Age and origin of uranium mineralization in the Camie River deposit (Otish Basin, Québec, Canada)

Accepted Manuscript Age and origin of uranium mineralization in the Camie River deposit (Otish Basin, Québec, Canada) M. Lesbros-Piat-Desvial, G. Beaudoin, J. Mercadier, R. Creaser PII: DOI: Reference:

S0169-1368(17)30438-9 https://doi.org/10.1016/j.oregeorev.2017.10.006 OREGEO 2364

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Ore Geology Reviews

Received Date: Revised Date: Accepted Date:

1 June 2017 3 October 2017 9 October 2017

Please cite this article as: M. Lesbros-Piat-Desvial, G. Beaudoin, J. Mercadier, R. Creaser, Age and origin of uranium mineralization in the Camie River deposit (Otish Basin, Québec, Canada), Ore Geology Reviews (2017), doi: https:// doi.org/10.1016/j.oregeorev.2017.10.006

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Age and origin of uranium mineralization in the Camie River deposit (Otish Basin, Québec, Canada) M. Lesbros-Piat-Desvial, G. Beaudoin, J. Mercadier and R. Creaser

Marion Lesbros-Piat-Desvial Département de géologie et de génie géologique, 1065 avenue de la Médecine, Université Laval, Québec, QC, Canada G1V 0A6 [email protected]

Georges Beaudoin Département de géologie et de génie géologique, 1065 avenue de la Médecine, Université Laval, Québec, QC, Canada G1V 0A6 [email protected] Julien Mercadier Université de Lorraine, CNRS, CREGU, GeoRessources, Faculté des Sciences et Technologies, rue Jacques Callot, BP 70239, Vandœuvre-Lès-Nancy, 54506, France [email protected] Robert Creaser Department of Earth and Atmospheric Sciences, 1-26 Earth Sciences Building, University of Alberta, Edmonton, AB, Canada T6G 2E3 [email protected]

Corresponding author: Georges Beaudoin ([email protected])

Oct. 3, 2017

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Abstract The Camie River uranium deposit is located in the southeastern part of the Paleoproterozoic Otish Basin (Québec). The uranium mineralization consists of disseminated and vein uraninite and brannerite precipitated close to the unconformity between Paleoproterozoic fluviatile, pervasively altered, sandstones and conglomerates of the Matoush Formation and the underlying sulfide-bearing graphitic schists of the Archean Hippocampe greenstone belt. Diagenetic orange/pink feldspathic alteration of the Matoush Formation consists of authigenic albite cement partly replaced by later orthoclase cement, with the Na2O content of clastic rocks increasing with depth. Basin-wide green muscovite alteration affected both the Matoush Formation and the top of the basement Tichegami Group. Uraninite with minor brannerite is mainly hosted by subvertical reverse faults in basement graphitic metapelites ± sulfides and overlying sandstones and conglomerates. Uranium mineralization is associated with chlorite veins and alteration with temperatures near 320 ºC, that are paragenetically late relative to the diagenetic feldspathic and muscovite alterations. Re-Os geochronology of molybdenite intergrown with uraninite yields an age of 1724.0 ± 4.9 Ma, whereas uraninite yields an identical, although slightly discordant, 1724 ± 29 Ma SIMS U-Pb age. Uraninite has high concentrations in REE with flat REE spectra resembling those of uraninite formed from metamorphic fluids, rather than the bell-shaped patterns typical of unconformity-related uraninite. Paragenesis and geochronology therefore show that the uranium mineralization formed approximately 440 million years after intrusion of the Otish Gabbro dykes and sills at ~2176 Ma, which constrains the minimum age for the sedimentary host rocks. The post-diagenetic stage of uraninite after feldspathic and muscovite alterations, the paragenetic sequence and the brannerite-uraninite assemblage, the relatively high

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temperature for the mineralizing event (~320 °C) following the diagenetic Na- and K-dominated alteration, lack of evidence for brines typical of unconformity-related U deposits, the older age of the Otish Basin compared to worldwide basins hosting unconformity-related uranium deposits, the large age difference between basin fill and mineralization, the older age of the uranium oxide compared to ages for worldwide unconformity-related U deposits, and the flat REE spectra of uraninite do not support the previous interpretation that the Camie River deposit is an unconformity-associated uranium deposit. Rather, the evidence is more consistent with a PaleoProterozoic, higher-temperature hydrothermal event at 1724 Ma, whose origin remains speculative.

Introduction The Camie River uranium deposit is located in north-central Québec, Canada, approximately 250 km northeast of Chibougamau, in the Paleoproterozoic Otish Basin (Fig. 1). It was discovered by Uranerz Exploration and Mining Limited Company during the 1970’s and 1980’s (Gatzweiler, 1987; Höhndorf et al., 1987). The Camie River deposit was discovered along with more than 30 other uranium showings in the Otish Basin and underlying basement, (Fig. 1), representing a variety of uranium deposit types. Gatzweiler (1987) and Höhndorf et al. (1987) grouped the Otish uranium deposits into four types according to their geological context (Fig. 2): (1) stratiform mineralization in basement units (e.g., Takwa River mineralized boulders), (2) vein mineralization in basement units (e.g., Beaver Lake and Lorenz Gully), (3) vein mineralization near the unconformity (e.g., Camie River), and (4) vein mineralization in Otish Supergroup sediments spatially associated with mafic dykes and faults (e.g., Matoush and Indicator Lake). The Matoush deposit has indicated and inferred resources estimated at 0.586 Mt at 0.95 wt.% U3O8 and 1.686 Mt at 0.44 wt.% U3O8, respectively (Ross, 2012). Based on the geology of the 3

basement, the controls of mineralization, the polymetallic association with U, and the ages obtained on uranium oxides (Höhndorf et al., 1987), the various types of U mineralization in the Otish Basin were interpreted as unconformity-associated uranium deposits (Gatzweiler, 1987; Höhndorf et al., 1987; Beyer et al., 2012). The Camie River uranium mineralization (up to 2.95 wt.% U3O8 over 7.35 m; Aubin et al., 2012a; 2012b) occurs near the unconformity between Paleoproterozoic fluviatile basin sedimentary rocks of the Otish Supergroup and Archean basement rocks of the Tichegami Group. Gatzweiler (1987) identified a zoned alteration halo composed of an inner Fe-Mg chlorite and carbonate zone and an external zone comprising albite, chlorite and pyrite in veins and as disseminations. Höhndorf et al. (1987) and Beyer et al. (2012) showed that the uranium mineralization formed at 1723±16 and 1721±21 Ma (207Pb/206Pb age), respectively, and argued that intrusion of the apparently coeval Otish Gabbro suite caused basinal brine movement and uranium deposition in conditions similar to those known for unconformityassociated U deposits in the Athabasca and Thelon basins (Canada) and McArthur Basin (Australia). A similar, but poorly constrained age of 1695 ± 110 Ma (U-Pb on uranium oxides; Alexandre et al., 2015) was proposed for the formation of the Matoush U deposit, for which the spatial association with a mafic dyke, the Matoush dyke, with an unusual Cr-rich signature, was used to support an unique model of formation different from other uranium mineralization in the Otish Basin and from the unconformity-associated U systems. The Otish Gabbro suite and of their sedimentary hostrocks, however, has been shown to be much older than previously assumed, with U-Pb ages from 2172-2162 Ma (Hamilton and Buchan, 2016; Milidragovic et al., 2016). In the absence of an age for the Matoush dyke, this indicates that uranium was precipitated nearly 450 Ma after intrusion of the Otish Gabbro suite, such that the interpretation of Camie River as a diagenetic, unconformity-associated, uranium deposit driven by the intrusion of the Otish Gabbro suite must be re-evaluated.

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In this paper, we reassess the genetic model for the Camie River U deposit by constraining the spatial and timing relationships between diagenetic and hydrothermal alterations, and uranium mineralization, based on detailed paragenetic sequence and mineral chemistry of hydrothermal alteration and uranium mineralization. Lithogeochemistry of altered sedimentary rocks and mineralized basin and basement rocks is used to estimate the mass balance of major elements during hydrothermal alteration, and to identify elements specifically associated with uranium mineralization. Sulfur and lead isotope geochemistry of pyrite and galena is used to infer formation temperatures and sources. The REE contents of uraninite are compared to those from other deposit types. Re-Os and U-Pb geochronology of molybdenite and uraninite is used to reassess the age of the uranium mineralization at the Camie River deposit. A new genetic interpretation for the Camie River deposit is then discussed. 1. Regional geology The Camie River deposit is hosted by the Paleoproterozoic Otish Basin, located near the southeastern margin of the Archean Superior Province, north of the Grenville Tectonic Front (Fig. 1). The Otish Basin represents the continental portion of the Otish-Mistassini depositional system, which formed an epicontinental sea encroachment on the margin of the Superior Province (Genest, 1989). The current 150 by 50 km current basin, showing a NE-trending elongation (Fig. 1; Gatzweiler, 1987; Chown and Caty, 1973), hosts the Otish Supergroup (Fig. 1), which records, from the NE to the SW, depositional facies ranging from high-energy fluviatile, syn-rift facies, to lower-energy, marginal, marine depositional environments, as a result of a marine transgression. The maximum age for the Otish Basin is provided by unconformably overlain NW- to NNW-trending Mistassini diabase dykes swarm, cutting basement rocks, with a baddeleyite U-Pb age of 2515 ± 3 Ma (Hamilton, 2009). The minimum age of the Otish 5

Supergroup is constrained by 2172-2162 Ma baddeleyite U-Pb ages for Otish Gabbro dykes and sills that intruded the Otish Supergroup (Fig. 1; Hamilton and Buchan, 2016; Milidragovic et al., 2016). 2.1 Otish Basin basement and sedimentary rocks The Otish Supergroup unconformably overlies an Archean basement showing a locallydeveloped regolith (Chown and Caty, 1983). The basement consists of the high-grade metamorphic Epervanche Complex and the EW-trending foliated metavolcano-sedimentary Tichegami Group, both cut by younger granitic intrusions (Fig. 1; Chown, 1984). Genest (1989) divided the Otish Supergroup into the predominantly reduced, clastic fluviatile sedimentary rocks of the Indicator Group and the conformably overlying oxidized clastic deltaic sedimentary rocks of the Peribonca Group. The Indicator Group is 225 to 1,200 m thick, and consists of the basal Matoush and summital Shikapio formations. The Peribonca Group crops out in the northern part of the Otish Basin and shows a minimum thickness of 1,200 m. It consists of the Laparre, Gaschet and Marie-Victorin formations. Genest (1989) suggested that the Peribonca Group was formed in a coastal sabkha depositional environment. 2.2 Intrusive rocks The Otish Gabbro suite is comprised of three sills cropping out mainly in the northern part of the basin, as well as variably NE to NW-trending dykes cutting basement rocks (Fig. 1). The sills have a shallow dip to the south, show rhythmic layering and cumulate textures, and terminate along NE-trending fedeer dykes (Milidragovic et al., 2016). The sills and dykes show thicknesses from <300 m to >500 m and from 30 to 200 m, respectively. Both sills and dykes are plagioclase and clinopyroxene-dominated, show variable grain size and texture, are weakly to strongly

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altered, and are metamorphosed to the greenschist facies (Milidragovic et al., 2016). Sills and dykes were emplaced as a result of a single magmatic pulse ca. 2172-2162 Ma (Hamilton and Buchan, 2016; Milidragovic et al., 2016). The Otish Gabbro suite formed from an olivinetholeiitic mantle-derived parental magma modified by crustal assimilation (Milidragovic et al., 2016). The Matoush dyke, hosting the Matoush deposit, has a fractionated trace element composition and LREE-enriched composition with a lamprophyric affinity, unrelated to the Otish Gabbro suite (Miligdragovic et al., 2016). Alexandre et al. (2015) described a mafic component of magnetite and biotite porphyric crystals in a plagioclase groundmass cut by a felsic and pegmatitic facies. The dyke is strongly altered to sericite and chlorite where it hosts the uranium mineralization of the Matoush deposit. The age of the Matoush dyke is unknown but it must be older than the Matoush deposit discordant uraninite U-Pb age of 1695±110 Ma (Alexandre et al., 2015). 2.3 Deformation during the Grenville Orogeny At the end of the Mesoproterozoic, the Otish Basin and proximal basement rocks were faulted during the Grenville Orogeny (1090-980 Ma; Hynes and Rivers, 2010). Chown (1979) suggested that the basin was sliced into blocks by a N050°-N060° discontinuous southward-dipping thrust and by subvertical N060°, dextral and, later N010°, sinistral strike-slip faults. The northwestern part of the basin has undergone brittle deformation and sub-greenschist grade metamorphism limited to fault zones, whereas the southeastern part of the basin underwent brittle-ductile deformation and increasing regional metamorphic grade from the greenschist to the amphibolite facies close to the Grenville Front (Chown, 1979; 1984). Such events likely favored fluid flow,

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diagenetic alteration, and U mobilization in the basin, as suggested by younger ages, such as the 1663-1409 Ma muscovite Ar-Ar age (Beyer et al., 2012), and a ca. 1070 Ma age for U mineralization (Höhndorf et al., 1987).

2. Camie River deposit geology The Camie River deposit is located along the southwestern rim of the Otish Basin (Figs. 1 and 3). Basement rocks comprise EW-trending and subvertical chlorite-altered metavolcanic and metavolcaniclastic rocks of the Hippocampe Belt (Tichegami Group; Genest, 1987). The southern rim of the Otish Basin is affected by discontinuous, EW-striking, steeply south-dipping reverse faults cut by younger NE-trending normal faults (Gatzweiler, 1987; Höhndorf et al., 1987). These two separate fault systems offset the regional unconformity and uplifted basements wedges into the sandstones, leading to a “staircase” pattern. These reverse and normal faults are commonly localized in or along margins of basement graphitic metapelite units and appear to be parallel to the subvertical foliation, with steep dips to the north or to south (Gatzweiler, 1987). These structures also extend into the sedimentary cover which comprises the Matoush Formation and the B member of the Laparre Formation. Adding to this, several NS-trending faults also appear to displace the basin edge. The Matoush A-member is composed of massive, polymictic conglomerate, progressively overlain by quartz-pebble monomictic conglomerate which is interbedded with conglomeratic sandstones towards the top of the unit. This polymictic conglomerate is composed of quartzpebbles and lithic fragments of mafic composition. Polymictic and monomictic conglomerates are generally matrix supported and poorly sorted. The Matoush B-member, overlying the

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conglomeratic sandstones of the Matoush A-member, consists of massive sandstones with some interbedded conglomeratic sandstones. The Matoush Formation is overlain by the conglomeratic sandstones and sandstones of the Laparre Formation B-member that are quartzarenitic to arkosic, moderately well-sorted and rounded, and have cross- and parallel-laminations. At the Camie River prospect, quartz is the predominant detrital phase. Beyer et al. (2012) estimated that feldspar grains constitute up to 25% of the detrital fragments in the Indicator Group (Fig. 4a). Minor detrital minerals are rutile, zircon, pyrite; and rare muscovite, apatite and monazite (Figs. 4b, 4c). The Camie River uranium mineralization is located near the fault-offset unconformity between the Matoush Formation and the basement rocks of the Hippocampe Belt. Uranium mineralization is mainly hosted by subvertical reverse faults cutting graphitic (± sulfides) metapelites (Gatzweiler, 1987; Beyer et al., 2012), and overlying fluviatile sandstones and conglomerates, near basement wedges (Genest, 1987). These post-Otish Supergroup reverse faults are interpreted to be reactivated Archean faults (Genest, 1987). According to Gatzweiler (1987) and Höhndorf et al. (1987), the main mineralized zone is a 550 m long, EW-trending body plunging to the east and extending locally from the unconformity down 20-30 m into the basement rocks and up to 30-50 m upwards along fractures into the basin sedimentary rocks. Uranium grades include 1.71 wt.% U3O8 over 10.75 m, including 25.9 wt.% U3O8 in diamond drill hole (DDH) OTS-15, and 2.95 wt.% U3O8 over 7.35 m, including 13.2 wt.% U3O8 in DDH CAM-11 (Aubin et al., 2012a; 2012b). The uranium mineralization mainly consists of uraninite (colloform pitchblende) and brannerite (Ruzicka and Le Cheminant, 1983; Gatzweiler, 1987; Höhndorf et al., 1987) with a brannerite/uraninite ratio of 1/10 (Beyer et al., 2012). Uraninite has been dated at 1 723 ± 16 Ma by U/Pb (Höhndorf et al., 1987) and ay1 721 ± 20 Ma by the Pb-Pb

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method (Beyer et al., 2012). Muscovite yielded variable Ar-Ar ages ranging from 1663 to 1409 Ma, the younger age being considered representative of the age of a thermal or fluid event that reset the muscovite K-Ar isotopic system (Beyer et al., 2012). According to Gatzweiler (1987) and Höhndorf et al. (1987), the uranium mineralization is polymetallic with Mo, Cu, Co, Ni, As, Se, Nb, V, Ag, and Au (± Th) and coincides with a mushroom-shaped, and zoned, alteration halo, composed of an inner Fe-Mg chlorite, and Fe-dolomite zone, and an outer albite, chlorite and pyrite zone. The location of the deposit at the interface between basin and basement, structural control by reverse faults, the age of the U mineralization and related alteration, the O and H isotope composition calculated for the mineralizing fluids, and the polymetallic nature of the mineralization, have led to the classification of the Camie River deposit as an unconformityassociated uranium deposit (Gatzweiler, 1987; Höhndorf et al., 1987; Beyer et al., 2012).

3. Materials and analytical methods The study is based on the description of fifteen diamond drill holes (DDH), both mineralized (n=6) and non-mineralized (n=9). Fourteen of the DDH are located in the Camie River deposit, along six N-S sections spaced over 1.3 km from east to west (Fig. 3). The fifteenth DDH (OTS01) is located approximately 18 km NW of the Camie River deposit and is used as a regional reference located far from known uranium mineralization (Fig. 1). The petrographic study is based on 149 polished thin sections, using optical and electron microscopy using a JEOL JSM840A Scanning Electron Microscope (SEM) in backscattered electron (BSE) and energy dispersive spectroscopy (EDS) modes, with a 15 kV accelerating voltage. Electron Probe MicroAnalyses (EPMA) were conducted using the 5-WDS CAMECA SX-100 instrument at Université Laval. Analytical conditions were a 5 µm diameter electron beam, a 15 kV accelerating voltage

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and a 20 nA current. Calibration was achieved using natural and synthetic standards, and data reduction used Cameca PAP software. The whole rock chemical composition of 59 samples from basin sedimentary rocks (n=53) and basement rocks (n=6) was measured after pulverization in an agate motar followed by total digestion using a mixture of concentrated HF:HNO3:HCIO4, and by analysis by ICP-OES for major and minor elements, and by ICP-MS for trace elements (SRC Geoanalytical Laboratories, Saskatoon, Canada). FeO concentrations were measured by titration and, C and S using a Leco induction furnace, and B by Na2O/NaCO3 fusion followed by ICP-OES (SRC Geoanalytical Laboratory, Saskatoon, Canada). The Rare Earth Elements (REE) and yttrium concentrations in uranium oxides were measured at GeoRessources (Université de Lorraine, Nancy, France) using a laser ablation-inductively coupled plasma mass spectrometer (LA-ICP-MS) system composed of a GeoLas excimer laser (ArF, 193 nm, Microlas) coupled to an Agilent 7500c quadrupole ICP-MS. The LA-ICP-MS system was optimized to have the highest sensitivity for all elements (from 7Li to 238U), THO/Th ratio < 0.5% and Th/U ratio of ~1. Samples were ablated with laser spot sizes of 24 µm. The least-altered zones were selected based on SEM images and EPMA results, and analyzed to obtain the primary rare earth element concentrations. The REE analyses by LA-ICP-MS were done in the same location as the U-Pb dating by SIMS. A fluence of ~ 7.5 J.cm-2 and a repetition rate of 5 Hz were used. The carrier gas used was helium (0.5 l/min) which was mixed to argon (0.5 l/min) gas before entering the ICP-MS. The ICP-MS settings were the following: ICP RF Power at 1550 W, Cooling gas (Ar) at 15 l/min, auxiliary gas (Ar) at 0.96 l/min and dual detector mode was used. For each analysis, acquisition time was 30 s for background, 30 s for external standards using NIST 610 and NITS 612 silicate glasses for concentrations (Pearce and al., 1997)

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and in-house UO2 standard Mistamisk for REE and Y (Lach et al., 2013), and 30 s for uranium oxide minerals. The external standard was NIST610 and

238

U was mainly used as internal

standard, as described in Lach et al. (2013). NIST612 and Mistamisk uranium oxides were analyzed and considered as cross-calibration samples to control the quality of the analyses (precision, accuracy, repeatability), as described in Lach et al. (2013). U contents in uranium oxides from Camie River were measured before LA-ICP-MS analyses using an electronic microprobe (Cameca SX-100, GéoRessources, Université de Lorraine, Nancy, France). The U content is relatively constant between the analyzed zones (variation < 5 wt.%,) and the same concentration of 66.1 wt. % U was consequently used for all LA-ICP-MS analyses. Acquisition times were the following: 0.01 s for all elements except U (0.005 s). Total cycle time was 220 ms. Data treatment was done using “Iolite” (Paton et al., 2011), following Longerich et al. (1996) for data reduction. The U–Pb isotopic composition of uranium oxides were determined using a CAMECA ims 1280-HR Secondary Ion Mass Spectrometer (SIMS) at CRPG-CNRS (Nancy, France). The O− primary ion beam was accelerated at 13 kV, with an intensity ranging between 3.5 and 5 nA. The primary beam was set in Gaussian mode with a raster of 10 µm. The size of the spot on the uranium oxides was ~ 15 µm. Positive secondary ions were extracted with a 10 kV potential, and the spectrometer slits were set for a mass resolving power of ~6,000 to separate isobaric interferences of rare earth element (REE) dioxides from Pb. The field aperture was set to 2,000 µm, and the transfer optic magnification was adjusted to 80. Rectangular lenses were activated in the secondary ion optics to increase the transmission at high mass resolution. The energy window was opened at 30 eV, and centred on the low energy side, 5eV before the maximum value. Ions were measured by peak jumping in monocollection mode using the axial Faraday cup (FC) for

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238

U and

238

UO and the axial electron multiplier (EM) for

204

Pb,

206

Pb,

207

Pb,

208

Pb, and

248

ThO.

Each analysis consisted of 8 successive cycles. Each cycle began with measurement of the mass 203.5 and 203.6 for backgrounds of the FC and the EM respectively, followed by 207

Pb, 208Pb, 238U, 248ThO, and

238

204

Pb,

206

Pb,

UO, with measurement times of 4, 4, 10, 6, 20, 4, 4, 3, and 3 s,

respectively (waiting time of 1 s). The beam centering, mass, and energy calibrations were checked before each measurement, after a 60 s presputtering by rastering the primary beam over a 30×30 µm area to clean the gold coating and avoid pollution. Several spot analyses (at least five) were measured on the Zambia reference uraninite (concordant age of 540 ± 4 Ma; Cathelineau et al., 1990) before and after each sample for sample bracketing. To define the relative sensitivity factor for Pb and U used for samples, an empirical linear relationship was defined between UO+/U+ and Pb+/U+ from all the measurements performed on the reference mineral (Zambia). The error on the calibration curve is reported in the error given for each analysis. To achieve good reproducibility, each analysis was preceded by automated centering of the sample spot image in the field aperture and contrast aperture (Schuhmacher et al. 2004) and of the magnetic field values in scanning the 206Pb peak. The high 204Pb/206Pb ratio (> 0.0005 and up to 0.002) obtained for all analyses suggests that there is a significant amount of common lead. Such high common lead contribution was previously detected for other uranium oxides from Camie River (Beyer et al., 2012). A correction for common lead was consequently made for each analytical spot by precisely measuring the calculated at the

207

204

Pb amount and the common lead composition was

Pb/206Pb measured age, using the Pb isotopic composition calculated from

Stacey and Kramers (1975) model at the age of uranium oxide (1723 Ma). Ages and error correlations were calculated using the ISOPLOT flowsheet of Ludwig (1999). Uncertainties in the ages are reported at the 1σ level.

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Molybdenite geochronology was performed at the University of Alberta and used the Re-Os Negative Thermal Ionization Mass Spectrometry (NTIMS) method described in Selby and Creaser (2004) and Markey et al. (2007). Molybdenite was purified from a mineralized graphitic metapelite sample after crushing followed by gravity and magnetic concentration. The 187Re and 187

Os concentrations in molybdenite were determined by isotope dilution mass spectrometry,

using the Carius-tube, solvent extraction, anion chromatography and NTIMS techniques. A mixed double spike containing known amounts of isotopically enriched

185

Re,

190

Os, and

188

Os

analysis was used. Isotopic analysis was made using a ThermoScientific Triton mass spectrometer by Faraday collector. Total procedural blanks for Re and Os were less than <3 pg and 2 pg, respectively, which are insignificant for the Re and Os concentrations in molybdenite. The Chinese molybdenite powder HLP-5 (Markey et al., 1998), was analyzed as a standard. For this control sample over a period of two years, an average Re-Os date of 221.56 ± 0.40 Ma (1SD uncertainty, n=10) is obtained. This Re-Os age date is identical to that of 221.0 ± 1.0 Ma reported by Markey et al. (1998). The age uncertainty is quoted at 2σ level, and includes all known analytical uncertainty, including uncertainty in the decay constant of 187Re.

4. Results 4.1 Hydrothermal alteration and uranium mineralization in basements rocks At the Camie River deposit, the Otish Supergroup and the underlying Tichegami Group rocks have been affected by several stages of hydrothermal alteration and are cut by several types of veins (Fig. 5). Mafic to intermediate volcanic rocks, intermediate to felsic tuffs and beds of graphitic schists of the Hippocampe Belt are moderately to strongly chloritized (Chl1) with disseminated epidote (Ep1; Figs. 5 and 6a). The Fe-rich Chl1 chlorite is parallel to foliation and

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considered to be metamorphic in origin. Fe-rich basement chlorite (Table 1) forms a tight cluster in the ternary Al-Fe-Mg diagram (Fig. 7a), and plot in the ripidolite field with the exception of one analysis which has the composition of pycnochlorite (Lesbros-Piat-Desvial, 2014). Using the chlorite geothermometer proposed by Cathelineau (1988), Chl1 yields an average formation temperature of 357 ± 23°C. The Chl4 chlorite (Fig. 7a), present in mineralized basement rocks, plots at the intersection of the sheridanite, ripidolite, clinochlore and pycnochlorite fields (Table 1; Lesbros-Piat-Desvial, 2014). The temperature of formation is estimated at 317 ± 37°C using Cathelineau (1988). Camie River chloritized basement rocks also contain coarse-grained pyrite (Py1) and pyrrhotite (Po1) with minor galena (Gn1), and are cut by numerous, centimeter-scale, quartz (Qz1) and calcite (Cal1) veins parallel to foliation, with disseminated pyrite (Py2) and pyrrhotite (Po2; Fig. 5). The quartz (Qz1) and calcite (Cal1) veins are cut by later calcite (Cal2) and/or epidote (Ep2) veins containing rare disseminated pyrite (Py3) and pyrrhotite (Po3; Fig. 5). In graphitic schists, the quartz (Qz1) and calcite (Cal1 and Cal2) veins are less abundant, and epidote veins (Ep2) are absent. As Beyer et al. (2012) noted, after basement erosion and deposition of the Otish Supergroup, pervasive coarse- to fine-grained muscovite alteration (Fig. 5) affected the uppermost several meters of the Tichegami Group, in which the altered basement rocks are bleached. The intensity of bleaching is strongest at the unconformity and decreases rapidly with depth. The muscovite alteration is composed of coarse- to fine-grained muscovite, likely intergrown with fine-grained illite as determined by XRD analyses and ASD TerraSpec Short-Wave Infrared (SWIR) relectance spectroscopy (S. Rogers, personal communication, 2011). Muscovite from the pervasive muscovite alteration in basement rocks plots into the phengitic field (Table 2; Fig. 7b).

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The pervasive muscovite alteration of basement rocks is cut by quartz veins (Qz2) followed by chlorite and epidote veins (Chl2 and Ep3, respectively; Fig. 5). Chl2 chlorite and Ep3 epidote veins are then cut by rare ankerite veins (Ank), themselves being cut by later calcite veins (Cal3; Fig. 5). Pyrite (Py4), pyrrhotite (Po4) and minor chalcopyrite fill fractures in earlier calcite (Cal3) and epidote veins (Ep3; Fig. 5). In one occurrence, a fourth generation of calcite vein (Cal4) is surrounded by an internal albite (Ab) and external chlorite (Chl3) alteration halo (Lesbros-PiatDesvial, 2014). In basement rocks, uranium mineralization comprises a first generation of disseminated uraninite (Urn1), rich in lead (6 to 19 wt.% PbO), that is replaced by brannerite ((U,Ca,Ce)(Ti,Fe)2O6), that is low in cerium and iron, but rich in uranium (average UO2/TiO2 ratio of 1.66), lead, silicon, and niobium (Table 3; Figs. 5 and 6b-c-e-f). The Urn1 uraninite commonly shows a mottled and/or replacement texture (Figs. 6b-c-e). Brannerite contains abundant inclusions of Urn1 and few inclusions of galena (Gn1; Fig. 6d). It is possible that the high U-content of brannerite could be caused by sub-microscopic inclusions of Urn1 although care was taken to avoid these inclusions during EPMA (Figs. 6c-f). Urn1 is also replaced by an unidentified uraniferous phosphate mineral (Figs. 5 and 6e) that contains Si, Ca, Pb, and Th as determined by SEM-EDS (Lesbros-Piat-Desvial, 2014). This U-phosphate mineral contains rare galena inclusions (Gn1) and is replaced by rare apatite (Ap; Fig. 5). No paragenetic relationship between brannerite and the U-phosphate was determined. Beyer et al. (2012), however, attributed the U-phosphate to a post-mineralization hydrothermal event (i.e., "post-diagenetic alteration"). Disseminated and euhedral uraninite (Urn1) is intergrown with molybdenite, forming irregular aggregates and veins (Figs. 5 and 6f). Molybdenite forms fine-grained (< 10 by 4 µm), euhedral crystals commonly cemented by galena (Gn2; Figs. 5 and 6f). Molybdenite is molded by a second

16

generation of Pb-rich uraninite (Urn2) (Figs. 5 and 6g). Uranium mineralization is cut by carbonate veins (Dol1 and Cal5; < 10 cm; Fig. 5), that are only found in rocks containing uranium mineralization. Galena (Gn2) cements fractures in disseminated pyrite (Py5) and pyrrhotite (Po5) in dolomite (Dol1) and calcite (Cal5) veins (Figs. 5 and 6h). In dolomite veins (Dol1), minor amounts of late coffinite (U(SiO4)1-n(OH)4n; Table 3) replaces and fills fractures in pyrrhotite (Po5) and galena (Gn2; Figs. 5 and 6h). Uranium mineralization and molybdenite are also cut by small pyrite (Py5) and siderite veins, dolomite, and calcite veins (Fig. 5). 4.2 Diagenetic and hydrothermal alteration and uranium mineralization in Otish Basin sedimentary rocks The Otish Basin sedimentary rocks show three diagenetic and hydrothermal alteration zones (Fig. 8): a (1) a lower zone of dark green muscovite alteration in polymictic and quartz-pebble monomictic conglomerates, interbedded with sandstones and conglomeratic sandstones at the base of the stratigraphic sequence, overlain by (2) a zone of interdigitated orange/pink feldspathic and green muscovite alterations in sandstones and conglomeratic sandstones, overlain by (3) an upper zone of pale green muscovite alteration in sandstones and conglomeratic sandstones (Fig. 8). In the Otish Basin, the pervasive feldspathic alteration of sedimentary rocks has been only identified at the Camie River deposit, whereas the pervasive muscovite alteration is widely distributed throughout the basin, such as in the regional DDH OTS-01. Petrography shows the feldspathic cement is found at Camie River from the unconformity up to, at least, 347 m above it. In the interdigitated zone, feldspathic and muscovite alterations show gradual or sharp contacts (Lesbros-Piat-Desvial, 2014). The type of contact appears to be related to the host lithology: gradual contacts are common between feldspathic and muscovite alterations in the same lithology, whereas sharp contacts typically occur at lithological contacts. Commonly, feldspathic

17

alteration in coarser-grained rocks, such as conglomeratic sandstone, is in sharp contact with muscovite alteration in finer-grained rocks, such as sandstone. At Camie River, early diagenetic quartz is rare and forms overgrowths on detrital quartz grains (Fig. 5). The feldspathic and muscovite alterations almost completely replace this early diagenetic quartz cement (Figs. 5 and 9a to c). The orange/pink feldspathic cement (Figs. 9a-c) is composed of two feldspars (Figs. 9d-e): (1) an early albite (Ab1) cement was progressively replaced by (2) a later potassic feldspar (Kf1; Fig. 9e). Albite appears orange/pink due to numerous sub-microscopic Fe oxide inclusions (Ruhlmann et al., 1986; Beyer et al., 2012). The albitic cement has compositions between Ab100 and Ab83, with an average of Ab98 (Table 2; Fig. 7c). The K-feldspar has compositions between Or99 and Or73, with an average of Or96 (Table 2; Fig. 7c). It is likely, however, that the albite and orthoclase analyses with lower proportions of Ab and Or molecules, respectively, represent areas where either small inclusions or intergrowths of the other mineral were excited by the EPMA beam. The feldspathic cement, and rare detrital K-feldspar and muscovite, are replaced by a coarseto fine-grained green muscovite throughout the stratigraphic sequence (Figs. 4, 5 and 9c-f). The muscovite appears to be intergrown with fine-grained illite as determined by XRD and SWIR reflectance spectroscopy (S. Rogers, personal communication, 2011). The muscovite-illite forms pore-filling aggregates of irregular, platy crystals or minute flakes (Duffin et al., 1989), resulting in a "floating-grain" texture for muscovite-altered sandstones (Beyer et al., 2012). In sandstones far from the Camie River deposit (DDH OTS-01; Fig. 1), pervasive muscovite alteration and detrital muscovite show compositions close to the "pure muscovite" end-member (Tappert et al., 2013; Fig. 7b). At the Camie River deposit, pervasive muscovite alteration and detrital muscovite compositions plot in the "muscovite" and "phengite" fields (Fig. 7b). Pervasive muscovite from

18

the lower zone and basement rocks shows a small enrichment in Na2O in comparison to muscovite from the overlying interdigitated and upper zones (Table 2; Fig. 10; Lesbros-PiatDesvial, 2014). Rare muscovite veins cut the pervasive muscovite alteration and have similar compositions to pervasive muscovite, with the exception of a small enrichment in MgO (Fig. 7b). The muscovite alteration is cut by narrow (< 5 mm) chlorite veins (Chl1) associated with moderate brick-red hematite alteration (BRHem1), commonly located near fracture zones (Fig. 5). The chlorite veins (Chl1) plot at the boundary between the clinochlore and penninite fields (Table 1; Lesbros-Piat-Desvial, 2014). The average Chl1 chlorite formation temperature is estimated at 210 ± 20°C using Cathelineau (1988). Rare quartz veins (Qz2; <10 cm) are associated with brick-red hematite alteration (BRHem1; Fig. 5), as well as paragenetically later albite (Ab2; Fig. 5), K-feldspar (Kfs2; Fig. 5) and calcite (Cal1) veins (Fig. 5). However, the brick-red hematite alteration (BRHem1) often occurs isolated from other minor alterations and veins. In sedimentary rocks, uranium mineralization and fractured zones commonly contain narrow dolomite veins, with disseminated, euhedral pyrite, with fractures cemented by galena with minor chalcopyrite and pentlandite. Dolomite in veins is partly replaced by calcite (Cal2; Fig. 5). Fractured zones also contain a spatially associated chlorite alteration (Chl2; Figs. 5 and 7a). Chlorite (Chl2) plots at the boundary between the ripidolite and pycnochlorite fields with the exception of two analyses in the sheridanite and diabanite fields (Table 1; Lesbros-Piat-Desvial, 2014). The average Chl2 chlorite formation temperature is 323 ± 26°C using Cathelineau (1988). In sandstones, the same disseminated, Pb-rich uraninite (Urn1) found in basement rocks is progressively replaced by Nb-rich brannerite (Figs. 5 and 9g), which is partly replaced by rare apatite (Figs. 5 and 9h). The U-phosphate and the younger molybdenite and uraninite (Urn2)

19

observed in mineralized basement rocks are not found in sedimentary rocks. In sedimentary rocks, uranium mineralization is cut by small calcite veins (Cal3), themselves cut by brick-red hematite veins (BRHem2) which are cut by later green clay veins (Fig. 5). The paragenetic sequence ends with limonite in veins and limonite alteration. Limonite weathers the Otish Supergroup rocks over several meters below the current erosion surface, and infiltrates deeper in fractured zones. 4.3 Lithogeochemistry Fifteen

feldspathic

and

twenty-six

muscovite-altered,

non-mineralized,

sandstones/conglomeratic sandstones from the Camie River deposit and five muscovite-altered and four "least-altered", non-mineralized, sandstones/conglomeratic sandstones from DDH OTS01 were analyzed. Among the twenty-six muscovite-altered sandstones from the Camie River prospect, eleven are from the upper alteration zone, ten from the interdigitating alteration zone and five from the lower alteration zone near the unconformity. In DDH OTS-01, five muscovitealtered sandstones are used for comparison with Camie River muscovite-altered sandstones. The DDH OTS-01 "least-altered" sandstones are used as a reference because they show a stronger preservation of the early quartz cement than other altered sandstones, however, they also contain considerable amounts of fine- to coarse-grained muscovite. The DDH OTS-01 "least-altered" sandstones and Camie River feldspathic sandstones have similar lower Al2O3 and K2O, but higher SiO2 contents compared to muscovite-altered sandstones from DDH OTS-01 and Camie River (Fig. 11). Camie River feldspathic sandstones are characterized by the lowest LOI values (Lesbros-Piat-Desvial, 2014) and the highest Na2O contents decreasing from 3.4 wt.%, near the base of the stratigraphic sequence, to lower values of 1.02 wt.%, 126 m above the unconformity (Fig. 11c). Feldspathic sandstones are characterized by

20

lower Al2O3 and K2O, but higher SiO2 contents (Fig. 11). DDH OTS-01 "least-altered" sandstones and muscovite-altered sandstones from DDH OTS-01 and Camie River commonly have higher LOI values (Lesbros-Piat-Desvial, 2014) and lower Na2O contents (Fig. 11c). The Camie River muscovite-altered sandstones from the upper, interdigitating and lower alteration zones show similar major element compositions. Using the isocon method (Grant, 1986), the Camie River feldspathic sandstones show a mass loss of about 75%, higher than those for the muscovite-altered sandstones from the interdigitating (about 40%) and upper (about 30%; alteration zones (Lesbros-Piat-Desvial, 2014). Camie River muscovite-altered sandstones show an increasing mass loss with depth, towards the unconformity (Lesbros-Piat-Desvial, 2014). In comparison to the "least-altered equivalent", the Camie River feldspathic and muscovite altered sandstones are characterized by an enrichment in U and W; a depletion in SiO2, FeOt, Pb, Sr, and B; and insignificant mass changes for MnO, Ag, Bi, Cd and S (Lesbros-Piat-Desvial, 2014). Feldspathic sandstones are also depleted in TiO2 , Sc, Tb, V, Cr and Cs, whereas muscovitealtered sandstones from the three different alteration zones are characterized by an enrichment in TiO2, Sc, V, Cr, Cs, and immobile Tb. Muscovite-altered sandstones from the upper and interdigitating zones are similar to those from the lower alteration zone: sandstones from the upper and interdigitating alteration zones are commonly enriched in Al2O3, MgO, Na2O, K2O, La, Ce, Pr, Nd and Sm (LREE), Be, Co, Ga, Nb, Ni, Rb, Sn and Ta, and show insignificant changes in CaO, C and Th. Sandstones from the lower alteration zone are depleted in Al2O3, MgO, CaO, K2O, LREE (La, Ce, Pr), C, Co, Ga, Rb, Sn and Th, whereas Na2O, LREE (Nd, Sm), Be, Nb, Ni and Ta show no significant mass changes (Lesbros-Piat-Desvial, 2014). Most of the altered sandstones are depleted in REE in comparison to PAAS (Fig. 12). "Leastaltered" sandstones from DDH OTS-01 commonly show flat LREE patterns and depleted HREE

21

patterns (Fig. 12a). The Camie River feldspathic sandstones commonly show flat LREE patterns, but enriched, flat, or depleted HREE patterns (Fig. 12b). The Camie River and DDH OTS-01 muscovite-altered sandstones show patterns similar to DDH OTS-01 "least-altered" sandstones (Figs. 12c and 12d). However, two samples from the Camie River muscovite lower alteration zone show a HREE enrichment compared to other muscovite-altered sandstones. At the Camie River deposit, a majority of feldspathic and muscovite-altered sandstones show moderate negative anomalies in Eu, unlike the "least-altered" sandstones from DDH OTS-01 (Fig. 12). 4.4 Chemical composition of Urn1 uraninite The major and trace element concentrations of Camie River uraninite (Urn1) have been measured by EPMA and LA-ICP-MS (Tables 3 and 4). For REEs, four analyses were carried by LA-ICP-MS on two uraninite crystals, less than 1 cm apart on a polished thin section, from a moderately graphitic metapelite sample from a 30 cm zone grading 3.49 wt.% U and 1.44 wt.% Pb, representative of higher grade uranium mineralization (DDH OTS-04; Cameco Corp. data). Urn1 is characterized by relatively high Pb contents (mean: 14.37 ± 7.19 wt.% PbO with a maximum value of 19 wt.% PbO) and significant concentrations for Si (1.26 ± 1.86 wt.% SiO2), Ca (1.1 ± 0.72 wt.% CaO), Fe (1.23 ± 1.22 wt.% FeO) and Th (1.68 ± 0.88 wt.% ThO). Camie River Urn1 is characterized by high concentrations in REEs and by a flat chondrite-normalized REE patterns, with high REE/chondrite values near 1000 for all the REEs (Fig. 13). In this sample, Urn1 is disseminated and shows a molted texture, and it is partly replaced by brannerite and U-phosphate, itself replaced by apatite. Uraninite Urn1 is cut by veins of pyrite or pyrrhotite. 4.5 Uraninite U-Pb geochronology

22

U-Pb isotope analyses were performed using SIMS on the same two Camie River uraninite (Urn1) crystals analysed by LA-ICP-MS for the REE compositions. 204Pb/206Pb ratios are high, at least two orders of magnitude higher than the standard, indicating a high content in common lead. After correction (see Methods), U-Pb isotope compositions yield

207

Pb/206Pb ages ranging from

1540 to 1704 Ma. The discordant U-Pb upper intercept age is 1724 ± 29 Ma and the lower intercept is at 66±24 Ma, with a MSWD of 1.3 (Table 5, Fig. 14). 4.6 Molybdenite Re-Os geochronology A molybdenite sample was extracted from a mineralized graphitic metapelite interval grading 9.55 wt.% U and 1.37 wt.% Mo. Molybdenite is disseminated, with a typical prismatic habit, and is intergrown with Urn1 uraninite (Fig. 15). Analyses of two fractions from the sample yield similar compositions in Re (58 and 48 ppm) and Os (1055 and 886 ppb, Table 6). The Re-Os model ages are 1725 ± 7 Ma and 1723 ± 7 Ma, indistinguishable within analytical error. A weighted mean age of 1724.0 ± 4.9 Ma is computed for molybdenite, at the 95% level of confidence.

5. Discussion 5.1 Diagenetic and hydrothermal alteration of basement and sedimentary cover rocks at Camie River U deposit The early feldspathic alteration at Camie River occludes porosity in basal clastic sedimentary rocks near the unconformity. The feldspathic alteration is accompanied by an increase in whole rock Na2O content towards the base of the sedimentary rock sequence (Fig. 11c). This suggests that the feldspathic alteration is a diagenetic cement formed by a downward percolating fluid,

23

likely with an evaporated seawater composition (Beyer et al. 2012). The shift from albite to potassium feldspar in the feldspathic alteration records an increase of the potassium activity (aK+ /aH+) as a result of sodium depletion through albite precipitation (Lesbros-Piat-Desvial, 2014). The replacement of the feldspahic alteration by the muscovite alteration likely reflects depletion in alkalis (Na, K), through feldspar precipitation, that led to an increase in the activity of H+, causing the precipitation of muscovite (Lesbros-Piat-Desvial, 2014). Replacement of albite by muscovite in Na2O-enriched sandstone (Fig. 11c) likely explains the high sodium content of muscovite in sandstones close to the unconformity (Fig. 10), as suggested by Beyer et al. (2012). Beyer et al. (2012) attributed the late chlorite and sulfides in basement and sedimentary cover rocks (Fig. 5) to a late diagenetic stage associated with the uranium mineralization. The higher temperature, above 300º C, of that proposed U-bearing, late diagenetic stage was interpreted by Beyer et al. (2012) to be related to intrusion of the Otish Gabbro suite at 1730 Ma, on the basis of a poorly documented Sm/Nd isochron age (Gatzweiler, 1987; Hönhdorf et al., 1987). The baddeyleite U/Pb ages for the Otish Gabbro suite range from 2172 to 2162 Ma (Hamilton and Buchan, 2016; Milidragovic et al., 2016), which makes this interpretation unlikely, because it shows a time gap of nearly 440 Ma between sill intrusion and the 1721 Ma Pb/Pb age for uraninite reported by Beyer et al. (2012). Later stages of chlorite, carbonate, and sulfides in veins and alteration are likely unrelated to diagenesis. The numerous cross-cutting relations (Fig. 5) perhaps indicate a number of discrete, small fluid infiltration events followed by uranium mineralization, and ending with late calcite (in basin) and carbonate-sulfides (in basement) stages, before overprinting by supergene claylimonite alteration (Fig. 5). Post-uranium mineralization, late hydrothermal events are indicated by the 1663-1409 Ma Ar-Ar muscovite ages, which have been ascribed to metamorphic fluid

24

infiltration, followed by late meteoric water ingress after the Grenville orogeny (Beyer et al., 2012). 5.2 Age of Camie River uranium mineralization The Otish Basin rests on an Archean basement cut by the 2515 Ma Mistassini dykes (Hamilton, 2009). The sedimentary rocks of Otish Basin were intruded by mafic dykes and sills of the Otish Gabbro suite between 2162 and 2172 Ma (Hamilton and Buchan, 2016; Milidragovic et al., 2016), constraining the age of the sedimentary host rocks to the Camie River uranium deposit between 2515 and 2162 Ma. This age is far more older than those for the sedimentary basins hosting unconformity-related U deposits worldwide, which have maximum sedimentation ages younger than 1.85 Ga for Athabasca, Thelon and McArthur basins (Kyser and Cuney, 2015; Furnaletto et al., 2016). Höhndorf et al. (1987) reported a U-Pb TIMS 1723±16 Ma upper intercept for uraninite, whereas Beyer et al. (2012) reported a LA-ICP-MS discordant Pb-Pb age of 1721±20 Ma for the uranium mineralization at the Camie River deposit (Fig. 15). These ages are indistinguishable, within error, with our SIMS U-Pb uraninite discordant upper intercept age of 1724±29 Ma (Fig. 14) and

Re-Os molydenite age of 1724.0±4.9 Ma (Table 6). Thus,

geochronology shows that the uranium mineralization formed at ~1724 Ma, at least 440 Ma after deposition of the Otish Group sedimentary rocks. The geochronology is consistent with the late paragenetic stage of the uranium mineralization, which cuts diagenetic and hydrothermal feldspathic and muscovite alteration in veins and faults with late vein and alteration chlorite (Fig. 5). The age of the Camie River deposit uranium mineralization is similar to other geochronological data for uranium in the Otish Basin. Höhndorf et al. (1987) reported TIMS U-

25

Pb ages of 1717±20 Ma and 1711±2 Ma for the Lorenz Gully and L occurrences, respectively (Fig. 1). At the Matoush deposit (Fig. 1), Alexandre et al. (2015) reported an imprecise uraninite LA-ICP-MS U-Pb age of 1695±110 Ma, reset at 1010±25 Ma, which, they suggested, was likely a result of the Grenville Orogeny. At Matoush (Fig. 1), Hörndorf et al. (1987) reported a TIMS U-Pb discordant 1359±28 Ma age, whereas they obtained a “Grenvillian” age (1072±5 Ma) at the Indicator Lake occurrence (Fig. 1). The hydrothermal alteration event that formed the uranium mineralization at ca. 1724 Ma, also likely caused alteration of the Otish Gabbro suite to yield the 1730±10 Ma Sm-Nd isochron age reported by Hörndorf et al. (1987). The evidence thus suggests that a significant part of the U mineralization in the Otish Basin was formed at ca. 1720 Ma, approximately 440 Ma after the intrusion of the Otish Gabbro suite, such that it is unlikely the Otish Gabbro suite contributed in any way to the formation of the uranium mineralization. The relation between dyke emplacement and U mineralization, however, remains a permissive hypothesis for the Matoush U deposit, because the Matoush dyke is undated and has a different origin than the Otish Gabbro suite (Milidragovic et al., 2016), and because of the Cr-rich alteration associated with the U mineralization (Alexandre et al., 2015). 5.3 Origin of the Camie River uranium mineralization The uranium mineralization at Camie River formed almost 440 Ma after deposition of the sedimentary rocks covering the Archean basement. Mineralization therefore formed long after diagenesis of the Otish Basin, after the porosity of the rocks was reduced by early feldspar cementation. This timing is rather different than those known for sedimentary basins hosting unconformity-related U deposits in the Athabasca and Thelon (Canada) and McArhtur (Australia; Jefferson et al., 2007). In these basins, the porosity is considered high at the time of primary mineralization (ca. 1680 Ma and 1590 Ma for McArthur and Athabasca, respectively; Skirrow et

26

al., 2016; Alexandre et al., 2009) with sandstone and conglomerate almost purely made of detrital quartz (Hiatt et al., 2007; Beyer et al., 2011), which could be related to the relatively short period of time (< 150 Ma) between the deposition of the basin and the formation of the U deposits for the Athabasca, Thelon and McArthur basins, compared to the Otish Basin (< 440 Ma). The metamorphic chlorite in basement rocks yields temperatures (~360 ºC) consistent with the greenschist grade rocks. Post-diagenetic chlorite associated with hematisation in basin sedimentary rocks yields temperatures near 210 ºC. The uranium is paragenetically late and associated with chlorite which yields temperatures near 320 ºC both in basement rocks and the sedimentary cover. This suggests an increase in temperature, from ~210 ºC after diagenesis, to ~320 ºC at the time of the uranium mineralization, similar to results reported by Beyer et al. (2012). This temperature is much higher than temperatures typical of unconformity-related uranium deposits, which formed from 100-220 ºC brines, based on fluid inclusion microthermometry (Renac et al., 2002; Derome et al., 2005; Derome et al., 2007; Chu and Chi, 2016; Chi et al., 2017; Richard et al., 2012; Richard et al., 2016) or from 150-250°C brines, based on illite and chlorite geothermometry (Alexandre et al., 2005; Cloutier et al., 2011; Chu and Chi, 2016) or O isotopic equilibrium between coeval hydrothermal minerals (Kotzer and Kyser, 1995). In addition, the dominant chlorite associated with uranium at Camie River has a composition that range from clinochlore to chamosite. Such composition differs from chlorite paragenetically associated with uranium oxides in unconformity-related deposits, which is sudoitic in composition (Hoeve and Quirt, 1984; Kötzer and Kyser, 1995, Billault et al., 2002, Skirrow et al., 2016). Beyer et al. (2012) reported rare, post-Otish Basin uranium, late (“postdiagenetic”) sudoitic chlorite in basement rocks that was not found in our study. From a global perspective, the paragenetic association with the uranium oxides in the Otish Basin differs from what is known for all the unconformity-related U deposits in the Athabasca, McArthur, and 27

Thelon basins, for which U is temporally associated with illite, sudoite, Mg-foitite/dravite tourmaline and alumino-phosphate-sulfate minerals (Kyser and Cuney, 2015; Jefferson et al., 2007). The formation of Mg-rich minerals such as sudoite and Mg-foititic/dravitic tourmaline is considered to directly reflect the high-Mg content of the mineralizing brines derived from the evaporation of seawater (Mercadier et al., 2012). The rare occurrence of these minerals in the Otish Basin does not support infiltration of basinal brines, but this hypothesis cannot be ruled out considering the O and H isotope compositions calculated for the fluids, which suggest mixing between a seawater-derived basinal brine and fluids of metamorphic origin (Beyer et al., 2012). APS are ubiquitous alteration minerals in all known unconformity-related U deposits in both Canada and Australia (Gaboreau et al., 2005; 2007; Adlakha et Hattori, 2015). These minerals are present in the alteration halo directly related to U mineralization but also distally in the sedimentary basins, and this spatial repartition is correlated to a change in chemical composition (Gaboreau et al., 2007). The coeval crystallization of APS, sudoite and Mg foitite/dravite in unconformity-related U systems worldwide is considered to be directly linked to the physicochemical characteristics of the mineralizing fluids, i.e. highly-saline (> 20 wt.% eq. NaCl) basinal brines derived for subaerial evaporation of seawater (Richard et al., 2011; Richard et al., 2014). The composition of chlorite and lack of the typical unconformity-related mineralogical association within the alteration halo associated to the Camie River U deposit, and other U deposits in the Otish Basin, seems to preclude the implication of basinal brines in the genetic processes at the origin of the U mineralization in the Otish Basin. This hypothesis is reinforced by the geochemical signature of alteration at Camie River compared to that typical for unconformity-related U deposits characterized by an enrichment in Na, Si and REEs at Camie River compared to an enrichment in B, Mg, K, and losses in Na and Si (Fig. 12; Polito et al., 2011; Fisher et al., 2013; Skirrow et al., 2016).

28

The high temperature of formation for uranium-related chlorite is consistent with the chemical composition of the uranium oxides, wherein Urn1 is characterized by significant incorporation of thorium at the time of its crystallization (Table 3), which has not been documented for the uranium oxides from unconformity-related U deposits (Alexandre and Kyser, 2005; Alexandre et al., 2015), and which is considered to be directly related to the temperature of crystallization of the uranium oxides (Frimmel et al., 2014). A high temperature of crystallization is also likely based on the REEs concentrations and REE patterns in Urn1 (Table 4 and Fig. 13). The high concentrations of REEs and the flat REE pattern from Camie River are more typical of hightemperature, syn-metamorphic uranium mineralization (Mercadier et al., 2011; Eglinger et al. 2013; Frimmel et al., 2014; Alexandre et al., 2015). Syn-metamorphic uranium deposits (Mistamisk, Lufilian Belt), with REE patterns similar to those of Camie River, formed at temperatures near 300-350º C (Mercadier et al., 2011; Eglinger et al., 2013).

In contrast,

unconformity-related uranium deposits, whatever the host basin, typically show a bell-shape pattern depleted in LREE (Fig. 13, Mercadier et al., 2011; Alexandre et al., 2015). The physicochemical parameters explaining the unconformity-related U deposits REE pattern remain poorly constrained, but it is now well demonstrated that uranium oxides that formed under the geological conditions of unconformity-related U deposit are only characterized by such bell-shape REE pattern. The high concentration in common lead (Pbc> 1 wt.% PbO based on

204

Pb/206Pb ratio,

Table 5) is another argument discriminating Camie River U mineralization from unconformityrelated U deposits where common lead is not detectable in uranium oxides (204Pb/206Pb < 10-6 ; Skirrow et al., 2016). There are few examples of common lead-bearing uranium oxides similar to Camie River Urn1, which are from granite-related vein deposits in the European Hercynian Massifs (Ballouard et al., 2017a; 2017b). These deposits formed from relatively low-salinity fluids (< 6wt.% eq. NaCl) and at temperature between 250-350°C (Ballouard et al., 2017a). Even 29

if the mechanisms of incorporation of the common lead in uranium oxide remains poorly understood, the occurrence of common lead in both Camie River and Hercynian U deposits could indicate that their genetic model bear similarities, and which are different from unconformityrelated U deposits. Therefore, considering the presence of common lead, the U/Th ratios and REEs concentrations of Urn1, the crystallization temperature is inferred to be > 300°C (Frimmel et al., 2014), consistent with chlorite thermometry. Our study does not provide direct constraints on the nature, origin and chemistry of the mineralizing fluids. Beyer et al. (2012), based on the O and H isotopes, suggested the fluids indicated mixing between seawater-derived basinal brines, similar as those invoked in unconformity-associated U deposits, with a high δ18O fluid of metamorphic origin. It is noteworthy that the two chlorite samples analysed by Beyer et al. (2012) have the higher δ18O values that define the metamorphic fluid end-member, whereas the earlier muscovite plot along the mixing trend with the seawater-derived basinal brine composition. This is consistent with the sodium enrichment of sedimentary rocks near the unconformity that suggests evaporative concentration of a seawater-derived basinal brine during early diagenesis. The ~1724 Ma age for mineralization, >440 Ma after deposition of the sedimentary rocks of the Otish Group, the high temperature of chlorite associated to uraninite, the incorporation of thorium and common lead, and the high content in REEs and the REE pattern of Camie River uraninite, however, are unlike that from unconformity-related uranium deposits, as previously suggested by Beyer et al. (2012). Instead, an origin by high temperature (~320 ºC) hydrothermal fluids, long after sedimentation and diagenesis of the Otish Group rocks, perhaps related to a metamorphic event, is a more likely hypothesis for the origin of the Camie River uranium deposit. High temperature (>300°C) conditions for the formation of the Camie River U mineralization is also supported by the

30

paragenetic association of uraninite and brannerite, the crystallization of brannerite being considered to require higher temperatures compared to uranium oxides formed in unconformityrelated systems. Brannerite has not been described with uranium oxides in classical, unconformity-related U deposits (Kyser and Cuney, 2015). The event that caused the ca. 1724 Ma flow of ~300º C hydrothermal fluids and deposition of the Camie River mineralization remains highly speculative. Possible events are the emplacement of the Kivalliq Igneous Suite and Timpton Large Igneous Province (LIP) across Siberia and Laurentia at 1750 Ma (Ernst et al. 2016), and the 1744-1704 Ma Yavapai Orogeny at the southern boundary of the Superior Province (Raharimahefa et al. 2014). At Matoush, the uranium mineralization replaces an altered mafic dyke (Alexandre et al., 2015) that has an alkaline geochemical affinity, akin to lamprophyres and kimberlites, and which is unrelated to the Otish Gabbro suite (Milidragovic et al. 2016). The age of the Matoush dyke is unknown, but it could be an early manifestation of a large-scale mantle-derived magmatic event that heated the crust and set into motion high temperature hydrothermal fluids that formed the Otish Basin uranium mineralization. The 1744-1704 Ma Yavapai Orogeny affected the southern part of the Paleoproterozoic Southern Province, as shown near Sudbury by Raharimahefa et al. (2014). The Yavapai Orogeny resulted in crustal thickening and granitic magmatism under metamorphic pressures of 2.8-4 kbar and temperatures of 540-565º C (Raharimahefa et al., 2014). At peak upper amphibolite grade metamorphism, conditions reached 5.1-7.6 kbar and 580-615º C (Raharimahefa et al., 2014). The Yapavai Orogeny is a major northeast-trending accretionary event of dominantly juvenile crust to the south of Laurentia (Whitmeyer and Karlstrom, 2007) that could have initiated hydrothermal fluid flow that formed the Camie River uranium deposit farther to the north, in the Otish Basin.

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6. Conclusions The Camie River uranium deposit formed at ~1724 Ma, after diagenesis, about 440 Ma after deposition of the Otish Group sedimentary rocks. The mineralization is associated with higher temperature chlorite (~320 ºC), and occurs late in the paragenetic sequence that differs from that typical of unconformity-related U deposits in the Athabasca, Thelon, and McArthur basins, with presence of brannerite and lack of Mg-rich minerals such as sudoite, Mg-foitite/dravite, or alumino-phosphate-sulfate minerals commonly attributed the highly-saline basinal brines forming unconformity-related U systems. The geochemistry of the alteration is also different with Na and Si enrichment in contrast to the Mg, B and K alteration typical for unconformity-related U deposits. The high common Pb, Th and REE contents, and the flat REE pattern of Camie River uranium oxides differs from typical unconformity-related U deposits worldwide. These patterns are similar to that from syn-metamorphic uraninite elsewhere, indicating a higher formation temperature compared to unconformity-related U deposits. The age difference with basin fill and uranium mineralization, the high temperature of chlorite, rare and paragenetically late relative to uranium mineralization, sudoitic chlorite, the lack of Mg and B geochemical anomalies and the REE pattern of uraninite are unlike those of unconformity-related uranium deposits worldwide. Instead, the Camie River mineralization is more likely related to hydrothermal fluid flow, from a distant magmatic or metamorphic event, at about 1724 Ma.

Acknowledgements This research was funded by Cameco Corporation and by the Targeted Geoscience Initiative 4 of the Geological Survey of Canada. We thank Alexandre Aubin, Scott Rogers and Gerard

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Zaluski, for their comments and contributions to field work, and Mike Hamilton for discussions on the Otish Basin geology. We thank David Quirt and Guoxiang Chi for their helpful comments.

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Whitmeyer, S.J., and Karlstrom, K.E. 2007. Tectonic model for the Proterozoic growth of North America. Geosphere, v. 3, p. 220-259. Figure Captions Figure 1 – Geological map of the Otish Basin showing location of selected uranium occurrences and the OTS-01 regional reference drill hole.

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Figure 2 – Schematic cross-section of the Otish Basin with position of known and speculative uranium occurrences (modified from Gatzweiler (1987) and Genest (1989)). Figure 3 – Camie River geological map showing 14 of the 15 studied DDH, located along 6 N-S sections (modified from Cameco Corp.). The 15th drill hole (OTS-01) is located approximately 18 km NW of the Camie River deposit and is used as a regional reference located far from known uranium mineralization (Fig. 1). Figure 4 – Photomicrographs of detrital minerals in the Matoush Formation at Camie River: (a) detrital K-feldspar grains (Det. Kfs) in muscovite-altered sandstone (Ms; CAM.78933 – 104.70 m – CAM-01; plane-polarized transmitted light), (b) detrital muscovite (Det. Ms) in muscovitealtered polymictic conglomerate (CAM.78915 – 433.00 m – OTS-08; plane-polarized transmitted light); and (c) detrital Ce-monazite (Det. Mnz; disseminated small white crystals) and apatite (Det. Ap; light grey) in a feldspar cement (Fsp cmt) interbedded quartz-pebble monomictic conglomerate showing partial fine-grained muscovite alteration (sample CAM. 79000 – 112.50 m – CAM-11; BSE-SEM image). Figure 5 – Paragenetic sequence of minerals associated with metamorphism, diagenesis and hydrothermal alteration at the Camie River deposit. Numbers indicate the different generations of a mineral. Vertical lines into basin and basement paragenesis indicate hydrothermal alteration and mineralization events that took place at the same time in both sedimentary and basement rocks. Figure 6 – Photomicrographs of hydrothermal alteration and uranium mineralization basement rocks at Camie River: (a) Representative chloritized (Chl1) metapelite showing disseminated epidote (Ep1; sample CAM.78979 – 239.20 m – CAM-02; plane-polarized transmitted light) (b) Disseminated uraninite (Urn1) replaced by brannerite (Brt) in pyrite (Py1-4; sample CAM.78961 – 175.35 m – OTS-04; BSE-SEM image); (c) Replacement front between Urn1 uraninite and brannerite, with Urn1 inclusions in brannerite (sample CAM.78961 – 175.35 m – OTS-04; BSESEM image); (d) Galena (Gn1) inclusions in brannerite (sample CAM.78987 – 180.10 m – CAM-09; BSE-SEM image) ; (e) Uraninite (Urn1) replaced by an unidentified U-phosphate (Pb±Th±Ca; U-Ph), itself replaced by apatite (Ap; sample CAM.78958 – 170.35 m – OTS-04; BSE-SEM image) ; (f) Uraninite (Urn1) and brannerite (with uraninite (Urn1) inclusions) cemented by fine laths of molybdenite (Mol) intergrown with Urn1 uraninite and galena (Gn2;

45

sample CAM.78607 – 196.80 m – CAM-11; BSE-SEM image) ; (g) Molybdenite molded by uraninite (Urn2; sample CAM.78958 – 170.35 m – OTS-04; BSE-SEM image); (h) Dolomite (Dol1) vein with fractures filled by pyrrhotite (Po5) and galena (Gn2); coffinite cement (Cof) residual porosity around pyrrhotite grains (Po5) and cleavage plans and fractures in galena (Gn2; sample CAM.78607 – 196.80 m – CAM-11; BSE-SEM image). Figure 7 – (a) Al-Fe-Mg ternary diagram of the different types of chlorite at the Camie River deposit, with reference to pure chamosite, clinochlore (Deer et al., 1992), and sudoite compositions (Lin and Bailey, 1985). Red stars represent the average chlorite compositions of Beyer et al. (2012; Table 1); (b) Fe-Si, Al-Mg ternary diagram of the different types of muscovite at the Camie River deposit and in DDH OTS-01 (modified from Tappert et al., 2013). The red star represents the average composition of Beyer et al. (2012) pervasive muscovite in basement rocks. The grey star represents the average composition of Beyer et al. (2012) pervasive muscovite in basin sedimentary rocks (Table 2); and (c) Na-K-Ca ternary diagram of the different types of feldspar at the Camie River deposit and in DDH OTS-01. The red star represents the average composition of albite cement (n=8) analyzed by Beyer et al. (2012). a = number of rock samples; n = number of analyses. Figure 8 – Schematic cross-section of the distribution of the two principal types of pervasive hydrothermal alteration in sedimentary rocks at the Camie River deposit, with core pictures associated: (a) Lower zone: dark green muscovite alteration of polymictic conglomerate, above the unconformity (DDH OTS-08); (b) Interdigitation zone: orange/pink feldspar and green muscovite alteration of sandstones and conglomeratic sandstones (DDH CAM-09); (c) Upper zone: pale green muscovite alteration of sandstones and conglomeratic sandstones (DDH CAM09). Figure 9 – Photomicrographs of hydrothermal alteration and uranium mineralization in Camie River and OTS-01 sedimentary rocks: (a) Quartz cement (Qz cmt) replaced by feldspathic cement (Fsp cmt) in Camie River feldspar altered sandstone (sample CAM.78703 – 108.20 m – OM-45; plane-polarized transmitted light); (b) Quartz cement replaced by fine-grained muscovite (Ms) in arkosic sandstone 18 km from the Camie River deposit (sample CAM.78679 – 409.40 m – OTS01; plane-polarized transmitted light); (c) Quartz cement replaced by feldspathic cement and finegrained muscovite (sample CAM.78650 – 215.90 m – OTS-15; plane-polarized transmitted

46

light); (d) Albite (Ab1) and K-feldspar (Kfs1) cements in a sandstone after feldspars coloration by HF and cobaltinitrite (sample CAM.78698 – 92.15 m – CAM-07); (e) Albite cement replaced by K-feldspar cement (sample CAM.78649 – 192.50 m – OTS-15; BSE-SEM image); (f) Feldspathic cement replaced by fine-grained muscovite (sample CAM.79000 – 112.40 m – CAM-11; plane-polarized transmitted light); (g) Uraninite (Urn1) cemented by brannerite (Brt) in a strongly chloritized (Chl2) quartz-pebble monomictic conglomerate (sample CAM.78621 – 179.85 m – CAM-13; BSE-SEM image); (h) Brannerite replacement by apatite (Ap), showing galena (Gn) inclusions, in a strongly chloritized (Chl2) quartz-pebble monomictic conglomerate (sample CAM.78621 – 179.85 m – CAM-13; BSE-SEM image). Figure 10 – Diagram of Na2O contents (wt.%) in pervasive muscovite alteration in Camie River and DDH OTS-01 altered sandstones and basement rocks versus distance from the unconformity. Figure 11 – Diagram of (a) Al2O3, (b) K2O, (c) Na2O and (d) SiO2 contents (wt.%) of Camie River and DDH OTS-01 altered sandstones versus distance to the unconformity Figure 12 – PAAS-normalized Rare Earth Elements (REE) patterns of Camie River and DDH OTS-01 altered sandstones: (a) "least-altered" sandstones of DDH OTS-01, (b) feldspathic sandstones of the Camie River deposit, (c) muscovite-altered sandstones of DDH OTS-01, and (d) muscovite-altered sandstones of the three alteration zones of the Camie River deposit. PAAS values from Taylor and McLennan (1985). Figure 13 – Chondrite-normalized Rare Earth Elements (REE) patterns of Camie River uraninite (Urn1), compared to other types of uranium deposits (Mercadier et al., 2011). Chondrite values are from Anders and Grevesse (1989). Figure 14 – Camie River U-Pb uraninite (Urn1) geochronology: (a) BSE-SEM image showing SIMS spots on one of the two uraninite crystals analyzed; (b) U-Pb concordia diagram from SIMS analyses of Camie River uraninite (Urn1), also showing Höhndorf et al. (1987) TIMS and Beyer et al. (2012) LA-ICP-MS isotopic analyses. Figure 15 – Secondary electron image of 1724 ± 4.9 Ma Re-Os molybdenite age intergrown with 1724 ± 29 Ma uraninite (Urn1) SIMS age.

47

Table 1 – Camie River chlorite composition. Basement – Chl1 (a=3; n=13)

Oxides (wt%)

Basin veins – Chl1 (a=2; n=37)



s

Range



s

Range

SiO2

25.260

0.844

24,201-26,975

33.445

0.973

31,893-37,34

TiO2

0,063

0.015

D.L.-0,083

0,266

0.412

D.L.-2,400

Al2O3

21.161

0.957

19,447-22,381

17.711

0.660

16,565-20,787

Cr2O3

0,320

0.187

D.L.-0,512


-

-

MgO

12.082

0.955

10,736-13,753

29.239

1.327

22,739-30,520

CaO

0,133

0.077

D.L.-0,277

0,116

0.057

D.L.-0,313

MnO

n.a.

n.a.

n.a.

0,064

-

D.L.-0,064

FeO

29.517

0.964

28,186-31,014

6.868

0.422

4,855-7,780

NiO

0,078

0.009

D.L.-0,085


-

-

ZnO


-

-


-

-

Na2O

0,027

0.013

D.L.-0,058

0,033

0.008

D.L.-0,042

K2O

0,064

0.108

D.L.-0,312

0.327

0.669

0,024-4,084

H2Ocalculated

11.227

0.107

11,093-11,410

12.707

0.172

12,234-13,070

Total

99.656

98,522-100,703

100.784

Oxides (wt%)

Basin matrix – Chl2 (a=3; n=42)

97,052-104,189

Mineralized basement – Chl4 (a=4; n=18)



s

Range



s

Range

SiO2

27.767

1.130

26.067-32.866

27.840

1.097

26.268-29.811

TiO2

0.068

0.017

D.L.-0.098

0.053

0.008

D.L.-0.065

Al2O3

20.999

0.630

19.469-22.682

21.271

1.080

18.833-23.096

Cr2O3


-

-

0.120

-

D.L.-0.120

MgO

19.774

1.640

17.619-27.584

19.820

3.078

13.093-23.871

CaO

0.037

0.010

D.L.-0.059

0.079

0.075

D.L.-0.227

48

MnO

0.095

0.021

D.L.-0.152

n.a.

n.a.

n.a.

FeO

19.352

2.081

9.369-22.632

17.915

3.207

14.266-23.701

NiO

0.073

-

D.L.-0.073

0.066

-

D.L.-0.066

ZnO

0.213

-

D.L.-0.213

0.285

0.078

D.L.-0.392

Na2O

0.233

0.448

D.L.-1.62

0.026

0.013

D.L.-0.051

K2O

0.163

0.252

D.L.-12.444

0.082

0.146

D.L.-0.427

H2Ocalculated

11.878

0.187

11.512-12.444

11.834

0.189

11.523-12.141

Total

100.121

97.574-101.925

99.036

Oxides (wt%)

95.665-100.718

Mineralized basin matrix – Chl2 (a=1; n=6) 

s

Range

SiO2

28.951

0.685

28.440-30.036

TiO2


-

-

Al2O3

21.809

1.019

20.381-22.649

Cr2O3


-

-

MgO

27.696

0.704

27.046-28.767

CaO

0.027

0.008

D.L.-0..035

MnO

n.a.

n.a.

n.a.

FeO

8.999

0.427

8.418-9.687

NiO


-

-

ZnO


-

-

Na2O

0.014

0.007

D.L.-0.020

K2O

0.004

-

D.L.-0.004

H2Ocalculated

12.460

0.043

12.430-12.543

Total

100.92

99.644-100.549

a = number of samples; n = number of analyses;  = average, based on the number of analyses; s = standard deviation; D.L. = detection limit; n.a. = "not analyzed".

49

Table 2 – Camie River albite and orthoclase, and muscovite compositions. Albite cement (a=13; n=55)

Oxides (wt%)

Orthoclase cement (a=15; n=67)



s

Range



s

Range

SiO2

69.80

4.37

63.50-89.43

64.32

1.11

59.86-66.10

Al2O3

19.37

2.89

7.76-24.70

18.90

0.71

17.93-22.52

FeO

0.43

0.64

0.04-2.33

0.29

0.45

0.02-2.29

MgO

0.09

0.17

0.02-0.62

0.12

0.22

0.01-0.68

MnO

0.09

-

D.L-0.09

0.08

0.02

0.06-0.09

CaO

0.05

0.05

0.02-0.33

0.04

0.04

0.02-0.19

K2O

0.28

0.56

0.03-3.00

15.52

0.95

11.69-16.37

Na2O

10.98

1.51

4.90-12.01

0.41

0.53

0.11-2.92

TiO2


-

-

0.12

-

0.12-0.12

BaO

0.08

0.03

0.05-0.10

0.15

0.08

0.04-0.37

SrO

0.06

0.04

0.03-0.10

0.03

-

0.03-0.03

Total

100.74

97.47-102.63

99.47

97.48-101.27

Basin pervasive muscovite Oxides (wt%)

Basement pervasive muscovite (a=2; n=6) lower alteration zone (a=2; n=14) 

s

Range



s

Range

SiO2

46.71

0.65

46.01-47.74

49.10

2.29

46.02-54.68

TiO2

0.27

0.07

0.20-0.38

0.28

0.10

0.12-0.42

Al2O3

34.10

1.42

32.15-35.84

35.99

3.22

25.28-37.93

FeO

1.83

1.19

0.78-3.68

0.76

0.78

0.33-3.29

MnO


-

-


-

-

MgO

1.20

0.32

0.87-1.66

0.62

0.17

0.39-1.05

CaO

0.09

0.07

0.03-0.16

0.06

0.08

0.02-0.28

Na2O

0.34

0.11

0.24-0.47

0.30

0.09

0.06-0.42

K2O

10.44

0.22

10.15-10.71

9.24

1.29

7.35-11.03

50

BaO

0.61

0.20

0.47-0.76


-

-

Cr2O3


-

-


-

-

NiO


-

-

0.14

-

0.14-0.14

F

0.13

0.08

0.07-0.25


-

-

Cl

0.02

-

0.02-0.02

0.02

0.01

0.02-0.03

H2Ocalculated

4.44

0.06

4.40-4.54

4.61

0.10

4.43-4.75

Total

99.74

99.17-100.25

100.98

99.33-102.77

Basin pervasive muscovite

Basin pervasive muscovite

interdigitation alteration zone (a=8; n=48)

upper alteration zone (a=5; n=22)

Oxides (wt%)



s

Range



s

Range

SiO2

48.86

1.90

45.99-54.55

48.30

2.66

44.60-56.40

TiO2

0.39

0.15

0.17-0.81

0.27

0.12

0.14-0.53

Al2O3

34.38

2.38

28.35-38.47

34.44

4.19

23.35-38.95

FeO

1.05

0.48

0.25-2.19

0.75

0.78

0.19-4.08

MnO


-

-


-

-

MgO

1.26

0.78

0.25-3.00

1.43

1.92

0.29-9.28

CaO

0.05

0.03

0.02-0.16

0.06

0.03

0.03-0.17

Na2O

0.14

0.06

0.05-0.46

0.16

0.06

0.04-0.34

K2O

9.47

1.40

7.36-12.97

9.92

1.80

5.42-14.43

BaO

0.25

0.11

0.10-0.49

0.15

0.04

0.08-0.19

Cr2O3

0.16

-

0.16-0.16

0.09

-

0.09-0.09

NiO

0.07

-

0.07-0.07


-

-

F

0.22

0.04

0.16-0.25


-

-

Cl

0.03

0.03

0.00-0.06

0.03

0.01

0.02-0.04

H2Ocalculated

4.54

0.08

4.36-4.73

4.53

0.11

4.30-4.71

Total

100.33

98.23-102.92

99.94

97.11-102.76

a = number of samples; n = number of analyses;  = average, based on the number of analyses; s = standard deviation; D.L. = detection limit; n.a. = "not analyzed".

51

Table 3 – Camie River uranium minerals composition. Uraninite (Urn1; a=1; n=3)

Oxides (wt%)

Brannerite (a=2; n=7)



s

Range



s

Range

UO2

76.31

1.91

74.91-78.49

43.06

5.62

36.07-53.17

PbO

14.37

7.19

6.08-18.84

4.90

4.15

2.40-14.22

TiO2

0.37

0.13

0.22-0.48

25.97

5.22

16.04-31.00

SiO2

1.26

1.86

0.17-3.40

4.77

1.33

3.41-7.16

CaO

1.10

0.72

0.56-1.92

2.22

0.49

1.74-3.02

Nb2O3

0.87

0.20

0.65-1.00

2.38

0.98

1.75-4.51

FeO

1.23

1.22

0.40-2.64

2.21

0.90

1.66-4.16

ThO2

1.68

0.88

0.66-2.22

0.47

0.15

0.33-0.79

SO2

0.59

0.74

0.16-1.44

0.97

1.16

0.14-3.31

Al2O3


-

-

0.62

0.51

0.05-1.63

P2O5

0.20

0.13

0.10-0.34

0.08

0.03

0.05-0.13

Y2O3

0.59

0.17

0.42-0.75

0.44

0.28

0.06-0.74

ZrO2

0.31

0.30

0.09-0.52

0.38

0.40

0.10-1.20

MgO


-

-

0.24

0.47

0.04-1.30

MnO


-

-

0.17

0.04

0.13-0.22

Cl

0.05

0.002

0.05-0.05

0.16

0.15

0.01-0.36

Ce2O3


-

-


-

-

K2O


-

-


-

-

Total

99.02

97.14-100.28

88.93

Oxides (wt%)

Coffinite (a=1; n=2) 

s

Range

UO2

67.63

3.25

65.33-69.93

PbO


-

-

TiO2


-

-

52

86.58-91.52

SiO2

16.60

1.73

15.38-17.82

CaO

3.27

1.20

2.42-4.12

Nb2O3

0.18

0.04

0.15-0.21

FeO

2.11

1.37

1.14-3.08

ThO2


-

-

SO2

0.22

-

0.22-0.22

Al2O3

0.52

0.24

0.34-0.69

P2O5

2.94

0.22

2.78-3.10

Y2O3

0.56

-

0.56-0.56

ZrO2

0.08

-

0.08-0.08

MgO

0.13

-

0.13-0.13

MnO


-

-

Cl

0.02

-

0.02-0.02

Ce2O3


-

-

K2O


-

-

Total

93.96

93.56-94.36

a = number of samples; n = number of analyses;  = average, based on the number of analyses; s = standard deviation; D.L. = detection limit; n.a. = "not analyzed".

53

Table 4 – Camie River uraninite REE and Y composition. camie_1

±2σ

camie_2

±2σ

camie_3

±2σ

camie_4

±2σ

1466

51

1458

74

1597

82

1328

55

La

197.0

7.7

213.8

8.0

207.4

9.1

213.6

7.2

Ce

372

14

358

13

332

12

324

13

141

Pr

60.4

2.3

54.4

2.0

55.9

2.7

50.7

2.0

146

Nd

336

10

300

11

320

15

289.9

7.9

Sm

162.9

6.9

155.0

6.3

162.8

6.2

154.2

6.1

153

Eu

62.5

2.7

61.7

2.3

66.0

2.4

56.2

1.9

157

Gd

176.3

6.0

165.3

6.1

182.7

7.4

163.1

6.4

159

Tb

39.4

1.4

37.2

1.5

42.5

1.5

36.77

0.94

163

Dy

304

11

280

12

301

13

278

11

165

Ho

63.0

2.6

61.6

3.4

67.6

2.3

57.6

2.1

Er

177.5

7.0

175.0

9.7

191.3

8.5

168.6

5.4

Tm

26.07

0.97

26.8

1.3

28.2

1.0

26.20

0.96

156.1

5.3

164.4

7.7

170.5

7.6

153.8

6.1

13.79

0.55

14.8

1.0

15.86

0.80

14.79

0.65

89

Y

139

140

147

166

169

172

Yb

175

Lu

54

Table 5. Camie River uraninite U-Pb geochronology

SIMS raw ratio spot

207

206

Pb/ Pb

1σ

204

Fractionation and common Pb corrected measured isotopic ratios 206

Pb/ Pb

1σ

204

206

Pb/ Pb

1σ

207

235

Pb/ U

1σ

206

238

Pb/ U

1σ

Correlation error

Age (Ma) 207

206

Pb/ Pb

1σ

207

235

206

Pb/ U 1σ

Pb/238U

207

1σ

Pb/206Pb

1σ

CAMIE-1@2

0.10795

0.00128

0.00069

0.00003

0.00068

0.00003

1.15206

0.03558

0.08472

0.00233

0.8

0.09862

0.01406

778

17

524

14

1596

21

CAMIE-1@3

0.10937

0.00052

0.00063

0.00002

0.00063

0.00002

1.53008

0.03482

0.11005

0.00240

0.8

0.10084

0.00664

943

14

673

14

1639

9

CAMIE-1@4

0.11193

0.00032

0.00059

0.00002

0.00059

0.00002

2.06766

0.03245

0.14434

0.00217

0.8

0.10390

0.00446

1138

11

869

12

1694

5

CAMIE-1@5

0.11083

0.00045

0.00101

0.00004

0.00101

0.00004

1.07409

0.03510

0.08033

0.00252

0.8

0.09698

0.00925

741

17

498

15

1565

7

CAMIE-1@6

0.13279

0.00089

0.00252

0.00014

0.00251

0.00014

1.40779

0.05264

0.10394

0.00265

0.8

0.09823

0.02739

892

22

637

15

1590

12

CAMIE-1@7

0.13278

0.00075

0.00270

0.00006

0.00269

0.00006

0.90045

0.02548

0.06826

0.00167

0.8

0.09568

0.01425

652

14

426

10

1540

10

CAMIE-2A@1

0.11359

0.00119

0.00067

0.00006

0.00067

0.00006

1.71120

0.05732

0.11878

0.00352

0.8

0.10449

0.01556

1013

21

724

20

1704

19

CAMIE-2A@2

0.11550

0.00052

0.00092

0.00003

0.00091

0.00003

1.91422

0.02917

0.13475

0.00183

0.8

0.10303

0.00696

1086

10

815

10

1679

8

CAMIE-2A@3

0.12367

0.00291

0.00146

0.00011

0.00146

0.00011

1.57049

0.10915

0.10972

0.00662

0.8

0.10381

0.03455

959

42

671

38

1692

41

CAMIE-2A@4

0.13121

0.00165

0.00234

0.00012

0.00233

0.00012

0.58896

0.02691

0.04308

0.00154

0.8

0.09916

0.02838

470

17

272

10

1606

22

CAMIE-2B@1

0.11541

0.00223

0.00081

0.00013

0.00081

0.00013

1.77747

0.05779

0.12347

0.00125

0.8

0.10441

0.03089

1037

21

751

7

1703

34

CAMIE-2B@2

0.12299

0.00082

0.00170

0.00008

0.00170

0.00008

0.94969

0.02775

0.06907

0.00165

0.8

0.09972

0.01687

678

14

431

10

1618

12

CAMIE-2B@3

0.10962

0.00126

0.00050

0.00006

0.00050

0.00006

1.51914

0.05762

0.10715

0.00368

0.8

0.10283

0.01607

938

23

656

21

1675

21

CAMIE-2B@4

0.11093

0.00046

0.00052

0.00005

0.00051

0.00005

2.13389

0.03129

0.14887

0.00174

0.8

0.10396

0.00886

1160

10

895

10

1695

8

55

56

Table 6 – Camie River molybdenite Re-Os geochronology Sample

Re (ppm)

± 2s

85780

57.58

85780-2

48.40

187

Re (ppm)

± 2s

0.15

36.19

0.09

0.13

30.42

0.08

187

Os (ppb)

± 2s

Model age (Ma)

± 2s (Ma)

1 055

1

1725

7

886.1

0.6

1723

7

56

Lorenz Gully

N

Indicator Lake

20 km Beaver Lake

Matoush OTS-01 CAMIE RIVER

LEGEND

Superior Province Mistassini Group (Proterozoic)

Grenville Province

Otish Supergroup (Proterozoic)

Albanel Formation

Marie-Victorin Formation

Superior

Gaschet Formation

Quaternary Grenville Front

Proterozoic mafic dikes Fault

Fold axis

Uranium occurences OTS-01

DDH

Basement

Peribonca Group

Archean to Proterozoic

René Group (Archean)

Tichegami Group (Archean)

Béthoulat Complex (Archean)

Inferior

Laparre Formation

Epervanche Complex (Archean)

Cheno Formation

Shikapio Formation

Pambrun Complex (Archean)

Papaskwasati Formation

Matoush Formation

Indicator Group

Bohier Group (Archean)

(modified from SIGEOM, 2012)

Beaver Lake type Lorenz Gully type

A.A. Matoush

Beaver Lake Lorenz Gully

Indicator Lake Gordon Lake CAMIE RIVER Coon showing

A.A. Matoush type

NW

SE

Takwa River A.A. Matoush type

LEGEND Basement (Archean)

Otish Supergroup (Proterozoic)

Fault

Marie-Victorin Formation Gaschet Formation

Unconformity

Laparre Formation

Proterozoic mafic dikes

Shikapio Formation Matoush Formation

Tichegami Group Peribonca Group

Epervanche Complex

900 m

Uranium occurences

Indicator Group 5 km

71°59'N

71°58'N

North

SECTION 1

SECTION 6 OM-52

SECTION 5

SECTION 4 OM-45

CAM-07 51°47'N

OM-58

OTS-04

CAM-13 CAM-01 CAM-02 CAM-09 CAM-11

SECTION 2

OTS-08

SECTION 3 OM-50

51°47'N

OTS-15 OTS-07

LEGEND Studied drill hole with trace Other drill hole Water system Fault Edge of sedimentary Basin Schistosity Laparre Formation Tichegami Group

0

100

200

300

meters

400

500

a Det. Qz Det. Kfs

Ms

Det. Qz 300 μm

b Det. Qz

Det. Ms Det. Qz Ms 600 μm

c

Det. Ap Ms

Det. Qz Det. Mnz

Fsp cmt Fsp cmt Fsp cmt

Ms

Post-Otish Supergroup deposition

Chlorite - Chl Epidote - Ep Tremolite - Tr Pyrite - Py Pyrrhotite - Po Galena - Gn Quartz - Qz Calcite - Cal Muscovite - Ms Ankerite - Ank Chalcopyrite - Ccp Albite - Ab Hematite - Hem Uraninite - Urn Brannerite - Brt U-Phosphate - U-Ph Molybdenite - Mol Apatite - Ap Dolomite - Dol Coffinite - Cof Siderite - Sd

1

2 1

2 2

1

Otish Supergroup deposition

BASIN BASEMENT

Syn-metamorphism Quartz - Qz Albite - Ab K-Feldspar - Kfs Muscovite - Ms Brick-red hematite - BRHem Chlorite - Chl Calcite - Cal Dolomite - Dol Pyrite - Py Galena - Gn Chalcopyrite - Ccp Pentlandite - Pn Uraninite - Urn Brannerite - Brt Apatite - Ap Green Clay - GC Limonite - Lm 1 1

1

2 2

?

2 1

3

2

3

2 3

2 1 1 1

2

1

4

4 4

3 3

5 5 2

1 1

2 2

3

4

5

1

6

7

2

1

2

a

b Qz

Chl Ur Ep Py Brt Ur 300 μm

c

d

Brt Ur

Gn Ur Brt

e

f Mo

Ur

Ur

Gn

U-Ph Brt

g

h Mo

Mo Po Ur

Dol

Cof Gn

Fe

b

Muscovite 3.1 Si, 1.9 AlVI

IV

Mg Si

Fe

Si

Chamosite

VI Al Al

Al VI Al IV

a

K(Al2VI)AlVISi3O10(OH)2 3 Si, 2 AlVI

Phengite Series

Clinochlore

Sudoite

Al

Mg

3.5 Si Mg-number

50% Fe, 1 Fe KFe2Si4O10(OH)2

KAlSi3O8

c

50% Mg, 1 Mg KMg2Si4O10(OH)2

Mineralized host-rock

Non-mineralized host-rock CHLORITE Vein

Matrix BASIN

Chl2

BASEMENT

Chl1

MUSCOVITE

pa Alk ali

fel

ds

BASIN

Anorthoclase

-

Andesine

Labradorite

Bytownite

-

NaAlSi3O8

CaAl2Si2O8

Camie OTS-01

Albite-hematite

Vein

a=15 Camie n=84 a=1 OTS-01 n=6 a=2 Camie n=6

Albite-hematite

Anorthite BASEMENT

Plagioclases

Matrix

Detrital

K-feldspar Albite Oligoclase

Chl4

Mineral

BASIN

a=1 n=6 a=4 n=18

Chl2

-

Detrital

BASEMENT

FELDSPARS

a=2 n=37

Chl1

a=3 Camie n=4 a=1 OTS-01 n=1

rs

Sanidine

a=3 n=42 a=3 n=13

Matrix

-

Camie

-

Cement

Vein

a=13 n=55 a=14 n=66

Camie a=1 n=1 a=1 n=3

a=3 n=8

Camie

Camie

a=4 n=5 a=7 n=21

Camie

a=2 n=3

Camie

-

c

b LEGEND

Otish Supergroup Alteration zones

Overburden Upper pale green muscovite alteration Interdigitation of feldspathic and muscovite alterations Lower dark green muscovite alteration

a

Mineralization Tichegami Group Lithological contact Extrapolated alteration surface Fault Unconformity

0

10

20

30

40

50

60 m

v

a

b

Fsp cmt

Det. Qz

Det. Qz Det. Qz Det. Qz

K-Fsp

Ab-Hem

Det. Qz Det. Qz Qz cmt

Det. Qz

150 μm

Det. Qz

Ms-ill

c

Ab-Hem

d

Ab-Hem

Det. Qz Fsp cmt

Ms-ill Det. Qz Det. Qz

0.25 cm

K-Fsp

e

300 μm

f

Chl

Chl

Gn

Ap

Ur Brt Brt

Na2O (wt%) 0

0.1

0.2

0.3

0.4

0.5

360

Distance to the unconformity (m)

300

240

180

120

60

0

-60

Pervasive muscovite - Basin (OTS-01) Pervasive muscovite - Upper zone (Camie) Pervasive muscovite - Interdigita on zone (Camie) Pervasive muscovite - Lower zone (Camie) Pervasive muscovite - Basement (Camie)

a

Distance to the unconformity (m)

0

5

10

15

20

25

30 500

400

400

300

300

200

200

100

100

0

0

0.5

1

1.5

2

1

2

3

d

Na2O (wt%) 0

K2O (wt%) 0

500

c

Distance to the unconformity (m)

b

Al2O3 (wt%)

2.5

3

500

500

400

400

300

300

200

200

100

100

0

0

5

6

7

8

9 10

SiO2 (wt%) 60

3.5

4

Muscovite-altered sandstone - Upper zone (Camie) Muscovite-altered sandstone - Interdigita on zone (Camie) Muscovite-altered sandstone - Lower zone (Camie)

65

70

75

80

85

90

95

Feldspathic sandstone (Camie) Least-altered sandstone (OTS-01) Muscovite-altered sandstone (OTS-01)

Least-altered sandstones (OTS-01)

a

10

Altered sandstone/PAAS

Altered sandstone/PAAS

10

1

0,1

0,01

La

Ce

c

Pr

Nd

Sm

Eu

Gd

Tb

Dy

Ho

Er

Yb

Muscovite-altered sandstones (OTS-01)

0,1

La

Ce

Pr

Nd

Sm

Eu

Gd

Tb

Dy

Ho

Er

Yb

Muscovite-altered sandstones (Camie River)

d 10

Altered sandstone/PAAS

Altered sandstone/PAAS

1

0,01

10

1

0,1

0,01

Feldspathic sandstones (Camie River)

b

La

Ce

Pr

Nd

Sm

Eu

Gd

Tb

Dy

Ho

Er

Yb

1

0,1

0,01

Muscovite-altered sandstone - Upper zone (Camie) Muscovite-altered sandstone - Interdigita on zone (Camie) Muscovite-altered sandstone - Lower zone (Camie)

La

Ce

Pr

Nd

Sm

Eu

Gd

Tb

Feldspathic sandstone (Camie) Least-altered sandstone (OTS-01) Muscovite-altered sandstone (OTS-01)

Dy

Ho

Er

Yb

100000

Sample/Chondrite

10000

1000

100 Camie River Unconformity-related

10

Magmatic Synmetamorphic 1

La

Ce

Pr

Nd

Sm

Eu

Gd

Tb

Dy

Ho

Er

Tm

Yb

Lu

data-point error ellipses are 68.3% . conf

0 .5

Beyer et al. (2012) UI: 1692±32 Ma LA-ICP-MS

0 .4

2200

SIMS spots

206 P b / 238 U

1800

0 .3 1400

0 .2

Höhndorf et al. (1987) UI: 1714±22 Ma TIMS

1000

0 .1 6 0 0

Intercepts at 66 ±24 & 1724 ±29 [ ±30] M a M S W D = 1,3

Camie River_Urn1 0 .0

0

2

4

207 P b/ 235 U

6

8

10

Urn1

Mol

Mol

Urn1

100000

Camie River uraninite

Sample/Chondrite

10000

1000

100 Camie River Unconformity-related

10

Magmatic Synmetamorphic 1

La

Ce

Pr

Nd

Sm

Eu

Gd

Tb

Dy

Ho

Er

Tm

Yb

Lu

57 Highlights

-

SIMS U/Pb and Re/Os age of uranium mineralization is 1724 Ma Uranium mineralization is ~440 Ma younger than sedimentation : uranium is not unconformityrelated Flat REE patterns in uraninite indicate uranium mineralization is related to high temperature (~320º C) fluids

-

57