1. Introduction
Carbonate-hosted Pb–Zn deposits account for a high proportion of the world’s Pb–Zn resources and are mainly hosted by siliceous clastic rocks and carbonates that generally show no direct association with igneous rocks (Leach et al. Reference Leach, Sangster, Kelley, Large, Garven, Allen, Gutzmer and Walters2005, Reference Leach, Bradley, Huston, Pisarevsky, Taylor and Gardoll2010a; Mudd et al. Reference Mudd, Jowitt and Werner2017). These deposits display a broad range of affiliations to the enclosing host rocks and include stratiform, stratabound, and discordant ores (Leach et al. Reference Leach, Sangster, Kelley, Large, Garven, Allen, Gutzmer and Walters2005). The Himalayan-Tibetan and Zagros mountain ranges, as the youngest and most extensive continental-collision orogens in the Tethyan domain, host several major sediment-hosted Pb–Zn deposits, including the world-class Jinding, Huoshaoyun, Mehdiabad, and Angouran deposits (Fig. 1; Reynolds & Large, Reference Reynolds and Large2010; Rajabi et al. Reference Rajabi, Rastad and Canet2012, Reference Rajabi, Canet, Rastad and Alfonso2015; Hou & Zhang, Reference Hou and Zhang2015; Song et al. Reference Song, Hou, Liu and Zhang2017). Many of these deposits were only recently discovered and are poorly documented, especially in Iran. Although more than 300 carbonate-hosted Zn–Pb±Ba deposits/occurrences have been discovered in Iran, there is no general agreement regarding their genetic models of ore formation. It is plausible that these deposits range from sedimentary exhalative (SEDEX) to Mississippi Valley-type (MVT) (Rajabi et al. Reference Rajabi, Rastad and Canet2012, Reference Rajabi, Canet, Rastad and Alfonso2015, Reference Rajabi, Alfonso, Canet, Rastad, Niroomand, Modabberi and Mahmoodi2020). The majority of these deposits are hosted in carbonates of Devonian to Cretaceous age (Ehya, Reference Ehya2014). They largely occur in the Malayer-Esfahan Metallogenic Belt and Yazd-Anarak Metallogenic Belt (YAMB), which is located in the Yazd Block along the northern margin of the Central-East Iranian microcontinent (Rajabi et al. Reference Rajabi, Rastad and Canet2012) (Fig. 2). Those in the YAMB include many world-class Pb–Zn deposits, such as the Mehdiabad Zn–Pb–Ba–(Cu–Ag) deposit (45.2 Mt oxide @ 7.15% Zn and 2.47% Pb) and 116.5 Mt sulphide (7.3% Zn and 2.3% Pb) (Maghfouri et al. Reference Maghfouri, Hosseinzadeh, Choulet, Alfonso, Azim Zadeh and Rajabi2019, Reference Maghfouri, Hosseinzadeh and Choulet2020a, Reference Maghfouri, Hosseinzadeh, Lentz, Tajeddin, Movahednia and Shariefi2021). The Yazd block encompasses major Zn-Pb deposits including the Nakhlak Pb–(Ag) (Jazi et al. Reference Jazi, Karimpour and Malekzadeh Shafaroudi2017), Darreh-Zanjir Zn–Pb (Maghfouri & Choulet, Reference Maghfouri and Choulet2021), Mansourabad-Farahabad Zn–Pb–(Ag) (Maghfouri & Hosseinzadeh, Reference Maghfouri and Hosseinzadeh2018; Maghfouri et al. Reference Maghfouri, Hosseinzadeh, Lentz and Choulet2020b), Hovzesefid and Anjireh Zn–Pb (Rajabi et al. Reference Rajabi, Rastad and Canet2012) and Chahmileh Pb–Zn deposits (Technoexport, report, 1984) (Fig. 2). Most Zn-Pb deposits in this region are considered to be MVT deposits, formed in a platform carbonate succession, typically in passive margins (Rajabi et al. Reference Rajabi, Rastad and Canet2012).
Figure 1. Distribution of the major sediment-hosted Pb–Zn deposits from China to Iran in the Tethyan domain (modified from Hou & Zhang, Reference Hou and Zhang2015; Song et al. Reference Song, Liu, Hou, Fard, Zhang and Zhuang2019). Paleo-Tethyan sutures (green curves): (a) North Turkey; (b) Lesser Caucasus; (c) Kopet Dagh; (d) North Pamir; (e) Kunlun; (f) Garzȇ-Litang; (g) western Jinshajiang; (h) eastern Jinshajiang; (i) Longmu Co-Shuanghu; (j) Changning-Menglian; (k) Inthanon; (l) Bentong-Raùb; (m) Ailaoshan. Neo-Tethyan sutures (red curve): 1-Izmir-Ankara-Erzincan; 2-Alborz; 3-Zagros; 4-Nain; 5-Sabzevar; 6-Sistan; 7-Bela-Waziristan-Quetta; 8-Bangonghu-Nujiang; 9-Shan Boundary; 10-Indus-Yarlung-Zangbo; 11-Burma. MVT: Mississippi Valley-Type, CD: Clastic-Dominated, CRD: Carbonate Replacement Deposit.
Figure 2. Distribution map of sediment-hosted Zn–Pb (±Ag ± Cu ± Ba) deposits in the Malayer-Esfahan Metallogenic Belt and the Yazd-Anarak Metallogenic Belt (modified after Rajabi et al. Reference Rajabi, Rastad and Canet2012). Most of the deposits occur on both sides of the Nain-Baft back-arc basin, bordered by the Nain-Baft ophiolites. PB: Posht-e Badam Block, SSZ: Sanandaj–Sirjan Zone, Za: Zagros fold and thrust belt.
The Chahmileh Pb-Zn deposit is located 30 km northeast of Anarak and 12 km south of the Nakhlak Pb-(Ag) deposit, which is one of the oldest mining areas in Iran. Technoexport Co. carried out detailed regional-scale geology and structural investigation in the Anarak area, including the Chahmileh prospect from 1975 to 1985. A detailed exploration program was conducted by Kan-Azin Mining Consultant Company from 2014 to 2017 drilling five boreholes (total length of 1500 m). Previous studies of the deposit focused on geophysical and geochemical exploration features (Technoexport, report, 1984), while the metallogenic mechanism of this deposit remains the subject of considerable debate. However, the sulphur and metal sources, evolution of the ore-forming fluids and the genesis of the deposit remain poorly understood.
In this contribution, we present the results of a comprehensive investigation of the Chahmileh deposit that involves geological field studies, ore mineralogy, fluid inclusions, stable (C–O–S) and radiogenic (Pb) isotope compositions. The aims of the study are to (1) investigate the physical–chemical conditions and metal transport mechanisms; (2) determine the characteristics of the ore-forming fluids and sulphide precipitation mechanism; and (3) evaluate the metal sources and discuss the ore genesis. Studies of the Chahmileh deposit will help to elucidate the ore-forming processes in comparable geologic settings (e.g., MVT and SEDEX), especially where mineralization is associated with subsequent modifications by later hydrothermal activity. In order to facilitate prospecting in the area and improve our understanding of the regional metallogeny in the YAMB, it is crucial to gain further insight into the formation of the Chahmileh deposit.
2. Regional geological setting
The geotectonic history of Iran was affected by the development and evolution of three Tethyan Oceans: The Proto-Tethys Ocean in the Late Neoproterozoic-Cambrian (Pan-African orogeny), the Paleo-Tethys Ocean in the Paleozoic (Cimmerian orogeny) and the Neo-Tethys Ocean during the Mesozoic and Cenozoic (Alpine orogeny) (Bagheri & Stampfli, Reference Bagheri and Stampfli2008). The Chahmileh Pb-Zn deposit is situated in the Anarak region, in the northwest corner of the Central Iran Zone (CIZ). The CIZ is the most complicated and largest geological unit in Iran and is an area of continuous continental deformation in response to the ongoing convergence between the Arabian (Gondwanan) and Turan (Eurasian) plates. A series of tectonic events that shaped early evolution of CIZ is associated with the Peri-Gondwanan or Proto-Tethyan episode. At least two further episodes of orogenic activity, one in the Early Triassic and another in the Late Tertiary, affected the CIZ before its final incorporation into the Alpine–Himalayan Belt (Stöcklin, Reference Stöcklin, Burke and Drake1974). The CIZ geology consists of Precambrian to Miocene sedimentary rocks, Palaeozoic to Cenozoic ultramafic-acid igneous rocks and Palaeozoic to Mesozoic metamorphic rocks (Balini et al. Reference Balini, Nicora, Berra, Garzanti, Levera, Mattei, Muttoni, Zanchi, Bollati, Larghi, Zanchetta, Salamati, Mossavvari, Brunet, Granath and Wilmsen2009). The CIZ was a stable platform during the Palaeozoic until Late Triassic movements resulted in the formation of a series of horsts and grabens (Zanchi et al. Reference Zanchi, Zanchetta, Garzanti, Balini, Berra, Mattei, Muttoni, Brunet, Wilmsen and Granath2009). Major structural trends were created during the Mesozoic when a contiguous platform of the CIZ was divided into small segments (e.g., Stöcklin, Reference Stöcklin1968; Ramezani & Tucker, Reference Ramezani and Tucker2003). The CIZ consists of three major crustal domains from east to west: the Lut, Tabas, and Yazd Blocks (e.g., Alavi, Reference Alavi1991) which are separated by a series of intersecting regional-scale faults (Berberian, Reference Berberian, Gupta and Delany1981). It is delimited to the north by an E-W trending left-lateral Doruneh (or Great Kavir) fault that interacted with dextral N-S trending faults (Nozaem et al. Reference Nozaem, Mohajjel, Rossetti, Della Seta, Vignaroli, Yassaghi, Salvini and Eliassi2013), inherited from the Paleozoic evolution. The Doruneh fault, which is one of the longest and most prominent faults in Iran (Wellman, Reference Wellman1966; Farbod et al. Reference Farbod, Bellier, Shabanian and Abbassi2011), plays an important role in the regional tectonics (Torabi, Reference Torabi2010).
The Chahmileh Pb-Zn deposit is located in the Anarak Metamorphic Complex (AMC) in the CIZ. The AMC consists of intricate polyphase thrust stacks, containing low-grade metapelites, metabasites, and marbles with a greenschist to blueschist metamorphic overprint. It is associated with slivers of ultramafic rocks and metamorphosed-pillow lavas that formed under high-pressure/low-temperature conditions (M Sharkovski et al., report, 1984; Bagheri & Stampfli, Reference Bagheri and Stampfli2008; Zanchi et al. Reference Zanchi, Zanchetta, Garzanti, Balini, Berra, Mattei, Muttoni, Brunet, Wilmsen and Granath2009; Buchs et al. Reference Buchs, Bagheri, Martin, Hermann and Arculus2013; Zanchi et al. Reference Zanchi, Malaspina, Zanchetta, Berra, Benciolini, Bergomi, Cavallo, Javadi and Kouhpeyma2015). It is in tectonic contact with other metamorphic complexes and sedimentary successions of various ages and palaeogeographic affinities. To the west, the Great Kavir-Doruneh fault system represents the contact between AMC and Cretaceous ophiolites that border the Central-East Iranian Microcontinent (Ghasemi & Talbot, Reference Ghasemi and Talbot2006). The southern limit of the AMC coincides with the Palaeozoic to Mesozoic sequences of the Yazd block, which geologically are similar to the Alborz region (Wendt et al. Reference Wendt, Kaufmann, Belka, Farsan and Bavandpur2005; Leven & Gorgij, Reference Leven and Gorgij2006; Zanchi et al. Reference Zanchi, Malaspina, Zanchetta, Berra, Benciolini, Bergomi, Cavallo, Javadi and Kouhpeyma2015). Towards the north, it is bordered by the non-metamorphic Nakhlak ophiolite and associated sedimentary complex (Balini et al. Reference Balini, Nicora, Berra, Garzanti, Levera, Mattei, Muttoni, Zanchi, Bollati, Larghi, Zanchetta, Salamati, Mossavvari, Brunet, Granath and Wilmsen2009). Sedimentary rocks found in the Nakhlak area include a 2400 m-thick forearc succession of turbiditic, shallow-marine and fluvial deposits (i.e. ‘Nakhlak Group’ after Balini et al. Reference Balini, Nicora, Berra, Garzanti, Levera, Mattei, Muttoni, Zanchi, Bollati, Larghi, Zanchetta, Salamati, Mossavvari, Brunet, Granath and Wilmsen2009) recording the erosion of a nearby volcanic arc and metamorphic basement, presumably the AMC (Bagheri & Stampfli, Reference Bagheri and Stampfli2008; Balini et al. Reference Balini, Nicora, Berra, Garzanti, Levera, Mattei, Muttoni, Zanchi, Bollati, Larghi, Zanchetta, Salamati, Mossavvari, Brunet, Granath and Wilmsen2009; Zanchi et al. Reference Zanchi, Zanchetta, Garzanti, Balini, Berra, Mattei, Muttoni, Brunet, Wilmsen and Granath2009). Forearc tectonic features of the Nakhlak area are characterized by supra-subduction and boninitic gabbros dated at ∼387 Ma (S Bagheri, unpub. PhD thesis, Univ. Lausanne, 2007; Bagheri & Stampfli, Reference Bagheri and Stampfli2008). The contact between the AMC and Nakhlak complex is not exposed, so the timing relationship between them is unclear (Balini et al. Reference Balini, Nicora, Berra, Garzanti, Levera, Mattei, Muttoni, Zanchi, Bollati, Larghi, Zanchetta, Salamati, Mossavvari, Brunet, Granath and Wilmsen2009; Zanchi et al. Reference Zanchi, Zanchetta, Garzanti, Balini, Berra, Mattei, Muttoni, Brunet, Wilmsen and Granath2009). To the east of the AMC, the Jandaq Metamorphic Complex, containing medium to possibly high-grade)likely Carboniferous or pre-Carboniferous(metamorphic rocks, was intruded by early Mesozoic granites and pegmatite (Bagheri & Stampfli, Reference Bagheri and Stampfli2008; Berra et al. Reference Berra, Zanchi, Angiolini, Vachard, Vezzoli, Zanchetta, Bergomi, Javadi and Kouhpeyma2017).
The AMC is composed of several subunits, including the Morghab, Chah Gorbeh, Patyar, Lakh Marble, Palhavand Gneiss, Doshak and Bayazeh complexes (Fig. 3; M Sharkovski et al., report, 1984; Bagheri & Stampfli, Reference Bagheri and Stampfli2008; Zanchi et al. Reference Zanchi, Zanchetta, Garzanti, Balini, Berra, Mattei, Muttoni, Brunet, Wilmsen and Granath2009, Reference Zanchi, Malaspina, Zanchetta, Berra, Benciolini, Bergomi, Cavallo, Javadi and Kouhpeyma2015). They display heterogeneous structural and metamorphic histories and are cross-cut by small mafic to felsic intrusive bodies, mainly trondhjemite dykes and Late Permian stocks (Bagheri & Stampfli, Reference Bagheri and Stampfli2008). The AMC is interpreted to be an allochthonous crustal fragment that was part of an accretionary wedge developed along the southern Eurasian margin, in the hanging wall of the Palaeo-Tethys subduction zone (Zanchi et al. Reference Zanchi, Malaspina, Zanchetta, Berra, Benciolini, Bergomi, Cavallo, Javadi and Kouhpeyma2015) and preceding the collision of the Iran plate with Eurasia during the Cimmerian orogenic event (Zanchetta et al. Reference Zanchetta, Malaspina, Zanchi, Benciolini, Martin, Javadi and Kouhpeyma2017).
Figure 3. Geological map of the northern part of the Anarak Metamorphic Complex with N–S trending cross-section (A-B) (modified after Zanchi et al. Reference Zanchi, Malaspina, Zanchetta, Berra, Benciolini, Bergomi, Cavallo, Javadi and Kouhpeyma2015). Radiometric ages of various rocks are adopted from Bagheri & Stampfli (Reference Bagheri and Stampfli2008).
The Anarak complex is cross-cut by an array of E-W and NW-SE striking faults that are truncated to the NW against the southern terminal branch of the Doruneh fault system. These faults cut the Cenozoic sedimentary deposits that overly the metamorphic and ultramafic basement. The main structure in the area is the 45 km long, E-W to NNW-SSE striking, SW dipping Ashin thrust fault, which juxtaposes the hanging wall of the metamorphic basement, and ultramafic mantle rocks of the AMC with the Cenozoic sedimentary and volcanic rocks in its footwall (Javadi et al. Reference Javadi, Esterabi Ashtiani, Guest, Yassaghi, Ghassemi, Shahpasandzadeh and Naeimi2015).
3. Ore deposit geology
The Chahmileh area mainly consists of ultramafic rocks and the Carboniferous to Permo-Triassic low-grade metamorphic sequence of the Morghab and Chah Gorbeh complexes (Fig. 4). Ultramafic rocks are composed of coarse-grained hornblende gabbro in close association with serpentinized ultramafic rocks and minor blueschist and small trondhjemite stocks (Fig. 5a). The blueschists have an ocean island basalt geochemical affinity (Bagheri & Stampfli, Reference Bagheri and Stampfli2008; Torabi, Reference Torabi2011), whereas the trondhjemites have a supra-subduction origin and yielded a U-Pb zircon age of 262.3 ± 1.0 Ma (Bagheri & Stampfli, Reference Bagheri and Stampfli2008; Torabi, Reference Torabi2012).
Figure 4. A simplified geological map of the Chahmileh Pb-Zn deposit showing the mineralization and dolomitic marble host rock of the Chah Gorbeh Complex (modified after Kan-Azin Mining Consultant Company, report, 2014).
Figure 5. Field photographs and photomicrograph of representative rocks at the Chahmileh. (a) Serpentinite southeast of Mazra-e Deraz, (b) Quartz veins in micaschist and phyllite of the Morghab Complex (Cmr sch), (c) Muscovite chlorite schist unit (Tch sch) and its contact with dolomitic marble (Tch mb) of the Chah Gorbeh Complex, (d) Dolomitic marble (Tch mb) of the Chah Gorbeh Complex; e Muscovite chlorite schist, showing nematoblastic and granoblastic textures, f NW-trending normal fault in the Kuh-e Mileh tunnel. F: Fault.
The Morghab Complex is present throughout the area in close association with the Chah Gorbeh Complex (Figs. 3, 4). It consists of a monotonous assemblage of low-grade metapelites, varying from phyllite to mica schist alternating with quartzite and metabasites layers, and thin intercalations of marble (Fig. 5b). Garnet-biotite-mica schists locally occur close to the Chah Karbouzeh area (Fig. 3). Bagheri and Stampfli (Reference Bagheri and Stampfli2008) reported an Ar-Ar age of 318.9 ± 1.63 Ma for metamorphism of the Morghab Complex.
The Chah Gorbeh Complex consists of two units: (1) quartzite-rich phyllite, micaschist and metabasites and (2) interlayers of thick-bedded dolomitic marble (Figs. 3, 4, 5c–e). These rock units are very similar in composition to those in the Morghab Complex. Meta-cherts occur within the marble layers and along their contacts. Metabasite, which was metamorphosed to the greenschist facies is intercalated with this unit. The Chah Gorbeh Complex is generally concordant with the Morghab Complex, at least around Chah Gorbeh mountain (Fig. 3). Zanchi et al. (Reference Zanchi, Malaspina, Zanchetta, Berra, Benciolini, Bergomi, Cavallo, Javadi and Kouhpeyma2015) reported a tectonic contact with the Lakh Marble in the Doldol mountain (southern Anarak) and around the Chah Gorbeh mountain (Fig. 3). Metabasalt and meta-greywacke samples from the Chah Gorbeh Complex yield an age of 232.8 ± 2.35 Ma using Ar-dating of stilpnomelane (Bagheri & Stampfli, Reference Bagheri and Stampfli2008; Buchs et al. Reference Buchs, Bagheri, Martin, Hermann and Arculus2013). Previous K-Ar radiometric dating on mineral separates and bulk rock samples range in age from 420 to 208 Ma with a main cluster between 375 and 300 Ma (M Sharkovski et al., report, 1984).
In the Chahmileh prospect area, NW-trending faults dipping towards the NE are the most prominent structural feature (Fig. 5f) in association with NE- to N-trending faults that dip steeply to the NW or SE. These normal faults have a moderate sinistral strike-slip component. Exposed reverse faults in the Mazra-e Deraz area displaced ultramafic units over the younger sequences.
A thick-layered dolomitic marble unit of the Chah Gorbeh Complex with a maximum exposed thickness of 700 m is the main host rock of the Chahmileh Pb–Zn deposit (Fig. 4). Field evidence indicates that ore mineralization is controlled mainly by faults (Fig. 6a, b). In the study area, there are three zones of Pb-Zn mineralization, Kuh-e Mileh, Seilacho and Mazra-e Deraz.
Figure 6. Photograph showing ore textures and mineralization features at the Chahmileh deposit. (a) Mineralization located in the footwall of the NW-trending normal fault (F), (b) Mineralization in fault zone, (c) Quartz-galena vein hosted in the dolomitic marble unit (Tch mb) of the Chah Gorbeh Complex, (d) Vein-type galena mineralization associated with minor malachite at the Kuh-e Mileh tunnel, (e) Clasts of dolomitic host rock replaced by galena, (f) Dolomite breccia clasts associated with quartz-galena, (g) Cerussite and (h) Mimetite in the oxidized zone. Abbreviation of minerals adopted from Warr (Reference Warr2021): Cer: Cerussite, Dol: Dolomite, Gn: Galena, Mlc: Malachite, Mim: Mimetite, Qz: Quartz.
Kuh-e Mileh is the largest deposit, located in the eastern part of the Chahmileh area (Fig. 4), with a reserve of ∼1 Mt @ 2.15% Pb + Zn (Kan-Azin Mining Consultant Company, report, 2015). The ore body is stratabound with a lenticular shape. Mineralization mainly occurs in structurally controlled open-space fillings and locally as a replacement of the host rock where breccia- and vein-type ores are the major style of mineralization (Fig. 6c–f). Breccia clasts are poorly sorted, angular to subangular in shape and range in size from a few centimetres to a few tens of centimetres in size. Quartz, galena and haematite are the main minerals in the breccia cement (Fig. 6e, f). Disseminated and open-space filling is prominent and likely formed by hydrothermal fluid percolation or replacement. In some cases, disseminated galena and pyrite fill the open-space between dolomitic crystals in the altered host rock. Massive ore is not common and is only rarely found as thick veins of galena (Fig. 6d). Ore mineral assemblages are relatively simple and are dominated by galena and sphalerite with minor amounts of chalcopyrite and pyrite. Oxidation and weathering processes lead to the dissolution and alteration of sulphide minerals and development of non-sulphide ore that consists of cerussite, hemimorphite, mimetite, wulfenite, smithsonite, malachite and iron oxy-hydroxides (Fig. 6g, h).
Galena is the main ore mineral and is characterized by medium- to coarse-grained (0.1 mm to >1 cm in size), subhedral to euhedral crystals mainly as disseminations (Figs. 7a, b, 10c), veins and veinlets (Fig. 7c, d) and as grains in breccia (Fig. 7f). Disseminated galena formed at the same time as intergrowths with other sulphides (Fig. 7b). Sphalerite is found as subhedral to anhedral inclusions (40–60 μm in size) within galena crystals (Fig. 8a, i) whereas chalcopyrite is a minor ore mineral and occurs as disseminated, medium to coarse-grained (0.1–3 mm), subhedral to anhedral crystals. Locally, chalcopyrite was replaced by secondary covellite, chalcocite, goethite and malachite (Fig. 7b, e). Pyrite is a minor phase and occurs as euhedral and/or subhedral fine grains (2–5 µm in size), disseminated in the host rock and among other ore minerals (Fig. 7a) where it is partially or completely altered to goethite.
Figure 7. Photomicrographs of sulphide, non-sulphide, and gangue minerals at the Chahmileh deposit. (a) Disseminated pyrite accompanied by subhedral disseminated galena (Gn-1) partly replaced by cerussite (Crt-1) (PPL), (b) Intergrowth of chalcopyrite, galena (Gn-1), and chalcocite replacing chalcopyrite (PPL), (c) Galena (Gn-3)-quartz (Qz-2) veinlet within dolomitic marble (Dol-2) (XPL), (d) Conjugate galena (Gn-3) veinlets (PPL), (e) Secondary minerals formed on rims of chalcopyrite (PPL), (f) Covellite bladed crystals and cerussite (Crt-1) replacing galena (Gn-2) (PPL), (g) Wulfenite crystals (XPL), (h) Banded haematite+goethite associated with calcite (Cal-1) and malachite (XPL), (i) Acicular baryte (XPL). Abbreviation of minerals adopted from Warr (Reference Warr2021): Brt: Baryte, Cal: Calcite, Ccp: Chalcopyrite, Cc: Chalcocite, Cer: Cerussite, Cv: Covellite, Dol: Dolomite, Gth: Goethite, Gn: Galena, Hem: Haematite, Mlc: Malachite, Py: Pyrite, Qz: Quartz, Wul: Wulfenite.
Figure 8. BSE images of sulphide and non-sulphide minerals at the Chahmileh deposit. (a) Sphalerite inclusions in galena (Gn-3) and replacement of cerussite (Crt-1) on galena rims, (b) Galena (Gn-3) cleavage and cerussite (Crt-1) replacement, (c) Galena (Gn-1) boundary replacement by cerussite (Crt-1) and litharge, (d) Mimetite with open-space filling texture in dolomitic marble, (e) Chalcopyrite with platy hemimorphite and disseminated cerussite (Crt-2), (f) Platy euhedral crystals of hemimorphite and second generation of cerussite (Cer-2), (g) Hemimorphite inclusions within galena (Gn-3), (h) Assemblage of hemimorphite, willemite and colloform smithsonite, (i) Sphalerite inclusions within galena (Gn-3) and sphalerite replacement by smithsonite, (j) Needle shape radial haematite and pyramidal malachite as open-space filling texture, (k) Atacamite inclusions within linarite, (l) Coronadite subhedral crystals. Abbreviation of minerals adopted from Warr (Reference Warr2021): Ata: Atacamite, Ccp: Chalcopyrite, Cer: Cerussite, Cor: Coronadite, Dol: Dolomite, Gn: Galena, Hem: Haematite, Hmp: Hemimorphite, Lna: Linarite, Mlc: Malachite, Lit: Litharge, Mim: Mimetite, Qz: Quartz, Smt: Smithsonite, Sp: Sphalerite, Wlm: Willemite.
Supergene minerals are divided into two categories: sulphide and non-sulphide minerals. Covellite is the most abundant supergene sulphide mineral where it occurs as bladed crystals (10–200 µm in size) and is derived from the alteration of chalcopyrite along fractures (Fig. 7e, f). Cerussite, the most common oxidized lead mineral in the deposit, occurs as fine-grained cerussite that replaced galena (Figs. 7a, d, f, 8a–c, i) and coarse-grained cerussite filling open fractures and cavities (Fig. 8e, f). Mimetite (Pb5(AsO4)3Cl) occurs in the supergene mineralization and is spatially associated with secondary Pb minerals (Fig. 8d). Wulfenite (PbMoO4) is associated with cerussite as open-space filling textures (Fig. 7g), and is likely an oxidation product of molybdenum in galena (e.g., Graton & Harcourt, Reference Graton and Harcourt1935; Takahashi, Reference Takahashi1960). Supergene Zn minerals are generally less common than secondary Pb minerals and consist of hemimorphite and smithsonite with minor, hydrozincite and willemite. Hemimorphite has a medium to coarse-grained, elongate shape, relatively high relief and distinct longitudinal cleavage, with a mosaic-type and blocky texture, up to 3 mm in size (Fig. 8e–h). Smithsonite is generally less abundant than hemimorphite and formed in open-space cavities in the carbonate rocks where it replaced sphalerite (Fig. 8h, i). Haematite occurs in banded, scaffold, reticulate, colloform, needle and radial textures (Figs. 7h, 8j). Other supergene minerals associated with the non-sulphide ore identified by scanning electron microscopy and energy dispersive spectroscopy (SEM-EDS) are litharge (PbO) (Fig. 8c), minium (Pb3O4), linarite (PbCu[(OH2)-SO4] (8k), atacamite (Cu2Cl(OH)3) (8k), coronadite (8l), malachite (Fig. 8j), azurite, cuprite, chrysocolla and magnetite.
Gangue minerals are dominated by quartz, dolomite, calcite and baryte. Quartz typically occurs as fine to coarse-grained (< 20 μm to >1 cm in size), anhedral to euhedral crystals and cryptocrystalline to holocrystalline forms, and locally contains carbonate inclusions. Dispersed platy and prismatic crystals, subparallel and radial growths and swallow-tail bundles and stellate aggregates of baryte, a few mm to cm in length, fill open spaces and vugs in the host rocks (Fig. 7i). Calcite occurs as a minor gangue mineral formed during the main and late-stages of mineralization.
Based on mineral assemblages, ore textures and cross-cutting relationships, the ore-forming process at the Chahmileh deposit can be divided into three stages: (1) sedimentary and diagenetic stage (pre-ore stage), (2) hydrothermal mineralization (main-ore stage) and (3) post-ore stage (Fig. 9).
Figure 9. Paragenetic sequence of the Chahmileh deposit, thickness of line representing the minerals frequency.
Figure 10. Photomicrographs of various alterations at the Chahmileh deposit. (a) Type I (Dol-1) and III dolomite (Dol-3) (XPL), (b) Medium to coarse-grained, subhedral to euhedral type II dolomite (Dol-2) (XPL), (c) The second type of dolomite (Dol-2), which is replaced by disseminated galena (Gn-1) (XPL) (d) Silicification alteration (XPL). Abbreviation of minerals adopted from Warr (Reference Warr2021): Dol: Dolomite, Gn: Galena.
The sedimentary and diagenetic stage (pre-ore stage) is represented by dolomite crystals formed during diagenesis, prior to the dolomitization associated with the main-ore stage. This stage is locally difficult to identify due to hydrothermal overprinting by the main-stage minerals. The pre-ore stage is also associated with the formation of dolomite that replaced calcite in limestones, a process that generates porosity for further fluid flow and deposition of other minerals. The sedimentary and diagenetic stage also contains very fine-grained disseminated pyrite.
Hydrothermal mineralization (main-ore stage) is the main ore-forming stage in terms of both volume and grade.
Post-ore stage includes barren calcite, quartz or dolomite veins and veinlets, which cross-cut stages 1 and 2 and fill secondary fractures and voids, together with supergene ores, derived from the oxidation of primary ores in surficial environments.
Hydrothermally altered rocks occur together with sulphide mineralization mainly along and/or near faults or fractures, indicating that both the mineralized and altered rocks are structurally controlled. Dolomitization and silicification are the main types of alteration, where it is commonly accompanied by brecciation. Based on grain size, shape and colour, dolomite can be classified into three types (Dol-1, Dol-2 and Dol-3). Dol-1, which mainly occurs as cement, consists of a mosaic of grey unimodal, nonplanar dolomite crystals less than 20 μm size, formed by the late diagenetic replacement of limestone (Fig. 10a). The most pervasive type of dolomite, Dol-2, is characterized by medium- to coarse-grained, subhedral to euhedral crystals. Euhedral, polymodal crystals are rarely observed but are characterized by cloudy centres and clear rims. This zonation reveals fluctuations in the chemistry of the dolomitising fluid. The Dol-2 intercrystalline porosity is infilled by galena and quartz and/or is replaced by galena and quartz (Fig. 10b, c). The Dol-3 occurs as veins containing anhedral, nonplanar and medium-grained crystals (Fig. 10a).
Silicification formed massive to vuggy bodies of microcrystalline to cryptocrystalline quartz generally as a cement in mineralized breccia or individual veins. There is a close spatial association between sulphides and quartz, with sulphides formed as disseminations and veinlets within the quartz bodies (Fig. 10c, d).
4. Sampling and analytical methods
One hundred and fifty representative samples of the Pb–Zn ore and host rocks were collected, and after preparing thin and polished sections, they were examined by a ZEISS Axioplan-2 transmitted and reflected light microscope. After ore petrography, sulphide and non-sulphide minerals were checked by SEM-EDS using a VP-SEMEDS ZEISS 3700 at the Central Laboratory of Kharazmi University (Tehran, Iran). Mineralogical abbreviations used throughout the manuscript are according to Warr (Reference Warr2021).
4.1. Fluid inclusion analysis
Petrographic and microthermometric studies were carried out on fluid inclusions in dolomite (n = 3) and quartz (n = 2) from the main-ore stage and late calcite veins (n = 1) of the post-ore stage. Microthermometry of fluid inclusions was carried out in the Geochemistry Department of Kharazmi University (Tehran, Iran), using a Linkam THMS600 heating–freezing stage (−190 to +600 °C) mounted on a ZEISS Axioplan-2 microscope. Temperature was calibrated using Synflinc synthetic fluid inclusion standards. The estimated accuracy is ±0.5 °C for temperatures below 100 °C and ±1.0 °C for temperatures in the range of 100–600 °C. Heating/cooling rates were restricted to 5–10 °C/min and were reduced to 0.1–0.5 °C/min near phase transformations. Freezing measurements were conducted before heating measurements. The following parameters were measured in aqueous inclusions (based on the nomenclature of Diamond, Reference Diamond, Samson, Anderson and Marshall2003); first ice-melting temperature (Tfm), final ice-melting temperature (Tmice) and total homogenisation temperature (Th). Salinity is expressed as wt.% NaCl eq. calculated from (Tmice) using the equations of Bodnar (Reference Bodnar1993) for aqueous inclusions. Molar volumes, compositions and density were calculated using the FLINCOR software (Brown, Reference Brown1989).
4.2. Stable isotope analysis
Eleven dolomite samples of the main-ore stage, deposited contemporaneously with galena, were analysed for carbon (δ13C) and oxygen (δ18O) isotopic composition. Samples were analysed using a Kiel III device connected to a Finnigan MAT 252 isotope ratio mass spectrometer in the Department of Geological Sciences at the University of Florida. Carbon isotope data expressed in the δ notation in per mil (‰) relative to Vienna Pee Dee Belemnite (VPDB), whereas those for oxygen isotopes are reported relative to Vienna Standard Mean Ocean Water (VSMOW). The precision of the technique was measured with an internal standard of Carrera Marble calibrated with NSB-19 and found to be ± 0.04‰ for δ18O and ± 0.08‰ for δ13C.
Eight sulphide samples including galena (n = 4), chalcopyrite (n = 2) and sphalerite (n = 2) from the main-ore stage and baryte (n = 2) from the post-ore stage were selected for sulphur isotope analyses. Separates of sulphide minerals were prepared by handpicking under a binocular microscope to achieve a purity of > 99%. The δ34S values were measured in SO2 gas using a continuous-flow gas-ratio mass spectrometer (Thermo Quest Finnigan Delta Plus XL) at the Environmental Isotope Laboratory of Arizona University. Sulphur isotopic composition was determined after combustion at 1030 °C in oxygen (O2 or V2O5) using an elemental analyser (Costech) coupled to the mass spectrometer. Standardization was based on international standards OGS-1 and NBS123, and several other sulphide and sulphate materials that have been compared between laboratories. The data are presented in delta (δ) notation as per mill (‰) deviations relative to the Vienna Canyon Diablo Troilite (VCDT) standard for sulphur. Calibration is linear in the range –10 to +30‰. Precision is estimated to be ± 0.15‰ or better (1σ), based on repeated measurement of internal standards.
4.3. Lead isotope analysis
Lead isotope compositions of galena (n = 2, stage I; n = 2, stage II and n = 2, stage III of main-ore stage) were measured at the University of Oslo. Lead was leached from the samples with dilute acid, mixed with phosphoric acid and silica gel and loaded directly on outgassed Re filaments. Ratios were measured by thermal ionization mass spectrometry on a MAT262 instrument using multiple Faraday cups in static mode (Corfu, Reference Corfu2004). Data are corrected for fractionation of 0.10 ± 0.06% per atomic mass unit. Reproducibility of the fractionation (based on NBS982) is propagated into the uncertainty of the corrected ratios. The Isoplot 4.1 programme was used for plotting Pb-isotope results.
5. Results
5.1. Fluid inclusions petrography
Fluid inclusions were studied in dolomite (Dol-2), quartz (Qz-1 and Qz-2) and calcite (Cal-2) from the main-ore and post-ore stages. No workable fluid inclusions were identified in minerals from the pre-ore stage. Based on the criteria of Roedder (Reference Roedder1984), there are both primary and secondary fluid inclusions in the samples. Fluid inclusions that are clustered or isolated are considered primary, whereas those aligned along microfractures in transgranular trails were designated as secondary, which are mostly less than 2 μm in size. Only primary fluid inclusions associated with the main and post-ore stages of mineralization processes were selected for fluid inclusion studies. Fluid inclusion shapes are rectangular, elliptical, circular, rod-shaped or elongated and rarely irregular. Diameter of the investigated inclusions was in range of 5–30 μm, mostly around 5–15 μm. With exception of a few single-phase aqueous and vapour inclusions, almost all inclusions are two-phase (L + V), liquid-rich with a 10–20 volume per cent of vapour bubble that homogenized to liquid upon heating (Fig. 11).
Figure 11. Fluid inclusions photomicrographs in the Chahmileh deposit. (a) Primary cluster of fluid inclusions relative to the quartz grain boundary, (b) Primary cluster of fluid inclusions hosted by calcite, (c) A row of secondary fluid inclusions in calcite, (d) Primary liquid-rich two-phase fluid inclusions with secondary fluid inclusions in quartz, (e) Association of liquid-rich two-phase (L+V) with liquid monophase (L) fluid inclusions hosted in quartz, (f) Primary liquid-rich two-phase fluid inclusions in dolomite.
5.2. Microthermometry
Microthermometric data of the primary fluid inclusions are listed in Table 1 and shown in Fig. 12. The Tfm of primary inclusions in dolomite crystals from the main-ore stage I varies from −35.0 to −29.0 °C (avg. = −31.4 °C, n = 24), suggesting the presence of appreciable amount of CaCl2 in addition to NaCl and KCl (Van den Kerkhof & Hein, Reference Van den Kerkhof and Hein2001). Inclusions homogenized into the liquid phase between 133.0 and 248.0 °C (avg. = 205.3 °C, n = 27). The Tmice value for these inclusions varies from −17.7 to −12.3 °C (avg. = −14.3 °C, n = 27), corresponding to a salinity of 16.2 to 20.8 wt.% NaCl eq. (avg. = 18.0 wt.%, n = 27). Density of these inclusions is 0.89−1.04 g/cm3.
Table 1. Microthermometric measurements of fluid inclusions from the Chahmileh deposit
Abbreviations: Data are reported as averages of fluid inclusions assemblages; n = the number of available analyses, Tfm = first ice-melting temperature, Tmice = final ice-melting temperature, Th = homogenization temperature, LV = liquid-rich type, (L) = homogenization to liquid. nd = not detected.
Figure 12. Histograms of total homogenization temperatures and calculated salinities based on microthermometric data of fluid inclusions from different mineralization stages.
Primary inclusions in quartz crystals from the main-ore stage II show Tfm of −38.0 to −32.0 °C (avg. = −34.3 °C, n = 17), suggesting the presence of appreciable amounts of CaCl2 in addition to NaCl and KCl (Van den Kerkhof & Hein, Reference Van den Kerkhof and Hein2001). Inclusions homogenized into the liquid phase between 136.0 and 211.0 °C (avg. = 173.9 °C, n = 28). The Tmice value for these inclusions varies from −16.8 to −12.2 °C (avg. = −14.2 °C, n = 28), corresponding to a salinity of 16.2–20.1 wt.% NaCl eq. (avg. = 17.9 wt.%, n = 28). Fluid density estimated for these fluid inclusions ranges from 0.96 to 1.06 g/cm3.
The Tfm of primary liquid-rich fluid inclusions in quartz-galena veins from the main-ore stage III is in the range of −48.0 to −39.0 °C with an average of −42.1 °C, suggesting the presence of appreciable amount of CaCl2, in addition to NaCl (Van den Kerkhof & Hein, Reference Van den Kerkhof and Hein2001). Inclusions homogenized into the liquid phase between 147.0 and 198.0 °C (avg. = 171.1 °C, n = 21). Tmice values vary from −14.5 to −10.3 °C, with an average of −12.1 °C (n = 21), corresponding to a salinity of 14.3–18.2 wt.% NaCl eq. (avg. = 16.0 wt.%). Fluid density estimated for the fluid inclusions in these veins is in range of 0.97–1.07 g/cm3.
The Tfm of primary liquid-rich fluid inclusions in late calcite veins of the post-ore stage is in range of −25.0 to −21.0 °C with an average of −22.6 °C. These values are close to the first ice-melting temperature of the NaCl−H2O system (−20.8 °C), suggesting that NaCl is the principal salt in the solution. Inclusions homogenized into the liquid phase between 88.0 and 115.0 °C (avg. = 103.5 °C, n = 12). Tmice values for these inclusions vary from −9.3 to −5.6 °C with an average of −7.1 °C (n = 12), corresponding to a salinity of 8.7–13.2 wt.% NaCl eq. (avg. = 10.6 wt.%, n = 12). Fluid density estimated for the fluid inclusions in late calcite veins is in range of 0.88–0.95 g/cm3.
5.3. Stable (C−O−S) isotopes
The δ13C and δ18O isotopic composition of eleven dolomite crystals (Dol-2), associated with the main-ore stage, is listed in Table 2 and shown in Fig. 14. Dolomite samples have δ13CVPDB and δ18OVSMOW values of −0.99 to +1.99‰ (avg. = +0.31‰) and +20.74 to +25.48‰ (avg. = +23.79‰), respectively. The δ13CCO2 and δ18Ofluid values for dolomite samples range from −0.89 to +2.09‰ and +11.04 to +15.78‰, respectively. The C−O isotope values are similar to those of carbonates in carbonate-hosted Pb-Zn deposits of the YAMB (Table 3), (S Maghfouri, unpub. PhD thesis, Univ. Tabriz, 2017; Maghfouri & Choulet, Reference Maghfouri and Choulet2021), suggesting a similar source of carbon and oxygen.
Table 2. Carbon and oxygen isotopic composition of main-ore stage dolomite (Dol-2) at the Chahmileh deposit
a δ18OVSMOW= 1.03091 (δ18OVPDB) + 30.91 (Friedman and O’Neil 1977),
b 1000lnα(CO2-dolomite) = −1.637× 106/(T+ 273.15)2+ 7.290 (Horita Reference Horita2014),
c 1000lnα(dolomite-fluid) = 4.60× 106/(T+ 273.15)2- 4.650× 103 /(T+ 273.15)+ 1.710 (Zheng Reference Zheng1999). T= 205.3 °C, based on microthermometry analysis of fluid inclusion in dolomite (Dol-2) from the main sulphide stage
Figure 13. (a) Salinity versus homogenization temperature of fluid inclusions hosted in dolomite, quartz and calcite from the Chahmileh deposit, (b) Salinity versus total homogenization temperature and possible fluid composition of the Chahmileh deposit (Kesler, Reference Kesler2005). Fluid evolution trends are represented by arrows from Shepherd et al. (Reference Shepherd, Rankin and Aiderton1985). Notes: trend 1 represents primitive fluid A mixed with cold and low salinity fluid B; trends 2 and 2’ represent the result of fluid A isothermally mixing with different salinity fluid B; trend 3 represents the salinity of residual phase increased, caused by boiling of fluid A; trend 4 represents cooling of fluid A; trend 5 represents leakage of fluid inclusions during heating; and trend 6 represents necking down of the fluid inclusion.
Figure 14. δ13CVPDB versus δ18OVSMOW diagram showing isotopic composition of dolomitic host rock of the Chahmileh deposit and isotopic composition of mineral separates from Mediabad, Farahabad, Mansourabad and Darre Zanjir deposits of YAMB. Given range by Taylor & McLennan (Reference Taylor and McLennan1985) and Hoefs (Reference Hoefs2015).
Table 3. Carbon and oxygen isotopic composition of mineral separates from the Pb–Zn deposits of Yazd-Anarak Metallogenic Belt
a Maghfouri (2017).
b Maghfouri & Choulet (Reference Maghfouri and Choulet2021).
Sulphur isotopic compositions of sulphides and baryte samples are presented in Table 4 and shown in Figs. 15, 16. The δ34SVCDT values of galena samples from the Chahmileh deposit show a narrow range between +6.3 and +8.2‰ with an average value of +7.2‰. The δ34SVCDT values of sphalerite and chalcopyrite range from +5.9 to +6.2‰ (avg. = +6.0‰, n = 2) and +1.4 to +3.4‰ (avg. = +2.6‰, n = 3), respectively. The δ34SVCDT values for baryte samples vary between +15.0 and +17.4‰ (avg. = +16.2‰, n = 2). The δ34S values of equilibrated fluid were calculated using δ34S value of sulphides and mineral-H2S equilibrium isotopic fractionation factors (1000 lnα) based on Li and Liu (Reference Li and Liu2006), assuming H2S as the main sulphur species in the fluid (Table 4). Calculated δ34S values of reduced sulphur (H2S) in equilibrium with the sulphides range from a maximum of 11.0‰ in galena to a minimum of 1.2‰ in chalcopyrite.
Table 4. Sulphur isotopic composition of sulphide minerals and barytes from the Chahmileh deposit
Figure 15. Histogram of sulphur isotopic compositions of various sulphides and sulphate minerals of the Chahmileh deposit.
Figure 16. (a) δ34SVCDT values in sulphides of the Chahmileh deposit in comparison with range and median δ34S values of sulphides in a selection of orogenic-related MVT deposits (data from Leach et al. 2010b; Ehya et al. Reference Ehya, Lotfi and Rasa2010; Ehya, Reference Ehya2014; Jazi et al. Reference Jazi, Karimpour and Malekzadeh Shafaroudi2017; Nejadhadad et al. 2018; Fazli et al. Reference Fazli, Taghipour, Moore and Lentz2019; Rajabi et al. Reference Rajabi, Canet, Alfonso, Mahmoodi, Yarmohammadi, Sharifi, Mahdavi and Rezaei2022), (b) Distribution of δ34S values of baryte and sulphide minerals from the Chahmileh deposit in relation to age curve for sulphur (Claypool et al. Reference Claypool, Holser, Kaplan, Sakai and Zak1980; Bottrell & Newton, Reference Bottrell and Newton2006).
5.4. Pb-isotope analyses
Lead isotopic compositions of galena samples from the Chahmileh deposit are presented in Table 5 and shown in Fig. 17. Galena samples (n = 6) have 206Pb/204Pb ratios ranging from 18.546 to 18.576, 207Pb/204Pb ratios ranging from 15.650 to 15.688 and 208Pb/204Pb ratios ranging from 38.795 to 38.918. 206Pb/204Pb, 207Pb/204Pb and 208Pb/204Pb ratios for galena samples from the Pb–Zn deposits of the CIZ show the following ranges, respectively: 18.427 to 19.081, 15.586 to 15.722 and 38.500 to 38.910 (Table 5; Mirnejad et al. Reference Mirnejad, Simonetti and Molasalehi2015), respectively. The Pb-isotope data of CIZ carbonate-hosted Pb–Zn deposits are similar to those obtained for the Chahmileh deposit, suggesting a similar lead source.
Table 5. The Pb isotopic composition of galena samples from the Chahmileh deposit and Pb–Zn deposits of the Central Iran Zone
a This study.
b Mirnejad et al. (Reference Mirnejad, Simonetti and Molasalehi2015).
Figure 17. (a) and (b) Pb isotopic ratios of galena samples on a ‘plumbotectonic’ diagram (Zartman & Doe, Reference Zartman and Doe1981). The Pb-isotope data of galena from the Central Iran Zone (Mirnejad et al. Reference Mirnejad, Simonetti and Molasalehi2015) are presented for comparison.
6. Discussion
6.1. Evolution of ore-forming fluids and ore controls
Carbonate-hosted Pb–Zn deposits are an important global source of Zn and Pb, which formed through fluid circulation within carbonate rocks (Leach et al. Reference Leach, Macquar, Lagneau, Leventhal, Emsbo and Premo2006, Reference Leach, Bradley, Huston, Pisarevsky, Taylor and Gardoll2010a; Li et al. Reference Li, Chen, Hollings, Zhang, Sun, Lu, Wang and Fang2018). Temperature, salinity, pH and redox state (oxygen fugacity, fO2) of fluids are crucial factors to control Zn and Pb mineralization (Cooke et al. Reference Cooke, Bull, Large and McGoldrick2000; Gómez-Fernández et al. Reference Gómez-Fernández, Both, Mangas and Arribas2000; Leach et al. Reference Leach, Bradley, Lewchuk, Symons, de Marsily and Brannon2001, Reference Leach, Sangster, Kelley, Large, Garven, Allen, Gutzmer and Walters2005, Reference Leach, Bradley, Huston, Pisarevsky, Taylor and Gardoll2010a; Conliffe et al. Reference Conliffe, Wilton, Blamey and Archibald2013).
The salinity-homogenization temperature diagram (Fig. 13a) of fluid inclusions shows a wide range of temperatures in dolomite and quartz samples (Th = 133.0–248.0 °C) and their salinities (14.3–20.8 wt.% NaCl eq.). Range of fluids salinities probably reflects mixing between high-temperature fluid (Fore) (up to 248.0 °C) and high salinity (up to ∼20.8 wt.% NaCl eq.) with a fluid (Fsub) characterized by high temperature and low salinity. These two fluids likely mixed at the site of sulphide deposition. Temperature and salinity of fluid inclusions at Chahmileh fall within the range of Irish-type and MVT Pb–Zn deposits (salinity = 10–35 wt.% NaCl eq. and temperature = 80–250 °C) and are typical of basinal brines (Wilkinson, Reference Wilkinson2001; Leach et al. Reference Leach, Sangster, Kelley, Large, Garven, Allen, Gutzmer and Walters2005, Reference Leach, Bradley, Huston, Pisarevsky, Taylor and Gardoll2010a). Relationship between homogenization temperature and salinity (Fig. 13b) indicates that basinal brine with some possible input from seawater fluid was responsible for ore mineralization at Chahmileh.
The ore precipitation mechanisms, common in numerous MVT and other sediment-hosted Pb-Zn ore deposits, include (i) cooling of hydrothermal fluid, (ii) pH decrease and (iii) mixing between two or more fluids (Anderson, Reference Anderson1973; Sverjensky, Reference Sverjensky1981). Upward migration of a metal-bearing, sulphur-depleted, hot brine (Fore), originating from the Paleozoic basement, most likely mixed with a shallow, warm, metal-depleted, sulphur-rich reservoir (Fsub). Such mixing event accounts for the relatively wide range of salinity recorded in dolomite and quartz-hosted fluid inclusions.
Corbella et al. (Reference Corbella, Ayora and Cardellach2004), based on reactive transport modelling, demonstrated that fluid-mixing concomitantly triggers carbonate dissolution and precipitation of sulphides in MVT/sediment-hosted Pb–Zn deposits. The pH increases due to carbonate dissolution and a temperature decrease most likely contributed to the precipitation of the ore minerals. Such a fluid-mixing model has been proposed for numerous sediment-hosted Pb–Zn ore deposits where the mixing of a hot, metal-bearing fluid with a cooler, diluted fluid triggered ore precipitation (Beales & Jackson, Reference Beales and Jackson1966; Anderson, Reference Anderson1973; Sverjensky, Reference Sverjensky1986; Plumlee et al. Reference Plumlee, Leach, Hofstra, Landis, Rowan and Viets1994; Leach et al. Reference Leach, Sangster, Kelley, Large, Garven, Allen, Gutzmer and Walters2005). In consideration of this evidence, it is likely that fluid dilution may have had an important role in ore deposition in the Chahmileh district (Fig. 13b).
Ore-forming hydrothermal fluids transport metals as ions and molecular complexes when migrating through the Earth’s crust (Seward et al. Reference Seward, Williams-Jones and Migdisov2014). Complexing with chloride and hydrosulphide/sulphide ligands is generally considered the important transport forms for silver, lead and zinc in the fluid systems, as demonstrated in some detail by several researchers (Seward, Reference Seward1976; Ruaya & Seward, Reference Ruaya and Seward1986; Sverjensky et al. Reference Sverjensky, Shock and Helgeson1997; Tagirov et al. Reference Tagirov, Suleimenov and Seward2007; Tagirov & Seward, Reference Tagirov and Seward2010; Mei et al. Reference Mei, Sherman, Liu, Etschmann, Testemale and Brugger2015; Zhong et al. Reference Zhong, Brugger, Chen and Li2015). Metal complexing by chloride and bisulphide complexes is controlled by sulphide solubilities and the hydrothermal fluid temperature (Zhong et al. Reference Zhong, Brugger, Chen and Li2015). In low to medium temperature (< 350 °C) hydrothermal systems (MVT and SEDEX) (e.g., Chahmileh deposit), low sulphide solubilities dictate that the ore fluids cannot carry both reduced sulphur and metals. Hanor (Reference Hanor and Sangster1996) and Reed (Reference Reed2006) proposed that in ore-forming brines, transportation of Pb and Zn was mostly controlled by chloride complexes. Precipitation of sulphide minerals is probably a consequence of chloride complexe destabilization during fluid mixing and dilution. A chloride threshold of 100 g.L−1 (salinity ∼17 wt.% NaCl eq.) was calculated for metal transportation in the basinal brines (at temperatures of <150 °C) when the reduced sulphur concentrations were lower than 0.02 mg.L−1 (Kharaka et al. Reference Kharaka, Maest, Carothers, Law, Lamothe and Fries1987; Sicree & Barnes, Reference Sicree and Barnes1996; Giordano, Reference Giordano2000). Therefore, it is reasonable to imply that concentration of chloride (110 g·L−1) calculated based on fluid inclusion salinity data is enough to act as a complex for Pb and Zn migration. First ice-melting temperature measurements in fluid inclusions suggest the presence of Na+, K+, Ca+2 and Mg+2 as dissolved cations in the ore-forming fluid inclusions in the Chahmileh deposit. These cations were probably leached from the sedimentary units and transported as chloride complexes (Sverjensky, Reference Sverjensky, Boyle, Brown, Jefferson, Jowett and KirKham1989) in the hydrothermal solution.
Possible processes that significantly invoked increasing the chlorine content of seawater trapped in the sediments include; shale membrane filtration (Graf, Reference Graf1982; AMF Garavito Rojas, unpub. PhD thesis, Univ. Vrije, 2006), maturation of white mica (sericite and muscovite) (Michalik, Reference Michalik and Papunen1997) and hydration of detrital clastic minerals (biotite) to sheet silicates (sericite or chlorite) via diagenesis (Gleeson et al. Reference Gleeson, Yardley, Munz and Boyce2003). Identifying the exact mechanism for the source(s) of Cl in the Chahmileh deposit needs further scrutiny, but high chloride content of the metalliferous hydrothermal fluid in the Chahmileh deposit is inferred from fluid inclusion data.
6.2. Source of CO2
There are three principal sources of carbon and oxygen in hydrothermal fluids: (I) mantle, (II) marine carbonate rocks and (III) sedimentary organic matter (Taylor et al. Reference Taylor, Frechen and Degens1967; Veizer & Hoefs, Reference Veizer and Hoefs1976; Demény & Harangi, Reference Demény and Harangi1996; Liu & Liu, Reference Liu and Liu1997; Demény et al. Reference Demény, Ahijado, Casillas and Vennemann1998). The δ13CVPDB and δ18OVSMOW values of the mantle, marine carbonate and organic matter range from −8.0 to −4.0‰, +6.0 to +10.0‰ (Taylor et al. Reference Taylor, Frechen and Degens1967), −4.0 to +4.0‰ and +20.0 to +30.0‰ (Veizer & Hoefs, Reference Veizer and Hoefs1976) and −30.0 to −10.0‰ and +24.0 to +30.0‰ (Liu & Liu, Reference Liu and Liu1997), respectively. The δ18OVSMOW values plotted against δ13CVPDB for dolomite samples of the main-ore stage are higher than typical values for igneous carbonate and somewhat lower than those of organic sediments, but are similar to those of marine carbonate rocks (Fig. 14). Thus, the carbon and oxygen isotopic data indicate that the CO2 in the ore-forming fluid (hydrothermal dolomite) likely originated from the dissolution of marine carbonates. Therefore, the Triassic dolomitic marble of the Chah Gorbeh Complex was likely the main source of CO2 in the ore-forming fluid. The CO2 produced by dissolution of Triassic dolomitic marble will increase H2CO3 content and activity, making the fluid more acidic and dissolving further carbonates until it reaches chemical equilibrium with dolomitic marble (Spangenberg et al. Reference Spangenberg, Fontboté, Sharp and Hunziker1996). We plotted both δ13CVPDB and δ18OVSMOW values of carbonate samples from Mediabad, Farahabad, Mansourabad and Darre Zanjr deposits in Fig. 14. The isotopic data plot near those of marine carbonates and/or between the marine carbonate and sedimentary organic matter field. Thus, the similarity in C-O isotope ratios for the Pb-Zn deposits located within YAMB suggests a similar source for CO2 in the ore-forming fluids, which likely originated from dissolution of marine carbonate rocks or dihydroxylation of sedimentary organic matter during mineralization.
Calcite and dolomite are two main C-bearing hydrothermal minerals in the Chahmileh deposit. H2CO3 (present as CO2 (aqueous)) and HCO3 ¯ are two key C-bearing species in hydrothermal fluids. Given that the C isotope fractionation between H2CO3 or HCO3 ¯ (aq.) and CO2 (gas) is negligible, i.e., δ13Cfluid ≈ δ13CCO2 (Ohmoto, Reference Ohmoto1972), it is possible to estimate the theoretical δ13Cfluid value from calculated value of δ13CCO2. Assuming an average homogenization temperature of 205 °C (fluid inclusions in dolomite samples associated with the main-ore stage, Fig. 9), the δ13CCO2 values are in the range of −0.89 to +2.09‰, calculated using the 1000lnα(CO2-dolomite)= −1.637× 106/(T+ 273.15)2 + 7.290 function (Horita, Reference Horita2014; T = 205.3 °C). Similarly, δ18Ofluid values vary from +11.04 to +15.78‰ and were calculated using the 1000lnα(dolomite-fluid) = 4.60 × 106/(T+ 273.15)2 − 4.650 × 103/(T+ 273.15) + 1.710 function (Zheng Reference Zheng1999; T = 205.3 °C). The theoretical δ13CCO2 and δ18Ofluid values suggest that source of C is likely the 13C-enriched marine carbonate, while O is possibly derived from a mixed source of 18O-depleted metamorphic water and 18O-enriched dolomitic marble.
6.3. Source of sulphur and mechanisms of sulphide deposition
Sulphur isotopic compositions of sulphur-bearing minerals are mostly affected by the source and fractionation processes (Ohmoto & Goldhaber, Reference Ohmoto, Goldhaber and Barnes1997; Seal et al. 2006; Hoefs, Reference Hoefs2015). Sulphides in carbonate-hosted Pb-Zn deposits show wide ranges of δ34S (Fig. 16a), with values from lower than −25‰ to higher than +35‰ (Leach et al. Reference Leach, Sangster, Kelley, Large, Garven, Allen, Gutzmer and Walters2005). However, the Chahmileh Pb-Zn deposit sulphides show a relatively narrow range of δ34S values, from +1.4 to +8.2‰, with an average of +5.4‰ (Figs. 15, 16). Although this narrow range may be caused by the limited number of analysed samples of sulphides, the δ34S values also may suggest a uniform isotopic composition of the source and uniform conditions governing the isotopic fractionation between sulphur species in the ore-forming fluids during mineralization. The observed δ34SPbS > δ34SZnS trend suggests that ore minerals were precipitated under disequilibrium conditions, which are typical of ore formation at temperatures well below 250 °C (Ohmoto & Rye, Reference Ohmoto, Rye and Barnes1979; Ohmoto, Reference Ohmoto1986).
The original fluid δ34SH2S can be estimated from δ34S values of hydrothermal ore minerals. Under physical and chemical conditions (T < 300 °C, low pH, and Eh) envisaged for the main-stage fluids, the major sulphur species would be H2S (Ohmoto & Rye, Reference Ohmoto, Rye and Barnes1979). At these temperatures, sulphide dominance causes sulphide mineral δ34S values to be close to the original fluid δ34SH2S (Ohmoto & Rye, Reference Ohmoto, Rye and Barnes1979). The original fluid δ34SH2S values in equilibrium with sulphide minerals were estimated to be in range of +1.2‰ to +11.0‰ (avg. = +6.6‰).
Predominant source of sulphur in sediment-hosted Zn−Pb−(Cu−Ag−Ba) deposits is seawater sulphate (Leach et al. Reference Leach, Sangster, Kelley, Large, Garven, Allen, Gutzmer and Walters2005, Reference Leach, Bradley, Huston, Pisarevsky, Taylor and Gardoll2010a; Magnall et al. Reference Magnall, Gleeson, Stern, Newton, Poulton and Paradis2016), and sulphur isotopic composition of the Triassic seawater suphate ranges from +11‰ to +20‰ (Fig. 16b) (Claypool et al. Reference Claypool, Holser, Kaplan, Sakai and Zak1980; Bottrell & Newton, Reference Bottrell and Newton2006). The Triassic age of carbonates hosting the Chahmileh deposit and the δ34S values for baryte (+15.0‰ and +17.4‰) (Fig. 16b) fall in the range of Triassic marine sulphate. Sulphide can be produced from seawater sulphate either by biogenic processes (bacterial sulphate reduction, BSR) or abiogenically via thermochemical processes (thermochemical sulphate reduction, TSR) (Machel, Reference Machel2001). Values of δ34S are also dependent on other factors, such as the sulphate reduction rate (Leavitt et al. Reference Leavitt, Halevy, Bradley and Johnston2013), sulphate concentration (Habicht et al. Reference Habicht, Gade, Thamdrup, Berg and Canfield2002), temperature (Sawicka et al. Reference Sawicka, Jørgensen and Brüchert2012), content of organic-rich matter in sediments (Goldhaber et al. 1995) and rate of sulphate replacement by sulphides (Rajabi et al. Reference Rajabi, Alfonso, Canet, Rastad, Niroomand, Modabberi and Mahmoodi2020). BSR usually takes place at 60–80 °C (Machel, Reference Machel1989), although its occurrence at 110 °C has also been reported (Jørgenson et al. Reference Jørgenson, Isaksen and Jannasch1992). BSR commonly produces a wide δ34SVCDT range due to the large isotopic fractionation (15–66‰) between sulphate and sulphide (Rees, Reference Rees1973; Leach et al. Reference Leach, Sangster, Kelley, Large, Garven, Allen, Gutzmer and Walters2005; Basuki et al. Reference Basuki, Taylor and Spooner2008; Sim et al. Reference Sim, Bosak and Ono2011; Li et al. Reference Li, Liu, Xue and Li2019). In contrast, TSR (occur at 150–350 °C) produces a relatively narrow range for sulphur isotopic fractionation, yielding <15‰ in the presence of organic matter (Ohmoto, Reference Ohmoto1972; Ohmoto & Rye Reference Ohmoto, Rye and Barnes1979; Worden et al. Reference Worden, Smalley and Oxtoby1995; Wang et al. Reference Wang, Mi, Zhou and Luo2018). Moreover, mixing of hot and cold ore-forming fluids could form S2– from SO4 2- by TSR through the SO4 2– + 2C = S2– + 2CO2, SO4 2– + CH4 = H2S + CO3 2– + H2O, or SO4 2– + 2CH2O = H2S + 2HCO3- reaction (Worden et al. Reference Worden, Smalley and Oxtoby1995; Leach et al. Reference Leach, Sangster, Kelley, Large, Garven, Allen, Gutzmer and Walters2005). TSR has been invoked as an important sulphate reduction process in MVT deposits (Leach et al. Reference Leach, Sangster, Kelley, Large, Garven, Allen, Gutzmer and Walters2005; Wilkinson, Reference Wilkinson, Holland and Turekian2014). In the Chahmileh deposit, the narrow range of positive δ34S values of the sulphide minerals (+1.4 to +8.2‰) suggests that sulphur was probably supplied by the TSR process since Th values during ore formation at the Chahmileh (133.0 °C to 248.0 °C) are unfavourable for BSR (Kyle & Saunders, Reference Kyle, Saunders and Sangster1996; Leach et al. Reference Leach, Sangster, Kelley, Large, Garven, Allen, Gutzmer and Walters2005).
6.4. Source of metals
The Pb isotopic compositions of galena in Chahmileh Pb-Zn deposit are quite homogeneous, which implies that Pb was supplied from either a completely homogenized mixed source or a single source. On a thorogenic diagram, Pb compositions show a positive linear correlation between the lower and upper crust curves, reflecting a likely mixing between multiple endmembers while the uranogenic isotope diagram shows more complexity. Data plot between the orogen and upper crust growth curves, suggesting a possible heterogeneous source for Pb (Fig. 17). These variations likely reflect a contribution from different crustal sources to the Pb mineralization. Crustal sources of Pb are also reflected by higher U/Pb ratios and high 206Pb/204Pb ratios, as revealed by calculated μ values (Zartman & Doe, Reference Zartman and Doe1981). Since Pb-isotope data of possible source rocks in the region are not available, the ultimate source(s) of Pb is unclear. However, like other Pb-Zn deposits in the CIZ (Fig. 17), the source(s) of Pb were likely continental crust or pelagic sediments developed during orogenic activities (Mirnejad et al. Reference Mirnejad, Simonetti and Molasalehi2015). In Pb−Zn deposits of the CIZ and Alborz Zone (AZ), most of the galena yield Pb model ‘ages’ of ∼140 and ∼250 Ma, indicating that mineralization resulted from the extraction of ore-bearing fluids from Upper Triassic−Lower Jurassic sequences (Mirnejad et al. Reference Mirnejad, Simonetti and Molasalehi2015). The similarity in Pb-isotope ratios for the Pb-Zn deposits located within these zones suggests similar crustal evolution histories. Mirnejad et al. (Reference Mirnejad, Simonetti and Molasalehi2015) argued that Pb-Zn mineralization in sedimentary and igneous rocks of the CIZ and AZ tectonic regions occurred following the Late Cretaceous-Tertiary accretionary stage of crustal thickening.
Different types of Pb-Zn deposits have different Pb isotopic signatures (Sangster et al. Reference Sangster, Outridge and Davis2000), though the great majority of MVT deposits have 206Pb/204Pb values of 17.5–23 and 207Pb/204Pb values of 15.4–16.2 (Leach et al. Reference Leach, Sangster, Kelley, Large, Garven, Allen, Gutzmer and Walters2005). The Pb isotopic composition of galena samples from Chahmileh also falls in this broad range (average 206Pb/204Pb = 18.566 and 207Pb/204Pb = 15.668). The Pb-isotope data suggest an orogenic reservoir coupled with a large contribution from crustal basement rocks as metal sources for the Chahmileh deposits; a pattern reported for most MVT deposits (Leach et al. Reference Leach, Sangster, Kelley, Large, Garven, Allen, Gutzmer and Walters2005).
6.5. Proposed genetic model
Three distinct categories of sediment-hosted Pb–Zn deposits have been proposed by Leach et al. (Reference Leach, Sangster, Kelley, Large, Garven, Allen, Gutzmer and Walters2005), DL Leach et al., report (2010b) and Wilkinson (Reference Wilkinson, Holland and Turekian2014) based on their tectonic setting, host-rock sequence and geochemical characteristics: MVT-type, SEDEX-type and Irish-type. The Pb–Zn mineralization in the Chahmileh deposit is hosted by the Middle Triassic dolomitic marble of the Chah Gorbeh Complex. Due to the association of mineralization with carbonate rocks, this deposit is comparable to MVT- and Irish-type deposits. The Chahmileh deposit does not share many of the main features of the Irish-type deposits, as outlined by Leach et al. (Reference Leach, Sangster, Kelley, Large, Garven, Allen, Gutzmer and Walters2005) and Wilkinson (Reference Wilkinson, Holland and Turekian2014). Although it is hosted by carbonate rocks and shows open-space filling textures, mineralization is not a sub-seafloor replacement associated with synsedimentary normal faults, it did not form during diagenesis and the sulphides in the deposit do not show any synsedimentary textures. Instead, the Chahmileh Pb–Zn deposit has many features that match closely with those of most MVT deposits (Sangster, Reference Sangster1990; Leach et al. Reference Leach, Sangster, Kelley, Large, Garven, Allen, Gutzmer and Walters2005). These features include: (1) the ore mineralization is epigenetic and stratabound; (2) there is no relationship between mineralization and igneous activity; (3) thrust and normal faults are indicative of a compression and extension geodynamic setting; (4) ore is mineralogically simple and composed dominantly of galena, sphalerite, pyrite, chalcopyrite and dolomite, quartz and calcite; (5) ore bodies show open-space filling, brecciated and replacement textures; (6) lack of exhalative processes or laminated ores; (7) hydrothermal alteration mainly consists of dolomitization and silicification, associated with host-rock dissolution and brecciation; (8) moderate to high Th values up to 248 °C, reflects the relatively high temperatures of ore formation that are not typical of many MVT deposits; (9) C-O isotopes suggest that CO2 is originated from carbonate host rock; (10) the average δ34S value of sulphide is lighter than contemporaneous seawater and (11) Pb isotopic data indicate crustal sources for the metals. Most of these evidences indicate that the Chahmileh deposit is similar to a typical MVT deposit (e.g., Leach et al. Reference Leach, Bradley, Lewchuk, Symons, de Marsily and Brannon2001, Reference Leach, Sangster, Kelley, Large, Garven, Allen, Gutzmer and Walters2005, 2010b; Pirajno, Reference Pirajno2009, Reference Pirajno2013; Wilkinson, Reference Wilkinson, Holland and Turekian2014) (Table 6).
Table 6. Comparison between MVT deposits and the Chahmileh Pb–Zn deposit
Rajabi et al. (Reference Rajabi, Rastad and Canet2012, Reference Rajabi, Rastad and Canet2013) proposed that the main Cimmerian (Upper Triassic) and Laramide (Late Cretaceous-Tertiary) orogenic collisions led to development of discordant, stratabound MVT deposits in different Pb–Zn metallogenic belts of Iran adjacent to the Paleo-Tethys and Neo-Tethys suture zones. The main Cimmerian orogenic event transformed the northern margin of the Iranian Plateau into a collisional foreland basin (Wilmsen et al. Reference Wilmsen, Fürsich, Seyed-Emami, Majidifard and Taheri2009). This compression ‘squeezed’ fluid from the foreland basin towards the Triassic carbonate platforms, and brines migrated in basin large-scale regional faults and fissures in the basement and overlying strata, extracted metals (i.e., Pb and Zn) and finally migrated to suitable host rocks (Chah Gorbeh Complex). Simultaneously, extensive hydrothermal fluid flow and circulation of reduced sulphur of sulphate in the overlying strata into thiosulphuric acid and hydrogen sulphuric acid migrated with the infiltrating fluid, eventually mixed with the reduced sulphur generated by pyrolysis of sulphur-containing organic matter in the wall rock. Once metal-bearing and reduced sulphur-bearing fluids mixed in favourable fracture zone(s), by changes in ore-forming conditions, metal sulphides precipitated and formed the ore bodies.
7. Summary and conclusions
The Chahmileh Pb–Zn deposit is a fault-controlled deposit located in the YAMB of the CIZ. It is composed of sulphide and non-sulphide ores hosted in Middle Triassic dolomitic marble of the Chah Gorbeh Complex. The main sulphides are galena, sphalerite and minor amounts of chalcopyrite and pyrite, which were partially or completely transformed into non-sulphide minerals by supergene processes. Cerussite, hemimorphite, wulfenite, malachite, mimetite, smithsonite and iron oxy-hydroxides are the main non-sulphide ore minerals. Fluid mixing and dilution is the most probable mechanism of ore-forming fluids evolution. Carbon and oxygen isotopic compositions of dolomite suggest that CO2 in the ore-forming hydrothermal fluids mainly originated from marine carbonate rocks. TSR is the most likely process of supplying reduced sulphur for sulphide deposition. Pb in the deposit originated from an orogenic source, which was dominated by upper crustal rocks with high 238U/204Pb and 232Th/204Pb ratios. Ore precipitation is mostly initiated by interaction of the ore-bearing fluid with carbonate host rocks, accompanied by decreasing temperature, increasing pH and perhaps a drop-in oxygen fugacity, and therefore, with a concomitant increase in the reduced sulphur content by a TSR mechanism. The Chahmileh deposit system is likely an MVT-type Pb–Zn deposit, related to the thrust compression-driven fluid flow, developed in Middle Triassic carbonate strata.
Acknowledgements
This paper is part of the second author’s PhD research thesis, supported by partial grant from the Kharazmi University (Iran). The authors acknowledge financial support from the Iranian Mines and Mineral Industries Development and Renovation Organization. We would like to thank Dr. J.H. Curtis from the Stable Isotope Mass Spectroscopy Laboratory, University of Florida, for his assistance in carbon-oxygen isotope analysis, and Dr. D. Dettman from the Environmental Isotope Laboratory at the University of Arizona for his help in sulphur isotope analysis. Author would like to thank Prof. Paul Spry for his thorough review, constructive comments and massive improvement of the manuscript.
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