Constrains from magmatic and hydrothermal epidotes on crystallization of granitic magma and sulfide mineralization in Paleoproterozoic Malanjkhand Granitoid, Central India

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Constrains from magmatic and hydrothermal epidotes on crystallization of granitic magma and sulfide mineralization in Paleoproterozoic Malanjkhand Granitoid, Central India Dinesh Pandit a,∗ , Mruganka K. Panigrahi b , Takeru Moriyama c,d a

National Centre for Antarctic & Ocean Research, Vasco-Da-Gama 403804, Goa, India Department of Geology & Geophysics, Indian Institute of Technology Kharagpur 721302, India c Institute of Geo-Resources and Environment, National Institute of Advanced Industrial Science and Technology (AIST), Tsukuba, Japan d Metal and Mineral Resources Department, Toyota Tusho Corporation, Japan b

a r t i c l e

i n f o

Article history: Received 14 November 2012 Accepted 23 April 2014 Handling Editor K. Heide Keywords: Epidotes Granitic magma Transport rate Hydrothermal fluid Oxygen fugacity Sulfide mineralization

a b s t r a c t The Paleoproterozoic Malanjkhand Granitoid (MG) pluton in the Central India was studied to constrain the depth of emplacement, upward magma transport rate, and quantification of physicochemical condition of sulfide deposition. In this study, the magmatic and hydrothermal epidotes are of two varieties, reported from a mineralized granitoid. In the MG, composition of magmatic epidotes (pistachite – Ps: 21.6–31.1 mole%) and hydrothermal epidotes (Ps: 22.6–31.1 mole%) are overlapping in terms of mole percent of pistachite component. It does not provide any significant discrimination between the two varieties. Presence of oscillatory zoning in magmatic epidotes indicates that there was cyclic change in the oxygen fugacity or bulk composition of granitic magma during crystallization. Al-in-hornblende barometry indicates that the MG crystallized under 2–5.6 kbar pressure and high oxidation state (FMQ-HM) conditions, inferred from Fe/(Fe + Mg) ratio in hornblende (0.36–0.51) over wide range of temperature (800–650 ◦ C). Partial dissolution of epidote indicates an average 6 years time that corresponds to upward transport rates 0.45 km/year for magma migration in the crust. Rapid upward magma migration in most cases was probably through dyke mechanism, which is also the most appropriate model to understand the emplacement of granitic magma responsible for the formation of Malanjkhand pluton. In the Malanjkhand ore deposit, hydrothermal epidotes associated with major sulfide phases (chalcopyrite and pyrite) suggest that they equilibrated with the mineralizing ore fluid. Hydrothermal epidotes were formed over a wide range of temperature (147–424 ◦ C). From mineral–fluid equilibria modeling it was inferred that low to moderate temperature, moderate to high fO2 (>HM buffer) and low fS2 conditions were favorable for formation of hydrothermal epidotes. Interaction between hydrothermal epidote with mineralizing ore fluid in the wall rock would raise the log(aCa2+ /a2H+ ) ratio that brings a fall in pH values, followed by potassic alteration, which promotes the deposition of sulfide ores at Malanjkhand. Sulfide mineralization in the MG represents a unique Paleoproterozoic granite ore system. © 2014 Elsevier GmbH. All rights reserved.

1. Introduction Epidote is a hydrous rock forming silicate mineral reported from various plutonic rocks such as tonalite, dacite, diorite (Schmidt and Poli, 2004), granite–granodiorite (Sial et al., 2008), gabbro (Korinevskii, 2008), volcanic rocks, calcareous sediments, and geothermal environments (Bird and Spieler, 2004). It is also reported as metamorphic mineral and extending its stability from

∗ Corresponding author. Tel.: +91 9332425587; fax: +91 8322520877. E-mail address: [email protected] (D. Pandit).

lower greenschist to amphibolite, blueschist, and eclogite facies rocks (Poli and Schmidt, 2004). Chemical composition of the epidotes depends on various factors such as bulk composition, temperature, pressure, pH, fugacities of CO2 , S2 , and O2 (Arnason et al., 1993; Schmidt and Poli, 2004; Bird and Spieler, 2004). It occurs in granitic rocks as single holocrystalline grain, euhedral crystal, where the cenotypal appearance indicates magmatic origin (Schmidt and Poli, 2004; Korinevskii, 2008). Magmatic epidotes are formed at a pressure of about 10 kbar and sometimes are stable above 3 kbar pressures at relatively low-temperature (Schmidt and Thompson, 1996). It crystallizes before biotite within a narrow range of temperature in granitic magma (Schmidt and Poli,

http://dx.doi.org/10.1016/j.chemer.2014.04.008 0009-2819/© 2014 Elsevier GmbH. All rights reserved.

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2004). Presence of magmatic epidote in granitic rock indicates that the hydrous magma is crystallized at high oxidation state in greater depth or at high pressure (Dawes and Evans, 1991), and is emplaced in the upper crust due to rapid transportation (Brandon et al., 1996). Epidotes occur in veins and cavities along with secondary mineral phases such as altered biotite, chlorite, pyrite, chalcopyrite, magnetite, and hematite indicating its hydrothermal origin (Bird et al., 1988; Bird and Spieler, 2004). A physicochemical environment with low to moderate temperature and high oxidation state is more favorable for the formation of hydrothermal epidotes. It is very sensitive to redox conditions and aqueous speciation in hydrothermal fluid. The composition of hydrothermal epidotes is influenced by a wide range of octahedral substitution Al3+ –Fe3+ in low-pressure hydrothermal systems in order to become stable (Bird and Spieler, 2004). Formation of granitic plutons through incremental addition of multiple pulses of magma from contrasting sources partially preserves the primary heterogeneities (Bea, 2010; Clemens et al., 2010; Brown, 2013). Heterogeneities in granitic magma from largescale (exposed outcrop) to micro-scale (individual mineral grains) promote investigations to constraint the petrogenesis and crystallization evolution of mineral grains (Gagnevin et al., 2008). Slow cooling rate during crystallization of granitic magma promotes continuous exsolution of H2 O and growth of equant mineral grains (Nabelek et al., 2010). The crystallization sequence of minerals in granitic magma is strongly dependent upon the upward transport rate (Brandon et al., 1996; Petford et al., 2000) and emplacement mechanism during pluton formation (Whitney, 1988; Hutton et al., 1990). The timescale of magma crystallization and ore deposition (Shinohara et al., 1995; Lowenstern et al., 2000), multiple events of the magmatic-hydrothermal fluid–rock interaction in the formation of significant large ore deposits (Ballard et al., 2001), may be shorter compared to the total time span of pluton formation (Halter et al., 2004). Granite related ore systems represent one of the most complex processes of magma crystallization and multiple episodes of hydrothermal activities. It can be also influenced by various physicochemical factors such as magma composition, oxidation state, crystallization history, ascent rate, exsolution of volatile phases, depth of emplacement, and so on (Candela, 1997). The post-magmatic multiple episodes of hydrothermal activities may easily overprint the texture and chemistry of the existing magmatic and hydrothermal mineral assemblages (Schaltegger et al., 2005). Post-emplacement tectonic stresses develop extensive fault and fracture network in granitic plutons, and allow circulation of hydrothermal fluid (Marques et al., 2010). Infiltration of hydrothermal fluid along the fracture network and micro-cracks favors dissolution–precipitation processes and the subsequent formation of hydrothermal or secondary minerals (Nishimoto and Yoshida, 2010). Typical hydrothermal epidote-group minerals also occur in granitic rocks (Bird and Spieler, 2004; Vlach, 2012). Therefore, it is difficult to understand the crystallization sequences of epidotes in granitic rocks, and they need a systematic study. Investigation of the sequence of magma crystallization reactions and proper identification of hydrothermal mineral assemblages in granitic rocks could reveal the physicochemical conditions of pluton formation. Epidote is well recognized as one of the common silicate minerals present in the granitic rocks, formed during magmatic crystallization and hydrothermal alteration processes. Phase equilibria studies of magmatic and hydrothermal mineral assemblages would provide a better opportunity to constrain the history of granite evolution and related ore mineralization (Marks et al., 2003). The present study deals with two varieties of epidotes (magmatic and hydrothermal types) found in a Paleoproterozoic mineralized granitoid, located in the Central India and their fundamental contributions to the granite related ore system. The objective is divided into three parts: (1) petrographic and mineral

chemical discrimination between magmatic and hydrothermal epidotes; (2) investigating the crystallization trends of magmatic epidotes, evaluating the depth of epidote formation and estimation of transport rate of granitic magma; and (3) constraining the physicochemical conditions of hydrothermal epidotes formation and sulfide deposition at the Malanjkhand. For this purpose, five samples from the mineralized zone (mine pit) and seven samples from outside the mineralized zone were selected from the MG. Primarily, we have attempted to constrain the P–T conditions for crystallization of magmatic epidotes during the granite emplacement and contextualize the significance of hydrothermal epidotes during sulfide mineralization in a granite related ore system.

2. Geological setting The Bastar or Bhandara craton dominantly occupies the Central India, bordered by the Pranhita–Godavari Rift in the South, the Mahanadi Rift in the northeast, the Satpura Mobile Belt in the North, the Eastern Ghats Mobile Belt in the East, and Deccan Traps cover in the West (Naqvi and Rogers, 1987). It consists of mainly Precambrian granites and basement granitic gneisses, supracrustal sequences, mafic dyke swarms and sedimentary basins (Meert et al., 2010). The Malanjkhand and Dongargarh granitoids are two prominent units of Precambrian/Paleoproterozoic granitoids (Pandit and Panigrahi, 2012). Supracrustal sequences are part of the Dongargarh supergroup (Nandgaon and Khairagarh group), Sakoli Group and Sausar Group (part of Central Indian Tectonic Zone, i.e. CITZ). Multiple episodes of Precambrian mafic magmatism formed dyke swarms, mafic volcanic rocks and mafic dykes (French et al., 2008; Srivastava and Gautam, 2009). There are two major sedimentary basins (Chhattisgarh and Indravati Basin) and other six minor basins in the Bastar Craton. The ENE–WSW striking CITZ divides the Precambrian Indian Peninsula into two parts, i.e. northern and southern crustal blocks (Radhakrishna and Naqvi, 1986; Acharyya, 2003; Naganjaneyulu and Santosh, 2010; Bhowmik et al., 2012; Mandal et al., 2013). The MG situated along the southern boundary of CITZ (Fig. 1a), hosts one of the largest copper deposit in India (Sikka, 1989; Sarkar et al., 1996; Panigrahi and Mookherjee, 1997; Bhargava and Pal, 1999; Stein et al., 2004). It is comprised of coarse-grained hornblende–biotite bearing granite–granodiorite and emplaced as a single episode of early Paleoproterozoic (ca. 2.48 Ga) felsic magmatism in the Central India exposed over ∼1400 km2 area (Panigrahi et al., 2004, 2009). Two subordinate units of leucogranite are found with sporadic occurrences at Birsa (close to copper mines) and Devgaon (∼12 km south of the mine pit) within the Malanjkhand batholith (Fig. 1b). They are considered as a separate phase of granitic activity, which provide a Rb–Sr whole rock isochron age ∼2.11 Ga (Panigrahi et al., 1993), rendering uncertain their origin and emplacement. Amphibolitic enclaves in MG are one of the common features in which modal hornblende is >60%. This suggests mixing of mafic and felsic magma in various proportions within a dynamic magma chamber resulting in the formation of granitic magma and emplacement at shallow crustal level (Kumar et al., 2004; Kumar and Rino, 2006). However, Re–Os model ages (∼2.49 to 2.44 Ga) of molybdenites obtained from the Malanjkhand deposit are interpreted as discrete deformation episodes and molybdenite formation (Stein et al., 2004, 2006) which overlap with the zircons ages (∼2.48 Ga) possibly due to the extent or episodic hydrothermal activities (Panigrahi et al., 2006, 2008, 2013). Further, microstructural studies as well as the anisotropy of magnetic susceptibility reveal that emplacement of MG and tectonic evolution of CITZ was synchronous during the Neoarchean/Paleoproterozoic era (Majumder and Mamtani, 2009a,b). Recently, Pandit and Panigrahi (2012) suggested that the origin of the Palaeoproterozoic granitoids

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Fig. 1. (a) Study area adjacent to the CITZ in the map of India; and (b) Geological map of the MG in Central India with sample locations.

in the Central India is attributed to the mixing process between crustal derived felsic and mafic magma contributed by basaltic underplating in a continental rift environment. Intrusive aplitic dykes in the host granitoid are only exposed in the mine pit whereas mafic dykes occurred throughout the Malanjkhand pluton. These two igneous bodies are later phases of magmatic activities. Younger Chilpi Group of sediments partially covers the MG. 3. Formation of ore deposit at Malanjkhand The Malanjkhand copper deposit is an arcuate shape mineralized quartz reef nearly ∼1.8 km long, with a N-S strike ∼70o E dip, and ∼70 m maximum width in the central part. The quartz reef varies from a single thin tabular body at the northern end to a composite of parallel sheeted, interconnected veins in the central and southern part. It occupies a major fracture zone in the enclosed granitoid (Panigrahi and Mookherjee, 1997). Inside the mine pit, the granitoid is characterized by extensive deuteric alteration in the form of saussuritization of plagioclase, breakdown of hornblende to biotite + epidote + titanite, chloritization of hornblende, and chloritization of biotite (Pandit, 2014). K-feldspar megacrysts show conspicuous pink color typical for this granitoid, which invariably contains inclusions of saussuritized plagioclase, and recurrently occurs in the mine pit. They are often associated with myrmekite intergrowth texture, which indicates that K-feldspar deposition was more favorable after the magmatic–hydrothermal transition (Hibbard, 1979). K-feldspar impregnation also occurred in significant quantity within the mineralized quartz reef. Chalcopyrite and pyrite are dominant ore mineral phases in the mineralized quartz reef whereas sphalerite, bornite, hematite, and magnetite occur as minor phases (Sikka et al., 1991; Panigrahi and Mookherjee, 1997; Panigrahi et al., 2008). Dominating sulfide phases (chalcopyrite and pyrite) are intimately associated with hydrothermal epidote along with other secondary silicate mineral phases (biotite, chlorite, Kfeldspar, hematite, and altered plagioclase, etc.) in the mineralized wall rock (Sikka et al., 1991; Panigrahi et al., 2008). Stein et al. (2004, 2006) argued that the Malanjkhand deposit is a subductionrelated stockwork-porphyry-style deposit based on petrography and high Re concentration in molybdenites. Panigrahi et al. (2009) have summarized in more detail the discussion on the sulfide mineralization and hydrothermal fluid activities at Malanjkhand. A

conclusive model based on the fluid inclusion and chemistry of secondary silicate minerals suggests that internally evolved late stage hydrothermal fluid caused deuteric alteration with remobilization of metals (K, Si, Fe and Cu), and deposited in a fracture zone to finally form a mineralized quartz reef. The fluid exsolved from the subordinate leucogranite was S and CO2 -rich and interacted with host granitoid to favor sulfide precipitation (Pandit et al., 2008). Sulfur isotope composition of pyrite, molybdenite, and chalcopyrite from the Malanjkhand copper deposits allows to infer that the hydrothermal fluid was derived from a single magmatic source (Panigrahi et al., 2013). The multiple sources of fluids and metals associated with the two phases of granitic activities represent a fossil hydrothermal system at the Malanjkhand (Panigrahi and Mookherjee, 1997; Panigrahi et al., 2008, 2013). 4. Petrography More than 150 fresh rock samples were collected for detailed petrography and geochemical studies (Pandit, 2008; Pandit and Panigrahi, 2012) from the Malanjkhand pluton on three occasions (7th to 24th Dec. 2004; 6th to 23rd Dec. 2005 and 4th to 10th April 2006). More than 30 polished thin sections have been prepared from the collected samples during the present investigation and more than 80 thin sections were available from the previous studies (Panigrahi, 1992; Naik, 2006). All these thin sections have been used for comprehensive petrographic studies. Twelve polished thin sections containing high modal abundance of epidotes are slides # 284, 304, 319, 19/472, and 2916 from the mine pit and slides # 56, 72, 104, 106, 714, 1-2/12, and 4-18/12 from other parts of the pluton (Fig. 1b). In the MG, two distinct mineral assemblages have been observed i.e. primary/magmatic and secondary/alteration phases. Magmatic phases comprise biotite, hornblende, epidote, plagioclase, and Kfeldspar whereas late- to post-magmatic secondary phases are associated with sulfide ore mineralization such as chlorite, epidotes, altered biotites, saussuritized plagioclase, secondary titanite, and so on. Earlier workers have discussed the petrography of quartz (Pandit, 2012) and mineral chemistry of hornblende, plagioclase, biotites, and chlorites from the MG (Sarkar et al., 1996; Kumar et al., 2004; Kumar and Rino, 2006; Panigrahi et al., 2008; Pandit and Panigrahi, 2012; Pandit, 2014). In this study, we are reporting

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Fig. 2. Photomicrograph of different types of magmatic epidote and their textural relationships in the MG under crossed polarized light: (a) epidote within biotite, (b) Epidote in unaltered hornblende, (c) Zoned epidote with biotite partially resorbed, (d) Zoned epidote, (e) Euhedral epidote with biotite and K-feldspar, and (f) Prismatic epidote with biotite.

for the first time the occurrences of magmatic and hydrothermal epidotes in a single granitoid body. Magmatic epidotes (Schmidt and Poli, 2004; Dawes and Evans, 1991) are commonly euhedral hexagonal in shape (Fig. 2a–f), showing a greenish-pink to pinkish-yellow color (Fig. 2c–e), and sometimes colorless with oblique extinction. Occurrences of independent equant subhedral epidotes with biotite, hornblende, and plagioclase are common in granitoids. It is recurrently associated with hornblende and biotite in the MG (Fig. 2b). Korinevskii

(2008) reported magmatic epidotes as individual minerals up to 1.5 mm grain size, isometric or with short-prismatic habit, displaying smoothed and rounded grains in contacts with plagioclase, amphibole, and pyroxene in gabbro from eastern shore of Lake Bol’shoi Ishkul near Cape Osinovyi, Chelyabinsk Region, Russia. In present study, two types of magmatic epidote were identified based on textural criteria. Type I epidote occurs as euhedral phenocrysts, some having allanite cores and partially rimmed by biotite. Type II epidote forms smaller euhedral to subhedral grains, which has

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Fig. 3. Photomicrographs of various types of hydrothermal epidote and their textural relationships in the MG under crossed polarized light: (a) variable colored aggregates of epidote grains, (b) aggregates of epidote grains with K-feldspar and quartz, (c) epidote in association with chalcopyrite and quartz in the mineralization zone at Malanjkhand, (d) epidote associated with euhedral pyrite and altered plagioclase, (e) clots of epidote in K-feldspar along with quartz, and (f) epidote in association with chlorites.

no core component, is totally rimed by biotite or hornblende, and is partially resorbed by the host mineral (Fig. 2a–c). Sometimes, epidotes display oscillatory zoning. The oscillatory zones in epidote grains are mostly continuous and sometimes contain allanite cores (Fig. 2c and d). These oscillatory zones are of variable thickness. One of the possibilities is that the growth of magmatic epidote around allanite core occurred at the early stage of magma crystallization. Magmatic epidotes seem to have a moderate to large grain size (200–1000 ␮m) and have survived from dissolution. Thus, it

indicates that crystallization of granitic magma was relatively a slow process. Occurrence of magmatic epidote indicates crystallization of a hydrous granitic melt at high pressures (Dawes and Evans, 1991). Secondary mineral phases such as quartz, K-feldspar, chlorite, sulfides (chalcopyrite, pyrite), and oxides minerals (hematite and magnetite) are found in association with epidotes around the mineralized zone in the MG (Fig. 3a–f). This epidote occurs in clots with irregular grain boundaries, characteristic of secondary origin and

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more akin to hydrothermal epidote (Bird and Spieler, 2004). Hydrothermal epidotes show variable strong anisotropic colors (green to pinkish-yellow) and small to moderate grain size (100–500 ␮m). These are associated with various mineral assemblages such as quartz + K-feldspar (Fig. 3a and b), chalcopyrite + pyrite (Fig. 3c and d), quartz + K-feldspar + altered plagioclase (Fig. 3d). Secondary epidotes mainly occur in the veins along with incipient fractures in the mineralized quartz reef (Fig. 3e and f). They are variable in color but easily distinguished from magmatic epidote based on the mineral association and textural criteria discussed by Schmidt and Poli (2004). Sulfide ore minerals are commonly associated with hydrothermal epidotes in the mineralized quartz reef. Based on mineral fluid equilibria modeling, Panigrahi et al. (2008) inferred that the presence of hydrothermal epidotes strongly influenced copper deposition at Malanjkhand.

5. Mineral chemistry Two varieties of epidotes and some co-existing pairs of epidotes and hornblende associated with sulfide phases in selected polished thin sections of granitoid rock samples have been analyzed. A JEOL 8900 Superprobe microanalyser at National Institute of Advanced Industrial Science and Technology, Tsukuba, Geological Survey of Japan has been used to obtain mineral chemical data of various silicate minerals. The beam current has been 12 nA, the beam diameter 4 ␮m and the acceleration voltage 15 kV with a counting time set at 20 s. All natural silicate standards have been used for routine calibrations with high accuracy and precision (<1%). Mineral standards used for calibration are as follows: wollastonite (Si, Ca), orthoclase (K), albite (Na), spinel (Al, Mg), hematite (Fe), rutile (Ti), rhodonite (Mn), topaz (F), and sodalite (Cl). The composition of magmatic epidotes showed significant variation in their mineral chemistry: SiO2 34.56–40.01 wt%, Al2 O3 16.97–25.10 wt%, Fe2 O3 10.60–16.12 wt%, CaO 19.27–23.79 wt% and other elements such as TiO2 and MnO were mostly less than 1 wt% but in few cases slightly higher than this (Appendix Table A1). The magmatic epidote from gabbroic rock (Korinevskii, 2008) has very restricted composition but within the range of the data obtained in this study. Overall, there was a slightly restricted variation in the composition of hydrothermal epidotes from the MG: SiO2 36.32–39.65 wt%, Al2 O3 20.90–23.51 wt%, Fe2 O3 11.94–15.96 wt%, CaO 20.72–23.77 wt% and other elements such as TiO2 and MnO mostly <1 wt% (Appendix Table A2). There is a range of variation in the chemical composition of magmatic and hydrothermal epidotes but they do not show any systematic trend in their major element oxide contents. A wider range in composition of magmatic epidote was observed compared to hydrothermal variety. However, CaO content has limited range in both the cases and does not show much distinction between magmatic and hydrothermal varieties. In the present study, SiO2 content was higher than the values obtained from earlier reported values in hydrothermal epidotes (Panigrahi et al., 2008) and magmatic epidotes (Sial et al., 1999, 2008). However, Al2 O3 , CaO, and Fe2 O3 content in epidotes from the MG were within the range comparable to other magmatic epidotes (Sial et al., 1999, 2008; Korinevskii, 2008) and hydrothermal epidotes (Panigrahi et al., 2008). Holdaway (1972) expresses the epidote composition in terms of mole fraction of pistachite [XPs = molar {Fe3+ /(Fe3+ + Al)}]. Overall, the compositional variation recorded as end member component of epidote (XPs ) range from 0.101 to 0.311 mole fraction. Electron microprobe data indicate that the XPs in magmatic epidote range is 0.216–0.311 and for hydrothermal epidotes, the range is 0.226–0.301. In the MG, however, there were exceptions, as a few samples showed low values (XPs = 0.11). The range of compositional variation in magmatic and hydrothermal epidotes is overlapping

Fig. 4. Histograms of mole fractions of pistachite (XPs ) in magmatic and hydrothermal epidotes from the MG. The compositional ranges of epidote in association with alteration of plagioclase and biotite considered from Tulloch (1979).

(Fig. 4). It also overlaps with the epidote compositions reported by Sial et al. (1999, 2008), and Panigrahi et al. (2008), which does not corroborate with the arguments put forth by Tulloch (1979). Therefore, there is no clear discrimination observed between the two varieties of epidotes (magmatic and hydrothermal) based on their mineral chemistry obtained in this study. Morrison (2004) suggested that the stable isotope as well as trace elements data of epidote possibly provide reliable information on their magmatic origin and may be helpful in discriminating from secondary epidote. 5.1. Magmatic epidotes and crystallization of granitic magma Keyes (1893), for the first time reported that epidote occurs as a magmatic phase in granitic rocks. Laboratory experiments by Naney (1983) also demonstrated that epidote could be stable above the granite and granodiorite solidus. However, identification of magmatic epidote in granitic rock is a matter of debate. According to Cornelius (1915), field observations and microscopic textures are the only possibility to recognize magmatic epidotes, and this argument is further supported by Schmidt and Poli (2004 and references their in). A summary of the textural and compositional criteria to distinguish magmatic and hydrothermal epidotes as suggested by Zen and Hammarstrom (1984), Zen (1985), Tulloch (1979), Evans and Vance (1987), Dawes and Evans (1991), Schmidt and Poli (2004) and Sial et al. (2008) is given below: (i) commonly, magmatic epidotes are euhedral to subhedral in shape with sharp and clear grain boundary such as hexagonal or prismatic; (ii) mineralogical associations with primary hornblende and biotite; sometimes along with apatite, titanite and zircons as accessories phases; (iii) oscillatory zoning in epidote, allanite-rich core, embayed contacts with plagioclase and quartz; (iv) chemical composition in terms of pistachite [Ps = XPs × 100] content in epidote > 25%; and (v) magmatic epidote <0.2 wt% TiO2 content.

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Fig. 5. (a) Photomicrograph of zoned magmatic epidote with allanite core associated with biotites and titanite as inclusions, and (b) compositional (Ps) variation along X–Y profile within the epidote grain.

The criteria used for identification of magmatic epidotes based on chemical compositions (Tulloch, 1979; Evans and Vance, 1987) are not useful in this study. Based on the textural criteria, in the present study magmatic epidotes are distinguished from the hydrothermal epidotes as discussed above. The chemical variation in epidotes is observed in terms of optical oscillatory zoning (Figs. 5a and 6a), which is supposed to be magmatic origin, in which the composition of epidotes in terms of Fe2 O3 ranges from 10.3 to 13.9 wt%. A typical magmatic epidote contains allanite core (Type-I, sample no 56) with repetitive oscillatory rims around it (Fig. 5a). This epidote grain co-exists with biotite along with titanite inclusions. A line profile (X–Y) obtained across the grain not considering the allanite core, shows low iron content (Ps) at the grain boundary whereas higher values are observed in the internal part. The compositional variation recorded in terms of Ps component of epidote from core to rim is in the range of 23.7–30.2 mole%. The profile of Ps variation in mole percent against distance (D in micrometer) plotted shows oscillatory variation, and successive decreases in Ps values as away from the core (Fig. 5b). From the decrease in Ps-content from core toward rim it is inferred that there is significant depletion of Fe3+ in the granitic melt during the crystallization of magma. Although some epidotes have nucleated on allanite cores, indicating that the epidote formation started when at least one-third of the magma was already crystallized in a deep-seated magma chamber, survived from dissolution in the melt, and to be transported upwards (Dawes and Evans, 1991). The back-scattered electron (BSE) image of magmatic epidote without allanite core (Type-II) shows characteristic oscillatory zoning and inclusions of accessory minerals like apatite, titanite, and zircon (Fig. 6a). In this case, Ps content in epidote is low in the inner and outer parts but higher in the intermediate rims with a rhythmic variation. Inclusions of accessory minerals indicate that the formation of epidote is a later event, at least after crystallization of zircon, apatite, and titanite. The compositional profile (A–B) across this epidote grain (Fig. 6b) shows complex rhythmic variation in Ps content compared to another magmatic epidote with allanite core (Fig. 6b). Oscillatory zoning in magmatic epidote indicates that successive rims over grew one after another during crystallization either due to cyclic change in the composition of granitic magma or oxygen fugacity or physicochemical conditions. In both the cases (TypeI and Type-II epidotes), lower Ps-content in outer rims compared to inner one suggests that the granitic magma was depleted with respect to Fe3+ in the late stage of epidote crystallization. From the

cyclic pattern in the compositional profiles of the epidote grains it is inferred that there was no complete miscibility between allanite and epidote growth during crystallization of magma. Type-I epidote is compositionally discontinuous; it formed as a later phase during crystallization and nucleated on the pre-existing allanite core. However, Type-II epidote was formed without core because the granitic magma was depleted in REE content with further progress in crystallization (Dawes and Evans, 1991). 5.2. Hornblende barometry In the MG, hornblende occurs as megacrysts (up to 4 mm in size) and shows variable degrees of break down reactions. In thin sections, hornblende grains range from euhedral to anhedral in shape, pleochroic from bright bluish green to yellow, and often show twinning. The textural evidences from the MG suggest that hornblende equilibrated with granitic melt in the presence of plagioclase and biotite. Occurrences of both hornblende and biotite together are possibly due to discontinuous incongruent hydration crystallization of hydrous silicic melt (Beard and Lofgren, 1991; Rushmer, 1991; Beard et al., 2004). Hornblende megacrysts wherever present retain sharp grain boundary and sometimes co-exists with magmatic epidotes (Fig. 2b). The chemical compositions of some co-existing pairs of hornblendes and magmatic epidotes have been analyzed using Electron Probe Micro Analyzer (EPMA) and are presented in Appendix Table A3. Hornblende composition varies in a wide range: SiO2 42.74–49.48 wt%, Al2 O3 5.99–10.29 wt%, FeOtot 13.15–18.74 wt%, MgO 6.68–16.48 wt%, and other oxides (TiO2 , Na2 O and K2 O) also present in trace amount (<1 wt%). In one spot, exceptionally high value of TiO2 (∼9.72 wt%, sample no 714 spot 46) was recorded, possibly because of a titanite inclusion within the epidote influenced the composition. Notable amounts of F and Cl were also recorded in hornblende, with a maximum concentration of 1.09 wt% and 0.21 wt%, respectively. In the present study, a ternary (Ca + Na)B –(Fe2+ + Fe3+ )–NaB discrimination diagram (Fig. 7) in accordance with the International Mineralogical Association (IMA, Leake et al., 1997) and proposed by Yavuz (1999) was used to classify most of the hornblendes as calcic amphiboles (Field A). The Fe/(Fe + Mg) ratios in hornblende (0.36–0.51) were reasonably high inferring that granitic magma crystallized under high fO2 conditions (Speer and Becker, 1992; Oliveira et al., 2010). The Al-in-hornblende geobarometer (Schmidt, 1992) was used for the estimation of pressure at which the granitic magma crystallized. In this study, the calculated crystallization pressure ranged

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Fig. 6. (a) Backscattered image of zoned magmatic epidote without core associated with titanite, apatite and zircon as inclusions, and (b) compositional (Ps) variation in zoned epidote along A–B profile.

between 2.04 and 5.63 kbar for the MG. Pandit and Panigrahi (2012) estimated P–T conditions for the formation of MG in the range of 2.16–3.88 kbar and 743–773 ◦ C, respectively, using hornblendeplagioclase thermobarometry. From this it can be inferred that the MG was formed at relatively higher pressure as compared to the previous study. Hence, it is assumed that the limiting P–T condition for epidote formation must be <5.63 kbar and <800 ◦ C, respectively, whereas the lower limit is bound by the wet granite solidus curve as shown in Fig. 8. According to Naney (1983), the magmatic epidote is stable in the presence of 4 wt% water along with hornblende + plagioclase + biotite + melt at high pressure (∼8 kbar). However, with decreasing temperature, hornblende and plagioclase disappear and the assemblage changes to plagioclase + epidote + biotite + melt. Therefore, these magmatic epidotes were formed when the MG experienced high pressure. In this granitoid, magmatic epidote crystallized from the granitic melt as a primary phase in H2 O-saturated system around ∼800 ◦ C temperature at 5.63 kbar pressure and plausibly appeared until

Fig. 7. Plot of recalculated analyses of hornblendes from the MG on ternary (Ca + Na)B –(Fe2+ + Fe3+ )–NaB group classification diagram (after Yavuz, 1999). A = calcic amphibole field; B = sodic–calcic amphibole field; C = sodic amphibole field; D = Fe–Mg–Mn–Li amphibole field.

the magma cooled down to a temperature ∼650 ◦ C at 2.04 kbar pressure (Fig. 8). Presence of magmatic epidote in the granitoid indicates that the parental magma originated from a minimum depth of ∼18.75 km (assuming lithostatic pressure 5.63 kbar and crustal density 2.7 g/cm3 ) or possibly from even deeper in the crust (Zen and Hammarstrom, 1984; Zen, 1985).

Fig. 8. P–T diagram showing stability of magmatic epidotes in the MG; solidi for wet granite and granodiorite with excess H2 O as well as appearance of biotite and hornblende for granodioritic composition (after Liou, 1993).

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5.3. Magma transport rate Magma transport rate can be estimated using the dissolution kinetics of magmatic epidotes in granitic rocks (Brandon et al., 1996). Epidote is stable in granitic magmas, which crystallized at shallow crustal levels and derived from the deep crust through dyke mechanism (Sial et al., 1999, 2008; Brasilino et al., 2011). The variable size of epidote grains indicates a long duration of continued crystallization of granitic magma. Resorption and dissolution of crystallized epidote is possible due to change in the physicochemical conditions during transportation of granitic magma in the crust. A qualitative estimation of dissolution and transportation rate can be possible by measuring the amount of resorption in magmatic epidotes. Generally, crystal nucleation and crystal growth rate increases with crystallization of magma but decreases in the presence of H2 O rich vapor phase (Swanson, 1977). In other words, low crystal nucleation and high growth rate favored crystallization of coarse grains in magma. Hogan et al. (2000) suggest that the crystallization of coarse grain granitic rock occurs over a wide range of temperature in the presence of high H2 O contents in magma. In long-term crystallization history of magma, a broad range of crystal sizes can be formed as long as nucleation continues (Marsh, 1988). Grain size and dissolution width of magmatic epidotes have been measured in thin sections of MG using a Leica DM4500 petrological microscope with an attached Leica QWin image analysis software. The grain size of magmatic epidotes shows a wide range from 100 to 650 ␮m and the corresponding dissolution width varies between 59 and 178 ␮m (Appendix Table A4). An average time of ∼6.3 years is estimated for partial dissolution of magmatic epidotes using the formulation of Brandon et al. (1996), which is in the range 2.2–20.1 years (Supplementary Appendix S1). Schmidt and Thompson (1996) experimentally proved that epidote is stable at >10 kbar pressure in tonalitic magmas which corresponds to a depth ∼33.3 km in the crust. Pandit and Panigrahi (2012) suggested that the MG emplaced in a tectonic environment with pressure ranges of 2.16–3.88 kbar and temperature ranges of 743–773 ◦ C based on thermobarometric calculations, which is consistent with the results obtained in this study. An average pressure of ∼3.5 kbar is obtained after considering both the results. Therefore, it is assumed that the magmatic epidote was formed at ∼10 kbar pressure in the MG. It survived from dissolution during transportation in the upper crust. The route length is ∼21.7 km, which is the difference between the depth of crystallization (∼10 kbar) and depth of emplacement (∼3.5 kbar). The estimated time of partial dissolution range is 2.2–20.1 years, which corresponds to a magma ascent rate 0.95–9.72 km/year (average ∼0.45 km/year) in this study. Assumptions and limitations of epidote dissolution kinetic discussed in Sial et al. (2008), have also been considered for this study. Overall, the method provides a good qualitative approach of granite magma transport in the continental crust. Hence, it can be inferred that the upward ascent rate of granitic magma was rapid enough for survival of magmatic epidotes against complete dissolution and only possible through a dyke/conduit flow mechanism in a predominantly extensional tectonic setting (Petford et al., 2000). 5.4. Hydrothermal epidotes and sulfide deposition Intermediate compositions within the epidote solid solution appear to be stable at relatively low temperature under conditions of low fO2 (Keith et al., 1968). Most commonly, hydrothermal epidotes reported from low temperature (200–250 ◦ C) active geothermal systems are associated with calc-silicate rocks (Cho et al., 1988; Bird et al., 1988; Caruso et al., 1988; Liou, 1993). However, secondary epidotes in the hydrothermal system can be easily recognized based on the mineralogical assemblages and other

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characteristic features for identification are also available in literatures (Bird and Spieler, 2004 and reference therein). A summary of important characteristics features of hydrothermal epidotes are discussed below: (i) hydrothermal epidotes are subhedral to anhedral in shape, hazy with rough grain boundary and occur as clots or aggregates of grains, associated with low temperature mineral phases such as quartz, K-feldspar, altered plagioclase, chlorites, altered biotites, and in few cases with sulfide minerals such as chalcopyrite, pyrites, bornite, and so on, secondary and more Fe-rich epidotes occur with altered mineral assemblages of igneous origin such as biotite and hornblende (Dawes and Evans, 1991), pistachite (Ps) content of epidote should be <25% or >35%, and (ii) most of the hydrothermal epidotes have TiO2 >0.6 wt%. The stabilities of hydrothermal epidotes depend on various factors. Therefore, the composition of hydrothermal epidotes associated with sulfide minerals is an important tool for characterizing the environment of ore mineralization (Bird and Helgeson, 1981; Bird et al., 1984, 1988; Panigrahi et al., 2008). 5.5. Temperature of sulfide mineralization (i) Sulfide mineralization at Malanjkhand occurs in the form of a quartz reef and mainly contains chalcopyrite–pyrite phases. Jairath and Sharma (1986), Ramnathan et al. (1990), and Panigrahi and Mookherjee (1997) have carried out detailed fluid inclusion microthermometry studies. According to Jairath and Sharma (1986), a CO2 -bearing hydrothermal fluid, a temperature range of 210–470 ◦ C with an adequate quantity of Na and K in the fluid promotes sulfide deposition. The assemblages of pyrite–magnetite–chalcopyrite–hematite are the main product of sulfide mineralization and predict log(fO2 ) in the range of −30.2 to −34.4, log(fS2 ) in the range of −8.8 to −11.6, pH in the range of 4.5–6.5 and pressure in the range of 0.225–0.44 kbar. Panigrahi and Mookherjee (1997) advocate a two-stage evolution of ore fluid under a pressure range of 0.55–0.179 kbar. In the first stage, sulfide deposition was due to mixing of fluids derived from two different sources i.e., F1 and F2 with higher proportion where, F1 is high temperature (∼375 ◦ C), low saline fluid (4–8 wt% NaCl equivalent) with appreciable amounts of S + CO2 and (ii) F2 is low temperature (180–200 ◦ C), moderately saline (20–24 wt% NaCl equivalent) metal-rich fluid. In the second stage, the CO2 bearing fluid phase separates because of an increase in the salinity of aqueous fluids. Panigrahi et al. (2008) conclusively suggested a more appropriate evolution path of the mineralizing fluid with the revised paragenetic sequence of ore and silicate minerals. The mineral chemistry associated with the ore body indicates low oxidation state (FMQ-NNO) in which biotites get re-equilibrated with the ore fluid at ∼300 ◦ C and chlorite at ∼200 ◦ C, but the authors do not provide any detailed account on hydrothermal epidote. It is further discussed that the interaction of ore fluid with epidote in the wall rock favored the deposition of chalcopyrite and enhanced Ca2+ ions activity because of a decrease in pH, and subsequent acid alteration of K-feldspar in the mineralized zone. On the other hand, Bhargava and Pal (1999) suggested that the mineralized reef and veins/stringers in the granitoids represent two different stages of quartz deposition. Stein et al. (2004) advocate multiple episodes of fracturing and molybdenite deposition at the Malanjkhand. The textural relationship of secondary silicate and ore minerals indicates that the hydrothermal epidotes were formed in the presence of fluid phase during sulfide mineralization at

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Fig. 9. Isobaric (at 1 kbar) temperature versus oxygen fugacity plot for hydrothermal epidote in the MG. Unit activities used unless specified otherwise. The stability constraint for hydrothermal epidotes during the mineralization compared with oxides assemblages is shown as solid lines, sulfide assemblages as broken lines and shaded area represents the conditions in which hydrothermal epidotes equilibrated with ore fluids.

Malanjkhand (Panigrahi et al., 2008). Compositions of these epidotes are useful in constraining the temperature of their formation in the hydrothermal environment. Temperature of epidote formation can be calculated using epidote solid solution model (Bird and Helgeson, 1980), considering Al–Fe substitution in the octahedral site. The mineral composition of hydrothermal epidotes obtained in this study provides an estimated temperature range of 267–424 ◦ C which is indicative of a slightly higher temperature range (147–333 ◦ C) that was reported in the previous study (Panigrahi et al., 2008). Therefore, it can be inferred that the hydrothermal epidotes were stable in a wide temperature range of 147–424 ◦ C during hydrothermal mineralization. The fugacity of oxygen and sulfur changes with changes in the temperature of hydrothermal fluids during mineralization generated a wide compositional variation in hydrothermal epidotes.

Fig. 10. The relationship between temperature and pH of ore fluid in the Malanjkhand granite-ore-system at 1 kbar for various secondary mineral phases associated with epidotes during mineralization corresponding to aSiO2 = aH2 O = 1 and other metal ions activities i.e. log(aM i+ ) = −6. Dotted lines indicate the equilibrium conditions of various secondary mineral phases. The crosshatched area indicates physic-chemical condition of Cu-mineralization at Malanjkhand (Panigrahi et al., 2008) and rationalized with an active geothermal system in Iceland represented by the shaded area (S&A: Stefansson and Arnorsson, 2002).

Appendix) was invariably close to the hematite–magnetite (HM) buffer and is shown with a gray field. The occurrence of hematite in epidote-bearing assemblages restricts the field to the hematitestable side of HM buffer curve. Fig. 9 shows log(fO2 ) versus T diagram constructed over the range of temperature obtained from hydrothermal epidote, representing phase boundaries and stability of ore minerals. In the hematite reaction, the range of calculated log(fO2 ) is −21.64 to −46.83 (Appendix Table A5), considering log(fS2 ) values of −15 to −5 based on previous estimation (Panigrahi et al., 2008). The values of log(fO2 ) have also been obtained in the presence of pyrite considering reaction (8) for same log(fS2 ) values, which are not reliable for the stability of hydrothermal epidotes at Malanjkhand. 5.7. pH-condition

5.6. Oxidation state The stability of hydrothermal epidote is depicted in the isobaric diagram (Fig. 9, temperature versus oxygen fugacity), considering a number of standard oxidation and sulfidation reactions of ore and silicate minerals (Supplementary Appendix S2). The stability of the epidote-bearing assemblage in terms of log(fO2 ) and T at constant P was investigated, considering various phases such as epidote (Ep), clinozoisite (Czo), tenorite (Tn), cuprite (Cpr), hematite (Hem), magnetite (Mag), chalcopyrite (Ccp), pyrite (Py) and native copper (Cu) in association with fluid constituents such as SiO2(aq) , H2 O, H2 and O2 . The activity of hydrothermal epidote was calculated based on the formulations of Bird and Helgeson (1980). Since, pressure has a negligible effect on the equilibrium constant (log K), the activities of SiO2(aq) and H2 O were taken as unity in the temperature range obtained from the chemistry of hydrothermal epidotes at 1 kbar pressure, which brings out very small changes in the position of the reactions in the log(fO2 )–T field. The oxygen fugacity was calculated by determining the activity-corrected log K values of oxidation and sulfidation reactions using the DSLOP thermodynamic database and SUPCRT92-routine (Johnson et al., 1992) at 1 kbar with unit activities of H2 O and SiO2 . The stability fields of sulfide-bearing and sulfide free epidote-assemblages were constructed at 1 kbar (Fig. 9). The reaction (5: Supplementary

Simple hydrolysis reactions involving the epidote, clinozoisite, hematite, magnetite, end members of plagioclase (albite and anorthite), low temperature K-feldspar (microcline) and kaolinite were considered at isobaric (1 kbar) conditions in order to plot pH diagram at unit activity of H2 O; and other metal ion activity (Mi+ ) was assumed to be log(aM i+ ) = −6. The temperature obtained from the chemistry of hydrothermal epidote was chosen for construction of this stability diagram. Consequently, much attention has been focused on determining the pH condition of the ore fluids. Presence of a K-feldspar grain with chalcopyrite (Fig. 3c) indicates that the acid alteration of wall-rock minerals does not bring in a significant fall in pH values of the ore fluids during sulfide deposition (Jairath and Sharma, 1986; Panigrahi et al., 2008). The major phases associated with hydrothermal epidotes are K-feldspar, hematite, pyrite, chalcopyrite (Fig. 3a–f), biotite, and chlorite (Panigrahi et al., 2008) in the mineralized zone at Malanjkhand. Assuming, that the ore fluid was in equilibrium with the hydrothermal epidote–plagioclase–K-feldspar–hematite, which then was altered to kaolinite (saussuritization?) has been represented by various hydrolysis reactions (Supplementary Appendix S3). The equilibrium constant of those reactions was calculated using DSLOP thermodynamic database with the help of SUPCRT92 program (Johnson et al., 1992) and the result is shown in Fig. 10. However,

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epidote + hematite and epidote + pyrite are two important assemblages that intimately influence the chemistry of ore fluids. The formation of an epidote + pyrite assemblage (reaction R8 in Supplementary Appendix S5) depends on the sulfur activity (fS2 ) whereas the epidote + hematite assemblage (reaction R11 in Supplementary Appendix S5) strongly depends on the activity of hydrogen ions (aH+ ). The hydrolysis of epidotes is directly related to the acidity of the ore fluids that influenced the sulfide mineralization at Malanjkhand. In the present study, an epidote + hematite assemblage is more suitable to establish the relationship between the chemistry of hydrothermal epidotes and pH of ore fluids. The hydrolysis of epidote increases the activity of Ca2+ ions in the fluid that results in a fall in pH. According to Panigrahi et al. (2008), the biotite + chalcopyrite assemblage re-equilibrated in a pH range 4.7–6.7 that corresponds to log(aCa2+ ) values of −6.44 to −5.32. The pH condition of the ore fluid slightly increases with the decrease in temperature in the presence of epidote but strongly depends on the log(aCa2+ ) activity. A more favorable pH condition of 2.7–5.4 has been estimated using the equilibrium reaction (R11) for the stability of hydrothermal epidote assuming log(aCa2+ ) values of −7.0 to −4.0 (Appendix Table A6). This indicates that a lower pH and higher Ca2+ activity in the ore fluid possibly influenced the precipitation of sulfide minerals. The assessment of the equilibrium state of secondary minerals with hydrothermal fluid is fraught with considerable uncertainties due to aqueous speciation in the system.

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Fig. 11. Logarithmic activity–temperature diagram for epidote–clinozoisite– chalcopyrite assemblage at 1 kbar corresponding to aSiO2 = aH2 O = 1. The stability relation involving epidote with a composition of aEp = 0.33 and rationalized with the situation of active Iceland Geothermal System (S&A: Stefansson and Arnorsson, 2002). The broken line indicates various buffer limits using different combination of log(fO2 ) and log(fS2 ); shaded area represents the most suitable conditions in which hydrothermal epidote equilibrated with ore fluids.

6. Discussion

5.8. Cation activities The association of hydrothermal epidotes with chalcopyrite in the mineralized zone at Malanjkhand significantly influences the chemistry of the ore fluid. In the Malanjkhand fossil hydrothermal system, K-feldspar was the earliest mineral to form followed by hydrothermal epidote similar to that in active Salton Sea geothermal system (Caruso et al., 1988). In an active geothermal system such as in Iceland (Arnorsson, 1983; Stefansson and Arnorsson, 2002) and Salton Sea (Cho et al., 1988; Caruso et al., 1988; Bird et al., 1988; Charles et al., 1988), the redox chemistry and Ca2+ ions activities in the fluid is strongly dependant on the fluid temperature. According to Stefansson and Arnorsson (2002), the interaction of ore fluids with the host rock is considered as a simplified process like titration in which the primary minerals are treated as bases and the fluid behaves as acid, and the reaction progresses with the deposition of ore and the formation of secondary minerals. The formation of secondary minerals increases the log(aCa2+ /a2 + ) ratio in the fluid. In the present study, the H epidote–clinozoisite–chalcopyrite assemblage has been taken into account to examine the influence of fluid temperature during Cu mineralization in terms of log(aCa2+ /a2 + ) ratio. The theoretH

ical stability relations (Fig. 11) in terms of log(aCa2+ /a2 + ) ratio H has been calculated at 1 kbar using the epidote–clinozoisite equilibrium reaction (R14) as a function of log(aCu+ /aH+ ), log(fO2 ) and log(fS2 ) whereas aH2 O = 1 and aEp = 0.33. The log(aCu+ /aH+ ) values varies from −6 to −2 during the deposition of ore minerals as inferred from the abridged activity diagram suggested by Panigrahi et al. (2008, Fig. 8). Therefore, a minimum value for log(aCa2+ /aH+ ) = −6 has been chosen in the calculation of log(aCa2+ /a2 + ) ratio (Appendix Table A7). The stability diagram H (Fig. 10) illustrates that the hydrothermal epidote was equilibrated with ore fluids over a wide range of temperature due to simultaneous changes in fO2 and fS2 conditions in the Malanjkhand ore system. The observed textural relationship and mineral–fluid equilibria modeling that involves hydrothermal epidote indicates that temperature, log(fO2 ) and log(fS2 ) decrease with chalcopyrite deposition.

The MG which is emplaced at a shallow level in the crust (Kumar and Rino, 2006; Pandit and Panigrahi, 2012), has undergone extensive deuteric alteration during the late stage hydrothermal activities (Panigrahi et al., 2008, 2013) and displays fascinating records of two generations of epidotes. The magmatic epidote was formed in the later stage of crystallization whereas the hydrothermal epidote was the product of deuteric alterations during sulfide precipitation. A comparative study of the two varieties of epidotes from MG has been carried out to discriminate between magmatic and hydrothermal types based on their origin. Petrographically, the two types of epidotes can be easily distinguished based on mineral association, texture, shape, and color. Magmatic epidotes are associated with primary minerals (hornblende and biotite), euhedral in shape, sharp grain boundary, high order interference colors, and oscillatory zoning with or without allanite core. On the other hand, hydrothermal epidote is associated with secondary minerals (chlorite, altered biotite, altered plagioclase, low temperature K-feldspar, chalcopyrite, pyrite, hematite, etc.), anhedral in shape, hazy grain boundaries, variegated colors, and occurs in aggregates of grains in various size. The major element mineral chemistry does not reveal any significant differences between the magmatic and hydrothermal epidotes in this study. Magmatic epidotes show a range of 10.1–31.1 mole% whereas hydrothermal epidotes have a range of 22.6–30.1 mole% in Ps values. The proposed criteria for discrimination between the two varieties of epidotes reported in many literatures (Tulloch, 1979; Evans and Vance, 1987, Dawes and Evans, 1991; Schmidt and Poli, 2004; Sial et al., 1999, 2008) could be possible based on major element oxides, mineral chemistry and pistachite compositions. However, these propositions have not been found suitable in this study. Composition of epidotes is strongly dependant on the oxidation state, and can be estimated in terms of Fe-content wherein high Fe3+ indicates oxidizing conditions during crystallization and vice versa (Holdaway, 1972; Liou, 1973; Raith, 1976). The presence of magmatic epidote in the MG indicates a higher level of oxidation state at the time of crystallization of magma. However, the hydrothermal epidote has been formed

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Fig. 12. log(fO2 )–pH diagrams are showing the stability relations among minerals in the system Cu–Fe–S–O. Equilibrium constant for the reaction at 400, 300, and 200 ◦ C calculated from DSLOP thermodynamic database using SUPCRT92 computer program (Johnson et al., 1992). Shaded area represents the stability of hydrothermal epidote and cross-hatched area indicates sulfide deposition at Malanjkhand (see Panigrahi et al., 2008). Mineral abbreviations are Hem – hematite; Py – pyrite; Bn – bornite; Ccp – chalcopyrite; Mag – magnetite; Po – pyrrhotite.

at low temperature conditions, as far as reported above 150 ◦ C with a minor changes in fO2 and fS2 conditions of the hydrothermal system. Photomicrographs of magmatic epidote (Fig. 2a–f) show that grains were partially resorbed, corroded, and shielded by host minerals (biotite, hornblende, and K-feldspar). Unusual partially resorbed magmatic epidote grains indicate that they were unstable in the granitic magma and survived complete dissolution possibly due to a rapid magma transport rate. In other words, surrounding minerals crystallized faster than epidote dissolution. The physicochemical condition of the granitic magma changes with progressive crystallization within the epidote stability field. Magmatic epidote is directly crystallized from granitic melt by nucleation and some grow on a pre-existing allanite grain. It can be inferred that the presence of zoning in magmatic epidotes and the survival from complete dissolution not only explains the rapid upward magma transport but also reflects the changes in the physicochemical conditions during the crystallization of pluton. Transport of granitic magma at different crustal level has been modeled by diapirism (Ramburg, 1972; Mahon et al., 1988) as well as by fracture propagation through a dike mechanism (Hutton et al., 1990; Clemens, 1998; Cruden, 1998; Petford et al., 2000). Most of the granitic magmas emplaced in the middle to the upper crust (Cruden, 1998) were transported through a dyke mechanism, which is a rapid and thermally efficient way to form tabular batholiths (Clemens, 1998). Sulfide mineralization at Malanjkhand has been accounted by various workers (Sikka et al., 1991; Panigrahi and Mookherjee, 1997; Bhargava and Pal, 1999; Panigrahi et al., 2008, 2009, 2013), but the presence of epidote in association with sulfide and

secondary silicate phases has not been discussed. Composition of hydrothermal epidotes from the mineralized zone is quite similar to that of magmatic epidote obtained from other parts of the Malanjkhand pluton. A temperature range of 147–424 ◦ C is more reliable for the stability of hydrothermal epidote at Malanjkhand. Various important physicochemical factors that have acted to maximize ore deposition have been examined over the temperature range in which hydrothermal epidote has been in equilibrium with the ore fluid. The quantification of such parameters has been carried out using mineral–fluid equilibria modeling. The values log(fO2 ) ranges −21.64 to −46.83 and range of pH values from 2.7 to 5.4 have been calculated for sulfide deposition during late stage hydrothermal fluid activities. Theoretically, the log(aCa2+ /a2 + ) ratio decreases H with a decrease in fluid temperature and strongly depends on log(fO2 ) and log(fS2 ) in the presence of the epidote. According to Stefansson and Arnorsson (2002), it is possible that a low saline fluid has a higher pH and lower Ca2+ activity or vice versa, and would have similar log(aCa2+ /a2 + ) ratios at the same temperature. H Therefore, it is concluded that a low to moderate temperature saline fluid interacted with the hydrothermal epidote present in the host rock and released adequate quantities of Ca2+ ions into the fluid, which increased the values of log(aCa2+ /a2 + ) ratio and H promoted mineralization. Chlorite is one of the characteristic secondary mineral formed in the Paleoproterozoic granite ore system at the Malanjkhand because of extensive hydrothermal alteration primary biotite and hornblende (Pandit, 2014). Thus, a sequence of stability diagrams (Fig. 12a–c) is presented to describe the role of epidote, biotite, and chlorite in the sulfide mineralization at Malanjkhand.

Please cite this article in press as: Pandit, D., et al., Constrains from magmatic and hydrothermal epidotes on crystallization of granitic magma and sulfide mineralization in Paleoproterozoic Malanjkhand Granitoid, Central India. Chemie Erde - Geochemistry (2014), http://dx.doi.org/10.1016/j.chemer.2014.04.008

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7. Conclusions

Acknowledgements

A comparative study of magmatic and hydrothermal epidotes revealed that the two varieties can be discriminated based on petrography, texture, and associated mineral assemblages. However, the major element composition does not provide significant clues for their discrimination. Occurrences of both epidotes in the MG reflect a complex sequence of physicochemical processes such as crystallization, resorption, and dissolution during magmatic evolution; followed by replacement and subsequent alteration during hydrothermal mineralization. Ore deposition in a granite related hydrothermal ore system at Malanjkhand can simply be explained by considering three silicate minerals (epidote, biotite and chlorite), which have equilibrated with the mineralizing fluid (Panigrahi et al., 2008). The ore fluid interacted with the hydrothermal epidote at ∼400 ◦ C temperature and above the HM buffer (Fig. 12a). As a result, activity of Ca2+ ions increased and pH decreased in the hydrothermal fluid. The fluid cooled down to ∼300 ◦ C temperature at which biotite re-equilibrated (Fig. 12b) and a further fall in oxidation state (FMQ to NNO buffer) triggered the deposition of chalcopyrite. Finally, chlorite registered equilibrium with the ore fluid at ∼200 ◦ C temperature (Fig. 12c) and accelerated ore deposition with a rise in pH.

DP acknowledges the Council of Scientific and Industrial Research, (CSIR), Government of India for the financial support in the form of research fellowship. The research work was also supported in terms of project sponsored by the Department of Science of Technology (DST), Government of India (Project No. ESS/16/246/2005) to MKP. Sincere acknowledgment to Dr. Rajesh K. Naik, Geological Survey of India, Kolkata for his consistent support in research at IIT Kharagpur and Dr. Achuthankutty C. T., Eminent Scientist (NCAOR) for improving the grammar. The National Institute of Advanced Industrial Science and Technology (AIST), Geological Survey of Japan is also acknowledged for facilitating Electron Probe Micro Analyzer. The authors express their sincere thanks to Dr. Ursula Kelm for editorial support and two anonymous referees for their constructive comments to improve the scientific quality on our previous version of the manuscript.

Appendix A. Appendix See Tables A1–A7.

Table A1 Representative EPMA analyses of magmatic epidote from the Malanjkhand granitoid and Fe content reported as total FeO. Analysis in wt% and number of cations calculated on the basis of 12 oxygens. Slide No Point#

106 29

106 30

106 31

106 32

106 33

106 34

106 35

106 36

106 37

106 38

106 39

106 40

106 41

106 42

106 43

SiO2 TiO2 Al2 O3 Fe2 O3 MnO CaO

39.00 0.06 22.77 13.53 – 23.52

38.19 0.11 22.91 13.95 – 23.06

39.37 0.09 23.05 13.76 – 23.42

39.13 0.08 23.02 13.89 – 23.54

38.85 0.12 22.25 13.70 – 23.14

38.30 0.06 22.02 14.41 – 22.76

36.95 0.11 21.58 15.21 – 22.63

37.54 0.22 21.90 15.58 – 22.86

38.33 0.09 22.42 14.47 – 23.46

39.51 0.16 22.55 14.25 – 23.58

37.98 0.23 22.13 15.00 – 23.65

38.08 0.35 22.81 13.76 – 23.36

38.54 0.11 22.63 13.97 – 23.10

38.68 0.28 22.47 13.95 – 23.61

37.57 0.22 22.73 14.00 – 23.07

Total

97.63

96.84

98.40

98.29

96.73

96.40

95.48

96.61

97.36

98.65

97.58

97.56

97.05

97.63

96.23

Number of cations calculated on the basis of 12 oxygens Si 3.096 3.036 3.080 3.069 2.131 2.147 2.125 2.128 Al Fe 0.740 0.827 0.792 0.796 Mn – – – – Ca 2.001 1.964 1.963 1.978 Sum 7.968 7.974 7.960 7.971 0.258 0.278 0.271 0.272 XPs

3.113 2.101 0.757 – 1.987 7.958 0.265

3.069 2.080 0.852 – 1.954 7.955 0.291

3.013 2.074 0.925 – 1.977 7.989 0.308

3.012 2.071 0.933 – 1.965 7.981 0.311

3.055 2.106 0.818 – 2.003 7.982 0.280

3.108 2.091 0.774 – 1.987 7.960 0.270

3.063 2.103 0.780 – 2.044 7.990 0.271

3.070 2.167 0.727 – 2.018 7.982 0.251

3.059 2.117 0.827 – 1.965 7.968 0.281

3.113 2.131 0.693 – 2.036 7.973 0.245

3.022 2.155 0.824 – 1.988 7.989 0.277

Slide No Point#

106 44

106 45

106 46

106 47

106 48

106 49

106 50

106 51

106 52

106 53

106 54

106 55

106 56

106 57

106 58

SiO2 TiO2 Al2 O3 Fe2 O3 MnO CaO

38.10 0.04 22.95 14.17 – 23.11

39.70 0.06 22.94 13.97 – 23.15

38.29 0.03 23.04 13.45 – 23.09

38.78 0.00 23.54 13.22 – 23.15

38.33 0.07 23.98 12.80 – 23.71

40.01 0.01 23.87 12.60 – 23.47

39.56 0.04 23.67 12.70 – 23.40

39.86 0.05 23.78 13.06 – 23.20

39.18 0.04 24.04 12.63 – 23.40

39.17 0.02 23.45 13.02 – 23.39

37.91 0.05 23.51 13.15 – 23.53

38.36 0.28 21.53 15.45 – 23.07

38.88 0.20 21.53 15.62 – 23.18

39.08 0.07 22.52 14.92 – 23.36

38.70 0.11 22.46 14.72 – 23.70

Total

97.07

98.47

96.73

97.42

98.20

98.81

98.15

98.67

98.16

97.81

96.93

97.21

97.94

98.48

98.27

Number of cations calculated on the basis of 12 oxygens 3.026 3.094 3.046 3.053 Si 2.148 2.107 2.160 2.184 Al 0.840 0.812 0.798 0.776 Fe – – – – Mn 1.967 1.933 1.968 1.953 Ca 7.981 7.946 7.972 7.966 Sum XPs 0.281 0.278 0.270 0.262

3.030 2.234 0.720 – 2.008 7.992 0.244

3.097 2.177 0.727 – 1.946 7.947 0.250

3.086 2.176 0.739 – 1.956 7.957 0.254

3.090 2.173 0.756 – 1.927 7.946 0.258

3.059 2.212 0.736 – 1.958 7.965 0.250

3.071 2.167 0.762 – 1.965 7.965 0.260

3.020 2.207 0.761 – 2.008 7.996 0.256

3.078 2.036 0.869 – 1.983 7.966 0.299

3.087 2.014 0.888 – 1.972 7.961 0.306

3.058 2.077 0.871 – 1.959 7.965 0.295

3.066 2.097 0.808 – 2.012 7.983 0.278

Slide No Point#

106 59

106 60

106 61

106 62

106 63

106 64

106 65

106 66

106 67

106 68

106 69

106 70

106 71

106 72

106 73

SiO2 TiO2 Slide No

39.31 0.39 106

39.35 0.21 106

39.43 0.30 106

39.86 0.33 106

38.06 0.30 106

39.47 0.11 106

39.53 0.08 106

38.98 0.02 106

35.09 0.08 106

39.93 0.07 106

39.28 0.02 106

39.28 0.07 106

37.55 1.49 106

39.48 0.04 106

37.80 0.04 106

Please cite this article in press as: Pandit, D., et al., Constrains from magmatic and hydrothermal epidotes on crystallization of granitic magma and sulfide mineralization in Paleoproterozoic Malanjkhand Granitoid, Central India. Chemie Erde - Geochemistry (2014), http://dx.doi.org/10.1016/j.chemer.2014.04.008

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14 Table A1 ( Continued ). Point#

59

60

61

62

63

64

65

66

67

68

69

70

71

72

73

Al2 O3 Fe2 O3 MnO CaO

22.80 13.45 – 23.69

23.15 13.69 – 23.44

22.92 13.32 – 23.64

22.10 15.43 – 23.57

21.74 15.21 – 23.53

22.28 15.17 – 23.35

23.07 13.99 – 23.26

22.87 14.15 – 23.43

24.10 12.48 – 23.48

25.10 10.92 – 23.61

22.66 13.86 – 22.86

23.39 13.50 – 23.45

21.80 15.64 – 22.55

23.64 13.36 – 23.66

23.54 13.39 – 23.52

Total

98.37

98.54

98.33

99.77

97.38

99.04

98.70

98.10

94.19

98.59

97.33

98.42

97.53

98.88

97.00

Number of cations calculated on the basis of 12 oxygens Si 3.149 3.086 3.138 3.112 2.153 2.140 2.150 2.034 Al Fe 0.625 0.766 0.654 0.835 – – – – Mn 2.034 1.970 2.016 1.972 Ca 7.961 7.962 7.958 7.953 Sum XPs 0.225 0.264 0.233 0.291

3.085 2.077 0.781 – 2.043 7.986 0.273

3.077 2.047 0.882 – 1.951 7.957 0.301

3.080 2.119 0.813 – 1.942 7.954 0.277

3.058 2.115 0.828 – 1.969 7.970 0.281

2.912 2.357 0.699 – 2.088 8.056 0.229

3.086 2.286 0.629 – 1.955 7.956 0.216

3.095 2.104 0.815 – 1.930 7.944 0.279

3.065 2.151 0.786 – 1.961 7.963 0.268

3.106 2.125 0.736 – 1.998 7.965 0.257

3.064 2.162 0.774 – 1.967 7.967 0.264

3.002 2.203 0.793 – 2.001 7.999 0.265

Slide No Point#

106 75

106 76

106 77

72 61

72 62

72 63

72 64

4 18/12 24

4 18/12 25

4 18/12 26

714 38

714 39

714 40

104 41

104 42

SiO2 TiO2 Al2 O3 Fe2 O3 MnO CaO

39.88 0.16 21.92 14.96 – 22.86

39.24 1.81 21.69 13.75 – 23.51

39.83 0.05 24.59 11.52 – 23.79

37.93 0.58 20.34 15.35 – 22.91

37.89 0.38 19.86 16.12 – 23.32

37.76 0.45 20.16 15.47 – 23.35

37.29 0.24 21.23 14.86 – 23.26

37.36 0.05 22.05 13.54 – 22.19

38.50 0.12 21.31 15.08 – 22.89

38.42 0.10 22.05 15.07 – 22.66

37.48 0.02 22.12 13.85 – 23.24

37.47 0.06 21.91 14.89 – 23.18

38.54 0.05 22.04 14.67 – 23.09

39.10 0.04 23.34 13.39 – 23.48

37.89 0.07 23.23 14.51 – 23.33

Total

98.34

98.88

98.64

95.69

95.99

95.68

95.54

93.94

96.48

96.89

95.51

96.07

97.01

98.34

97.78

Number of cations calculated on the basis of 12 oxygens Si 3.120 3.351 3.097 3.179 2.021 2.183 2.253 2.009 Al 0.873 0.246 0.630 0.714 Fe – – – – Mn 1.916 2.151 1.982 2.057 Ca 7.930 7.931 7.962 7.959 Sum 0.302 0.101 0.219 0.262 XPs

3.177 1.963 0.735 – 2.095 7.970 0.272

3.189 2.007 0.665 – 2.113 7.974 0.249

3.094 2.076 0.751 – 2.068 7.989 0.266

3.057 2.127 0.826 – 1.946 7.956 0.280

3.095 2.019 0.873 – 1.971 7.958 0.302

Slide No Point#

01 2/12 124

56 22

56 23

56 24

56 19

56 20

56 21

56 22

56 23

56 24

56 25

56 27

56 28

56 29

56 30

SiO2 TiO2 Al2 O3 Fe2 O3 MnO CaO

37.66 0.01 24.68 10.60 0.21 22.89

38.49 0.06 23.40 13.40 – 23.30

38.17 0.09 23.25 13.82 – 23.45

38.65 0.08 23.73 12.78 – 23.65

35.42 0.06 21.14 12.89 0.1 21.89

36.41 0.02 21.95 13.34 0.11 22.98

36.23 0.08 22.27 13.75 0.13 22.46

34.56 0.21 19.67 11.60 0.24 19.27

37.10 0.06 21.85 13.60 0.02 21.94

35.27 0.08 22.70 12.03 0.12 20.49

36.45 0.00 22.20 13.93 0.01 22.24

37.04 0.05 22.52 14.57 0.2 22.89

36.80 0.00 22.21 13.86 0.15 23.08

37.05 0.04 22.13 13.97 0.22 22.56

37.22 0.03 22.26 13.97 0.19 23.05

Total

95.02

97.36

97.47

97.79

90.82

93.69

93.68

84.68

93.40

90.16

93.68

95.85

94.80

94.64

95.46

Number of cations calculated on the basis of 12 oxygens 3.027 3.039 3.021 3.057 Si 2.338 2.177 2.168 2.212 Al 0.636 0.789 0.810 0.709 Fe 0.014 – – – Mn 1.971 1.971 1.988 2.004 Ca 7.986 7.976 7.987 7.982 Sum 0.214 0.266 0.272 0.243 XPs

3.085 2.170 0.676 0.007 2.042 7.980 0.237

3.062 2.175 0.687 0.008 2.071 8.003 0.240

2.991 2.167 0.847 0.009 1.986 8.000 0.281

3.134 2.102 0.785 0.018 1.873 7.912 0.272

3.058 2.123 0.836 0.001 1.938 7.956 0.283

3.059 2.069 0.895 – 1.933 7.956 0.302

3.012 2.285 0.767 0.009 1.875 7.948 0.251

3.063 2.130 0.761 – 2.035 7.989 0.263

3.029 2.087 0.869 – 2.008 7.993 0.294

3.005 2.157 0.857 0.001 1.964 7.984 0.284

3.064 2.065 0.871 – 1.967 7.967 0.297

2.989 2.142 0.878 0.014 1.979 8.002 0.291

3.028 2.154 0.776 0.01 2.035 8.003 0.265

3.063 2.155 0.779 – 1.971 7.968 0.266

3.022 2.127 0.850 0.015 1.972 7.986 0.286

2.996 2.165 0.856 – 1.976 7.993 0.283

3.035 2.139 0.794 0.013 2.014 7.995 0.271

Table A2 Representative EPMA analyses of hydrothermal epidotes from Malanjkhand granitoid. Fe content reported as total FeO and all analysis presented in wt%. Note: XPs –mole fraction of pistachite component in epidotes (Fe-rich end member) and temperature estimated using epidote solid solution model (Bird and Helgeson, 1980). Slide No Point#

284 129

319 88

19/472 100

19/472 101

19/472 102

2916 13

2916 14

2916 15

2916 16

2916 17

2916 18

304 99

SiO2 TiO2 Al2 O3 Fe2 O3 MnO CaO

38.46 0.26 22.95 14.87 – 22.53

37.37 0.08 23.51 12.05 0.23 23.09

36.84 0.00 20.98 15.48 0.05 23.14

36.34 0.08 22.51 14.22 0.01 22.80

37.04 0.00 22.31 14.96 0.06 23.77

36.32 0.02 22.95 13.25 – 22.25

37.87 0.08 22.70 13.45 – 22.00

38.32 0.06 22.97 12.77 – 22.43

37.77 0.04 22.95 13.08 – 22.45

37.50 0.11 22.51 13.22 – 22.13

37.85 0.07 22.44 13.35 – 21.97

38.20 0.11 22.60 14.34 – 23.14

Total

97.65

95.17

94.98

94.62

96.68

93.57

95.09

95.31

95.08

94.21

94.41

97.01

Number of cations calculated on the basis of 12 oxygens 3.037 3.046 3.041 Si 2.136 2.258 2.041 Al 0.876 0.659 0.868 Fe

2.973 2.170 0.867

3.000 2.130 0.825

2.988 2.225 0.813

3.064 2.164 0.812

3.079 2.175 0.765

3.049 2.184 0.787

3.057 2.162 0.804

3.075 2.148 0.809

3.038 2.118 0.850

Please cite this article in press as: Pandit, D., et al., Constrains from magmatic and hydrothermal epidotes on crystallization of granitic magma and sulfide mineralization in Paleoproterozoic Malanjkhand Granitoid, Central India. Chemie Erde - Geochemistry (2014), http://dx.doi.org/10.1016/j.chemer.2014.04.008

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15

Table A2 ( Continued ). Slide No Point#

284 129

319 88

19/472 100

19/472 101

19/472 102

2916 13

2916 14

2916 15

2916 16

2916 17

2916 18

304 99

Mn Ca Sum XPs T (◦ C)

– 1.906 7.955 0.291 419

0.016 2.016 7.995 0.226 274

0.003 2.047 8.000 0.298 424

0.001 1.998 8.009 0.286 414

0.004 2.063 8.022 0.279 387

– 1.961 7.987 0.268 362

– 1.907 7.947 0.273 360

– 1.931 7.950 0.260 316

– 1.942 7.962 0.265 337

– 1.933 7.956 0.271 353

– 1.912 7.944 0.274 358

– 1.972 7.978 0.286 396

304 111

304 112

Slide No Point#

304 100

304 116

304 118

304 119

304 129

304 135

2916 19

2916 20

2916 21

2916 22

SiO2 TiO2 Al2 O3 Fe2 O3 MnO CaO

38.90 0.05 22.68 13.43 – 23.27

37.96 0.01 22.36 13.63 – 23.28

38.38 0.07 22.28 13.49 – 23.10

39.04 0.19 20.90 15.25 – 22.08

38.47 0.26 20.90 15.96 – 23.27

38.80 0.03 20.95 14.17 – 23.03

39.10 0.08 21.12 15.51 – 23.28

39.65 0.08 22.52 14.39 – 23.37

37.33 0.01 22.38 13.13 – 21.95

37.21 0.11 22.02 13.90 – 20.72

37.96 0.06 22.83 13.69 – 22.37

36.60 0.06 22.02 14.06 – 22.49

Total

97.27

95.93

96.15

97.27

97.33

95.61

97.57

98.61

93.53

93.75

95.64

93.88

Number of cations calculated on the basis of 12 oxygens 3.094 3.074 3.099 3.161 Si 2.126 2.134 2.120 1.994 Al 0.758 0.753 0.748 0.801 Fe Mn – – – – 1.983 2.020 1.999 1.915 Ca 7.961 7.981 7.966 7.871 Sum 0.263 0.261 0.261 0.287 XPs 304 310 300 350 T (◦ C)

3.112 1.993 0.844 – 2.017 7.966 0.297 390

3.176 2.021 0.731 – 2.020 7.948 0.265 282

3.122 1.987 0.855 – 1.992 7.956 0.301 400

3.093 2.070 0.837 – 1.953 7.953 0.288 384

3.060 2.162 0.803 – 1.928 7.953 0.271 352

3.051 2.128 0.850 – 1.820 7.849 0.285 395

3.050 2.162 0.821 – 1.926 7.959 0.275 368

3.010 2.134 0.863 – 1.982 7.989 0.288 407

Slide No Point#

2916 27

2916 28

2916 30

2916 31

2916 33

2916 34

2916 23

2916 24

2916 25

2916 26

2916 29

2916 32

SiO2 TiO2 Al2 O3 Fe2 O3 MnO CaO

38.00 0.06 22.65 13.39 – 22.77

38.43 0.07 22.78 13.92 – 22.86

38.30 0.00 23.33 11.94 – 22.71

38.25 0.07 22.12 13.19 – 22.64

37.39 0.05 21.87 14.07 – 22.82

38.24 0.08 22.89 13.65 – 22.78

38.03 0.13 22.49 13.30 – 22.76

38.55 0.15 22.62 13.36 – 22.86

38.73 0.03 22.60 13.20 – 22.21

38.34 0.03 22.67 13.48 – 22.74

38.68 0.19 22.92 13.38 – 22.82

38.09 0.14 22.80 13.53 – 22.91

Total

95.57

96.76

95.10

94.97

95.24

96.34

95.47

96.42

95.54

95.95

96.78

96.17

3.061 2.110 0.808 – 2.001 7.980 0.277 357

3.051 2.153 0.812 – 1.948 7.964 0.274 361

3.071 2.141 0.783 – 1.970 7.965 0.268 333

3.087 2.135 0.772 – 1.961 7.955 0.266 323

3.103 2.134 0.789 – 1.907 7.933 0.270 339

3.068 2.138 0.804 – 1.949 7.959 0.273 353

3.072 2.146 0.793 – 1.942 7.953 0.270 342

3.048 2.150 0.808 – 1.964 7.970 0.273 356

Number of cations calculated on the basis of 12 oxygens 3.056 3.056 3.078 3.100 Si 2.147 2.135 2.210 2.113 Al Fe 0.803 0.826 0.716 0.777 – – – – Mn 1.962 1.948 1.955 1.966 Ca Sum 7.968 7.965 7.959 7.956 0.272 0.279 0.245 0.269 XPs 352 373 267 328 T (◦ C)

Table A3 Representative electron probe analysis of amphibole co-existing with magmatic epidotes from Malanjkhand granitoid. Fe content as total FeO and all analysis presented in wt% and. Pressure estimated using the Al-in-hornblende barometry (Schmidt, 1992). Slide No Point #

714 11

714 12

714 13

714 14

714 21

714 22

714 23

714 31

714 32

714 33

714 34

714 35

714 36

SiO2 TiO2 Al2 O3 FeOtot MgO CaO Na2 O K2 O F Cl

48.12 0.73 6.87 16.32 12.54 11.97 0.89 0.84 0.37 0.07

47.18 0.63 7.44 15.72 12.95 10.75 0.79 1.63 0.57 0.07

47.24 1.32 7.49 15.60 12.32 11.87 0.94 0.93 0.19 0.11

46.62 0.84 7.75 15.50 11.78 11.90 1.03 0.86 0.42 0.09

46.53 0.22 8.26 17.64 11.34 11.41 1.01 0.79 0.37 0.11

48.93 0.32 6.14 15.62 13.33 12.03 0.92 0.43 0.00 0.06

47.27 0.46 7.91 17.34 11.74 11.86 0.99 0.82 0.29 0.12

46.30 1.12 8.09 17.29 11.56 11.88 1.14 1.06 0.42 0.21

47.57 1.26 6.91 14.42 12.20 11.96 0.84 0.83 0.10 0.12

45.62 0.84 8.73 17.97 10.91 11.83 1.12 1.03 0.34 0.14

46.00 0.27 9.40 18.74 10.77 11.70 0.99 0.98 0.32 0.14

47.26 0.61 7.09 16.98 11.86 11.94 0.79 0.83 0.48 0.20

47.48 0.60 7.38 17.07 11.90 11.98 0.73 0.80 – 0.13

Total

98.54

97.47

97.91

96.59

97.50

97.76

98.65

98.83

96.15

98.34

99.14

97.80

98.04

Number of cations calculated on the basis of 23 oxygens 7.080 7.040 6.990 7.010 Si AlIV 0.920 0.960 1.010 0.990 VI 0.270 0.340 0.300 0.380 Al 0.080 0.070 0.150 0.090 Ti 0.070 – – – Fe3+ 1.940 1.960 1.930 1.950 Fe2+

6.960 1.040 0.410 0.020 0.130 2.070

7.180 0.820 0.240 0.040 0.170 1.750

6.980 1.020 0.350 0.050 0.130 2.010

6.870 1.130 0.280 0.120 0.070 2.080

7.120 0.880 0.330 0.140 0.030 1.770

6.810 1.190 0.340 0.090 0.140 2.100

6.800 1.200 0.430 0.030 0.240 2.070

7.040 0.960 0.280 0.070 0.160 1.950

7.020 0.980 0.300 0.070 0.180 1.930

Please cite this article in press as: Pandit, D., et al., Constrains from magmatic and hydrothermal epidotes on crystallization of granitic magma and sulfide mineralization in Paleoproterozoic Malanjkhand Granitoid, Central India. Chemie Erde - Geochemistry (2014), http://dx.doi.org/10.1016/j.chemer.2014.04.008

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16 Table A3 ( Continued ). Slide No Point #

714 52

2814 11

2814 12

2814 14

2814 41

2814 42

2814 43

2814 44

2814 45

106 31

106 32

106 33

106 34

Mg Ca Na K Sum P (kbar) Fe/(Fe + Mg)

2.750 1.890 0.250 0.160 15.410 2.66 0.42

2.880 1.720 0.230 0.310 15.510 3.21 0.40

2.720 1.880 0.270 0.180 15.430 3.21 0.42

2.640 1.920 0.300 0.160 15.440 3.52 0.42

2.530 1.830 0.290 0.150 15.440 3.92 0.47

2.920 1.890 0.260 0.080 15.340 2.04 0.40

2.580 1.880 0.280 0.150 15.440 3.54 0.45

2.560 1.890 0.330 0.200 15.530 3.72 0.46

2.720 1.920 0.240 0.160 15.320 2.79 0.40

2.430 1.890 0.320 0.200 15.520 4.30 0.48

2.370 1.850 0.280 0.180 15.470 4.78 0.49

2.630 1.900 0.230 0.160 15.390 2.91 0.45

2.620 1.900 0.210 0.150 15.360 3.11 0.45

Slide No Point #

714 37

714 38

714 39

714 310

714 311

714 312

714 41

714 42

714 43

714 44

714 45

714 46

714 51

SiO2 TiO2 Al2 O3 FeOtot MgO CaO Na2 O K2 O F Cl

46.65 0.74 8.09 17.20 11.55 11.90 0.84 0.94 – 0.19

47.20 0.97 7.53 17.03 11.82 11.77 0.70 0.79 0.12 0.12

47.31 0.52 7.44 16.60 11.99 11.98 0.89 0.79 0.35 0.11

46.60 0.94 7.85 16.80 11.76 11.97 0.76 0.88 0.43 0.17

47.54 1.10 7.11 16.26 12.50 11.97 0.74 0.89 0.19 0.16

48.57 0.21 6.98 16.69 12.73 11.97 0.68 0.66 – 0.07

46.37 1.16 8.29 16.50 11.89 11.66 0.91 0.98 0.21 0.15

46.53 0.25 8.32 16.98 11.63 11.94 0.67 0.86 – 0.11

46.82 1.11 8.12 17.26 11.48 11.79 0.76 1.03 0.07 0.21

49.48 0.37 5.99 13.37 13.35 12.07 0.61 0.53 0.41 0.07

47.61 0.62 7.16 16.83 12.11 11.53 0.69 0.76 0.28 0.10

42.74 9.72 6.07 13.15 8.93 16.48 0.53 0.48 0.62 0.08

43.37 1.07 10.29 15.55 11.96 6.68 0.38 4.77 1.00 0.06

Total

98.06

97.97

97.81

97.92

98.35

98.54

97.99

97.27

98.57

96.07

97.53

98.51

94.68

7.010 0.990 0.240 0.120 0.130 1.870 2.750 1.890 0.210 0.170 15.380 2.87 0.42

7.090 0.910 0.290 0.020 0.250 1.790 2.770 1.870 0.190 0.120 15.320 2.71 0.42

6.890 1.110 0.340 0.130 0.070 1.980 2.630 1.850 0.260 0.190 15.450 3.90 0.44

6.920 1.080 0.380 0.030 0.290 1.820 2.580 1.900 0.190 0.160 15.360 3.93 0.45

6.930 1.070 0.340 0.120 0.070 2.060 2.530 1.870 0.220 0.190 15.410 3.73 0.46

7.320 0.680 0.360 0.040 0.140 1.520 2.940 1.910 0.170 0.100 15.190 – 0.36

7.070 0.930 0.320 0.070 0.120 1.970 2.680 1.830 0.200 0.140 15.340 2.96 0.44

6.400 1.070 0.000 1.090 – 1.650 1.990 2.000 0.150 0.090 15.090 2.09 0.45

6.740 1.260 0.630 0.130 – 2.020 2.770 1.110 0.110 0.950 15.720 – 0.42

Number of cations calculated on the basis of 23 oxygens 6.920 6.990 7.020 6.920 Si 1.080 1.010 0.980 1.080 AlIV VI Al 0.340 0.310 0.320 0.300 0.080 0.110 0.060 0.110 Ti 0.160 0.140 0.130 0.180 Fe3+ 1.980 1.970 1.930 1.900 Fe2+ 2.550 2.610 2.650 2.600 Mg 1.890 1.870 1.910 1.910 Ca 0.240 0.200 0.260 0.220 Na K 0.180 0.150 0.150 0.170 Sum 15.420 15.350 15.410 15.390 3.72 3.25 3.19 3.53 P (kbar) Fe/(Fe + Mg) 0.46 0.45 0.44 0.44 Slide No Point #

714 52

2814 11

2814 12

2814 14

2814 41

2814 42

2814 43

2814 44

2814 45

106 31

106 32

106 33

106 34

SiO2 TiO2 Al2 O3 FeO MgO CaO Na2 O K2 O F Cl

46.59 0.88 7.59 16.08 11.78 11.95 0.77 0.91 0.58 0.17

43.35 0.44 10.00 18.19 9.77 11.38 1.03 1.32 0.62 0.13

44.78 0.64 8.89 18.08 10.16 11.19 0.68 1.07 0.73 0.10

46.02 1.95 7.64 16.74 10.23 12.25 0.85 0.96 0.71 0.08

45.92 0.89 7.54 17.28 10.95 11.93 0.63 0.81 0.11 0.07

44.89 0.43 8.44 17.93 10.61 11.55 0.94 0.94 0.50 0.09

45.31 0.41 8.64 18.52 10.99 11.67 0.83 1.04 0.46 0.10

43.58 1.80 8.49 17.61 10.23 12.33 0.92 0.98 0.60 0.10

44.95 0.45 8.50 17.05 10.43 11.63 0.74 1.04 0.47 0.11

47.72 0.65 7.83 16.80 12.00 12.05 1.09 0.92 1.09 0.10

47.68 0.55 7.99 16.84 11.77 12.06 1.05 0.85 – 0.09

48.67 0.35 7.52 16.44 12.25 12.12 1.05 0.69 – 0.08

48.17 0.60 7.73 16.82 11.81 11.85 1.11 0.85 0.22 0.15

Total

97.02

95.93

95.99

97.12

96.07

96.08

97.74

96.36

95.14

99.77

98.85

99.14

99.17

6.960 1.040 0.300 0.100 0.200 1.990 2.470 1.940 0.190 0.160 15.340 3.4 0.47

6.850 1.150 0.360 0.050 0.230 2.060 2.410 1.890 0.280 0.180 15.460 4.21 0.49

6.790 1.210 0.310 0.050 0.370 1.950 2.450 1.870 0.240 0.200 15.440 4.25 0.49

6.670 1.330 0.200 0.210 0.200 2.050 2.330 2.000 0.270 0.190 15.490 4.28 0.49

6.910 1.090 0.450 0.050 0.110 2.080 2.390 1.920 0.220 0.200 15.420 4.32 0.48

7.000 1.000 0.350 0.070 0.020 2.040 2.620 1.890 0.310 0.170 15.480 3.43 0.44

7.010 0.990 0.390 0.060 0.020 2.050 2.580 1.900 0.300 0.160 15.460 3.58 0.45

7.100 0.900 0.400 0.040 – 2.010 2.660 1.890 0.300 0.130 15.430 3.15 0.43

7.060 0.940 0.400 0.070 – 2.060 2.580 1.860 0.320 0.160 15.440 3.35 0.44

Number of cations calculated on the basis of 23 oxygens 6.990 6.670 6.850 6.950 Si 1.010 1.330 1.150 1.050 AlIV 0.330 0.490 0.460 0.310 AlVI 0.100 0.050 0.070 0.220 Ti 0.090 0.170 0.130 0.000 Fe3+ 1.930 2.170 2.190 2.110 Fe2+ 2.630 2.240 2.320 2.300 Mg 1.920 1.880 1.840 1.980 Ca 0.220 0.310 0.200 0.250 Na 0.170 0.260 0.210 0.180 K Sum 15.400 15.570 15.410 15.360 3.38 5.63 4.62 3.46 P (kbar) 0.43 0.51 0.50 0.48 Fe/(Fe + Mg)

Please cite this article in press as: Pandit, D., et al., Constrains from magmatic and hydrothermal epidotes on crystallization of granitic magma and sulfide mineralization in Paleoproterozoic Malanjkhand Granitoid, Central India. Chemie Erde - Geochemistry (2014), http://dx.doi.org/10.1016/j.chemer.2014.04.008

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Table A4 Calculated rates of granitic magma transport for Malanjkhand granitoid using the epidote dissolution kinetics (Brandon et al., 1996). Apparent diffusion coefficient = 5 × 10−17 m2 /s; average pressure of epidote formation = 3.5 kbar; maximum emplacement pressure = 10 kbar; and lithostatic pressure variation is 3 kbar at a depth of 10 km. Sample No. (Spot No.)

Grain width (␮m)

Maximum width of dissolution (␮m)

Partial dissolution time (years)

Transport rate (m/year)

Sample No. (Spot No.)

Grain width (␮m)

Maximum width of dissolution (␮m)

Partial dissolution time (years)

Transport rate (m/year)

10-520(2) 10-520(3)

239.4 194.6 171.0 185.8 267.9 181.7 139.1 327.3 207.0 343.3 102.9 225.9 282.0 136.6 192.0 239.9 246.8 323.6 141.3 576.3 237.3 316.6 649.1 283.3 296.1 145.6 370.0 216.4 263.2 285.0 244.4 214.5 334.8

91.6 77.3 80.9 71.5 68.1 63.4 64.4 145.1 64.6 96.9 62.3 92.1 76.1 79.5 117.9 78.6 100.2 116.7 71.5 93.7 62.9 83.3 178.1 87.2 109.7 86.2 113.6 120.0 85.7 107.3 59.8 64.6 106.1

5.3 3.8 4.2 3.2 2.9 2.5 2.6 13.4 2.6 6.0 2.5 5.4 3.7 4.0 8.8 3.9 6.4 8.6 3.2 5.6 2.5 4.4 20.1 4.8 7.6 4.7 8.2 9.1 4.7 7.3 2.3 2.6 7.1

4072 5718 5220 6683 7367 8499 8238 1623 8187 3638 8802 4028 5899 5405 2458 5530 3403 2509 6683 3891 8635 4924 1077 4493 2839 1832 1055 946 1854 1183 3808 3263 1210

319(1) 319(2) 319(4)

310.6 508.4 201.4 253.2 243.1 182.7 203.0 211.0 224.4 271.1 367.1 158.0 313.6 263.1 462.3 402.6 553.5 263.6 247.4 398.8 215.1 452.7 279.7 544.2 623.5 280.8 435.4 445.3 280.5 227.1 544.2 623.5 168.7

139.7 64.7 59.3 75.9 89.4 60.9 71.5 81.9 65.6 109.9 137.7 62.3 91.7 60.7 100.0 174.5 99.5 75.6 132.1 116.8 103.7 81.1 70.4 112.8 118.1 69.5 76.8 149.9 133.7 125.1 143.2 164.8 60.6

12.4 2.7 2.2 3.7 5.1 2.4 3.2 4.3 2.7 7.7 12.0 2.5 5.3 2.3 6.3 19.3 6.3 3.6 11.1 8.7 6.8 4.2 3.1 8.1 8.8 3.1 3.7 14.3 11.3 9.9 13.0 17.2 2.3

1751 8161 9715 5930 4275 9212 6683 5093 7939 2829 1802 8802 4063 9272 3416 1122 3451 5978 1958 2504 3177 5194 6893 2685 2449 7073 5792 1520 1911 2183 1666 1258 9303

714(1)

714(2)

714(3) 714(4) 714(5) 714(6) 284(1)

284(2) 284(4) 284(5) 284(6) 284(7) 284(8) 006-16-12(1)

006-16-12(2) 006-16-12(3) 006-16-12(6) 006-16-12(7)

104(1) 104(2)

104(3) 104(4) 104(5) 304(1) 304(2) 304(3) 304(4) 304(5) 304(6)

304(7) 304(8) 304(9) 304(10) 304(11) 304(12) 304(13) 304(14)

Table A5 Range in log(fO2 ) at different temperature obtained from the reaction involving epidote–clinozoisite–hematite and assuming aEp = 0.33 (see appendix for detail). T (◦ C)

log(fS2 ) log K5

−15 log(fO2 )

−10 log(fO2 )

−5 log(fO2 )

150 200 250 300 350 400 450

293.2 254.0 222.8 197.4 176.2 158.0 142.0

−43.5 −37.0 −31.8 −27.5 −24.0 −21.0 −18.3

−45.2 −38.6 −33.4 −29.2 −25.7 −22.6 −20.0

−46.8 −40.3 −35.1 −30.9 −27.3 −24.3 −21.6

Table A6 Range in pH at different temperature obtained from the reaction involving epidote–clinozoisite–hematite and assuming aEp = 0.33. T (◦ C)

log(aCa2+ ) log K11

150 200 250 300 350 400 450

17.6 15.3 13.3 11.7 10.3 9.2 8.2

−7 pH 5.4 5.1 4.9 4.7 4.5 4.4 4.3

−6 pH 4.9 4.6 4.4 4.2 4.0 3.9 3.8

−5 pH 4.4 4.1 3.9 3.7 3.5 3.4 3.3

−4 pH 3.9 3.6 3.4 3.2 3.0 2.9 2.8

Please cite this article in press as: Pandit, D., et al., Constrains from magmatic and hydrothermal epidotes on crystallization of granitic magma and sulfide mineralization in Paleoproterozoic Malanjkhand Granitoid, Central India. Chemie Erde - Geochemistry (2014), http://dx.doi.org/10.1016/j.chemer.2014.04.008

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Table A7 Range in log(aCa2+ /a2 + ) at different temperature obtained from the reaction involving epidote–clinozoisite–chalcopyrite, assuming aEp = 0.33 and log(aCa2+ /aH+ ) = −6. H

T (◦ C)

log(fS2 ) log(fO2 ) log K14

−12 −35 log(aCa2+ /a2 + )

−10 −35 log(aCa2+ /a2 + )

−12 −36 log(aCa2+ /a2 + )

−10 −36 log(aCa2+ /a2 + )

150 200 250 300 350 400 450

−58.9 −53.9 −50.1 −47.2 −44.8 −42.9 −41.2

−4.5 11.5 13.4 14.9 16.0 17.0 17.8

−1.5 1.0 2.9 4.4 5.5 6.5 7.3

−3.0 −0.5 1.4 2.9 4.0 5.0 5.8

13.5 16.0 17.9 19.4 20.5 21.5 22.3

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Please cite this article in press as: Pandit, D., et al., Constrains from magmatic and hydrothermal epidotes on crystallization of granitic magma and sulfide mineralization in Paleoproterozoic Malanjkhand Granitoid, Central India. Chemie Erde - Geochemistry (2014), http://dx.doi.org/10.1016/j.chemer.2014.04.008