geology and age of the morrison porphyry cu-au-mo 1 1 geology and age of the morrison porphyry...

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Geology and age of the Morrison porphyry Cu-Au-Mo

deposit, Babine Lake area, British Columbia

Journal: Canadian Journal of Earth Sciences

Manuscript ID cjes-2015-0231.R1

Manuscript Type: Article

Date Submitted by the Author: 21-Mar-2016

Complete List of Authors: Liu, Lijuan; University of Alberta, Earth and Atmospheric Sciences Richards, Jeremy P.; University of Alberta, Earth and Atmospheric Sciences Creaser, Robert; University of Alberta, Earth and Atmospheric Sciences DuFrane, S. Andrew; University of Alberta, Earth & Atmospheric Sciences Muehlenbachs, Karlis; University of Alberta, Earth and Atmospheric Sciences

Larson, Peter; Washington State University, School of the Environment

Keyword: Porphyry Cu-Au-Mo deposit, Eocene, Babine Lake, British Columbia

https://mc06.manuscriptcentral.com/cjes-pubs

Canadian Journal of Earth Sciences

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1

Geology and age of the Morrison porphyry Cu-Au-Mo deposit, Babine Lake 1

area, British Columbia 2

3

Lijuan Liu1, Jeremy P. Richards

2*, Robert A. Creaser

3, S. Andrew DuFrane

4, Karlis 4

Muehlenbachs5, and Peter B. Larson

6 5

6

1. Dept. Earth and Atmospheric Sciences, University of Alberta 7

Edmonton, Alberta, Canada, T6G 2E3, [email protected] 8

9



12



15



18



21

6. Washington State University, School of the Environment, Pullman, WA, USA, 99164-2812, 22

[email protected] 23

24

* Corresponding author: Jeremy P. Richards, Dept. Earth and Atmospheric Sciences 25

Earth Sciences Building, Rm. 3-02, University of Alberta, Edmonton, Alberta, Canada, T6G 2E3 26

Tel: 780-492-3430 27

Fax: 780-492-2030 28

E-mail: [email protected]

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Geology and age of the Morrison porphyry Cu-Au-Mo deposit, Babine Lake area, British 31

Columbia 32

33

Lijuan Liu, Jeremy P. Richards, Robert A. Creaser, S. Andrew DuFrane, Karlis Muehlenbachs, 34

and Peter B. Larson 35

36

Abstract 37

The Morrison porphyry Cu-Au-Mo deposit is genetically and spatially related to Eocene 38

plagioclase-hornblende-biotite porphyry intrusions. One porphyry intrusion yielded a U-Pb age 39

of 52.54 ± 1.05 Ma. Mineralization occurs in three stages: (1) vein-type and disseminated 40

chalcopyrite and minor bornite (associated with potassic alteration and gold mineralization); (2) 41

vein-type molybdenite (associated with weak phyllic alteration); and (3) polymetallic sulfide-42

carbonate veins (dolomite ± quartz-sphalerite-galena-arsenopyrite-chalcopyrite, associated with 43

weak sericite-carbonate alteration). Re-Os dating of molybdenite yielded ages of 52.54 ± 0.22 44

and 53.06 ± 0.22 Ma, similar to the age of the host porphyry intrusion. 45

Stage 1 vein fluids were predominantly magmatic origin: Th = 400° to 526°C; salinity = 39.8 46

to 47.8 wt.% NaCl equiv.; δ18

Ofluid = 3.7 to 6.3‰; disseminated chalcopyrite-pyrite δ34

SCDT = 0.2 47

and -0.8‰. Stage 2 fluids were a mixture of magmatic and meteoric water: Th = 320° to 421°C; 48

salinity = 37.0 to 43.1 wt.% NaCl equiv.; δ18

Ofluid values range from 0.3 to 3.4‰; molybdenite 49

and pyrite δ34

SCDT = -2.1 and -1.2‰. Stage 3 fluids were predominantly of meteoric water origin: 50

Th = 163° to 218°C; salinity = 3.1 to 3.9 wt.% NaCl equiv.; δ18

Ofluid = -2.3 to 3.9‰ for early 51

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vein quartz, and 1.1 to 6.1‰ for late vein dolomite; sphalerite and pyrite δ34

SCDT = -7.1 to -52

5.6‰. 53

Morrison is interpreted to be a typical porphyry Cu-Au-Mo deposit related to a calc-alkaline to a 54

high-K calc-alkaline diorite to granodiorite intrusive suite, generated in a continental arc in 55

response to early Eocene subduction of the Kula-Farallon plate beneath North America. 56

Keywords: Porphyry Cu-Au-Mo deposit; Eocene; Babine Lake; British Columbia. 57

58

Introduction 59

The Morrison porphyry deposit is located in the northern Babine Lake area of central British 60

Columbia (Fig. 1). Several porphyry copper deposits related to early Cenozoic dioritic to granitic 61

porphyry intrusions occur in this area, including the Bell, Granisle, Nakinilerak, Dorothy, North 62

Newman, South Newman, Hearne Hill, and Morrison deposits (Fig. 2; Carson and Jambor 1974; 63

Carter 1982; Zaluski et al. 1994; Nokleberg et al. 2005). However, only three of these deposits 64

contain resources that are considered economic or close to economic: the Bell (77.2 million 65

tonnes @ 0.48% Cu), Granisle (52.7 million tonnes @ 0.43% Cu), and Morrison (207 million 66

tonnes @ 0.39% Cu) (Carson and Jambor 1974; Carson and Jambor 1976; Fahrni et al. 1976; 67

Dirom et al. 1995; Simpson and Geo 2007). The Bell and Granisle deposits were important past 68

Cu-Au producers but both closed in 1992, whereas the Morrison deposit has not yet been mined 69

(Carter et al. 1995; Simpson and Geo 2007). 70

The Morrison porphyry Cu-Au-Mo deposit is located 65 km northeast of Smithers in central 71

British Columbia (55°11'N, 126°18'W), and is spatially associated with Eocene dioritic to 72

granodioritic plagioclase-hornblende-biotite porphyry stocks and dikes, which intruded siltstones 73

and greywackes of the Middle to Upper Jurassic Ashman Formation of the Bowser Lake Group 74

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(Carson and Jambor 1976; Simpson and Geo 2007). Noranda Mines Limited first discovered the 75

deposit in 1963 during a stream sediment sampling program (Simpson and Geo 2007). Follow-up 76

exploration from 1963–1973 involved geological mapping and diamond core drilling (95 holes 77

with a total length of 13,890 m; Carson and Jambor 1976; Simpson and Geo 2007). Two 78

mineralized zones were outlined northwest and southeast of a small central lake, which are 79

interpreted to be offset by a fault (the East Fault in Fig. 3). Pacific Booker Minerals acquired the 80

Morrison property in 1997, and drilled a further 96 holes (26,202 m) by 2007 (Simpson and Geo 81

2007). This work has defined a measured and indicated resource of 207 Mt with average grades 82

of 0.39% Cu, 0.2 g/t Au, and 0.005% Mo (0.3% Cu equivalent cut-off; Simpson and Geo 2007). 83

Previous research at Morrison includes an early description by Carson and Jambor (1976), 84

and geochronological investigations that reported a K-Ar age of 52.1 ± 2.1 Ma for a porphyry 85

intrusion (Carter 1982), and a similar but more precise 40

Ar/39

Ar age of 53.2 ± 0.5 Ma 86

(Maclntyre et al. 2001). 87

In this study we present new petrological and whole-rock geochemical data for the porphyry 88

intrusions associated with the deposit, geochronological data for one of the intrusions (zircon, U-89

Pb) and hydrothermal molybdenite (Re-Os), and stable isotope and fluid inclusion data from 90

different stages of mineralized veins. The objective of this work is to place the Morrison 91

porphyry system in the context of other Eocene porphyry systems in central British Columbia 92

(related to the Babine Lake plutonic suite), and to determine the mode of ore formation. 93

94

Tectonic Setting 95

The Morrison deposit is located in the central Stikinia terrane of the Canadian Cordillera (Fig. 96

1). The Cordillera comprises a number of allochthonous island arcs, oceanic terranes, and 97

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pericratonic terranes, which accreted to the western margin of the North American craton during 98

the Middle Jurassic to late Mesozoic (Monger 1977; Monger and Irving 1980; Monger et al. 99

1982; Gabrielse et al. 1991; Monger and Price 2002; Nelson and Colpron 2007). As one of the 100

allochthonous island arcs, the Stikinia terrane hosts a significant number of porphyry copper 101

deposits, which formed both prior to and after the accretion events. 102

The Stikinia island arc was initiated in the Late Devonian, in response to subduction of the 103

Panthalassa oceanic plate beneath ancestral North America (Nelson and Colpron 2007; Logan 104

and Mihalynuk 2014). The arc was separated from ancestral North America by a back-arc basin 105

(Nelson and Colpron 2007). Growth of the arc ceased during the Late Permian to Middle Triassic 106

following collision with the Kutcho island arc, which caused uplift and erosion (the Tahltanian 107

orogeny; Souther 1971, 1972; English et al. 2003; Logan and Mihalynuk 2014). Arc construction 108

recommenced in the Late Triassic with deposition of new volcanic and sedimentary rocks. 109

Consequently, the Stikinia terrane consists of a Devonian to Permian assemblage of volcanic 110

rocks and carbonate sedimentary rocks (named the Stikinia assemblage), overlain by a Triassic to 111

Jurassic sequence of volcanic and associated sedimentary rocks (the Late Triassic Takla Group, 112

and the latest Triassic to Middle Jurassic Hazelton Group in the Babine Lake area; Massey et al. 113

2005; Logan and Mihalynuk 2014; Barresi et al. 2015). A suite of Late Triassic to Middle 114

Jurassic intrusive rocks is associated with an early pulse of porphyry Cu mineralization in the 115

Stikinia terrane (McMillan et al. 1995; Barresi et al. 2015). 116

Several island arc terranes including Stikinia, Cache Creek, and Quesnellia collided with the 117

continental margin in the Middle Jurassic, resulting in closure of the middle Paleozoic-Mesozoic 118

back-arc basin (Monger and Price 2002; Nelson and Colpron 2007). The Stikinia terrane then 119

became part of a continental arc, with renewed deposition of volcanic and sedimentary rocks, 120

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including the Jurassic to Cretaceous Bowser Lake Group, Lower Cretaceous Skeena Group, 121

Cretaceous to Tertiary Sustut Group, and the Paleogene Ootsa Lake and Endako Groups in the 122

Babine Lake area (Fig. 2; Massey et al. 2005). The tectonic regime of the Cordillera at this time 123

(Early Jurassic to late Paleocene) was characterized by compression and transpression, but this 124

changed to extension and transtension during the late Paleocene to Eocene (Parrish et al. 1988; 125

Dostal et al. 2001). 126

Late Cretaceous to early Cenozoic intermediate-to-felsic plutons intruded the Mesozoic 127

volcanic and sedimentary rocks along deep-seated strike-slip faults (Nokleberg et al. 2005; 128

Nelson and Colpron 2007), and are associated with a second pulse of post-accretionary porphyry 129

Cu deposits in the Stikinia terrane (McMillan et al. 1995). Several porphyry deposits constitute 130

the Skeena Arch metallogenic belt in central British Columbia, many of which are related to 131

Eocene intrusive systems. These include: the Ajax, Bell Moly, and Kitsault porphyry Mo 132

deposits associated with the 54−48 Ma Alice Arm intrusive suite; the Berg porphyry Cu-Mo, and 133

Ajax porphyry Mo deposits associated with the 54−48 Ma Nanika intrusions; the Equity Silver 134

and Prosperity-Porter Idaho Ag polymetallic vein deposits associated with the 50−57 Ma Goosly 135

intrusions; the Bell, Granisle, and Morrison porphyry Cu-Au-Mo deposits associated with 54−50 136

Ma Babine Lake igneous suites; and the Endako porphyry Mo deposit associated with the 50.5 ± 137

0.5 Ma Francois Lake plutonic suite (Church 1970, 1972; Carter 1982; Dostal et al. 1998, 2001; 138

Grainger et al. 2001). These intrusions are calc-alkaline, quartz dioritic to granitic in composition, 139

and form small stocks or dikes (typically less than 1 km in diameter; Carter 1976), which are 140

mostly emplaced along northwest- and northeast-striking faults (Woodsworth et al. 1991; Carter 141

1982; Nokleberg et al. 2005). 142

143

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Geology of the Babine Lake Area 144

Mesozoic to early Cenozoic volcanic and associated sedimentary rocks form the country 145

rocks to porphyry deposits in the Babine Lake area. At Morrison, these rocks are composed of 146

three main volcanic and sedimentary units: the Lower to Middle Jurassic Hazelton Group, the 147

Middle Jurassic to Upper Jurassic Bowser Lake Group, and the Lower Cretaceous Skeena Group. 148

These stratified rocks are intruded by Cretaceous to early Cenozoic stocks and dikes (Fig. 2). 149

The Hazelton Group consists mainly of andesitic volcanic rocks and marine to non-marine 150

sedimentary rocks (Carter 1976; Massey et al. 2005), whereas the Bowser Lake and Skeena 151

Groups are composed mainly of marine to non-marine clastic sedimentary rocks deposited in a 152

fluvial-deltaic to near shore shelf environment (Maclntyre 2006). Late Cretaceous to Eocene 153

transpressional to transtensional tectonics resulted in uplift, faulting, and tilting to generate 154

several linear horsts and grabens. The Morrison deposit is located in one of these grabens, and is 155

bounded by an unnamed fault to the west and the Morrison Fault to the east (Fig. 2; Simpson and 156

Geo 2007). Within this graben, the Ashman Formation of the Bowser Lake Group forms the 157

immediate host to the Morrison porphyry deposit. 158

Two groups of faults are recognized in the Morrison area, with dominantly NNW, and ENE 159

trends (Fig. 2; Massey et al. 2005). Carter (1976) suggested that the faults controlled the 160

emplacement of the Cretaceous to early Cenozoic plutonic rocks, and that the intrusions 161

generally occur along the NNW-trending structures. 162

The Cretaceous to early Cenozoic plutonic suites mainly range from quartz diorite, quartz 163

monzonite, granodiorite, to granite in composition (Carter 1976; Maclntyre 2006). In particular, 164

the plagioclase-hornblende-biotite porphyry stocks and dikes of the Babine Lake intrusive suite 165

are quartz-dioritic to granodioritic in composition (Carson and Jambor 1974). These intrusions 166

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are associated with several porphyry deposits in the northern part of the Babine Lake area, 167

including the Bell, Granisle, Nakinilerak, Dorothy, North Newman, South Newman, Hearne Hill, 168

and Morrison deposits (Carson and Jambor 1974). 169

Geology of the Morrison Porphyry Cu-Au-Mo Deposit 170

The central part of the Morrison deposit is relatively well exposed, whereas the surrounding 171

area is mostly covered by Quaternary alluvium and glaciofluvial till, and is heavily forested. In 172

this study, we relied on the previous mapping and geological descriptions of Carson and Jambor 173

(1976), supplemented by our own observations and sampling of a limited number of preserved 174

drill cores. Six drill holes were selected for study (Fig. 3), which contain typical examples of the 175

main lithological units, alteration styles, and mineralization types. 176

177

Eocene plagioclase-hornblende-biotite porphyry 178

A small stock of plagioclase-hornblende-biotite porphyry and associated dikes intrude the 179

Ashman Formation in the centre of the Morrison property (Fig. 3). The porphyry stock is 180

interpreted to have been circular with a diameter of about 600 meters, but is now bisected and 181

offset by a north-trending strike-slip fault, named the East Fault, which has a dextral offset of 182

approximately 300 m (Simpson and Geo 2007). 183

The plagioclase-hornblende-biotite porphyry is dark grey in color, dioritic to granodioritic in 184

composition, and has a porphyritic texture (Fig. 4). The porphyry mainly consists of biotite, 185

plagioclase, and hornblende phenocrysts with accessory magnetite and apatite, set in a fine-186

grained quartzofeldspathic matrix. Carson and Jambor (1976) suggested that the porphyry stock 187

consists of multiple phases, and in this study three distinct phases of porphyry (phases A, B, and 188

C) were identified based on contrasting abundances and different grain sizes of phenocrysts, and 189

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relative timing (Fig. 4A−D; summarized in Table 1). Contacts between the three phases were 190

rarely observed in drill core, and the only boundary observed in this study indicated that phase C 191

porphyry intruded into phase A porphyry as shown in Figure 4D. However, cross-cutting 192

relationships between phase B and phases A and C were not clear. 193

194

Sedimentary host rocks 195

The Ashman Formation sedimentary sequence in the Morrison area strikes N to NW, with a 196

steep dip (Carson and Jambor 1976). The sequence consists of marine pebble conglomerate at 197

the base overlain by siltstone, sandstone, and greywacke (Simpson and Geo 2007). The siltstone, 198

which is dominant, is grey to dark grey, fine- to medium-grained, and mainly composed of 199

detrital quartz and feldspars. 200

201

Structure 202

The most significant structure affecting the Morrison porphyry is the East Fault, which trends 203

NNW, dips vertically, and shows dextral strike-slip offset of the main porphyry stock of ~300 m 204

(Fig. 3; Carson and Jambor 1976). Based on the description of the faults and sedimentary rocks 205

by Carson and Jambor (1976), the fault is likely to be bedding parallel. Some vertical 206

displacement is also thought to be present because the two segments of the porphyry stock do not 207

fully match (Carson and Jambor 1976). The East Fault ranges from a few meters to ~50 meters in 208

width, and rocks are strongly fractured along its length. Sericite-carbonate alteration associated 209

with polymetallic vein mineralization is associated with this structure. 210

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Simpson and Geo (2007) also suggested the existence of a second NNW-trending fault 211

cutting sedimentary rocks to the west of the main Morrison deposit (the West Fault), but this 212

fault was not observed in this study. 213

214

Hydrothermal Alteration 215

Hydrothermal alteration (potassic, propylitic, sericite-carbonate, and argillic) has affected the 216

central porphyry stock and surrounding siltstones at Morrison. Potassic alteration mainly occurs 217

in the central porphyry stock (Fig. 5A, B), with minor development in adjacent sedimentary 218

rocks, whereas propylitic alteration mainly occurs in the sedimentary rocks, and only locally 219

affects the intrusion as an overprint on potassic alteration (Fig. 5C). The potassic alteration is 220

closely related to Cu mineralization, which occurs as vein-hosted and disseminated chalcopyrite 221

and bornite in altered plagioclase-hornblende-biotite porphyry. In contrast, the propylitic-altered 222

rocks only contain minor pyrite. Classic phyllic alteration (sericite-quartz-pyrite), which is 223

common in many other porphyry Cu deposits (Lowell and Guilbert 1970), does not occur at 224

Morrison. Instead, minor sericite-carbonate alteration is restricted to the East Fault, and is 225

associated with late-stage polymetallic sulfide-carbonate veins (Fig. 5D). Argillic alteration is 226

also restricted to the East Fault, and overprints all other alteration styles; it may be supergene in 227

origin (i.e., related to groundwaters permeating the fault zone). 228

229

Potassic alteration 230

The potassic alteration is characterized by secondary biotite, which replaces hornblende 231

phenocrysts and occurs as fine-grained crystals in the matrix (Fig. 6A). In strongly altered rocks, 232

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igneous biotite phenocrysts also show overgrowths of hydrothermal biotite (Fig. 6B). Early 233

biotite veins are locally present in the potassic zone, with widths between 1–3 mm. 234

The degree of potassic alteration decreases from the center of the porphyry stock outwards. 235

Hydrothermal biotite in strongly altered rocks is deep brown, coarse-grained, and well developed 236

in the matrix, whereas in more weakly altered rocks it is greenish brown, fine-grained, and only 237

present as alteration of hornblende phenocrysts. These observations are consistent with 238

the earlier descriptions of Carson and Jambor (1976). 239

240

Propylitic alteration 241

Propylitic alteration is mainly characterized by secondary chlorite and carbonate minerals. 242

Secondary chlorite replaces hornblende and locally biotite, and carbonate replaces plagioclase 243

phenocrysts (Fig. 6C). 244

245

Sericite-carbonate and argillic alteration 246

Sericite-carbonate alteration is mainly characterized by secondary sericite and carbonate. The 247

sericite is present as halos a few millimeters wide around polymetallic sulfide-carbonate veins, 248

whereas dolomite occurs in the center of the veins (Fig. 6D). 249

Argillic alteration is characterized by kaolinite and carbonate minerals (Figs. 6E, F), which 250

replace biotite, hornblende, and plagioclase phenocrysts, as well as the matrix of the porphyries. 251

The clay and carbonate minerals in the argillic-altered rocks are primarily kaolinite and dolomite, 252

with minor ankerite, siderite, and rare calcite (based on X-ray diffraction analyses). 253

254

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Vein relationships 255

A total of five vein types were identified at Morrison based on vein mineralogy and cross 256

cutting relationship, including (1) early biotite (EB) veins; (2) stockwork veinlets of chalcopyrite 257

± quartz-bornite-pyrite; (3) quartz-chalcopyrite-bornite-pyrite veins with centerline sulfides; (4) 258

quartz-molybdenite ± pyrite veins; (5) polymetallic sulfide-dolomite veins composed of dolomite 259

with quartz-sphalerite-galena-arsenopyrite ± chalcopyrite (Fig. 7). These veins can be grouped 260

into three paragenetic stages based on similarities in timing, mineral assemblages, and related 261

alteration assemblages (Fig. 8). Veins types 1 to 3 are associated with potassic alteration, type 4 262

veins display weak sericitic selvedges, and type 5 veins are related to weak sericite-carbonate 263

alteration. 264

The early biotite (EB) veins (type 1) are rare, and are only present in the strongly biotite-265

altered porphyry (Fig. 7A, B). These veins are typically 1–3 mm in width, and are mainly 266

composed of dark brown biotite without sulfides. The type 2 and 3 veins are widespread in the 267

biotite-altered porphyry, and are associated with the bulk of the Cu mineralization. Type 2 veins 268

are typically 1–2 mm in width, and consist of chalcopyrite, minor bornite, magnetite, and pyrite 269

with or without quartz (Fig. 7B, C), whereas type 3 veins are straight, 3–5 mm in width, and 270

mainly composed of coarse-grained quartz with a centerline of sulfides (mainly chalcopyrite and 271

minor pyrite; Fig. 7C, D). The type 2 and 3 veins are similar to A and B veins of Gustafson and 272

Hunt (1975). The type 4 quartz-molybdenite ± pyrite veins are straight and 5–10 mm in width, 273

and formed later than the type 2 and 3 veins; they display weak sericitic alteration selvedges 274

where they cut potassic-altered rocks (Fig. 7D, E). 275

Type 5 veins formed during the last hydrothermal stage at Morrison, and only occur in 276

fractured zones around the East Fault, as infillings of fractures and breccias. These polymetallic 277

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sulfide-dolomite veins are 1 cm to a few cm in width, and consist of minor quartz at the vein 278

margin, coarse-grained sulfides (sphalerite, galena, minor arsenopyrite, pyrite, and rare 279

chalcopyrite), and late coarse-grained dolomite in the vein center and in cross-cutting veinlets 280

(Figs. 6D and 7F). Narrow (0.5–2 mm) alteration halos consist of sericite-carbonate. 281

282

Mineralization 283

Most Cu mineralization at Morrison is related to potassic alteration within the plagioclase-284

hornblende-biotite porphyries. All three phases of porphyry are mineralized, suggesting that they 285

all immediately predated the mineralizing event. Two semicircular copper zones, termed the 286

northwest and southeast zones, have average grades of 0.39% Cu (0.3% Cu cutoff), and are cut 287

and offset by the East Fault (Fig. 3). The copper zones are surrounded by well-developed annular 288

pyrite halos (Fig. 3). Pyrite is typically associated with phyllic alteration in porphyry deposits, 289

but at Morrison it is mainly associated with the chlorite-carbonate (propylitic) alteration. The 290

ratio of chalcopyrite to pyrite decreases from the potassic to the propylitic zone. Weak copper 291

mineralization (≤0.3% Cu) also occurs in the siltstone country rocks in the northwest zone, but is 292

mainly restricted to the porphyry in the southeast zone (Carson and Jambor 1976). 293

The mineralization can be divided into three stages (Figs. 8, 9). The first stage contains the 294

bulk of the Cu mineralization, and includes vein types 1−3 (Fig. 9A–C). Carson and Jambor 295

(1976) noted that seventy to eighty percent of the copper mineralization is hosted by stockwork 296

veins and small veinlets in this first stage. Sulfides in stage 1 mineralization mainly consist of 297

chalcopyrite, minor bornite, and pyrite, with minor magnetite in veins and disseminations 298

associated with potassic alteration. 299

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The second stage of mineralization consists of molybdenite-pyrite veins with weak sericitic 300

alteration selvedges (Fig. 9D, E). Molybdenum grades do not correlate with copper, and mainly 301

occur at the edge of the copper zone (Simpson and Geo 2007). The molybdenum mineralization 302

is not considered to be economic (Simpson and Geo 2007). 303

The third stage of mineralization consists of polymetallic sulfide-dolomite veins (sphalerite-304

galena-arsenopyrite-pyrite ± chalcopyrite; Fig. 9F–H), which are present along the East Fault 305

zone. These veins are also not considered to be economic (Simpson and Geo 2007). 306

Gold was detected in the porphyry by assay analyses (207 million tonnes grading 0.2 g/t 307

gold; Simpson and Geo 2007), but is not visible in hand samples or under the microscope. Gold 308

appears to correlate with copper mineralization based on assay data (Ogryzlo et al. 1995). 309

310

Fieldwork and Analytical Methods 311

Sample selection 312

A total of 33 drill-core samples from 6 selected diamond drill-holes were collected for this 313

study. The samples are representative of the main lithological units, alteration facies, and 314

mineralization styles in the Morrison deposit. Detailed samples descriptions are provided in 315

Table A1, and sampled drill holes are marked on the geological map in Figure 3. 316

317

Lithogeochemistry 318

Nine samples of least-altered porphyry intrusions were collected for lithogeochemical 319

analysis. These least-altered samples still contain some secondary biotite indicating weak 320

potassic alteration, and many samples also have a weak chlorite overprint. Whole-rock 321

geochemical analyses were conducted by Activation Laboratories (Ancaster, Ontario, Canada), 322

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using instrumental neutron activation analysis and fusion ICP-MS methods (Actlabs code 4E 323

Research + ICP/MS). As determined by reproducibility of standards and duplicates, accuracy is 324

within 5 and 10 relative percent for major and trace elements, respectively. All major oxide 325

compositions were recalculated to a volatile-free basis for plotting on lithological classification 326

diagrams. 327

328

Geochronology 329

A drill-core sample of phase A plagioclase-hornblende-feldspar porphyry (sample MO128 330

from drillhole MO-01-26, 215.8 to 217.3 m) was crushed and zircons were separated using 331

standard gravimetric and magnetic methods followed by hand-picking. The zircons were 332

mounted in epoxy, and polished to expose the crystal cores. 333

Zircons were dated using a multiple collector inductively coupled plasma-mass spectrometer 334

(MC-ICP-MS; Nu-Plasma, Nu Instruments, UK) coupled to a frequency quintupled (λ = 213 nm) 335

Nd:YAG laser ablation system (New Wave Research, USA) at the Canadian Center for Isotopic 336

Microanalysis (CCIM), University of Alberta. The analytical approach was modified from 337

Simonetti et al. (2005), and laser pits were approximately 30 µm in width and 20–30 µm in 338

depth. Zircon reference materials GJ-1-32 (Jackson et al. 2004) and LH94-15 (Ashton et al. 339

1998) were used for correction of laser induced U-Pb fractionation, correction of 340

instrument/mass bias, and assessment of data quality during the analytical session. The 2σ 341

reproducibility of the standards is ~3% for U/Pb, and 1% for 207

Pb/206

Pb; these external errors are 342

propagated quadratically into the within-run measurement errors. A concordia intercept age was 343

calculated by anchoring to a common Pb 207

Pb/206

Pb value of 0.83 ± 0.06 (Stacey and Kramers 344

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1975) and using a 2 dimensional York linear regression algorithm (within the Isoplot software of 345

Ludwig 2003). 346

Two molybdenite samples were analyzed for Re-Os geochronology at the Radiogenic Isotope 347

Facility, University of Alberta. The molybdenite occurs in 1–2 cm-wide type 4 quartz-348

molybdenite-pyrite veins. Pure molybdenite was separated from the veins by metal-free crushing, 349

gravity, and magnetic separation. Then the concentrations of 187

Re and 187

Os in molybdenite 350

were analyzed by isotope dilution mass spectrometry using methods described by Selby and 351

Creaser (2004) and Markey et al. (2007). The model age of molybdenite was calculated based on 352

the equation: 187

Os = 187

Re*(eλt

-1), where the decay constant (λ187

Re) used was 1.666× 10-11

a-1

353

(Smoliar et al., 1996). Two sigma uncertainties of the molybdenite Re-Os data are attributed to 354

analytical uncertainty, decay constant uncertainty, and calibration uncertainties. 355

356

Stable isotopes 357

Sulfur isotopes: Sulfide minerals, including chalcopyrite, pyrite, molybdenite, and sphalerite, 358

were separated by crushing, sieving, and hand picking from mineralized rocks. Approximately 359

5–10 mg of each sample was separated for sulfur analysis, and the purity of each sulfide sample 360

was greater than 90%. Seven sulfide samples were sent for analysis to the Isotope Science 361

Laboratory at the University of Calgary, and data are reported in the usual per mil notation 362

relative to the Canyon Diablo Troilite (CDT) standard. The accuracy of sulfur isotope 363

measurements is ±0.3‰. 364

Oxygen and carbon isotopes: Pure quartz was separated from quartz-sulfide veins by 365

crushing, sieving, and hand picking. Seven quartz samples of approximately 10 mg each were 366

sent to the Geoanalytical Laboratory at the Washington State University for oxygen isotope 367

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analyses, using methods described by Takeuchi and Larson (2005). Isotopic compositions were 368

analyzed using a Finnigan MAT Isotope-Ratio Mass Spectrometer, attached to a high-vacuum 369

laser fluorination line for the extraction of oxygen. The UWG-2 garnet standard was used in this 370

analysis (Valley et al. 1995). All results are reported in per mil notation relative to Vienna 371

Standard Mean Ocean Water (VSMOW). The accuracy and precision are better than ± 0.1 and 372

0.05‰, respectively. The oxygen stable isotopic compositions of hydrothermal fluids were 373

calculated using the fractionation equation of Clayton et al. (1972), assuming oxygen isotope 374

equilibrium was reached between quartz and hydrothermal fluids at crystallization temperatures 375

estimated from fluid inclusion microthermometry. 376

Pure dolomite was separated from two stage 3 polymetallic sulfide-dolomite veins by 377

crushing and hand picking. Samples were analyzed for their oxygen and carbon isotope 378

compositions in the Stable Isotope Laboratory at the University of Alberta, using a Finnigan 379

MAT 252 dual inlet mass spectrometer. Results are reported in the usual per mil notation relative 380

to the Vienna-Pee Dee Belemnite (VPDB) standard for carbon isotopes, and the Vienna-Standard 381

Mean Ocean Water (VSMOW) standard for oxygen isotopes. Measurement accuracy is ± 0.1‰. 382

Oxygen isotopic compositions of hydrothermal fluids, which are assumed to be in equilibrium 383

with the dolomite, were calculated using the fractionation equation of Horita (2014), at 384

temperatures estimated from fluid inclusion microthermometry. 385

386

Fluid inclusion samples and methodology 387

Doubly polished thin sections (~100 µm thick) of quartz and dolomite vein samples 388

representing the main mineralization stages were prepared for fluid inclusion analysis. A 389

Linkham THMSG600 heating and freezing stage mounted on an Olympus BX50 microscope was 390

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used for microthermometric analyses, and the stage was calibrated using synthetic fluid inclusion 391

standards from Syn Flinc. The analytical precision is ±0.1°C for low temperature measurements, 392

and ±1°C for high temperature (above ambient) measurements. The measurement accuracy was 393

±0.2°C for low temperature measurements, and ±2°C above 10°C. Sample chips containing fluid 394

inclusions to be analyzed were cooled to -100°C, and then reheated progressively, recording the 395

temperatures of phase changes, until total homogenization. During the reheating process the 396

melting temperatures of CO2 (where present; TmCO2), ice (Tmice), sylvite (Tmsylvite), and halite 397

(Tmhalite), and total homogenization temperature (Th) were recorded. The salinities of fluid 398

inclusions, reported in weight percent NaCl equivalent, were derived from ice melting or halite 399

dissolution temperatures. 400

401

Whole-Rock Geochemistry 402

Due to the widespread hydrothermal alteration in the area, almost all rocks observed in drill 403

core were at least weakly altered. Nine least-altered (weak biotite alteration ± minor chlorite 404

overprint) samples of the various porphyry phases were selected and analyzed for their whole-405

rock geochemical composition. Major and trace element geochemical data are presented in Table 406

2. 407

408

Major element geochemistry 409

All major element oxide data were recalculated to 100% volatile free for the purposes of 410

plotting on lithological discrimination diagrams. Most porphyry samples plot within the diorite 411

and granodiorite fields on the total alkali-silica diagram (Fig. 10), but two samples (MO001 and 412

MO038) straddle the boundary between the monzonite and diorite fields. On a plot of K2O 413

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versus SiO2, most of the suite plots within the calc-alkaline field, with some extending into the 414

high-K calc-alkaline field (Fig. 11). The moderately elevated K2O contents of some of the 415

samples in this suite might be due in part to weak potassic alteration. The compositional range of 416

the Morrison samples is similar to that of other reportedly fresh Babine Lake intrusions (Carson 417

and Jambor 1974; Ogryzlo 1995). 418

419

Trace element geochemistry 420

Trace element data for the plagioclase-hornblende-biotite porphyry samples are plotted on 421

primitive mantle-normalized trace element and chondrite-normalized rare earth element (REE) 422

diagrams in Figure 12. In the primitive mantle-normalized diagram (Fig. 12A), the samples show 423

enrichments in large ion lithophile elements (LILE) but relatively low concentrations of high 424

field strength elements (HFSE) and middle and heavy REE (MREE, HREE). Phase A porphyry 425

samples are somewhat more enriched in LILE (especially Ba, La, Ce, and Sr) than the other 426

samples, whereas the porphyry C sample is slightly more enriched in MREE-HREE. All samples 427

display negative Ta, Nb, and Ti anomalies, which are characteristics of arc-related igneous rocks 428

(Brenan et al. 1994; Foley et al. 2000). On the chondrite-normalized rare earth element diagram 429

(Fig. 12B), all samples show weakly listric-shaped REE patterns, which slope downward from 430

light REE (LREE) to MREE (high [La/Sm]n ratios from 3.1 to 8.3), and then flatten from MREE 431

to HREE (low [Dy/Yb]n ratios from 1.4 to 1.6). 432

All porphyry samples plot in the field of volcanic arc granites on tectonic discrimination 433

diagrams (Fig. 13). Thus, the plagioclase-hornblende-biotite porphyry suite is concluded to be of 434

calc-alkaline, dioritic to granodioritic composition, and of arc-related magmatic affinity. 435

436

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Geochronology 437

U-Pb zircon dating 438

Zircons from a sample of phase A plagioclase-hornblende-biotite porphyry (MO128) are 439

typical of magmatic crystals: clear, euhedral, and colorless. Weak oscillatory zoning is evident in 440

the backscattered electron images (Fig. 14). Thirty zircons were analyzed, and all U-Pb isotopic 441

data are summarized in Table 3 and plotted on a Concordia diagram in Figure 15. The data show 442

a relatively homogeneous population, and yielded a U-Pb concordia intercept age of 52.54 ± 0.37 443

Ma (MSWD = 1.13, probability of fit 0.29) using the 2D York regression algorithm of Isoplot. It 444

has been observed over several years that a systematic offset of ~2-3% relative to TIMS U-Pb 445

ages can occur when normalizing one standard to another (Hanchar 2009; Klötzli et al. 2009; 446

Košler et al. 2013). While the ultimate causes of this disagreement are debated, suggestions 447

include differences in α dose radiation damage (Allen and Campbell 2012), differences in trace 448

element composition (Black et al. 2004), relative orientation of the crystallographic plane and/or 449

slight differences in laser focus (Marillo-Sialer et al. 2014), and the variable presence of oxygen 450

in the plasma (Košler et al. 2014). To account for this potential inaccuracy we assume a 451

minimum uncertainty of 2%, leading to an ultimate age uncertainty of ±1.05 m.y. Thus, our 452

preferred age for this sample is 52.54 ± 1.05 Ma. This age estimate independent of any possible 453

geologic complexity that may be present, and does not include associates uncertainties in the 454

TIMS measurements of the standards. 455

456

Molybdenite Re-Os dating 457

Molybdenite at Morrison occurs as fine-grained crystals in type 4 quartz-molybdenite veins 458

(Fig. 7D, E) commonly associated with pyrite. Two samples (MO094 and MO097) of 459

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molybdenite-bearing quartz veins were dated by the Re-Os technique. The veins are hosted by 460

clay-carbonate altered plagioclase-hornblende-biotite porphyry (phase B), which is probably an 461

overprint on previous phyllic alteration. They were collected from drill hole MO-01-031 at 462

depths of 184.5 and 198.5 m respectively. The Re and Os concentrations and model ages of the 463

samples are presented in Table 4, and yielded Re-Os ages of 52.54 ± 0.22 and 53.06 ± 0.22 Ma, 464

which are consistent with the U-Pb zircon age for the phase A porphyry intrusion (52.54 ± 1.05 465

Ma). 466

467

Fluid inclusions 468

Fluid inclusion types 469

In this study, primary and pseudosecondary fluid inclusions in quartz and dolomite were 470

analyzed, whereas secondary and necked or leaked inclusions were avoided (using the criteria of 471

Roedder 1984, and Goldstein and Reynolds 1994). Primary inclusions were typically found in 472

growth zones of quartz and dolomite crystals, whereas pseudosecondary inclusions generally 473

occur in healed microfractures within crystals; these groupings constitute fluid inclusions 474

assemblages (terminology of Goldstein and Reynolds 1994). Most measured inclusions have 475

sizes of 5–10 µm. 476

Two type 3 quartz-chalcopyrite-pyrite veins, two type 4 quartz-molybdenite-pyrite veins, and 477

two type 5 polymetallic sulfide-dolomite veins samples were selected for fluid inclusion analysis 478

after petrographic examination. Three types of fluid inclusions were recognized based on their 479

features at room temperature: multiphase hypersaline brine inclusions (H) with halite ± sylvite ± 480

opaque minerals and anhydrite daughter crystals (Fig. 16A−C, E), two-phase (liquid + vapor) 481

vapor-rich inclusions (V; Fig. 16D), and two-phase (liquid + vapor) liquid-rich inclusions (L; 482

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Fig. 16F). These fluid inclusions occurred in three assemblage types: (1) brine plus vapor-rich 483

inclusions (H + V; Fig. 16A); (2) vapor-rich inclusions (V only; Fig. 16D); and (3) liquid-rich 484

inclusions (L only; Fig. 16F). The first and second assemblage types occur in quartz in both the 485

type 3 quartz-chalcopyrite-pyrite and type 4 quartz-molybdenite-pyrite veins, whereas the third 486

assemblage only occurs in carbonate in the type 5 polymetallic sulfide-dolomite veins. 487

The hypersaline inclusions are irregular to rounded in shape, with cubic halite daughter 488

crystals that account for 20−40% of the inclusion volume. Sylvite daughter crystals are smaller 489

and rounded in shape, and are only observed in a few inclusions. Upon heating, the sylvite 490

crystals dissolve at lower temperatures than halite. Anhydrite crystals are the least common 491

daughter crystals, and are recognized by their rectangular shape and high birefringence. Minor 492

opaque daughters are too small to be distinguished, and are likely sulfides (chalcopyrite?). 493

Neither the anhydrite nor opaque daughter phases dissolved upon heating. 494

495

Homogenization temperature/salinity data 496

A total of 102 fluid inclusions were analyzed from the vein samples, and all 497

microthermometric results are listed in Table A2 and illustrated in Figures 17 and 18. Fluid 498

inclusions were measured in assemblages, and the homogenization temperature and salinities 499

within a single fluid inclusion assemblage were found to be similar (± 50°C, ± 5 wt.% NaCl 500

equiv.), but varied between different assemblages. No data could be recorded from vapor-rich 501

inclusions because of difficulties in observing phase changes in these predominantly dark 502

inclusions (cf. Bodnar et al. 1985). No liquid CO2 was observed in any inclusions at or below 503

room temperature, but faint melting events close to -56.6°C (interpreted to be the melting of 504

small crystals of solid CO2) were observed in some vapor-rich inclusions, suggesting the 505

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presence of minor amounts of CO2 in the vapor phase. However, no clathrate melting events 506

were observed in liquid-rich fluid inclusions, so the CO2 contents of the liquid phase are thought 507

to be low. 508

Type 3 quartz-chalcopyrite-pyrite vein samples (MO135 and MO144): Forty-two hypersaline 509

inclusions from eleven fluid inclusion assemblages homogenized finally by vapor-bubble 510

disappearance at temperatures between 400–526°C, with a slightly bimodal distribution and an 511

average of 464° ± 42°C. Halite dissolution temperatures ranged from 320 to 403°C, and 512

calculated salinities range from 39.8 to 47.8 wt.% NaCl equiv. (mean = 43.6 ± 1.8 wt.% NaCl 513

equiv., n = 42; equation of Sterner et al. 1998). Sylvite was observed in several inclusions, and 514

the sylvite dissolution temperatures ranged from 102° to 122°C. If the hydrothermal fluid is 515

modeled in the NaCl-KCl-H2O system, then the composition can be estimated to be 32 wt.% 516

NaCl and 21 wt.% KCl based on the phase diagram shown in Figure 19 (Roedder 1984). 517

Several fluid inclusion assemblages included vapor-rich fluid inclusions. Although 518

homogenization temperatures and salinities could not be measured for these inclusions due to 519

their dark appearance, the coexistence of vapor-rich and liquid-rich inclusions indicates that 520

trapping conditions were on the two-phase liquid-vapor curve, and that final homogenization 521

temperatures approximate the trapping temperature. Small amounts of CO2 were detected in two 522

of these inclusions during melting at ~-56.6°C. 523

Type 4 quartz-molybdenite-pyrite vein samples (MO094 and MO097): Eleven fluid inclusion 524

assemblages containing hypersaline fluid inclusions were measured from this sample. Fluid 525

inclusions in nine of these assemblages homogenized finally by vapor bubble disappearance, 526

with an average Th(L-V)L of 370° ± 31°C and TmNaCl of 40 ± 1.2 wt.% NaCl equiv. Fluid 527

inclusions in two assemblages homogenized finally by halite melting at temperatures of 327°–528

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357°C, a few degrees above the bubble disappearance temperature of 292°–330°C. The average 529

final homogenization temperature for all of these fluid inclusions is 363 ± 30°C (n = 36). 530

Salinities calculated from halite dissolution temperatures range from 37.0 to 43.1 wt.% NaCl 531

equiv. with an average of 40.0 ± 1.4 wt.% NaCl equiv. (n = 36). Sylvite daughter crystals are rare 532

in fluid inclusions from these veins, and no sylvite dissolution temperatures were recorded. 533

Vapor-rich fluid inclusions also occurred in some of these fluid inclusion assemblages, and 534

final homogenization temperatures are therefore interpreted to approximate the trapping 535

temperature. Halite-bearing inclusions in these fluid inclusion assemblages homogenized by 536

vapor disappearance. 537

Type 5 polymetallic sulfide-dolomite vein samples (MO072 and MO083): Twenty-four 538

liquid-rich fluid inclusions from 9 assemblages in dolomite crystals were measured. 539

Homogenization temperatures range from 163 to 218°C with an average of 185 ± 16°C (n = 24). 540

Fluid salinities calculated from final ice melting temperatures range from 3.1 to 3.9 wt.% NaCl 541

equiv. (mean = 3.5 ± 0.3 wt.% NaCl equiv., n = 8; equation of Bodnar 1993). 542

543

Pressure corrections 544

No pressure corrections were applied to the homogenization temperatures of fluid inclusions 545

from type 3 and 4 veins, because of the coexistence of liquid- and vapor-rich fluid inclusions, 546

indicating trapping under two-phase (boiling) conditions, such that homogenization temperature 547

equals the trapping temperature (and pressure). 548

However, no vapor-rich fluid inclusions occur in the type 5 polymetallic-sulfide dolomite 549

veins, and so the homogenization temperature is presumed to be a minimum estimate of the 550

trapping temperature, and a pressure correction is required. No independent estimate of pressure 551

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is available for this late stage of mineralization, but the trapping pressure of stage 2 fluids can 552

provide a maximum constraint, assuming that pressures fell from stage 2 to stage 3 (perhaps by 553

transition from near lithostatic to hydrostatic pressure conditions, and/or by uplift and erosion). 554

The trapping pressure of stage 2 fluid, was estimated using the program HOKIEFLINCS_H2O-NaCl 555

(Steele-MacInnis 2012), and yielded a value of ~75 bar. This value is likely to be a minimum 556

estimate, given the observation of small amounts of CO2 in some vapor-rich fluid inclusions, but 557

nevertheless suggests that fluid pressures during the porphyry stage of mineralization were quite 558

low. Assuming this pressure estimate is approximately correct, a pressure correction of 4°C is 559

indicated for the ~3.5 wt.% NaCl equiv. fluids in stage 3 polymetallic-sulfide dolomite veins. 560

Given that this small correction could be too large (if pressures were significantly less than 561

during phase 2) or too small (if CO2 added significantly to stage 2 pressures), is seems arbitrary 562

to make any specific correction to stage 3 fluid temperatures, except to note that the trapping 563

temperature might be higher than the homogenization temperature by up to ~5°C, thus indicating 564

an average trapping temperature for stage 3 fluids of ~190° ± ~20°C. 565

566

Oxygen isotope compositions 567

Quartz was separated from three type 3 (stage 1), two type 4 (stage 2), and two type 5 (early 568

stage 3) vein samples for oxygen isotopic analysis. Dolomite was separated from two type 5 (late 569

stage 3) vein samples for oxygen and carbon isotopic analysis. Isotopic values are listed in Table 570

5. Oxygen isotopic compositions of quartz from type 3 and 4 veins range from δ18

O = 6.1 to 571

8.4‰, and 9.4 and 10.7‰ in type 5 veins. Dolomite from type 5 veins has δ18

O values of 13.6 572

and 16.6‰, and δ13

C values of 0.7 and 0.6‰, respectively. 573

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Assuming that O isotopic equilibration occurred between quartz and dolomite and the 574

hydrothermal fluids during crystallization, the δ18

O compositions of the hydrothermal fluids can 575

be derived using the fractionation equations of Clayton et al. (1972) and Horita (2014), and 576

temperatures derived from fluid inclusion studies. Evidence for phase separation in type 3 and 4 577

veins means that measured fluid inclusion homogenization temperatures approximate the 578

trapping temperatures for these fluids. The average temperature of fluids inclusions in type 3 579

veins is 464° ± 42°C, and 363° ± 30°C for type 4 veins. δ18

Ofluid values have been calculated at 580

one standard deviation above and below these average temperatures (Table 5 and Fig. 20). 581

Quartz at the margin of type 5 polymetallic sulfide-dolomite veins was deposited from early 582

stage 3 hydrothermal fluids, whereas dolomite in the center of the veins was deposited from late 583

stage 3 fluid. No suitable fluid inclusions were found in the quartz, but inclusions in dolomite 584

yielded an average trapping temperature of 185° ± 16°C. Adjusting for pressure, the trapping 585

temperature of these fluid inclusions is estimated to be 190° ± 16°C was used to calculate 586

δ18

Ofluid compositions. For early stage 3 fluids, we have calculated fluid isotopic compositions at 587

300°C and 200°C in order to bracket the likely range of values. 588

Calculated δ18

Ofluid values for stage 1 (type 3 veins) fluids range from 3.7 to 6.3‰ 589

(incorporating the error range of the fluid inclusion data), whereas fluid compositions for stage 2 590

(type 4 veins) fluids range from 0.3 to 3.4 ‰. The composition of early stage 3 fluids (type 5 591

veins) is roughly estimated to be in the range of -2.3 to 3.9‰ (based on the likely fluid 592

temperature range), whereas the composition of late stage 3 fluids is estimated to range between 593

1.1 and 6.1‰. 594

595

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Sulfur isotope compositions 596

Seven sulfide samples were analyzed for their sulfur isotopic composition: two mixed 597

chalcopyrite and pyrite samples from stage 1 disseminated sulfides in potassic alteration, one 598

molybdenite and one pyrite sample from type 4 quartz-molybdenite-pyrite veins of stage 2 599

mineralization, and one pyrite and two sphalerite samples from type 5 polymetallic-sulfide-600

dolomite veins of stage 3 mineralization. The results are reported in Table 6. 601

The two mixed chalcopyrite and pyrite samples (MO095 and MO096) occur as finely 602

intergrown 1−3 mm composite grains that could not be separated, and the analyses therefore 603

represent a mixture of chalcopyrite and pyrite. However, it is noted that the fractionation of S 604

isotopes between these two minerals, and between these minerals and HS-, is not large (<1‰ at 605

300-400°C; Ohmoto 1972, 1986; Ohmoto and Rye 1979). The δ34

S compositions of these stage 1 606

sulfides are -0.8 and 0.2‰. 607

Molybdenite (sample MO042) and pyrite (sample MO115) from type 4 quartz-molybdenite-608

pyrite veins (stage 2) yielded δ34

S values of -2.1 and -1.2‰, respectively. Pyrite (samples 609

MO072 and MO083) and sphalerite (sample MO072) from type 5 polymetallic veins (stage 3), 610

and yielded δ34

S values of-5.8 and -5.6‰, and -7.1‰, respectively. 611

612

Discussion 613

Petrogenesis of the plagioclase-hornblende-biotite porphyries 614

The U-Pb age of the phase C plagioclase-hornblende-biotite porphyry at Morrison obtained 615

in this study is 52.54 ± 1.05 Ma. This age is consistent within analytical uncertainties of 616

previously published K-Ar (52.1 ± 2.1 Ma; Carter 1982) and laser 40

Ar/39

Ar ages for porphyry 617

intrusions at Morrison (53. 2 ± 0.5 Ma; Maclntyre et al. 2001), but it is not known which 618

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intrusive phase (A, B, or C) these previous dates relate to. The phase C porphyry is younger than 619

phase A, but its relationship to phase B is not know; all porphyry phases are altered and 620

mineralized, so can be considered to be pre- or syn-mineralization. These ages for intrusive rocks 621

at Morrison (~53 Ma) are similar to those of other intrusions in the Babine Lake igneous suite, 622

which range from 55–48 Ma (K-Ar and laser 40

Ar/39

Ar ages: Carter 1976, 1982; Villeneuve and 623

MacIntyre 1997; Maclntyre et al. 2001). 624

Whole-rock geochemical data indicate that the plagioclase-hornblende-biotite porphyries at 625

Morrison are calc-alkaline to high-K calc-alkaline diorites to granodiorites, and were generated 626

in an arc setting. This interpretation is consistent with the tectonic setting of central British 627

Columbia in the early Eocene, which involved oblique convergence between the Kula-Farallon 628

and North America plates (Nokleberg et al. 2000). 629

The normalized trace element patterns of the porphyry samples from Morrison are similar to 630

those from other Babine Lake intrusions and coeval andesitic–dacitic volcanic rocks (SiO2 < 65%) 631

from the Eocene Ootsa Lake Group (47 to 53 Ma) in central British Columbia (Fig. 12), although 632

the overall trace element contents of the volcanic rocks are somewhat higher than the intrusive 633

rocks (Grainger 2000). These patterns show enrichments in LILE and depletions in Nb, Ta, and 634

Ti (Fig. 12A), characteristic of arc-related igneous rocks (Brenan et al. 1994; Foley et al. 2000). 635

The listric-shaped chondrite-normalized REE patterns (Fig. 12B) are likely caused by 636

fractionation of hornblende and/or titanite, because hornblende and titanite preferentially 637

partition MREE (Gromet and Silver 1987; Klein et al. 1997; Bachmann et al. 2005; Prowatke 638

and Klemme 2006). Negative Eu anomalies are absent from most of the samples suggesting 639

either that plagioclase did not fractionate extensively from these magmas (Eu2+

can substitute for 640

Ca2+

in plagioclase; Hanson, 1980) because magmatic water content was high (Naney 1983; 641

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Merzbacher and Eggler 1984; Rutherford and Devine 1988; Moore and Carmichael 1998), 642

and/or that magmatic oxidation state was high (such that most of the Eu was present as Eu3+

, and 643

was not incorporated into fractionating plagioclase; Philpotts 1970; Housh and Luhr 1991; 644

Sisson and Grove 1993). 645

The geochemical similarities between the Morrison and Babine Lake suites, combined with 646

their similar ages and locations, suggest that they are broadly comagmatic, and were generated in 647

an arc environment. More felsic volcanic rocks (SiO2 >65%) from the Eocene Ootsa Lake Group 648

show increased depletions in Sr, P, and Ti, and distinct negative Eu anomalies (especially in the 649

rhyolites) with increase in SiO2 (Fig. 12), suggesting extensive fractionation of plagioclase, 650

apatite, and magnetite in these more evolved rocks. Early amphibole fractionation (prior to 651

plagioclase crystallization) is also suggested by the listric shapes of the REE patterns and Sr/Y 652

ratios >20 (46–164; Table 2). 653

Overall, these data are consistent with a hydrous (hornblende-porphyritic), relatively 654

oxidized (magnetite-dominant) magmatic suite generated in a post-accretionary continental arc 655

environment. Such magmas are considered to be fertile for porphyry Cu ore formation (Blevin 656

and Chappell 1992; Candela 1992; Richards 2003, 2011; Loucks, 2014). 657

658

Timing of ore-formation at Morrison 659

Two molybdenite Re-Os model ages (52.54 ± 0.22 Ma, 53.06 ± 0.22 Ma) were obtained in 660

this study, which are in broad agreement with the U-Pb age of the host porphyry (phase C: 52.54 661

± 1.05 Ma), suggesting that Mo (and Cu) mineralization was coeval with magmatism (within the 662

error of these geochronological methods, and bearing in mind small systematic differences 663

between different radiometric systems). 664

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665

Ore fluids and genetic model 666

The coexistence of vapor-rich inclusions and hypersaline inclusions in type 3 and 4 quartz-667

sulfide veins indicate the occurrence of fluid immiscibility, and permit the interpretation that 668

homogenization temperatures represent actual trapping temperatures. The microthermometric 669

data from type 3 (stage 1) quartz-chalcopyrite-pyrite vein samples show that the early fluids were 670

hot and saline (400° to 526°C; 39.8 to 47.8 wt.% NaCl equiv.), whereas data from type 4 (stage 2) 671

quartz-molybdenite-pyrite veins are slightly cooler and less saline (320° to 421°C; 37.0 to 43.1 672

wt.% NaCl equiv.; Fig. 18, Table A2). Sulfur isotopic compositions of stage 1 disseminated 673

chalcopyrite-pyrite mixtures are +0.8 and -0.2‰, whereas stage 2 vein molybdenite and pyrite 674

are -2.1 and 1.2‰, respectively (Table 6). Calculated fluid oxygen isotopic compositions range 675

from δ18

Ofluid = 3.7 to 6.3‰ in stage 1, and 0.3 to 3.4‰ in stage 2 (Table 5). Taken together, 676

these data support a predominantly magmatic origin for the porphyry fluids and contained sulfur, 677

with a possible minor contribution from meteoric water in stage 2. By comparison, Ohmoto and 678

Rye (1979), Ohmoto (1986), Marini et al. (2001), and Simon and Ripley (2011) report a range 679

for magmatic sulfur of δ34

S = 0 ± 2‰, and Sheppard (1977) suggested a range for magmatic 680

fluids of δ18

Owater = +5.5 to +10.0‰. 681

The third stage polymetallic sulfide-carbonate mineralization is related to localized sericite-682

carbonate alteration (commonly overprinted by argillic alteration) spatially controlled by the East 683

Fault. Early stage 3 fluids have δ18

Ofluid compositions estimated to be roughly between -2.3 and 684

3.9‰ (imprecisely constrained due to a lack of fluid inclusion temperature information; Table 5), 685

which is lower than in the stage 1 fluids, but overlaps with stage 2. Sulfur isotopic compositions 686

are also lower than earlier sulfides (δ34

S = -7.1 to -5.6‰; Table 6), suggesting the involvement 687

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of a sedimentary sulfide component. 688

Fluids related late stage 3 dolomite precipitation are cooler and much more dilute than in 689

stages 1 and 2 (163° to 218°C; 3.1 to 3.9 wt.% NaCl equiv.; Table A2), and calculated δ18

Ofluid 690

compositions from two dolomite samples have a wide range from 1.1 to 6.1‰, which overlaps 691

the entire range of earlier stages. This wide range can be explained either by mixing of magmatic 692

water and meteoric water, or by wallrock reaction. The latter explanation is supported by carbon 693

isotope compositions (δ13

C = 0.6 to 0.7‰), which are consistent with a marine carbonate origin 694

(δ13

C ~ 0‰; Hoefs, 1987), and may reflect reaction between groundwater and the surrounding 695

sedimentary wall rocks. 696

In summary, these data suggest that the stage 3 fluids associated with polymetallic 697

mineralization were dominantly of meteoric groundwater origin, and reacted with the 698

surrounding sedimentary rocks, whereas the earlier fluids associated with porphyry-type Cu and 699

Mo mineralization were dominantly of magmatic origin. 700

701

Comparison with other porphyry deposits in the Babine Lake area 702

Key characteristics of the Bell, Granisle, and Morrison porphyry deposits are listed in Table 703

7. Based on geological descriptions from Carson and Jambor (1974), Dirom et al. (1995), and 704

Ogryzlo (1995), all three deposits show similarities in host rock ages and compositions, 705

hydrothermal alteration, mineralization style, and ore mineralogy. All deposits are genetically 706

and spatially related to small Eocene plagioclase-hornblende-biotite porphyry stocks, which 707

intruded into Jurassic-Cretaceous volcanic and sedimentary rocks. All porphyries in the Babine 708

Lake area have similar monzonite-diorite-granodiorite compositions (Fig. 10) and ages (~52 Ma). 709

Each of the three deposits is centered on a potassic-altered porphyry stock, which is surrounded 710

Page 31 of 90



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by a peripheral chlorite-carbonate (propylitic) zone. However, other alteration styles differ in the 711

three deposits: sericite-carbonate alteration and quartz-sericite alteration occur at Bell, sericite-712

carbonate alteration occurs at Granisle, and minor sericite-carbonate and argillic alteration occur 713

at Morrison. Cu mineralization is mainly associated with potassic alteration generated by 714

magmatic fluids at Granisle and Morrison, whereas it mainly occurs in quartz-sericite 715

stockworks formed by mixtures of magmatic water and groundwater at Bell (Wilson et al. 1980). 716

The three deposits are similar in many respects to the classic Lowell & Guilbert (1970) 717

model, but differ at Granisle and Morrison in the absence of a large quartz-sericite-pyrite (phyllic) 718

alteration zone. These differences likely reflect contrasts in the composition of the country rocks 719

at the three deposits. The Bell porphyry is hosted by Skeena Group siltstone on its western side, 720

which underwent quartz-sericite alteration, and dark green marine tuff (Hazelton Group) on its 721

eastern side, which underwent sericite-carbonate alteration (Carson and Jambor 1974; Dirom et 722

al. 1995). At Granisle, the porphyry is mainly hosted by intermediate-composition tuffs and 723

breccias interlayered with pebble conglomerate of the Lower Jurassic Hazelton Group, which 724

underwent sericite-carbonate alteration (Carson and Jambor 1974; Dirom et al. 1995). In contrast, 725

at Morrison, the country rocks are mainly siltstones of the Upper Jurassic Ashman Formation, 726

which predominantly underwent chlorite-carbonate alteration (Carson and Jambor 1976; Ogryzlo 727

et al. 1995). The different proportions of sericite, chlorite, and carbonate in these alteration zones 728

may reflect different proportions of feldspathic and ferromagnesian minerals in the protoliths, 729

with sericite being more abundant in feldspathic rocks, and chlorite and carbonate in more mafic 730

rocks. Such differences are observed in other porphyry deposits in British Columbia, such as at 731

Schaft Creek, where chloritic alteration predominates due to the mafic nature of the country 732

rocks (Scott et al. 2008). 733

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Hollister (1975) proposed a diorite model for some British Columbia low-silica alkaline 734

porphyries, in which the phyllic alteration zone is poorly developed or absent because iron is not 735

totally consumed by sulfur in the hydrothermal fluid; consequently, chlorite-rich alteration 736

minerals are formed rather than sericite. A similar explanation may apply to Morrison. 737

738

Conclusions 739

Porphyry Cu mineralization at Morrison is shown to be spatially, temporally, and probably 740

genetically related to Eocene-age calc-alkaline plagioclase-hornblende-biotite porphyry stocks 741

and dikes with continental arc affinity, which were intruded into sedimentary rocks of the Upper 742

Jurassic Ashman Formation of the Bowser Group. Two molybdenite samples yielded Re-Os ages 743

of 52.54 ± 0.22 Ma and 53.06 ± 0.22 Ma, in good agreement with the intrusive age (52.54 ± 1.05 744

Ma for phase A porphyry). Potassic (biotite), localized phyllic (sericite), and widespread 745

propylitic (chlorite-carbonate) alteration was developed during the early stages of hydrothermal 746

fluid circulation. The potassic alteration mainly occurs in the plagioclase-hornblende-biotite 747

porphyry, and is closely associated with Cu mineralization. The main Cu ore minerals, 748

chalcopyrite and minor bornite, are primarily located in stockwork veinlets and quartz-sulfide 749

veins, but also occur as disseminations within the altered porphyry. Propylitic alteration is 750

primarily present in peripheral sedimentary rocks, and does not carry Cu mineralization, whereas 751

sericite-carbonate and argillic alteration are associated with a later stage fluid of possible 752

meteoric groundwater origin, and are localized along the East Fault. Sericite-carbonate alteration 753

is restricted to halos around late stage polymetallic sulfide-carbonate veins, and is locally 754

overprinted by supergene argillic alteration. 755

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The vein paragenesis at Morrison can be classified into five types associated with three main 756

stages of mineralization. Stage 1 veins are closely related to strong potassic alteration, and 757

comprise: early type 1 biotite veins consisting of fine-grained biotite but lacking sulfides; type 2 758

stockwork veinlets of chalcopyrite ± quartz-bornite-pyrite; and type 3 veins chalcopyrite-quartz-759

bornite-pyrite veins with sulfide in the center line. Type 4 veins (stage 2) are quartz-rich, with 760

minor molybdenite and pyrite, and are related to weak sericitic (phyllic) alteration. Type 5 veins 761

(stage 3) are dolomite-rich with polymetallic sulfides (sphalerite-galena-arsenopyrite-pyrite ± 762

chalcopyrite) and minor quartz, and are related to sericite-carbonate alteration. 763

Oxygen isotope (δ18

Ofluid = 3.7 to 6.3‰), sulfur isotope (δ34

S = -0.2 to 0.8‰), and fluid 764

inclusion data indicate that the first stage of mineralization involved a high-temperature (400° to 765

526°C) and saline (39.8 to 47.8 wt.% NaCl equiv.) fluid of likely magmatic origin, which was 766

responsible for potassic alteration and Cu precipitation. The second stage, with minor 767

molybdenum mineralization, was related to mainly magmatic fluids, but possibly with a minor 768

cooler and dilute groundwater component as indicated by oxygen isotope (δ18

Ofluid = 0.3 to 769

3.4‰), sulfur isotope (δ34

S = -2.1 to -1.2‰), and fluid inclusion data (Th: 320° to 421°C; 770

salinities: 37.0 to 43.1 wt.% NaCl equiv.). In contrast, the third stage of polymetallic sulfide-771

carbonate veining was formed predominantly from meteoric groundwater which had undergone 772

partial isotopic exchange with country rocks: δ18

Ofluid = -2.3 to 3.9‰ (derived from early vein 773

quartz); δ18

Ofluid = 0.8 to 6.3‰ (derived from late vein dolomite); δ34

S = -7.1 to -5.6‰; δ13

C = 774

0.6 to 0.7‰; fluid inclusion homogenization temperatures = 163° to 218°C, and salinities = 3.1 775

to 3.9 wt.% NaCl equiv.). 776

The Morrison porphyry deposit is similar to other Late Cretaceous-early Cenozoic porphyry 777

deposits in central British Columbia in terms of host rocks, age, alteration styles, and ore 778

Page 34 of 90



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mineralogy. These deposits are all related to subduction along the western margin of North 779

America in the Paleogene. Plate tectonic reconstructions for the Late Cretaceous to Eocene show 780

oblique subduction between the Kula-Farallon and North American plate, which generated 781

intermediate-to-felsic continental-arc magmatism with calc-alkaline affinity, and with high 782

potential for the formation of porphyry Cu ± Mo ± Au deposits such as Morrison. 783

784

Acknowledgements 785

This work was funded by a Strategic Projects Grant from the Natural Sciences and 786

Engineering Research Council of Canada (STPGP413264-11). Pacific Booker Minerals Inc. is 787

thanked for providing access to drill cores of the Morrison property. We also thank Martin von 788

Dollen, Diane Caird, and Robert Dokken for their help with sample preparation, XRD analysis, 789

and zircon U-Pb dating, respectively. Peter Hollings and an anonymous reviewer are thanked for 790

their helpful comments on the manuscript, and Tony Barresi is thanked for advice on 791

stratigraphic relationships.792

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793

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Economic Geology, 98: 1515-1533. 1008

Richards, J.P. 2011. High Sr/Y magmas and porphyry Cu ± Mo ± Au deposits: Just add water. 1009

Economic Geology, 106: 1075–1081. 1010

Roedder, E. 1984. Fluid inclusions. Mineralogical Society of America, Reviews in Mineralogy, 1011

12, 644 p. 1012

Rutherford, M.J., Devine, J.D. 1988. The May 18, 1980, eruption of Mount St. Helens. 3. 1013

Stability and chemistry of amphibole in the magma chamber. Journal of Geophysical 1014

Research, 93: 11949–11959. 1015

Scott, J.E., Richards, J.P., Heaman, L.M., Creaser, R.A., and Salazar, G.S. 2008. The Schaft 1016

Creek porphyry Cu-Mo-(Au) deposit, Northwestern British Columbia. Exploration and 1017

Mining Geology, 17: 163–196. 1018

Selby, D., and Creaser, R.A. 2004. Macroscale NTIMS and microscale LA-MC-ICP-MS Re–Os 1019

isotopic analysis of molybdenite: Testing spatial restrictions for reliable Re–Os age 1020

determinations, and implications for the decoupling of Re and Os within molybdenite. 1021

Geochimica et Cosmochimica Acta, 68: 3897–3908. 1022

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Simpson, R.G., and Geo, P. 2007. Mineral resource update Morrison project. Pacific Booker 1030

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Simon, A.C., and Ripley, E.M. 2011. The role of magmatic sulfur in the formation of ore 1032

deposits. Reviews in Mineralogy and Geochemistry, 73: 513–578. 1033

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laser heating. Geochimica et Cosmochimica Acta, 59: 5223–5231. 1062

Villeneuve, M.E., and MacIntyre, D.G. 1997. Laser 40

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the Granisle and Bell porphyry Copper deposits, British Columbia: Economic Geology, 75: 1068

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systematics of the Babine porphyry Cu deposits, British Columbia: Implications for fluid 1074

evolution of porphyry systems. Economic Geology, 89: 1518–1541. 1075

1076

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Figure Captions 1077

Fig. 1. Terrane map of the British Columbian Cordillera, showing the location of the Morrison 1078

and other early Cenozoic porphyry Cu deposits in the central Stikinia terrane. Abbreviation: TFI 1079

= Takla-Finlay-Ingenika fault system. Modified from Nelson and Colpron (2007). 1080

1081

Fig. 2. Regional geological map of the Babine Lake area and its surroundings. Modified from 1082

Massey et al. (2005). 1083

1084

Fig. 3. (a) Geological map of the Morrison porphyry Cu deposit; (b) Copper contour map. 1085

Modified from Carson and Jambor (1976) and Simpson and Geo (2007). 1086

1087

Fig. 4. Hand samples of the main intrusive units in the Morrison deposit: (A) Phase A 1088

plagioclase-hornblende-biotite porphyry (sample MO001). (B) Phase B plagioclase-hornblende-1089

biotite porphyry (sample MO059). (C) Phase C plagioclase-hornblende-biotite porphyry (sample 1090

MO128). (D) Phase C plagioclase-hornblende-biotite porphyry (centre) cutting phase A 1091

plagioclase-hornblende-biotite porphyry (sample MO012). 1092

1093

Fig. 5. Hand samples showing different alteration types at Morrison: (A) Potassic-altered phase 1094

A porphyry (sample MO100). (B) Potassic-altered phase A porphyry after staining. The yellow 1095

represents the secondary potassium feldspar (sample MO014). (C) Chlorite alteration 1096

overprinting potassic alteration in phase B porphyry (sample MO004). (D) Clay-carbonate-1097

altered phase C porphyry (sample MO037). 1098

1099

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Fig. 6. Alteration styles in thin section (A–D and F in plane-polarized light; E in cross-polarized 1100

light). (A) Potassic alteration: hydrothermal biotite replaces a hornblende phenocryst, and also 1101

occurs in the matrix, whereas a biotite phenocryst was not affected (sample MO014). (B) Strong 1102

potassic alteration: the rims of igneous biotite phenocrysts are overgrown by hydrothermal 1103

biotite (sample MO010). (C) Chlorite-carbonate alteration: chlorite mainly replaces hornblende 1104

phenocrysts, whereas plagioclase and biotite phenocrysts are affected by hydrothermal carbonate 1105

minerals (sample MO004). (D) Sericite halo of a quartz-carbonate-polymetallic vein, and a 1106

carbonate veinlet cutting the sericite halo (sample MO072). (E) (F) Clay-carbonate alteration: 1107

hornblende, biotite, and plagioclase phenocrysts are all replaced by clay and carbonate minerals 1108

(sample MO020). 1109

1110

Fig. 7. Vein relationships in hand samples: (A) Type 1 biotite veins in strongly potassic-altered 1111

porphyry (sample MO010). (B) Type 2 stockwork veinlets of chalcopyrite-bornite-pyrite ± 1112

quartz in potassic-altered porphyry (sample MO135). (C) Type 3 quartz-chalcopyrite-bornite-1113

pyrite vein cutting type 2 stockwork veinlets of chalcopyrite-bornite-pyrite ± quartz in potassic-1114

altered porphyry (sample MO135). (D) Type 4 quartz-molybdenite-pyrite vein cutting a type 3 1115

vein in weakly potassic-altered porphyry (sample MO099). (E) Type 4 quartz-molybdenite-1116

pyrite vein in clay-carbonate-altered porphyry, which probably overprints earlier phyllic 1117

alteration (sample MO098). (F) Type 5 dolomite-sphalerite-galena vein in clay-carbonate-altered 1118

porphyry, overprinting previous sericite-carbonate alteration (sample MO072). 1119

1120

Fig. 8. Mineralized vein paragenesis based on hand sample observations and petrographic 1121

studies. 1122

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1123

Fig. 9. Paragenetic relationships between sulfide minerals from the three mineralization stages: 1124

(photomicrographs taken in plane-polarized reflected light). Stage 1: (A) Disseminated 1125

chalcopyrite intergrown with bornite, accompanied by magnetite (sample MO144). (B) Type 3 1126

chalcopyrite-quartz vein (sample MO135). (C) Disseminated chalcopyrite with minor pyrite 1127

(sample MO031). Stage 2: (D) Pyrite overgrowing molybdenite in a type 4 quartz vein (sample 1128

MO094). (E) A type 4 quartz-molybdenite vein (sample MO094). Stage 3: (F) Pyrite overgrown 1129

by chalcopyrite and then sphalerite in a type 5 polymetallic sulfide-dolomite vein (sample 1130

MO072). (G) Sphalerite intergrown with galena in a type 5 polymetallic sulfide-dolomite vein 1131

(sample MO072). (H) Pyrite overgrown by arsenopyrite with late dolomite infill in a type 5 1132

polymetallic sulfide-dolomite vein (sample MO071). 1133

Abbreviations: Apy = arsenopyrite, Bo = bornite, Cpy = chalcopyrite, Gn = galena, Mo = 1134

molybdenite, Py = pyrite, Sp = sphalerite. 1135

1136

Fig. 10. Total alkali versus silica diagram showing the compositions of weakly potassic-altered 1137

plagioclase-hornblende-biotite porphyry intrusions from the Morrison area (after Middlemost, 1138

1994). Fresh plagioclase-hornblende-biotite porphyry samples from other Babine Lake intrusions 1139

are also shown for comparison (data from Carson and Jambor, 1974, and Ogryzlo et al., 1995). 1140

1141

Fig. 11. K2O versus SiO2 diagram showing the chemical compositions of weakly potassic-altered 1142

plagioclase-hornblende-biotite porphyry intrusions from the Morrison area (after Peccerillo and 1143

Taylor, 1976). Fresh plagioclase-hornblende-biotite porphyry samples from other Babine Lake 1144

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intrusions are also shown for comparison (data from Carson and Jambor, 1974, and Ogryzlo et 1145

al., 1995). 1146

1147

Fig. 12. (A) Primitive mantle-normalized trace element diagram, and (B) C1 Chondrite-1148

normalized REE diagram for plagioclase-hornblende-biotite porphyry samples from Morrison 1149

(primitive mantle normalization values from Sun and McDonough, 1989). Other Babine Lake 1150

igneous rocks from the Babine Lake area, and volcanic rock samples from the Eocene Ootsa 1151

Lake Group from nearby areas are shown for comparison (data from Grainger, 2000). 1152

1153

Fig. 13. Tectonic discrimination diagrams for plagioclase-hornblende-biotite porphyry from the 1154

Morrison deposit (Pearce, 1984). Abbreviations: ACM: active continental margins; MORB: 1155

Mid-ocean ridge basalt; WPB: Within plate basalts; WPVZ: within-plate volcanic zones. 1156

1157

Fig. 14. Backscattered electron images of typical zircons from plagioclase-hornblende-biotite 1158

porphyry sample MO128. Zircons show weak magmatic oscillatory zoning from center to 1159

margin, and dark areas (inclusions) were avoided during analysis. 1160

Fig. 15. U-Pb Concordia diagram for zircon laser ablation ICPMS data. The error ellipses are 1161

shown at 2σ. 1162

1163

Fig. 16. Transmitted light photomicrographs showing primary fluid inclusions from veins at 1164

Morrison. (A) A fluid inclusion assemblage with hypersaline inclusions and vapor-rich 1165

inclusions from a type 3 quartz-chalcopyrite-pyrite vein (MO135). (B, C) Hypersaline inclusions 1166

with halite + liquid + vapor ± additional daughter crystals including sylvite, anhydrite, and 1167

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opaque crystals from a type 3 quartz-chalcopyrite-pyrite vein (MO135). (D) A vapor-rich 1168

inclusion assemblage from a type 3 quartz-chalcopyrite-pyrite vein (MO135). (E) Hypersaline 1169

and vapor-rich inclusions from a type 4 quartz-molybdenite-pyrite vein (MO097). (F) Liquid-1170

rich inclusions in a dolomite from a type 5 polymetallic-sulfide-dolomite vein (MO072). 1171

1172

Fig. 17. Histograms showing homogenization temperatures and salinities of fluid inclusions from 1173

type 3 and 4 quartz-sulfide veins (stage 1 and 2), and type 5 dolomite-sulfide veins (stage 3). 1174

1175

Fig. 18. Salinity versus homogenization temperature plot of fluid inclusions from type 3 and 4 1176

quartz-sulfide veins (stages 1 and 2), and type 5 polymetallic sulfide-dolomite veins (stage 3). 1177

The green line is the halite-saturation curve, and the black line shows the evolution trend of the 1178

hydrothermal fluids from stage 1 to 2. 1179

1180

Fig. 19. Vapor-saturated NaCl-KCl-H2O phase diagram (after Roedder, 1984), showing the 1181

evolution trend of fluid inclusions containing both halite and sylvite from a type 3 quartz-1182

chalcopyrite-pyrite vein (MO135) (assuming halite is pure NaCl). The average sylvite melting 1183

temperature was 114°C, and the average halite melting temperature was 362°C. 1184

1185

Fig. 20. Evolution of the δ18

O composition of the hydrothermal fluids from stages 1 to 3 (vein 1186

types 3–5). 1187

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Fig. 1. Terrane map of the British Columbian Cordillera, showing the location of the Morrison and other early Cenozoic porphyry Cu deposits in the central Stikinia terrane. Abbreviation: TFI = Takla-Finlay-Ingenika

fault system. Modified from Nelson and Colpron (2007).

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Fig. 2. Regional geological map of the Babine Lake area and its surroundings. Modified from Massey et al. (2005).

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Fig. 3. (a) Geological map of the Morrison porphyry Cu deposit; (b) Copper contour map. Modified from Carson and Jambor (1976) and Simpson and Geo (2007).

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Fig. 4. Hand samples of the main intrusive units in the Morrison deposit: (A) Phase A plagioclase-hornblende-biotite porphyry (sample MO001). (B) Phase B plagioclase-hornblende-biotite porphyry (sample MO059). (C) Phase C plagioclase-hornblende-biotite porphyry (sample MO128). (D) Phase C plagioclase-hornblende-biotite porphyry (centre) cutting phase A plagioclase-hornblende-biotite porphyry (sample

MO012).

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Table 1. Characteristics of 3 phases plagioclase-hornblende-biotite porphyry

Plagioclase-hornblende-biotite porphyry Biotite phenocrysts

1-3 mm

1-3%

2-4 mm

2-5%

2-5 mm

8-10%Phase C

Phase B

Phase A

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Table 1. Characteristics of 3 phases plagioclase-hornblende-biotite porphyry

Hornblende phenocrysts Plagioclase phenocrysts Matrix

3-5 mm 3-5 mm fine-grained

4-7% 10-15% 75-85%


8-10% 15-20% 65-75%


12-15% 20-25% 50-60%

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Table 2. Major and trace element analyses of plagioclase-hornblende-biotite porphyry from the Morrison deposit

Sample ID Analytical

method

Det.

lim. MO001 MO017 MO038 MO093 MO120B MO128 MO139 MO142 MO149

Phase A C C B B C C C C

Weight %

SiO2 FUS-ICP 0.01 54.91 59.93 60.04 59.38 60.11 61.34 61.18 58.43 60.73

Al2O3 FUS-ICP 0.01 14.59 16.42 15.96 15.91 16.06 15.73 15.69 15.09 15.42

Fe2O3(Total) FUS-ICP 0.01 6.37 5.69 5.66 5.74 5.88 4.88 4.90 5.86 5.36

MnO FUS-ICP 0.001 0.06 0.04 0.07 0.02 0.03 0.04 0.02 0.04 0.05

MgO FUS-ICP 0.01 4.45 3.26 3.27 3.82 3.34 2.83 2.74 3.05 2.79

CaO FUS-ICP 0.01 4.99 4.72 5.13 4.24 4.06 4.38 3.86 4.15 4.45

Na2O FUS-ICP 0.01 3.18 4.65 4.29 3.37 3.86 4.33 4.00 4.20 4.24

K2O FUS-ICP 0.01 2.65 1.27 2.27 2.23 1.65 1.60 1.63 1.59 2.16

TiO2 FUS-ICP 0.001 0.84 0.84 0.77 0.81 0.78 0.79 0.76 0.71 0.69

P2O5 FUS-ICP 0.01 0.26 0.33 0.29 0.34 0.29 0.29 0.23 0.26 0.24

LOI 5.37 1.62 1.90 3.26 3.67 2.44 3.80 3.19 2.81

Total 97.67 98.77 99.65 99.12 99.73 98.65 98.81 96.57 98.94

ppm

Cs FUS-MS 0.1 3 2 2.4 3.8 2.4 3 2.4 4 3

Tl FUS-MS 0.05 0.7 0.38 0.35 1.25 0.6 0.4 0.38 0.34 0.42

Rb FUS-MS 1 76 30 39 72 39 39 39 39 38

Ba FUS-ICP 1 765 687 1716 546 889 1171 2306 11850 1313

Th FUS-MS 0.05 4.21 4.98 5.65 5.91 5.18 5.07 5.07 5.18 5.86

U FUS-MS 0.01 1.45 2.39 2.19 2.66 2.37 1.51 1.19 1.15 1.59

Nb FUS-MS 0.2 6 6.6 6.8 8.1 6.5 6.8 7.1 6.3 7.9

Ta FUS-MS 0.01 0.39 0.52 0.45 0.51 0.39 0.41 0.43 0.45 0.42

La FUS-MS 0.05 21.6 27.4 22.5 27.8 20.8 19 51 25.1 27.9

Ce FUS-MS 0.05 40.4 52.4 42.2 50 39 37.5 75.9 44.1 53.1

Pr FUS-ICP 0.01 4.99 6.41 5.35 5.95 4.85 4.83 7.79 5.24 6.45

Sr FUS-MS 2 502 905 877 775 738 866 928 1639 850

Nd FUS-MS 0.05 20 25.1 21.5 22.3 19.8 19.5 26.8 20.3 25

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Zr FUS-ICP 1 94 113 113 120 118 114 112 112 114

Hf FUS-MS 0.1 2.3 2.8 2.9 3 3 2.9 2.7 2.7 2.9

Sm FUS-MS 0.01 3.89 4.63 3.88 4.09 3.75 3.92 3.99 3.67 4.46

Eu FUS-MS 0.005 1.1 1.27 1.12 1.14 1.05 0.938 1.18 1.08 1.13

Gd FUS-MS 0.01 3.1 3.37 2.95 3.04 2.85 3.01 2.88 2.87 3.2

Tb FUS-MS 0.01 0.43 0.44 0.4 0.4 0.39 0.41 0.39 0.38 0.43

Dy FUS-MS 0.01 2.35 2.16 2.06 2.03 2.01 2.12 2.09 2.04 2.13

Y FUS-ICP 1 11 10 10 10 10 10 10 10 10

Ho FUS-MS 0.01 0.44 0.38 0.36 0.37 0.37 0.36 0.37 0.36 0.39

Er FUS-MS 0.01 1.22 0.97 1.01 1 0.95 0.96 1.04 0.95 1.01

Tm FUS-MS 0.005 0.181 0.147 0.153 0.148 0.139 0.142 0.145 0.135 0.141

Yb FUS-MS 0.01 1.11 0.89 0.95 0.91 0.89 0.86 0.94 0.85 0.93

Lu FUS-MS 0.002 0.149 0.133 0.142 0.138 0.131 0.125 0.138 0.125 0.145

V FUS-ICP 5 164 139 130 134 130 130 131 127 115

Ga FUS-MS 1 20 21 20 21 21 21 21 21 20

Ge FUS-MS 0.5 1.4 1.4 1.3 1.6 1.4 1.7 1.6 1.2 1.3

Sn FUS-MS bdl bdl bdl bdl 2 bdl 2 bdl bdl bdl

Co INAA 0.1 18.9 20.3 16.1 28.3 13.2 13.8 13.6 16.4 12.4

Cr INAA 0.5 159 77.2 93.3 123 83.4 74.4 72.2 80.3 78.7

Ni TD-ICP 1 77 42 45 67 40 47 49 46 44

Pb TD-ICP 5 7 < 5 < 5 < 5 < 5 < 5 < 5 < 5 < 5

Sc INAA 0.01 14.6 8.98 10.2 11.6 9.75 8.87 10.3 9.38 8.53

Sr/Y 45.6 90.5 87.7 77.5 73.8 86.6 92.8 163.9 85.0

[La/Yb]n 14.0 22.1 17.0 21.9 16.8 15.8 38.9 21.2 21.5

[La/Sm]n 3.6 3.8 3.7 4.4 3.6 3.1 8.3 4.4 4.0

[Dy/Yb]n 1.4 1.6 1.5 1.5 1.5 1.6 1.5 1.6 1.5

Abbreviations: bdl = at or below detection limit; FUS-ICP = fusion inductively-coupled plasma; FUS-MS = fusion inductively-coupled plasma mass

spectrometry; ICP = inductively-coupled plasma; INAA = instrumental neutron activation analysis; TD-ICP = total acid digestion inductively-coupled plasma.

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Table 3. Zircon U-Pb data for phase C plagioclase-hornblende-biotite porphyry sample MO128206

Pb/238

Ua

2 σ207

Pb/206

Pb 2 σ238

U/206

Pb 2 σ age (Ma) error (Ma)

MO128-1 0.0574 0.0014 121.23 4.09 53.0 1.8

MO128-3 0.0522 0.0012 122.41 4.82 52.5 2.1

MO128-4 0.0535 0.0010 123.11 4.34 52.2 1.8

MO128-5 0.0558 0.0016 122.53 4.63 52.4 2.0

MO128-6 0.0522 0.0009 123.68 4.06 51.9 1.7

MO128-7 0.0605 0.0015 121.49 4.38 52.8 1.9

MO128-8 0.0560 0.0011 124.56 4.68 51.5 1.9

MO128-9 0.0624 0.0017 121.71 4.38 52.7 1.9

MO128-10 0.0594 0.0018 123.84 4.78 51.8 2.0

MO128-11 0.0554 0.0012 120.34 4.07 53.3 1.8

MO128-12 0.0611 0.0047 116.89 4.30 54.9 2.0

MO128-13 0.0521 0.0011 124.44 4.61 51.6 1.9

MO128-14 0.0566 0.0016 123.13 4.80 52.1 2.0

MO128-15 0.0538 0.0012 119.83 4.19 53.6 1.9

MO128-16 0.0525 0.0008 122.00 4.46 52.6 1.9

MO128-17 0.0553 0.0012 123.78 4.84 51.9 2.0

MO128-18 0.0598 0.0015 117.24 4.13 54.8 1.9

MO128-19 0.0513 0.0010 122.22 4.30 52.5 1.8

MO128-20 0.0530 0.0012 123.87 4.70 51.8 2.0

MO128-21 0.0502 0.0009 120.14 4.37 53.4 1.9

MO128-22 0.0678 0.0020 115.13 4.48 55.7 2.2

MO128-23 0.0548 0.0016 116.57 4.84 55.1 2.3

MO128-24 0.0546 0.0011 121.67 4.29 52.8 1.9

MO128-25 0.0522 0.0010 119.26 4.72 53.8 2.1

MO128-26 0.0541 0.0013 119.95 4.64 53.5 2.1

MO128-27 0.0521 0.0012 117.20 4.70 54.8 2.2

MO128-28 0.0569 0.0016 118.34 4.80 54.2 2.2

MO128-29 0.0609 0.0014 118.18 4.26 54.3 1.9

MO128-30 0.0645 0.0036 117.42 5.54 54.7 2.6

aThese are the apparent ages and have not been corrected for common Pb, though a

common Pb correction is negligible for these data and is within reported errors.

Sample name

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Table 3. Zircon U-Pb data for phase C plagioclase-hornblende-biotite porphyry sample MO128

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Table 4. Molybdenite Re-Os data

Sample Re ppm ± 2s187Re ppm ± 2s

187Os ppb ± 2s

MO-094 431.7 1.1 271.3 0.7 237.6 0.2

MO-097 559.5 1.5 351.7 0.9 311 0.2

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Model Age (Ma) ± 2s (Ma)

52.54 0.22

53.06 0.22

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Table 5. Oxygen isotope data for quartz and dolomite from vein types 3–5

Sample Vein type

MO135-Quartz 3

MO137-Quartz 3

MO144-Quartz 3

MO151-Quartz 4

MO098-Quartz 4

MO072-Quartz 5 (Early)

MO083-Quartz 5 (Early)

MO072-Dolomite 5 (Late)

MO083-Dolomite 5 (Late)

Note: 1Maximum oxygen compositions of fluid are calculated with the mean homogenization temperature plus the standard deviation of the mean.

2Minimum oxygen compositions of fluid are calculated with the mean homogenization temperature minus the standard deviation of the mean.

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Table 5. Oxygen isotope data for quartz and dolomite from vein types 3–5

Temperature used for calculation of fluid composition δ18OVSMOW (‰)

8.4

8.2

7.3

6.1

7.7

10.7

9.4

13.6

16.6

Maximum oxygen compositions of fluid are calculated with the mean homogenization temperature plus the standard deviation of the mean.

Minimum oxygen compositions of fluid are calculated with the mean homogenization temperature minus the standard deviation of the mean.

464° ± 42°C

363° ± 30°C

200° and 300°C

190° ± 16°C

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Maximum δ18Ofluid (‰)

1 Minimum δ

18Ofluid (‰)

2

6.3 4.8

6.1 4.6

5.2 3.7

1.9 0.3

3.4 1.9

3.9 -1.0

2.5 -2.3

3.1 1.1

6.1 4.0

Maximum oxygen compositions of fluid are calculated with the mean homogenization temperature plus the standard deviation of the mean.

Minimum oxygen compositions of fluid are calculated with the mean homogenization temperature minus the standard deviation of the mean.

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Table 6. Sulfur isotope data for sulfides from stage 1-3 mineralization

Sample Mineral δ34

SCDT (‰)

MO95 Mixture of chalcopyrite and pyrite -0.8

MO96 Mixture of chalcopyrite and pyrite 0.2

MO042 Molybdenite -2.1

MO115 Pyrite -1.2

MO072 Pyrite -5.8

MO072 Sphalerite -7.1

MO083 Pyrite -5.6

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Table 6. Sulfur isotope data for sulfides from stage 1-3 mineralization

Comments

Stage 1, disseminated sulfides

Stage 1, disseminated sulfides

Stage 2, from type 4 quartz-molybdenite-pyrite vein

Stage 2, from type 4 quartz-molybdenite-pyrite vein

Stage 3, from type 5 polymetallic-sulfide-dolomite vein



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Table 9 Characteristics of three economic porphyry deposits (Bell, Granisle, and Morrison) in the Babine Lake area

Deposit nameResource

(MT)Average grades Country rock

Porphyry

age

Morrison 270

0.39% Cu; 0.2

g/t Au; 0.005%

Mo

Sedimentary rocks

(Siltstones, silty argillites,

and minor conglomerates)

of Upper Jurassic Ashman

Formation of the Bowser

Group

52.2 ± 0.37

Ma (U-Pb

zircon age)

Bell 77.2

0.48% Cu; 0.17

g/t Au; <0.01%

Mo

Siltstone of the Skeena

Group in the west of the

main intrusion, and dark

green marine tuff of the

Hazelton Group in the east

of the main intrusion

52.0 ± 0.5

Ma (Biotite

Ar-Ar age)

Granisle 52.7

0.43% Cu; 0.13

g/t Au; 0.005%

Mo

Intermediate tuff and

breccia interlayered with

pebble conglomerate of the

Lower Jurassic Hazelton

Group

51.2 ± 0.6

Ma (Biotite

Ar-Ar age)

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Table 9 Characteristics of three economic porphyry deposits (Bell, Granisle, and Morrison) in the Babine Lake area

Porphyry intrusion

compositionAlteration types

Diorite to granodiorite

Potassic alteration; chlorite-

carbonate alteration; argillic

alteration; lack of typical phyllic

alteration except as narrow halos

to molybdenite-bearing veins.

Copper mineralization is closely

related to the potassic alteration.



carbonate alteration; sericite-

carbonate alteration; quartz-

sericite alteration. Copper

mineralization is closely

associated with the quartz-

sericite alteration.



carbonate alteration; sericite-

carbonate alteration. Copper

mineralization is closely related

to the potassic alteration.

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Mineralization styles Ore mineralsHythermal fluid related

with mineralization

Stockwork veins and

dissemination

Primarily chalcopyrite,

minor bornite, and rare

molybdenite, sphalerite,

galena, gold, and silver

Magmatic water

Stockwork veins and

dissemination


moderate bornite, minor

chalcocite, and rare

molybdenite, sphalerite, and

galena, gold, and silver.

Mixture of magmatic

water and groundwater

Stockwork veins and

dissemination


minor bornite, and rare

molybdenite, gold and silver

Magmatic water

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Draft

Source(s)

Carson and Jambor (1974);


Ogryzlo et al. (1995)


Carson et al. (1976); Wilson et

al. (1980); Zaluski et al.

(1994); Dirom et al. (1995);

Maclntyre et al. (2001)


Fahrni et al. (1976); Wilson et

al. (1980); Zaluski et al.

(1994); Dirom et al. (1995);

Maclntyre et al. (2001)

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geology and age of the morrison porphyry cu-au-mo 1 1 geology and age of the morrison porphyry...

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