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Article

Chemistry and Fe Isotopes of Magnetites in the Orbicular Bodies in the Tanling Diorite and Implications for the Skarn Iron Mineralization in the North China Craton

by
Ruipeng Li
,
Shangguo Su
* and
Peng Wang
School of Earth Sciences and Resources, China University of Geosciences (Beijing), Beijing 100083, China
*
Author to whom correspondence should be addressed.
Minerals 2025, 15(10), 1061; https://doi.org/10.3390/min15101061
Submission received: 12 August 2025 / Revised: 6 October 2025 / Accepted: 7 October 2025 / Published: 9 October 2025
(This article belongs to the Special Issue Using Mineral Chemistry to Characterize Ore-Forming Processes)
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Abstract

Skarn-type iron ore is economically significant, and numerous skarn ore deposits have been identified in the North China Craton. The newly discovered orbicular diorite in this region is distinguished from other analogous rocks due to the accumulation of large magnetite particles, which may shed new light on the genesis of this ore type. The magnetite in different parts of the orbicular structure exhibits distinct compositional differences. For example, magnetite at the edge has a small particle size (200 μm) and is associated with the minerals plagioclase and hornblende, indicating that it crystallized from normal diorite magma. By contrast, magnetite in the core has a relatively large particle size (>1000 μm), is associated with apatite and actinolite, and contains apatite inclusions as well as numerous pores. The size of magnetite in the mantle falls between that of the edge and the core. The syngenetic minerals of magnetite in the mantle include epidote and plagioclase. The magnetites in the cores of orbicules have a higher content of Ti, Al, Ni, Cr, Sc, Zn, Co, Ga, and Nb than those in the rim. The δ56Fe value of the core magnetite (0.46‰–0.78‰) is much higher than that of the mantle and rim magnetite in orbicules. Moreover, the δ56Fe value of magnetite increases as the V content of magnetite gradually decreases. This large iron isotope fractionation is likely driven by liquid immiscibility that forms iron-rich melts under high oxygen fugacity. The reaction between magma and carbonate xenoliths (Ca, Mg)CO3 during magma migration generates abundant CO2, which significantly increases the oxygen fugacity of the magmatic system. Under the action of CO2 and other volatile components, liquid immiscibility occurs in the magma chamber, and Fe-rich oxide melts are formed by the melting of carbonate xenoliths. Iron oxides (Fe3O4/Fe2O3) will crystallize close to the liquidus due to high oxygen fugacity. These characteristics of magnetite in the Tanling orbicular diorite (Wuan, China) indicate that diorite magma reacts with carbonate xenoliths to form “Fe-rich melts”, and skarn iron deposits are probably formed by the reaction of intermediate-basic magma with carbonate rocks that generate such “Fe-rich melts”. A possible reaction is as follows: diorite magma + carbonate → (magnetite-actinolite-apatite) + garnet + epidote + feldspar + hornblende + CO2↑.

1. Introduction

Skarn-type iron deposits are an important type of iron ore deposit in China. This type of deposit is commonly characterized by large reserves, high grades, and facilitated mining, thus holding important economic value. At present, the genesis of skarn-type iron deposits remains controversial, with two primary genetic models proposed: (1) the contact hydrothermal metasomatism model [1,2]. According to this model, after the upper part of the intermediate-to-acidic intrusions is consolidated, the residual magma exsolves a fluid containing volatiles; this fluid rises along structurally weak zones and reacts with surrounding country rocks to form the iron-bearing hydrothermal solution, which crystallizes to form large amounts of magnetite under reduced pressure and eventually accumulates to form a deposit [3]. However, the iron in the ore and that in the magma in this model cannot achieve mass balance [4], and the phenomena of vesicular structure and flat contact between the ore body and the surrounding rocks cannot be explained; (2) the genetic model of iron-rich magma: Iron-rich magma forms via differentiation of parent magma at depth, and this iron-rich magma then intrudes and crystallizes along weak structural zones to form magnetite [3,5,6,7]. This model can explain the occurrence of some iron veins and the porous structure of ores, but it is difficult to explain the rapid ascent of high-density iron-rich magma.
It can be seen that there are still issues with the two proposed genetic mechanisms of skarn-type iron deposits; thus, their genesis remains to be further clarified. In recent years, we have discovered an orbicular diorite in the Tanling area of the North China Craton. This orbicular diorite is distinguished from other orbicular rocks by the accumulation of large magnetite particles and the presence of dark actinolite aggregates in its core. Because the study area is located in the famous “Hanxing-type iron ore belt” and the Tanling orbicular diorite occurs near the contact zone between diorite intrusions and carbonate rocks, it is inferred that iron enrichment in this orbicular diorite is related to regional skarn iron mineralization.
Previous studies have shown that the major and trace element compositions of magnetite, as well as its iron isotopic composition, are helpful for understanding the processes of magnetite formation [8,9,10,11,12,13,14,15,16] and associated metallogeny. Magnetite in the Tanling orbicular diorite exhibits compositional variations from core to edge; therefore, this study analyzes the petrographic and geochemical characteristics of the magnetite in the Tanling area to clarify the origin of the magnetite and further explores the mechanism of iron enrichment in the orbicular diorite as well as its implications for regional skarn iron mineralization.

2. Regional Geology

The Wuan area is located in the central part of the North China Craton and the southern segment of the Taihang Mountains [17], as shown in Figure 1. Magmatic rocks are widely distributed in the region. Based on their distribution characteristics, these magmatic rocks can be divided into three representative magmatic belts from east to west: the Hongshan, Wuan, and Fushan belts. Previous zircon U-Pb dating results show that the intrusive complex in the Wuan area has an age between 120 Ma and 136 Ma [18,19,20,21], and that it formed during the Early Cretaceous magmatic event in eastern China. This indicates that the Wuan complex is located within the tectonic setting of strong lithospheric extension and thinning in the eastern North China Craton. The North China Craton, a major Archean craton in China, was destroyed during the Mesozoic—possibly due to the collision between the North China and South China blocks and the collapse of associated collisional orogens [22]. The Early Cretaceous magmatism—including the Tanling diorite in this study—marks the culmination of the cratonic destruction of the North China Craton [23].
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Figure 1. Regional geological map of the study area (modified after [24]).
Figure 1. Regional geological map of the study area (modified after [24]).
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The earliest strata exposed in this area are the Lower Zanghuang Group of the Archean [25], whose lithology mainly consists of gneiss, schist, and marble. The overlying Mesoproterozoic, Paleozoic, Mesozoic, and Cenozoic strata are in unconformable contact with it.
The region is tectonically controlled by NNE- and NEE-trending faults. The exposed strata in the Tanling area are mostly limestone and dolomite of the Ordovician Majiagou Formation, a small amount of siltstone of the Carboniferous Benxi Formation, as well as minor shale and mudstone, etc. [25].

3. Petrographic Characteristics

3.1. Orbicular Diorite

There is only one occurrence of orbicular diorite in the Wuan area. It is mainly composed of a core and a shell. The proportion varies with the degree of weathering, with the core accounting for approximately 15% of the volume of the orbicule and the outer shell for 85% (Figure 2a,b). Field observation shows that the orbicular diorite has the following characteristics: (1) the orbicular diorite can be divided into three parts: magnetite-actinolite-apatite orbicular core, epidote orbicular mantle, and single diorite orbicular shell; (2) the obvious vesicular structure is observed in the dark mass area of the orbicule’s core (Figure 3b); (3) carbonate xenoliths are present in the rock mass adjacent to the orbicular diorite (Figure 2c).

3.2. Characteristics of Magnetite in Each Part of the Orbicular Diorite

The petrographic characteristics and mineral assemblages of magnetite from the core to the rim of the orbicular diorite are as follows:
The magnetite in the orbicular core has a large particle size (approximately 1000 μm), a low degree of idiomorphism, and is mostly in the form of hypidiomorphic-heteromorphic spheroids, with curved grain edges. Apatite grains are commonly observed within both magnetite and actinolite (Figure 3b), and these inclusions are long columnar. The mineral assemblage in the core is mainly magnetite, actinolite, apatite, and titanite (Figure 3b). The core contains an extremely small number of albite and hornblende grains (Figure 4). In TIMA images, the voids are widespread within the orbicule’s core, and some of them are distributed between the inclusions and magnetite grains (Figure 5).
The magnetite at the rim of the orbicule has a small particle size (<200 μm), a high degree of idiomorphism, and mostly occurs as idiomorphic to semi-idiomorphic granular grains; its magnetite particle edges are relatively flat. The degree of magnetite idiomorphism in the mantle is lower than that in the rim, and the grain boundaries are less flat than those in the rim. The mineral assemblage of the rim is magnetite, albite, hornblende, quartz, titanite, epidote, and apatite (Figure 3d).
The particle size of magnetite in the orbicule’s mantle is intermediate between that of the core and the rim, and the mineral assemblage in the mantle is mainly magnetite, epidote, and plagioclase (Figure 3c). From the core to the rim of the orbicule, the pore area within the orbicule gradually decreases (Figure 5).

4. Methods

The major element analysis of minerals was conducted using an electron microprobe in the laboratory of Hebei Institute of Regional Geology and Mineral Resources. The analytical instrument model is JXA-8230, with an accelerating voltage of 15 kV, a beam current of 1 × 10−8 A, and a beam spot size of 1–10 μm. The analytical standard samples consist of 53 mineral species from SPI Company (USA), and the analytical accuracy is better than 1%.
Double-polished thin sections (with a thickness range of 60–100 μm) were prepared for LA-ICP-MS (Laser Ablation-Inductively Coupled Plasma-Mass Spectrometry) at the State Geological Laboratory and Testing Center, Academy of Geological Sciences, Beijing China. In situ laser ablation was performed using a Coherent GeoLasPro 193 nm laser ablation system coupled with an Agilent 7700 quadrupole ICP-MS. After a 20 s gas blank monitoring step, 160 successive laser pulses (4 Hz, spot size: 44 μm) were used to ablate the sample surface for approximately 40 s. The generated aerosols were carried by a helium carrier gas, then mixed with argon make-up gas via a T-connector before entering the ICP-MS for the acquisition of ion signal intensities. For the detailed analytical process, refer to [26].
Elemental distribution maps were obtained using a TESCAN Integrated Mineral Analyzer (TIMA) at the Institute of Geology, Chinese Academy of Geological Sciences (CAGS), Beijing. Automated mineralogical analyses for phase mapping were conducted on the thin sections, with an accelerating voltage of 25 kV, a beam current of 8.47 nA, a working distance of 15 mm, and a spot size of 100.50 nm.
In situ Fe isotopic analyses were obtained using a NEPTUNE Plus MC-ICP-MS (Thermo Fisher Scientific, Bremen, Germany), which was combined with an NWR FemtoUC femtosecond system (New Wave Research, Fremont, CA, USA). For the in situ Fe isotope analyses, the laser ablation system was operated in single-spot mode, with a spot size of 40 μm, frequency of 3 Hz, and laser fluence of ~3.5 J·cm−2.
For in situ Fe isotope analysis, the laser ablation system was operated in single-spot ablation mode, with a spot size of 20–40 μm, a frequency of 3–5 Hz, and a laser fluence of ~2.5–4.5 J·cm−2. Ultrapure water was introduced into the ICP to form a “wet” plasma atmosphere, which effectively suppressed the matrix effects during Fe isotope analysis. LA-MC-ICP-MS Fe isotope measurements were performed at high mass resolution (M/ΔM ≈ 7000, defined by 5%–95% peak width) to resolve molecular interferences from argon nitrides and argon oxides on Fe isotopes. The Faraday cup detector array was set to acquire the ion beams of 53Cr, 54Fe, 56Fe, 57Fe, 58Fe, and 60Ni. 53Cr and 60Ni were monitored and used to correct for isobaric interferences on 54Fe and 58Fe. The SSB (Sample-Standard Bracketing) method was employed for mass fractionation correction. The international Fe isotopic standard IRMM-014 (pure iron, 99.9% purity), from the Institute for Reference Materials and Measurements (IRMM) of the European Commission, was used as an external standard, and the in-house reference material magnetite MT-8 was used as a quality control sample for in situ Fe isotope analyses. Standards and reference materials were measured before every set of 8 samples analyzed for drift monitoring. Fe isotope data reduction was performed using the Iso-Compass software [27]. Details of the LA-MC-ICP-MS operating conditions and measurement parameters are described in Feng et al. [28].

5. Results

5.1. Major Elements of Magnetite

All electron probe microanalysis (EPMA) data for the major elements of magnetite are shown in Table 1. For core magnetite (magnetite in the cores of orbicules): SiO2 content is 0.03%–0.47%, CoO is 0.06%–0.25%, MgO is 0%–0.05%, and Cr2O3 is 0.03%–1.48%. For mantle magnetite (magnetite in the mantles of orbicules): SiO2 is 0.03%–0.56%, CoO is 0.05%–0.16%, MgO is 0%–0.04%, and Cr2O3 is 0.01%–0.78%. For rim magnetite (magnetite in the rims of orbicules): SiO2 is 0.02%–0.36%, CoO is 0.02%–0.17%, MgO is 0%–0.08%, and Cr2O3 is 0%–1.12%.

5.2. Trace Elements of Magnetite

Major element data for magnetite in various parts of orbicular diorite are shown in Table 1. Trace element data for magnetite in these parts are shown in Supplementary Table S1.
In the TIMA elemental distribution maps (Figure 6), the Fe and Ca elemental contents in the core of orbicular diorite are significantly higher than those in the rim. Core magnetite has higher contents of Ti (4140–17,361 ppm), Al (2163–8732 ppm), Ni (30.7–68.5 ppm), Cr (64–3147 ppm), Sc (2.14–5.19 ppm), Zn (50–1117 ppm), Co (65.7–453 ppm), Ga (34.4–82 ppm), Nb (0.46–1.29 ppm) than those in the mantle and rim (Figure 7, Figure 8 and Figure 9). On the bulk continental crust-normalized multi-element spider diagram, core magnetite is enriched in Pb, Ge, W, Mo, Sn, Ga, Ti, Zn, Co, V, and Cr, but depleted in Si, Ca, Mg, and high field strength elements (Y, P, Nb, Ta, Zr, Hf) (Figure 10). Generally, core magnetite is more enriched in highly compatible elements with respect to magnetite.
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Figure 8. Binary diagrams of magnetite from different parts of the Tanling orbicular diorite: (a) V vs. Ti, (b) V vs. Cr, (c) Al vs. Ti, (d) Al vs. Ni. The direction of oxygen fugacity change is referenced from [29,30,31,32].
Figure 8. Binary diagrams of magnetite from different parts of the Tanling orbicular diorite: (a) V vs. Ti, (b) V vs. Cr, (c) Al vs. Ti, (d) Al vs. Ni. The direction of oxygen fugacity change is referenced from [29,30,31,32].
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Figure 9. Binary diagrams of magnetite from different parts of the Tanling orbicular diorite: (a) Sc vs. Zn, (b) Co vs. Ni, (c) Ga vs. Zn, (d) Y vs. Nb, (e) Mn vs. P, (f) Ni/Cr vs. Ti.
Figure 9. Binary diagrams of magnetite from different parts of the Tanling orbicular diorite: (a) Sc vs. Zn, (b) Co vs. Ni, (c) Ga vs. Zn, (d) Y vs. Nb, (e) Mn vs. P, (f) Ni/Cr vs. Ti.
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Figure 10. Multi-element spider diagrams for different types of magnetite in the Tanling orbicular diorite. All data are normalized to bulk continental crust values (after [33]). (a) core magnetite; (b) mantle magnetite; (c) rim magnetite; (d) comparison of the average values of three types of magnetite. Note: The selected elements in this figure are ordered by their increasing compatibility with magnetite, based on compiled experimental and empirical partition coefficients between magnetite and silicate magmas. More details can be found in [34] and references therein.
Figure 10. Multi-element spider diagrams for different types of magnetite in the Tanling orbicular diorite. All data are normalized to bulk continental crust values (after [33]). (a) core magnetite; (b) mantle magnetite; (c) rim magnetite; (d) comparison of the average values of three types of magnetite. Note: The selected elements in this figure are ordered by their increasing compatibility with magnetite, based on compiled experimental and empirical partition coefficients between magnetite and silicate magmas. More details can be found in [34] and references therein.
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The content range of trace elements in the rim and mantle magnetite is narrower than that in core magnetite. Most of the rim magnetite has relatively low contents of Ti (249–2522 ppm), Al (556–1266 ppm), Ni (7.9–24.6 ppm), Cr (57–195 ppm), Sc (0.19–1.89 ppm), Zn (33–258 ppm), Co (54–170 ppm), Ga (15.3–34.2 ppm), and Nb (0–0.69 ppm) (Figure 8 and Figure 9). On the bulk continental crust-normalized multi-element spider diagram, rim magnetite is enriched in P, Pb, Ge, W, Mo, Co, and V, but depleted in Hf, Al, Nb, Ta, Sn, Mg, Ni, and Cr. (Figure 10). Mantle magnetite generally exhibits elemental characteristics intermediate between the core and rim magnetite, with Cr from 74 to 640 ppm, V from 2648 to 3685 ppm, Ti from 1197 to 4611 ppm, Al from 652 to 2859 ppm, and P from 29 to 5260 ppm. On the bulk continental crust-normalized multi-element spider diagram, mantle magnetite has similar characteristics to the rim magnetite (Figure 10). The content of Al+Mn gradually increases from the rim magnetite to the core magnetite (Figure 11).
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Figure 11. (Al + Mn) vs. (Ti + V) plots for discriminating among magmatic-type, iron/magma type, and skarn type magnetite (after [35]; data from: [8,36,37]).
Figure 11. (Al + Mn) vs. (Ti + V) plots for discriminating among magmatic-type, iron/magma type, and skarn type magnetite (after [35]; data from: [8,36,37]).
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5.3. In Situ Fe Isotopes

δ56Fe test data of magnetite in different parts of the orbicular diorite are shown in Table 2.
The overall distribution of δ56Fe in the studied magnetite ranges from −0.66‰ to 0.78‰. For core magnetite, δ56Fe values range from 0.46‰ to 0.78‰ with an average value of 0.58‰; for mantle magnetite, δ56Fe values range from 0.12‰ to 0.76‰, with an average of 0.56‰; for rim magnetite, δ56Fe values range from −0.66‰ to 0.27‰, with an average of 0.05‰. These results exhibit three characteristics:
  • The δ56Fe values of magnetite in the orbicular diorite tend to decrease from the core to the rim (Figure 12).
  • The Fe isotopic composition of rim magnetite falls within the range of Fe isotopic composition of magnetite from magmatic deposits [12] reported in previous studies, but the rim magnetite is slightly enriched in light Fe isotopes (Figure 13).
  • Compared with magnetite in skarn hydrothermal deposits and magmatic deposits [16,38], the core and mantle magnetite are more enriched in heavier Fe isotopes, with maximum δ56Fe values reaching 0.78‰ (Figure 13).
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Figure 13. Iron isotope distribution diagrams of magnetite from different parts of Tanling orbicular diorite. Shown is the distribution of iron isotopes (δ56Fe) in magnetite, incorporating data from this study and the available literature [12,14,15,39,40,41,42,43,44]. Reference fields for common hydrothermal magnetites [12,13,45] and magmatic magnetites [14,15,41,46] are included for comparison (literature iron isotope data are provided in Table 3). The orange area denotes the range of common igneous magnetites, the light grey area denotes the range of hydrothermal ores.
Figure 13. Iron isotope distribution diagrams of magnetite from different parts of Tanling orbicular diorite. Shown is the distribution of iron isotopes (δ56Fe) in magnetite, incorporating data from this study and the available literature [12,14,15,39,40,41,42,43,44]. Reference fields for common hydrothermal magnetites [12,13,45] and magmatic magnetites [14,15,41,46] are included for comparison (literature iron isotope data are provided in Table 3). The orange area denotes the range of common igneous magnetites, the light grey area denotes the range of hydrothermal ores.
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6. Discussion

6.1. Controls on Compositions of Magnetite

Factors controlling the trace element contents in magmatic magnetite include silicate melt composition, magnetite crystallography, coexisting minerals, oxygen fugacity, temperature, cooling rate, sulfur fugacity, and silica activity [29,47,48,49,50]. Magmatic magnetite crystallizes from evolved silicate melts, sulfide melts, or Fe-Ti-(P)-rich melts and thus contains various trace elements. Coexisting minerals, magma composition, and oxygen fugacity are the dominant controlling factors. Next, we will discuss these major factors that may influence the trace element distribution of magnetite in the Tanling orbicular diorite.

6.1.1. Magma Composition and Coexisting Minerals

The minerals associated with magnetite can reflect the crystallization environment and conditions of the magnetite. In the Tanling orbicular diorite, rim magnetite has a small particle size (<200 μm), and its associated minerals are mostly plagioclase and hornblende, indicating that it crystallized from normal diorite magma. Core magnetite has a large particle size (about 1000 μm) and is associated with apatite and actinolite. This mineral association is similar to that in Chile ore magma flow [36,51], and the P content of magnetite in the core and apatite is 1–2 orders of magnitude lower than that in the mantle and rim (Figure 9e). These observations suggest that core magnetite may have crystallized from an evolved iron-rich melt [36,51,52,53].
The trace elements of magnetite may carry important information about magmatic composition [8,9,10,11,12,13,14,15,30,35,47,49,54,55,56]. Magnetites in the cores of orbicules have higher Ti, Al, Ni, Cr, Sc, Zn, Co, Ga, and Nb than those in the mantle and rim (Figure 8 and Figure 9), indicating that the core magnetite is significantly different from rim magnetite and is not directly crystallized from normal diorite magma. The Cr content of the core magnetite shows a multi-peaked distribution (Figure 8b), which is consistent with the data reported by Wang and Zhou [57]. The highly compatible element Cr (DCr magnetite/melt = 100–600); [58] is inconsistent with the assumption of fractional crystallization of magnetite in diorite magma, where its content is expected to decline sharply and continuously. In addition, Detao He et al. noted that Fe-rich melts formed under high oxygen fugacity may have relatively high Al2O3 contents [59,60]. The Al content of magnetites in the cores of orbicules is an order of magnitude higher than that of the rim and mantle magnetites (Figure 8c,d), suggesting that the magnetites in the cores of orbicules may crystallize from iron-rich melts with high oxygen fugacity. The enriched characteristics of compatible elements in the magnetites in the cores of orbicules of the Tanling orbicular diorite in Wuan, Hebei Province (Figure 10) indicate that they crystallize from a homogeneous iron-rich melt, which is consistent with the latest model proposed by Liu et al. [61] based on evidence from multiphase solid inclusions.
To gain a more concise and intuitive understanding of the genesis of magnetite, many researchers have addressed this need by establishing genetic discrimination diagrams for magnetite. Dupuis and Beaudoin [35] first proposed the identification diagram of magnetite origins. Based on the compositions of characteristic elements of magnetite from different deposits, they constructed the (Al + Mn) vs. (Ti + V) diagram to identify different types of magnetite. While collecting data on magnetite in our study, we found a new trend in the (Al + Mn) vs. (Ti + V) diagram (Figure 11) of magmatic origin towards iron-rich magma origin, and the changes from rim to core of magnetite in orbicular diorite are consistent with this trend. Therefore, we propose that the (Al + Mn) vs. (Ti + V) chart can be used to distinguish among iron-magma-type, magmatic-type, and skarn-type magnetite. Additional data will be required to verify the broad applicability of this chart.

6.1.2. Oxygen Fugacity

Vanadium in magnetite can be used to trace melt-fluid evolution and identify magma evolution processes [62]. Published papers [27,30,31,32] have shown that only V3+ is easily incorporated into the lattice of magnetite, and the content of V3+ in magnetite is negatively correlated with fO2. In the Wuan orbicular diorite in Hebei Province, the V content varies as follows: it decreases from the rim magnetite (average: 653 ppm) to the mantle magnetite (average: 0.30 ppm), then slightly increases in the core magnetite (average: 2.98 ppm) (Figure 8a,b, Supplementary Table S1). This variation indicates that the core magnetite formed under high fO2 conditions. This is consistent with the phenomenon of the dissolution of ilmenite lamellae in the large-grained magnetite of the core. In addition, tremolite-actinolite aggregates occur in the core center. Yang Et Al. [63] calculated the Mg/(Mg + Fe2+) ratio using electron probe microanalysis (EPMA) data of calcium amphibole, and the results were concentrated in the range of 0.88~0.9, further confirming the relatively high oxygen fugacity of the core.
The above observations show that the oxygen fugacity increases gradually from the rim of the orbicule to the core, suggesting that the core part must involve the participation of oxidizing agents. However, what causes the oxygen fugacity of the core to be higher than that of the rim? According to previous studies [64], when the gas phase consists of CO2-CO, H2O-H2, S2-H2S-SO2, the following reaction occurs when CO2 is injected into the silicate melt-gas system containing FeO:
CO2 + 2FeO = CO + Fe2O3
This reaction affects the distribution of Fe3+/Fe2+ in the melt and modifies the oxygen fugacity. Meanwhile, researchers have also found that the total amount of CO2 involved was sufficient to significantly increase oxygen fugacity by several log units [64,65]. They proposed that the interaction between basic magma and carbonate rocks forms calcium silicate rocks and releases large amounts of CO2, and that the interaction between the basic magma and surrounding carbonate rocks triggers mineralization. The CO2-rich fluid released by rock decarbonization of carbonate rocks oxidizes the invading magma and promotes the crystallization of iron and titanium oxides, thereby leading to the formation of ore deposits [66].
Therefore, the crystallization of large-grained magnetite in the core of the orbicular diorite occurs under high fO2 conditions, followed by that of magnetite in the mantle, whereas magnetite at the rim of the orbicule forms under relatively low fO2 conditions. Therefore, we propose that the accumulation of large-grained magnetite in the core is caused by the addition of CO2, which increases the oxygen fugacity of the magma. This indicates that the oxygen fugacity of the magma may play an important role in controlling the crystallization of magnetite in the orbicular diorite.

6.2. Reasons for High δ56Fe in Magnetites in the Cores of Orbicules

The δ56Fe values of magnetite in the orbicular diorite are mostly >0‰, exhibiting igneous rock characteristics. The δ56Fe of the core and mantle magnetite are much higher than those of other related magmatic deposits except the marginal magnetite (Figure 13). The high δ56Fe of magnetites in the cores of orbicules may be the result of crystallization of light 54Fe silicate minerals [15,16,38], high oxygen fugacity [67,68], magma differentiation/immiscibility [16,39,40,46,69], exsolution of ilmenite [26,41,70,71]. These issues are discussed below in the context of Tanling orbicular diorite magnetite.

6.2.1. Crystallization of Light δ56Fe Silicate Minerals

The fractional crystallization of minerals in magma can lead to changes in the isotopic composition of iron. Minerals enriched in light δ56Fe values—such as olivine and pyroxene—crystallize, a process that leads to higher δ56Fe values in the residual magma [15,38,72]. In contrast, magnetite tends to enrich in heavy iron isotopes during its crystallization [16,46]. For magnetite in different parts of the Tanling orbicular diorite, there is little variation in the Fetotal, and no correlation is observed between Fetotal and δ56Fe (Figure 12a). This indicates that the fractional crystallization of minerals enriched in the light 54Fe isotope is not the main mechanism driving Fe isotopic fractionation in magnetite from the Tanling orbicular diorite.

6.2.2. Hydrothermal Alteration

Hydrothermal alteration is another mechanism for iron isotope fractionation [40,41,70]. The δ56Fe values of iron ores associated with hydrothermal fluids show significant variations, with a maximum range of 5‰ [73,74,75,76,77]. For the orbicular diorite, no strong correlation was observed between δ56Fe and the CaO/(Na2O + K2O) ratio, and the δ56Fe values of magnetites in the orbicule cores are more concentrated (Figure 12c and Figure 13). It can be concluded that the high δ56Fe values of core magnetites in the orbicules are unlikely to result from hydrothermal alteration.
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Figure 14. Triangular diagram of the end-member content of garnet in orbicular diorite (Nanminghe data from [78]). Garnet abbreviations: Gro—grossular, And—andradite, Alm—almandine.
Figure 14. Triangular diagram of the end-member content of garnet in orbicular diorite (Nanminghe data from [78]). Garnet abbreviations: Gro—grossular, And—andradite, Alm—almandine.
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6.2.3. High Oxygen Fugacity

As mentioned above, the oxygen fugacity (fO2) of magnetite in different parts of the orbicular diorite varies greatly during its formation. The crystallization of large-grained magnetite in the core of orbicular diorite occurs under high fO2 conditions, followed by that of magnetite in the mantle and then the rim. Since the partitioning of vanadium (V) in magnetite depends on fO2 [31], the V content of magnetite decreases from the rim to the core, while its δ56Fe value gradually increases (Figure 12c), which suggests that the higher δ56Fe values of core magnetites (relative to the low values of the rim magnetites) may result from crystallization under high oxygen fugacity. Therefore, fO2 may play an important role in controlling the iron isotopic system of the orbicular diorite and the crystallization of magnetite.

6.2.4. Magma Differentiation/Immiscibility

Yue Wang et al. [12,40] analyzed extensive Fe isotope data and showed that iron isotopic fractionation occurs not only during hydrothermal processes, but also during magma differentiation and liquid immiscibility. Magma immiscibility results in the enrichment of heavy Fe isotopes in the immiscible Fe-rich melt fraction and light Fe isotopes in the Si-rich melt fraction. Yonghua Cao et al. [39] proved via their Fe isotope data that magnetite in the Panzhihua deposit was crystallized from iron-rich melts formed by liquid immiscibility from parent magma under high oxygen fugacity conditions. When the melt evolves toward a Fe-rich trend, the residual melt becomes enriched in light Fe isotopes [69], and the more Fe-rich melt had higher Fe3+, more heavy iron enrichment, and higher fO2 [64]. As noted earlier, Fe-rich melts formed at high fO2 conditions may have relatively high Al2O3 contents [59,60]. Consistent with this, the Al content of core magnetite (2163–8732 ppm) is one order of magnitude higher than that in the rim (556–1266 ppm) and mantle (652–2859 ppm) magnetites (Figure 8c,d). There is a significant positive correlation between δ56Fe and Al content of magnetite (Figure 12d and Figure 13). The δ56Fe values of different types of magnetite commonly found in the world are summarized. The high δ56Fe values of core magnetite in the Tanling orbicular diorite in Wuan, Hebei Province, are similar to those in the Panzhihua Pluton. This evidence shows that the core magnetites were already enriched in heavy Fe isotopes when they formed, which is called Fe-rich magma. This Fe enrichment may be driven by magmatic degassing (a common process in many volcanic systems) or more likely by Fe oxide-rich melts formed via Fe precipitation [79,80].

6.2.5. Exsolution Lamellae of Ilmenite

Ilmenite generally occurs within the magnetite in the core of orbicules (Figure 4c), and the δ56Fe values in ilmenite are indeed very low [38]. The magnetites in the orbicule cores have ilmenite exsolution lamellae, while mantle magnetites do not have such lamellae, but there is no significant difference in δ56Fe values between the two (average δ56Fe of the magnetites in the core magnetites: 0.58‰; the average δ56Fe of mantle magnetites: 0.56‰). Although the exsolution of ilmenite (enriched in light 54Fe) would cause an increase of δ56Fe values in the host magnetite, instead, the two remain in iron isotope equilibrium. There is no significant difference in δ56Fe values between core and mantle magnetites, indicating that ilmenite exsolution has little effect on the high δ56Fe values of the magnetites.

6.3. Origin of Skarn Iron Ore

6.3.1. Melting of Carbonate Xenoliths by Magma

The petrographic observations show that the minerals in the core of the orbicule are obviously different from those in the rim. The main minerals in the rim include plagioclase, hornblende, quartz, apatite, and magnetite. The core minerals are mainly magnetite, apatite, tremolite-actinolite, epidote, and garnet (mainly grossular). The minerals in the core are more enriched in Ca than those in the rim (Figure 5). A large amount of skarn minerals and skarn alteration features were discovered in the WuanIron Mine, as well as typical skarn minerals garnet and epidote developed in the mantle and edge of the orbicules. This leads us to believe that the formation of minerals in the orbicules has a type similar to that of skarn deposits. However, magnetite within orbicules differs from typical skarn-type magnetite in that it contains internal vesicles. The presence of vesicles in magnetite indicates that the magnetite particles formed under conditions of rapid magma crystallization, and the gas was unable to escape, becoming trapped in their lattice. Previous studies have shown that pores in magnetite are important evidence for iron-rich magma mineralization [81], and the development of vesicles within or between minerals is consistent with the characteristics of viscous melt containing a large amount of volatile matter (i.e., ore-bearing melt-fluid flows [44,82]). Therefore, the formation process of magnetite in orbicular diorite differs from that of skarn-type magnetite.
The composition and its variation characteristics of garnet can usually reflect the changes in physicochemical conditions such as temperature, pressure, oxygen fugacity, and pH during mineralization [83,84,85]. Calculations of the elemental composition of garnet (Table 4) show that it consists mainly of grossular (Gro), andradite (And), and almandine (Alm) (with an average elemental formula is Alm16.04And23.68Gro59.88Pyr0.10Spe0.50Ura0.06). The garnet in the orbicular diorite is mainly calcium-aluminum garnet (Figure 14), which differs from the garnet in typical iron skarns and that in the Nanminghe iron ore of this area. This indicates that this garnet does not have a hydrothermal origin typical of traditional skarns, and when combined with the previous analysis, it is inferred that this garnet may have formed during the evolution stage of iron-rich melts. The garnet associated with magnetite is mainly calcium-aluminum garnet, indicating that the magnetites in the cores of orbicules formed in advance, consumed Fe3+, inhibited the formation of calcium-iron garnet, and made it become calcium-aluminum garnet. The formation of this calcium-aluminum-rich garnet may be related to the formation of carbonate rocks [86]. At relatively low temperatures, decarbonation reactions occur in the system to form iron-rich garnet (Alm75Prp17Grs8), magnesiowüstite (Mg# ≤ 0.13), and CO2-rich fluids. However, Mg-bearing wüstite is unstable under high-pressure CO2-rich fluids and in the presence of a carbonate-silicate melt: It is either completely oxidized or is dissolved in the melt or fluid phase, leading to the formation of carbonate-silicate melts enriched in Fe2+ and Fe3+. These melts thus act as potential metasomatic agents in the mantle [87].
Based on the field distribution characteristics of the strata, orbicular diorite is found to occur at the contact area between diorite intrusions and limestone, and carbonate xenoliths are present in the adjacent rock masses (Figure 2b). Detao He [59] recorded the recycling of sedimentary carbonates to the deep mantle by studying the abundant carbonate xenoliths in the Dali Lake basalts. Qiu Jiaxiang (1985) showed that intermediate-acidic magmas have a strong ability to assimilate surrounding rocks, while assimilative mixing with carbonates is the most likely to occur, resulting in an increase in the contents of Ca, Mg, and Fe, while a decrease in the contents of Si, Na, and K in the magma [88]. According to Yang’s [63] whole-rock principal weight data and TIMA element distribution map (Figure 6), orbicular calcium content increased significantly, while silicon, sodium, and potassium decreased significantly. It is hypothesized that the increase in calcium content in the core is attributed to the assimilation of the diorite intrusion and mixing of some limestone fragments, and this process alters the magma composition.
By studying the isotopic composition of rock masses and ores in this area, numerous researchers have found that: ① the εNd values in this area show characteristics of mixing between Archaean TTGs (εNd = −30) and Mesozoic intrusions (εNd = −8~−15) [89]; ② the Sr and Nd isotopic compositions are roughly negatively correlated, which also implies significant mixing of lower crustal components [90]; ③ there is little difference in Pb isotopic composition between the Pb isotopic composition of ores and that of the diorite pluton in this area, and these isotopic compositions exhibit significant characteristics of mixed Pb—with Pb mainly derived from the crust beneath the lead source area, and a small amount of mantle-derived Pb incorporated [91]. The magnetite in the cores of orbicules in the orbicular diorite of this area exhibits a strong positive Pb anomaly (Figure 10a), indicating that its formation is driven by the addition of crust-derived components. This might indicate that the mineralization process or iron enrichment occurred near the carbonate formations.
Therefore, both petrographic and geochemical characteristics show that there is evidence of carbonate xenoliths mixing in the formation of orbicular diorite. During the upwelling process of magma, the carbonate formations were softened. Carbonate xenoliths fell into the magma and were melted into orbicules, eventually forming orbicular diorite. A possible reaction is:
Diorite magma + carbonate → iron-rich magma (Magnetite + actinolite + apatite) + garnet + epidote + feldspar + hornblende + CO2↑

6.3.2. Formation Mechanism of Fe-Rich Melt

The ore magma concept was first proposed by J.E. Spurr in 1923. R. Fischer conducted experiments on the immiscibility of high-iron oxides, silicates, and phosphates, while A. R. Philpotts carried out experiments on the immiscibility of apatite-rich liquids, magnetite-apatite melts, and silicate melts [92,93,94]. These achievements enabled the iron ore magma concept to be firmly established and promoted its solid development in mineral deposit science. At the same time, the Gushan iron deposit, the Meishan main orebody, and the Hanxing area are also considered as ore magma type iron ore deposits, rather than skarn type iron ore deposits in the traditional sense [24,44].
Orbicular diorite occurs within diorite around the Wu’an iron ore. This special structure may represent a small processing plant and reflect the mineralization process of the Wu’an Iron Mine. Orbicular diorite has the pseudomorph of carbonate xenoliths, reflecting the process of its interaction with magma. A part of the carbonate xenoliths reacted completely and not recorded, while another part demonstrated the process of their mutual reaction, namely the presence of iron-rich melts. As for large-grained magnetite, it is mainly enriched in the center of the orbicules (Figure 5). We believe that the magma melted the carbonate xenoliths. During this process, an iron-calcium exchange occurred. When the reaction was complete, the center of the orbicules was the original location where the carbonate xenoliths existed, and also the center where iron was enriched. The mantle and rim of the orbicules are the contact points between the carbonate xenoliths and the magma, eventually forming orbicular diorite. Actinolite is one of the important minerals in magmatic ore deposits. Many previous studies have often directly assumed that actinolite in ores is a product of hydrothermal alteration, similar to actinolite formed during the wet skarn stage in skarn deposits, and thus inferred that this actinolite originated from diopside alteration, leading to the conclusion that these deposits are of hydrothermal origin. However, the metasomatism of diopside by actinolite has not been observed within the mining areas of some magmatic ore deposits in Chile; instead, idiomorphic actinolite phenocrysts or aggregates are present in fresh volcanic lavas and clastic rocks [95]. Similarly, in the Tanling orbicular diorite of Wuan, Hebei Province, actinolite often occurs in columnar or fibrous forms and is associated with magnetite and apatite. The evidence suggests that at least some actinolite, especially the part of the massive ore associated with magnetite and apatite, is likely to have crystallized directly from magma.
Regarding the exsolution mechanism of Fe-rich melt, previous studies suggested that the addition of phosphorus in the orbicule’s core caused the exsolution of Fe-rich melts when the magma evolved to the diorite stage. Philpotts’ experiments on the diorite-magnetite-apatite system demonstrated that iron can be relatively concentrated in a liquid phase to form iron-rich magma. The FeO-Ca5(PO4)3F-NaAlSiO4-CaMgSiO6 experiment conducted by Su Lianghe and the FeO-CaMgSi2O6-KMg3(AlSi3O10)F2 experiment conducted by Yu Xuehui both proved that Fe-rich melt can be formed by separation with the participation of volatile components [96,97,98]. During the intrusion process of magma, it reacts with carbonate rocks to generate components such as CO2. During this process, physical and chemical conditions (e.g., oxygen fugacity) change, resulting in immiscibility between iron-rich melts and the magma. After multiple enrichments, iron-rich magma is formed.
The results of this study show that core magnetite in the Wuan orbicular diorite (Hebei Province, China) crystallized from iron-rich melts. Variations in magnetite composition indicate that the genesis of the Hanxing-type iron deposits is the result of a series of reactions between magma and carbonate rocks. Therefore, the following conceptual model is proposed to explain this step-by-step process (Figure 15):
① Iron primarily exists as Fe2+ in normal silicate magma. During magma intrusion, the hot magma reacts with carbonate xenoliths [(Ca, Mg) CO3], generating large amounts of CO2, which greatly increases the oxygen fugacity of the magmatic system.
② Under the action of volatile components (e.g., CO2), liquid immiscibility occurs between the iron-rich magma and the silicate melt, iron-rich oxide melts are formed via melting and subsequent enrichment, and iron-rich magma is finally formed after multiple rounds of enrichment [6,94,99].
③ Iron oxides (Fe3O4/Fe2O3) will crystallize close to the liquidus due to high oxygen fugacity [64,99], accompanied by the formation of numerous iron-poor silicate minerals (e.g., tremolite and actinolite).
Minerals 15 01061 g015
Figure 15. Metallogenic model of orbicular diorite in Wuan (phase diagram from [64]). During (a) upward intrusion of magma, (b) hot magma reacts with the carbonate xenoliths to gen-erate abundant CO2, which greatly increases the oxygen fugacity of the magmatic system; (c) liquid immiscibility occurs in the magma chamber, and the Fe-rich oxide melt is formed by melting away. Iron oxides will crystallize close to the liquidus due to high oxygen fugacity; Finally, (d) it cools and crystallizes to form orbicular diorite.
Figure 15. Metallogenic model of orbicular diorite in Wuan (phase diagram from [64]). During (a) upward intrusion of magma, (b) hot magma reacts with the carbonate xenoliths to gen-erate abundant CO2, which greatly increases the oxygen fugacity of the magmatic system; (c) liquid immiscibility occurs in the magma chamber, and the Fe-rich oxide melt is formed by melting away. Iron oxides will crystallize close to the liquidus due to high oxygen fugacity; Finally, (d) it cools and crystallizes to form orbicular diorite.
Minerals 15 01061 g015

7. Conclusions

Comprehensive analysis of petrographic and geochemical characteristics of magnetite in the Tanling orbicular diorite yields the following conclusions:
  • Magnetite composition varies significantly across the orbicule: rim magnetite has a small particle size (about 200 μm) and is associated mainly with plagioclase and hornblende, indicating that it crystallizes from normal diorite magma. The magnetite in the core has a large particle size (>1000 μm), and is associated with apatite and actinolite, with apatite inclusions and a large number of vesicles in the core. The magnetites in the cores of orbicules have higher Ti, Al, Ni, Cr, Sc, Zn, Co, Ga, and Nb than those in the rim. The multi-peak distribution of Cr and the (Al + Mn) vs. (Ti + V) diagram of the magnetites in the cores of orbicules indicate that they crystallized from iron-rich melts.
  • The δ56Fe value of the core magnetite (0.46‰–0.78‰) is much higher than that of the mantle and rim magnetite in orbicules. The δ56Fe values of magnetite in different parts have no correlation with Fetotal and CaO/(Na2O + K2O), but they increase as the V content of magnetite gradually decreases. This large fractionation of iron isotopes is not caused by mineral fractionation, crystallization, or fluid alteration, but may be driven by liquid immiscibility that forms iron-rich melts under high oxygen fugacity.
  • The iron enrichment mechanism of the Wuan orbicular diorite is as follows: During upward intrusion of magma, hot magma reacts with the carbonate xenoliths [(Ca, Mg) CO3] to generate abundant CO2, which greatly increases the oxygen fugacity of the magmatic system; Under the action of CO2 and other volatile components, liquid immiscibility occurs in the magma chamber, and the Fe-rich oxide melt is formed by melting away. Iron oxides (Fe3O4/Fe2O3) will crystallize close to the liquidus due to high oxygen fugacity.
  • The aforementioned characteristics of magnetite in the Tanling orbicular diorite (Wuan, Hebei Province) indicate that diorite magma reacts with carbonate xenoliths to form “Fe-rich melts”; furthermore, skarn iron ore deposits are probably formed by the reaction of intermediate-to-basic magma with carbonate rocks to generate such “Fe-rich melts”. A possible reaction is as follows:
Diorite magma + carbonate → Fe-rich melt (Magnetite + actinolite + apatite) + garnet + epidote + feldspar + hornblende + CO2↑

Supplementary Materials

The following supporting information can be downloaded at: https://www.mdpi.com/article/10.3390/min15101061/s1. Table S1. Trace elemental compositions of magnetite (ppm).

Author Contributions

Conceptualization, R.L. and S.S.; Methodology, S.S.; Software, P.W.; Validation, R.L., S.S. and P.W.; Formal analysis, R.L.; Investigation, R.L.; Resources, S.S.; Data curation, P.W.; Writing—original draft preparation, R.L.; Writing—review and editing, R.L.; Visualization, P.W.; Supervision, R.L.; Project administration, S.S.; Funding acquisition, S.S. All authors have read and agreed to the published version of the manuscript.

Funding

This research was funded by the Strategic Key Metals Supernormal Enrichment Kinetics (92162213) Project of the Major Research Program.

Data Availability Statement

Data are contained within the article and Supplementary Materials.

Conflicts of Interest

The authors declare no conflict of interest.

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Figure 2. Field and structural characteristics of the Tanling orbicular diorite: (a) outcrop of orbicular diorite; (b) internal structure of orbicular diorite; (c) marble xenoliths in the rock mass adjacent to orbicular diorite.
Figure 2. Field and structural characteristics of the Tanling orbicular diorite: (a) outcrop of orbicular diorite; (b) internal structure of orbicular diorite; (c) marble xenoliths in the rock mass adjacent to orbicular diorite.
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Figure 3. Microscopic characteristics of the orbicular diorite at different positions: (a) microphotograph of the orbicular diorite; (b) core mineral association (plane-polarized light, −); (c) mantle mineral association (cross-polarized light, +); (d) rim mineral association (plane-polarized light, −). Mag—magnetite; Ap—apatite; Act—actinolite; Ep—epidote; Qz—quartz; Pl—plagioclase.
Figure 3. Microscopic characteristics of the orbicular diorite at different positions: (a) microphotograph of the orbicular diorite; (b) core mineral association (plane-polarized light, −); (c) mantle mineral association (cross-polarized light, +); (d) rim mineral association (plane-polarized light, −). Mag—magnetite; Ap—apatite; Act—actinolite; Ep—epidote; Qz—quartz; Pl—plagioclase.
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Figure 4. Magnetite BSE photos of orbicular diorite at different positions: (a) core mineral association; (b) magnetite and apatite around the voids; (c) exsolution of ilmenite; (d) magnetite inclusions in the core’s apatite; (e) mantle mineral association; (f) rim mineral association. Mag—magnetite; Ap—apatite; Act—actinolite; Ilm—ilmenite; Ep—epidote; Pl—plagioclase; MI—melt inclusions.
Figure 4. Magnetite BSE photos of orbicular diorite at different positions: (a) core mineral association; (b) magnetite and apatite around the voids; (c) exsolution of ilmenite; (d) magnetite inclusions in the core’s apatite; (e) mantle mineral association; (f) rim mineral association. Mag—magnetite; Ap—apatite; Act—actinolite; Ilm—ilmenite; Ep—epidote; Pl—plagioclase; MI—melt inclusions.
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Figure 5. Distribution diagram of major mineral phases of orbicular diorite (based on TESCAN Integrated Mineral Analyzer). Mag—magnetite; Ap—apatite; Act—actinolite; Ep—epidote; Qz—quartz; Ab—albite.
Figure 5. Distribution diagram of major mineral phases of orbicular diorite (based on TESCAN Integrated Mineral Analyzer). Mag—magnetite; Ap—apatite; Act—actinolite; Ep—epidote; Qz—quartz; Ab—albite.
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Figure 6. Elemental distribution maps of Si, Na, Fe, and Ca of orbicular diorite (acquired using TESCAN Integrated Mineral Analyzer, TIMA).
Figure 6. Elemental distribution maps of Si, Na, Fe, and Ca of orbicular diorite (acquired using TESCAN Integrated Mineral Analyzer, TIMA).
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Figure 7. Element variation in magnetite from different parts of the Tanling orbicular diorite.
Figure 7. Element variation in magnetite from different parts of the Tanling orbicular diorite.
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Figure 12. Binary diagrams of magnetite from orbicule in diorite: (a) FeOt vs. δ56Fe, (b) CaO/(Na2O + K2O) vs. δ56Fe, (c) V vs. δ56Fe, (d) Al vs. δ56Fe.
Figure 12. Binary diagrams of magnetite from orbicule in diorite: (a) FeOt vs. δ56Fe, (b) CaO/(Na2O + K2O) vs. δ56Fe, (c) V vs. δ56Fe, (d) Al vs. δ56Fe.
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Table 1. Major composition of magnetite (wt.%).
Table 1. Major composition of magnetite (wt.%).
Core Magnetite
Sample NameK2OCaOTiO2CoOP2O5Na2OMgOAl2O3SiO2Cr2O3MnOFeOTotalCaO/(Na2O + K2O)
Q22-A-Q1-10.03 0.07 0.16 0.06 0.01 0.06 0.05 0.53 0.47 1.48 0.02 86.24 89.68 0.76
Q22-A-Q1-20.02 0.03 0.20 0.12 n.d.0.04 n.d.0.42 0.22 0.50 0.08 87.56 89.72 0.43
Q22-A-Q1-3n.d.0.03 0.14 0.14 0.02 0.00 0.03 0.31 0.24 0.20 0.04 87.73 89.43 33.00
Q22-A-Q2-1n.d.0.04 0.15 0.14 0.01 0.01 0.03 0.22 0.19 0.23 n.d.88.78 90.35 2.73
Q22-A-Q2-2n.d.0.01 0.11 0.13 n.d.0.06 0.04 0.36 0.20 0.19 0.09 88.43 90.12 0.24
Q22-A-Q2-3n.d.0.05 0.15 0.30 0.03 n.d.0.02 0.42 0.39 0.80 0.02 88.17 90.87 -
Q22-A-Q3-1n.d.0.03 0.16 0.09 n.d.n.d.n.d.0.36 0.18 0.07 0.06 88.73 90.16 -
Q22-A-Q3-2n.d.0.04 0.18 0.11 n.d.n.d.0.03 0.42 0.21 0.10 0.04 88.04 89.70 -
Q22-A-Q3-3n.d.0.04 0.12 0.07 0.01 n.d.0.03 0.19 0.22 0.38 0.06 88.52 90.08 -
Q22-A-Q4-1n.d.0.03 0.09 0.06 n.d.0.04 n.d.0.38 0.17 0.34 0.07 87.83 89.53 0.55
Q22-A-Q4-2n.d.0.04 0.14 0.13 n.d.0.01 0.02 0.30 0.05 0.14 0.05 87.58 88.93 3.00
Q22-A-Q4-3n.d.0.06 0.13 0.16 0.02 n.d.0.02 0.22 0.06 0.08 0.05 87.92 89.20 -
Q22-A-Q5-1n.d.0.04 0.16 0.09 0.02 0.06 0.06 0.15 0.16 0.06 0.16 87.22 88.47 0.60
Q22-A-Q5-2n.d.0.04 0.12 n.d.n.d.0.15 n.d.0.21 0.12 0.10 0.03 92.02 92.98 0.28
Q22-A-Q6-1n.d.0.13 0.07 n.d.n.d.0.01 n.d.0.07 0.16 0.03 0.02 91.27 92.02 6.19
Q22-A-Q6-2n.d.0.12 0.03 n.d.n.d.0.01 n.d.0.05 0.03 0.26 0.03 90.82 91.51 8.92
Q22-A-Q6-30.01 0.07 0.16 0.06 0.01 n.d.0.04 0.13 0.18 0.41 0.04 87.62 89.28 5.58
Q22-A-Q7-1n.d.0.02 0.12 0.23 0.05 n.d.n.d.0.38 0.22 0.13 0.08 88.28 90.10 4.00
Q22-A-Q7-2n.d.0.07 0.03 n.d.n.d.0.01 n.d.0.03 0.04 0.03 0.04 92.24 92.66 5.29
Q22-A-Q7-3n.d.0.04 0.10 0.25 n.d.n.d.0.08 0.46 0.29 0.46 0.06 86.77 89.12 4.30
Middle Magnetite
Sample NameK2OCaOTiO2CoOP2O5Na2OMgOAl2O3SiO2Cr2O3MnOFeOTotalCaO/(Na2O + K2O)
Q22-B-1n.d.0.13 0.11 0.06 n.d.n.d.0.02 0.13 0.33 0.78 0.03 87.73 89.75 19.14
Q22-B-20.02 0.16 0.14 0.12 0.03 0.03 0.04 0.27 0.56 0.71 0.03 88.06 90.53 3.42
Q22-B-3n.d.0.06 0.05 n.d.n.d.0.06 n.d.n.d.0.09 0.04 n.d.92.16 92.62 0.88
Q22-B-4n.d.0.22 0.06 0.08 n.d.0.02 0.04 0.21 0.53 0.51 0.06 88.61 90.70 11.42
Q22-B-5n.d.0.07 0.14 0.08 n.d.n.d.0.03 0.18 0.25 0.56 0.03 88.34 90.16 -
Q22-B-6n.d.0.11 0.25 0.13 0.01 0.04 0.03 0.28 0.42 0.72 0.08 87.76 90.31 2.43
Q22-B-7n.d.0.10 0.04 0.09 0.05 0.05 0.01 0.05 0.11 0.04 n.d.89.40 90.29 1.83
Q22-B-80.01 0.12 0.06 0.05 0.01 0.03 0.03 0.14 0.22 0.02 0.04 87.73 88.81 3.03
Q22-B-9n.d.0.03 0.03 0.11 0.02 n.d.0.01 n.d.0.07 0.03 0.03 91.71 92.35 8.50
Q22-B-100.01 0.07 0.05 0.11 n.d.0.01 n.d.0.04 0.09 0.02 0.06 91.89 92.65 3.25
Q22-B-12n.d.0.05 0.06 0.16 n.d.0.05 0.03 0.02 0.05 n.d.n.d.91.29 92.07 1.00
Q22-B-130.01 0.04 0.07 0.09 n.d.0.03 n.d.0.04 0.06 0.05 0.03 90.22 90.99 0.97
Q22-B-14n.d.0.12 0.06 0.09 0.01 n.d.0.04 0.06 0.19 0.05 0.03 88.39 89.34 -
Q22-B-15n.d.0.04 0.03 0.09 n.d.0.08 n.d.0.01 0.03 0.01 0.02 91.48 92.11 0.47
Rim Magnetite
Sample NameK2OCaOTiO2CoOP2O5Na2OMgOAl2O3SiO2Cr2O3MnOFeOTotalCaO/(Na2O + K2O)
Q22-C-Q1-10.01 0.16 0.18 0.04 0.06 0.11 0.04 0.19 0.36 0.15 0.05 87.12 88.85 1.28
Q22-C-Q1-2n.d.0.16 0.16 0.07 n.d.0.07 0.02 0.09 0.27 0.13 0.03 87.18 88.62 2.41
Q22-C-Q1-3n.d.0.07 0.12 0.06 n.d.0.11 n.d.0.11 0.19 0.30 0.09 88.72 90.15 0.61
Q22-C-Q1-40.03 0.07 0.07 0.08 n.d.0.08 n.d.0.08 0.39 0.06 0.03 86.56 87.83 0.63
Q22-C-Q1-5n.d.0.14 0.05 0.12 0.02 0.04 0.04 0.08 0.33 0.03 0.03 88.49 89.71 3.71
Q22-C-Q1-6n.d.0.05 0.06 n.d.n.d.n.d.0.01 0.08 0.05 0.01 n.d.93.11 93.54 7.67
Q22-C-Q1-70.02 0.03 0.02 0.11 0.02 0.03 0.03 0.13 0.06 0.05 0.06 91.70 92.59 0.71
Q22-C-Q2-10.08 0.11 0.05 0.06 0.02 0.32 0.02 0.18 0.42 0.45 n.d.87.87 89.88 0.28
Q22-C-Q2-2n.d.0.06 0.04 0.17 n.d.0.00 n.d.n.d.0.03 n.d.0.02 92.01 92.55 -
Q22-C-Q2-30.02 0.08 0.24 0.13 n.d.0.13 0.08 0.24 0.41 1.12 0.06 86.53 89.48 0.54
Q22-C-Q2-4n.d.0.09 0.06 n.d.n.d.0.05 0.02 0.06 0.06 0.07 n.d.91.91 92.43 1.67
Q22-C-Q2-5n.d.0.05 0.07 0.10 n.d.0.07 n.d.0.06 0.12 0.04 0.05 88.84 89.78 0.66
Q22-C-Q2-60.02 0.04 0.07 0.08 0.02 0.05 0.03 0.16 0.30 0.02 0.06 87.93 89.11 0.65
Q22-C-Q3-1n.d.0.10 0.10 0.13 n.d.0.11 0.03 0.07 0.17 0.10 0.02 87.57 88.78 0.87
Q22-C-Q3-2n.d.0.03 0.13 0.02 0.01 0.07 n.d.0.02 0.06 0.05 0.01 93.41 94.20 0.42
Q22-C-Q3-30.01 0.07 0.04 0.08 0.01 n.d.n.d.0.01 0.04 0.01 0.02 91.86 92.51 7.20
Q22-C-Q3-4n.d.0.07 0.04 0.12 0.01 0.02 0.01 0.05 0.06 n.d.0.04 85.80 86.58 3.36
Q22-C-Q3-5n.d.0.02 0.05 0.14 0.02 0.06 n.d.0.07 0.02 0.04 0.03 92.76 93.60 0.28
Q22-C-Q3-6n.d.0.11 0.06 0.10 n.d.0.08 0.01 0.04 0.03 0.05 0.02 88.53 89.39 1.32
Note: all Fe was recalculated as FeO in the microprobe analyses.
Table 2. The analysis results of the Fe isotope for magnetite.
Table 2. The analysis results of the Fe isotope for magnetite.
Sample NameRemarksδ56FeIRMM(‰)Delta-2SE
MT-8Standards−0.200.06
MT-8Standards−0.190.06
MT-10Standards0.250.06
MT-10Standards0.240.06
MT-8Standards−0.080.06
MT-10Standards0.300.04
MT-8Standards−0.130.06
Q22-A-Q1-1Core magnetite0.670.05
Q22-A-Q1-2Core magnetite0.540.04
Q22-A-Q1-3Core magnetite0.580.04
MT-8Standards−0.190.05
Q22-A-Q2-1Core magnetite0.460.05
Q22-A-Q2-2Core magnetite0.460.05
Q22-A-Q4-1Core magnetite0.620.04
MT-8Standards−0.180.06
Q22-A-Q4-2Core magnetite0.610.04
Q22-A-Q5-1Core magnetite0.520.04
Q22-A-Q5-2Core magnetite0.560.05
MT-8Standards−0.110.05
Q22-A-Q6-1Core magnetite0.650.05
Q22-A-Q6-2Core magnetite0.780.04
Q22-A-Q6-3Core magnetite0.700.04
MT-8Standards−0.180.07
Q22-A-Q7-1Core magnetite0.610.06
Q22-A-Q7-2Core magnetite0.460.05
Q22-A-Q7-3Core magnetite0.510.05
MT-8Standards−0.130.06
MT-10Standards0.270.06
MT-8Standards−0.100.06
Q22-B-1Middle magnetite0.670.05
Q22-B-2Middle magnetite0.380.06
Q22-B-3Middle magnetite0.570.05
MT-8Standards−0.140.05
Q22-B-4Middle magnetite0.740.04
Q22-B-6Middle magnetite0.670.06
MT-8Standards−0.230.07
Q22-B-7Middle magnetite0.120.05
Q22-B-9Middle magnetite0.710.04
MT-8Standards−0.110.06
MT-8Standards−0.200.04
MT-8Standards−0.090.06
MT-10Standards0.310.06
MT-8Standards−0.120.06
Q22-B-12Middle magnetite0.560.05
MT-8Standards−0.100.04
Q22-B-13Middle magnetite0.760.06
Q22-B-15Middle magnetite0.400.05
MT-8Standards−0.090.07
MT-10Standards0.310.05
MT-10Standards0.300.05
MT-8Standards−0.140.06
MT-10Standards0.340.09
MT-10Standards0.320.07
MT-8Standards−0.070.05
MT-8Standards−0.090.06
MT-10Standards0.230.07
MT-10Standards0.230.06
MT-8Standards−0.100.06
MT-10Standards0.280.05
MT-8Standards−0.170.06
Q22-C-Q1-1Rim magnetite0.210.06
Q22-C-Q1-2Rim magnetite0.100.05
Q22-C-Q1-3Rim magnetite0.190.06
MT-8Standards−0.110.08
Q22-C-Q2-1Rim magnetite−0.660.04
Q22-C-Q2-3Rim magnetite0.160.05
MT-8Standards−0.070.05
Q22-C-Q3-1Rim magnetite0.020.08
Q22-C-Q3-2Rim magnetite0.070.09
Q22-C-Q3-3Rim magnetite0.130.08
MT-8Standards−0.170.06
MT-8Standards−0.150.05
Q22-C-Q3-4Rim magnetite0.230.07
Q22-C-Q1-4Rim magnetite0.060.08
Q22-C-Q1-5Rim magnetite0.240.05
MT-8Standards−0.170.07
Q22-C-Q1-6Rim magnetite0.090.06
Q22-C-Q2-4Rim magnetite−0.030.07
MT-8Standards−0.180.07
MT-10Standards0.270.07
MT-10Standards0.230.05
MT-8Standards−0.110.06
MT-8Standards−0.120.05
Q22-C-Q1-7Rim magnetite0.270.07
Q22-C-Q2-6Rim magnetite−0.230.06
Q22-C-Q3-5Rim magnetite−0.060.07
Q22-C-Q3-6Rim magnetite0.020.10
MT-8Standards−0.190.06
MT-10Standards0.210.09
MT-8Standards−0.140.06
Table 3. Iron isotope analysis for magnetite from reference materials.
Table 3. Iron isotope analysis for magnetite from reference materials.
SampleSample DescriptionSample Provenanceδ56Fe in ‰2σ
Apatite-iron oxide ore
Kiruna Mining District (KMD), Northern Sweden
Kiruna-17NY28Banded massive (Ap-)magnetite oreKiirunavaara mine, Kiruna0.19±0.03
K-Mt-1 (907/75)Massive magnetite oreKiirunavaara mine, Kiruna0.2±0.03
Ki-Mi-2a (1079/251)Massive magnetite oreKiirunavaara mine, Kiruna0.21±0.02
Ki-Mi-2b (1079/251)Massive magnetite oreKiirunavaara mine, Kiruna0.22±0.04
K-Mt-1079/303Massive magnetite oreKiirunavaara mine, Kiruna0.27±0.04
K-Mt-1079/437Massive magnetite oreKiirunavaara mine, Kiruna0.16±0.02
M1931Massive magnetite oreKiirunavaara mine, Kiruna--
M1937Skeletal magnetite oreKiirunavaara mine, Kiruna--
LVA-3Massive magnetite oreLuossavaara mine, Kiruna0.23±0.02
LVA-FW-1Massive magnetite oreLuossavaara mine, Kiruna0.12±0.04
LVA-FW-2Massive magnetite oreLuossavaara mine, Kiruna0.27±0.03
K-Mt-3Massive magnetite oreMertainen mine, Kiruna0.41±0.03
K-Mt-4Massive magnetite oreMertainen mine, Kiruna0.29±0.03
M7557*Massive magnetite oreRektorn mine, Kiruna--
Grängesberg Mining District (GMD), Central Sweden
DC717-KES090068Massive (Ap-)magnetite oreGrängesberg mine, Grängesberg0.4±0.03
DC717-KES090070Massive (Ap-)magnetite oreGrängesberg mine, Grängesberg0.24±0.03
DC717-KES090072Massive (Ap-)magnetite oreGrängesberg mine, Grängesberg0.33±0.03
DC717-KES090084Magnetite vein in intermediate volcanic rockGrängesberg mine, Grängesberg0.11±0.03
DC690-KES090011Massive (Ap-)magnetite oreGrängesberg mine, Grängesberg0.31±0.03
DC690-KES090012Massive (Ap-)magnetite oreGrängesberg mine, Grängesberg0.31±0.04
DC690-KES090020Massive Ap-veined magnetite oreGrängesberg mine, Grängesberg0.3±0.04
DC690-KES090024Massive (Ap-)magnetite oreGrängesberg mine, Grängesberg0.26±0.03
DC690-KES090027Ap-veined/banded massive magnetite oreGrängesberg mine, Grängesberg0.29±0.03
DC690-KES090030Silicate-spotted massive (Ap-)magnetite oreGrängesberg mine, Grängesberg0.39±0.04
DC690-KES090034Coarse-grained Ap-spotted massive magnetite oreGrängesberg mine, Grängesberg0.27±0.03
DC690-KES090044Disseminated magnetite in intermediate volcanic rockGrängesberg mine, Grängesberg0.24±0.03
DC575-KES103011Magnetite-dominated massive oreGrängesberg mine, Grängesberg0.31±0.03
DC575-KES103016Coarse, massive (Ap-)magnetite oreGrängesberg mine, Grängesberg0.27±0.04
DC575-KES103003Magnetite-dominated massive oreGrängesberg mine, Grängesberg1±0.03
KES091013bMassive (Ap-)magnetite oreBlötberget mine, Blötberget0.33±0.03
El Laco Ap-Fe-oxide deposit, Chile
EJ-LS-11-1Massive magnetite oreLaco Sur, El Laco0.28±0.03
EJ-LS-11-2Massive magnetite oreLaco Sur, El Laco0.24±0.03
EJ-LS-11-3Massive magnetite oreLaco Sur, El Laco0.36±0.03
EJ-LS-11-4Massive magnetite oreLaco Sur, El Laco0.34±0.03
LS-2Massive magnetite oreLaco Sur, El Laco0.27±0.03
LS-52Massive magnetite oreLaco Sur, El Laco0.28±0.03
Plutonic reference material
RuoutevareTi-magnetite, layered igneous intrusionKvikjokk, Norrbotten, Sweden 0.31±0.03
UlvönTi-magnetite, layered igneous intrusionUlvön island, Ångermanland, Sweden 0.13±0.03
TabergTi-magnetite, layered igneous intrusionIron mine, Taberg, Småland, Sweden 0.23±0.04
EM419Massive Fe-Ti magnetite oreNorthern pit, Panzhihua, China0.61±0.05
EM424Massive Fe-Ti magnetite oreNalahe, Panzhihua, China 0.12±0.04
BushveldMassive magnetite oreUpper Zone, Bushveld Complex, South Africa--
GabbrobombMagnetite from a gabbro xenolithNW-Flank, Skjaldbreiður, Iceland 0.46±0.03
Volcanic reference material
TEF-NER-18Magnetite from an ankaramite dykeNE Rift Zone, Tenerife, Spain0.07±0.05
TEF-NER-57BMagnetite from an ankaramite dykeNE Rift Zone, Tenerife, Spain0.16±0.02
TEF-NER-70Magnetite from a pyroxene phyric dykeNE Rift Zone, Tenerife, Spain0.1±0.02
MG-07Igneous magnetite from daciteS-Flank, Mt. Ruapehu, New Zealand 0.32±0.03
MG-09Igneous magnetite from daciteS-Flank, Mt. Ruapehu, New Zealand 0.29±0.03
Kelut A1Igneous magnetite from basaltic andesiteMt. Kelut, Java, Indonesia0.1±0.04
GD-D-2Igneous magnetite from basaltic andesiteGede Dome, Java, Indonesia0.12±0.03
AK-B1Igneous magnetite from basaltic andesiteSE-Flank, Anak Krakatau, Indonesia0.06±0.03
AK-B3Igneous magnetite from basaltic andesiteSE-Flank, Anak Krakatau, Indonesia0.16±0.03
A-BA-1Igneous magnetite from basaltic andesiteMt. Agung, Bali, Indonesia0.18±0.05
M-BA06-KA-3Igneous magnetite from basaltic andesiteMt. Merapi, Java, Indonesia0.17±0.03
83/CRS/6Igneous magnetite from a dolerite dykeAgros, Troodos Massif, Cyprus 0.34±0.03
Low-temperature or hydrothermal magnetites
KES091007BCalcite bearing-magnetite oreBjörnberget, Sweden−0.02±0.03
DM-1Iron-skarn magnetite oreBotenhäll, Dannemora, Sweden −0.36±0.03
DM-2Iron-skarn magnetite oreNorrnäs 3, Dannemora, Sweden 0.01±0.03
DM-3Iron-skarn magnetite oreKonstäng, Dannemora, Sweden −0.43±0.03
DM-4Iron-skarn magnetite oreStrömsmalmen, Dannemora, Sweden−0.35±0.03
EJ092008Magnetite from a banded iron formation depositStriberg, Bergslagen, Sweden−0.57±0.03
Panzhihua Layered Intrusion
LJ1482Leuco-gbthe Lanjiahuoshan open pit, Panzhihua, China0.340.05
LJ1487aLeuco-gbthe Lanjiahuoshan open pit, Panzhihua, China0.360.03
LJ1487Leuco-gbthe Lanjiahuoshan open pit, Panzhihua, China0.320.01
LJ1434Ap-gbthe Lanjiahuoshan open pit, Panzhihua, China0.290.03
LJ1438Ap-gbthe Lanjiahuoshan open pit, Panzhihua, China0.40.04
LJ1453Ap-gbthe Lanjiahuoshan open pit, Panzhihua, China0.570.01
LJ1457Ap-gbthe Lanjiahuoshan open pit, Panzhihua, China0.30.03
LJ1458Ap-gbthe Lanjiahuoshan open pit, Panzhihua, China0.620.02
LJ1461Mela-gbthe Lanjiahuoshan open pit, Panzhihua, China0.530.04
LJ1462Mela-gbthe Lanjiahuoshan open pit, Panzhihua, China0.460.04
LJ1463Mela-gbthe Lanjiahuoshan open pit, Panzhihua, China0.490.03
LJ1430aMela-gbthe Lanjiahuoshan open pit, Panzhihua, China0.230.03
LJ1430Mela-gbthe Lanjiahuoshan open pit, Panzhihua, China0.250.04
LJ1427Mela-gbthe Lanjiahuoshan open pit, Panzhihua, China0.280.01
LJ1426Oxide orethe Lanjiahuoshan open pit, Panzhihua, China0.220.01
LJ1424Oxide orethe Lanjiahuoshan open pit, Panzhihua, China0.220.02
LJ1420Oxide orethe Lanjiahuoshan open pit, Panzhihua, China0.170.05
LJ1410Micro-gbthe Lanjiahuoshan open pit, Panzhihua, China0.280.05
LJ1407Micro-gbthe Lanjiahuoshan open pit, Panzhihua, China0.440.03
Gushan iron deposit, China
10GS_4_HmMassive ore(vesicular structure)Gushan iron deposit, China0.24
10GS_5_HmMassive ore (vesicular structure)Gushan iron deposit, China0.32
10GS_6_HmMassive ore (vesicular structure)Gushan iron deposit, China0.53
10GS_17_HmMassive ore (vesicular structure)Gushan iron deposit, China0.24
10GS_19_HmMassive ore (vesicular structure)Gushan iron deposit, China0.09
10GS_20_HmMassive ore (vesicular structure)Gushan iron deposit, China0.49
Xishimen iron deposit
XSM2-130503-01Massive magnetite oreXishimen iron deposit, China0.0080.04
XSM3-130504-06Massive magnetite oreXishimen iron deposit, China0.0340.04
XSM3-130504-09Massive magnetite oreXishimen iron deposit, China0.1150.04
XSM3-130504-14Massive magnetite oreXishimen iron deposit, China0.0470.04
XSM3-130504-16Massive magnetite oreXishimen iron deposit, China0.0540.04
XSM3-130504-18Massive magnetite oreXishimen iron deposit, China0.1110.04
XSM3-130504-33Massive magnetite oreXishimen iron deposit, China0.1130.04
XSM4-130506-22Massive magnetite oreXishimen iron deposit, China0.1050.052
XSM4-130506-23Massive magnetite oreXishimen iron deposit, China0.0790.052
XSM4-130506-26Massive magnetite oreXishimen iron deposit, China0.060.052
XSM4-130506-29Massive magnetite oreXishimen iron deposit, China0.0490.04
XSM4-130506-34Massive magnetite oreXishimen iron deposit, China0.0670.052
Table 4. Major composition of Grt (wt.%).
Table 4. Major composition of Grt (wt.%).
Sample NameSiO2TiO2Al2O3Cr2O3P2O5V2O3FeOMnOMgOCaONiONa2OK2OTotal
Grt140.20n.d.26.02n.d.n.d.n.d.10.170.220.0223.280.020.060.02100
Grt240.01n.d.23.88n.d.n.d.n.d.12.940.060.0223.040.010.04n.d.100
Grt340.140.0425.540.01n.d.n.d.11.020.02n.d.23.16n.d.0.030.0399.99
Grt439.590.0226.200.05n.d.n.d.10.680.420.0223.030.01n.d.n.d.99.95
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Li, R.; Su, S.; Wang, P. Chemistry and Fe Isotopes of Magnetites in the Orbicular Bodies in the Tanling Diorite and Implications for the Skarn Iron Mineralization in the North China Craton. Minerals 2025, 15, 1061. https://doi.org/10.3390/min15101061

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Li R, Su S, Wang P. Chemistry and Fe Isotopes of Magnetites in the Orbicular Bodies in the Tanling Diorite and Implications for the Skarn Iron Mineralization in the North China Craton. Minerals. 2025; 15(10):1061. https://doi.org/10.3390/min15101061

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Li, Ruipeng, Shangguo Su, and Peng Wang. 2025. "Chemistry and Fe Isotopes of Magnetites in the Orbicular Bodies in the Tanling Diorite and Implications for the Skarn Iron Mineralization in the North China Craton" Minerals 15, no. 10: 1061. https://doi.org/10.3390/min15101061

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Li, R., Su, S., & Wang, P. (2025). Chemistry and Fe Isotopes of Magnetites in the Orbicular Bodies in the Tanling Diorite and Implications for the Skarn Iron Mineralization in the North China Craton. Minerals, 15(10), 1061. https://doi.org/10.3390/min15101061

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