Geological and geochemical constraints on the Farahabad vent-proximal sub-seafloor replacement SEDEX-type deposit, Southern Yazd basin, Iran

Journal Pre-proof Geological and geochemical constraints on the Farahabad ventproximal sub-seafloor replacement SEDEX-type deposit, Southern Yazd basin, Iran

Sajjad Maghfouri, Mohammad Reza Hosseinzadeh, David R. Lentz, Flavien Choulet PII:

S0375-6742(19)30389-9

DOI:

https://doi.org/10.1016/j.gexplo.2019.106436

Reference:

GEXPLO 106436

To appear in:

Journal of Geochemical Exploration

Received date:

9 July 2019

Revised date:

23 November 2019

Accepted date:

1 December 2019

Please cite this article as: S. Maghfouri, M.R. Hosseinzadeh, D.R. Lentz, et al., Geological and geochemical constraints on the Farahabad vent-proximal sub-seafloor replacement SEDEX-type deposit, Southern Yazd basin, Iran, Journal of Geochemical Exploration (2019), https://doi.org/10.1016/j.gexplo.2019.106436

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Journal Pre-proof Geological and geochemical constraints on the Farahabad vent-proximal sub-seafloor replacement SEDEX-type deposit, Southern Yazd basin, Iran Sajjad Maghfouria*, Mohammad Reza Hosseinzadehb, David R. Lentzc, Flavien Chouletd a b

Department of Earth Sciences, Faculty of Natural Sciences, University of Tabriz, Tabriz, Iran

Department of Earth Sciences, University of New Brunswick, New Brunswick, Canada

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Chrono-Environnement, Université de Franche-Comté/CNRS, 25030 Besançon Cedex, France

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c

Department of Geology, Faculty of Basic Sciences, Tarbiat Modares University, Tehran, Iran

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_________________________________________________________________________________________

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Abstract

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Farahabad is a Lower Cretaceous-age sandy dolomite, sandstone and silty limestone hosted zinc-

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lead deposit in the southern Yazd basin, Iran. This deposit occurs in the lower part of Taft

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Formation. The Farahabad deposit is a stratiform and stratabound accumulation of hydrothermal sulfides. The mineralized zone at Farahabad is lens shaped and has a relatively flat top. It also

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ranges greatly in thickness, from a few meters to more than 30 m. Three styles of ore facies have

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been differentiated at the deposit. Feeder zone, bedded ore facies, whereas the bulk of ore is contained within sphalerite and galena breccias that record a prolonged and texturally complex history of ore replacement (massive-replacement ore facies). The shape and size of the mineralized body suggest that dolomites replaced with sulfide minerals in a submarine channel. Dolomitization and sulfide mineralization’s are the major hydrothermal products in the Farahabad deposit, and they occur in all host rocks, but only in the immediate vicinity of the synsedimentary normal fault that served as the main conduit for the mineralizing fluids. Based on microscopic studies, sulfide mineralization in the Farahabad deposit emplaced during two stages: fine-grained sulfide bands (stage I) are intricately interlayered with organic matter-rich beds and 1

Journal Pre-proof thin turbidite beds. They exhibit classic sedimentary textures, such as laminae, cross bedding, graded bedding and bedding, indicative of a synsedimentary origin. Main-stage (stage II) mineralization involved the progressive replacement of preexisting sulfides and the dolomite breccias, initially by replacement of the breccia matrix and ultimately by replacement of clasts. Dolomite minerals from the feeder zone and massive-replacement ore facies have 115–234 °C trapping temperature with salinities ranging from 0.99 to 16.70 eq. wt.% NaCl eq. δ 34 S values of

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sulfide minerals range from –19.47 to +3.9‰, suggesting that sulfur in the hydrothermal fluids

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was derived predominantly from reduction of seawater sulfate by bacteriogenic sulfate reduction.

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The δ 13 CPDB and δ 18 OSMOW values of host limestone’s and hydrothermal dolomites plot the near marine carbonate rocks field in a plot of δ 13 CPDB vs. δ 18 O SMOW diagram. It suggests that CO 2 in

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the hydrothermal fluids was mainly originated from marine carbonate rocks. Abrupt lateral

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changes in facies and thickness, along with the existence of synsedimentary breccia’s and debris

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flows within the ore sequence, suggest the proximity of synsedimentary faults and tectonic activity contemporaneous with the sedimentation in the Lower Cretaceous, favorable to the

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formation of deposit. So this deposit formed by combination of sub-seafloor replacement and on

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the seafloor processes. Characteristics of the Farahabad deposit are compatible with a ventproximal sub-seafloor replacement SEDEX -type classification.

Keywords: Lower Cretaceous; Farahabad Zn-Pb (- Ag) deposit; SEDEX-type, Southern Yazd basin; Iran _________________________________________________________________________

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1. Introduction It is known that the sediment-hosted Pb-Zn deposits have been classified in two main classes by Leach et al. (2010). The first class is “clastic-dominated Pb-Zn (CD Pb-Zn) deposits” (SEDEX-type) that occur within the shale, sandstone, siltstone and/or carbonates belonging to clastic sedimentary sequences. The tectonic settings of the CD Pb-Zn deposits are passive

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margins, back-arc, continental rifts and sag basins (Leach et al. 2010). The second class is the

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Mississippi Valley-type (MVT) Pb-Zn deposits which occur in platform carbonates of passive

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margin tectonic environments. Together, these contain around 48% of the global resources of zinc and 52% of the lead (Singer 1995), so they are the most important repositories of these

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metals on Earth. About 600 sediment-hosted Zn-Pb-Ba deposits and occurrences are presently

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known in Iran (Fig. 1). These Zn-Pb (- Ag) deposits formed dominantly in the Yazd-Anarak

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metallogenic belt (YAMB), the Malayer-Esfahan metallogenic belt (MEMB), and the Central Alborz metallogenic belt (CAMB) (Rajabi et al. 2012a; 2104) (Fig. 1). Only a few deposits have

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been actually explored and/or exploited although <70 deposits are currently being mined (Rajabi

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et al. 2012a, b and 2014). The age of mineralization is commonly considered to be synsedimentary (Kuoshk and Chahmir ((Yaghubpour and Mehrabi 1997; Rajabi et al. 2012b)), however, some deposits formed in a sub-seafloor replacement environment (i.e. Mehdiabad and Mansourabad (Maghfouri et al. 2015; Maghfouri 2017; Maghfouri et al. 2018)) and some formed during burial diagenesis (i.e. Gushfil (Boveiri and Rastad 2016; Boveiri et al. 2017)). Nevertheless, the age of formation is assumed to be the same or very close to the age of the host rocks.

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Journal Pre-proof The southern Yazd basin is located in the southern part of the YAMB, approximately 20 km south of Yazd city (Figs. 1 and 2). Abundant sediment-hosted Zn–Pb (-Ag) mineralization is present in this basin, including the Mehdiabad world-class deposit, Mansourabad, Farahabad and Darreh-Zanjir deposits (Maghfouri et al. 2015; Maghfouri 2017; Maghfouri et al. 2018b; Maghfouri et al., 2019). With estimated global reserves and resources exceeding 394 Mt at ~9 percent Zn, ~4 percent Pb, and ~50 g/t Ag, the southern Yazd basin deposits contains one of the

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largest basin of zinc in the world (Leach et al. 2005; Maghfouri 2017; Maghfouri et al. 2018a)

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(Fig. 2). The Farahabad deposit is located approximately 71 km southwest of the Yazd city in the

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southern Yazd basin (Fig. 2). Despite having been studied and explored during the last three decades, some aspects of the geology of the Farahabad deposit, and particularly its genesis, are

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still very poorly understood. Very little information has been published on the mineralization of

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the Farahabad deposit (Ghasemi 2006; Maghfouri 2017), and most information comes from

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reconnaissance and broad regional studies (Ghasemi 2006; Maghfouri 2017; Maghfouri et al. 2018b). This paper describes the local stratigraphic and geological features together with, the

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relationships between the Zn-Pb (-Ag) mineralization and the host rocks. Descriptions are

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supplemented by petrographic, sulfur, carbon and oxygen isotopes, fluid inclusion and mineralogical studies to constrain the conditions for the genesis of the Farahabad deposit. 2. Geologic Setting 2.1. Geology of Yazd Block and southern Yazd basin The study area is located in the western part of the Central-East Iranian Microcontinent (CEIM; Takin 1972) (Fig. 1). The CEIM, together with central Iran and the Alborz Mountains, forms the Iran Plate, which occupies a structural key position in the Middle Eastern Tethysides

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Journal Pre-proof (Sengor et al. 1988). As an element of the Cimmerian microplate assemblage, it became detached from Gondwana during the (Late) Permian and collided with Eurasia (Turan Plate) in the late Middle– early Late Triassic, thereby closing the Palaeo-Tethys (e.g. Berberian and King 1981; Boulin 1991; Sengor et al. 1988; Sengor 1991; Stampfli and Borel 2002; Fursich et al. 2009a). The CEIM consists of three north–south oriented structural units, called the Lut, Tabas and Yazd Blocks, which are today aligned from east to west, respectively (Fig. 1). The Yazd

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Block is bounded by the Anar Fault to the east, the Chapedoni Fault to the northeast and the

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Naein-Dehshir-Baft faults to the south and to the west (Alavi 1994). The stratigraphy of the Yazd

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Block is very similar to other parts of Central Iran; exposing sequences of Neoproterozoic to Middle Triassic sedimentary rocks, covered by extensive coal-bearing Upper Triassic to Mid-

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Jurassic sandstones, siltstones and shales (Shemshak Formation). It has a Neoproterozoic–

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Cambrian basement consolidated during the late Pan-African tectonic event, followed by a

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Middle Cambrian molasse (Lalun sandstone). This basement was affected by Late Ordovician– Early Devonian rifting events (continental tholeiitic flood basalts, evaporites, dolomites, and

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sandstones) (Bagheri 2007), covered by a relatively continuous shallow-water platform-type

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Devonian to Middle Triassic passive margin mega-sequence (Stöcklin 1968; Stampfli 1978; Sharkovski et al. 1984; Wendt et al. 2005; Leven et al. 2006). In the southern Yazd basin, Ordovician to Jurassic rocks of the Badamo, Shemshak and Esfandiar Formations disconformably overlie the late Neoproterozoic metamorphic rocks of the Boneh- Shurow basement complex (Aghanabati 2004) (Fig. 2). The Boneh-Shurow Complex is the most widely exposed metamorphic unit of the southern Yazd basin with ridge-forming outcrops directly adjacent to, and to the east of, the Deh Shir Fault (Haghipour 1977) (Fig. 2). It exhibits a distinct metamorphic layering composed of an alternation of pink quartz-feldspathic 5

Journal Pre-proof gneisses, greenish-gray mica-schists and dark-colored amphibolites. Other subordinate, but important, constituents of the Boneh-Shurow Complex are dolomitic marble interlayers and latestage, mafic-intermediate magmatic intrusions (for example, diabase dikes) of limited volume. The Badamu Formation consists of commonly dark-grey, often oolitic and/or bioclastic limestone beds of variable thickness, with intercalated sandy marls, shales and fine-grained calcareous sandstones. The Badamu Formation is fairly fossiliferous, and mainly yields

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ammonites, bivalves, corals and serpulids (Huckriede et al. 1962; Seyed-Emami 1967; 1971;

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Seyed-Emami et al. 2000; Wilmsen et al. 2009). The Esfandiar Limestone Formation entirely

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consists of medium-bedded to massive carbonates of light-grey–brownish-yellow colour, which were deposited on an extensive carbonate platform, the ‘Esfandiar Platform’ of Fursich et al.

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(2003). Most of the Esfandiar limestones, however, are fine-grained mudstones or wackestones

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with an impoverished fauna of the interior platform. The Shemshak Formation is a 2000 m-thick,

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well stratified sequence of weakly metamorphosed to unmetamorphosed sedimentary rocks. The overlying 1500−2300 m-thick Lower Cretaceous Sedimentary Sequence (LCSS) consists of

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unmetamorphosed interlayered, sedimentary rocks, including micro-conglomerates, sandstones,

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pyritic black siltstones, shales, dolomitic limestones, and dolomites (Fig. 3). Representative LCSS stratigraphic columns of the southern Yazd basin are shown in Fig. 4. In this basin, the lower Cretaceous Sangestan syn-rift sequence (with an overall maximum thickness of 400 m) unconformably overlies the Jurassic Shir-Kuh granite and metamorphic Shemshak Formation (Figs. 3 and 4). This Formation (Sangestan) displays abrupt changes in thickness, attaining 400 m in the Kharkuh and Farahabad area but tapers rapidly westward (<10 m) as it approaches the edge of the basin (Fig. 4). The post-rift sequence (sag phases) of the LCSS is marked by Taft and Abkuh Formations (Fig. 4). This sequence, deposited in marine environments, is well exposed

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Journal Pre-proof from Mehdiabad area in the east to Mansourabad area in the west (Fig. 4). It varies in thicknesses from 80 to over 450 m. The transition from the syn-rift to the post-rift sequence of the LCSS implies major changes in the sedimentary regime, from dominantly siliciclastic to calcareous (Fig. 4). The economically most important Zn-Pb (-Ag) deposits of southern Yazd basin occur within the Taft Formation of this sag phase (Fig. 4).

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2.2. Local stratigraphy of the Farahabad area

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In the Farahabad area, the LCSS includes, from the base to the top, the following

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lithostratigraphic units (Nabavi 1972; Maghfouri et al. 2017; Maghfouri 2017) (Figs. 4 and 5): (1) the Sangestan Formation; (2) the Taft Formation; and (3) the Abkuh Formation (or Darreh-

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Zanjir Formation).

The succession starts with conglomerates and sandstones (Sangestan Formation) filling a

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pronounced palaeo-relief of basement rocks of various ages belonging to the Shir Kuh granite and the metamorphic Shemshak Group (Upper Triassic–Middle Jurassic) (Fig. 4). A basal clast-

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supported, polymict, pebble-boulder conglomerate unit consists of slate, marble, black cherts,

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milky quartz, quartzites, sandstones, and carbonates (red oolites, dolostones) (Fig. 5). The matrix is coarse-grained, arkosic sandstone. The thickness of this Formation is very variable, reaching several hundred metres at the Farahabad area (Maghfouri 2017; Maghfouri et al. 2017) (Fig. 4). Locally, there are small exposures of basic volcanic rocks along with the sediments in the lower part of Sangestan Formation (Fig. 4). The Sangestan Formation is conformably overlain by Taft Formation (Fig. 5A). The Taft Formation consists of thin-bedded to massive, dark-coloured sandy dolomite and limestone with abundant orbitolinids and rudists. The Formation is up to 300 m thick, cliff-forming, and often hardly accessible; complete sections are thus difficult to

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Journal Pre-proof measure. The base of this Formation is predominate silty–sandy dolomite, limestone and siliciclastic sediments (Fig. 5B). The Taft Formation includes two units: the lower unit (30–40 m in thickness) is composed of organic matter-rich limy sandstone, sandy dolomite, shale and sandstone (host of mineralization in the Farahabad deposit) (Figs. 4, 5C and 6A), while the upper unit (30–40 m in thickness) includes dolomite and thick bedded limestone (Fig. 5A). Finegrained euhedral to subhedral dolomite (regional dolomite (DR)) (lower unit) occurred as a

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replacement in the Taft limestone (Fig. 6B). This regional dolomitization (DR) forms very sharp

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fronts, and sometimes layers that parallel the bedding (Fig. 6B). DR dolomite mainly occurred as

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a replacement of the micritic matrix (Fig. 6B). Regarding the development of porosity in limestone due to dolomitization, we propose that this regional event had a key role in providing

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the necessary open spaces for the later precipitation of hydrothermal dolomite (D H) and

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associated sulfide minerals (Fig. 6B,C). The Taft Formation is conformably overlain by thick-

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bedded carbonates of the Albian Abkuh Formation (Nabavi 1972) (Fig. 4), consisting of thickbedded to massive chert-bearing shallow-water limestones. The Abkuh Formation reaches up to

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50 m in thickness and is part of a large-scale carbonate platform system that characterized wide

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parts of the Yazd Block at that time (Wilmsen et al. 2013). 3. Sampling and analytical methods This study focused on the mineralogical and geochemical of hypogene mineralization in the Farahabad deposit. Determination of the mineralogy and paragenesis of the ore horizone in the Farahabad deposit is based on logging of drill cores and petrographic studies of over 45 polished thin and thick sections, supplemented by XL30 scanning electron microscopy conducted at Tabriz University, Iran.

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Journal Pre-proof Sulfur isotope data were obtained from pyrite, sphalerite, and galena separated from mineralized drill core samples. Nearly pure sulfides were separated using a microdrill under a binocular microscope. Samples were analyzed by mass spectrometry using a Delta C Finnigan MAT continuous flow isotope-ratio mass spectrometer with a TC-EA elemental analyzer. These analyses were carried out at Institute of Geochemistry, Chinese Academy of Sciences (n=20). The results are given as δ 34 S‰ values relative to the V-CDT (Vienna—Canyon del Diablo

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Troilite) standard. The analytical precision is within ±0.1‰ at 1 σ.

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C and O isotope analyses were carried out at the State Key Laboratory of Environmental

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Geochemistry, Institute of Geochemistry, Chinese Academy of Sciences, by using a Finnigan

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MAT-253 mass spectrometer. Carbonate reacts with 100% phosphoric acid (H3 PO4 ) to produce

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CO2 . The analytical precision (2 δ) is ±0.2‰ for δ 13 C value and ±2‰ for δ 18 O value. Carbon and oxygen isotope compositions are reported relative to Pee Dee Belemnite (PDB). δ 18 O SMOW =

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1.03086 * δ 18 OPDB + 30.86 (Friedman and O’Neil, 1977).

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Doubly polished wafers using standard techniques were prepared from seven samples Micro-thermometric

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collected in the presumed feeder zone of the Farahabad deposit.

measurements of fluid inclusions were performed on a Linkam THMS 600 combined heating/freezing stage with a German Zeiss microscope at the Iranian Mineral Processing Research Center (IMPRC), Karaj, Iran. The system was calibrated using synthetic fluid inclusion standards for CO 2 (– 56.6 °C), the freezing point of H2 O (0 °C) and critical point of H2 O (374.1 °C). This device can measure temperatures ranging from –196 °C to +600 °C. The precision of measurements below 0 °C was ± 0.2 °C and for homogenization temperatures was ± 0.6 °C. Salinities of aqueous inclusions were calculated using the equation of Bodnar and Vityk (1994) for the NaCl–H2 O system. 9

Journal Pre-proof 4. General features of ore facies in the Farahabad deposit Zn-Pb (-Ag) mineralization of the Farahabad deposit occurs along one horizon that extend over a length of >220 m (Fig. 5). The trend of the mineralization is SW−NE, dipping 10°-15° to the northwest, consistent with the bedding of the host rock (Fig. 7). Ore horizon is hosted in the organic matter-rich limy sandstone, sandy dolomite, shale and sandstone of the Taft Formation

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(Figs. 5C, 6A and 7). Through examination of drill holes and cross sections derived from drill hole information, a stratigraphic architecture model has been developed (Fig. 7). The Taft

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Formation at Farahabad displays a complex stratigraphic architecture interpreted to be controlled

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by synsedimentary faults (Fig. 7). The lower member siliciclastic to mixed siliciclastic-carbonate

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sandstone and turbidites (Fig. 7) occur in the hanging wall of the normal synsedimentary fault.

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The deposit lies north of, but close to, the normal fault, which is believed to have been active during deposition of the Taft Formation (Maghfouri 2017) (Fig. 7). The mineralized zone is

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located about 80 to 100 m below the surface and is concentrated along the contact between the

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shale and sandstone unit and the dolomite unit, forming a stratabound lens of ore zone which is approximately 20–30 m thick and 250 m long (Fig. 7). The extent of mineralization is limited by

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normal fault to the north (Fig. 7), although drilling is needed to test the presence of the mineralization below and northward of the fault. The Farahabad deposit consists of three ore facies, which are, when moving away from synsedimentary normal fault (Fig. 7): (1) a stockwork zone, formed by sulfide veins accompanied by carbonate or quartz; (2) a massivereplacement ore, formed by a massive sulfide zone with sulfide breccias; and (3) a bedded ore, consisting of a banded body of pyrite, galena, and disseminated sphalerite. - The stockwork zone (or feeder zone) consists of irregular veins of fine-grained dolomite, quartz, pyrite, sphalerite, galena, and dolomite, cutting altered and brecciated limy sandstone, 10

Journal Pre-proof sandy dolomite, shale and sandstone (Fig. 8A, D). These veins are mainly observed as adjacent to the syn-sedimentary normal fault (Fig. 7). The feeder zone is defined by numerous veins (from 1 mm to 5 cm thick) of quartz with pyrite and irregular proportions of galena, and minor sphalerite (Fig. 8A). This feeder or stockwork zone underlying the massive-replacement sulfides is similar in texture as that described elsewhere and is interpreted as the zone through which hot ascending and mineralizing hydrothermal fluids passed (Valdes-Nodarse 1998; Lydon 2000).

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Subaqueous debris flow deposits are locally abundant adjacent to the synsedimentary normal

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fault. The debris flow deposits contain shale, siltstone, carbonate and sometimes sandstone clasts

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of the stratigraphically underlying turbidites and are either matrix or clast supported with sandy to muddy matrix; matrix-poor breccias are also present. Clasts are commonly less than 1 to 7 cm to elongate,

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in diameter and generally rounded

occasionally displaying compaction or

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synsedimentary folding (Fig. 8A, B). No clasts of Sangestan Formation sedimentary rocks have

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been recognized, but clasts of mineralized and hydrothermally altered rock of the Taft Formation are locally present. Correlation of the ore- bearing debris flow in the Taft Formation with the

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major ore mineralization in the hanging wall of the normal fault (Fig. 7) suggests that both

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parameters reflect periods of major normal synsedimentary faulting and subsidence, although on significantly different scales. The presence of ore mineralization at Farahabad and along the northern margin suggests synsedimentary faults formed pathways for escape of hydrothermal fluids. -

The massive-replacement ore facies overlie a feeder zone which is characterized by

carbonate-quartz alteration and was likely related to a major hydrothermal brecciation event (Fig. 7). The massive sulfide has irregular edges and commonly shows dissolution contacts with the host rocks, suggesting that the sulfides replaced host rock minerals (Figs. 7 and 9). Locally, the 11

Journal Pre-proof sulfides fill cavities of the host rock (Fig. 9C). Sulfide minerals in this zone comprise 10 to 50 vol percent of the rock. Interbedded with the massive-replacement ore facies are discontinuous lenses of variably altered dolomitic breccias (Fig. 9A, D). Hydrothermally altered relicts of the host rocks can also be observed within the massive sulfides (Fig. 9A, B); some of these relicts include dolomite and sandstone rocks. The massive-replacement ore facies, located in the hanging wall of the normal fault, occur above the main debris flow deposits (Fig. 7) and form a

One of the most characteristic features of the Farahabad stratiform ore lenses (bedded ore

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-

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~60-m-long blanket that extends from the debris flow north to near the bedded ore facies.

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facies) is the occurrence of regular, fine layered sulfide-rich sedimentary rocks (Figs. 7 and 10).

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Compared to the massive-replacement ore, the bedded ore facies is generally thinner (about 2-3

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m thick) and characterized by low grade ores in the northern part of the deposit (Figs. 7 and 10). This facies displays stratiform banded sulfide-rich layers, concordant with the bedding (Fig. 10).

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Carbonate beds in the bedded ore facies are laterally discontinuous, whereas carbonate beds in

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the massive-replacement ore facies are mappable throughout the deposit. These beds range from <1-cm-thick carbonate layers in the mixed siliciclastic-carbonate turbidites to massive carbonate

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beds in excess of 1 m thick. There are many sedimentary structures in the bedded ore facies as follows: Bedding, lamination, laterally grades, graded bedding and cross bedding are the major characteristic sedimentary structures that are visible at different scales (Figs. 10 and 11). Normally in the bedded ore facies ore-bearing graded turbidites with bases containing coarse clasts are common and grade upward from siliciclastic sedimentary material to dolomitic caps (Figs. 10 and 11); siltstone and fine sandstone beds occur locally. Individual turbidites form orebearing siltstone beds that are typically several millimeters to several centimeters thick and are mostly planar bedded, although disruption due to slumping is locally prominent (Figs. 10 and 12

Journal Pre-proof 11). The ore-bearing sandstones are normally graded, have thicknesses between 0.1 and 5 cm and have sharp bases (Fig. 11). Detrital silt-sized quartz, dolomite, and mica constitute the main components of the beds; minor feldspar is locally present. 5. Ore Mineralogy and Paragenesis Galena and Sphalerite are the principal ore minerals in a relatively simple mineral

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assemblage that also contains iron sulfide, including pyrite (Fig. 12 and 13). Based on the broad

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textural and mineralogical associations observed within the deposit, mineralization has been

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classified into two stages: stage I and stage II (Fig. 13).

(Sph1),

fine-grained

disseminated

and

laminated

galena

(Gn1),

and

minor

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sphalerite

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The first mineralization stage (stage I) consists of framboidal and disseminated pyrites (Py1),

disseminated dolomite (Figs. 12A, B and 13). This stage of mineralization formed in the bedded

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ore facies (Fig. 13). Individual beds are generally thicker lower in the stratigraphy. The pyrite (Py1) in this ore facies is predominantly very fine grained (Fig. 12A). Pyrite is commonly

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concentrated along wispy dark organic matter-rich laminations. Framboidal pyrite (Py1) is the

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earliest formed generation and the range of framboid sizes (approximately 5–50 μm) (Fig. 12A) indicate formation from H2 S-bearing seawater within shallow, unconsolidated sediments and/or within the water column during syngenesis or in early diagenesis (Berner et al. 2013). The pyrite (Py1) is hosted in the limy sandstone and some of them were replaced by sphalerite (Sph2) and galena (Gn2) veinlets. The presence of framboidal pyrite reflects rapid crystallization on the seafloor caused by mixing of ascending hydrothermal fluids with cold ambient seawater (Herzig and Hannington 1995).

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Journal Pre-proof The second paragenetic stage (stage II) is the main ore stage and is characterized by the development of the massive-replacement textures (Figure 9C) and vein mineralization. This stage includes coarse-grained pyrite (Py2), sphalerite (Sp2) and galena (Gn2) (Figs. 12B, C, D). A decrease in carbonate content within the more strongly sulfurized zones suggests sulfide (stage II) replaced fine-grained sedimentary dolomite. The host rock clasts are cemented by galena (Gn2) and sphalerite (Sph2). This stage is characterized by abundant massive galena (Gn2) co-

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precipitated with medium-grained hydrothermal dolomite (DH). Stage II in the feeder zone is

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dominated by coarse-grained galena occurring as small veins cutting across the host rock of

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dolomitized limestone (Fig. 8D), or locally cutting the earlier pyrite (Py1), is indicating that the galena probably precipitated slightly later than the Py1. The textural relationship of the base

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metals to the two types of minerals led Maghfouri (2017) to conclude that stage I minerals

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formed by exhalation and precipitated from an overlying water column during variable input of

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clastic sediments during early diagenesis by biogenic reduction of sulfate (stage I), whereas stage

II).

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II minerals were precipitated later during a sub-seafloor hydrothermal replacement event (stage

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In addition, a supergene mineralization is also reported at Farahabad deposit (Fig. 13). Smithsonite, hemimorphite, hydrozincite, cerussite, anglesite and iron (hydr-) oxide occur as the secondary minerals, in accordance with the common Zn supergene mineralogy reported in the South Yazd Basin (Maghfouri et al. 2018b). These minerals mainly occur within veins, which postdate the faults and fractures related to the uplift tectonic event, which affected the South Yazd Basin during the Tertiary (Maghfouri et al. 2018b).

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Journal Pre-proof 6. Sulfur isotope compositions Sulfur isotopes were analyzed in sphalerite (n=9), galena (n=9), and pyrite (n=2), from the bedded (stage I), massive-replacement and stockwork ores (stage II) of the Farahabad deposit (Table 1). Measured δ 34 S values from sulfide minerals show an extensive distribution (Table 1, Fig. 14) with one population ranging from –19.47 to –14.6 ‰, and the other ranging from –8.8 to

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+3.9 ‰. Galena rich massive-replacement ore facies has δ 34 S values that range from -17.6 to 16.1 ‰ (mean of -16.98.2‰) (Fig. 14). Massive-replacement ore facies sphalerite minerals have

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δ34 S values that range from -8.8 to +3.8 ‰ (mean of -9.3‰) (Fig. 14). Sulfide minerals from

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stringer veins have δ 34 S values that range from -14.6 to +3.9 ‰ (mean of +13.6‰). Pyrite-rich

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bedded ore, in which pyrite has primary depositional textures such as very fine grained layers

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and framboids, have low δ 34 S values ranging from –19.0 to -19.2 ‰ (Fig. 14). Results from bedded ore facies samples follow: –19.0 to -18.5 per mil in sphalerite, and -19.47 to -18.16 ‰ in

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galena from bedded ore facies (Fig. 14).

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5.5. Carbon and oxygen isotope compositions

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The C–O isotopic composition of carbonates is listed in Table 1 and illustrated in Fig. 15. The host Taft Formation limestones have δ 13 C values ranging from 1.4 to 2.1‰ (mean = 1.83‰) (Fig. 15B), within the range typical of marine Cretaceous limestone (Veizer and Hoefs 1976; Land 1980). The δ 18 O values of host limestones range from 6.5 to 9.4‰ (Fig. 15A). The δ 13 C values of the ore-dolomite (DH), associated with the sulfide mineralization, vary between 1.6 and 2.7‰ (Fig. 15B). The δ 18 O values for the ore-dolomite range from 10.4 to 13.1‰ (mean = 12.17‰), and are significantly higher than those measured in the host limestone (Fig. 15A).

15

Journal Pre-proof 5.6. Classification, distribution and microthermometry of fluid inclusions Fluid inclusions suitable for microthermometric study were found mostly in dolomite; with rare examples observed in sphalerite. The criteria suggested by Roedder (1984) were used as a standard for distinguishing the origin of inclusions. Primary fluid inclusions (following the criteria of Roedder 1984) were observed as array and isolated along growth zones in the crystals.

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Secondary fluid inclusions were observed in healed fractures and were trapped after crystal growth. Only primary fluid inclusions that are associated with the main stage of mineralization

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processes were selected for fluid inclusion studies and secondary fluid inclusions were

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disregarded. A systematic description of fluid inclusion types based on the phases on estimated

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volumetric proportion of vapor to liquid present at room temperature is summarized in figure 1.

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Type I inclusions: Type I inclusions are characterized by a dominant mono-phase fluid (Fig.

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16). They can be subdivided into type Ia mono-phase vapor-rich and type Ib mono-phase liquidrich on the basis of gas and liquid contents. Type Ia and type Ib fluid inclusions are

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petrographically indistinguishable at room temperature, being of similar size and shape, and

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exhibiting a comparable apparent monophases character. Fluid inclusions within this group are generally between 4-8 μm in size and show equant to negative crystal morphologies. In most cases, they are isolated or occur in clusters of three or four. Type II inclusions: Type II inclusions are aqueous, two-phase inclusions with variable vaporto-liquid ratios and are dominant in nearly all exanimated samples (Fig. 16). They can be divided into two basic categories in terms of volumetric proportion of vapor and liquid phase in the inclusions along with consideration of their origin. Type IIa inclusions are relatively vapor-rich and probably primary in origin, although the unambiguous distinction normally appears

16

Journal Pre-proof impossible (Roedder 1984). In general, a vapor bubble constitutes typically 50 to 60 vol percent of the inclusion cavity, but the vapor phase in some inclusions was found to exceed 70 percent by volume of the inclusion (Fig. 16). Fluid inclusion of this type usually occurs as isolated individuals or aligned along a growth zone. Most such inclusions are relatively small and have rounded to ellipsoidal shapes, but others, notably the larger inclusions (>15 μm), are quite irregular (Fig. 16). Most commonly, type IIa inclusions are observed in the core of the host

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minerals, though some can occur at the grain boundaries. In contrast, type IIb inclusions are

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liquid-rich and almost exclusively primary origin. They are generally present along microcrystals

-p

and range from 8 to 15 μm in diameter (Fig. 16). At room temperature, fluid inclusions within this subgroup contain a small vapor bubble (less than 25% of the inclusion volume) within a

lP

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dominant aqueous liquid phase, together with an elongate morphology. Fluid inclusions of type II inclusions freeze to ice upon cooling below -100°C. During

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heating initial melting was observed between -21.5°C and -59.0°C, with a population centered on

ur

-23°C. Such initial melting temperatures indicate that NaCl±KCl probably is the dominant dissolved salt in the inclusion fluids. Other salts must be percent because the initial melting

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temperatures of some inclusions were considerably depressed relative to NaCl. Given the presence of carbonate alteration associated with mineralization as well as the occurrence of dolomite in the paragenesis, the most important of these was possibly CaCl2 . Eutectic temperatures are around -52 and -26°C, indicating that the solution was dominated by the H2 O– NaCl or H2 O–NaCl–CaCl2 (cf. Crawford 1981; Hall et al. 1988; Davis et al. 1990). Final melting temperatures (Tmice) range from −0.7 to −15.8 °C (Table 4). Salinities were calculated using the equations of Bodnar and Vityk (1994) for the simplified H2 O– NaCl and H2 O-NaCl+CaCl2 systems. Measured final melting temperatures correspond to salinities of 0.99 to 16.70 wt% NaCl 17

Journal Pre-proof eq., with an average of 9.50 wt% NaCl eq. (Fig. 17B, Table 2). Homogenization temperatures (Th ) range from 115 to 234 °C, with an average value of 177.10 °C (Fig. 17A, Table 2). 6. Discussion 6.1. Origin of sulfur In most cases, the predominant source of sulfur in stratiform, sediment-hosted Zn-Pb-(Cu-

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Ag-Ba) deposits is interpreted to be seawater sulfate (Leach et al. 2005, 2010; Ohmoto 1996).

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Juvenile sulfur originated from a magmatic source, as interpreted from most igneous copper

-p

deposits (porphyry), has δ 34 S composition of sulfide minerals ranging from -5 to +5 per mil (Rye

re

and Ohmoto 1974, 1996; Seal 2006). There is no geological evidence for igneous activity linked

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with the Farahabad deposit, wither as tuffaceous horizons related to volcanism or granitic intrusions associated with magmatic activity. The δ 34 S values of sulfides range from –19.47 to

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+3.9 ‰ (Fig. 14), well low values expected for magmatic source and less than values expected for coexisting Cretaceous seawater sulfate (+15 to +21‰). However, the processes by which

ur

sulfate is reduced to sulfide and delivered to the depositional site are still debated and remain

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contentious because there are several possible mechanisms to generate the reduced sulfur (Ohmoto and Goldhaber 1997). The transformation of sulfate to sulfide either by bacteriogenic sulfate reduction (BSR) or by thermochemical sulfate reduction (TSR) implies isotopic fractionations greater than 20‰ (Ohmoto et al. 1990). Sulfide generated during BSR may possess δ 34 S values that range from extremely negative to positive, which is a function of whether the environment is open (e.g., water column) or closed (e.g., sediment-pore waters) to the diffusion of sulfate (Gadd et al., 2016). Negative δ 34 S values exist because of the large kinetic fractionation between sulfate and sulfide associated with BSR. Biogenic sulfide is recognized as

18

Journal Pre-proof a major component in some sediment-hosted Zn-Pb districts (e.g., Irish ore fields), particularly where a vast majority of the base-metal sulfides possess negative δ 34 S values (Fallick et al., 2001; Wilkinson et al., 2005). Biogenic sulfide has also been implicated as the predominant source of sulfur for base-metal sulfides of the Mansourabad deposit (Maghfouri et al., 2018a). δ34 S values of sulfide minerals in the different facies of Farahabad deposit are negative, with the highest values corresponding to sulfides of the feeder zone (Fig. 14). So in the Farahabad

of

deposit, sulfide source is reduction of sulfate biogenically in the sediment or water column

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through which the hydrothermal system has vented. By mixing there would be a decrease of δ 34 S

-p

in the sulfides as is observed in Figure 14. Bacterial reduction of seawater sulfate resulates in a δ34 S values for H2 S 15±5 ‰ lower than that for SO 4 -2 (orr 1975). The fractionation observed for

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a typical sample from different ore facies is 18.8 ‰. This observation, combined with the

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interpretation of a biogenic origin for the framboidal pyrite (Gadd et al. 2016), gives support to

na

the thesis that sulfate-reducting bacteria contributed to sulfide formation. In-situ biogenic sulfate reduction would only have been possible below 120°C (Jorgenson et al. 2004) although some

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modern bacteria groups attain optimum reduction rates at 85-102°C (Elsgaard et al. 1994). Fluid

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inclusion from ore-bearing dolomite range from 115 to 234 °C and it is unlikely that sulfurreducing bacteria could survive (Trudinger et al. 1972). Sulfide could have been produced biogenically in cooler environment (the basal water column) or between hydrothermal pulses. Metal-bearing hydrothermal fluids may then have vented onto the sea floor, mixed with reduced sulfur and precipitated sulfide minerals. Maghfouri et al. (2018a) suggested that dissolved sulfate was a major component of metalliferous brines that formed the Lower Cretaceous Mansourabad Zn-Pb(-Ag) deposit in the Southern Yazd basin; Iran. In their model, dense brines exhaled into an anoxic water column and ponded on the seafloor. Metal sulfide precipitation was triggered by

19

Journal Pre-proof biogenically reduced sulfur that diffused into the brine from sediments below and from the water column above. It is likely that BSR played an important role in the formation of sulfide minerals in the Farahabad in a manner similar to Mansourabad. 6.2. Water depth during sulfide deposition and nature of the ore fluids Few reliable criteria exist for the quantification of water depth for exhalative deposits:

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palaeontological, sedimentological, and geochemical criteria can only provide an indication of

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water depth (Plimer 1981). Evidence for deep-water settings of exhalative deposits (> 500 m) is

-p

generally the lack of shallow-water features, although deposits present in thick sequences that display graded bedding (e.g. Broken Hill-Laing 1980) are probably the best examples of deep-

re

water deposits. The presence of associated fossiliferous clastic rocks (e.g. crinoidal carbonates at

lP

Tynagh and Silvermines-Russell 1975), stromatolitic carbonates (McArthur River-Murray 1975,

na

Walker et al. 1978), evaporites (McArthur River-Williams 1978), reefs (Meggen-Gwosdz and Krebs 1977), and arkoses and conglomerate (Queen Bee-Sangster 1979) provides evidence of

ur

shallow-water settings (Plimer 1981). The intimate association of acid pyroclastic rocks (Iberian

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Pyritic Belt-Strauss and Madel 1974; Kuroko deposits-Lambert and Sato 1974) may be taken to indicate that hydrothermal phreatic explosions took place when the total vapor pressure exceeded the total confining pressure plus the mechanical strength of the rocks. Such explosions would certainly take place in shallow water, where the total confining pressure is low, but, even at depths at or below the carbonate compensation depth (presently~4000 m), the hydrostatic pressure of sea water is only in the order of 0.5 kb, and, hence, such explosions could still take place in very deep water.

20

Journal Pre-proof Fluid inclusion measurements of hydrothermal dolomites associated with the sulfides in the Farahabad deposit give homogenization temperatures averaging 177.10°C. Salinity for these inclusions is estimated to average 9.50 wt% NaCl eq (Fig. 17). The fluid inclusions do not show evidence of boiling. Phase separation or boiling can result in the formation of high salinity fluids, but none of the adjacent primary fluid inclusions from Farahabad display variable liquid to vapor phase ratios which would provide evidence that boiling or phase separation occurred at the site

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of mineral deposition. A rapid temperature drop on contact with seawater could account for

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sulfide precipitation. This is possibly reflected by the disequilibrium in the sulfur isotope data.

-p

Because the fluids did not boil or undergo phase separation at the site of mineral deposition, a pressure correction corresponding to the weight of the water column overlying the deposit at the

re

time of formation must be applied to obtain true trapping temperatures for the inclusions (Potter

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1997; Peter and Scott 1993). The water depth is unknown; however, it is possible to constrain the

(Bischoff and

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pressure correction by considering the physical and chemical properties of hot NaCl solution Pitzer 1989;

Peter and

Scott 1993)

in light

of the

fluid

inclusion

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microthermometric data. For a fluid with a maximum temperature of 234°C and 9.50 wt.% NaCl

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equiv., a confining pressure of about 30 bar is required to suppress phase separation or boiling (Fig. 17C). Thus, the Farahabad deposit formed under a water column of at least 300 m (Fig. 17C). Such a depth is very similar to that observed for Mansourabad deposit (Maghfouri et al. 2018a). The presence of sandstone and micro-sandy conglomerate and abundant graded bedding (Fig. 11) in the host rocks of Farahabad deposit are sedimentological evidences from shallowwater condition in the formation of sulfide mineralization. The above data suggest that some exhalative deposits associated with felsic volcanism probably formed at seawater depths greater than 1000 m (Fig. 18A), whereas some associated with mafic volcanism formed at depths greater

21

Journal Pre-proof than 2000 m. Deposits associated with thick sequences of sedimentary rocks (Farahabad, Irish deposits and McArthur) probably formed over a very narrow depth range (Russell et al. 1981), whereas other shale hosted deposits (e.g. Gamsberg, South Africa; Broken Hill, Australia; Sullivan, Canada) probably formed over a greater range (Plimer 1981) (Fig. 18A). There are several possible sources for the carbon in the Farahabad carbonates: magmatic,

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seawater, diagenetic degradation of organic matter in the sediments, dissolution of primary marine carbonate, oxidation of methane, or combinations of these sources and processes. These

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sources may be distinguished isotopically. Modern seawater and primary marine carbonates have

-p

δ13 C compositions near 0‰. The δ 13 CPDB and δ 18 OSMOW values for mantle, marine carbonate and

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organic matter range from −4.0‰ to −8.0‰ and +6.0‰ to +10.0‰ (Taylor et al. 1967), −4.0‰

lP

to +4.0‰ and +20.0‰ to +30.0‰ (Veizer and Hoefs 1976), and −30.0‰ to −10.0‰ and +24.0‰ to +30.0‰ (Liu and Liu 1997), respectively (Fig.15C). The δ 18 O values of host

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limestones in the Farahabad deposit are lower than hydrothermal dolomites (Fig. 15C). In

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addition, recalculated δ 18 OH2O values are relatively low for host limestones (mean +7.83‰) compare to hydrothermal dolomites (mean 12.17‰). Host rock limestone from the Farahabad

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deposit has higher δ 13 CPDB and δ 18 OSMOW values than mantle and sedimentary organic matter, but similar to marine carbonate rocks and the ore hosting dolomite (Fig. 15C). This indicates that mantle and organic matter may not have contributed significant carbon to the hydrothermal fluids. Thus the carbon and oxygen isotopic data indicate that the CO 2 in the ore-forming fluid likely originated from marine carbonate rocks. Therefore the Lower Cretaceous Taft Formation carbonate rocks were likely the main supplier of the CO 2 in the ore-forming fluid.

22

Journal Pre-proof 6.3. Sub-seafloor replacement and factors controlling replacement Early venting of the exhaling hydrothermal fluid in the basin led to the precipitation of an early assemblage of the fine grain size of the minerals in the bedded ore facies (stage I, including framboidal pyrite, laminated sphalerite and galena) in reaction to the mixing with seawater (Fig. 18B). Although the observed mineralogy and textures of stage I mineralization in the Farahabad

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deposit are similar to those reported in many sediment-hosted Zn-Pb (SEDEX) deposits worldwide (Goodfellow et al. 1993; Lydon 1995; Large et al. 2002; Sangster 2018), textures of

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stage II are unlike those believed to have formed by exhalative processes. The fine alternations

-p

of monomineralic sulfide bands and sandstone that characterize shale-hosted Zn-Pb deposits

re

such as Sullivan (Lydon et al. 2000), Rammelsberg (Large and Walcher 1999), Mehdiabad

lP

(Maghfouri et al. 2015; 2019; Maghfouri 2017) and McArthur River (Large et al. 2002) are rarely met in the stage II mineralization at Farahabad (Fig. 18B). The stage II mineralization

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(massive-replacement ore facies) is the largest of the sulfide bodies located along the normal

ur

fault system. This body has a maximum thickness of 30 m adjacent to the fault, progressively thinning and decreasing in grade away from the fault (Fig. 18C). The location, geometry, and

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texture patterns of massive sulfide lenses, coupled with evidence of fluid flow (hydrothermal alteration) along this fault below the orebody, indicate that the major NW-SE fault in the relay zone between them provided numerous pathways for metal-bearing, hydrothermal fluids to enter the Taft Formation (Fig. 18). Replacement of carbonate layers in sandy limestone by sulfides (stage II) is evident in hand samples from the Farahabad deposit, and sulfide-replaced textures are common in carbonate rocks have been noted in some places in the Southern Yazd basin (Maghfouri 2017; Maghfouri et al. 2017; Maghfouri et al. 2018a). Hydrothermal fluids entering the sandy limestone along the normal synsedimentary fault system produced both hydrothermal 23

Journal Pre-proof dolomite (DH) and sulfides. The intimate association between replacement and ore strongly suggest a genetic link between replacing and sulfide mineralization (Fig. 18C). Overall, the sulfide and dolomite textures suggest that sulfides replaced carbonate rather than lithic turbidites. Breccia mineralization in a vent may be an acceptable explanation for the massive-replacement ore facies, where cross-cutting and replacement mineralisation occurred in the water saturated sediments immediately beneath the seafloor. Sub-seafloor replacement of carbonate within shale-

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dominated sedimentary sections has been invoked for numerous large zinc deposits that are

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generally considered part of the sedimentary exhalative (SEDEX) class, such as Century

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(Broadbent et al. 1998; Sangster 2018), Mount Isa in part (Chapman 2004), Mehdiabad (Maghfouri et al. 2015; Maghfouri 2017), Mansourabad (Maghfouri et al. 2018a), Rammelsberg

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(Eldridge et al. 1988) and Meggen (Geer 1988). At Meggen in Germany, both barite and sulfides

lP

are proposed to have replaced carbonates (Geer 1988). Textural evidence suggest that sulfide

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mineralization at Farahabad is formed by the replacement of carbonate-rich sediments rather than by exhalative processes (Fig. 18). Similar to Anarraaq deposit in the Red Dog district (Kelley et

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al. 2004), the Farahabad deposit exposes zones of massive-replacement high-grade ores, which

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are adjacent to a feeder zone close to a syn-sedimentary normal fault (Fig. 18B). Banded sulfide zones developed away from the feeder zone (Hitzman and Beaty 1996; Maghfouri 2017) (Fig. 18A). Such association has often been interpreted as a vent complex similar to that found in other well-known sediment hosted massive sulfide deposits; thus, it can be classified as a ventproximal sub-seafloor replacement SEDEX-type deposit (cf. Lydon 1995; Sangster 2002, 2018; Large et al. 2005; Leach et al. 2005; Goodfellow and Lydon 2007; Kelley et al. 2004; Wilkinson 2014). Interpretation of the timing of sulfide replacement in stratiform Zn-Pb-Ag sulfide deposits

24

Journal Pre-proof is still a debated question and opinions range from early syn-diagenetic to late diagenetic replacement (Kelley et al. 2004; Maghfouri 2017; Maghfouri et al. 2018a). The distribution of DR in the Farahabad deposit may have been important for controlling the replacement of carbonate layers by hydrothermal dolomite (DH) and sulfides. Pre-mineralization (DR) and syn-mineralization dolomitization (DH) have both been documented in the district. If

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some of the dolomite in the Farahabad deposit formed prior to introduction of metal-bearing fluids, dolomite-rich (DR) zones may have formed as permeable horizons to the fluids, thereby

ro

controlling fluid flow through carbonate layers. The close temporal and spatial association with

-p

synsedimentary faulting is the common controlling characteristic of Farahabad deposit. The

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significance of synsedimentary fault systems in the formation of this deposit is twofold; they

lP

provide the pathways from the aquifer toward the surface and control the generation of local basin, which act as morphological traps on the seafloor. The textural patterns, the presence of

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syn-sedimentary breccias, the increase of hydrothermal alteration and temperature of ore fluids

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toward the faults, suggest that this synsedimentary fault was active during ore formation and

7. Conclusion

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played a role as conduits for hydrothermal fluid flow.

The Farahabad deposit was formed in the sag-phase of an extensional sedimentary basin. In the Farahabad sequence >800 m of sediment accumulated in the Lower Cretaceous time. The deposit is hosted in the Lower Cretaceous Taft Formation. The lower half of the Taft Formation is rhythmically banded, consisting of sandy dolomite, dolomite, sandstone and silty limestone, contains the bulk of the sulfide mineralization at the Farahabad. The major components of the mineral deposit are hydrothermal dolomite, sphalerite, galena and pyrite.

25

Journal Pre-proof The Farahabad deposit shares characteristics with both on-saefloor mineralization and subseafloor replacement mineralization. The affinity of the Farahabad deposit to sub-seafloor replacement mineralization and its close spatial association to exhalative deposits in clastic units of the Taft Formation emphasizes the potential of sedimented extensional rift basins to host both types

of syngenetic-diagenetic

mineralization. This

deposit

could

reflect

sub-seafloor

replacement mineralization in the late stages of sedimentation, when dolomitic rocks had become

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predominant. The Farahabad deposit formed proximal to the source vent. Most of mineralization

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is replacement or breccia textured to indistinctly bedded and facies changes from one ore type to

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another occur over short distances (with away from synsedimentary fault). Sulfide veins occurring dominantly in the footwall sedimentary rocks are interpreted to be feeder veins. Host

lP

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rock replacement textures are common, particularly in dolomite rock. The temperature and salinity of the dolomite samples from Farahabad are remarkably similar

na

to those characteristics of dolomite samples from veins of the Mansourabad deposit in the southern Yazd basin. These data show that the ore fluids were probably less than 234°C and

ur

were moderately strongly saline (about 10–12 wt % NaCl equiv). The δ 13 CPDB and δ 18 OSMOW

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values of hydrothermal dolomites and host rock limestone’s shows that the CO 2 of the ore fluid was derived mainly from the marine carbonate rocks. δ34 S values of pyrite, sphalerite and galena range from –19.47 to +3.9‰. The overall range of δ 34 S is remarkably higher than typical magmatic values, suggesting that sulfides formed from the reduction of seawater sulfate by bacteriogenic sulfate reduction. Farahabad has many geologic, geochemical, and metallogenic similarities to the SEDEX-type deposits, formed by a combination of sediment replacement in the sub-seafloor and deposition on the sea floor processes (Fig. 17).

26

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Berberian M, King GCP (1981) Towards a palaeogeography and tectonic evolution of Iran. Can.

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Davis D, Lowenstein T, Spencer R (1990) Melting behavior of fluid inclusions in laboratorygrown halite crystals in the systems NaCl- H2 O, NaCl-KCl-H2 O, NaCl-MgCl2 –H2 O and NaCl-CaCl2 –H2 O. Geochim Cosmochim Acta. 54, 591–601. Eldridge CS, Compston W, Williams IS, Both RA, Ohmoto H, Walsche JL (1988) Sulfur isotopic variability in sediment hosted massive sulfide deposits as determined by using ion 28

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Gwosdz W, Krebs W (1977) Manganese halo surrounding Meggen ore deposit, Germany. Institution of Mining and Metallurgy Transactions. 86, 1373-1377. Haghipour A (1977) Geological Quadrangle Map of Posht-e-Badam. Tehran, Iran, Geological Survey of Iran. Hall DL, Sterner SM, Bodnar RJ (1988) Freezing point depression of NaCl-KC1-H2O solutions. Economic Geology. 83, 197–202.

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oxidation and a deep H2S sink generate isotopically heavy sulfides in Black Sea sediments.

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Geology. 99, 1577–1591.

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northern Alaska: Evidence for replacement of carbonate by barite and sulfides. Economic

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(editor), A guide to the stratigraphy and mineralization in the Broken Hill Block, New South Wales. Geological Survey of New South Wales, Record. 20, 71-85. Lambert IB, Sato T (1974) The Kuroko and associated ore deposits of Japan: a review of their features and metallogenesis. Economic Geology. 69, 1215-1236. Land LS (1980) The isotopic and trace element geochemistry of dolomite: the state of the art. In: Concepts and models of dolomitization, Zenger, D. H., Dunham, J.B., Ethington, R.L. (eds). Journal Sedimentary Petrology. 28, 87–110.

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Journal Pre-proof Large DE, Walcher E (1999) The Rammelsberg massive sulfide Cu-Zn-Pb-Ba deposit, Germany: An example of sediment-hosted, massive sulfide mineralization. Mineralium Deposita. 34, 522–538. Large RR, Bull SW, Yang J, Cooke DR, Garven G, McGoldrick PJ, Selley, D (2002) Controls on the formation of giant stratiform sediment-hosted Zn-Pb-Ag deposits with particular reference to the north Australian Proterozoic. University of Tasmania, Centre for Special Ore

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hosted lead-zinc deposits in Earth history: Economic Geology. 105, 593–625. Leven EJ, Davydov VI, Gorgij MN (2006) Pennsylvanian stratigraphy and fusulinids of central and eastern Iran. Palaeontologia Electronica. 9, 1-36. Lydon JW (1995) Sedimentary exhalative sulfides (SEDEX), in Eckstrand, O.R., Sinclair, W.D., and Thorpe, R.I., eds., Geology of Canadian mineral deposit types: Geological Survey of Canada. 8, 130–152.

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Cretaceous carbonate-clastic hosted Zn-Pb deposits in southern Yazd basin, with emphasis on

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Mehdiabad deposit. Unpublished Ph.D. thesis. Tabriz, University of Tabriz, Iran. 475 p. Maghfouri S, Hosseinzadeh MR, Rajabi A, Azimzadeh AM (2017) Facies analysis and stratighraphy position of carbonate-clastic hosted Zn-Pb-Ba mineralization horizons in the early cretaceous sedimentary sequence, Southern Yazd Basin: Scientific Quarterly Journal, Geosciences. 26, 102-115. Maghfouri S, Hosseinzadeh, MR (2018a) The early Cretaceous Mansourabad shale carbonate hosted Zn–Pb (-Ag) deposit, central Iran: An example of vent-proximal sub-seafloor replacement SEDEX mineralization. Ore Geology Reviews. 95, 20–39. 33

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Territory, Australia: II. The sulfide-S and organic-C relationships of the concordant deposits and their significance. Economic Geology. 73, 1036-56. Wilmsen M, Fürsich FT, Seyed-Emami K, Majidifard MR, Taheri J (2009) An overview of the stratigraphy and facies development of the Jurassic System on the Tabas Block, east-central Iran. The Geological Society, London, Special Publications. 312, 323–343. Wilmsen M, Fürsich FT, Majidifard J (2013) The Shah Kuh Formation, a latest Barremian – early Aptian carbonate platform of Central Iran (Khur area, Yazd Block). Cretac. Res. 39, 183–194. 38

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Ar/39 Ar Geochronology of Alteration

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Zarasvandi A, Liaghat S, Zentilli M, Reynolds PH (2007)

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and Petrogenesis of Porphyry Copper-Related Granitoids in the Darreh-Zerreshk and Ali-

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Abad area, Central Iran, Exploration and Mining Geology. 16, 11–24.

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Journal Pre-proof Figure captions Fig.1 Distribution map of sediment-hosted Zn-Pb deposits in the main tectonic elements of Iran; Al, Alborz zone; CIGS,Central Iranian geological and structural gradual zone; E, East Iran ranges; K, Kopeh-Dagh; KR, Kermanshah Radiolarites subzone; KT, Khazar-Talesh- Ziveh structural zone; L, Lut block; M, Makran zone; Oph, ophiolite belts; PB, Posht-e-Badam block; SSZ, Sanandaj-Sirjan zone; T, Tabas block; TM, tertiary magmatic rocks; UDMA, Urumieh-

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Dokhtar magmatic arc; Y, Yazd block; Z, Zabol area; Za, Zagros ranges (tectonic and structural

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map of Iran modified after Aghanabati 2004 and Alavi 1994).

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Fig. 2 Simplified geological map of the southern Yazd basin, showing the location of Zn-Pb

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deposits within this basin. Small rectangle shows the location of study area in the Fig. 3.

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Cretaceous sedimentary sequence.

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Fig. 3 Geology map of the Farahabad deposit showing the location of this deposit in the Lower

Fig. 4 Lithostratigraphic correlation diagram of the Lower Cretaceous Sedimentary Sequence

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(LCSS) in the southern Yazd basin showing the position of the deposits in the stratigraphic units.

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Fig. 5 Correlation between schematic cross section of the Farahabad deposit and sedimentary units of lower Cretaceous sequence. (A): General view of the Sangestan Formation, conformably covered by Taft Formation.

(B): Northeast-southwest cross section of the ore sequence with

extension of sulfide ore mineralization’s in the sandy dolomite, sandstone and silty limestones of Taft Formation. (C): Hand specimen photograph of sandstone, sandy dolomite with interbedded of silty limestone host rock in the Farahabad deposit. Fig. 6 (A): Hand specimen photograph of ore-bearing host rock facies in the Farahabad deposit. (B): Photomicrograph of silty limestone has been replaced by DR dolomite. (C): Microscopic 40

Journal Pre-proof picture represents replacement of DR dolomite with DH dolomite. (D): Photomicrographs of DH dolomite and sulfide minerals. Fig. 7 Southwest-northeast cross section of the ore sequence and extension of ore facies in the Farahabad deposit. Fig. 8 Hand specimen (A, B and C) and microscopic (D) photographs of sulfide vein-veinlet’s

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and mass flows textures in the stringer zone. (A): Feeder zone ore composed of veinlet’s of

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sulfide minerals crosscutting silty limestone and sulfide laminae. (B and C): Hand specimen photographs of ore-bearing (Sph2: sphalerite; Gn2: galena; Py2: pyrite) mass flows in the

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hanging wall of synsedimentary normal fault within feeder zone. (D): Veinlet of galena (Gn2)

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sulfide mineral in the feeder zone ore facies.

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Fig. 9 Hand specimen (A, B and D) and microscopic (C) photographs of the massive-

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replacement ore facies in the Farahabad deposit. (A): Hydrothermal dolomite (DH) host rock brecciated and replaced by stage II galena (Gn2) and sphalerite (Sph2) sulfide minerals (sulfides the primary carbonate folding).

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have retained

(B): Massive-replacement sulfide

texture

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consisting of sphalerite (Sph2) and galena (Gn2). (C): Photomicrograph showing textural banded replacement galena (Gn2) in the massive-replacement ore facies. (D): Hand

specimen

photograph of hydrothermal dolomite (DH) replaced by stage II galena (Gn 2) and sphalerite (Sph2) sulfide minerals. Fig. 10 Hand specimen photographs of the bedded ore facies in the Farahabad deposit. (A): Photograph of typical pyritic (Py1) laminae associated with galena (Gn1) and sphalerite (Sph1)rich laminae. (B, C, D and E): Photographs of sedimentary breccia and laminated ore-bearing limy siltstone and sandstone in the bedded ore facies. Textural and sedimentological relationship 41

Journal Pre-proof between the laminated ore-bearing limy siltstone and

sandstone and

overlying graded

sedimentary breccia bed, indicates that the orebody is synsedimentary in origin. (Sph: sphalerite; Gn: galena; DH: hydrothermal dolomite). Fig. 11 Main sedimentary structures in the bedded ore facies. (A and B): Ore-bearing sandstone laterally grades into silty sandstone. (C): Hand specimen photograph of graded bedding in the ore-bearing host rocks of bedded ore facies. (D): Cross- bedding of ore-bearing host rocks in a

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hand specimen scale.

Fig. 12 Reflected light photomicrographs of sulfide minerals in the Mehdiabad deposit. (A):

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Reflected light photomicrograph of framboidal pyrite (Py1) with sphalerite (Sph1) in bedded ore

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facies of the Farahabad deposit. (B): Early galena (Gn1) and sphalerite (Sph1) minerals crosscut

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by late stage pyrite (Py2). (C and D): Reflected light image showing intergrowth of galena (Gn2)

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with sphalerite (Sph2).

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Fig. 13 Paragenetic sequence of mineralization stages in the Farahabad deposit. Fig. 14 (A): Histogram of δ 34 S values in the different facies of the Farahabad deposit. (B):

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Frequency distribution of sulfur isotope values for sulfides from the Farahabad deposit. Fig. 15 (A): Frequency distribution of oxygen isotope values for limestone host rocks and hydrothermal dolomite from the Farahabad deposit. (B): Frequency distribution of carbon isotope values for limestone host rocks and hydrothermal dolomite from the Farahabad deposit. (C): Plot of δ 13 CPDB vs. δ 18 OSMOW for the Farahabad Zn-Pb deposit. Fig. 16 Photomicrographs of fluid inclusion types. (A): TIIb fluid inclusion in the dolomite mineral. (B): TIIb and TIb fluid inclusions. (C, D and E): TIIb fluid inclusion of dolomite mineral in the feeder zone. (F). TIIa fluid inclusion with irregulars shape. 42

Journal Pre-proof Fig. 17 (A): Histogram of homogenization temperatures of fluid inclusions in Farahabad deposit. (B): Histogram of salinity for fluid inclusions in dolomite minerals. (C): Temperature-pressurecomposition relations for seawater (3.2 wt.% NaCl; Bischoff and Rosenbauer, 1985) and a fluid of 10 wt.% NaCl equiv. salinity (after Roedder and Bodnar, 1980). Primary Farahabad fluid inclusions display constant phase ratios and indicate that the fluids were not boiling at the site of mineral deposition. This allows constraints to be placed on the minimum water depth at which

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the inclusion data, a representative primary inclusion fluid at 234 and 9.50 wt.% NaCl eq.

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(salinity only a minor effect at this temperature) must have been trapped at a minimum pressure

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of 30 bar in order for boiling to be suppressed.

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Fig. 18 (A): Schematic relation between water depth at the site of exhalation in the different

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types of exhalative ore deposits (modified after Plimer, 1981). (B and B): Schematic diagram showing the inferred depositional environment of the Farahabad deposit. (A): Early venting of

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the exhaling hydrothermal fluid on the seafloor would have led to the precipitation of an early assemblage of the fine grain size minerals (stage 1) in the bedded ore facies. (B): Continued up-

ur

flow of high-temperature, metal-bearing, dense hydrothermal fluid into and along the bedded ore

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facies (ore stage I) and dolomite in the sub-sea floor show that the ore stage I minerals and dolomitic rocks were replaced by massive sulfide (ore stage II), normally towards their base. The temperature and salinity of hydrothermal fluids of this stage is 115-234 o C and 9.50% NaCl.

43

Journal Pre-proof Table captions Table. 1 Sulfur, oxygen and carbon isotopic (‰) compositions of mineral separates from the Farahabad deposit.

Jo

ur

na

lP

re

-p

ro

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Table. 2 Summary of fluid inclusion micro-thermometric data of the Farahabad deposit.

44

Journal Pre-proof

Mineral

δ34SCDT

Gn Gn Gn

Bedded ore facies

-14.6

Dol

11.7

3.9

Dol

10.4

-7.2

Dol

12.8

Dol

13.1

Dol

12.1

Sph

2.8

Sph

-3.9

Hydrothermal dolomite

Sph

-1.1

Dol

12.7

Gn

-17.25

Dol

12.4

Gn

-17.6

Lim

7.6

Gn

-16.1

Lim

9.4

Sph

3.8 -8.8

Gn

-19.47

Gn

-18.47

Gn

-18.16

Sph

-19

Sph

-18.5

Sph

-19.3

Py

-19

Lithology

Hydrothermal dolomite

-19.2

na

Py

of

2.2

Sph

Lim

ro

Sph

Host rock limestone

-p

Massivereplacement ore zone

δ18OSMOW

re

Stringer zone

Mineral

Lithology

lP

Facies

Host rock limestone

6.5 13

Mineral

δ CPDB

Dol

1.7

Dol

1.6

Dol

2.7

Dol

1.9

Dol

2.4

Dol

2.6

Dol

1.6

Lim

1.4

Lim

2.1

Lim

2

Jo

Farahabad deposit.

ur

Table. 1 Sulfur, oxygen and carbon isotopic (‰) compositions of mineral separates from the

45

Tm-hh (C) ….

wt% NaCl 0.99

wt% CaCl2 ….

wt % NaCl/(NaCl+CaCl2) ….

wt % NaCl+CaCl2 0.99

Thv-l (۫ C) 118

L+V

P

-52

-12.3

-22.3

12.9

3.83

0.77

16.7

146

9

L+V

P

….

-6.1

….

9.34

….

….

9.34

139

11

L+V

P

….

-4.8

….

7.59

….

….

7.59

159

10

L+V

P

….

-3.8

….

6.16

….

….

6.16

154

5

L+V

P

….

-4

….

6.62

….

….

6.62

189

8

L+V

P

….

-2.9

….

4.8

….

….

4.8

192

6

L+V

P

….

-5

….

8.07

….

….

8.07

216

10

L+V

P

-52

-11.3

-22

14.1

0.71

0.95

14.8

125

9

L+V

P

….

-6

….

9.41

….

….

9.41

220

13

L+V

P

….

-2

….

3.39

….

ro

4

….

3.39

196

5

L+V

P

….

-4

….

6.62

….

….

6.62

197

8

L+V

P

-52

-15.8

-22.5

15.2

3.9

0.79

19.1

234

3

L+V

P

….

-5.5

….

8.75

….

….

8.75

232

7

L+V

P

….

-6.2

….

9.47

….

….

9.47

217

16

L+V

P

….

-1.5

….

2.5

….

….

2.5

234

11

L+V

P

….

-5.2

….

8.35

….

….

8.35

160

10

L+V

P

….

-6.5

….

10.05

….

….

10.05

170

8

L+V

P

….

-11.5

….

15.47

….

15.47

186

8

L+V

P

-52

-12.2

-22

14.25

1.8

0.88

16.05

203

9

L+V

P

….

-10.5

….

14.49

….

….

14.49

174

11

L+V

P

….

-7.5

….

11.27

….

….

11.27

190

3

L+V

P

….

-6.2

….

9.67

….

….

9.67

185

8

L+V

P

….

-3.5

….

5.86

….

…..

5.86

162

5

L+V

P

….

-8.2

….

11.81

….

…..

11.81

122

7

L+V

P

….

-6.4

….

9.73

….

…..

9.73

125

9

L+V

P

….

-4.5

….

7.17

….

…..

7.17

115

6

L+V

P

….

-4.5

….

7.36

….

…..

7.36

152

5

L+V

P

….

-10.2

….

14.15

….

…..

14.15

221

-p

L+V

of

Tm-ice (۫C) -0.7

re

Origin

ur

Type

Jo

Size (µm) 12

na

P

Te (۫C) ….

lP

Journal Pre-proof

Tm-ice ( ۫ C): final melting temperature of ice; Te ( ۫ C): first melting temperature of ice; Tm-hh ( ۫ C): temperature melting of hydrohalite; Th l-v (°C): homogenization temperatures of primary fluid inclusions contains liquid and vapor (l-v) phase; L: liquid; V: vapor; P: primary.

Table. 2 Summary of fluid inclusion micro-thermometric data of the Farahabad deposit.

46

Jo

ur

na

lP

re

-p

ro

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47

Journal Pre-proof Highlights ► Farahabad is a Lower Cretaceous-age sandy dolomite, sandstone and silty limestone hosted zinc-lead deposit in the southern Yazd basin, Iran. sandy

dolomite,

sandstone

and

silty

limestone.

► Mineralization is mainly hosted in

►

Sulfur

isotopes,

together

with

sedimentological, textural, mineralogical, and geochemical evidences, suggest that this deposit

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ro

of

should be classified as a vent-proximal sub-seafloor replacement SEDEX ore deposit.

48

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Conflict of Interest:

Department of Geology, Faculty of Basic Sciences, Tarbiat Modares University, Tehran, Iran Department of Earth Sciences, Faculty of Natural Sciences, University of Tabriz, Tabriz, Iran Department of Earth Sciences, University of New Brunswick, New Brunswick, Canada

Jo

ur

na

lP

re

-p

ro

of

Chrono-Environnement, Université de Franche-Comté/CNRS, 25030 Besançon Cedex, France

49

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