Paleoproterozoic Mississippi Valley Type Pb–Zn mineralization in the Ramah Group, Northern Labrador: Stable isotope, fluid inclusion and quantitative fluid inclusion gas analyses

Chemical Geology 362 (2013) 211–223

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Paleoproterozoic Mississippi Valley Type Pb–Zn mineralization in the Ramah Group, Northern Labrador: Stable isotope, fluid inclusion and quantitative fluid inclusion gas analyses J. Conliffe a,⁎, D.H.C. Wilton b, N.J.F. Blamey b,c, S.M. Archibald d a

Geological Survey of Newfoundland and Labrador, St John's, NL, Canada Dept. of Earth Sciences, Memorial University of Newfoundland, St John's, NL, Canada c Dept. Earth and Environmental Science, New Mexico Tech, Socorro, NM 87801, USA d Aurum Exploration Services Canada, Whitby, ON, Canada b

a r t i c l e

i n f o

Article history: Accepted 19 August 2013 Available online 24 August 2013 Keywords: Ramah Labrador Paleoproterozoic MVT mineralization Fluid inclusions Isotopes

a b s t r a c t Mississippi Valley Type (MVT) Pb–Zn sulfide mineralization is common in Phanerozoic rocks but relatively few MVT deposits have been reported from Paleoproterozoic rocks. This study investigates the genesis of numerous Pb–Zn showings in the ca.2.0 Ga Ramah Group, northern Labrador. Pb–Zn mineralization is hosted within the Reddick Bight Dolomite Member (RBDM), Ramah Group. The RBDM is characterized by a well developed secondary porosity associated with solution dissolution of tepee structures in the dolomite. Mineralization occurs as open fill in this secondary porosity, with dark-brown sphalerite, galena and pyrite with gangue quartz, dolomite, calcite and feldspar. The presence of blocks of mineralized RBDM in thrust faults indicates that mineralization predates propagation of these faults during the Torngat Orogeny from 1.85 to 1.87 Ma. Fluid inclusion analysis of sphalerite has identified H2O–NaCl–CaCl2 + CO2 ± N2 ± CH4 mineralizing fluids with high salinities (up to 19.4 eq. wt.% NaCl + CaCl2) and homogenization temperatures of 104 to 177 °C. The presence of significant volatiles in the mineralizing fluids has been confirmed through quantitative fluid inclusion gas analysis of sphalerite, which has recorded 5.9 mol% CO2. Fluid inclusion data from quartz suggests high temperature (135 to 231 °C), high salinity (up to 20.5 eq. wt.% NaCl + CaCl2) fluids with significant volatile contents (quantitative fluid inclusion gas analysis up to 21.5 mol% CO2). Carbon isotope data show a decrease in δ13C from host dolomite (− 1.3 ± 0.9‰) through pre-ore dolomite (−4.7 ± 1.5‰) to post-ore calcite (−5 ± 1.7‰), indicating an influx of isotopically light hydrocarbon-bearing fluids. The δ34S isotopic ratios range from 8.3 to 11.1‰ for early pyrite mineralization, 23.3 to 31.8‰ for late stage pyrite, 16.7 to 32.9‰ for galena and 23.2 to 33.8‰ for sphalerite. The relatively high δ34S values associated with Pb–Zn mineralization, high fluid temperatures (N175 °C) and presence of CO2 fluids indicate ore deposition was associated with thermochemical sulfate reduction. Pb isotope data from galena separates suggests that the Pb was derived from a relatively non-radiogenic source consistent with the Archean basement rocks of the Nain Province. The Pb–Zn mineralization in the RBDM shares many features with classic MVT deposits including geological setting, mineralogy, fluid characteristics and crustal sources for both metals and sulfur. However, MVT mineralization in the RBDM is unusual when compared with Phanerozoic MVT deposits, with elevated fluid temperatures, high CO2 contents of mineralizing fluids and minimum trapping pressures for fluid inclusions corresponding to mineralization depths of N 4.5 km. Similar characteristics have been described from other Paleoproterozoic MVT deposits (e.g. Pering Zn–Pb deposit, South Africa, Kamarga Pb–Zn deposit, Australia). Crown Copyright © 2013 Published by Elsevier B.V. All rights reserved.

1. Introduction Carbonate-hosted Mississippi Valley Type (MVT) deposits are common in Phanerozoic and Neoproterozoic rocks, but relatively rare in Mesoproterozoic, Paleoproterozoic and Archean rocks (b 10 known ⁎ Corresponding author at: Geological Survey of Newfoundland and Labrador, PO Box 8700, St. John's, NL, A1B 4J6, Canada. Tel.: +1 709 729 4014. E-mail address: [email protected] (J. Conliffe).

deposits; Singer et al., 2009; Taylor et al., 2009; Leach et al., 2010). This study reports on carbonate-hosted epigenetic Pb–Zn showings from the Paleoproterozoic Ramah Group in northern Labrador, ~600 km north–northwest of Happy Valley-Goose Bay, Labrador (Fig. 1). Carbonate-hosted Pb–Zn mineralization in the Ramah Group was first reported by Morgan (1975), who noted sphalerite–galena– chalcopyrite mineralization in veins crosscutting the Reddick Bight Dolomite Member (RBDM). ESSO Minerals conducted reconnaissance geological work and geochemical stream sediment heavy mineral

0009-2541/$ – see front matter. Crown Copyright © 2013 Published by Elsevier B.V. All rights reserved. http://dx.doi.org/10.1016/j.chemgeo.2013.08.032

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J. Conliffe et al. / Chemical Geology 362 (2013) 211–223

Fig. 1. Geological maps of the study area. (a) Simplified lithotectonic of northern Labrador, showing the location of the Rae and Nain provinces, and of the Ramah and Mugford supracrustal sequences (adapted from Knight and Morgan, 1981). (b) Generalized map of the Ramah Group, showing the location of the main Pb–Zn showings in the Reddick Bight Dolomite Member (see text for details). BF = Branagin Fault. Adapted from Archibald (1995), geology from Morgan (1975).

sampling over the Ramah Group in the 1980s (Mac Leod, 1985). This work concentrated on the potential of SEDEX deposits in the Nullataktok Formation shale, but reported a number of Pb–Zn showings in the underlying RBDM (MacLeod #1 showing; Fig. 1). In 1993 and 1994, metallogenic investigations carried out by a Memorial University of Newfoundland research group recognized 21 new Pb–Zn occurrences in the RBDM, ranging from minor indications to promising exploration targets (Wilton et al., 1993, 1994). Grab samples from these occurrences returned assays ranging up to 14% Zn and 10.5% Pb (Table 1; Wilton et al., 1993). All of the Pb–Zn showings are now located within the Torngat Mountains National Park of Canada. The current study presents a detailed description of the Pb–Zn mineralization in the Ramah Group. Fluid inclusion, stable (C, S) and radiogenic (Pb) isotopic studies were conducted on samples from a number

Table 1 Location of Pb–Zn showing discussed in text. Assay data from Wilton et al. (1993) and Archibald (1995). Showing Panda Daniels Point MacLeod #1 Pine Harbour 3-0 Saor Alba Blue Sky Reddick Bight V-8

Sample No.

Longitude

Latitude

% Cu

% Pb

% Zn

SA93-058 W93-039A W93-039A W93-039B W93-036D W93-026A SA93-066 SA93-060 W93-056B W93-016 W93-040

63°10.408′ 63°10.51′ 63°10.51′ 63°10.094′ 63°21.944′ 63°13.876′ 63°11.371′ 63°11.743′ 63°22.321′ 63°14.174′ 63°10.991′

58°43.222′ 58°43.225′ 58°43.225′ 58°43.013′ 58°54.262′ 58°56.903′ 58°38.625′ 58°40.457′ 58°55.009′ 58°55.81′ 58°44.195′

0.003 – 0.003 – – 0.004 – – – 0.001 –

0.002 10.5 0.006 4.3 0.97 0.004 0.02 – 0.001 0.004 0.42

14 3.7 14 0.03 8.1 5 0.26 – 0.2 0.2 0.067

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of the most significant showings, in order to characterize ore-forming fluids and identify the processes responsible for ore deposition. The geochemical characteristics of Pb–Zn mineralization in the Ramah Group are compared with previous studies of other Paleoproterozoic MVT deposits, which have identified a wide range of geochemical characteristics (Jones et al., 1999; Huizenga et al., 2006a,b; Kesler et al., 2007; Muhling et al., 2012). 2. Geological setting The Ramah Group in northern Labrador consists of predominantly Paleoproterozoic sedimentary rocks, and is located between the Archean basement of the Nain Province in the east and the late Archean Southeast Rae Structural Province (part of the Churchill Province) in the west (Fig. 1a). The Ramah Group has been mapped discontinuously for a strike length of ~190 km, from Okak Bay in the south to Nachvak Fiord in the North (Fig. 1a). It is best preserved between Saglek Fiord and Nachvak Fiord (Fig. 1b), where it is exposed as a north–northwest trending synclinorium up to 16 km wide (Morgan, 1975; Knight and Morgan, 1981). South of Saglek Fiord, the Ramah Group narrows considerably and is highly deformed and poorly preserved (Ermanovics and Van Kranendonk, 1990). It has been correlated with the Mugford Group, approximately 100 km southeast of the southern-most exposure of the Ramah Group (Fig. 1a), and together these deposits are believed to have been deposited on the western margin of the Nain Province. The deposition of the Ramah Group is thought to have occurred prior to the collision of the Nain and the Churchill provinces during the Torngat Orogeny (1.85 to 1.87 Ma; Wardle et al., 2002). The onset of the Torngat Orogeny is marked by subduction-related ca. 1880 Ma arc magmatism in the margin of the Rae Structural Province, west of the Ramah Group (Bertrand et al., 1993). Upper sedimentary units within the Ramah Group were intruded by extensive diabase and gabbro sills from which Sahin et al. (2012) report dates of 1888 ± 5 and 1887 ± 4 Ma respectively for two sills. The ages of these dykes suggest that the Ramah Group was deposited prior to ca. 1.9 Ga. No direct age for the Ramah Group has yet been derived, but unpublished geochronological data for detrital zircons derived by Wilton and Cox from an the upper strata (the Cameron Brook Formation) indicate youngest ages of ca. 1935 Ma, thus providing an upper bracket on Ramah Group deposition. An unpublished U–Pb titanite age of ca. 1.9 Ga was derived by Wilton for the correlative basal sedimentary unit in the Mugford Group. Deposition of the Ramah Group thus is bracketed between the late Archean (ca. 2.5 Ga) age of the Nain craton and the Torngat Orogeny deformation. The Ramah Group has a total thickness of ~1700 m, and it has been subdivided into six formations (Morgan, 1975; Knight and Morgan, 1981; Fig. 2). Overall the group consists of two sequences, which represent an upward change from shallow, siliclastic shelf to deep basinal deposition (Knight and Morgan, 1981). The lower sequence consists of the Roswell Harbour Formation and the Reddick Bight Formation. The Roswell Harbour Formation varies in thickness from 250 to 470 m (Knight and Morgan, 1981) and consists of quartzite, sandstones and mudstones with a thin (b 20 m) volcanic unit. Knight and Morgan (1981) showed that deposition of the Roswell Harbour Formation occurred in a sandy, shallow shelf and shoreline setting, with minor muddy shelf deposition and sub-aerial volcanism. This is overlain by turbidite sandstones and mudstones of the Reddick Bight Formation (53–143 m thick), which were interpreted to represent northward advancing deltaic complex (Knight and Morgan, 1981). The Reddick Bight Formation is capped by a 4 to 17 m thick dolomite member (Reddick Bight Dolomite Member; RBDM), which can be traced for over 65 km, from Nachvak to Saglek fiords. The RBDM has an irregular gradational base, with a lower dolomitic sandstone and an upper 3–4 m of pure dolomite. This dolomite represents a depositional hiatus, with shallow carbonate deposited over the delta plain in a peritidal enviroment (Knight and Morgan, 1981). Jefferson (1973)

Fig. 2. Simplified stratigraphic column for the Ramah Group (adapted from Swinden et al., 1991, based on Knight and Morgan, 1981).

described evidence for soft-sediment deformation in the dolomite that included liquifaction and intraformational brecciation. The upper pure dolomite is commonly brecciated with a well developed secondary porosity which was in turn cemented by quartz, dolomite and calcite ± sulfides. Knight and Morgan (1981) and Archibald (1995) attributed this brecciation to tepee formation during the diagenesis of precursor carbonate to dolomite, similar to those described by Assereto and Kendall (1977) and Kendall and Warren (1987). Archibald (1995) reported on several generations of tepee formation in the RBDM, with the presence of rounded dolomitic sands (probable aeolian-derived) indicating sub-aerial exposure of the dolomite during deposition. Dissolution of calcite rich areas, probably associated with influxes of meteoric freshwater, was responsible for solution collapse brecciation of the tepee structures.

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The RBDM is unconformably overlain by the upper basin sequence, which consists of the Nullataktok, Warspite, Typhoon Peak and Cameron Brook formations (Fig. 2). The 595 m thick Nullataktok Formation comprises of pyritic shales, varicolored mudstones, calcareous and dolomitic mudstones, sedimentary chert, pyrite iron formation, siliceous dolomite and argillite (Knight and Morgan, 1981). The presence of organic rich, pyritiferous black shales directly overlying the RBDM suggests that subsidence was rapid, with the establishment of stagnant deep water condition (Knight and Morgan, 1981). The Nullataktok Formation passes upwards into a sequence of deep water slope deposits, encompassing the dolomite and dolomitic sandstones of the Warspite Formation, shales and sandstones of the Typhoon Peak Formation and the turbidite dominated Cameron Brook Formation (Knight and Morgan, 1981). Deformation and metamorphism of the Ramah Group are highly variable along its length, (Mengel et al., 1991; Mengel and Rivers, 1994). Between Nachvak Fiord and Saglek Fiord (Fig. 1) the Ramah Group is well preserved in a foreland fold and thrust belt, with the grade of metamorphism increasing from north to south and from east to west (Mengel et al., 1991; Mengel and Rivers, 1994). In this region the Ramah Group can be subdivided into two main zones separated by the Branagin Fault (Fig. 1b), which marks the eastern limit of the Torngat Thrust Zone (Calon and Jamison, 1993). West of the Branagin Fault there is a westward increase in the intensity of metamorphism associated with the Torngat orogeny, from lower greenschist facies

metamorphism close to the Branagin Fault to lower amphibolite facies along the western margin of the Ramah Group (Mengel et al., 1991; Mengel and Rivers, 1994). Rocks to the east of the Branagin Fault (including major Pb–Zn showings in the RBDM; Fig. 1b) are only mildly deformed as open folds with no metamorphic mineral assemblages produced in pelitic rocks (Knight and Morgan, 1981; Mengel et al., 1991). This supports the observation that Pb–Zn mineralization in the RBDM (i.e. east of the Branagin Fault) has not undergone significant metamorphism. 3. Pb–Zn mineralization in the RBDM Numerous Pb–Zn showings have been described in the RBDM (Wilton et al., 1993, 1994; Archibald, 1995). Two main types of Pb–Zn mineralization have been documented; carbonate-hosted Pb–Zn and vein-hosted Pb–Zn (Archibald, 1995). Carbonate-hosted Pb–Zn deposits are open fill style mineralization found within large irregular voids in the RBDM (associated with solution collapse brecciation). These voids are filled with variable amounts of sphalerite, galena, pyrite, quartz and dolomite (Fig. 3a), with minor calcite and feldspar (albite and microcline) in some showings. Colloform sphalerite has also been recorded from some showings (Archibald, 1995). Carbonate-hosted Pb–Zn deposits are found throughout the study area, with highly variable degrees of sulfide mineralization (Table 1). Detailed descriptions

Fig. 3. (a) Colloform sphalerite in the Reddick Bight North Shore Showing. (b) Solid bitumen in calcite from the Pine Harbour Showing. (c)Primary fluid inclusion assemblages in quartz, with trails parallel to crystal boundary. (d) Two-phase (LH2O + VCO2) H2O–CO2–NaCl–CaCl2 inclusions in sphalerite. (e) Three phase (LH2O + LCO2 + VCO2) H2O–CO2–NaCl–CaCl2 inclusions in late stage quartz. (f) Same view as (e) at −130 °C, showing brown granular ice typical of CaCl2 rich fluids.

J. Conliffe et al. / Chemical Geology 362 (2013) 211–223

of these showings can be found in Wilton et al. (1993, 1994) and Archibald (1995). The most significant showings are the Panda and Daniels Point showings, which display a high degree of sulfide mineralization, a lack of deformation and well developed brecciation (Archibald, 1995). These showings are located on a 2.5 km continuous exposure of the RBDM south of Little Ramah Bay. The Panda Showing is approximately 30 m in length and 15 m thick, with mineralization best developed in the lower portions of the dolomite where sphalerite and galena account for up to 5% of the rock. Grab samples from the Panda Showings returned assays ranging up to 14% Zn and 10.5% Pb, (Table 1; Wilton et al., 1993). The Daniels Point Showing, located ~400 m south of the Panda Showing, has very similar sulfide mineralogy and carbonate textures (Wilton et al., 1993, 1994; Archibald, 1995) and may represent a stratigraphic continuation of the same deposit (Archibald, 1995). The following paragenetic sequence, (summarized in Fig. 4) is based on the Panda and Daniels Point showings. The earliest stage of mineralization is associated with a rim of fine-crystalline dolomite overgrowing the brecciated host dolomite, followed by a coarse pre-ore dolomite and a thin zone of early pyrite mineralization. This was followed by the main stage mineralization event, with coeval dark-brown sphalerite and galena, the sphalerite and galena sometimes contain euhedral pyrite grains. Pyrobitumen is common and is associated with early and main stage mineralization (Fig. 3b). The final stage of main stage mineralization is characterized by the deposition of quartz and minor feldspar. Quartz is locally intergrown with sphalerite, indicating that quartz and sphalerite precipitation occurred during the same mineralization event. Postdating the main mineralization event was deposition of late stage calcite which is present as linings in late vugs and cavities. Determining the timing of carbonate-hosted Pb–Zn mineralization is difficult due to the lack of minerals which can be dated within the RBDM. Archibald (1995), however, showed that mineralization was pre-deformational, based on the presence of mineralized blocks of Reddick Bight dolomite within thrust planes on the southern shore of Delabarre Bay (Fig. 1). This indicates that mineralization must have occurred prior to the propagation of the thrust faults during the 1.85 to 1.87 Ma Torngat Orogeny (Wardle et al., 2002). Vein-hosted Pb–Zn showings are rarer than carbonate-hosted Pb–Zn showings, and are most common in the southern portion of the study area (Fig. 1). Mineralization is associated with planar quartz ± dolomite veins up to 10 cm wide, which crosscut the host dolomite. Archibald (1995) concluded that vein-hosted Pb–Zn showings represent galena and sphalerite precipitation during late stage faulting and fluid movement, possibly associated with remobilization from carbonate-hosted Pb–Zn showings in the RBDM. The restriction of vein-hosted Pb–Zn showings to the southern section of the study area (where deformation was greater) and the association of mineralized veins with thrust faults indicate that this remobilization is associated with the Torngat Orogeny.

215

4. Analytical techniques 4.1. Fluid inclusion petrography and microthermometry Fluid inclusion petrography and microthermometric analyses were conducted at the Bruneau Innovation Center, Memorial University of Newfoundland. Microthermometric fluid inclusion analyses were performed using a Linkam THMSG600 heating–freezing stage. Following procedures outlined by Shepherd et al. (1985), the temperatures of clathrate melting (Tm (clath)), last ice melting (Tm (ice)) and the temperature of total homogenization (Th (tot)) were measured in two-phase (liquid + vapor) inclusions hosted in sphalerite and quartz. In addition, the temperatures of CO2 melting (Tm (CO2)), clathrate melting (Tm (clath)) and homogenization of the carbonic phases (Th (CO2)) were recorded in aqueous-carbonic inclusions. Aqueous fluid salinities were calculated using Tm (ice) and the equation of Bodnar (1993) and Oakes et al. (1990), and in aqueous-carbonic inclusions using Tm (clath) and the software package CLATHRATES (Bakker, 1997). 4.2. Gas analysis Quartz and sphalerite were separated using tweezers to provide mineral pure samples for analysis at the Fluid Inclusion Gas Laboratory at the New Mexico Institute of Mining and Technology. These separates were first cleaned with KOH to remove surface organic matter, rinsed with deionized water, and were then dried below 100 °C. About 0.2 g per sample was crushed incrementally under a vacuum of ~10−8 Torr. Fluid inclusion volatile analysis was done using the CFS (crush-fast scan) method (Norman and Blamey, 2001; Parry and Blamey, 2010; Blamey, 2012). The volatiles were analyzed by dual Pfeiffer Prisma quadrupole mass spectrometers, and the system was calibrated using commercial gas mixtures, synthetic inclusions filled with gas mixtures, and three in-house fluid inclusion gas standards as described by Norman and Blamey (2001). Precision is better than 5% for major gaseous species and 0.2% for water/gas ratios, whereas the detection limit for most species is b1 ppm based on in-house standards (Blamey et al., 2012). The system routinely analyzes for gaseous species including H2, He, CH4, H2O, N2, H2S, Ar, CO2, C2H4, C2H6, SO2, C3H6, C3H8, C4H8, C4H10, and benzene. The amount of each species was calculated using software developed by Nigel Blamey and the late David Norman to provide a quantitative analysis (Blamey, 2012). 4.3. Stable isotopes For C and O isotope analyses of calcite and dolomite, samples were handpicked and purity was determined by X-Ray diffraction (XRD) techniques. The crushed calcite and dolomite separates were left to react with anhydrous phosphoric acid for 24–48 h respectively (McCrea, 1950). Evolved CO2 gas was analyzed on a Finnigan MAT 252

Fig. 4. Paragenetic sequence of early stage, main stage and late stage mineralization in the RBDM, based on the Panda and Daniels Point showings (adapted from Archibald, 1995).

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mass spectrometer at the Department of Earth Sciences (EASC), Memorial University of Newfoundland. Sulfide minerals were handpicked and combusted at 1000 °C with CuO to liberate SO2 gas at EASC, the gasses were analyzed at the Ottawa-Carleton Geoscience Centre stable isotope facility. All isotope values are expressed in the delta notation in ‰ (per mil) relative to PDB (Pedee Belemnite Standard) for C isotopes, V-SMOW (Vienna Standard Mean Ocean Water) for O isotopes and CDT (Cañon Diablo Troilite) for S isotopes. Precision of the data is 0.1‰ for C and O isotopes and 0.2‰ for S isotopes. Galena grains were handpicked from the samples at EASC. Lead isotope ratios were determined at the GEOTOP laboratories, Université du Québec à Montréal., Data are reported as 206Pb/204Pb, 207Pb/204Pb and 207 Pb/206Pb with an analytical uncertainty of 0.05% amu-1 at the 1 σ level (Moritz and Malo, 1996). 5. Fluid inclusion petrography and microthermometry Samples of coarsely crystalline sphalerite and quartz were selected from the Panda Showing to determine the nature of the fluids associated with mineralization. Fluid inclusions were studied in samples from the Panda Showing due to the optical characteristic of quartz and sphalerite, absence of abundant of secondary inclusions trails, and because the inclusions exhibit no evidence of post-entrapment modification (i.e. stretching, leaking, necking down; Bodnar, 2003). This fluid inclusion petrographic study was based on the concept of fluid inclusion assemblages as described by Goldstein (2003), an approach that places fluid inclusions into assemblages interpreted to have been trapped penecontemporaneously. Fluid inclusion data were collected from both primary and pseudosecondary fluid inclusion assemblages, representing fluids trapped during crystal growth. Primary fluid inclusion assemblages occur either as clusters of inclusions in the core of sphalerite crystals or in trails that parallel crystal boundaries in both sphalerite and quartz (Fig. 3c). Rare pseudosecondary fluid inclusion assemblages were recorded in sphalerite and quartz, with trails of inclusions in fractures that terminate at growth zones (represent fracturing during crystal growth). Clusters of inclusions of unknown origin were also observed in both quartz and sphalerite. Microthermometric data were collected from two–three inclusions in each fluid inclusion assemblage. Phase changes in a single fluid inclusion assemblage generally occurred over 1–2 °C, consistent with petrographic observations that inclusions did not undergo post-entrapment modification (Bodnar, 2003). Based on the chemical composition of the inclusions, three main types were recorded (Table 2): (1) Two-phase H2O–NaCl–CaCl2 inclusions with no evidence of CO2 recorded on cooling; (2) Two-phase (LH2O + VCO2) H2O–CO2–NaCl–CaCl2 inclusions

forming clathrate, but no observable solid CO2, on cooling (Fig. 3d); and (3) Three phase (LH2O + LCO2 + VCO2) H2O–CO2–NaCl–CaCl2 inclusions forming solid CO2 on freezing (Fig. 3e). Microthermometric analyses were carried out on 66 inclusions, and the results are summarized in Table 2. 5.1. Type 1 inclusions Type 1 inclusions were recorded as primary (rare pseudosecondary) inclusions in sphalerite only. Upon freezing these inclusions formed a brown, granular ice at ~ −75 °C. First ice melting was observed in a single Type 1 inclusion at −50.7 °C, close to the eutectic point for the H2O–NaCl–CaCl2 system. Tm (ice) occurred between −12.6 and −16 °C, corresponding to salinities of 16.5 to 19.4 eq. wt.% NaCl (Fig. 5), according to the equations of Bodnar (1993) and Oakes et al. (1990). Inclusions of Type 1 homogenized into liquid phase (L + V → L) between 104 and 125 °C (Fig. 6). 5.2. Type 2 inclusions Type 2 inclusions occur as primary and pseudosecondary assemblages in sphalerite and quartz. Tm (eutectic) was recorded at −55.1 °C in one sphalerite-hosted inclusion. The formation of brown granular ice upon freezing in other Type 2 inclusions is consistent with the presence of Ca2+ salts in solution (Shepherd et al., 1985). Tm (ice) was recorded in Type 2 inclusions between −17.4 and −8.2 °C. Clathrate melting occurred between 0.8 and 4.2 °C (consistent with the presence of significant volumes of CO2 in these inclusions) and was used with last ice melting temperatures to calculate fluid salinities of 15.9 to 16.4 eq. wt.% NaCl for sphalerite-hosted inclusions and 15.8 to 20.5 eq. wt.% NaCl for Type 2 inclusions in quartz. Type 2 inclusions homogenize to the liquid phase between 121 and 227 °C (higher temperatures in quartz-hosted Type 2 inclusions; Fig. 6). 5.3. Type 3 inclusions Type 3 inclusions were observed as primary inclusions in quartz only. Upon cooling to approx. −130 °C, brown granular ice forms at ~ −70 °C (indicative of CaCl2) and the volatile freezes at ~ −120 °C (Fig. 3f). Melting of the volatile phase (Tm (CO2)) occurs between −61 and −57 °C, lower than the melting temperature of pure CO2 (−56.6 °C) indicating the presence of other volatile species (e.g., CH4, N2). Clathrate melting, in the presence of liquid CO2, occurred between 2.2 and 4.1 °C. In conjunction with the non-aqueous fluid composition (see below), the clathrate melting temperatures were used to calculate salinities between 12.5 and 14.7 eq. wt.% NaCl (using the CLATHRATES

Table 2 Classification of fluid inclusion types and summary of microthemometric data from sphalerite and quartz-hosted fluid inclusions. Te = eutectic temperature; Tm (CO2) = temperature of CO2 melting; Tm (ice) = last ice melting; Tm (clath) = temperatures of clathrate melting; Tm (clath) = temperature of clathrate melting; Th (CO2) = temperature of homogenisation of the carbonic phase; Th (tot) = temperature of total homogenization. Sphalerite

Quartz

Type 1

Type 2

Type 2

Type 3

Paragenesis

Primary, pseudosecondary trails H2O + NaCl + CaCl2

n Fill Microthermometry Te (°C) TmCO2 (°C) Tm ice (°C) Tmclath (°C) Salinity (eq. wt.% NaCl) ThCO2 (°C) Thtot (°C)

12 0.95 Min.

Max.

Mean −50.7

Primary, pseudosecondary trails and clusters of unknown origin H2O + NaCl + CaCl2 + CO2 ± N2 16 0.9 Min. Max. Mean

Primary growth zones

Composition

Pseudosecondary trails and clusters of unknown origin H2O + NaCl + CaCl2 + CO2 ± N2 26 0.9 to 0.95 Min. Max. Mean −55.1

−16

−12.6

−15.1

16.5

19.4

18.7

−16.4 2.7 15.9

−8.2 4.2 16.4

−11.9 3.2 16

−17.4 0.8 15.8

−11.8 4.1 20.5

−14.4 1.9 18.4

104

125

116

121

177

148

135

227

184

H2O + NaCl + CaCl2 + CO2 ± N2 12 0.85 to 0.9 Min. Max. Mean −57

−61

−58.8

2.2 12.5 15.2 207

4.1 14.7 21.2 231

3.1 13.5 18 219

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7. Stable isotopes 7.1. C and O of carbonates Carbon and oxygen isotope data from host dolomite, pre-ore dolomite and post-ore calcite are listed in Table 4. The host dolostones have a wide range of δ18O values (−21.5 to −9.8‰), with a mean of −13.6 ± 3.7‰. δ13C values range from −3.1 to −0.3‰ (mean − 1.3 ± 0.9‰). Pre-ore dolomite has a similar range of δ18O values (− 19.8 to − 10.8; mean − 15.1 ± − 2.6‰) but lower δ13C values (− 7.8 to − 2.9‰; mean − 4.7 ± − 1.5‰). Finally post-ore calcite has δ18O values between − 15.6 and − 13‰ (mean − 14.3 ± 1.3‰) and δ13C values between − 6.6 and − 3.4‰ (mean − 5 ± 1.8‰). Overall the carbon and oxygen isotope data show that carbonates associated with mineralization are depleted in 13C, with no corresponding change in 18O (Fig. 7). Fig. 5. Bivariate plot of X CO2 vs. salinity for sphalerite-hosted Type 1 and 2 inclusions and quartz-hosted Type 3 inclusions.

software packages; Bakker, 1997). Th (CO2) was recorded for the liquid vapor phases between 15.2 and 21.2 °C. Total homogenization to the liquid phase occurs between 207 and 231 °C. A number of inclusions decrepitated prior to total homogenization at ~200 °C.

6. Fluid chemistry The chemistry of the ore-forming fluid can be constrained from fluid inclusion gas analysis, supported by microthermometry. The fluid inclusion gas analyses are summarized in Table 3, the data being reported as burst-size weighted means which is equivalent to a bulk sample analysis. Sphalerite and quartz fluid inclusion gasses are generally similar in concentration with the exception of CO2 and He. CO2 contents are higher in quartz (8.5 to 21.5 mol%) than in sphalerite (0.9 to 5.9 mol%). The He content of sphalerite is almost two orders of magnitude greater than values from quartz (Table 3). This may reflect changes in the metal content of mineralizing fluids over time. Given the age of the inclusion fluids (~1.8 to 2.0 Ga) it is conceivable that He may have been generated in situ via the decay of radioactive elements within the sphalerite (e.g. K, U, and Th).

7.2. S of sulfides Sulfur isotopic data were obtained from pyrite, galena and sphalerite from the showings present in the RBDM and these data are summarized in Table 5. Care was taken to exclude analyses from remobilized sulfides due to possible isotopic reequilibration. Values of δ34S from sulfides in the RBDM show a wide range (Fig. 8), with δ34S ratios of early pyrite mineralization from 8.3 to 11.1‰, late stage pyrite from 21.6 to 31.8‰, galena from 16.7 to 32.9‰ and sphalerite from 23.2 to 33.8‰. The sulfur isotope data from the Ramah Group sulfides are generally higher than estimated δ34S values of Paleoproterozoic seawater (10 to 18‰; Strauss, 1993) and seem to indicate an increase from early pyrite to late stage pyrite, galena and sphalerite. 8. Radiogenic isotopes 8.1. Pb isotopes Lead isotope ratios were measured for galena separates and are listed in Table 6. The ratios range from 14.092 to 14.524 for 206Pb/204Pb, 14.546 to 14.636 for 207Pb/204Pb, and 34.117 to 34.366 for 208Pb/204Pb. Mu (μ) values are very non-radiogenic (6.617 to 7.037) and suggest that the lead source area had undergone uranium loss earlier in its history. Stacey and Kramers' (1975) model ages range from 1559– 1919 Ma. These ages are definitively younger than the host dolomites and are seemingly younger than the age of mineralization (as constrained from structural relationships). The discrepancy in the ages may in part be due to the model itself which is based on “average crust” growth curve rather than derivation from a demonstrably nonradiogenic source (Archibald, 1995). The Pb isotope data also suggest derivation of the Pb in the Ramah occurrences from a low μ source, most likely a low μ reservoir in the Archean crust of the Nain Province (Wilton, 1991; Ashwal, 1993). 9. Discussion 9.1. Nature of ore-forming fluids

Fig. 6. Frequency distribution histogram for homogenization temperatures of sphaleritehosted Type 1 and 2 inclusions and quartz-hosted Type 2 and 3 inclusions.

Type 1, Type 2 and Type 3 fluid inclusions in quartz and sphalerite from the RBDM Pb–Zn showings contain high salinity (up to 20 eq. wt.% NaCl) H2O–NaCl–CaCl2 brines. Such fluids are common in sedimentary basins (Hanor, 1994) and are frequently associated with sedimentary hosted Pb–Zn deposits (see Wilkinson, 2001). Although quartz postdates sphalerite precipitation, field relationships (presence of mineralized blocks of sphalerite and quartz in thrust fault breccias) and similarities between primary Type 1 and Type 2 fluid inclusions (CaCl2 rich brines with similar CH4, N2, and Ar contents) indicate that quartz precipitation was associated with the waning stages of mineralization. The slightly lower salinities of the quartz-hosted fluid inclusions

218

J. Conliffe et al. / Chemical Geology 362 (2013) 211–223

Table 3 Fluid inclusion gas analysis for sphalerite and quartz mineral separates, with the data reported as burst-size weighted means, equivalent to a bulk sample analysis. Sphalerite Sample

8176d

8176e

0.000 0.025 0.141 97.40 0.694 0.002 0.016 1.680 0.002 0.000 0.023 0.014 0.000 0.000 0.002 0.001 11.9 42.6 2.4

H2 He CH4 H2O N2 H2S Ar CO2 SO2 C2H4 C2H6 C3H6 C3H8 C4H8 C4H10 Benzene CO2/CH4 N2/Ar CO2/N2

8176f

0.000 0.017 0.144 98.04 0.795 0.004 0.024 0.953 0.001 0.002 0.015 0.006 0.000 0.001 0.000 0.000 6.6 33.4 1.2

0.000 0.009 0.093 96.35 0.843 0.004 0.022 2.661 0.001 0.000 0.013 0.008 0.000 0.000 0.000 0.000 28.7 38.2 3.2

8176g 0.000 0.009 0.211 94.46 0.861 0.005 0.022 4.412 0.001 0.000 0.015 0.009 0.000 0.000 0.001 0.000 20.9 38.9 5.1

8176h 0.000 0.013 0.147 92.80 0.969 0.011 0.041 5.994 0.002 0.000 0.014 0.005 0.000 0.001 0.001 0.000 40.8 23.6 6.2

8176j 0.000 0.011 0.136 95.03 1.013 0.006 0.037 3.743 0.001 0.000 0.011 0.007 0.000 0.000 0.001 0.000 27.6 27.6 3.7

8176k 0.000 0.007 0.118 97.25 0.978 0.005 0.035 1.596 0.001 0.001 0.008 0.004 0.000 0.000 0.000 0.000 13.5 27.9 1.6

8176l 0.000 0.013 0.134 97.10 0.820 0.002 0.019 1.875 0.000 0.000 0.024 0.015 0.000 0.001 0.000 0.001 14.0 43.0 2.3

8176m 0.000 0.011 0.112 97.63 0.816 0.002 0.022 1.378 0.001 0.000 0.016 0.009 0.000 0.000 0.000 0.000 12.3 36.3 1.7

8176n 0.000 0.020 0.186 96.55 1.042 0.001 0.023 2.138 0.001 0.000 0.027 0.013 0.000 0.001 0.001 0.000 11.5 45.7 2.1

Mean 0.0000 0.0135 0.1421 96.2595 0.8831 0.0042 0.0261 2.6431 0.0011 0.0003 0.0167 0.0090 0.0000 0.0005 0.0006 0.0003 18.7900 35.7180 2.9444

Quartz Sample

8177a

8177c

8177d

8177e

8177f

8177g

8177h

8177j

8177k

8177l

Mean

H2 He CH4 H2O N2 H2S Ar CO2 SO2 C2H4 C2H6 C3H6 C3H8 C4H8 C4H10 Benzene CO2/CH4 N2/Ar CO2/N2

0.000 0.000 0.095 86.13 0.677 0.004 0.018 13.01 0.001 0.000 0.031 0.027 0.000 0.000 0.000 0.000 137.6 38.1 19.2

0.000 0.000 0.079 89.51 0.697 0.005 0.013 9.61 0.002 0.000 0.041 0.042 0.000 0.001 0.000 0.000 122.2 55.7 13.8

0.000 0.000 0.086 88.08 0.623 0.003 0.014 11.14 0.001 0.000 0.029 0.029 0.000 0.001 0.000 0.000 129.4 45.5 17.9

0.000 0.000 0.110 82.97 0.590 0.008 0.024 16.26 0.001 0.000 0.023 0.011 0.000 0.001 0.000 0.000 148.1 24.7 27.6

0.000 0.000 0.140 90.56 0.675 0.003 0.013 8.53 0.002 0.000 0.039 0.036 0.000 0.001 0.001 0.000 60.9 50.6 12.6

0.000 0.000 0.124 83.95 0.633 0.007 0.025 15.20 0.001 0.000 0.030 0.026 0.000 0.001 0.001 0.000 122.5 25.0 24.0

0.000 0.000 0.114 87.88 0.771 0.005 0.020 11.17 0.001 0.000 0.024 0.017 0.000 0.000 0.000 0.000 98.4 39.5 14.5

0.000 0.001 0.067 77.69 0.578 0.010 0.030 21.55 0.001 0.000 0.021 0.000 0.047 0.001 0.000 0.000 322.4 19.2 37.3

0.000 0.000 0.068 80.34 0.646 0.011 0.029 18.88 0.002 0.000 0.020 0.006 0.000 0.001 0.001 0.000 279.2 22.4 29.2

0.000 0.000 0.118 78.39 0.657 0.005 0.024 20.77 0.001 0.000 0.027 0.005 0.000 0.001 0.001 0.000 176.0 27.8 31.6

0.0000 0.0003 0.0999 84.5500 0.6548 0.0062 0.0209 14.6123 0.0013 0.0000 0.0285 0.0199 0.0047 0.0006 0.0004 0.0002 159.6686 34.8300 22.7702

(Fig. 5) may be related to minor dilution of ore-forming fluids by meteoric fluids. Fluid inclusion microthermometry and quantitative fluid inclusion gas analyses have identified variable amounts of volatiles (mainly CO2 Table 4 Carbon and oxygen isotopes of carbonates. Carbonate tYPE

Sample number

δ13CPDB (‰)

δ18OSMOW (‰)

Showing

Host dolomite

W93-016 W93-056B W93-36D SA93-033 SA93-060 W93-039 W93-039 W93-039 W93-039 W93-039 W93-036D W93-039 W93-039 W93-039 W93-036D W93-039b SA93-058 W93-016

−1 −3.1 −1.1 −1.3 −1.8 −0.8 −0.3 −2.9 −4.4 −3.9 −4.7 −4.6 −4.6 −7.8 −6.5 −6.6 −3.5 −3.4

−21.5 −13.4 −12.7 −9.8 −13.2 −12.2 −12.2 −14.3 −15.4 −15.4 −10.8 −15.2 −14.4 −19.8 −13.0 −15.2 −15.6 −13.3

Reddick Bight N.S. Blue Sky MacLeod #1 n/a Saor Alba Panda Panda Panda Panda Panda MacLeod #1 Panda Panda Panda MacLeod #1 Panda Panda n/a

Pre-ore dolomite

Post-ore calcite

with minor CH4, N2 and H2S) in fluids from the RBDM. These fluids show an increase in volatile content over time, from main stage sphalerite (0.9 to 5.9 mol%) to paragenetically later quartz (8.5 to 21.5 mol%). The increase in CO2 is not matched by any other changes in fluid composition (e.g., CH4, N2, and Ar contents), suggesting that CO2 is a product of thermochemical sulfate reduction during mineralization (see below), rather than a late influx of CO2-rich fluids. Fluid inclusion gas analysis also indicates that magmatic fluids were not involved at any mineralization stage in the RBDM. Blamey (2012) showed that magmatic volatiles have N2/Ar ratios in the 100s to 1000s, much higher than meteoric (N2/Ar ~38) and basinal fluids (N2/Ar approaching 1). N2/Ar ratios from quartz and sphalerite from the RBDM range from 19.2 to 55.7 and 23.6 to 45.7 respectively (Table 3), and plot in the field typical of meteoric water on a CO2/CH4 vs. N2/Ar discrimination diagram (Fig. 9). However, this may not reflect the original fluid composition due to the increase in CO2 content during mineralization, and therefore these brines may reflect more evolved crustal fluids. Estimates of the minimum trapping temperatures and pressures of primary fluid inclusion assemblages have been calculated using measured homogenization temperatures and compositional data from fluid inclusion microthermometry and gas analysis (Fig. 10). Isochores for aqueous Type 1 inclusions and aqueous-carbonic Type 3 inclusions were constructed using FLUIDS (Bakker, 2003), according the equations of state of Zhang and Frantz (1987) and Bakker (1999 — after Bowers

J. Conliffe et al. / Chemical Geology 362 (2013) 211–223

Fig. 7. Carbon and oxygen stable isotope data from host dolomites and carbonates associated with mineralization in the RBDM.

and Helgeson, 1983), respectively. Partial pressures of H2O, CO2 (3.6 mol%), and CH4 (0.14 mol%) from 160 to 300 °C were calculated for Type 2 inclusions, according to the equations of Drummond (1981) and Henley (1984). These data yield minimum PT conditions of 430 bars at 160 °C for Type 2 inclusions in quartz and sphalerite and 1375 bars at 215 °C for Type 3 inclusions in quartz. Care must be taken in interpreting these pressures and temperatures, as they represent minimum conditions, and significant pressure corrections may be needed when discussing true trapping pressures and temperatures. Both Type 2 and Type 3 inclusions have been recorded in primary fluid inclusion assemblages in quartz, with isochores intersecting at ~225 °C and ~1750 bars (Fig. 10). However, this may not represent true PT conditions of mineralization, as mineralization in the RBDM may not have occurred at a constant temperature. If mineralization occurred over a range of temperatures (175–225 °C; Fig. 10), pressures of approximately 1250 to 1500 bars can be calculated. This is consistent with minimum trapping pressures of Type 3 inclusions (1375 bars), and corresponds to depths of 4.7 to 5.7 km (lithostatic fluid pressure conditions and a density of overlying rocks of 2.7 g/cm3).

Table 5 Sulfur isotope data for sulfide minerals. Mineral

Sample Number

δ34SCDT (‰)

Showing

Early Pyrite

SA93-066 SA93-060 W93-056A W93-039B W93-036D W93-039A W93-040 W93-056B W93-036D W93-036D SA93-058 W93-039A SA93-060 W93-056B W93-036D SA93-058

8.3 11.1 32.9 18 16.7 21.5 17.6 33.8 24.3 24.2 30.1 23.2 24.7 31.8 21.6 27.1

# 3-0 Saor Alba Blue Sky Daniel's Point MacLeod #1 Panda V-8 Blue Sky MacLeod #1 MacLeod #1 Panda Panda Saor Alba Blue Sky MacLeod #1 Panda

Galena

Sphalerite

Late Pyrite

219

Fig. 8. Sulfur isotope data for early stage pyrite, late stage pyrite, sphalerite and galena.

9.2. Ore deposition in the RBDM The increase in δ34S from early pyrite to late stage pyrite, galena and sphalerite in the RBDM suggests some form of Rayleigh type fractionation during the reduction of sulfate to sulfide. In a closed system, the ongoing extraction of sulfide minerals will result in an increase in the δ34S ratio of the residual sulfur reservoir. Mineralization temperatures of b 180 °C (based on fluid inclusion data) are above the limit of biogenic sulfate reduction but in the range where the thermochemical reduction of sulfate (TSR) can operate (Machel, 1987, 2001; Worden et al., 2000). Although slow reaction kinetics can inhibit TSR at low temperatures (Worden et al., 2000), Anderson and Thom (2008) and Thom and Anderson (2008) showed that at high temperature (N150 °C), TSR can occur at sufficient speeds to form significant sedimentary hosted Pb–Zn deposits. TSR is controlled by reactions of hydrocarbons with aqueous sulfate species, according to the schematic equation (Machel, 1987, 2001): 2−

hydrocarbons þ SO4 → altered hydrocarbons þ solidbitumen þ H2 S þ CO2 þ H2 O þ heat Carbon isotope data from carbonates associated with mineralization supports an influx of isotopically light hydrocarbons (Fig. 7). The δ13C values of pre-ore dolomite (− 7.8 to − 2.9‰) and post-ore calcite (−6.6 and −3.4‰) in the RBDM are lower than the δ13C values of the host dolomites (−3.1 to −0.3‰). Therefore authigenic carbonates have incorporated a component of isotopically light carbon, most likely from an organic source (Machel et al., 1995). The source of these hydrocarbons is likely related to maturation of the overlying Nullataktok Formation shale (1.7 to 2.8% organic carbon; Hayashi et al., 1997). The source of sulfur in the RBDM mineralizing system remains unknown. In Phanerozoic MVT deposits, hydrogen sulfide is commonly derived from reduction of dissolved sea water sulfate and/or associated with the dissolution of solid calcium sulfate. Paleoproterozoic seawater from 2.0 to 1.8 Ga, however, is thought to have had low sulfate concentrations (Bekker and Holland, 2012), and there is no direct evidence for the presence of gypsum or anhydrite in the Ramah Group (Archibald, 1995). One possible source is in the overlying Nullataktok Formation. An 8–20 m thick pyrite-chert bed is recorded ~60 m from the base of

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J. Conliffe et al. / Chemical Geology 362 (2013) 211–223

Table 6 Lead isotope ratios calculated from galena separates; model ages calculated using Stacey and Kramers (1975) growth curve. Sample

206

Pb/204Pb

207

Pb/204Pb

208

Pb/204Pb

Mu (μ)

Model Age (Ma)

Showing

W93-39A W93-39B W93-40 W93-36D W93-56

14.092 14.137 14.524 14.152 14.168

14.546 14.548 14.636 14.582 14.612

34.146 34.121 34.117 34.27 34.366

6.671 6.617 6.722 6.84 7.037

1861 1810 1559 1871 1919

Panda Daniel's Point V-8 Macleod #1 Bluesky

the Nullataktok Formation (Knight and Morgan, 1981), which formed as a result of bacterial activity (Archibald, 1995). This bacterial activity may have released sulfate-rich fluids which would have been trapped within the pores of the Nullataktok Formation shales. These fluids would have subsequently been released due to tectonic pressures during the initiation of the Torngat Orogeny (Archibald, 1995) and transported to the site of mineralization. The products of TSR reactions include H2S (responsible for galena and sphalerite precipitation in the RBDM) and the CO2 recorded by fluid inclusion gas analysis. In addition solid pyrobitumen is commonly associated with Pb–Zn showings in the RBDM (Archibald, 1995). 9.3. Classification of Pb–Zn mineralization in the RBDM Pb–Zn mineralization in the RBDM shares many features with MVT mineralization (as defined by Leach et al., 2005). These include: 1) restriction of mineral deposits to dolomite horizons; 2) lack of any visible relationship to magmatic rocks; 3) location of showings on the flanks of a basin; 4) relatively simple mineralogy, dominated by sphalerite, galena, pyrite, dolomite and quartz (minor calcite and feldspar); 5) ore controlled by faults and dissolution breccias; 6) coarsely crystalline sulfides, with colloform sphalerite locally observed in outcrop; 7) mineralizing fluids are high salinity (10 to 30 eq. wt.% NaCl) CaCl2rich brines, and 8) crustal source of both metals and sulfur. Therefore Pb–Zn mineralization in the RBDM can be classified as MVT style. However, mineralization in the RBDM also has a number of unusual features when compared with Phanerozoic MVT mineralization. Although CH4 (±CO2) has been recorded from some MVT deposits (e.g. Jones and Kesler, 1992), concentrations of non-aqueous volatiles are generally low (b1 mol%) when compared to the data for sphalerite (2 to 7.2 mol%) and quartz (9.4 to 22.3 mol%) from the RBDM.

Fig. 9. CO2/CH4 vs. N2/Ar discrimination plot for handpicked sphalerite (circles) and quartz (diamonds) from the Ramah deposit. The fluid inclusion gasses from both minerals have b100 N2/Ar ratios and plot in the shallow meteoric field (template from Norman and Moore, 1999).

Mineralizing fluids in the RBDM also have higher temperatures (175– 225 °C) than the majority of Phanerozoic MVT deposits (generally b150 °C; Wilkinson, 2001). 9.4. Comparisons with other Precambrian MVT deposits Relatively few MVT deposits have been described from Precambrian rocks, with ten MVT deposits reported from carbonate rocks older than 1.0 Ga (Table 7). This uneven distribution of carbonate MVT deposits through time remains enigmatic, and has been attributed to a number of factors, including changes in the chemistry of the atmospheric and hydrosphere, changes in tectonic processes throughout the Archean and Paleoproterozoic and tectonic recycling and destruction of ore deposits (Leach et al., 2001, 2005, 2010; Kesler and Reich, 2006; Farquhar et al., 2010). Similarities between the geological settings and mineralogies of these deposits and Phanerozoic MVT deposits suggest the general model for the formation of MVT deposits was similar throughout most of Earth history, associated with tectonically related flow of basinal brines during ocean closure (Leach et al., 2001). This is supported by the studies of late Archean and Paleoproterozoic MVT deposits in the Transvaal Basin, South Africa (Kesler et al., 2007) and the Earaheedy Basin, Western Australia (Muhling et al., 2012) which have similar geochemical characteristics to modern MVT deposits. However, a number of Archean and Proterozoic MVT deposits share unusual geochemical characteristics with mineralization. Jones et al. (1999) reported high mineralization temperatures (270–320 °C) and CO2-rich inclusions associated with high salinity (16–22 eq. wt.% NaCl) aqueous inclusions in sphalerite from the Kamarga MVT deposit in Australia (hosted by 1.65 Ga Paradise Creek Formation). Fluid inclusion studies from the Bushy Park and Pering MVT deposits in South Africa (2.05 Ga; Gleason et al., 2011) have documented a CH4–CO2–HS− fluid which was important during the

Fig. 10. Representative isochores for sphalerite-hosted Type 1 and 2 inclusions and quartzhosted Type 3 inclusions. Estimated PT conditions of mineralization shown by gray shaded box. See text for details.

J. Conliffe et al. / Chemical Geology 362 (2013) 211–223

221

Table 7 Late Archean, Paleoproterozoic and Mesoproterozoic MVT deposits. Data compiled from Taylor et al. (2009), and Leach et al. (2010), with the exception of the Transvaal Basin (Kesler et al., 2007) and Earaheedy Basin (Muhling et al., 2012). Deposit

Host-rock age (Ga)

Mineralization age (Ga)

Size (Mt)

Reference

Late Archean MVT mineralization Bushy Park South Africa

Location

2.52–2.65

2.0–2.1

0.63

Pering

South Africa

2.52–2.65

2.0–2.1

18

Transvaal Basin

South Africa

2.4–2.6

2.0–2.35

not available

Wheatley et al. (1986), Greyling (2000), Duane et al. (2004) Wheatley et al. (1986), Greyling (2000), Duane et al. (2004) Kesler et al. (2007)

Paleoproterozoic MVT mineralization Ramah Canada Earaheedy Basin Australia Esker Canada Kamarga Australia Bulman Australia Coxco Australia

~2.0 2.0–2.2 1.89 1.6–1.8 ? 1.6–1.68

N1.85 1.81–1.83 ? ? ? 1.6–1.68

Not available Not available 80.6 50 0.38 Not available

this study Muhling et al. (2012) Gummer et al. (1997) Jones et al. (1999) Plumb et al. (1990) Walker et al. (1983)

Mesoproterozoic MVT mineralization Nanisivik Canada

1.2

?

17.9

Sherlock et al. (2004)

mineralization process (Huizenga et al., 2006a,b). This fluid was trapped with high salinity aqueous fluids under conditions of immiscibility, and PT modeling estimated trapping conditions of 200–210 °C and 1.1–1.4 kbar (Huizenga et al., 2006a). These pressures correspond to mineralization depths of 4.1 to 5.2 km, similar to the minimum mineralization depths calculated from the RBDM (N 4.5 km) and greatly exceeding predictions for Phanerozoic MVT deposits (Leach et al., 2001). Huizenga et al. (2006a) also noted that TSR was an important process in the mineralizing system at the Pering deposit. With the higher temperatures (N150 °C) and depths proposed for the RBDM, Bushy Park/Pering and Kamarga deposits, TSR would have been able to overcome kinetic constraints and could account for sufficient production of reduced sulfate within the deposits. Therefore, MVT mineralization in the RBDM, Bushy Park/Pering and Kamarga deposits is associated with the migration of metal-bearing brines to deeper crustal levels than Phanerozoic MVT deposits. This may be related to temporal changes in the concentration of seawater sulfate, with lower concentrations of seawater sulfate and the rarity of marine evaporate sequences in sedimentary rocks from 2.1 to 1.2 Ga (Kah et al., 2001; Bekker and Holland, 2012) limiting the potential sources of sulfate necessary for ore formation (Kesler and Reich, 2006). MVT mineralization in late Archean and Paleoproterozoic rocks would occur when suitable local conditions for ore deposition are encountered (e.g., rare evaporite sequences, local accumulations of sulfate-rich marine brines, sulfate-rich fluids trapped in connate fluids, magmatic sulfur etc.). Kesler and Reich (2006) noted that the Kamarga deposit was located close to evaporite sequences, and sulfate at Pering may be associated with the dissolution of locally sourced magmatic sulfides (Huizenga et al., 2006a; Kim et al., 2009), possible associated with distal hydrothermal circulation around the Bushveld Igneous Complex (Gleason et al., 2011). Leachate and decrepitate compositions of inclusion fluids from MVT mineralization in the Transvaal Basin indicate that sulfate is sourced in unusual sulfate-rich marine brines in the overlying Transvaal Basin (Kesler et al., 2007). In the RBDM, a suitable sulfate source remains enigmatic, but may have been shale-hosted connate water in the overlying Nullataktok Formation (Archibald, 1995). However such local conditions would be rare in carbonate platforms and this could account for the scarcity of MVT deposits in Precambrian rocks. 10. Conclusions Carbonate-hosted Pb–Zn mineralization in the ~2.0 Ga Ramah Group shares many characteristics of Phanerozoic MVT mineralization,

including restriction of mineral deposits to dolomite horizons, relatively simple mineralogy (dominated by sphalerite, galena, pyrite, dolomite, quartz and calcite), high salinity (10 to 30 eq. wt.% NaCl) CaCl2-rich mineralizing fluids (brines), and crustal sources for both metals and sulfur. Based on these similarities Pb–Zn mineralization in the Ramah Group can be classified as MVT style, based on the criteria of Leach et al. (2005). However, mineralization also exhibits a number of unusual features when compared with MVT mineralization hosted in Phanerozoic terranes. Fluid inclusion and isotopic data indicate that mineralization occurred at much higher high temperatures (175–225 °C) and pressures (1250 to 1500 bars) than Phanerozoic deposits. The presence of aqueous-carbonic fluids has also been recorded, which is attributed to the production of CO2 during thermochemical sulfate reduction (TSR). Similar unusual characteristic have been noted in other Paleoproterozoic MVT deposits, notably the Pering deposit in South Africa (Huizenga et al., 2006a) and the Karmarga deposit in Australia (Jones et al., 1999).

Acknowledgments The crew of the MV Robert Bradford, skipper Kevin and Roy, are thanked for facilitating access during fieldwork; R. Butler, Jr., A. Hussey, and M. Gates for assistance in the field. Major funding for the field work and initial analyses was provided by a grant to Wilton from the Comprehensive Labrador Cooperation Agreement (ACOA-ENL); analytical costs were also borne partly by an NSERC grant and MUN Metallogenic Research account for Wilton.

References Anderson, G.M., Thom, J., 2008. The role of thermochemical sulfate reduction in the origin of Mississippi Valley-type deposits. II. Carbonate–sulfide relationships. Geofluids 8 (1), 27–34. Archibald, S.M., 1995. The Geology, Mineralization, Geochemistry, and Metallogeny of the Paleoproterzoic Ramah Group, Northern Labrador. Unpublished MSc. thesis, Memorial University of Newfoundland, St John's, Canada, 424 pp. Ashwal, L.D., 1993. Anorthosites, v. 21, Series on Minerals and Rocks. Springer-Verlag, New York, Berlin. Assereto, R.L., Kendall, C.G.St.C., 1977. Nature, origin and classification of peritidal tepee structures and related breccias. Sedimentology 24, 153–210. Bakker, R.J., 1997. Clathrates: computer programs to calculate fluid inclusion V–X properties using clathrate melting temperatures. Comput. Geosci. 23, 1–18. Bakker, R.J., 1999. Adaption of Bowers & Helgeson (1983) equation of state to isochore and fugacity coefficient calculation in the H2O–CO2–CH4–N2–NaCl fluid system. Chem. Geol. 154, 225–236. Bakker, R.J., 2003. Package FLUIDs 1. Computer programs for analysis of fluid inclusion data and for modelling bulk fluid properties. Chem. Geol. 194, 3–23.

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