Long-Lived Deformation History Recorded Along the Precambrian Thelon and Judge Sissons Faults, Northeast Thelon Basin, Nunavut

Canadian Journal of Earth Sciences

Long-lived deformation history recorded along the Precambrian Thelon and Judge Sissons faults, northeast Thelon Basin, Nunavut

Journal: Canadian Journal of Earth Sciences

Manuscript ID cjes-2020-0108.R2

Manuscript Type: Article

Date Submitted by the 08-Nov-2020 Author:

Complete List of Authors: Hunter, Rebecca; Laurentian University, Harquail School of Earth Sciences; British Columbia Geological Survey Branch Lafrance, Bruno; Laurentian University, Harquail School of Earth Sciences Draft Heaman, Larry; University of Alberta, Earth and Atmosphere Sciences Thomas, David; Cameco Corp

Aberdeen Lake, strike-slip fault, U-Pb zircon geochronology, Rae domain, Keyword: Thelon fault, Judge Sissons fault

Is the invited manuscript for consideration in a Special Not applicable (regular submission) Issue? :

1 Long-lived deformation history recorded along the Precambrian Thelon and Judge Sissons

2 faults, northeast Thelon Basin, Nunavut

4 R.C. Hunter, B. Lafrance, L.M. Heaman, and D. Thomas

6 R.C. Hunter (corresponding author): Mineral Exploration Research Centre, Harquail School of

7 Earth Sciences, Goodman School of Mines, Laurentian University, Sudbury, ON, Canada P3E

8 2C6; e-mail: [email protected]; phone: 306-371-0020

10 Present address: British Columbia Geological Survey, 1810 Blanshard Street, Victoria, BC Canada 11 V8T 4J1; e-mail: [email protected]; phone: 250-419-8744 12

13 B. Lafrance: Mineral Exploration Research Centre, Harquail School of Earth Sciences, Goodman

14 School of Mines, Laurentian University, Sudbury, ON, Canada P3E 2C6; e-mail:

15 [email protected]; phone: 705-675-1151

17 L.M. Heaman: Department of Earth and Atmospheric Sciences, University of Alberta, AB,

18 Canada T6G 2E3; e-mail: [email protected]; phone: 780-492-2778

20 D. Thomas: 760 Trent Crescent, Saskatoon, SK, Canada S7H 4S5; e-mail:

21 [email protected]; phone: 306-291-6093

22 23 24

1 © The Author(s) or their Institution(s) Canadian Journal of Earth Sciences Page 2 of 82

25 Abstract 26

27 The ENE-striking Thelon and Judge Sissons faults of south-central Nunavut are well-preserved, and

28 record long-lived dextral transcurrent movement with complex reactivation and fluid flow

29 histories. The faults cut across Archean gneisses, Paleoproterozoic plutons, and a Mesoproterozoic

30 sedimentary basin in the Rae domain of the western Churchill Province. They formed and were

31 reactivated during multiple deformation events beginning with an initial faulting event at 1830-

32 1760 Ma, followed by an epithermal faulting event at 1760-1750 Ma and late reactivation events

33 at 1600-1300 Ma. The initial faulting event produced the core-damage zone architecture of the

34 faults. Damage zones are characterized by multiple fracture sets, quartz veins and hydrothermal 35 crackle breccias, surrounding core zonesDraft defined by multiple mosaic to chaotic breccias and 36 cataclasites with dextral slip indicators. The epithermal faulting event is expressed by the presence

37 of crosscutting comb, crustiform-cockade and lattice-bladed quartz  hematite  carbonate veins,

38 and is likely associated with a magmatic event of similar age. The late reactivation events resulted

39 in the formation of irregular, non-cohesive crackle to mosaic breccias and gouges, which became

40 the primary pathways for uranium-bearing hydrothermal fluids and the formation of unconformity-

41 type uranium deposits. The Thelon and Judge Sissons faults are similar to other major continental

42 faults in the Rae domain (e.g. McDonald fault, Wager Bay shear zone), which formed during the

43 Paleoproterozoic Taltson-Thelon and Trans-Hudson orogenies, and to modern analogues, such as

44 the Karakorum, Altyn Tagh, and Hunan-Jaingxi faults, which formed during the Himalayan-

45 Tibetan orogeny and experienced prolonged hydrothermal and even hot spring activity.

47 Key words: Aberdeen Lake, strike-slip fault, U-Pb zircon geochronology, Rae domain, Thelon fault,

48 Judge Sissons fault.

2 © The Author(s) or their Institution(s) Page 3 of 82 Canadian Journal of Earth Sciences

49 1. Introduction 50 51 Studies on the architecture, kinematics, and alteration history of long-lived, strike-slip faults and

52 their related fracture networks provide important insights into the structural and hydrothermal

53 evolution of fault zones. Strike-slip faults are important crustal-scale continental structures that

54 accommodate horizontal movement as transform faults along tectonic plate boundaries and as

55 major transcurrent faults during collisional tectonics (Woodcock 1986; Sylvester 1988).

56 Continent-continent collision between the Indian and Eurasian plates led to the creation of the

57 Himalayan-Tibetan orogenic belt, which is bordered within its hinterland regions by a vast network

58 of transcurrent, right- and left-lateral, strike-slip faults (Molnar and Tapponnier 1975; Tapponnier

59 and Molnar 1977; Armijo et al. 1989). In such collisional tectonic settings, older pre-existing

60 structures including transcurrent faults (WoodcockDraft 1986) may be reactivated and undergo changes

61 in their slip kinematics over time, adding structural complexities to these long-lived structures. As

62 noted by Faulkner et al. (2003), the study of brittle fault rock products developed in the upper crust

63 is commonly impeded by their poor preservation as non-cohesive gouge, fractured rocks and

64 breccias, which are typically lost to erosion and weathering. As a result, the best studied faults are

65 relatively young structures (Eocene to Miocene) and located in arid climates where weathering is

66 minimal, for example, the individual faults of the large San Andreas fault system (Chester et al.

67 1993; Schultz and Evans 2000; Faulkner et al. 2003; Wibberley and Shimamoto 2003; Cembrano

68 et al. 2005; Jeffries et al. 2006; Faulkner et al. 2008; Bradbury et al. 2011). The challenges are even

69 greater for Precambrian strike-slip faults, including well-exposed examples such as the McDonald

70 and Bathurst faults, and Wager Bay shear zone of the northwest Canadian Shield, which are not

71 only affected by erosion and weathering but are also modified by multiple post-faulting tectonic,

72 magmatic, metamorphic, and hydrothermal events. Although, studies of these faults have provided

3 © The Author(s) or their Institution(s) Canadian Journal of Earth Sciences Page 4 of 82

73 detailed timing and tectonic constraints with respect to their ductile evolution (Thomas et al. 1976;

74 Henderson and Broome 1990; Therriault et al. 2017; Ma et al. 2020), the late brittle deformation

75 events remain much more difficult to resolve.

77 Brittle fault zones comprise three main components: (1) a core zone that accommodates

78 displacement along discrete slip surfaces; (2) a fractured damage zone that surrounds the core zone

79 and forms in response to repeated slip events; and (3) the undeformed protolith, which may be

80 preserved as lozenges between higher strain domains (Sibson 1977; Chester and Logan 1986;

81 Caine et al. 1996; Faulkner et al. 2003, 2010). Paleoseismic studies suggest that earthquake events

82 occur more frequently along strike-slip faults than along normal or reverse dip-slip faults (Sylvester

83 1988). Thus, large-scale strike-slip faultsDraft typically form during multiple slip and aftershock events

84 over periods of tens of millions of years. Crustal ruptures enhance fluid flow (Cox et al. 1986;

85 Sibson 1987; Micklethwaite and Cox 2006) and may result in the formation of ore deposits along

86 the faults (Henley 1985).

88 In this paper we describe the formation of internal structures along two well-exposed,

89 Paleoproterozoic faults named the Thelon and Judge Sissons faults that formed in the hinterland

90 regions of the Taltson-Thelon and Trans-Hudson orogenies. High crustal levels of these faults are

91 well-exposed at surface and their 3D geometry is constrained by multiple drill holes. The

92 investigation into the age and tectonic history of the Thelon and Judge Sissons faults provides

93 additional insights into how fractures, veins, and breccias formed and changed through time,

94 eventually creating the conditions that drove and focused hydrothermal fluid flow during

95 reactivation events.

4 © The Author(s) or their Institution(s) Page 5 of 82 Canadian Journal of Earth Sciences

97 2. Regional Geology 98 99 The western Churchill Province is divided into three Archean lithotectonic domains, namely, the

100 Rae and Hearne domains and the Chesterfield block, separated by a Paleoproterozoic suture zone

101 named the Snowbird tectonic zone (Fig. 1, Hoffman 1988; Ross et al. 2000; Mahan and Williams

102 2005; Berman et al. 2007; Martel et al. 2008). The Rae and Hearne domains consist of ca. 2750-

103 2650 Ma Neoarchean greenstone belts that were intruded by ca. 2667 Ma granitic plutons (Hinchey

104 et al. 2011) and invaded by a large number of ca. 2610-2580 Ma granitic and gabbroic plutons of

105 the Snow Island Suite. Snow Island Suite magmatism is further manifested as subvolcanic sills and 106 volcanic flows of the Pukiq Lake FormationDraft (Roddick et al. 1992; Peterson et al. 2015a). 107

108 The Archean Hearne and Rae cratons were reworked and metamorphosed to amphibolite-granulite

109 grade during four major Paleoproterozoic orogenies, including the Arrowsmith (ca. 2550-2300

110 Ma), Taltson-Thelon (ca. 2000-1930 Ma), Snowbird (ca. 1900 Ma), and Trans-Hudson (ca.

111 1870-1800 Ma) orogenies (Ashton 1988; Hrabi et al. 2003; Zaleski 2005; Berman et al. 2007).

112 During rifting of the cratons following the Arrowsmith orogeny, fluvial, marine, and lacustrine

113 sediments of the ca. 2300-1910 Ma Amer and Ketyet River groups were deposited in rift basins

114 unconformably overlying the greenstone belts (Heywood 1977; Tippett and Heywood 1978;

115 Patterson 1986; Rainbird et al. 2010; Calhoun et al. 2011, 2012; Pehrsson et al. 2013; Jefferson et

116 al. 2015). The greenstone belts and overlying Amer and Ketyet River groups were subsequently

117 deformed during the Taltson-Thelon, Snowbird, and Trans-Hudson orogenies. This was followed

118 by a new rifting event characterized by the deposition of the Dubawnt Supergroup (i.e. Baker Lake,

119 Wharton and Barrensland groups) in large extensional and intracratonic basins (Gall et al. 1992;

5 © The Author(s) or their Institution(s) Canadian Journal of Earth Sciences Page 6 of 82

120 Peterson et al. 2002; Rainbird et al. 2003). The Baker Lake Group comprises alluvial-fluvial to

121 lacustrine sedimentary rocks and coeval ultrapotassic volcanic rocks (Rainbird et al. 2003), which

122 are interpreted as the extrusive equivalent of the ca. 1845-1795 Ma granitic, syenitic, and

123 lamprophyric intrusions of the Hudson Intrusive Suite (Sandeman et al. 2000; Peterson et al. 2002;

124 van Breemen et al. 2005; Miller and Peterson 2015). The overlying Wharton Group consists of

125 alluvial to aeolian sandstone and conglomerate with intercalated rhyolite flows, which were fed by

126 the coeval ca. 1765-1750 Ma granitic to gabbroic Kivalliq Igneous Suite and related ca. 1750 Ma

127 McRae Lake gabbroic dike swarm (Rainbird et al. 2003; van Breemen et al. 2005; Ernst and

128 Bleeker 2010; Peterson et al. 2015b).

129

130 Late NW- and ENE- to NE-trending faultsDraft cut across the western Churchill province (Fig. 2). The

131 NW-trending faults include the sinistral Bathurst fault, and other unnamed minor faults, and the

132 ENE-to NE-trending faults are represented by several major, ductile to brittle, dextral faults,

133 including the McDonald, Chantrey Mylonite, Amer, Thelon, Judge Sissons, Wager Bay, Tyrell,

134 and Grease River faults. These faults formed and were reactivated during the Taltson-Thelon and

135 Trans-Hudson orogenies (Henderson and Broom 1990) prior to the deposition of the Barrensland

136 Group. From base to top, the latter consists of ca. <1740 Ma (U-Pb detrital zircon - Rainbird and

137 Davis (2007) and U-Pb apatite - Miller et al. (1989)) aeolian and fluvial to shallow marine

138 sandstone and conglomerate of the Thelon Formation, 1540 ±30 Ma (U-Pb baddeleyite –

139 Chamberlain et al. 2010) shoshonitic basalt of the Kuungmi Formation, and stromatolitic limestone

140 of the Lookout Point Formation (Cecile 1973; Gall et al. 1992; Hiatt et al. 2003; Palmer et al. 2004;

141 Rainbird and Davis 2007). North-northwest-trending dikes of the ca. 1267 Ma Mackenzie dike

142 swarm cut across all other rock units (LeCheminant and Heaman 1989).

6 © The Author(s) or their Institution(s) Page 7 of 82 Canadian Journal of Earth Sciences

143

144 3. Geology of the Aberdeen Lake area 145

146 The Aberdeen Lake greenstone belt is located south of the northeastern lobe of the Thelon Basin

147 in the central Rae domain, Nunavut (Fig. 1). It comprises three amphibolite-facies,

148 lithostratigraphic sequences: (1) a Lower Sequence of ca. 2750 Ma komatiite, metagabbro, TTG

149 intrusions, and TTG-contaminated mafic gneiss; (2) a Middle Sequence of ca. 2687 Ma to ca. 2680

150 Ma psammopelitic gneiss, iron formation, and intermediate to felsic gneiss; and (3) an Upper

151 Sequence of <2650 Ma pelitic to psammopelitic gneiss with minor iron formation and arkosic

152 gneiss (Hunter et al. 2018; Fig. 3). The greenstone belt is stratigraphically similar to the Woodburn 153 Lake group of the Meadowbank area andDraft to other Neoarchean supracrustal assemblages of the 154 Rae domain (Hunter et al. 2018). Overlying orthoquartzite and conglomerate of the

155 Paleoproterozoic Ketyet River group are locally exposed along the Thelon fault east of Aberdeen

156 Lake.

157

158 The Neoarchean supracrustal sequences in the Aberdeen Lake area are interfolded with and locally

159 crossed by a quartzofeldspathic orthogneiss composed of plagioclase and K-feldspar (25-35%),

160 quartz (20-25%), biotite (3-5%), hornblende (1-3%), and trace magnetite, titanite and zircon (Fig.

161 4a). It has a gneissic banding defined by 1-20 cm-thick leucosomes of quartz and feldspar

162 alternating with <1 cm thick paleosomes dominated by biotite and hornblende. Amphibolite layers

163 are parallel to the gneissic banding and are interpreted as rafts of the Lower Sequence.

164

165 The quartzofeldspathic gneiss is crossed by intrusions of the ca. 1845 Ma to 1795 Ma Hudson

166 Intrusive Suite (Fig. 3). In order of decreasing abundance, the latter consists of monzogranite,

7 © The Author(s) or their Institution(s) Canadian Journal of Earth Sciences Page 8 of 82

167 mafic syenite, quartz-feldspar porphyry, microsyenite, and lamprophyre (Miller and Peterson

168 2015). Monzogranite is exposed as medium- to coarse-grained equigranular plutons, which are

169 typically ≤5 km in diameter, and as metre-thick dikes. These intrusions are composed of K-feldspar

170 and plagioclase (75-90%), quartz (7-20%), hornblende (1-3%), and trace fluorite (Fig. 4b).

171 Associated with these intrusions are fine- to coarse-grained mafic syenite plugs and dikes,

172 consisting of K-feldspar (75-85%), hornblende (10-15%), quartz (3-7%) and trace titanite and

173 fluorite (Fig. 4c). Some syenite plugs contain K-feldspar megacrysts (1-2 cm in size) with inclusions

174 of hornblende and phlogopite in a coarse-grained phlogopite-rich groundmass (Fig. 4d).

175 Undeformed, 0.2-1.0 m thick, quartz-feldspar porphyry and very fine-grained to aphanitic, K-

176 feldspar-rich microsyenite dikes cut the monzogranite and mafic syenite intrusions. The porphyry

177 dikes consist of quartz (5-10%) and K-feldsparDraft (15-25%) phenocrysts (2-5 mm in size) in a fine-

178 grained quartzofeldspathic matrix with trace biotite. These porphyry dikes are in turn crossed by

179 E- to ENE-trending, fine- to medium-grained lamprophyre dikes, varying in thickness from 0.1 to

180 2.0 m (Fig. 4e). The lamprophyre dikes contain 10-15% phlogopite phenocrysts (1-3 mm in size)

181 with locally glomeroporphyritic and radiating textures, and 1-3% K-feldspar phenocrysts (3-5 mm

182 in size), surrounded by a fine-grained matrix of K-feldspar, quartz, and minor plagioclase and

183 garnet. Leucogabbroic dikes of the younger McRae Lake dike suite cut across all other intrusions

184 and can be traced across the map area by their strong magnetic signature. They are composed of

185 plagioclase (85-90%), magnetite (5-10%), and clinopyroxene (3-5%) (Fig. 4f).

186

187 4. Ductile deformation 188 189 Ductile structures in the Aberdeen Lake area predate the emplacement of the undeformed Hudson

190 Intrusive Suite and are grouped into four deformation events based on overprinting relationships

8 © The Author(s) or their Institution(s) Page 9 of 82 Canadian Journal of Earth Sciences

191 (Fig. 3). Unpublished high-resolution magnetic maps by Cameco Corporation are used to delineate

192 units and fold interference patterns in areas of poor exposure. Small-scale structures and

193 overprinting relationships are best exposed in metasedimentary rocks of the Middle and Upper

194 sequences. The earliest structures (D1) include a S1 gneissic foliation, which is axial planar to rare,

195 isoclinal, F1 folds (Figs. 5a, b). The S1 foliation is parallel to bedding in the metasedimentary rocks

196 and defined by alternating biotite-rich and recrystallized quartzofeldspathic layers. It is also

197 defined by recrystallized, isoclinally folded, and boudinaged quartz veins, varying in thickness

198 from <1-5 cm in thickness. As the S1 foliation is parallel to bedding in the supracrustal units, the

199 main foliation in the Aberdeen Lake area is best described as a composite S0-S1 foliation (Fig. 5c).

200

201 The three Neoarchean supracrustal assemblagesDraft underwent major folding and thrusting. The S1

202 foliation and F1 folds are overprinted by D2 structures, including NE- to NW-trending, tight to

203 isoclinal, inclined to recumbent, F2 folds. The latter are associated with an axial planar S2 cleavage

204 (Fig. 5d) and coaxial L2 lineation defined by biotite and hornblende. East of Marjorie Lake, banded

205 iron formation and psammopelite are folded by map-scale isoclinal F2 folds, striking northwest

206 and plunging (20◦ to 30◦) to the north-northwest (Figs. 3, 6). Thrust faults are expressed by

207 pronounced S2 and L2 fabrics in high-strain zones that are also characterized by transposed,

208 rootless, mesoscopic F2 folds. The thrust faults are parallel and cut across highly strained and

209 strongly attenuated limbs of large-scale F2 folds (Section A-A‘ on Fig. 3), resulting in the

210 juxtaposition across thrust faults of folded and overturned anticlinal and synclinal stacks of the

211 Lower, Middle and Upper sequences (Fig. 3).

212

213 Second generation fold structures (F2) are overprinted by small, outcrop-scale F3 chevron folds with

9 © The Author(s) or their Institution(s) Canadian Journal of Earth Sciences Page 10 of 82

214 angular hinges, straight limbs, and amplitudes of a few centimetres to decimetres (Fig. 5e). The

215 folds have an axial planar S3 crenulation cleavage defined by recrystallized biotite along the long

216 limb of F3 crenulations. On outcrop scale, the S2 and S3 cleavages are locally overprinted by open,

217 upright to steeply-inclined, NNE- to NE-striking and shallow-plunging, F4 folds (Fig. 5f). The

218 latter have a steeply-dipping axial planar cleavage (S4) defined by biotite. On map scale, the older

219 F1 and F2 folds are gently warped by the F4 folds.

220

221 5. Brittle deformation 222 223 Two major ENE-trending faults, the Thelon and Judge Sissons faults, are present in the map area 224 within the northeastern part of the Paleo-Draft to Mesoproterozoic Thelon Basin (Figs. 2, 3). The two 225 faults can be traced over strike lengths of up to 200 km from the Thelon Basin, across Archean

226 basement rocks, and through a ca. 1830 Ma pluton of Hudson Intrusive Suite. The pluton is

227 dextrally offset by 14 km and 5 km along the Thelon and the Judge Sissons faults, respectively. The

228 younger, strongly magnetic McRae Lake dikes show similar but smaller dextral offsets of 2.5 km and

229 0.5 km, respectively, along the same faults (Figs. 3, 7). Second order faults with dextral offsets of

230 tens to hundreds of metres are parallel to the Thelon and Judge Sissons faults, while less prominent

231 NW-to NNW-trending faults sinistrally displace the Hudson Intrusive Suite pluton and the McRae

232 Lake dikes by up to 400 m (Fig. 7).

233

234 5.1. Terminology 235 236 Brittle fault rocks are named following the classification of Woodcock and Mort (2008). Their

237 classification is based on the relative proportion of clasts and matrix, where clasts have a minimum

238 size of 2 mm and the matrix consists of fine-grained, comminuted, particulate material less than

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239 0.1 mm in size. Cataclasites contain <30% clasts and breccias contain >30% clasts. Cataclasites

240 are subdivided into protocataclasite (>50% matrix), mesocataclasite (50-90% matrix), and

241 ultracataclasite (>90% matrix). Breccias are subdivided into crackle (75-100% clasts), mosaic (60-

242 75% clasts), and chaotic (30-60% clasts) breccias. Crackle breccias have jigsaw-fit textures with

243 angular clasts that underwent no or minor rotation and are separated from each other by thin films

244 of matrix. Clasts in mosaic breccias underwent rotation but can be fitted with other nearby clasts,

245 whereas clasts in chaotic breccias have undergone more rotation and have lost their fit to other

246 nearby clasts. Hydrothermal breccias have a cement of hydrothermal minerals rather than a matrix

247 (e.g. Jébrak 1997) but otherwise follow the same classification based on percent clasts. Cataclastic

248 breccia is used as a general term that includes crackle, mosaic, and chaotic breccias with a

249 comminuted particulate matrix. Draft

250

251 5.2. Initial faulting event 252 253 Similar to other well-exposed faults (e.g. Faulkner et al. 2010), the Thelon and Judge Sissons faults

254 have a complex fault architecture consisting of several cm- to m-thick core zones surrounded by up

255 to 50 m wide damage zones, resulting in overall fault widths between 100 m and 500 m. The

256 damage zones consist of fractured and brecciated rocks cut by quartz veins and hydrothermal quartz

257 crackle breccias (Fig. 8). Fractures in the damage zones are dominantly NW-SE striking with an

258 average dip of 65◦ to the northeast or southwest (Figs. 9a, b). They are located throughout the

259 damage zone and become partially to completely filled with quartz near the core zones. Other

260 secondary fractures strike N to NNE, and ENE to ESE with similar steep dips. The N- to NNE-

261 striking fractures are late structures with strike lengths of <5 m. They cut across early breccias and

262 veins, which are sinistrally offset by 0.1-1.0 m along the fractures. The veins are dominantly WSW-

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263 WNW-striking with an average dip of 60◦ to the north (Figs. 9c, d). They vary in thickness from 1

264 mm to 50 cm and consist mainly of quartz with massive crystalline to saccharoidal textures. Quartz

265 crystals in crystalline-textured veins are equant to semi-equant with a grain size of 0.25-0.50 mm.

266 They have well-developed growth zoning defined by clear quartz alternating with bands of fluid

267 inclusion-rich cloudy quartz (Fig. 10a). The merging of these veins forms hydrothermal crackle

268 breccias with jigsaw fit fragment textures and hydrothermal chaotic and mosaic hydrothermal

269 breccias with less fragments and more quartz vein matrix material (Fig. 10b).

270

271 Contacts between the damage zones and fault cores are sinuous and roughly parallel to the ENE-

272 trending trace of the faults. They are irregular and defined by the presence of cataclastic breccias.

273 The core zones are characterized by a gradualDraft decrease in the percentage of large clasts (>2 mm)

274 from breccias with more than 30% clasts, to cataclasites with less than 30% clasts (Fig. 11a).

275 Breccia textures change from dominantly crackle type to chaotic type from the margins to the

276 centres of the core zones. Clasts are typically angular and vary in size from 0.5-10.0 cm. Ghost

277 quartz veins are represented by trains of quartz fragments that can be traced from the mosaic

278 breccias to the chaotic breccias but disappear within the protocataclasites and mesocataclasites due

279 to their complete disaggregation into small random fragments (Fig. 11b). Fragments in both types

280 of cataclasites are matrix-supported, angular to sub-angular, and vary in size from <1 mm to 1.5

281 cm. The larger fragments are clasts of broken-up quartz veins, crackle breccia, and hematized wall

282 rock. The matrix material is cohesive, silicified and red due to strong hematization, and consists

283 of comminuted, fine-grained, wall rock material. Cm-thick zones of closely-spaced fractures cut

284 across the cataclasites. They are parallel to the trend of the faults and are lined by striations with

285 an average rake of 7◦ to the east. Steps cutting across the striations, and the asymmetry of trailed

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286 material markings (Doblas, 1998) along the slip surfaces (Fig. 12a), suggest that the fractures

287 represent discrete slip surfaces that formed during dextral transcurrent slip along the faults (Fig.

288 12b).

289

290 Clasts of hydrothermal breccias similar to those described above are present in the basal

291 conglomerate of the Thelon Formation adjacent to the Thelon and Judge Sissons faults. This

292 indicates that faults with similar veins and structures as the Thelon and Judge Sissons faults formed

293 and were eroded prior to the deposition of the basal Thelon Formation (Figs. 10c, d).

294 295 5.3. Epithermal faulting event 296 Draft 297 Younger quartz veins overprint the initial fault-related veins, breccias and cataclasites. These

298 younger veins have comb, lattice-bladed, crustiform, and cockade textures indicative of open space

299 filling. Comb-textured veins vary in thickness from 2 mm to 5 cm and consist of euhedral quartz

300 crystals with hexagonal cross-sections oriented sub-perpendicular to the vein walls (Fig. 13a).

301 Quartz crystals contain hematite inclusions and are zoned with alternating clear and white growth

302 bands. Veins with lattice-bladed textures consist of millimetre-thick quartz blades that are up to 2

303 cm in length (Fig. 13b). The blades have a radiating habit, are composed of carbonate minerals

304 almost completely replaced by quartz, and are surrounded by interstitial fine-grained red hematite.

305 Other veins consist of alternating bands of comb-textured quartz, white chalcedonic to

306 saccharoidal crustiform quartz, and lattice-bladed quartz (Fig. 13c). Veins with cockade textures

307 have thicknesses of 1-5 cm. They contain rounded to angular, white to pink, quartz fragments,

308 which are surrounded by crustiform quartz with red hematite bands and fine-grained comb-

309 textured quartz bands (Fig. 13d).

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310

311 Clasts of quartz veins with comb and bladed textures (Figs. 13e, f) are intermixed with clasts of

312 quartzite, gneiss, granite, and rhyolite of the Pitz Formation within basal Thelon Formation

313 conglomerate. Similar to the initial faulting event, this provides evidence that an epithermal

314 faulting event occurred prior to the deposition of the basal Thelon Formation.

315

316 5.4. Late reactivation event(s) 317

318 Late reactivation of the Thelon and Judge Sissons faults extensively fractured the adjacent rocks and

319 formed non-cohesive, crackle to mosaic, cataclastic breccias. The fractures typically occur along

320 the margins of pre-existing planar anisotropiesDraft such as older veins, breccias, gneissic foliation

321 planes and lithological contacts (Fig. 14a). Fractures are locally striated and filled with a non-

322 cohesive gouge, consisting of very fine-grained sand- to clay-sized comminuted and altered wall

323 rock material. The cataclastic breccias are clast-supported and vary in thickness from 1-100 cm.

324 Clasts are angular to sub-angular, 0.1-5.0 cm in size, and are surrounded by a matrix composed of

325 0.2-2.0 mm crushed rock fragments (Fig. 14b).

326

327 These breccias are spatially associated with uranium mineralization and were intensely altered

328 during this hydrothermal event. Measurement of their orientation and that of the gouge-filled

329 fractures is hindered by the broken up and friable nature of the drill core. However, all planar

330 structures including fractures, breccia, and most lithological contacts, as well as striations along

331 fracture planes, are typically oriented at low angles to the drill core axis. As most drill holes have

332 steep to near vertical inclinations, we infer that pre-existing anisotropies were reactivated as

333 normal or reverse dip-slip faults during the uranium mineralizing event(s) (Figs. 15a, b). Where

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334 the rock is less altered and drill core is more continuous, the orientation of the gouge-filled

335 fractures can be estimated and they typically strike E to NE with moderate to shallow dips (Figs.

336 16a, b).

337

338 5.4.1. Uranium mineralization and alteration 339

340 Uranium mineralization is found at the Ayra prospect in the Thelon Formation and at the Qavvik

341 and Tatiggaq prospects in the Archean supracrustal rocks and to a lesser extent in monzogranite

342 and syenite of the Hudson Intrusive Suite (Fig. 3). The Ayra prospect is characterized by intense

343 clay alteration and weakly enriched uranium values (i.e. metasedimentary rocks: >50 ppm U; 344 sandstone: >2 ppm U). The Qavvik andDraft Tatiggaq prospects are located off the Thelon and Judge 345 Sissons faults. The two faults were reactivated and altered during the uranium mineralizing event

346 but are only weakly mineralized. The Tatiggaq prospect is the more prospective of the two and is

347 located in fractured and brecciated rocks along second order ENE-trending faults close to or

348 parallel to the Judge Sissons fault. The Qavvik prospect occurs along another second order ENE-

349 trending fault between the Thelon and Judge Sissons faults (Fig. 3). Uranium mineralization at

350 these two prospects consists of pitchblende, uraninite, coffinite, and uranophane along cm- to m-

351 scale reduction-oxidation fronts that extend horizontally from gouge-filled fractures and breccias.

352 Uranium minerals are largely precipitated as disseminations and blebs along fluid-driven

353 reduction-oxidation fronts (Figs. 17a, b), and as massive pitchblende and uraninite pods filling

354 open space cavities along fractures and breccias. Proximal, medial, distal, and outer distal

355 alteration zones surround and extend more than 100 m into the hanging wall of the mineralized

356 zones (Fig. 18). The proximal alteration zone is 1-15 m thick and consists of strong pervasive and

357 alternating banded to patchy red hematite, limonite, sooty sulphide, chlorite and clay alteration.

15 © The Author(s) or their Institution(s) Canadian Journal of Earth Sciences Page 16 of 82

358 The medial alteration zone is 3-70 m thick and is characterized by weak to strong, pervasive to

359 patchy limonite, sooty sulphide and clay alteration. The distal alteration zone is 20-80 m thick and

360 comprised weak to strong pervasive clay alteration and locally moderate to strong pervasive dark

361 purple and locally silver specular hematite alteration. The outer distal alteration zone extends 50-

362 100 m from the mineralized zones and is characterized by coarse (1-2 mm) specular hematite and

363 white clay-filled fractures. The specular hematite is dark purple to silver and forms 0.1-5.0 cm

364 wide vuggy veins filling fractures along pre-existing breccia contacts and quartz vein margins (Fig.

365 14c). The fractures are cut by 1-10 mm thick, white clay-rich fractures with irregular, serrated

366 walls (Fig. 14d). The latter fractures are NW- to ENE-striking, moderate- to steep-dipping, and

367 form stockworks overprinting the previously faulted rocks (Figs. 16c, d).

368 Draft 369 6. U-Pb LA-ICP-MS Geochronology 370

371 One sample of quartzofeldspathic orthogneiss and two samples of monzogranite from the Hudson

372 Intrusive Suite were collected for U-Pb zircon geochronology (Fig. 3). Age uncertainties are

373 reported at the 2σ level. The methodology is described in Supplementary Data, and the data are

374 listed in Supplementary Table S1.

375

376 6.1 Quartzofeldspathic orthogneiss (Sample 204961) 377

378 Sample 204961 was collected near the northeast end of Marjorie Lake from a fine- to medium-

379 grained quartzofeldspathic orthogneiss with 80-85% feldspar and quartz, 15-20% biotite and

380 hornblende and up to 2% magnetite. The orthogneiss is strongly deformed and contains

381 dismembered and highly attenuated F2 folds, a strong L2 stretching lineation, and overprinting F3

16 © The Author(s) or their Institution(s) Page 17 of 82 Canadian Journal of Earth Sciences

382 chevron folds. Zircons are colourless to light brown and prismatic with aspect ratios of 2:1 to 3:1

383 and maximum dimensions varying from 100-400 µm. Cathodoluminescence (CL) imaging show

384 that most zircons have weak, irregular, and oscillatory zoning and several zircons have non-

385 luminescent rims overprinting resorbed primary zoning (Fig. 19a). Most 207Pb/206Pb dates range

386 between 2634 Ma and 2550 Ma. On a Concordia diagram, the zircon analyses plot along a

387 discordia line with an upper intercept date of 2601.7 ±4.2 Ma (n=84; MSWD=4.0) (Fig. 20). Zircon

388 grains 2B, 4, 32, 44, 82, 84, 97, and 98 were excluded from this age calculation because they have

389 distinctly young or old 207Pb/206Pb dates (Supplementary Table S1). The 2601.7 ±4.2 Ma date is

390 interpreted as the crystallization age of the quartzofeldspathic orthogneiss.

391 392 6.2 Monzogranite (Sample 204964)Draft 393

394 A coarse-grained, massive monzogranite was sampled between the Thelon fault and Judge Sissons

395 fault. The zircons vary in appearance from colourless to light brown with orange to red staining.

396 They are 50-250 µm in size and equant to elongate with aspect ratios between 2:1 and 4:1. Most

397 207Pb/206Pb dates are between 2709 Ma and 1805 Ma, but a few distinct younger dates are between

398 1794 Ma and 1680 Ma (Fig. 21a). The former dates plot on two separate discordia lines, which

399 yield upper intercept dates of 1832.9 ±9.6 Ma (n=25; MSWD=7.2) and 2592.4 ±7.1 Ma (n=18;

400 MSWD=3.1) (Fig. 21a). The younger date is interpreted as the crystallization age of the

401 monzogranite and the older date is interpreted as the age of inherited zircons. The younger ca.

402 1794 Ma and 1680 Ma dates are inferred to reflect the influence of Pb-loss during Wharton Group

403 volcanism and subsequent hydrothermal events.

404 405 6.3 Monzogranite (Sample 204974) 406

17 © The Author(s) or their Institution(s) Canadian Journal of Earth Sciences Page 18 of 82

407 Sample 204974 was collected from drill hole LOB-002 that cuts through the Judge Sissons fault.

408 The zircons are colourless to light brown with local red and yellow staining. They form prismatic

409 crystals with aspect ratios of 2:1 to 3:1 and long dimensions of 100-400 µm. They show complex

410 zonation patterns defined by rounded, resorbed, oscillatory zoned cores mantled by younger zircon

411 overgrowth (Figs. 19b, c). Most 207Pb/206Pb dates are between 2752 Ma and 1793 Ma, but younger

412 subsets of dates between 1773 Ma and 1702 Ma, and at 1641 Ma and 1395 Ma (Fig. 21b), are also

413 present. These younger zircons display either fragmented fluorescence patterns or large non-

414 fluorescent regions, which overprint older oscillatory zircon growth (Figs. 22a, b, c). On a

415 Concordia plot, the analyses plot along two discordia lines with intercept dates of 1831 ±11 Ma

416 (n=19; MSWD=10.8) and 2580 ±21 Ma (n=13; MSWD=34) (Fig. 20b). As with sample 204964,

417 the younger ca. 1830 Ma date is interpretedDraft as the crystallization age of the monzogranite, and the

418 ca. 2580 Ma date is interpreted as the age of inherited zircons. The younger ca. 1773 Ma to 1395

419 Ma zircon ages are interpreted to reflect the influence of Pb-loss from overprinting hydrothermal

420 events.

421

422 7. Discussion 423 424 7.1. Neoarchean magmatism and Paleoproterozoic ductile deformation 425

426 The emplacement of the igneous protolith of the quartzofeldspathic gneiss at ca. 2600 Ma marked

427 a period of extensive late Archean magmatism that ended with the emplacement of plutonic and

428 volcanic rocks of the ca. 2610 Ma to 2580 Ma Snow Island Suite (Roddick et al. 1992; Hinchey et

429 al. 2011; Peterson et al. 2015a). This “granite bloom” has been related to the build-up of a

430 continental arc on the margin of the Rae-Chesterfield block craton above a northwest-dipping

18 © The Author(s) or their Institution(s) Page 19 of 82 Canadian Journal of Earth Sciences

431 subduction zone (Berman et al. 2010, 2013). Subsequent Paleoproterozoic D1 and D2 deformation

432 events produced the strong gneissic banding in these rocks as well as the regional fabrics, folds,

433 and thrust faults in the supracrustal rocks of the Aberdeen Lake greenstone belt. These structures

434 and deformation events correlate with similar features in the nearby Meadowbank area, where they

435 formed between ca. 1910 Ma and ca. 1840 Ma during the Trans-Hudson orogeny (Pehrsson et al.

436 2013) and are thought to be responsible for thrusting of Archean supracrustal rocks over the

437 Paleoproterozoic Ketyet River Group (Ashton 1988; Zaleski et al. 2000; Sherlock et al. 2004;

438 Pehrsson et al. 2013).

439

440 The undeformed ca. 1830 Ma monzogranitic and syenitic plutons in the Aberdeen Lake area

441 belong to the extensive ca. 1845 Ma to Draft1795 Ma Hudson Intrusive Suite (Sandeman et al. 2000;

442 Miller and Peterson 2015). These plutons cut the D1 and D2 fabrics and related folds and are not

443 overprinted by D3 and D4 structures. Thus, as these plutons are offset by the Thelon and Judge

444 Sissons faults, they provide a minimum age of ca. 1830 Ma for the regional ductile deformation

445 and a maximum age for the initiation of brittle faulting in the Aberdeen Lake area.

446 447 7.2. Initial faulting event (ca. 1830-1760 Ma) 448

449 Several lines of evidence suggest that the ENE-trending Thelon and Judge Sissons faults formed

450 as dextral transcurrent structures. First, a large monzogranitic pluton of the Hudson Intrusive Suite

451 is dextrally offset by up to 14 km along the two faults as showed on geological and airborne

452 magnetic maps (Figs. 3, 7). Second, in the damage zone of the Judge Sissons fault, quartz veins

453 are steeply-dipping and occupy fractures oriented either parallel or clockwise to the fault, and are

454 offset by sinistral fractures oriented at a high-angle clockwise to near-perpendicular to the ENE-

19 © The Author(s) or their Institution(s) Canadian Journal of Earth Sciences Page 20 of 82

455 striking fault. This is consistent with their formation as extensional veins that were offset by

456 antithetic Riedel (R’) shear fractures during NW-directed shortening and dextral slip along the

457 fault (Figs. 8, 9a, b). Third, the shallow rake of slickenlines along slip surfaces in the fault core

458 zones, and the asymmetry of slickenline-parallel trailed material markings and cross-cutting steps,

459 are further indicative of dextral transcurrrent movement. During faulting, the breccias, cataclasites,

460 and fractures acted as pathways for the infiltration of fluids, the formation of quartz veins, and

461 pervasive hematization (Figs. 23a, b). Continued movement along the faults fractured and

462 fragmented the veins and breccias and resulted in the development of progressively younger

463 generations of cross-cutting quartz veins and breccias (Fig. 23c). The formation of the faults is

464 bracketed between ca. 1830 Ma, the crystallization age of the monzogranite pluton, and ca. 1760

465 Ma, the age of Wharton Group magmatismDraft (Fig. 24 - initial faulting event), as discussed below.

466

467 Damage zones are areas of highly fractured rock surrounding one or multiple core zones (Chester

468 and Logan 1986; Faulkner et al. 2003). Processes involved in the formation of damage zones

469 include: 1) coalescence of tensile microcracks during fault initiation; 2) linkage of tips of adjacent

470 parallel faults; 3) tensile and compressive fracturing in process zones at the tips of propagating

471 faults; 4) cumulative fracture damage caused by slip along irregular fault surfaces; and 5) co-

472 seismic dynamic fracture propagation (Mitchell and Faulkner 2009; Faulkner et al. 2010). Damage

473 zones likely form from a combination of those processes and thus interpreting fracture patterns is

474 commonly difficult especially for large and long-lived fault zones (Faulkner et al. 2010). Within

475 the damage zones of the Thelon and Judge Sissons faults, the fracture and vein pattern is relatively

476 simple as discussed above. Antithetic R’ structures and tensile fractures are more prominent than

477 synthetic Riedel (R) structures, which would be anomalous if this were not also observed along

20 © The Author(s) or their Institution(s) Page 21 of 82 Canadian Journal of Earth Sciences

478 other strike-slip faults such as the Malta Fault in the Rio do Peixe Basin (Kim et al. 2006).

479 Abundant antithetic R' structures develop at the tip of the master fault and commonly form in

480 association with branch faults producing block rotation and increased seismic activity (Nicholson

481 et al. 1986; Kim et al. 2003). The presence of fault-parallel extensional veins with massive to

482 comb textures suggests that the fractures hosting the veins opened against the normal stresses

483 acting on the faults. This is not typically described in brittle fault zones but can be explained by

484 the injection of high-pressure fluids following depressurization and even spalling of fault-parallel

485 sheets of wall rocks due to stress release immediately following seismic events (Cox 2020).

486

487 7.3. Epithermal faulting event (ca. 1760-1740 Ma) 488 Draft

489 The epithermal faulting event (Fig. 23d) marks the release of hydrothermal fluids at upper crustal

490 levels as described for epithermal (epizonal) precious metal deposits, which typically form at

491 temperatures between 160◦C and 270°C (Lindgren 1933; Henley 1985, 1991; Hedenquist et al.

492 2000; Henley and Berger 2000; Simmons et al. 2005; Sillitoe 2015). The Judge Sissons and Thelon

493 faults contain quartz veins with comb, crustiform, banded, cockade, and lattice-bladed textures

494 similar to those associated with epithermal deposits. Such veins form by the precipitation of silica

495 from hydrothermal fluids that become supersaturated in silica during cooling, fluid mixing, fluid-

496 rock interactions, boiling, and transient pressure drops (Buchanen 1981; Fournier 1985; Dong et

497 al. 1995), with the main mechanisms involving the cooling and boiling of supersaturated

498 hydrothermal fluids. “Gentle boiling” or adiabatic boiling involves the slow cooling and

499 decompression of boiling fluids during their ascent (Fournier 1985; Dobson et al. 2003; Moncada

500 et al. 2012; Shimizu 2014) and silica becomes saturated in the hydrothermal fluid as steam

21 © The Author(s) or their Institution(s) Canadian Journal of Earth Sciences Page 22 of 82

501 separates from it. “Intense boiling” involves the sudden decompression and fast cooling of

502 hydrothermal fluids due to rapid and transient pressure drops during faulting (Fournier 1985;

503 Dobson et al. 2003; Moncada et al. 2012; Shimizu 2014) and a large pressure drop can cause the

504 instantaneous and total conversion of the fluid into steam, a process called “flashing” (Brown

505 1986).

506

507 Comb-textured veins with zoned euhedral quartz crystals are indicative of protracted open space

508 filling at upper crustal levels typical of epithermal systems or peripheral to those systems (Oliver

509 and Bons 2001). This texture is thought to form during gentle boiling or slow cooling of a non-

510 boiling fluid that is slightly supersaturated in silica (Shimizu 2014). On the other hand, banded

511 crustiform veins typically comprise chalcedonyDraft bands that form by rapid nucleation and

512 precipitation of silica from a fluid with a high degree of silica supersaturation achieved through

513 intense boiling and rapid cooling (Fournier 1985; Henley and Hughes 2000; Dobson et al. 2003;

514 Moncada et al. 2012; Shimizu 2014). Chalcedonic and saccharoidal quartz layers may precipitate

515 directly as microcrystalline quartz or as amorphous silica that later recrystallizes into more stable

516 microcrystalline quartz (Fournier 1985). Lattice-bladed veins or bands in crustiform veins formed

517 during episodic flashing (Browne 1978; Buchanen 1981; Brown 1986) and precipitates first as

518 primary carbonate or barite crystals that are later dissolved and replaced by quartz (Buchanen

519 1981). The last type of vein, cockade-textured veins, are produced by the entrainment of host rocks,

520 fault rocks, and older vein fragments in ascending hydrothermal fluids (Shimizu 2014). The

521 breccias form during faulting, and the fragments may become coated by chalcedony bands, which

522 formed during intense boiling, and comb-textured and granular crystalline quartz, which

523 precipitated in bands and infills between fragments during gentle boiling (Dobson et al. 2003;

22 © The Author(s) or their Institution(s) Page 23 of 82 Canadian Journal of Earth Sciences

524 Shimizu 2014) or cooling from a non-boiling fluid (Moncada et al. 2012). Epithermal

525 hydrothermal systems are dynamic environments with maximum boiling depths that may shift over

526 time (Moncada et al. 2012), resulting in spatially overlapping veins with distinct different textures

527 as observed along the Judge Sissons and Thelon faults.

528

529 Similar to epithermal deposits, hydrothermal fluid flow along the Thelon and Judge Sissons faults

530 may have been driven by localized magmatic activity. Similar vein textures are described at the

531 Contact uranium prospect in the Kiggavik area, 15 km southeast of Aberdeen Lake (Grare et al.

532 2018). Possible hydrothermally-modified zircon grains from this study are dated at ca. 1760-1720

533 Ma, which correlates with the age of the Wharton Group felsic volcanic rocks that are exposed in

534 the Aberdeen Lake area along strike ofDraft the Judge Sissons fault (Fig. 3). The mineralogy of the

535 veins associated with this event is indicative of a low-sulphidation epithermal environment. The

536 age of the Wharton volcanic rocks and hydrothermally-modified zircon, together with the

537 intermixing of clasts of Wharton Group volcanic rocks and epithermal quartz veins at the base of

538 the Thelon Formation (Fig. 13f), suggests that the flow of hydrothermal fluids along the Thelon

539 and Judge Sissons faults is likely linked to a broader Wharton Group magmatic event that predates

540 the deposition of the Thelon Formation. This event is therefore bracketed between ca. 1760-1750

541 Ma, the emplacement age of the Wharton Group (Peterson et al. 2002; Rainbird et al. 2003;

542 Peterson et al. 2015b), and ca. 1720 Ma, the onset of Thelon Formation deposition (Rainbird et al.

543 2003; Palmer et al. 2004) (Fig. 24 - epithermal faulting event). A strong silicification-brecciation

544 event at ca. 1750 Ma was similarly proposed by Grare et al. (2018; 2020).

545

546 7.4. Thelon Basin sedimentation (ca. 1740-1540 Ma) 547

23 © The Author(s) or their Institution(s) Canadian Journal of Earth Sciences Page 24 of 82

548 Sandstone and conglomerate of the Thelon Formation were deposited between ca. <1740 Ma

549 (Miller et al. 1989; Rainbird et al. 2003; Rainbird and Davis 2007) and ca. 1540 Ma, the age of the

550 overlying shoshonitic basalt of the Kuungmi Formation (Gall et al. 1992). This 200 m.y. time

551 interval represents a period of relative tectonic quiescence dominated by intracratonic basin

552 formation across the entire Rae-Hearne craton (Fig. 24 - Thelon Basin sedimentation). This

553 includes the Thelon Basin and other basins that are similar in age and stratigraphy such as the

554 Athabasca, Hornby Bay and Ulu basins (Rainbird et al. 2003). The Thelon and Judge Sissons faults

555 may have been reactivated during basin formation but any structures that formed during this event

556 would be masked by the strong alteration and the younger structures associated with the uranium

557 mineralizing event(s).

558 Draft 559 7.5. Late reactivation event(s) (ca. 1600-1300 Ma) 560

561 Late, post-Thelon reactivation of the Thelon and Judge Sissons faults created an extensive network

562 of non-cohesive fractures, breccias and gouge (Fig. 23e), which acted as pathways for the migration

563 of hydrothermal fluids. In the study area, the fluids reacted with the basement rocks to produce the

564 extensive and pervasive alteration halos of clay and specular hematite that surround the Tatiggaq,

565 Qavvik, and Ayra prospects (Fig. 23f). Hematization and clay alteration of the basement rocks due

566 to the breakdown of feldspars and micas to illite and chlorite are typically associated with uranium

567 mineralization in the Thelon Basin (Fuchs and Hilger 1989; Hasegawa et al. 1990; Sharpe 2013;

568 Riegler et al. 2014; Sharpe et al. 2015; Shabaga et al. 2017).

569

570 The late, post-Thelon reactivation of the Thelon and Judge Sissons faults occurred after the

571 deposition of the Thelon Formation at ca. <1740 Ma (Miller et al. 1989; Rainbird et al. 2003;

24 © The Author(s) or their Institution(s) Page 25 of 82 Canadian Journal of Earth Sciences

572 Rainbird and Davis 2007) and before the emplacement of the cross-cutting Mackenzie diabase

573 dikes at ca. 1267 Ma (LeCheminant and Heaman 1989; Fig. 24 – Late reactivation event(s).

574 Tighter constraints on the age of the reactivation and uranium mineralization events may be

575 provided by the dated monzogranite samples. Zircons from these two samples yielded younger

576 dates of ca. 1740 Ma to ca. 1710 Ma, ca. 1680 Ma to ca. 1640 Ma, and ca. 1400 Ma. The ca. 1740

577 Ma to ca. 1710 Ma dates coincide with the formation of the Thelon Basin (Miller et al. 1989;

578 Rainbird and Davis 2007). The ca. 1680 Ma to 1640 Ma dates overlap with the age of diagenesis

579 in the basin which is constrained by ca. 1650 Ma illite (Renac et al. 2002) and ca. 1667 Ma

580 diagenetic apatite (Davis et al. 2011) in non-mineralized sandstone of the Thelon Formation. The

581 ca. 1400 Ma dates are similar to that of fluorite-bearing veins in the Mallery Lake area (Fig. 2)

582 located approximately 50 km south of theDraft study area. These veins yielded an Sm-Nd errorchron

583 age of 1434 ±23 Ma (Turner et al. 2003). They host low-grade precious metal (Au-Ag)

584 mineralization, which has been attributed to Wharton Group magmatism (Turner 2000; Turner et

585 al. 2003), but is more akin to unconformity-type uranium mineralized systems given their age, and

586 the high salinity (23-31 wt.% NaCl equivalent) and low temperatures (90◦C to 150◦C) of the

587 mineralizing fluids. The fluid chemistry measured in the Mallery Lake veins is similar to those of

588 other unconformity-type uranium deposits associated with Proterozoic basins overlying the Rae-

589 Hearne craton, which have salinities upwards of 15-30 wt.% NaCl equivalent and temperatures

590 ranging from 120◦C to 240◦C (Kotzer and Kyser 1995; Kyser et al. 2000; Derome et al. 2005).

591

592 Altered rocks at the Kiggavik, Bong and Andrew Lake deposits (Fig. 3), returned younger K-Ar

593 and Ar-Ar illite ages ranging mostly between ca. 1386 Ma to 1311 Ma with a few younger ages

594 ranging from ca. 1179 to 932 Ma (Weyer 1982; Miller and LeCheminant 1986; Riegler 2013;

25 © The Author(s) or their Institution(s) Canadian Journal of Earth Sciences Page 26 of 82

595 Shabaga et al. 2017). U-Pb geochronology of pitchblende/uraninite from the Kiggavik, Bong, End

596 and Andrew Lake deposits yielded ages at ca. 1520 Ma, ca. 1400 Ma, ca. 1300 to 1000 Ma, and a

597 few very young ages at ca. 500 and 300 Ma (Farkas 1984; Fuchs and Hilger 1989; Riegler 2013;

598 Sharpe et al. 2015; Shabaga et al. 2017; Grare et al., 2020). The ca. 1400 Ma date from the

599 monzonite sample and the ca. 1434 Ma Sm-Nd age from the Mallery Lake area, in conjunction

600 with the partially overlapping pitchblende/uraninite and illite ages listed above, suggests that

601 uranium mobilization and reactivation of the Thelon and Judge Sissons faults in the Aberdeen

602 Lake area occurred as late as ca. 1300 Ma as previously reported by Grare et al. (2020).

603

604 Sharpe et al. (2015) reported an older U-Pb pitchblende age of 1520 Ma for mineralization at the

605 Bong deposits (Fig. 3). This age is similarDraft to that of an early, ca. 1590 Ma to ca. 1520 Ma, post-

606 diagenetic hydrothermal event in the Athabasca basin (McGill et al. 1993; Alexandre et al. 2007),

607 although most other Athabasca uraninite and illite dates are younger and fall between ca. 1460 Ma

608 and 1350 Ma (McGill et al. 1993; Fayak et al. 2002; Alexandre et al. 2007; Chernonozhkin et al.

609 2020). Collectively, this suggests that both the Athabasca and Thelon basins experienced major,

610 long-lived, uranium-mineralizing and fault-reactivation events between ca. 1600 Ma and 1300 Ma.

611

612 7.6. Other major faults of the Rae Domain 613

614 The deformation history of the Judge Sissons and Thelon faults is similar to that of other major ENE-

615 trending faults in the Rae domain, including the McDonald fault and Wager Bay shear zone (Figs.

616 1, 2). The McDonald fault is a dextral mylonite zone that cuts through ca. 1840 Ma plutons

617 (Bowring 1985; Hildebrand et al. 2010; Cook 2011; Hildebrand 2011) and was reactivated at ca.

618 1735 Ma as a brittle dextral transcurrent fault (Henderson et al. 1990). Previous interpretations

26 © The Author(s) or their Institution(s) Page 27 of 82 Canadian Journal of Earth Sciences

619 have suggested that the McDonald fault formed during indentation of the Rae domain by the Slave

620 craton during the Taltson-Thelon orogeny (Gibb 1978; Hoffman 1989; Henderson et al. 1990).

621 However early movement along the fault has also been attributed to accretion of the Hottah

622 Terrane, a Paleoproterozoic terrane west of the Slave Craton (Ootes et al. 2015), and to the

623 Wopmay orogeny at ca. 1875-1850 Ma along the western margin of the Slave craton (Hoffman

624 1980; Hildebrand et al. 1987; Cook 2011; Jackson et al. 2013). The Wager Bay shear zone is a 0.2

625 to 3.0 km wide, dextral, ductile shear zone that extends for roughly 450 km across the east-central

626 Rae domain (Henderson and Broome 1990; Therriault et al. 2017). The fault initially formed at ca.

627 1850 Ma (Corrigan et al. 2009) during collision of the Sugluk and Superior cratons with the Rae-

628 Hearne craton (Henderson and Broome 1990) and continued to be active throughout the remainder

629 of the Trans-Hudson orogeny from ca. 1830Draft Ma to ca. 1800 Ma (Corrigan et al. 2009; Wodicka et

630 al. 2017). In summary, the McDonald and Wager Bay faults formed and were reactivated in

631 response to the far-field stresses and bulk east-west shortening generated during the indentation of

632 the Slave craton into the Rae-Hearne craton to the west during the Taltson-Thelon orogeny and

633 collision and continued convergence with the Sugluk, Sask. and Superior cratons to the east during

634 the Trans-Hudson orogeny (Ashton and Hunter 2004; Ashton et al. 2004, 2009). The Thelon and

635 Judge Sissons faults have similar deformation histories but their more brittle behavior suggests

636 that they formed at shallower crustal depths. However, it is likely that had these faults been

637 exhumed from greater depths, a more ductile structural character would be observed.

638

639 7.7. Comparisons with large Cenozoic strike-slip fault and fluid flow systems 640

641 Active strike-slip faults form in several different tectonic settings. The Karakorum and the Altyn

642 Tagh faults are located along the northwest margin of the Tibetian Plateau in southwest Asia. They

27 © The Author(s) or their Institution(s) Canadian Journal of Earth Sciences Page 28 of 82

643 are excellent examples of long-lived, crustal-scale, strike-slip faults, which formed during

644 continental collision. They have extensive strike-lengths (up to 2500 km) and significant offsets of

645 100 km and more and have been active over the last 15-50 m.y. during the Himalayan-Tibetan

646 orogeny (Murphy et al. 2000; Yin et al. 2002; Lacassin et al. 2004; Rutter et al. 2007; Cheng et al.

647 2016) and possibly underwent movement as far back as 250 Ma ago (Wang et al. 2005). The San

648 Andreas fault in California is another example of a long-lived, crustal-scale, strike-slip faults,

649 which acts as a transform fault between the Pacific plate and the south-migrating North American

650 plate. The fault has been active for 23-38 Ma (Powell 1993; Simms 1993; Sharman et al. 2013) for

651 a total offset of roughly 100 km. Another example is the Median Tectonic Line in Japan which

652 differs from the Karakorum, Altyn Tagh and San Andreas faults by its island arc setting. It formed

653 above a west-dipping subduction zoneDraft to accommodate the arc-parallel component of oblique

654 convergence between the arc and the subducting plate. It can be traced over >1000 km and has

655 been active as far back as the Cretaceous (Jeffries et al. 2006). Regardless of their tectonic setting,

656 these major strike-slip faults have been active over several millions of years and formed to

657 accommodate the lateral movement of tectonic plates or the lateral flow of rocks parallel to

658 collisional orogenic fronts.

659

660 Major faults in the upper crust either form conduits or barriers to fluid flow (Caine et al. 1996;

661 Gudmundsson 2000; Bense et al. 2013). The movement of fluids along faults is facilitated by

662 processes that increase permeability, such as fracturing and brecciation, and is hindered by

663 processes that reduce permeability, such as cataclasis, compaction, cementation, and alteration of

664 the host rocks to clay minerals (Bense et al. 2013). At depth, veins provide clues on the

665 passage of fluids along faults and on the faulting processes that affected the flow of these fluids

28 © The Author(s) or their Institution(s) Page 29 of 82 Canadian Journal of Earth Sciences

666 (Oliver and Bons 2001). At surface, fluid flow is manifested by hot springs punctuating the faults

667 at discharge areas (Forster et al. 1997). For example, hot spring activity is reported in Asia along

668 the strike-slip Hunan-Jaingxi fault, which accommodated the lateral escape of crustal slices away

669 from the colliding India and Eurasian plates (Li et al. 2001, 2002). Other examples include

670 crustal-scale brittle faults in the Canadian Cordillera, which acted as conduits for the downward

671 flow and penetration of cool meteoric water into the crust, but also allowed for hotter, deep-seated

672 lower crustal to mantle fluids to move up from below without requiring high-heat flow or a high-

673 level magmatic intrusion (Grasby and Hutcheon 2001). Similarly, the Alpine fault in New Zealand

674 allowed mantle fluids to migrate upwards through the crust (Menzies et al. 2016) as did the San

675 Andreas fault in the southwest U.S.A. as suggested by helium isotope values from hot springs

676 along the fault (Kennedy et al. 1997). TheseDraft examples reveal a strong association between deep

677 crustal faults and prolonged hydrothermal activity (Hochstein and Regenauer-Lieb 1998).

678 Injections of fluids may trigger seismic events (Cox 2016) and may contribute to the migration of

679 hydrothermal fluids from the deep crust to the surface. Pulsatory fluid regimes and deep

680 convection of brines downward into fault zones is interpreted as an important factor for the

681 formation of large U, Pb-Zn and F-Ba deposits at or near basement/cover unconformities (Boiron

682 et al. 2010).

683

684 The Thelon and Judge Sissons faults are Precambrian examples of major, long-lived, strike-slip

685 faults with upper crustal level hydrothermal activity. Dextral movement along the Thelon and

686 Judge Sissons faults began after ca. 1830 Ma, the age of the Hudson Intrusive Suite, continued

687 during Wharton Group magmatism at ca. 1760 Ma, and ended after the emplacement of the ca.

688 1750 Ma McRae Lake dikes. The latter are dextrally offset along the two faults, although less so

29 © The Author(s) or their Institution(s) Canadian Journal of Earth Sciences Page 30 of 82

689 than the Hudsonian pluton that is cut by the faults, suggesting that dextral movement along the two

690 faults continued after the emplacement of the dikes. The initial faulting event formed in

691 response to protracted northwest-directed shortening in the hinterland between the Taltson-

692 Thelon and the Trans-Hudson orogenies, and thus shares similarities to the Karakorum and

693 Altyn Tagh strike-slip faults, which formed in the hinterland of the Himalayan-Tibetan

694 orogen. During Wharton Group magmatism at ca. 1760 Ma to ca. 1750 Ma, hydrothermal fluid

695 activity at upper crustal levels led to the formation of comb, crustiform-cockade and bladed

696 textured veins. Erosion of the basement rocks and faults between ca. 1720 Ma and ca. 1540 Ma

697 shed detritus in the newly-formed Thelon Basin into which thick sequences of sandstone and

698 conglomerate accumulated and were later capped by thin deposits of shoshonitic basalt (Kuungmi

699 Formation) and stromatolite-bearing marineDraft dolomite and evaporite (Lookout Point Formation)

700 (Hiatt et al. 2003). Similar intracratonic basins such as the Athabasca, Hornby Bay and Ulu basins

701 were also deposited throughout northern Laurentia at roughly the same time (Rainbird et al. 2003;

702 Jefferson et al. 2007). These basins formed as the southern margin of the continent was under

703 compression during the ca. 1780 Ma to ca. 1650 Ma Yavapai and Mazatzal orogenies (Johansson

704 2009; Furlanetto 2015). A few million years later, the ca. 1640 Ma Wernecke basin formed during

705 rifting of the northwest margin of Laurentia (Thorkelson et al. 2001) and was deformed during the

706 Racklan-Forward orogeny at ca. 1640 Ma to ca. 1600 Ma (Brideau et al. 2002; Thorkelson et al.

707 2003, 2005; Laughton et al. 2005). Reactivation of the Thelon and Judge Sissons faults and the

708 emplacement of uranium mineralization was likely coeval with this orogenic event and with later

709 continent-wide extension of Laurentian (Anderson 1983; van Breemen and Davidson 1988;

710 Windley 1993; Frost et al. 2001), as hypothesized by the deposition of the Belt-Purcell Supergroup

711 in a passive rift setting between ca. 1470 Ma and ca. 1445 Ma (Chandler 2000; Price and Sears

30 © The Author(s) or their Institution(s) Page 31 of 82 Canadian Journal of Earth Sciences

712 2000) and the emplacement of voluminous suites of ca. 1400 Ma granitoid plutons from southern

713 California to Labrador, Canada (Anderson 1983; Windley 1993). The drivers of the late

714 reactivation and uranium mineralization fluid flow event(s) are interpreted as one or several of

715 these compressional to extensional events that post-date the deposition of the Thelon Formation.

716 The duration of fault movement and reactivation along the Thelon and Judge Sissons faults

717 exceeds the known movement histories of modern crustal faults such as the Karakorum and Altyn

718 Tagh faults. However, these faults will remain active for many years and even millions of years to

719 come and so are excellent modern analogues for the Thelon and Judge Sissons faults.

720

721 8. Conclusions 722 Draft 723 The Thelon and Judge Sissons faults are long-lived Paleoproterozoic to Mesoproterozoic faults with

724 complex reactivation and hydrothermal histories. The faults initially formed at <1830 Ma as dextral

725 transcurrent faults. They were reactivated at ca. 1760 to ca. 1750 Ma during the migration of

726 hydrothermal fluids to upper crustal levels and between ca. 1600 Ma and ca. 1300 Ma during the

727 emplacement of uranium mineralization during reverse or normal movement along the faults. The

728 structural history of the Thelon and Judge Sissons faults is similar to that of other large dextral

729 strike-slip faults in the Rae craton (e.g. McDonald and Wager Bay fault zones), which formed in

730 the hinterland of the Taltson-Thelon Orogen to the west and the Trans-Hudson Orogen to the east,

731 respectively. It is also broadly similar to that of the Karakorum and Altyn Tagh faults in Tibet, which

732 formed in the hinterland of the Himalayan-Tibetan Orogen during the collision of India with

733 Eurasia.

734

735 Large crustal-scale faults are commonly the loci of active hot springs produced by the discharge

31 © The Author(s) or their Institution(s) Canadian Journal of Earth Sciences Page 32 of 82

736 of hydrothermal fluids at surface. Along the Thelon and Judge Sissons faults, breccias and veins

737 with comb, crustiform, cockade and bladed textures, formed due to the discharge of hydrothermal

738 fluids possibly heated at depth by Wharton Group magmas. As with other large crustal-scale faults,

739 the fluids may have reached the surface to form hot springs. Reactivation of the two faults during

740 far-field accretionary events at ca. 1600 Ma and during Laurentia-wide extensional events at ca.

741 1500 Ma to ca. 1300 Ma channeled the flow of new pulses of hydrothermal fluids resulting in the

742 formation of unconformity-type uranium deposits.

743 744 Acknowledgements 745

746 Field and analytical work was funded by Cameco Corporation. Laurentian University is thanked

747 for providing support and use of their analyticalDraft facilities. We appreciate the help of Andy DuFrane,

748 Barry Herchuk, and Judy Schultz with U-Pb sample preparation and analyses at the University of

749 Alberta. Many thanks to Joel Lesperance, Ekaterina Savinova, Erika Bourlon and Lynde

750 Guillaume for cheerful field assistance and data collection. Dan Kontak is thanked for his help

751 collecting the cathodoluminescence imagery. Thoughtful reviews by Toby Rivers, Brendan

752 Murphy, Gerard Zaluski, and an anonymous reviewer greatly improved the quality of the

753 manuscript.

754

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1265 Thorkelson, D.J., Abbott, J.G., Mortensen, J.K., Creaser, R.A., Villeneuve, M.E., McNicoll, V.J., 1266 and Layer, P.W. 2005. Early Draft and Middle Proterozoic evolution of Yukon, Canada. 1267 Canadian Journal of Earth Sciences, 42: 223–239.

1268 Tippett, C.R., and Heywood, W.W. 1978. Stratigraphy and structure of the Northern Amer Group

1269 (Aphebian), Churchill Structural Province, District of Keewatin. Geological Survey of

1270 Canada Paper 78-1B, 7–11. doi:10.4095/103567.

1271 Tschirhart, V., Morris, W.A., and Oneschuk, D. 2011. Geophysical Series, Geophysical

1272 Compilation Project, Thelon Basin, Nunavut, NTS 66A, B and parts of 65N, O, P, 66C, F,

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1274 Turner, W. 2000. Geology and geochemistry of the Mallory Lake precious metal-bearing epithermal

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1319 doi:10.4095/211098.

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Fig. 1. Map of the regional geology of the western Churchill Province and bounding orogenic belts. The square represents the study area and the location of the Aberdeen Lake supracrustal belt. Abbreviations: ALb: Aberdeen Lake supracrustal belt, ASZ: Amer shear zone, BCB: Barclay belt, BF: Bathurst fault, CB: Chesterfield Block, CBb: Committee Bay belt, CSZ: Chantrey shear zone, CRS & SRG: Clarke River schist and Snow River gneiss (Dubawnt Lake area), ERGB: Ennadai-Rankin greenstone belt; GRSZ: Grease River shear zone, MBg: Murmac Bay group, MF: McDonald fault, MRg: Mary River group, PAb: Prince Albert belt, Pg: Penrhyn Group, Plg: Piling Group, STZ: Snowbird Tectonic Zone, TSZ: Tyrell shear zone, TTO: Taltson- Thelon Orogen, WBSZ: Wager Bay shear zone, Wg: Wollaston Group, and WLg: Woodburn Lake group. Red X’s represent the Meadowbank gold deposit and the Kiggavik uranium deposit. Map modified after Sanborn- Barrie et al. (2002), Hartlaub et al. (2004), and Pehrsson et al. (2013). CorelDraw®2017 (Version 19.1.0.419) was used in the preparation of this figure.

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Fig. 2. Annotated First Vertical Derivative of Magnetic Anomalies image modified from Miles and Oneschuk (2016) showing the location of the Thelon Basin and regional fault zones, including the interpreted traces of the Thelon and Judge Sissons faults. The location of the study area is represented by the red box. ESRI ArcGIS (Version 10.6) was used in the preparation of this figure.

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Fig. 3. Simplified geological map of the Aberdeen Lake area from the first author’s mapping and others at Cameco Corporation. Geology near Judge Sissons Lake is modified after Hadlari et al. (2004) and Jefferson et al. (2011). Shown on the map as a heavy black broken line is the location of a cross section A-A’ across folded and thrusted Archean supracrustal rocks. Abbreviations: T - Tatiggaq; Q – Qavvik; A - Ayra. Thin red marker units in the Middle and Upper sequences are banded iron formation. Map created with UTM coordinates using datum WGS84, Zone 14N. ESRI ArcGIS (Version 10.6) was used in the preparation of this figure.

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Fig. 4. Field photographs of rock units in the Aberdeen Lake area. (a) Quartzofeldspathic orthogneiss with strong gneissic banding (UTM 501187E, 7126187N). (b) Coarse-grained massive monzogranite (UTM 554414E, 7139323N). (c) Medium-grained massive hornblende-bearing mafic syenite (UTM 553984E, 7139065N). (d) Massive, megacrystic hornblende syenite (UTM 554414E, 7139323N). (e) Lamprophyre dike cutting across hornblende syenite of the Hudson Intrusive Suite, photograph taken looking east-northeast (UTM 554002E, 7137653N). (f) Coarse-grained McRae Lake gabbroic dike (UTM 536101E, 7141484N). Photo locations as UTM coordinates using datum WGS84, Zone 14N.

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Fig. 5. Aberdeen Lake area structures. (a) Isoclinal F2 folds refolding rare F1 isoclinal folds and associated S1 gneissic foliation. (UTM 505236E; 7122110N). (b) Close-up of refolded F1 isoclinal fold in Figure 4a outlined by red rectangle. (c) S2 cleavage defined by aligned volcaniclastic clasts and oriented oblique to the S0/S1 cleavage in the Lower Sequence (UTM 536054E; 7141708N). (d) S1 foliation folded by an F2 fold and overprinted by an axial planar S2 cleavage (UTM 506025E; 7121742N). (e) Isoclinal F2 folds defined by folded amphibolite layer of the Lower Sequence overprinted by angular to chevron-style F3 folds (UTM 536269E; 7141700N). (f) S2 fabric with garnet porphyroblasts folded by a gentle F4 fold with an axial plane S4 cleavage (UTM 534702E; 7136213N). Photo locations as UTM coordinates using datum WGS84, Zone 14N.

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Fig. 6. Geological map of the Marjorie Lake area. Location of map and map legend is shown in Fig. 3. Lower hemisphere, equal-area stereonet diagrams of poles to foliations and L2 lineation: S2 - contour interval in 2% of total measurements per 1% area of net, L2 - contour interval in 5% of total measurements per 1% area of net. ESRI ArcGIS (Version 10.6) and CorelDraw®2017 (Version 19.1.0.419) were used in the preparation of this figure.

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Fig. 7. Annotated Residual Total Magnetic Field map of the area shown in Fig. 3, showing the traces of the fault zones, plutons, McRae Lake dikes and geochronological samples (yellow stars). Magnetic map is modified from Tschirhart et al. (2011). Abbreviations: HIS – Hudson Intrusive Suite. ESRI ArcGIS (Version 10.6) and CorelDraw®2017 (Version 19.1.0.419) were used in the preparation of this figure.

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Fig. 8. Representative outcrop map showing part of the Judge Sissons fault with core zone defined by breccia with hematite-altered matrix and damage zone defined by quartz veins offset by sinistral fractures in hematite-altered wall rocks (UTM 559724E; 7136841N; WGS84, Zone 14N). Inset shows idealized Riedel shear network associated with a NW-directed σ1. CorelDraw®2017 (Version 19.1.0.419) was used in the preparation of this figure.

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Fig. 9. Uni-directional rose diagrams and equal area, lower hemisphere, contoured stereonet diagrams of fracture and vein orientations in the damage zones of the Thelon and Judge Sissons faults in the study area. The contoured stereonet diagrams were created using the Kamb contouring method (Kamb 1959) with an inverse area square smoothing technique from Vollmer (1995). Contour intervals are in 3 and 5 sigma standard deviations for the fractures and veins, respectively. (a) Rose diagram of fracture orientations with one orientation group containing 12% of all measurements. (b) Contoured stereonet diagram of poles to fractures. (c) Rose diagram of vein orientations with one orientation group containing 32% of all measurements. (d) Contoured stereonet diagram of poles to veins.

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Fig. 10. Photographs of quartz veins in the damage zone of the Thelon and Judge Sissons faults. (a) Cross- polarized light photomicrograph of quartz grains with growth zoning defined by inclusion-rich bands (UTM 559717E; 7136849N). (b) Field photograph of hydrothermal quartz crackle breccia (Q) cutting across strongly hematite-altered Archean supracrustal rocks with sporadic quartz veining (WR) (UTM 559710E; 7136843N). Lichen is denoted with LN. Pencil (15 cm in length) for scale. (c) Rounded clasts of Ketyet River Group quartzite (Qzite), hydrothermal quartz crackle breccia (Q), and quartz-veined hematite-altered Archean supracrustal rocks (WR), within brecciated Thelon Formation basal conglomerate. The red matrix consists of coarse-grained sandstone with small pebble-sized clasts (UTM 569010E; 7150698N). Pencil (15 cm in length) for scale. (d) Close-up of mosaic quartz breccia (Q) clast in (c). Cap of pencil (1 cm) for scale. Photo locations were recorded in UTM coordinates using datum WGS84, Zone 14N.

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Fig. 11. Photographs of breccia in the core zone of the Judge Sissons fault. (a) Field photograph of matrix- supported mesocataclasite with clasts (Q – quartz vein; WR – hematite-altered wall rock) varying in size from 2-7 mm (UTM 559482E; 7136735N). LN refers to lichen. (b) Photograph of a 2-3 cm wide hematized mesocataclasite (outlined by the medium dashed line) surrounded by hematite-altered chaotic and mosaic breccias. Ghost quartz veins (Q) outlined by solid black lines are brecciated and weakly rotated clockwise close to the mesocatalasite transecting the centre of the photo. NW-trending sinistral fractures (blue dotted lines) cut across the quartz veins (UTM 559696E; 7136817N). Pencil exposed over 5 cm for scale. Photo locations in UTM coordinates using datum WGS84, Zone 14N.

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Fig. 12. Field photograph and stereonet diagram of slickenlines and slip surfaces measured along the Thelon and Judge Sissons faults. (a) Hematite-altered ultracataclasite with well-developed trailed material slickenlines indicative of dextral displacement. Arrow denotes the movement direction of the missing hanging wall (UTM 560108E; 7147700N). Pencil cap is 0.8 cm wide. (b) Equal-area stereonet diagrams of fault slip surfaces plotted as great circles with rakes of slickenlines plotted as close circles. Arrows indicate the sense of movement. Photo locations in UTM coordinates using datum WGS84, Zone 14N. 188x74mmDraft (300 x 300 DPI)

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Fig. 13. Representative field and drill core photographs of epithermal quartz veins. (a) Coarse, prismatic comb-textured quartz crystals (hexagonal in cross-section) in vein cutting across hematite-altered breccia in drill core (DDH LOB-002, UTM 542032E; 7134163N, 223 m depth). (b) Quartz vein with bladed-textured quartz in the damage zone of the Judge Sissons fault (UTM 559488E; 7136745N). (c) Vein with fine-grained comb- and bladed-textured quartz and hematite along its margins and layered crustiform-banded quartz along its centre. Wall rock consists of hematite-altered ultracataclasite (DDH AYA-004, UTM 527833E; 7131730 N, 103 m depth). (d) Cockade-textured quartz vein with irregular crustiform quartz surrounding wall rock and quartz vein fragments (outlined by dotted line) along the Judge Sissons fault (UTM 559488E; 7136745N). (e) Angular clast of comb-textured quartz vein within strongly hematite altered Thelon Formation conglomerate (DDH: ABR-009B, 76 m; UTM 528044E; 7131844N). (f) Rounded clasts of hematite-altered rhyodacite of the Pitz Formation and white to clear comb- to bladed-textured quartz vein in clasts of clay-altered coarse-grained sandstone and conglomerate of the Thelon Formation (DDH ABR-009B, 69.8 m; UTM 528044E; 7131844N). Pencil (8 mm in width) for scale; LN = lichen; Qzite = Ketyet River group quartzite. Photo locations in UTM coordinates using datum WGS84, Zone 14N.

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Fig. 14. Drill core photographs of late reactivation fault structures along the Thelon and Judge Sissons faults. (a) Clay-altered gouge along margin of clay-altered mesocataclasite (DDH TUR-026, 75.3 m, UTM 548966E; 7135453N). (b) Non-cohesive, clay-altered breccia in Thelon Formation sandstone. Dotted lines outline some of the clasts (DDH ABR-004, 35.5 m, UTM 528036E; 7131868N). (c) Irregular cross-cutting specular hematite veins (DDH TUR-046, 154.1 m, UTM 549472E; 7135559N). (d) Irregular and discontinuous clay veins (DDH TUR-060, 127.8 m, UTM 548898E; 7135824N). Photo locations in UTM coordinates using datum WGS84, Zone 14N.

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Fig. 15. Drill core photographs of fracture surfaces with slickenlines in clay-altered units along the Judge Sissons fault zone. (a) DDH AYA-004, 36.4 m - steep dip-slip slickenlines along a thin gouge-filled fracture in clay- and hematite-altered sandstone of the Thelon Formation. (b) DDH AYA-004, 37.5 m - dip-slip slickenlines along limonite- and clay-altered fracture in Thelon Formation sandstone. AYA-004 drill hole location: UTM 527833 E; 7131731N. Photo locations in UTM coordinates using datum WGS84, Zone 14N.

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Fig. 16. Uni-directional rose diagrams and equal area, lower hemisphere contoured stereonet diagrams of structures related to the late reactivation event(s). The contoured stereonet diagrams were created using the Kamb contouring method (Kamb 1959) with an inverse area square smoothing technique from Vollmer (1995). Contour intervals are in 1 sigma standard deviations. (a) Rose diagram of gouge orientations with one orientation group representing 12% of all measurements. (b) Contoured stereonet diagram of poles to gouge structures. (c) Rose diagram of clay vein orientations with one orientation group representing 15% of all measurements. (d) Contoured stereonet diagram of poles to clay veins.

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Fig. 17. Drill core photographs of uranium mineralization within clay-altered and non-cohesively fractured and brecciated Archean supracrustal rocks. (a) Fracture-controlled and replacement pitchblende mineralization within clay-altered psammopelite gneiss. Assay values in vicinity of photograph = 3.17% U3O8 over 0.5 m (DDH SAN-006, 288 m, UTM 533545E; 7135796N). (b) Massive and fracture-controlled pitchblende mineralization in strongly clay altered and brecciated psammopelite gneiss. Assay values in vicinity of photograph = 3.01% U3O8 over 0.3 m (DDH TUR-017, 135 m, UTM 548875E, 7135500N). Photo locations in UTM coordinates using datum WGS84, Zone 14N. 217x83mmDraft (300 x 300 DPI)

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Fig. 18. NW-SE cross-section across the Tatiggaq prospect showing the geologic setting of the uranium mineralization and the surrounding alteration halo. See Fig. 3 for location of the Tatiggaq prospect.

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Fig. 19. Representative cathodoluminescence images of zircons from samples dated by Laser Ablation ICP- MS U-Pb geochronology. Ablation pits are outlined with yellow circle. 207Pb/206Pb dates for ablated material are reported on photographs and zircon number corresponds to data in Supplementary Table S1. (a) 204961- semi-prismatic grain with coarse unzoned core and weak outer zonation with black overgrowths. (b) 204974 - prismatic and strongly zoned zircon with irregular core region. (c) 204974 - rounded xenocrystic core with irregular zonation and strongly zoned overgrowths.

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Fig. 20. Laser Ablation ICP-MS U-Pb zircon Concordia diagram for Sample 204961 - quartzofeldspathic gneiss. Data point error ellipses are 2σ.

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Fig. 21. Laser Ablation ICP-MS U-Pb zircon Concordia diagrams for Hudson Intrusive Suite (HIS) monzogranite samples. Data point error ellipses are 2σ. (a) 204964 (monzogranite - HIS). (b) 204974 (monzogranite - HIS). Close-up plots are from the green circle areas on the first two Concordia diagrams and highlight the younger zircon dates for each sample.

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Fig. 22. Cathodoluminescence images of zircons from monzogranite sample 204974 displaying unusual non- fluorescent rim zonation patterns and < ca. 1800 Ma 207Pb/206Pb dates. Ablation pits are outlined with yellow circles. 207Pb/206Pb dates for ablated material are reported on photographs and zircon number corresponds to data in Supplementary Table S1. (a) Irregular zircon with large non-fluorescent zone cutting across an early zircon core domain. The non-fluorescent zone yielded a 207Pb/206Pb date of ca. 1395 Ma. (b) Zircon with unusual and irregular non-fluorescent growth zones that returned a 207Pb/206Pb date of ca. 1640 Ma. (c) Long prismatic and zoned zircon with thin non-fluorescent patches with a 207Pb/206Pb date of ca. 1770 Ma. 174x56mmDraft (300 x 300 DPI)

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Fig. 23. Paragenetic model of the different quartz and breccia events related to movement along the Thelon and Judge Sissons faults. (a) Initial faulting event - early fracturing and formation of the damage zone. (b) Initial faulting event- vein infill of early fractures, early pervasive maroon hematite alteration and subsequent formation of quartz hydrothermal crackle breccias. (c) Initial faulting event - ongoing fault movement led to the development of the core zone protocataclasites and mesocataclasites. (d) Epithermal faulting event - quartz veining event with comb, crustiform, cockade and bladed textures. (e) Late reactivation event - creation of non-cohesive fractures, breccia and gouge along the margins of and nearby older fault structures. (f) Late reactivation event - emplacement of specular hematite (HS) and clay veins in outer distal alteration envelope related to uranium mineralization.

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Fig. 24. Synopsis of U-Pb zircon age and events for the Aberdeen Lake area.

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