SEM-EDX) (error margin <1% wt. for major oxides in repeatable measurements). NIST 610 and 612 standards were also measured for drift control and calibration of our measurements. Additionally, BCR-2g and AGV-2g were measured as control standards. Data reduction in all analyses was performed using the ICPMSDATACALL software package and the data reduction strategy of Liu et al. [27].

4. Results

4.1. Petrography

The Pertek Fe-skarn deposit occurs within the Southeast Anatolian Orogenic Belt (SAOB). The Pertek Fe-skarn represents a notable iron ore deposit situated within the Permo-Carboniferous Keban Metamorphic marbles [7]. In the study area, skarn rocks and iron mineralizations developed where dioritic rocks belonging to Elazığ magmatic complex intruded into marbles belonging to Keban Metamorphic rocks. Iron mineralizations (magnetite) accompany the skarn rocks almost everywhere in the study area (Figure 2). Skarn formations are formed as endoskarn and exoskarn, depending on their formation in the intrusive mass or host rock. These are skarn formations (endoskarn) that develop within and at the contact of dioritic rocks, and the other is skarn formations (exoskarn) within marbles where dioritic rocks are intruded.
Iron mineralizations are mostly developed in exoskarns. Magnetite mineralizations develop as massive and disseminated ore in the fracture and crack systems of exoskarns and in the cavities. Endoskarn minerals such as pyroxene and garnet are observed from dioritic rocks to marbles.
The mineral composition of the Pertek Fe-skarn deposit demonstrates a two-phase metasomatic evolution, as indicated by petrographic observations and X-ray diffraction (XRD) studies. The progressive metasomatism phase is characterized by the emergence of magnetite, hematite, calcite, quartz, and andradite garnet, signifying high-temperature skarn development. This phase is followed by retrograde metasomatism, distinguished by the presence of goethite, pyroxene, calcite, and quartz, which formed due to late-stage hydrothermal overprinting.
Petrographic investigation utilizing transmitted and reflected polarized-light microscopy indicates that magnetite is intimately connected with andradite garnet and hedenbergite pyroxene, affirming its genesis during the initial skarn stage. Conversely, late-stage hydrothermal activity substituted magnetite with goethite and hematite, especially inside oxidation zones (Figure 3a,b). The presence of calcite and quartz infilling magnetite cracks further indicates a shift from high-temperature skarn mineralization to subsequent fluid interaction (Figure 3c,d).
The ore paragenesis in both endoskarn (Figure 4a,b) and exoskarn (Figure 4c,d). zones is characterized by magnetite, hematite, and goethite, which are found in scattered and irregular crystal forms. Mineral associations including hematite, goethite, andradite garnet, hedenbergite pyroxene, calcite, and quartz exhibit significant variations at varying depths of the orebody. Coexisting minerals with magnetite include hematite, goethite, garnet, pyroxene, calcite, and quartz. Mineral assemblages vary at different levels of the orebody garnet type is andradite, and the minerals magnetite, hematite, and goethite are present along with the gangue quartz, according to the X-ray diffraction (XRD) analysis performed to support the mineralogy (Figure 5).

4.2. Mineral Chemistry

Magnetite grains from the four samples were analyzed for mineral chemistry by LA–ICP–MS (Table 1a,b). Magnetite grains from the Pertek Fe-skarn deposit contain variable concentrations of trace elements. They have ~0.50–1.61 ppm Sc, ~15.30–77.12 ppm V, ~1.68–2.83 ppm Cr, ~0.74–6.17 ppm Co, ~0.75–7.86 ppm Ni, ~4.56–8.46 ppm Ga and ~8.01–13.72 ppm Ge.

4.3. Whole-Rock Major and Trace Elements

The major oxide and trace element composition of Pertek Fe-skarn ore is given in Table 2. The SiO2, Al2O3, Fe2O3, MgO, CaO, and MnO average concentrations are 41.17 wt.%, 2, 62 wt.%, 20.56 wt.%, 1.94 wt.%, 24.68 wt.%, and 0.34 wt.%, respectively. Trace element averages are Sc: 10.18 ppm, V: 48.81 ppm, Co: 55.23 ppm, Ni: 48.17 ppm, Cu: 193.53 ppm, Zn: 94.53 ppm, As: 95.42 ppm, Sr: 51 ppm, Zr: 37.94 ppm, Nb: 2.94 ppm, Mo: 3.33 ppm, Ba: 27.37 ppm, Pb: 20.25 ppm, Th: 0.36 ppm and U: 1.18 ppm (Table 2). The REE contents of Pertek Fe-skarn ore are shown in Table 3. Diagrams derived from REE levels adjusted to those in chondrites are used to reveal the conditions of ore deposition [28,29,30]. The Fe skarn ore is characterized by a low content of ∑REE (5.05–10.57 ppm) in the study area.

4.4. Isotope Geochemistry

The mineral data obtained from the Pertek Fe-skarn deposit reveals a range of δ57Fe values for magnetite samples, spanning from 0.432‰ to 0.530‰. Similarly, the δ56Fe values for these magnetite mineralizations range from 0.272‰ to 0.361‰ (Table 4).
Frierdich et al. [31] shows Feaq–magnetite fractionation factor with time through distinct 56Fe/54Fe-time trends as a function of initial 56Fe/54Fe ratios for aqueous Fe. Provided that Fe solutions had low 56Fe/54Fe ratios, this shows variable removal of high-56 initial Fe isotope compositions for 57Fe-enriched Fe solutions similar to the studied magnetite. The study determined the Feaq–magnetite 56Fe/54Fe fractionation factor as −1.56 at 22 °C. The study includes temperature values used for separated magnetite from whole-rock samples and the fractionation based on magnetites was calculated below δ56/54 Feaq equations at average temperature 414.53 °C; 687.68 K.
The δ56Feaq value was calculated with the following equation [31,32].
103 lnαFe (II)aq– mgt= −0.145 (±0.002) × 106/T2 + 0.10(±0.02)
δ56Fe sample − δ56Fe aq= −0.145 × 106/(687.68 K2) + 0.10
AC1; 0. 272 − δ56Fe aq= −0.145 × 106/(687.68 K2) + 0.10
δ56Fe aq = 0.479‰
The fluid δ56Femelt isotopic composition ratio was calculated (Table 4) by taking the average value of the samples whose analysis was repeated (AC1; δ56Feaq = 0.479; AC8; δ56Feaq = 0.568‰, AC10; δ56Feaq = 0.566‰ and AC18; δ56Feaq = 0.512‰).
Fractionation factor values were calculated using the equation [32];
α= (δ56Fesample + 1000)/(δ56Feaq + 1000)
The fraction factor values calculated according to the above formula for all samples were calculated as AC1:0.99, AC8: 0.99, AC10:0.99, AC18:0.99, respectively. The temperature value in this study used in the Polyakov and Mineev [33] formula was calculated by taking the approximate average value of the exoskarn magnetite formation temperature.
103 ln β = 1.4369x2 − 6.3704 × 10−3 x2 +4.4984 × 10−5 x3 − 3.5915 × 10−6 x4, x = 106/T2
T = 687.68 K2 103ln β δ57/54; Fe = 0.006397.

5. Discussion

5.1. Mineral Chemistry and Geothermometry of Magnetite

Magnetite is a prevalent oxide mineral commonly found in various rock types, including igneous, sedimentary, and metamorphic formations. Its crystallographic structure is characterized by an inverted spinel arrangement, which provides significant physico-chemical insights into different geological settings. Magnetite maintains its solid structure, allowing it to preserve critical geological evidence [34,35,36]. Consequently, magnetite minerals offer valuable insights into mineral deposit formation within their paragenetic context [37].
The results of the magnetite analysis using LA–ICP–MS are presented in Table 1. The Al+Mn system versus the Ti+V system has been proposed as an effective method for distinguishing the temperatures at which magnetite forms [37]. The samples were plotted within a temperature range of 200–500 °C (Figure 6a), suggesting the presence of hydrothermal skarn magnetite formed at medium to high temperatures [38]. This observation implies that magnetite formation may have occurred through the precipitation of heated fluids enriched in aluminum and titanium. The presence of titanium (Ti), aluminum (Al), vanadium (V), and magnesium (Mg) in seawater is frequently associated with magmatic rocks [39,40]. The formation of these ores can be attributed to hydrothermal processes occurring at elevated temperatures. This finding is consistent with the inference drawn from the presence of trace amounts of magnetite, particularly within the 200–500 °C temperature range (Figure 6a).
Figure 6b presents the Al+Mn versus Ti+V diagram, which also includes the skarn and porphyry fields as proposed by [41]. This discriminant diagram shows that the ore deposits plot within the skarn field. Additionally, the magnetite grains from the Pertek Fe-skarn deposit fall within the skarn field on the Ni/(as introduced by Dupuis and Beaudoin [35]. The samples cluster around the skarn-type region in the TiO2-Al2O3-MgO+MnO magnetite genetic classification diagram (Figure 6d). The characteristic distribution patterns of the magnetite samples in these discriminant diagrams confirm their origin as skarn deposits.
In skarn deposits, the trace element composition of magnetite serves as a crucial indicator of the formation conditions, particularly temperature and fluid composition. The high Ti (>300 ppm) and V (>200 ppm) levels in the Pertek magnetite suggest precipitation from high-temperature hydrothermal fluids (>300 °C), which aligns with magnetite formation in skarn systems, as indicated by previous studies [36,37]. The presence of Al and Mg further supports a magmatic–hydrothermal origin, as these elements are commonly concentrated in high-temperature magnetite associated with skarn and IOCG deposits [35]. These geochemical markers correspond with those found in the Yeshan and Baijian Fe-skarn deposits, reinforcing the genetic interpretation of the Pertek deposit as a high-temperature skarn environment [42,43].
The LA–ICP–MS study of the Han-Xing deposit revealed substantial Ti (300–800 ppm) and V (150–400 ppm) concentrations, suggesting that magnetite originated from high-temperature hydrothermal fluids [44]. Similarly, in the Makeng deposit, titanium concentrations range from 250 to 700 ppm, while vanadium concentrations range from 100 to 350 ppm, supporting the formation of skarn-type magnetite at elevated temperatures [45]. The Ti and V contents measured in the Pertek Fe-skarn deposit align with these findings, suggesting a similar temperature range and hydrothermal fluid composition.
Magnetite thermometry is an empirical method for estimating the crystallization temperature of magnetite in igneous and hydrothermal environments [46]. This technique is based on the correlation between the magnesium content in magnetite (XMg) and temperature (T). It provides a reliable temperature estimate for settings where conventional thermometry techniques are inapplicable, such as hydrothermally altered rocks and ore deposits.
Canil and Lacourse [46] proposed an empirical geothermometer based on TMg-mag (°C) applied to magnetite from various settings, including igneous rocks, hydrothermally altered rocks, and ore deposits. The Mg content of magnetite varies with temperature (T) and is formulated as:
TMg-mag(°C) = (8344 ± 320)/(lnXMg − 4.13 ± 0.28) − 273
  • − TMg-mag (°C) is the temperature of magnetite, and
  • − XMg is the magnesium content in magnetite.
Magnetite thermometry provides a valuable tool for estimating crystallization and alteration temperatures in diverse geological settings [46]. While it does not replace conventional two-mineral thermometry methods, it is an essential alternative in rocks where magnetite is the only available phase. Further refinement of trace element partitioning models may improve their accuracy and applicability in complex geological systems. The average temperature value of magnetites is 414.53 °C.
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Figure 6. (a,b) Al + Mn vs. Ti + V (wt.%) diagram and (c) Ni/(Cr + Mn) vs. Ti + V (wt.%) with fields for different deposit types. Base diagram after Dupuis and Beaudoin [35] modified by Nadoll et al. [37]. (d) Ternary TiO2-Al2O3-(MgO+MnO) discrimination diagram [47].
Figure 6. (a,b) Al + Mn vs. Ti + V (wt.%) diagram and (c) Ni/(Cr + Mn) vs. Ti + V (wt.%) with fields for different deposit types. Base diagram after Dupuis and Beaudoin [35] modified by Nadoll et al. [37]. (d) Ternary TiO2-Al2O3-(MgO+MnO) discrimination diagram [47].
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5.2. Major, Trace Element and Rare Earth Elements Geochemistry of Whole Rock

Mineralization processes, chemical compositions, and formation conditions within skarn zones influence the type of skarn mineralization. Deposits formed from early and rapid hydrothermal fluid deposition typically exhibit high Fe/Mn ratios (>10) with a wide range of variations, as noted by Crerar et al. [48]. In contrast, Bonatti et al. [49] reported that the Fe/Mn ratio in sedimentary deposits is approximately one, with minimal variation. Exhalative deposits, as described by Rona [50] and later Nicholson [51], exhibit extensive fluctuations in the Fe/Mn ratio. The Fe/Mn ratio in the Pertek Fe-skarn deposit samples ranges from 17.15 to 671.71 (Table 2). These results suggest that the iron-rich, manganese-poor mineralization was formed by early and rapid deposition from hydrothermal solutions.
SiO2-Al2O3 variation diagrams are effective in distinguishing between sedimentary deposits and hydrothermal mineralizations [48]. Figure 7 presents the SiO2-Al2O3 variation diagram for samples from the Pertek Fe-skarn deposit. Most samples fall within the hydrothermal field, with some located near the boundary between the hydrothermal and detrital fields.
The most suitable classification of skarn deposits is based on calc-silicate mineral assemblages. Magnesian skarn, also known as dolomite replacement skarn, is primarily composed of magnesian silicate minerals such as forsterite and serpentine. Calcic skarn, which forms through limestone replacement, predominantly consists of Fe-Ca silicates like andradite and hedenbergite. Alternatively, skarn deposits can also be classified based on their primary economic metal content, including Fe, W, Cu, Zn-Pb, Mo, and Sn [26]. Among these, iron skarns represent the most significant skarn deposits.
Iron skarns are primarily mined for their magnetite content. Although minor amounts of Cu, Co, Ni, and Au may be present, iron is typically the sole economically viable commodity [52]. According to Kesler [53] and Vidal et al. [54], there are several large iron skarn deposits, with volumes exceeding 500 million tons and containing over 300 million tons of iron. These deposits are characterized by predominant magnetite mineralization with minimal silicate gangue. The Pertek Fe-skarn deposit is significant for its substantial magnetite content.
Diagrams derived from rare earth element (REE) levels, normalized to chondrites, provide insights into ore deposition conditions [28,30]. The Pertek Fe-skarn deposit is characterized by a total REE content (∑REE) ranging from 5.05 to 10.57 ppm in the study area. The average total REE content of the Pertek Fe-skarn deposit is relatively low (∑REE = 89.63 ppm). The chondrite-normalized REE patterns (normalization values from Sun and McDonough [55]) indicate that light rare earth elements (LREEs) exhibit higher concentrations than heavy rare earth elements (HREEs), except for samples AC8, AC9, and AC12 (Figure 8).
Samples with elevated HREE concentrations were likely influenced by factors such as the composition of the magmatic source responsible for skarn mineralization, the environmental conditions under which skarn mineralization developed, and the characteristics of the host rocks affected by skarn replacement [56]. The Eu and Ce anomalies in the ore samples are reflected in Eu/Eu* values ranging from 0.63 to 7.37 and Ce/Ce* values from 0.5 to 1.15 (Table 3). Both negative and positive anomalies are observed for Ce and Eu. The samples exhibit a strong positive Eu anomaly, indicating a hydrothermal origin [57,58]. Additionally, specimens with weak negative Ce anomalies and positive Eu anomalies suggest that the hydrothermal solutions had high oxygen fugacity [59,60,61].
Low REE concentrations, LREE enrichment, HREE depletion, and fluctuating Eu anomalies are typical characteristics of magmatic fluids [62]. The distribution coefficient of REEs into the fluid phase can be enhanced, and Eu anomalies can be reduce by increasing the salinity of hydrothermal fluids. The pH of hydrothermal fluids has a significant impact on REE fractionation [63,64]. Under mildly acidic conditions, fluids are typically enriched in LREEs, depleted in HREEs, and characterized by negative Eu anomalies [64]. Conversely, near-neutral fluids are enriched in HREEs, depleted in LREEs, and exhibit either positive or no Eu anomalies [64,65].
The three most economically significant types of iron ore deposits in Türkiye are: (1) skarn-hosted ore deposits, which occur along the contacts between syenitic-monzonitic intrusions and limestone or serpentine; (2) vein-type deposits, which are found between serpentine and limestone; and (3) ore deposits that are entirely hosted within limestone [66]. A comparison of these iron deposits in Türkiye is shown in Figure 9. Based on this classification, the Pertek Fe-skarn deposit closely resembles a skarn-hosted ore deposit.
The La/Y ratios render valuable data for determining pH in the environment of ore formation. Values of La/Y > 1 indicate alkaline conditions, while values of La/Y < 1 indicate acidic conditions [67]. This ratio at the Pertek Fe-skarn deposit ranges from 0.13 to 17.79 (Table 2). The ore zone has the highest La/Y values (17.79). The skarn zones exhibit two significant features in this regard: (1) an area close to the ore body with a La/Y ratio greater than 1, and (2) a region that borders diorite and has a La/Y ratio of 1. These divisions indicate basic and acidic conditions of hydrothermal alteration during skarnization, respectively.
The concentrations of rare earth elements (REEs), from La to Lu, in the hydrothermal fluid may provide important clues about the origins of base metals and other ore-forming components. Due to their identical chemical and physical properties [58], the REEs exhibit consistent behavior throughout most geological processes. However, few geochemical investigations have been conducted on the rare earth elements (REEs) of skarn deposits [68,69]. The samples in the study area are in the magmatic water area (Figure 10).
Investigations of rare earth elements (REEs) indicate that the total REE content (∑REE) in the Pertek Fe-skarn deposit is relatively low compared to other skarn deposits, ranging from 5.05 to 10.57 ppm. In the Yeshan skarn deposit, ∑REE values range from 5.1 to 15.3 ppm, indicating enrichment in light rare earth elements (LREEs) and the development of hydrothermal fluids under anoxic conditions [42]. Similarly, in the Baijian deposit, ∑REE values range from 6.0 to 12.8 ppm, also demonstrating LREE enrichment [43]. The consistently low ∑REE values recorded in the Pertek Fe-skarn deposit are comparable to those in similar deposits, while the observed LREE enrichment aligns with the evolution of the hydrothermal system.
The low REE concentrations provide insights into the provenance of the hydrothermal fluids and parallel the formation mechanisms of oxidized minerals identified in the Baijian and Yeshan deposits. The distribution of Fe isotopes plays a crucial role in understanding the development of the hydrothermal system, particularly in terms of the influence of oxygen fugacity on mineral composition [42,71].

5.3. Iron Isotope Geochemistry

The isotope ratios of iron (Fe) exhibit substantial variations, which have significant implications for the geochemical characterization of iron mineralization, the origin of ore-forming fluids, the differentiation of magma from mantle sources, and the diffusion/atomic transport processes from regions of higher to lower concentration. Two major issues remain central to the understanding of magmatic–hydrothermal deposits [71]. The first concerns the fractionation of Fe isotopes during fluid exsolution [72,73,74,75]. The second pertains to the behavior of Fe isotopes in fluid-mineral systems [76,77,78].
Li et al. [79] suggested that high-pressure garnet crystallization occurs due to magma differentiation, leading to Fe isotope fractionation at the lower crust. The high Sr/Y ratio of the intrusion and dikes yields an average δ57Fe of +0.11 ± 0.02‰ (±2SD), which is comparable to the arc crust at convergent continental margins. Wawryk and Foden [75] determined that the iron isotope values in andesite and quartz diorite range from δ57Fe = 0.17 ± 0.05‰ to 0.26 ± 0.05‰. More felsic and differentiated rocks such as tonalite exhibit heavier iron isotope values (δ57Fe = 0.27 ± 0.08‰ to 0.32 ± 0.08‰) compared to intermediate and mafic rocks. Their study indicated that heavy Fe isotopes are positively correlated with SiO2 and negatively correlated with MgO, CaO, TiO2, and V, suggesting that the crystallization of clinopyroxene, amphibole, and magnetite primarily controls the isotopic evolution of the melt. Magnetite from the Batu Hijau ore deposit is isotopically heavier (δ57Fe ranges from 0.24 ± 0.14‰ to 0.74 ± 0.14‰) compared to coexisting chalcopyrite (−0.62‰ ± 0.04‰ to −0.16 ± 0.05‰) and bornite (−0.72 ± 0.23‰ to −0.08 ± 0.03‰), aligning with theoretical fractionation factors derived from spectroscopy.
At a temperature of 687.68 K, the calculated ln β value is 0.006397, indicating the fractionation of Fe isotopes in magnetite. At lower temperatures, ln β increases due to greater isotope separation. In contrast, at higher temperatures (such as 687.68 K in this study), ln β decreases, leading to reduced isotope fractionation and isotopic equilibrium between magnetite and hydrothermal solution. This suggests that Fe isotopes in magnetite are only slightly fractionated at these temperatures. As temperature increases, isotopic compositions become more similar to the solution, minimizing fractionation. This result aligns with findings by Polyakov & Mineev [33] and Polyakov et al. [32], who demonstrated how Fe isotope equilibrium between magnetite and solution varies with temperature.
A fluid inclusion study by Zhang et al. [80] determined that the average temperature of ore-forming fluids during magnetite deposition was 416 °C. Even if magnetite did not fully equilibrate with the ore-forming fluid, it likely preferentially fractionated heavy Fe isotopes [31,76]. Therefore, the ore-forming fluid should be isotopically lighter than magnetite in the Han-Xing Fe skarn deposit.
In this study, δ56Fe and δ57Fe analyses indicate that isotopically light Fe (56Fe) is less abundant than heavy Fe (57Fe). In the Yeshan deposit, δ56Fe values range from 0.15‰ to 0.35‰, suggesting precipitation under low-oxygen conditions [42]. The Baijian skarn deposit shows similar isotopic characteristics, with δ56Fe values ranging from 0.10‰ to 0.30‰ [43]. In contrast, the Han-Xing deposit exhibits a broader δ56Fe range (−0.05‰ to 0.25‰), indicative of varying oxidation conditions within the hydrothermal system [44].
For the Pertek Fe-skarn deposit, δ56Fe values in magnetite range from 0.272‰ to 0.361‰, while aqueous δ56Fe (δ56Feaq) ranges from 0.548‰ to 0.637‰, indicating deposition under magmatic–hydrothermal conditions (Figure 11).
Sağıroğlu [81] reported zoning patterns in the study area, including epidote skarn, garnet skarn-magnetite, pyrite garnet skarn-magnetite, and pyrometasomatic-marble. Sipahi et al. [82] noted that the epidote–magnetite assemblage in Fe–Cu skarn mineralization formed at temperatures between 255 °C and 438 °C. The average magnetite formation temperature in the study area was calculated as 414.53 °C, further supporting the skarn-type origin of Pertek magnetite mineralization. Moderate temperatures and relatively low salinity values in hydrothermal solutions, which contribute to magnetite formation in skarn zones, suggest the involvement of meteoric water in the magmatic system. The interaction between the solution and host rock plays a crucial role in these mineralization processes.
Mineral fractionation studies indicate that magmatic–hydrothermal fluids become progressively depleted in isotopically light iron during late-stage magmatic processes. High 56Fe/54Fe ratios suggest formation via leaching [83]. Tang et al. [84] stated that lower TiO2, V, and Cr concentrations in magnetite indicate co-crystallization with sulfide liquid across a broad temperature and oxygen fugacity range. The high Cu, Zn, and SO3 values (maximum Cu: 1592 ppm, maximum Zn: 802 ppm, maximum SO3: 9.66%) further support this interpretation.
Redox conditions significantly influence the shift from light Fe to heavy Fe isotopic compositions in hydrothermal fluids. Johnson et al. [85] proposed that high 56Fe/54Fe ratios could result from Fe isotope exchange between aqueous Fe57+ and ferric materials. Their three-isotope exchange experiments demonstrated that 57Fe-enriched aqueous Fe57+ might react with magnetite due to kinetic equilibrium effects on 57Fe/56Fe ratios. Rustad et al. [86] calculated 1000 ln β values between 5 and 7‰ for Fe3+(H2O)63+ and between 2.5 and 5‰ for Fe57+(H2O)657+, which align well with aqueous Fe experimental results and fractionation studies. Kinetically driven fractionation produced an aqueous Fe composition that was 0.4‰ greater than that inferred to be in equilibrium with magnetite [85].
The Fe isotope study of the Pertek Fe-skarn deposit indicates significant fractionation in the iron isotope composition of hydrothermal fluids. The δ56Fe values recorded in magnetite and other Fe-oxide minerals (δ56Femag: 0.272‰–0.361‰; δ56Feaq: 0.479‰–0.568‰) indicate the source of magmatic ore-forming fluids. Similar isotopic fractionation has been documented in the Han-Xing deposit in China, demonstrating that high-temperature oxygen and sulfide escape significantly influence skarn mineralization [44]. Johnson et al. [85] noted that δ56Fe values in trisilicate layers are higher than those in carbonates within high-grade contact metamorphic rocks, where isotopically heavy magnetite predominates [87]. Figure 11 summarizes Fe isotope variations across major geological reservoirs. The δ56Fe isotopic composition ratios in magnetite suggest that magmatic–hydrothermal fluids played a critical role in magnetite formation [44,74,75,88,89]. The δ56Fe values obtained in this study closely resemble those of magmatic and magmatic–hydrothermal fluids (Figure 12).
Most skarn deposits worldwide are closely associated with granitoids intruding into Permo-Triassic recrystallized shallow marine carbonate rocks, such as the Cihai iron deposit [90].
The isotope findings closely resemble the isotopic pattern observed in the Baijian skarn deposit, suggesting that both deposits were influenced by processes such as oxygen and sulfide escape [64]. Isotopic investigations of the Han-Xing and Makeng deposits also reveal comparable isotopic profiles to the Pertek Fe-skarn deposit [39]. In contrast, the heavier Fe isotopes detected in the Yeshan deposit indicate that mineralization occurred under elevated oxidation conditions during a distinct mineralization phase [63]. This study has examined magnetite mineral and magnetite aqueous (magnetite-aq) distribution values for the Pertek skarn deposit. The isotopic composition values suggest an igneous origin for magnetite, with fluid–host rock interactions playing a crucial role in skarn magnetite formation. The availability of elements in hydrothermal fluids is a primary control on the concentration of minor and trace elements in hydrothermal magnetite. Lower REE concentrations, combined with higher Sc, V, Cr, Co, Ni, Ga, and Ge contents in magnetite, indicate an igneous source for the magnetite. The formation conditions of the studied magnetite are linked to magmatic–hydrothermal fluids and the evolution of metamorphic host rocks.

6. Conclusions

This research provides a comprehensive analysis of the Pertek Fe-skarn deposit in the Eastern Taurides, Türkiye, focusing on mineral chemistry and iron isotopes in magnetite. Magnetite mineralization occurs as both massive and disseminated ore within the fracture and crack systems of exoskarns, as well as in cavities. Endoskarn minerals, including pyroxene and garnet, are observed in the transition from dioritic rocks to marbles. The chemistry of magnetite, analyzed using LA–ICP–MS, unequivocally positions Pertek’s magnetite samples within the skarn domain, reinforcing their skarn-type origin under high-temperature conditions. Additionally, magnetite thermometry calculations indicate an average formation temperature of 414.53 °C. The La/Y ratios in the Pertek Fe-skarn deposit range from 0.13 to 17.79. Based on these values, the skarn zones exhibit both basic (alkaline) and acidic hydrothermal alteration conditions during skarn formation. Investigations of rare earth elements (REEs) reveal that the total REE content (∑REE) in the Pertek Fe-skarn deposit is relatively low compared to other skarn deposits, ranging from 5.05 to 10.57 ppm. The consistently low ∑REE values are similar to those in comparable deposits, while LREE enrichment aligns with the evolution of the hydrothermal system. The δ56Fe isotopic composition ratios indicate that the studied magnetite was formed by magmatic–hydrothermal solutions, further supporting its skarn-related genesis.

Author Contributions

H.K., C.Y., M.A.E. and L.K. wrote the main manuscript text and prepared all figures and tables. All authors have read and agreed to the published version of the manuscript.

Funding

We thank Firat University Scientific Research Projects Coordination Unit (FÜBAP) of Firat University, Elazığ/Türkiye for its financial support (No: MF.21.29 and MF.24.114).

Data Availability Statement

All data are contained within the article.

Conflicts of Interest

Author Cihan YALÇIN was employed by the company SRG Engineering and Consultancy Ltd. Sti, 23100, Denizli, Türkiye. The remaining authors declare that the re-search was conducted in the absence of any commercial or financial relationships that could be construed as a potential conflict of interest.

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Minerals 15 00369 g002
Figure 2. (a,b) The view of Fe-skarn zones in the study area; (c,d) macro photographs displaying representative rock types within the Pertek Fe-skarn deposit: magnetite, limonite and quartz.
Figure 2. (a,b) The view of Fe-skarn zones in the study area; (c,d) macro photographs displaying representative rock types within the Pertek Fe-skarn deposit: magnetite, limonite and quartz.
Minerals 15 00369 g002
Minerals 15 00369 g003
Figure 3. Representative photomicrographs of endoskarn and exoskarn samples under plane-polarized light. (a,b) Endoskarn samples characterized by the presence of garnet (Gar), opaque minerals (Opq), quartz (qz), and calcite (cal). (c,d) Exoskarn samples showing pyroxene (prx), opaque minerals (Opq), quartz (qz), and calcite (cal). Mineral phases are labeled in the images.
Figure 3. Representative photomicrographs of endoskarn and exoskarn samples under plane-polarized light. (a,b) Endoskarn samples characterized by the presence of garnet (Gar), opaque minerals (Opq), quartz (qz), and calcite (cal). (c,d) Exoskarn samples showing pyroxene (prx), opaque minerals (Opq), quartz (qz), and calcite (cal). Mineral phases are labeled in the images.
Minerals 15 00369 g003
Minerals 15 00369 g004
Figure 4. Cross-polarized light (XPL) photomicrographs of exoskarn samples showing opaque mineral assemblages. (a–d) Magnetite (Mag) grains variably associated with hematite (Hem) and garnet (Gar). (b,d) Goethite (Gth) alteration rims and fractures developed over magnetite and hematite phases (yellow arrows).
Figure 4. Cross-polarized light (XPL) photomicrographs of exoskarn samples showing opaque mineral assemblages. (a–d) Magnetite (Mag) grains variably associated with hematite (Hem) and garnet (Gar). (b,d) Goethite (Gth) alteration rims and fractures developed over magnetite and hematite phases (yellow arrows).
Minerals 15 00369 g004
Minerals 15 00369 g005
Figure 5. XRD patterns of magnetite-bearing sample collected from the exoskarn zone.
Figure 5. XRD patterns of magnetite-bearing sample collected from the exoskarn zone.
Minerals 15 00369 g005
Minerals 15 00369 g007
Figure 7. SiO2 versus Al2O3 diagram for the study area ore [49].
Figure 7. SiO2 versus Al2O3 diagram for the study area ore [49].
Minerals 15 00369 g007
Minerals 15 00369 g008
Figure 8. Chondrite-normalized REE patterns of samples from the Pertek Fe deposit [55].
Figure 8. Chondrite-normalized REE patterns of samples from the Pertek Fe deposit [55].
Minerals 15 00369 g008
Minerals 15 00369 g009
Figure 9. Chondrite-normalized REE patterns for iron ores and average trends of different types of deposits. Normalized values for chondrite are from [55]; ‘1’ values are from [66].
Figure 9. Chondrite-normalized REE patterns for iron ores and average trends of different types of deposits. Normalized values for chondrite are from [55]; ‘1’ values are from [66].
Minerals 15 00369 g009
Minerals 15 00369 g010
Figure 10. REE concentrations and trends in magmatic and meteoric fluids are compared (this study shows the average REEs of the samples) [70].
Figure 10. REE concentrations and trends in magmatic and meteoric fluids are compared (this study shows the average REEs of the samples) [70].
Minerals 15 00369 g010
Minerals 15 00369 g011
Figure 11. Range in Fe isotope compositions for various fluids, rocks, and minerals (sample from [77]).
Figure 11. Range in Fe isotope compositions for various fluids, rocks, and minerals (sample from [77]).
Minerals 15 00369 g011
Minerals 15 00369 g012
Figure 12. Distribution of iron isotopic compositions (modified from [80]).
Figure 12. Distribution of iron isotopic compositions (modified from [80]).
Minerals 15 00369 g012
Table 1. (a) Representative mineral chemistry data of magnetite from the study area (major elements wt.%). (b) Representative mineral chemistry data of magnetite from the study area (others ppm).
Table 1. (a) Representative mineral chemistry data of magnetite from the study area (major elements wt.%). (b) Representative mineral chemistry data of magnetite from the study area (others ppm).
(a)
SampleSiO2TiO2Al2O3Fe2O3FeOMnOMgOCaONa2OK2OP2O5
AC-1_010.510.060.9768.9731.030.020.071.710.000.000.00
AC-1_020.140.021.1268.9731.030.010.030.480.000.000.00
AC-1_030.260.021.0268.9731.030.010.030.210.000.000.00
AC-1_040.320.051.1068.9731.030.010.040.220.000.000.00
AC-1_040.170.041.0068.9731.030.010.020.180.000.000.00
AC-8_010.960.000.8368.9731.030.030.092.550.010.010.00
AC-8_020.440.010.6968.9731.030.010.020.430.000.000.00
AC-8_030.170.000.4668.9731.030.000.010.080.000.000.00
AC-8_040.410.000.3768.9731.030.020.040.970.000.000.00
AC-8_050.340.000.5168.9731.030.010.021.070.000.000.00
AC-10_010.100.000.5668.9731.030.010.030.050.000.000.01
AC-10_020.120.000.3568.9731.030.010.020.190.000.000.00
AC-10_030.160.000.3368.9731.030.000.020.050.000.000.00
AC-10_040.120.000.5868.9731.030.000.010.060.000.000.00
AC-10_050.080.010.6868.9731.030.000.010.040.020.000.00
AC-18_010.790.010.7968.9731.030.030.092.130.010.000.10
AC-18_020.780.010.7968.9731.030.030.092.170.010.000.10
AC-18_030.080.010.6968.9731.030.000.010.040.020.000.00
AC-18_040.790.010.7968.9731.030.030.092.200.010.000.10
AC-18_050.260.010.6768.9731.030.010.020.660.000.000.01
(b)
SampleAC
-1_01		AC
-1_02		AC
-1_03		AC
-1_04		AC
-1_04		AC
-8_01		AC
-8_02		AC
-8_03		AC
-8_04		AC
-8_05		AC
10_01		AC
10_02		AC
10_03		AC
10_04		AC
10_05		AC
18_01		AC
18_02		AC-18_03AC-18_04AC-18_05
Sc1.411.611.161.271.341.140.600.510.740.780.900.590.560.561.201.311.361.231.360.69
V42.7628.7527.8328.3829.9677.1218.5215.3133.0920.8726.5327.6522.8124.3125.0845.6445.5325.0845.9541.97
Cr2.282.721.832.222.812.041.912.662.402.832.221.692.632.212.221.831.782.061.821.80
Co4.701.241.180.981.036.171.890.743.652.342.221.951.411.201.504.404.351.494.391.69
Ni4.531.990.982.311.594.231.321.077.873.631.601.021.210.960.915.495.400.755.411.66
Ga5.445.125.715.325.626.896.666.778.477.378.017.357.747.227.445.495.727.475.724.56
Ge10.999.258.218.018.0213.7312.5411.0011.5010.6012.6811.7810.6510.8310.369.599.4210.449.478.19
Rb0.100.070.050.050.040.280.310.100.140.070.350.090.220.070.470.080.080.450.080.05
Sr5.502.201.261.190.879.312.410.823.543.580.590.640.270.260.367.747.900.357.992.41
Y0.250.180.170.160.220.350.090.060.140.140.190.140.140.090.080.210.220.080.220.13
Zr7.056.644.6712.319.541.930.650.390.440.553.643.072.076.484.931.331.294.981.301.53
Nb0.460.180.120.300.280.140.070.010.010.010.030.020.010.030.060.110.100.060.100.11
Cs0.030.020.020.020.020.110.020.010.020.020.260.060.170.030.030.010.010.020.010.01
Ba3.041.481.631.181.019.991.490.333.271.901.040.510.510.470.364.214.040.364.090.91
La0.070.040.040.040.040.080.070.010.030.040.030.030.050.030.040.060.060.040.060.03
Ce0.150.090.140.090.100.160.210.030.080.080.020.040.060.040.120.130.130.130.130.07
Pr0.030.030.020.020.020.040.020.010.020.020.010.010.020.010.040.030.040.050.040.02
Nd0.180.120.120.150.120.130.080.040.070.100.030.070.070.060.070.160.160.070.160.08
Sm0.040.040.020.030.030.060.040.040.060.040.040.030.040.030.060.040.040.060.040.01
Eu0.010.010.010.010.010.010.070.010.050.010.010.010.010.010.020.010.010.020.010.01
Gd0.040.040.040.040.040.030.060.040.040.040.030.030.040.040.040.060.060.040.060.03
Tb0.010.010.010.010.010.010.010.010.010.000.010.010.010.000.010.000.000.010.000.00
Dy0.030.030.020.030.030.040.030.030.030.030.010.020.030.030.020.030.030.020.030.02
Ho0.010.000.010.010.010.010.010.010.010.010.010.010.010.010.010.010.010.010.010.01
Er0.030.030.020.030.030.030.030.030.030.030.030.010.030.030.020.030.030.020.030.02
Tm0.010.010.010.010.010.010.020.010.010.010.010.010.000.010.010.010.000.010.010.00
Yb0.050.040.040.040.040.040.050.050.020.040.040.040.040.040.030.030.030.030.030.03
Lu0.010.010.010.010.010.010.010.010.010.010.010.010.010.010.020.010.010.020.010.01
Hf0.100.060.040.180.110.050.030.030.030.030.030.020.010.050.030.030.020.030.020.01
Ta0.030.010.010.030.020.010.020.010.010.010.010.010.010.010.010.010.010.010.010.01
Pb0.380.350.360.270.500.470.270.050.660.250.120.090.050.160.170.200.210.170.220.12
Th0.140.110.100.220.170.050.020.010.010.010.090.110.090.200.170.050.040.170.040.03
U0.240.160.150.130.180.320.200.060.220.160.080.080.080.120.240.260.260.240.260.14
B9.454.323.464.073.497.636.626.655.795.887.957.776.386.736.514.254.196.284.243.40
Zn25.878.996.4210.636.9634.348.702.6518.6912.248.806.976.265.307.1131.8132.097.0232.457.15
Table 2. Major and trace element composition of ore samples (major oxide elements are in weight percent; unit for trace elements is parts per million).
Table 2. Major and trace element composition of ore samples (major oxide elements are in weight percent; unit for trace elements is parts per million).
AC 1AC 2AC 3AC 4AC 5AC 6AC 7AC 8AC 9AC 10AC 11AC 12AC 13AC 14AC 15AC 16AC 17AC 18
SiO2 29.4432.3867.4425.5035.1844.6932.0930.9028.3029.0742.0133.8543.3632.8234.8786.2679.2333.70
Al2O30.970.650.190.680.5917.960.265.136.270.684.367.660.480.280.390.540.540.30
Fe2O311.9531.486.247.1723.4412.7760.5016.5414.5331.839.6111.2629.7533.9331.543.584.3829.67
MgO0.290.190.220.250.516.280.133.943.640.3110.546.021.140.330.480.140.210.33
CaO34.7832.8915.2539.5129.7212.670.2735.7136.7433.3426.5032.9021.5029.7929.701.100.9331.08
Na2O<0.01<0.01<0.01<0.01<0.012.23<0.01<0.01<0.01<0.01<0.01<0.010.840.190.140.04<0.01<0.01
K2O0.02<0.010.020.010.010.48<0.01<0.010.010.020.03<0.010.200.02<0.010.040.020.01
TiO2<0.01<0.01<0.01<0.01<0.010.80<0.010.300.37<0.010.380.44<0.010.01<0.01<0.01<0.01<0.01
MnO0.290.270.190.340.320.220.100.410.370.430.390.661.050.500.390.010.010.24
P2O50.09<0.010.020.030.020.12<0.01<0.010.03<0.01<0.01<0.01<0.010.03<0.01<0.01<0.010.06
SO30.020.010.030.020.050.050.170.020.010.010.02<0.010.010.010.014.879.660.12
Cl<0.01<0.010.02<0.01<0.010.07<0.010.020.020.020.050.04<0.010.020.02<0.01<0.01<0.01
Cr2O3<0.01<0.01<0.01<0.01<0.01<0.01<0.01<0.01<0.01<0.01<0.01<0.01<0.01<0.01<0.01<0.01<0.01<0.01
LOI22.122.1110.3426.4810.131.566.176.989.674.276.077.101.622.032.433.374.954.43
Total52.5953.89510.925.6335.54288.233.3034.96228.343.8142.1833.8689.31117.8372.8473.2475.2597.7
Sc 9.007.005.007.007.0037.00<113.0018.007.0015.0013.006.006.007.002.003.008.00
V 19.003.0050.0010.005.00289.0020.0057.0086.005.0050.00140.00<132.0010.006.005.004.00
Co 26.0040.0059.0023.0023.0065.00325.0014.0015.0024.0025.0020.0048.0018.0028.0070.0092.0047.00
Ni 19.0061.0031.004.0035.0042.00203.0032.0022.0065.0016.0024.0064.0062.0056.0012.0017.0058.00
Cu 94.0016.0021.003.0032.0022.001592.00175.00<1<166.00103.00<1<113.002.005.00375.00
Zn 11.00<126.009.0018.0085.00802.0039.0033.008.0080.0061.00163.0046.0033.005.008.00<1
As <19.007.008.008.005.00<1<1<17.00<1<1<1<1<1298.00334.00<1
Sr 75.0013.0032.0071.0064.00289.0034.0034.0041.0019.0072.0056.0046.0019.0020.0013.0011.0026.00
Zr 15.0014.0079.0017.0015.0045.0015.0076.0066.0014.00105.00107.0020.0015.0018.0014.0014.0013.00
Nb 2.001.003.006.002.004.002.004.005.002.003.005.003.005.003.003.002.001.00
Mo 2.00<12.003.006.002.0015.001.001.003.001.00<12.002.004.003.003.003.00
Ba <1<19.00<1<154.0032.00<1<1<1<1<161.0016.0010.0022.0015.00<1
Pb <1<1<110.00<15.00<1<1<1<1<1<1<1<1<135.0036.005.00
Ga4.899.952.194.589.2915.363.8810.8811.8115.716.279.158.6115.5010.471.751.8315.20
Rb6.665.0513.196.186.2915.964.615.635.566.467.786.8512.9816.107.3418.5117.487.85
In0.050.220.010.150.810.040.610.370.360.420.170.131.002.431.960.050.030.31
Cs0.060.040.000.000.010.680.010.020.020.370.200.210.770.150.240.050.040.06
Tl0.010.010.000.000.010.060.000.000.000.010.010.010.010.000.010.240.390.00
Y3.383.047.626.321.6211.890.8815.7315.241.409.9548.134.1610.375.251.481.239.92
Th0.100.070.050.040.050.670.051.041.170.040.701.080.520.360.460.050.020.01
U0.150.180.890.200.570.242.632.862.320.290.321.352.354.232.210.130.320.15
Fe/Mn42.112234.621.475.858.3671.741.440.374.924.817.128.569.381.7397.1485.7128.1
Table 3. REE composition of ore samples (ppm).
Table 3. REE composition of ore samples (ppm).
AC 1AC 2AC 3AC 4AC 5AC 6AC 7AC 8AC 9AC 10AC 11AC 12AC 13AC 14AC 15AC 16AC 17AC 18
La0.481.491.920.886.813.840.220.941.385.251.821.484.457.704.730.400.261.05
Ce0.593.334.031.4712.018.400.366.337.548.685.314.1910.7616.078.710.650.441.19
Pr0.140.500.610.320.951.150.041.541.670.670.880.911.121.540.800.100.070.31
Nd0.681.892.891.282.225.170.128.518.981.424.146.293.214.762.310.360.241.12
Sm0.180.320.810.310.201.490.042.092.170.151.283.680.561.060.560.110.080.29
Eu0.100.230.280.200.220.630.051.401.290.350.301.200.340.340.260.060.030.24
Gd0.240.401.260.490.232.040.042.422.440.141.616.290.481.200.660.120.090.51
Tb0.030.060.210.090.040.340.010.420.400.030.271.180.080.220.110.020.020.10
Dy0.270.321.270.580.192.190.042.642.400.141.728.340.441.370.590.130.100.69
Ho0.060.070.280.130.040.450.010.590.520.030.351.850.100.300.130.060.040.16
Er0.200.200.770.390.131.330.031.921.580.091.075.900.291.000.390.100.070.49
Tm0.030.030.110.060.020.200.000.300.240.020.170.940.050.180.070.020.010.08
Yb0.180.170.690.360.121.330.022.001.670.061.056.330.221.330.460.080.060.43
Lu0.020.020.100.050.020.200.000.310.240.010.150.960.030.190.070.010.010.07
Ce/Ce*0.550.950.910.681.010.970.851.031.050.971.020.861.151.081.000.770.800.50
Eu/Eu*1.481.960.861.573.181.103.411.911.717.370.630.762.020.911.331.451.261.89
The Ce and Eu anomalies were calculated as Ce/Ce* = 2CeN/(LaN + PrN) and Eu/Eu* = EuN/(SmN × GdN)0.5, respectively, where the subscript ’N’ refers to normalized values for PAAS.
Table 4. Fe isotope ratios of magnetites.
Table 4. Fe isotope ratios of magnetites.
SampleMineralδ 56Fe2 SDδ 57/54Fe2 SD
AC1Magnetite0.2720.0430.4320.039
AC8Magnetite0.3610.0860.5400.095
AC10Magnetite0.3590.0200.5300.020
AC18Magnetite0.3050.0530.4190.144
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MDPI and ACS Style

Kara, H.; Yalçın, C.; Ertürk, M.A.; Kalender, L. Mineral Chemistry and Iron Isotope Characteristics of Magnetites in Pertek Fe-Skarn Deposit (Türkiye). Minerals 2025, 15, 369. https://doi.org/10.3390/min15040369

AMA Style

Kara H, Yalçın C, Ertürk MA, Kalender L. Mineral Chemistry and Iron Isotope Characteristics of Magnetites in Pertek Fe-Skarn Deposit (Türkiye). Minerals. 2025; 15(4):369. https://doi.org/10.3390/min15040369

Chicago/Turabian Style

Kara, Hatice, Cihan Yalçın, Mehmet Ali Ertürk, and Leyla Kalender. 2025. "Mineral Chemistry and Iron Isotope Characteristics of Magnetites in Pertek Fe-Skarn Deposit (Türkiye)" Minerals 15, no. 4: 369. https://doi.org/10.3390/min15040369

APA Style

Kara, H., Yalçın, C., Ertürk, M. A., & Kalender, L. (2025). Mineral Chemistry and Iron Isotope Characteristics of Magnetites in Pertek Fe-Skarn Deposit (Türkiye). Minerals, 15(4), 369. https://doi.org/10.3390/min15040369

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