Canadian Journal of Earth Sciences

Between the -- Deer Trail Group, an intermediate age unit between the Mesoproterozoic Belt-Purcell Supergroup and the Windermere Supergroup in northeastern Washington, U.S.A.

Journal: Canadian Journal of Earth Sciences

Manuscript ID cjes-2019-0188.R1

Manuscript Type: Article

Date Submitted by the 15-Jan-2020 Author: Complete List of Authors: Box, Stephen;Draft US Geological Survey Pacific Region, Mineral Resource Program; Pritchard, Chad; Eastern Washington University, Department of Geology Stephens, Travis; Eastern Washington University, Department of Geology O'Sullivan, Paul; GeoSep Services, ; GeoSep Services,

Columbia, , detrital zircon U-Pb geochronology, Keyword: , petrochemistry

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

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1 Between the supercontinents – Mesoproterozoic Deer Trail Group, an intermediate age 2 unit between the Mesoproterozoic Belt-Purcell Supergroup and the Neoproterozoic 3 Windermere Supergroup in northeastern Washington, U.S.A.

4 Box, Stephen E.1, Pritchard, Chad J.2, Stephens, Travis S. 2, and O’Sullivan, Paul B.3 5 1US Geological Survey, 904 West Riverside Ave., Rm. 202, Spokane, WA 99201 6 2Eastern Washington University, Department of Geology, Cheney, WA 99004-2439 7 3GeoSep Services, 1521 Pine Cone Rd., Moscow, ID 83843, USA

8 ABSTRACT

9 Mesoproterozoic and Neoproterozoic basins in western North America record the 10 evolving position of the Laurentian craton within two supercontinents during their growth 11 and dismemberment: Columbia (Nuna) and Rodinia. The westernmost exposures of the 12 Columbia rift-related Belt-Purcell Supergroup are preserved in northeastern 13 Washington, structurally overlain by Draftthe Deer Trail Group and depositionally overlying 14 Neoproterozoic Windermere Supergroup. It has been disputed whether the Deer Trail 15 Group is correlative with the Belt-Purcell Supergroup, or younger. To help resolve the 16 uncertain correlation of these units and their bearing on supercontinent evolution, we 17 characterized the detrital zircon age populations of units from the Deer Trail Group, the 18 Windermere Supergroup, and the Belt-Purcell Supergroup in northeastern Washington. 19 These data show that the western part of the Columbia supercontinent (now located in 20 Australia and eastern Antarctica) remained attached to western and continued 21 to supply 1600-1500 Ma detrital zircon grains to the Belt-Purcell Supergroup until after 22 ca. 1391 Ma. Deer Trail Group is younger than Belt-Purcell strata, with the basal unit 23 younger than ca. 1362 Ma and a middle unit younger than ca. 1300 Ma. The Deer Trail 24 Group has a pre-Grenville-age provenance from the southwestern US and possibly east 25 Antarctica. The Buffalo Hump Formation is younger than the Deer Trail Group, with 26 Grenville-age (ca. 1112 Ma) detrital zircon grains and a detrital zircon signature like that 27 of the overlying Neoproterozoic Windermere Supergroup. We interpret the Deer Trail 28 Group to have been deposited during the rift-demise of supercontinent Columbia and 29 before the Grenville-age assembly of supercontinent Rodinia.

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30 KEYWORDS: , Proterozoic, supercontinent Columbia, Rodinia, Laurentia, 31 detrital zircon U-Pb geochronology, petrochemistry, Belt-Purcell Supergroup, Deer Trail 32 Group, Windermere Supergroup, Washington

33

34 INTRODUCTION

35 Our understanding of how western Laurentia fit into the great Proterozoic 36 supercontinents (Columbia [sometimes called Nuna]: ca. 1.7-1.3 Ma, Zhao et al., 2004; 37 Rodinia: 1.1-0.75 Ma, Li et al., 2008) has improved in recent years through careful 38 correlation of Mesoproterozoic sedimentary sequences and their detrital sources (Doe 39 et al., 2013; Medig et al., 2014; Mulder et al., 2018). The Mesoproterozoic Belt-Purcell 40 Supergroup (1.46-1.39 Ga), a rift-related sequence straddling the western US- 41 border (Fig. 1), provides important evidence of an exotic western sediment source 42 during the late stages of the supercontinentDraft Columbia (Ross et al., 1992; Doe et al. 43 2013). The hypothesized western source, interpreted to be a craton that was formerly 44 adjacent to western Laurentia, provided abundant detrital zircons with igneous ages that 45 are rarely found in Laurentia (1.61-1.48 Ga, referred to as the North American Magmatic 46 Gap [NAMG]: Ross and Villeneuve, 2003). Igneous rocks in this age range are well 47 represented in present-day southeastern Australia, which most authors consider to be 48 the missing western Belt-Purcell sediment source (Ross et al., 1992; Karlstrom et al., 49 2001; Ross and Villeneuve, 2003; Belousova et al., 2009; Jones et al., 2015). This non- 50 Laurentian detrital zircon component disappeared from Belt-Purcell strata during the 51 transition from Piegan Group to Missoula Group of the Belt-Purcell Supergroup (ca. 52 1440; Evans et al., 2000), suggesting removal of the exotic western cratonic source. 53 The basin was overwhelmed with a flood of debris from the present-day southwestern 54 US at about the same time (Ross and Villeneuve, 2003; Link et al., 2007). Whether the 55 absence of western source debris represented rift-transport of that source away from 56 the western margin of the Belt-Purcell basin or simply drainage diversion, topographic 57 diminishment or dilution by influx of Laurentian debris is unclear.

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58 To better understand the evolution of detrital sources to the western Belt-Purcell basin, 59 we studied the westernmost exposures of the Belt-Purcell Supergroup and overlying 60 Proterozoic units near Chewelah, Washington, USA. Here the Belt-Purcell Supergroup 61 is structurally overlain by the Mesoproterozoic Deer Trail Group (Miller, 2001), which 62 had previously been interpreted as correlative western facies of the Belt-Purcell 63 Supergroup that was overthrust from the west (Miller and Whipple, 1989). The Deer 64 Trail Group is stratigraphically overlain by the Neoproterozoic Windermere Supergroup 65 (Miller et al., 1973), a widespread unit in western Laurentia recording late 66 Neoproterozoic rifting of supercontinent Rodinia (Stewart, 1972; Lund et al., 2010). 67 Cambrian , the basal unit of the Paleozoic passive margin sequence in western 68 Laurentia, unconformably overlies the Proterozoic strata (Gehrels and Pecha, 2014).

69 This study focuses on the detrital zircon geochronology of the Belt-Purcell Supergroup, 70 Deer Trail Group and Windermere Supergroup near Chewelah, Washington to constrain 71 both the age and provenance of theseDraft units. The results allow us to: (1) test the 72 correlation of the Deer Trail Group with the Belt-Purcell Supergroup, (2) test for the 73 presence of western-sourced NAMG zircons in the western Belt-Purcell section and in 74 potential stratigraphic equivalents (Deer Trail Group) overthrust from the west, and (3) 75 compare that provenance to that of the Neoproterozoic Windermere Supergroup. 76 Comparison of detrital zircon populations can also be made with published data from 77 Belt-Purcell strata farther east (e.g. Ross and Villeneuve, 2003; Lewis et al., 2010), from 78 other western North American Neoproterozoic units (Yonkee et al., 2014) and from the 79 unconformably overlying Cambrian Addy Formation (Linde et al., 2014; 2018).

80 REGIONAL SETTING

81 The Mesoproterozoic Belt-Purcell Supergroup underlies a broad area of western 82 , central and northern , and northeastern Washington in the United States 83 (Fig. 1), narrowing to the north into southern and in Canada 84 (e.g. Box et al, 2012). The Mesoproterozoic Deer Trail Group has only been mapped 85 west of the Belt-Purcell Supergroup in northeastern Washington and the northwestern 86 tip of Idaho (Miller et al, 1999). In contrast, Neoproterozoic Windermere Supergroup 87 and correlative rocks occur in narrow, fault-bounded blocks along a broadly north-south

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88 trend in the western US and western Canada (Lund et al., 2010) between the thinned 89 western margin of Laurentian crust (the initial Sr isotope 0.706-line of Armstrong, 1988) 90 and the eastern front of the Sevier fold-thrust belt (Fig. 1). Cordilleran Mesozoic 91 convergent deformation displaced the Proterozoic and Paleozoic sections and, in 92 places, their underlying basement, over 200 km to the east from Late Jurassic to early 93 Eocene time (DeCelles, 2004; Bedrosian and Box, 2016).

94 In the Chewelah area of northeastern Washington (Fig. 2), Belt-Purcell Supergroup 95 strata dip monoclinally to the west in the footwall of the Jumpoff Joe (Fig. 3). 96 This fault is the basal thrust of the east-vergent Mesozoic Colville fold-and-thrust belt 97 (Miller, 2001; Cheney and Zieg, 2016) and the eastern front of the arcuate 400 km long 98 Kootenay Arc (Watkinson and Ellis, 1987). Paleozoic strata with a basal Cambrian 99 quartzite unconformably overlie the Belt-Purcell strata in the lower plate below the 100 thrust, and typically all these strata lack even a low-grade structural fabric. The upper 101 plate of the fold-and-thrust belt consistsDraft of Mesoproterozoic Deer Trail Group strata 102 disconformably overlain by Neoproterozoic Windermere Supergroup and unconformably 103 overlain by lower Paleozoic strata with a basal Cambrian quartzite. Subsidiary thrust 104 faults imbricate the strata of the upper thrust plate and the strata typically bear a low 105 greenschist facies structural fabric.

106 Mesoproterozoic Belt-Purcell Supergroup

107 Belt-Purcell Supergroup strata crop out east of the town of Chewelah and consist of 4 108 groups (in ascending order): Lower Belt (Prichard Formation), Ravalli, Piegan and 109 Missoula groups (figs. 2, 3). The Prichard Formation consists of almost 4 km of pyritic 110 argillite with lesser interbedded siltite and quartzite, deposited by northeast-flowing 111 currents (Cressman, 1989; Miller, 2001). These strata are intruded by diabasic sills 112 (some intruded when strata were unconsolidated) with zircon U-Pb ages typically about 113 1.46 Ga (Anderson and Davis, 1995; Sears et al., 1998; Schandl and Davis, 2000). The 114 overlying Ravalli Group consists of (Fig. 4a) and lesser argillite horizons 115 deposited in northeast-flowing streams and marine or lacustrine deltaic systems (Mauk 116 and White, 2004). The overlying Piegan Group is comprised of bedded and 117 dolomite and calcareous quartzite, siltite and argillite (Winston, 2007). Detrital zircons

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118 from these 3 units at several localities to the east all record a prominent age population 119 peak between 1.6-1.5 Ga, within the NAMG (Ross and Villeneuve, 2003; Link et al., 120 2007; Lewis et al., 2010). The Missoula Group (and correlative Lemhi Group) records 121 the influx of mostly coarser, quartzofeldspathic detritus from the east and south after 122 initial bimodal volcanism. Deposition of the Missoula Group is thought to range from as 123 old as ca. 1450 Ma (Evans et al., 2000) to as young as ca. 1390 Ma (Link et al., 2016; 124 Constenius et al., 2017). Detrital zircons from the Missoula Group and correlative strata 125 lack the NAMG age population peak and have age population peaks between 1.8-1.6 126 Ga and 2.7-2.5 Ga, characteristic of Laurentian (North American) basement. 127 Deformation of the Belt-Purcell Supergroup is inferred to have ended sedimentation, 128 and deformed Belt-Purcell strata are crosscut by plutonic rocks with ages ca. 1370 Ma 129 (McMechan and Price, 1982; Doughty and Chamberlain, 1996; Aleinikoff et al., 2012; 130 McFarlane, 2015).

131 Mesoproterozoic(?) Deer Trail GroupDraft

132 The Deer Trail Group (Campbell and Loofbourow, 1962; Evans, 1987; Miller, 2001) is 133 exposed west and north of the town of Chewelah, Washington (Fig. 2). The Deer Trail 134 section consists of two shallowing upward sequences. The lowest unit, whose base is 135 not exposed, is the Togo Formation, an argillaceous unit with thinly interbedded siltite to 136 fine quartzose horizons. Locally, turbidite depositional features are 137 discernible, suggesting the unit was deposited below storm wave base (>100 m?). The 138 depositional environment appears to shallow upward into the Wabash-Detroit 139 Formation, comprised of dolomitic argillite and quartzites with common 140 beds that indicate deposition near sea level. It is capped by a thick, finely laminated 141 dark argillite with common soft-sediment deformation (McHale Slate), interpreted as 142 deposited in a subwavebase marine environment. This is overlain by the Stensgar 143 Dolomite, which locally preserves algal structures that again indicate the unit was 144 deposited near sea-level. Magnesite within the Stensgar Dolomite was mined for 145 magnesium content from early to mid-20th century (Campbell and Loofbourow, 1962; 146 Bamford, 1970).

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147 The Buffalo Hump Formation has been considered to be the youngest unit in the Deer 148 Trail Group. It overlies the Stensgar Dolomite in either conformable (Campbell and 149 Loofbourow, 1962; Miller and Whipple, 1989) or unconformable (Evans, 1987) contact. 150 Lithology of the Buffalo Hump Formation consists of interbedded argillite and quartzose 151 sandstone, locally with clasts up to cobble size. The unit is characterized by rapid 152 lateral thickness variations and facies changes between argillite and quartzose 153 sandstone. Detrital zircons from the Buffalo Hump analyzed by Ross et al. (1992) 154 included a few Grenville-age grains between 1.22-1.07 Ga, indicating a clearly younger 155 age than known Belt-Purcell strata.

156 Neoproterozoic Windermere Supergroup

157 The Neoproterozoic Windermere Supergroup in the Chewelah area consists of two 158 formations: Huckleberry Formation and overlying Monk Formation (Miller et al., 1973; 159 Miller, 2001). The Huckleberry FormationDraft consists of a lower conglomeratic member 160 and an upper, mafic-volcanic member. The conglomerate varies from pebbly sandstone 161 to clast-supported, cobble conglomerate composed of quartzose and calcareous 162 sedimentary clasts. The volcanic member consists of relatively massive fine-grained 163 basalt, generally finely porphyritic with plagioclase and clinopyroxene phenocrysts and 164 fine calcite or chlorite amygdules. Upper Huckleberry Formation basalt was dated by 165 the Sm-Nd isochron method at 762 + 44 Ma (Devlin and Bond, 1988). Sparse basaltic 166 dikes within the Deer Trail Group are lithologically similar to the Huckleberry volcanic 167 member and may be feeders to it. The Monk Formation is quite variable in thickness (0- 168 300 m) and lithology, ranging from argillite to quartz sandstone to carbonate- and 169 quartzite-clast, matrix- and clast-supported conglomerate to basaltic flows and flow 170 breccia. Both conglomeratic units have been interpreted as glacial marine deposits 171 (Link et al., 1994).

172 Cambrian Addy Quartzite

173 A thick lower Cambrian quartzite unconformably overlies each of the above units in 174 various places in the Chewelah area (Miller, 2001). The unit appears to thin (from ca. 175 1400 m to 200-400 m) and its base youngs (from early to middle Cambrian) to the east

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176 in the map area (Fig. 2; Cheney and Zieg, 2016). The locally crossbedded unit is 177 interpreted as beach and shallow marine sandstone and is overlain by lower and upper 178 Paleozoic marine carbonates and shales with less common sandstone. Detrital zircon 179 from the Addy Quartzite was presented by Linde et al. (2014, 2018) and will be 180 discussed later in this paper in comparison to our new data from underlying Proterozoic 181 strata.

182 ANALYTICAL METHODS

183 Sampling and preparation

184 Samples were obtained from roadcuts and bedrock outcrops in the Chewelah area and 185 located using GPS on published geologic maps (locations shown on Fig. 2; latitudes 186 and longitudes of sample sites given in Supplemental File 1; Evans, 1987; Miller, 2001). 187 Sample preparation included standard mineral separation including: 1) crush and sieve 188 with a 300 micron sieve cloth, 2) crushedDraft samples were washed thoroughly to remove 189 any clay and silt particulates and then dried, 3) heavy and light constituents were 190 separated via float separation, 4) magnetic constituents were then removed using a 191 Frantz magnetic separator, and 5) a second float separation was conducted to separate 192 zircons and apatite from lighter heavy minerals. Zircon grains (both standards and 193 unknowns) were then mounted in 1 cm2 epoxy wafers and ground down to expose 194 internal grain surfaces prior to final polishing. These wafers were then etched in 5.5N 195 HNO3 for 20.0 seconds at 21°C to thoroughly clean the surface of the grains prior to 196 analysis. Grains, and the locations for laser spots on these grains, were selected for 197 analysis from all sizes and morphologies present using transmitted light with an optical 198 microscope at a magnification of 2000x.

199 Detrital Zircon Analyses

200 U-Pb zircon ages were determined using a Finnigan Element2 ICP-MS combined with a 201 New Wave UP-213 laser ablation system at the Peter Hooper GeoAnalytical Lab at 202 Washington State University with GeoSep Services. The laser repetition rate was 5.0 203 Hz, spot size of 15-micron diameter, with argon and helium used as the carrier gas for 204 the ICP-MS and laser ablation system, respectively. Data collection is described in

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205 Chang et al. (2006). FC-1 (about 1099 Ma) and Tardree (about 61.23 Ma) were used as 206 standards.

207 Typically, about 100 grains were analyzed per sample (two samples yielded fewer 208 zircons). Detrital zircon results were excluded from subsequent analysis if concordance 209 (206Pb/238U age divided by the 207Pb/206Pb age) values were below 80% or above 105%. 210 Analyses were also excluded if either of the following applied: U value above 800 ppm, 211 or 2-sigma uncertainties in any age values >10% (all analyses shown in data tables in 212 Supplemental File 1). The excluded analyses were removed before the following 213 operations. The reported ages are based on the 206Pb/207Pb ages for grains > 1.0 Ga 214 (Gehrels and Pecha, 2014), which include all new ages reported here. Age populations 215 were identified using the AgePick macro for Microsoft Excel (Gehrels, 2009) and results 216 of this operation for all samples are shown in Table 1. The age distributions of the 217 detrital zircon samples are shown by probability distribution plots (Gehrels, 2010). 218 Metadata on analytical instruments andDraft settings, as suggested in Horstwood et al. 219 (2016), is presented in Supplemental File 2.

220 Geochemical Analyses

221 Three mafic igneous samples were crushed, powdered and analyzed by XRF (X-ray 222 fluorescence) and ICP-MS (inductively coupled plasma–mass spectrometry) at the 223 Peter Hooper GeoAnalytical Laboratory Washington State University (Johnson et al., 224 1999; Knaack et al., 1994).

225 RESULTS

226 Detrital zircon geochronology

227 New detrital zircon geochronological data were collected from nine samples from the 228 westernmost exposures of the Mesoproterozoic Belt-Purcell Supergroup and from the 229 structurally overlying Mesoproterozoic Deer Trail Group and Neoproterozoic 230 Windermere Supergroup in northeastern Washington. Data are presented as probability 231 distribution plots (Gehrels, 2010) in Fig. 5. Age population peaks are tabulated for each 232 of the samples shown in Fig. 5 in Table 1. Statistical data (2-sigma uncertainty, MSWD)

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233 for the youngest age population peak (Maximum Depositional Age, or MDA) from our 234 new samples are also given in Table 1. Complete data tables for each of our samples 235 are given in Supplemental File 1. Sample results are presented below by the three 236 stratigraphic groups, from oldest to youngest.

237 Belt-Purcell Supergroup

238 Samples from the western Belt-Purcell Supergroup (as mapped by Miller, 2001) were 239 collected from the Revett Formation of the Ravalli Group on the north shore of Deer 240 Lake (07SB306A) and from the Snowslip Formation (07SB203A) and the argillite of Half 241 Moon Lake (07SB305A), both mapped as Missoula Group, on the east slope of 242 Quartzite Mountain east of the town of Chewelah (Fig. 2). The Revett Formation 243 sample consists of medium-grained quartzitic sandstone from a thick-bedded sequence 244 (Fig. 4A). The sample has an MDA of 1579 ± 18 Ma (Table 1) from 46 zircon grains 245 (referred to as ‘n=46’ in subsequent Draftdiscussion). The sample has two older statistical 246 age population peaks (Table 1; Fig. 5) at ca. 1723 Ma (n=20) and ca. 1828 Ma (n=4). 247 The data was further analyzed with the Unmix program within Isoplot (Ludwig, 2008) 248 and an additional small peak of 1475 ± 16 Ma was detected (Unmix output shown in 249 Supplemental File 1) This suggests a younger MDA could be indicated for this sample.

250 The stratigraphically higher Snowslip Formation sample is a fine quartzitic sandstone in 251 a thin-bedded sequence of sandstone, siltstone and black argillite (Fig. 4B). Although 252 the sample yielded only 28 zircon grains (Table 1; Fig. 5), it records prominent age 253 population peaks of ca. 1578 Ma (n=13), ca. 1653 Ma (n=10), and a younger age 254 population peak (MDA) of 1391 ± 25 Ma (n=6). The Unmix program yields a similar 255 MDA (Supplemental File 1).

256 The stratigraphically highest unit sampled below the Cambrian quartzite east of 257 Chewelah is the argillite of Half Moon Lake (alternately interpreted as Bonner Quartzite: 258 Reed Lewis and Russ Burmester, written comm., 2013). This sample was collected 259 from a ledgy sequence of fine to medium-grained quartzitic on the east side 260 of Quartzite Mountain (Fig. 4C). It has an MDA of 1376 ± 27 Ma (n=5; Table 1) and a 261 more prominent age peak at about ca. 1666 Ma (n= 46; Table 1; Fig. 5). It differs from

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262 the two other mapped Belt-Purcell Supergroup samples discussed above in its younger 263 MDA and lack of any zircon grains with ages between ca. 1500 and 1600 Ma.

264 Deer Trail Group

265 Samples from the Deer Trail Group (as mapped by Miller, 2001) were collected from 266 road cuts along the road leading from Waitts Lake to the summit of Stensgar Mountain 267 (Fig. 2) and include the Togo Formation (Togo3), the lower (07SB302B) and upper 268 (07SB303A) Wabash-Detroit Formation, and the Buffalo Hump Formation (BH1). The 269 Togo Formation sample is from finely layered, fine sandstone, siltstone, and black slate 270 (Fig. 4D). It has an MDA of 1362 ± 35 Ma (Table 1) with n=4. The Togo Formation 271 sample also contains a large ca. 1726 Ma age population peak (n=56; Table 1; Fig. 5). 272 A shallowing upward sequence overlies the Togo Formation, culminating in the 273 stromatolite-bearing Wabash-Detroit Formation with a lower section of thin-bedded 274 quartzite, siltite and argillite (Fig. 4E)Draft with thick-bedded dolomite increasingly 275 interbedded in the upper section (Fig. 4F). The lower and upper Wabash-Detroit 276 Formation yield MDAs of 1300 ± 24 Ma (n=5) and 1359 ± 38 Ma (n=3), respectively 277 (Table 1). The lower Wabash-Detroit Formation also has older detrital zircon age 278 population peaks (Table 1; Fig. 5) of ca. 1403 (n=15), ca. 1684 Ma (n=19), ca. 1743 Ma 279 (n=32), ca. 2412 (n=3), and ca. 2629 Ma (n=5). The upper Wabash-Detroit Formation 280 (with only 17 grains used in analysis) has similar older detrital zircon age population 281 peaks of ca. 1670 Ma (n=4) and ca. 1742 Ma (n=4), and a different older age population 282 peak of ca. 1810 Ma (n=3). No sandstone samples were collected for detrital zircon 283 analyses from strata of the second deepening (McHale Slate) and shallowing (Stensgar 284 Dolomite) depositional cycle in the upper portion of the Deer Trail Group.

285 The Buffalo Hump Formation has traditionally been interpreted to be the highest unit in 286 the Deer Trail Group (Campbell and Loofbourow, 1962; Evans, 1987; Miller, 2001). The 287 Buffalo Hump Formation sample comes from thick-bedded, medium-grained quartzitic 288 sandstone (Fig. 4G). This sample has an MDA of 1112 ± 20 Ma (n=30; Table 1). Older 289 zircon age population peaks (Table 1; Fig. 5) from the Buffalo Hump formation are ca. 290 1427 Ma (n=29), ca. 1820 Ma (n=14), ca. 2452 Ma (n=4), and ca. 2695 (n=3). Our 291 more robust data set from the Buffalo Hump generally corroborates the limited data

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292 reported by Ross et al. (1992) but indicate a much younger depositional age than would 293 be expected if it were part of the Deer Trail Group.

294 Windermere Supergroup

295 Samples from the Windermere Supergroup (as mapped by Miller, 2001) were collected 296 from road cuts along the road leading from Waitts Lake to the summit of Stensgar 297 Mountain (Fig. 2). The lowest formation in the Windermere Supergroup is the 298 conglomeratic lower Huckleberry Formation (Fig. 4H). A sample of the lower 299 Huckleberry Formation (HB_1) has an MDA of 1143 ± 19 Ma (n=19). Older zircon age 300 population peaks (Table 1; Fig. 5) from the lower Huckleberry Formation are ca. 1434 301 Ma (n=32), ca. 1740 Ma (n=12), ca. 1851 Ma (n=10), and ca. 1972 Ma (n=3).

302 The highest unit in the Windermere Supergroup is the conglomeratic Monk Formation. 303 A matrix-supported metaconglomerate sample (fig 4i) of the Monk Formation (DT3), 304 proximal to the Lane Mountain thrustDraft fault (Fig. 2), has an MDA of 1252 ± 16 Ma. That 305 sample has two other zircon age population peaks (Table 1; Fig. 5) at ca. 1728 Ma 306 (n=21) and ca. 2653 Ma (n=4).

307 Whole rock geochemistry of mafic igneous rocks

308 Three mafic igneous samples were collected from the area west of Chewelah (Fig. 2). 309 A fine-grained basaltic (diabase) dike at least 5 m wide intruding the Togo Formation 310 was sampled about 11 km west-southwest of Waitts Lake. A 3-m-wide, fine mafic dike 311 (RM100) intruding the Stensgar Dolomite was sampled in the abandoned Red Marble 312 magnesite quarry along the Stensgar Mountain road. A plagioclase-phyric, 313 glomeroporphyritic, sparsely amygdaloidal fine basaltic flow from the middle of the 314 upper Huckleberry Formation was sampled near the top of the Stensgar Mountain road. 315 Geochemical data from these samples is given in Table 2. Major elements, determined 316 by XRF, are given as anhydrous weight percent values, while minor and trace elements 317 are given in ppm with no anhydrous correction.

318 All three samples are high titanium (3-5% TiO2) basalts with SiO2 contents between 319 48.7-50.2% and MgO between 6.0-7.6%. On chondrite-normalized, extended-rare-

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320 earth-element spidergrams (Fig. 6a), the upper Huckleberry lava and the dike cutting 321 the Togo Formation have very similar signatures with a relatively smoothly increasing 322 chondrite-normalized trace element contents from Yb (ca. 10 x chondritic value) to Ta 323 (ca. 70 x chondritic value). That smooth increase is broken by a sharply decreased 324 value for Sr. The Huckleberry lava sample decreases gradually to the left in Fig. 6a 325 from Ta to Ba (ca. 25 x chondritic value). The Togo dike sample also generally 326 decreases to the left of Ta in Fig. 6a, but with a more irregular pattern. The Stensgar 327 dike sample has a roughly parallel spidergram to that of the other two samples, but with 328 the values ca. 1.7 times higher.

329 DISCUSSION

330 Stratigraphic position and provenance based on detrital zircon age populations

331 The detrital zircon age populations from the stratigraphically constrained samples of 332 Belt-Purcell, Deer Trail and WindermereDraft units yield information on their MDAs and on 333 ages of materials in their source areas. These characteristics can also be used to 334 compare these units in northeastern Washington with potentially correlative units 335 elsewhere in western North America.

336 Belt-Purcell Supergroup

337 The detrital zircon age population of the Revett Formation of the Ravalli Group near 338 Chewelah is quite similar to those of published samples of the Ravalli Group (Ross and 339 Villeneuve, 2003), as well as from the Prichard Formation (Ross and Villeneuve, 2003; 340 Link et al., 2007; Lewis et al., 2010) and from the Piegan Group (Ross and Villeneuve, 341 2003; Lewis et al., 2007) in northern Idaho and western Montana. For comparison, 342 published data from a sample of Prichard Formation Member C from about 125 km east 343 of Chewelah (Lewis et al., 2010) is shown with our samples in Fig. 5 and in Table 1. All 344 these samples show the major age peak ca.1575 Ma, with subsidiary peaks between 345 1730 and 1830 Ga. Ross et al. (1992) were the first authors to recognize the 1600- 346 1500 Ma detrital zircon age population as falling within a gap in North American 347 magmatic ages (NAMG) and to suggest an exotic (i.e., non-Laurentian) outboard craton 348 to the west was the likely source for the Lower Belt (Prichard Formation), Ravalli and

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349 Piegan Groups of the western Belt-Purcell basin. Most authors (Ross and Villeneuve, 350 2003; Fanning and Link, 2003; Belousova et al., 2009; Lewis et al., 2010; Doe et al. 351 2013; Jones et al., 2015) favor the Gawler craton of southeastern Australia, as well as 352 the broader Mawson Continent and adjacent Northern Australia (Fig. 7a) as the most 353 likely source terrain of these NAMG zircons, suggesting that the Mesoproterozoic 354 Mawson Continent-Northern Australia was attached to the west side of Laurentia during 355 the deposition of the Lower Belt, Ravalli and Piegan Groups (Fig. 7a).

356 The stratigraphically higher sample mapped as the basal unit of the Missoula Group 357 (Snowslip Formation) overlying Piegan Group carbonates also records a prominent 358 1578 Ga peak of apparent western source, with an MDA of 1391 ± 25 Ma (Table 1). 359 This MDA is similar in age to, but slightly younger than, a tuff dated at 1401 ± 6 Ma in 360 the easterly sourced Libby Formation near the top of the Missoula Group ca. 180 km to 361 the east (Evans et al., 2000). Our Snowslip MDA is considerably younger than the 362 zircon U-Pb age (ca. 1443 Ma) reportedDraft for the Purcell Lava member of the Snowslip 363 Formation 180 km to the east in western Montana (Evans et al., 2000). However, that 364 older age was challenged by more recent work of Constenius et al. (2017), who 365 reported a new zircon U-Pb age for the Purcell Lava of 1386 +/-10 Ma, like the age 366 reported here. Constenius et al. (2017) suggested that the older age reported by Evans 367 et al. (2000) reflects a detrital zircon age peak from the peperite lavas misinterpreted as 368 rhyolitic lavas. Link et al. (2016) reported a similar MDA (ca. 1390 Ga) from a Lemhi 369 Group sample (Missoula Group-correlative). These overlapping ages (ca. 1405-1385 370 Ma) from basal to upper Missoula Group units may indicate that Missoula Group 371 deposition occurred during that shorter younger interval, rather than during the 1443- 372 1400 Ma interval suggested by Evans et al. (2000).

373 Elsewhere in the Belt-Purcell basin, the Missoula Group records a change in 374 provenance from the non-Laurentian western source, dominated by NAMG-age zircon 375 grains, to one or more Laurentian sources (Ross et al., 1992; Ross and Villeneuve, 376 2003). Our Snowslip Formation sample represents the youngest sediment with NAMG- 377 age non-Laurentian sources in the Belt-Purcell basin and, indeed, is the youngest 378 documented strata known within the entire Belt-Purcell basin to date. The Lemhi Group

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379 sample (Missoula Group-correlative) in central Idaho with an MDA of ca. 1390 Ma 380 reported by Link et al. (2016) also lacks any NAMG-age grains. If our Snowslip 381 Formation sample is correlative with the Lemhi Group, then the contrasting Laurentian 382 vs. non-Laurentian (NAMG-bearing) sources on opposing flanks of the Belt-Purcell 383 basin, documented for the lower three Belt-Purcell Groups by Ross and Villeneuve 384 (2003), continued at least through 1390 Ga.

385 In contrast to our Snowslip sample with distinct non-Laurentian NAMG-age detrital 386 zircon, other Missoula Group and correlative strata elsewhere in the Belt-Purcell basin 387 are derived primarily from juvenile 1.8-1.6 Ga source regions such as the Mazatzal- 388 Yavapai belts of southwestern Laurentia (Ross and Villeneuve, 2003; Lewis et al., 2010; 389 Link et al., 2016). A portion of the Missoula Group ca. 1.8–1.6 Ga detrital zircons have 390 epsilon Hf isotopic values that are too low to be compatible with the known isotopic 391 compositions of the Proterozoic accretionary provinces of the southwestern U.S. 392 (Stewart et al., 2010). The adjacent DraftMojave province (Fig. 1) could have supplied 393 somewhat more evolved (lower epsilon Hf) detrital zircons in that age range, but it is 394 uncertain whether it could have provided the lowest epsilon Hf detrital zircons in 395 samples of the Missoula Group and correlative strata. Stewart et al. (2010) suggest that 396 these most isotopically evolved zircons were more likely derived from sources in 397 eastern Antarctica that are suitable in age and have similar isotopic values (Goodge et 398 al., 2008). If true, this interpretation would require source regions in Antarctica to be 399 proximal with southwestern Laurentia until at least ca. 1.39 Ga and would imply that the 400 supercontinent Columbia was still somewhat intact on its western side (Zhao et al., 401 2004). Our data indicating a previously unrecognized non-Laurentian NAMG 402 contribution in ca. 1.39 Ga strata in Washington state provide additional support for that 403 model (Fig. 7b).

404 Deer Trail Group

405 The Togo Formation is the lowest unit of the Deer Trail Group. Although our reported 406 uncertainty is large, the MDA of the Togo sample (1362 + 35 Ma) is apparently younger 407 than any known Belt-Purcell strata and is similar in age to plutons cutting deformed Belt- 408 Purcell strata in southern Canada (Hellroaring stock, 1364 + 6 Ma: McFarlane, 2015)

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409 and in central Idaho (Big Deer Creek stock, 1377 + 4 Ma; Aleinikoff et al., 2012). This 410 suggests that the Togo Formation is younger than the deformational end of Belt-Purcell 411 sedimentation. The similarity of the MDA to the ages of the crosscutting plutons 412 suggests the youngest grains in the Togo Formation may be derived from the 413 magmatism that marks the end of Belt-Purcell sedimentation. An association of 414 widespread magmatism with black shale deposition at ca. 1380 Ma has also been noted 415 in other areas of western Laurentia (Zhang et al., 2018). The larger late 416 Paleoproterozoic age population peak in the Togo Formation of 1726 Ma (n=56) 417 appears to be primarily derived from the juvenile Mazatzal-Yavapai source in the 418 southwestern US. Preliminary detrital zircon Hf isotopic data (Brennan et al., 2019) 419 shows that some zircons of this age from the Togo Formation have more negative Hf 420 isotopic values than the Mazatzal-Yavapai province and seemingly require continued 421 contribution from a more evolved source like eastern Antarctica (Fig. 7c; discussed 422 further below). Draft 423 The highest sample in the nominal Belt-Purcell section east of Chewelah (argillite of 424 Half Moon Lake) has a similar MDA (1376 + 27 MY: Table 1) to the Togo sample. It 425 also has a large detrital zircon late Paleoproterozoic age population (1666 Ma; n=46) 426 appropriate for Mazatzal-Yavapai and possibly eastern Antarctica sources from the 427 south. Although our reported age uncertainties are large, the similarities of the MDAs of 428 the argillite of Half Moon Lake exposures east of Chewelah with the Togo Formation 429 farther west suggest these units are correlative and are post-Belt-Purcell Supergroup 430 strata. The argillite of Half Moon Lake is mapped as conformably overlying Missoula 431 Group strata (Shepard Formation) on its east side (Fig. 2), although it is mapped as 432 mostly faulted against the Missoula Group farther south. The argillite of Half Moon Lake 433 east of Chewelah is about 650 m thick beneath unconformably overlying Cambrian 434 Addy Quartzite (Miller, 2001).

435 West of Chewelah and of the Jumpoff Joe fault, a shallowing upward sequence overlies 436 the Togo Formation, culminating in the stromatolite-bearing Wabash-Detroit Formation. 437 The MDA of the lower Wabash-Detroit Formation (1300 + 24 Ma), represents 438 Mesoproterozoic sedimentation clearly younger than Belt-Purcell Supergroup

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439 deposition. A low-zircon-yield sample of the upper Wabash-Detroit Formation has an 440 MDA of 1359 + 38 Ma. Circa 1610-1490 Ma non-Laurentian grains were not recognized 441 in the Wabash-Detroit samples, suggesting a southern (southwestern US: Mazatzal- 442 Yavapai +/- eastern Antarctica) rather than a western sediment source. Older 443 Paleoproterozoic and late Archean grains in these samples likely originated from the 444 Laurentian craton to the east.

445 The distinctly younger MDA of the Buffalo Hump Formation (1112 + 20 Ma) suggests 446 that it is considerably younger (ca. 200 My) than other formations considered part of the 447 Deer Trail Group. Several lines of evidence suggest to us that the Buffalo Hump may be 448 the basal formation of the Windermere Supergroup, rather than the highest unit in the 449 Deer Trail Group, as suggested earlier by Evans (1987). Its detrital zircon age 450 signature, including its MDA, is like that of the overlying Huckleberry Formation of the 451 Windermere Supergroup and distinct from the underlying Deer Trail samples in its 452 significant Grenville age MDA. Its quartzose,Draft non-calcareous composition is similar to 453 the overlying quartzose Huckleberry Formation conglomerate and contrasts with the 454 generally dolomitic and argillitic character of the Deer Trail Group and indicates the 455 onset of a change in depositional framework. Its rapid lateral thickness and facies 456 changes between argillite and quartzose sandstone are more characteristic of 457 Windermere Supergroup strata than of either Deer Trail Group or Belt-Purcell 458 Supergroup strata (Campbell and Loofbourow, 1962; Miller et al., 1973; Miller and 459 Whipple, 1989; Miller, 2001). And, finally, the Buffalo Hump in the study area has been 460 interpreted to sit unconformably on the underlying Stensgar Dolomite of the Deer Trail 461 Group (Evans, 1987). However, since the ca. 1110 Ma MDA of the Buffalo Hump 462 sample would still allow the Buffalo Hump Formation to have a late Mesoproterozoic 463 and/or early Neoproterozoic depositional age, and considering the Neoproterozoic 464 depositional age of the Windermere Supergroup, we simply separate it from both the 465 underlying Deer Trail Group and the overlying Windermere Supergroup in figures 2, 3 466 and 5.

467 Strata coeval with the Deer Trail Group occur in the southwestern U.S. and 468 northwestern Canada (Figs. 1, 7). Basal strata in the Apache Group in southeastern

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469 Arizona (Fig. 1) contain tuffaceous layers from a syn-sedimentary volcanic source that 470 yield zircon U-Pb ages of 1328 + 5 Ma (Stewart et al., 2001) and 1338 Ma (Mulder et 471 al., 2017). The basal unit of the lower Pahrump Group in southeastern California 472 (Mulder et al., 2017) has a similar youngest detrital zircon peak (ca. 1323 Ma), while 473 higher strata in both the Pahrump, Apache and Unkar Group sections have MDAs 474 younging progressively upward from ca. 1230 Ma to as young as 1100 Ma. The 475 Pinguicula Group in northwestern Canada (Fig. 1) overlies the ca. 1380 Ma Hart River 476 volcanic complex and is clearly older than 1000 Ma (Thorkelson et al., 2005; Medig et 477 al., 2014, 2016) and possibly older than 1270 Ma (Medig et al., 2012).

478 Windermere Supergroup

479 The Windermere Supergroup in northeastern Washington is known to be 480 Neoproterozoic based on the age of 762 + 44 Ma of a basaltic flow in the upper 481 Huckleberry Formation (Devlin and Bond,Draft 1988). The gradational contact between the 482 upper Huckleberry Formation and conglomerates of the lower Huckleberry Formation 483 indicates their rough contemporaneity. Lower Huckleberry Formation conglomerate has 484 an MDA of 1143 + 19 Ma that suggests derivation from the Grenville province of eastern 485 and southern Laurentia (Fig. 1). Other detrital zircon age population peaks (ca. 1434, 486 1740 Ma) are characteristic of the A-type and Mazatzal-Yavapai belt of 487 southern Laurentia. Older detrital zircon age peaks (ca. 1851, 1972 Ma) are more 488 compatible with sources in the Trans-Hudson orogen and other Paleoproterozoic 489 terranes exposed north and east of northeastern Washington (Fig. 1: Fort Simpson, 490 Great Bear and Thelon magmatic arcs). As mentioned in the previous section, the 491 Buffalo Hump Formation has a similar detrital zircon signature to the lower Huckleberry 492 Formation, sharing a similar MDA (ca. 1125) and two other prominent detrital zircon age 493 peaks (ca. 1430 and 1835 Ma).

494 The conglomeratic Monk Formation comprises the highest unit of the Windermere 495 Supergroup in northeastern Washington. Its MDA of 1252 + 16 Ma and 1728 Ma age 496 population peak could reflect early Grenville arc magmatism in the Llano area of Texas 497 (Mosher, 1998) and Mazatzal-Yavapai magmatism in Arizona-New Mexico, all in 498 southern Laurentia (Fig. 1). Its lithologic character, including common pebbly mudstone

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499 (Fig. 4i), has led to the suggestion that it was deposited during the Neoproterozoic 500 Cryogenian glacial epoch (Link et al., 1994).

501 The prominent 1175 + 75 Ma detrital zircon age peak seen in the Windermere 502 Supergroup of northeastern Washington is common in all Neoproterozoic sedimentary 503 sequences from southeastern California and northern Arizona through Utah and Idaho 504 (Yonkee et al., 2014). A prominent Grenville age peak is also seen in correlative strata 505 in northern British Columbia (Eyster et al., 2018) and in the Mackenzie Mountains of 506 Northwest Territories (Rainbird et al., 2017). Pancontinental rivers (and glaciers?) 507 carrying sediment from the erosional unroofing of the Grenville orogen of eastern and 508 southern Laurentia across Laurentia to the Neoproterozoic rift basins of western 509 Laurentia have been proposed (Yonkee et al., 2014; Rainbird et al., 2017). Also, it has 510 been noted elsewhere (Moecher and Samson, 2006) that Grenville-age granitic 511 intrusions are exceptionally Zr-rich with high modal zircon and would contribute 512 disproportionately more zircon to detritalDraft zircon assemblages.

513 Finally, the Cambrian Addy Quartzite unconformably overlies all the Proterozoic 514 sequences, and published detrital zircon data (Linde et al., 2014, 2018) can be 515 compared with that from the Proterozoic sequences. The Addy Quartzite has an MDA 516 of ca. 1100 Ma, like the Grenvillean MDAs in the underlying Windermere Supergroup. 517 The other detrital zircon age population peaks in the Addy Quartzite (Table 1) could also 518 be a mix of Mazatzal-Yavapai sources (1750 + 50 Ma) and easterly derived cratonal 519 sources (1840 + 30 Ma; ca. 2455 Ma).

520 Huckleberry volcanism: Tectonic setting and apparent feeder dikes

521 The upper Huckleberry Formation volcanic rocks comprise a thick pile of basaltic lavas 522 that continue into northeasternmost Washington as the Leola volcanics and across the 523 border into Canada as the Irene volcanics (Miller et al., 1973). The best available 524 geochronology on these rocks was provided by Devlin and Bond (1988), who used the 525 Sm-Nd isochron method on clinopyroxene and whole rock to generate a best-fit age of 526 762 + 44 Ma. The upper Huckleberry lava and two underlying mafic dikes sampled here 527 (crosscutting the Stensgar dolomite and the Togo Formation) are geochemically similar

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528 to each other and have trace element chemistry (Fig. 6a) like Hawaiian and other 529 oceanic island basalt lavas (Harrison and Weis, 2018). On the Th-Ta-Hf diagram (Fig. 530 6b) of Wood et al. (1979), the Huckleberry lava and dike samples plot in the field of E- 531 MORB/Continental rift/Continental flood basalts. Epsilon Nd isotopic data (calculated 532 for 780 Ma) were reported from Huckleberry and Leola lavas north of the study area to 533 range from +5 to +6 (Mackinder et al., 2019), indicative of a depleted mantle source. 534 The trace element signature of the lavas, their source from the asthenospheric mantle, 535 and the temporal association with the rifting of western North America (Stewart, 1972; 536 Sandeman et al., 2014) support the setting of Huckleberry volcanism as continental rift- 537 related magmatism.

538 Mafic dikes and sills of similar age (mostly U-Pb ages of ca. 775 Ma) are documented in 539 exposures up to 500 km to the southeast of the Chewelah area across western 540 Montana into southeastern Idaho and northwestern Wyoming (Harlan et al, 1997), as 541 well as in northern British Columbia,Draft Yukon and Northwest Territories in northwestern 542 Canada (Park et al., 1995; Harlan et al., 2003; Mackinder et al., 2019). The northern 543 and southern intrusive clusters are referred to as the northern and southern Gunbarrel 544 intrusions, respectively (Mackinder et al., 2019). Although similar in age, these 545 intrusions are geochemically and isotopically distinct from the Huckleberry lavas. On 546 the Hf-Th-Ta diagram (Fig. 6b), Gunbarrel intrusive rocks (Mackinder et al., 2019) plot in 547 field D (subduction-related arc magmas or crustal contamination), in contrast to samples 548 of Huckleberry and associated units, which plot in field B and C (E-MORB/OIB/ 549 continental rift magmas). Epsilon Nd values for Gunbarrel rocks ca. 780 Ma are 550 considerably lower (ca. +2) than those of Huckleberry lavas (+5 to +6). Mackinder et al. 551 (2019) argue that these differences result from contamination of primary melts like 552 Huckleberry lavas by a significant volume (5-15%) of partial melts of upper continental 553 crust, while Sandeman et al. (2014) suggest primary Huckleberry-like magmas were 554 contaminated by lower continental crust to yield the Gunbarrel magmas.

555 More broadly, the widespread distribution of mafic magmatism in Rodinia from western 556 Laurentia to rifted blocks farther west between 860 and 570 Ma has been inferred to 557 reflect a giant mantle “superplume” beneath Rodinia that led to its breakup (Anderson,

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558 1982; Li et al., 2008). The divergent trends of the Gunbarrel dike swarms radiate from a 559 center presently located in the Pacific Ocean west of Vancouver Island, a location that 560 would be within the crustal block that was subsequently rifted away from Laurentia 561 (Park et al., 1995). Park et al. (1995) noted mafic dikes of correlative age are presently 562 located in southeastern Australia, while Li et al. (1999) cited similar age dikes in south 563 China. The mafic dikes in southeastern Australia have trace element signatures more 564 like the Huckleberry-Irene Formation lavas (transitional N-MORB/E-MORB: Park et al., 565 1995; Mackinder et al., 2019) and would seem to be a better fit for correlative, 566 uncontaminated, plume-related magmas.

567 Implications for Late Proterozoic normal fault motion on the Jumpoff Joe Fault

568 The thickness of preserved Deer Trail Group beneath the unconformably overlying 569 Cambrian Addy Quartzite is dramatically different across the Jumpoff Joe fault near 570 Chewelah (Fig. 3). According to MillerDraft (2001), the argillite of Half Moon Lake (which we 571 correlate here with the Togo Formation) underlies the Cambrian Addy Quartzite east of 572 the Jumpoff Joe fault and is 650 m thick, while the Deer Trail Group west of the Jumpoff 573 Joe fault is estimated to be over 3500 m thick. An additional 2300 m of Windermere 574 Supergroup strata overlie the Deer Trail Group and underlie the Cambrian quartzite 575 west of the Jumpoff Joe fault and are absent east of it. The presence of the ca 6-km- 576 thick section of Deer Trail Group and Windermere Supergroup on the west side 577 (hanging wall) of the Mesozoic, west-dipping Jumpoff Joe thrust fault and the presence 578 of only the lower 0.6 km of the Deer Trail Group east of the fault suggests that a sub- 579 Cambrian, down-to-the-west normal fault preserved these strata west of the fault and 580 that these strata were uplifted and removed by sub-Cambrian erosion east of the fault. 581 The Mesozoic east-vergent Jumpoff Joe thrust fault marks the site of this abrupt 582 change, and so we suggest that the Jumpoff Joe fault, the frontal thrust of the Mesozoic 583 Colville River fold and thrust belt, was reactivated along an earlier late Neoproterozoic, 584 down-to-the-west normal fault. Presumably this normal faulting occurred during or after 585 Windermere time (ca. 760 Ma) and erosion of the uplifted footwall occurred during the 586 ca. 200 m.y. interval before the onlap by Cambrian marine sediments.

587 Two rifting events or one?

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588 Some authors (Winston et al., 1984; Link et al., 1993) have suggested that the 589 disappearance of the western non-Laurentian source and the end of Belt-Purcell basin 590 sedimentation in general was the result of a failed intracontinental rift and that the 591 western source block simply subsided and remained in place. In this model renewed 592 rifting 800 m.y. later in the Late Neoproterozoic led to complete separation of the 593 western cratonic block from Laurentia and to generation of an intervening ocean basin 594 (Ross et al., 1992; Goodge et al., 2008; Lewis et al., 2010). Other authors (Harrison et 595 al., 1974; Gardner, 2008; Jones et al., 2015) suggest that the continental rifting implied 596 by Belt-Purcell basin sedimentation and magmatism did lead to mid-Mesoproterozoic rift 597 separation of the western source block (Australia), reflected by the demise of the 598 western sediment source. Do our new data from the Chewelah area bear on this issue 599 of one or two Proterozoic rifting events in western Laurentia?

600 Linking southeastern Australia as the now-removed western sediment source to the 601 Mesoproterozoic Belt-Purcell basin (RossDraft et al., 1992; Ross and Villeneuve, 2003) has 602 been extended to other correlative Mesoproterozoic sedimentary basins as far south as 603 the southwestern U.S. and as far north as the Yukon province in northwestern Canada 604 (Figs. 1, 7a). The 1.48-1.43 Ga successions in the Yankee Joe and Trampas basins in 605 Arizona-New Mexico (Fig. 1) also contain abundant detrital zircons between 1.6-1.5 Ga 606 which are correlative in age and Hf isotopic composition with rocks in the Mawson 607 Continent of east Antarctica and southern Australia and in the Northern Australia craton 608 (Fig. 7a; Doe et al., 2013; Jones et al., 2015). The PR1 basin in northern Yukon (Fig. 609 1), inferred to record deposition between 1.46-1.42 Ga, records a unimodal detrital 610 zircon age peak of ca. 1.5 Ga (Medig et al., 2014). These authors propose the Mt. Isa 611 region of the North Australia craton (Fig. 7a) as a unique source of these ca. 1.5 Ga 612 zircons. Together with the abundant input of 1.6-1.5 Ga detrital zircons to 1.48-1.44 Ga 613 units of the Belt-Purcell basin, these sources in southern and eastern Australia place 614 strong constraints on its position against Laurentia at 1.45 Ga (Fig. 7a).

615 New data presented here show that ca. 1.58 Ga sources to the west continued to feed 616 the westernmost Belt-Purcell basin until after 1.39 (+/- 0.025) Ga, while 1.43(?)-1.39 Ga 617 Missoula-Lemhi Group throughout the rest of the Belt-Purcell basin received sediment

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618 from a mix of Laurentian and southern Mawson Continent sources (Fig. 7b). During 619 Missoula-Lemhi Group deposition, regional orogenesis in the southwestern US (Picuris 620 : Daniel et al., 2013; Jones et al., 2015) resulted in deformation, 621 and uplift within the Yavapai-Mazatzal province, contributing to the flood of sediment 622 from this area into the Belt-Purcell basin. In the PR1 basin in northern Yukon, extensive 623 mafic sills were intruded (1.385-1.380 Ga) with juvenile isotopes and MORB-like 624 chemistry, which suggest the onset of continental rifting (Verbaas et al., 2018). 625 Although rifting of Australia-Antarctica from Laurentia may have begun in the north at 626 1.39-1.38 Ga, sediment sources to the Belt-Purcell basin require that at least the 627 Mawson craton still remained attached to Laurentia at this time (Fig. 7b).

628 Evidence for 1.37 Ga deformation in the area of the Belt-Purcell basin is scattered 629 around northern and central Idaho (Doughty and Chamberlain, 1996; Zirakparvar et al., 630 2010) and southern British Columbia (Leech, 1962; McMechan and Price, 1982; 631 McFarlane, 2015). In central Idaho,Draft Doughty and Chamberlain (1996) presented 632 evidence that metamorphic pressure increased during growth of garnet associated with 633 ca. 1.37 Ga bi-modal intrusive rocks and that this increase in pressure resulted from 634 about 6 km of sedimentation on the overlying crust during protracted metamorphism. If 635 Deer Trail Group sedimentation, with an onset after 1.37 Ga, was initially more 636 widespread than preserved today, its deposition could have contributed to that overlying 637 crustal load. McFarlane (2015) suggests an extensional core complex detachment 638 environment for metamorphic-deformational events in the interval 1.37-1.34 Ga within 639 the central Belt-Purcell basin in southern British Columbia, Canada. Extensional 640 tectonism may characterize the entire region between the two areas during this time 641 interval. However, further research into the kinematic environment of the 1.37-1.34 Ga 642 metamorphism at these and other dated sites is clearly needed.

643 Deer Trail Group sedimentation commenced after ca. 1.37 Ga and records a 644 provenance from similar southern sources to those of the Missoula Group. The four 645 Deer Trail Group samples analyzed for detrital zircon geochronology in this paper are 646 each dominated by 1.75-1.65 Ga zircon populations that could appropriately have 647 sources in the Mazatzal-Yavapai province of the southwest U.S. However, as

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648 discussed above, similar grain populations in Missoula and Lemhi Group samples from 649 the upper Belt-Purcell Supergroup were sourced not only from the juvenile Mazatzal- 650 Yavapai province but also from the older crustal province of eastern Antarctica (Goodge 651 et al., 2008; Stewart et al., 2010), which was still attached to the west side of Laurentia 652 as part of supercontinent Columbia (Zhao et al., 2004). Hf isotopic characterization of 653 detrital zircons from the Deer Trail Group strata would test whether their 1.75-1.65 Ga 654 zircon populations could have been derived entirely from the positive epsilon Hf zircons 655 of the Mazatzal-Yavapai province (supporting completed rifting of the western source) 656 or whether the western source of zircons with evolved Hf isotopic signature was still 657 present (indicating the southern Mawson craton of eastern Antarctica had not rifted 658 away). Preliminary epsilon Hf isotopic data reported by Brennan et al. (2019) from 659 ca.1.8-1.6 Ga detrital zircons in the Togo Formation of the Deer Trail Group range from 660 +7 to -9, similar to the range reported by Stewart et al. (2010) for coeval zircons from 661 the Missoula and Lemhi Groups of theDraft Belt-Purcell Supergroup. Although the detrital 662 zircons with positive epsilon Hf values could have been derived from the Mazatzal- 663 Yavapai province (Goodge and Vervoort, 2006), a different source is required for detrital 664 zircons with highly negative Hf isotopic values. Assuming the grains were not simply 665 recycled from eroded Missoula-Lemhi Group strata, then the southern Mawson 666 Continent is still the mostly likely source and, thus, must have been adjacent to 667 southwestern Laurentia at the onset of Deer Trail sedimentation. The change from a 668 northwestern (NAMG-bearing) detrital zircon source after 1.39 Ga to a southern source 669 after 1.37 Ga could be accommodated by counter-clockwise rotational rifting from 670 western Laurentia wherein Antarctica remains adjacent to southwestern Laurentia while 671 Australia is rifted away (Fig. 7c). Zircon Hf isotopic data from younger units in the Deer 672 Trail Group and Windermere Supergroup would test whether the eastern Antarctica 673 source contribution continues to be required, or it is lost. Paleomagnetic data 674 (summarized by Pisarevsky et al., 2014) supports the alignment of Australia against 675 western Laurentia from ca. 1.8 Ga to at least 1.58 Ga but requires significant separation 676 between the two cratons by 1.2 Ga.

677 Considerable evidence exists in the Belt-Purcell basin area for a second deformation 678 and metamorphic mineral growth episode with Grenvillean ages of ca. 1.1 Ga

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679 (Anderson and Parrish, 2000; Doughty and Chamberlain, 2008; Zirakparvar et al., 2010; 680 Nesheim et al, 2012; McFarlane et al., 2015). The kinematics of this deformation are 681 mostly uncertain. However, in northern Idaho, a well-documented syn-metamorphic 682 fabric in metamorphosed Belt-Purcell strata associated with garnet growth has been 683 dated at ca. 1.08 Ga (Nesheim et al., 2012). The structural fabric is axial planar to 684 northerly fold axes produced during east-west contractional deformation and regional 685 kyanite grade metamorphism (Lang and Rice, 1985). This deformation could be a mid- 686 crustal response to collision involving western Laurentia, or alternatively could reflect an 687 episode of transpressional tectonism within Laurentia. Thorkelson et al. (2017) suggest 688 the return of Australia-Antarctica block during the , perhaps in a 689 somewhat different alignment leading to Rodinia assembly. Other models invoke other 690 crustal blocks (e.g. south China: Li et al., 2008; Xu et al., 2019; Tarim craton: Wen et 691 al., 2018) being wedged between Australia and Laurentia during 1.1-1.0 Ga 692 construction of the Rodinia supercontinent.Draft 693 Late Neoproterozoic normal faulting and continental rift-related mafic magmatism are 694 associated in space and time with rifting of a continental mass from the western edge of 695 Laurentia as part of the breakup of the supercontinent Rodinia (Stewart, 1972; Lund et 696 al., 2003; Li et al., 2008). The Laurentian rifted margin is approximately traced by the

697 Sri=0.706 line in figure 1 (Armstrong, 1988). Many authors (e.g. Ross et al., 1992; 698 Karlstrom et al., 2001; Goodge et al., 2008) suggest that Australia and Antarctica had 699 remained attached to western Laurentia since Belt-Purcell basin deposition and were 700 finally rifted away in late Neoproterozoic time. Others authors suggest that the 701 Australia-east Antarctica crustal block had been rifted or translated (Thorkelson et al., 702 2016) away at 1.37 Ga but that either it or some other cratonal block (e.g. South China: 703 Li et al., 2008b) converged on western Laurentia near 1.1 Ga during the Grenville 704 orogeny, only to be rifted away again in the Late Neoproterozoic (Thorkelson et al., 705 2017). Our new detrital zircon geochronological data do not resolve this issue.

706 Conclusions

707 New information developed in this study lead to new interpretations of the ages and 708 sources of detritus for Meso- and Neoproterozoic clastic units in northeastern

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709 Washington. New geochemical data from Neoproterozoic lavas yield information about 710 their sources and their tectonic setting. This new information can be interpreted in 711 terms of supercontinent evolution in Proterozoic time.

712 The westernmost Belt-Purcell Supergroup continued to be fed by the non-North 713 American source of ca. 1.58 Ga zircons, likely eastern Australia, until at least 1391+/-25 714 Ma, showing this western source could not have been rifted from Laurentia until after 715 this time. This MDA is similar to other Missoula Group and Lemhi Group strata several 716 hundred kilometers to the east and southeast, but these are considerably thicker (>4.5 717 km thick) and are derived primarily from the southwestern Laurentia and eastern 718 Antarctica, which was still attached to southwestern Laurentia as part of supercontinent 719 Columbia.

720 Based on its detrital zircon age population, the Deer Trail Group is younger than the 721 Belt-Purcell Supergroup. The basal DraftDeer Trail Group unit (Togo Formation) is younger 722 than 1362+/-35 Ma, while the middle Deer Trail Group unit (Wabash-Detroit Formation) 723 is younger than 1300+/-24 Ma. Deer Trail Group strata lack NAMG grains and have 724 1800-1600 Ma sources similar to the Mazatzal-Yavapai Groups in the southwestern 725 U.S. Preliminary detrital zircon Hf isotopic data from the Togo Formation (Brennan et 726 al., 2019) suggest a source in eastern Antarctica was also still contributing sediment at 727 the onset of Deer Trail Group deposition. Deer Trail Group strata conspicuously lack 728 Grenville-age grains that are abundant in overlying Neoproterozoic strata. We interpret 729 this to indicate the Deer Trail Group was deposited after rifting had removed the 730 western crustal block that had supplied ca. 1580 Ma zircons to the Belt-Purcell strata, 731 but before regional Grenville age tectonism that signaled the assembly of 732 supercontinent Rodinia.

733 The Windermere Supergroup in northeastern Washington was previously dated at ca. 734 762 Ma by the age of upper Huckleberry Formation lavas. The lavas grade down into 735 lower Huckleberry Formation conglomerates with sources like the Mazatzal-Yavapai 736 Groups and the Grenville orogenic belt in southwestern Laurentia. These 737 conglomerates grade downward into Buffalo Hump Formation sandstones that are 738 similar in composition and detrital zircon signature. We interpret the Buffalo Hump

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739 Formation, previously considered to belong to the underlying Mesoproterozoic Deer 740 Trail Group, to be a separate stratigraphic unit resting unconformably on the older Deer 741 Trail Group. Its overall lithologic and detrital zircon population suggest the possibility 742 that the Buffalo Hump represents the basal unit of the overlying Neoproterozoic 743 Windermere Supergroup. However, the possibility remains that the Buffalo Hump is late 744 Mesoproterozoic and/or early Neoproterozoic and intermediate in age between the Deer 745 Trail Group and the Windermere Supergroup. The Monk Formation conglomerates, 746 overlying the Huckleberry lavas and the highest unit of the Windermere Supergroup, 747 have a similar post-Grenville, southwestern U.S. provenance and are thought to reflect 748 glacial deposition.

749 Huckleberry Formation lavas are high Ti-basalts with trace elements like those of 750 Hawaiian lavas. Published isotopic data show that they were derived from a depleted 751 mantle source. They are interpreted as continental rift basalts. Their chemistry supports 752 earlier interpretations that they are partDraft of a broad family of plume-related melts 753 associated with the onset of rifting of the western Rodinian supercontinent.

754 Previous work has shown the Jumpoff Joe thrust fault to be the easternmost thrust of 755 the Colville River fold-and-thrust belt, otherwise referred to as the Kootenay arc. 756 Contrasting thickness of Proterozoic strata across the Jumpoff Joe fault below the 757 Cambrian unconformity indicates the west-dipping Mesozoic thrust fault is a reactivated 758 late Neoproterozoic normal fault with about 6 km of normal displacement.

759 Acknowledgments

760 The first author is grateful to Dwight Bradley for his early interest in this project, and the 761 US Geological Survey for its support. Discussions with Fred Miller (USGS, retired) and 762 Reed Lewis (Idaho Geological Survey) refined our understanding of the regional 763 geology. Discussions with Jack Nisbet on the mining and settlement history of the area 764 were enlightening. Katy Maynard is thanked for her help with field photography. Jamey 765 Jones (US Geological Survey), Reed Lewis (Idaho Geological Survey) and Paul Link 766 (Idaho State University) are thanked for their helpful reviews.

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1034 Mulder, J.A., Karlstrom, K.E., Fletcher, K., Heizler, M.T., Timmons, J.M., Crossey, L.J., 1035 Gehrels, G.E., and Pecha, M., 2017, The syn-orogenic sedimentary record of the 1036 Grenville Orogeny in southwest Laurentia. Precambrian Research, 294: 33–52, 1037 https://doi.org/10.1016/j.precamres.201.03.006.

1038 Mulder, J.A., et al., 2018, Rodinian devil in disguise: Correlation of 1.25–1.10 Ga strata 1039 between Tasmania and Grand Canyon. Geology, 46: 1–4. 1040 https://doi.org/10.1130/G45225.1

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1041 Nesheim et al, 2012; Mesoproterozoic syntectonic garnet within Belt Supergroup 1042 metamorphic tectonites: Evidence of Grenville-age metamorphism and deformation 1043 along northwest Laurentia. Lithos 134–135: 91–107. 1044 doi:10.1016/j.lithos.2011.12.008

1045 Park, J.K., Buchan, K.L., and Harlan, S.S. 1995. A proposed giant radiating dyke swarm 1046 fragmented by the separation of Laurentia and Australia based on paleomagnetism 1047 of ca 780 Ma mafic intrusions in western North America. Earth and Planetary 1048 Science Letters, 132: 129–139. doi:10.1016/0012-821X(95)00059-L.

1049 Pisarevsky, S.A., Elming, S.A., Pesonen, L.J., and Li, Z.-X., 2014, Mesoproterozoic 1050 paleogeography: Supercontinent and beyond. Precambrian Research, 244: 207– 1051 225. http://dx.doi.org/10.1016/j.precamres.2013.05.014

1052 Rainbird, R.H., Rayner, N.M., Hadlari, T., Heaman, L.M., Ielpi, A., Turner, E.C., and 1053 MacNaughton, R.B., 2017, ZirconDraft provenance data record the lateral extent of 1054 pancontinental, early Neoproterozoic rivers and erosional unroofing history of the 1055 Grenville orogen. Geological Society of America Bulletin, 129: 1408–1423. 1056 doi:10.1130/B31695.1

1057 Ross, G.M., Parrish, R.P., and Winston, D., 1992, Provenance and U-Pb geochronology 1058 of the Mesoproterozoic Belt Supergroup (northwestern United States): implications 1059 for age of deposition and pre-Panthalassa plate reconstructions. Earth and Planetary 1060 Science Letters, 113: 57-76.

1061 Ross, G.M., and Villeneuve, M., 2003, Provenance of the Mesoproterozoic (1.45 Ga) 1062 Belt basin (western North America): Another piece in the pre-Rodinia 1063 paleogeographic puzzle. Geological Society of America Bulletin, 115: 1191-1217.

1064 Sandeman, H.A., Ootes, L., Cousens, B., and Kilian, T. 2014. Petrogenesis of 1065 Gunbarrel magmatic rocks: homogeneous continental tholeiites associated with 1066 extension and rifting of Neoproterozoic Laurentia. Precambrian Research, 252: 166– 1067 179. doi:10.1016/j.precamres.2014.07.007.

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1068 Schandl, E.S., and Davis, D.W., 2000, Geochronology of the Sullivan deposit: U-Pb and 1069 Pb-Pb ages of zircons and titanites, Chapter 8 in Lydon, J.W., Höy, T., Slack, J.F., 1070 and Knapp, M., eds., The Geological Environment of the Sullivan Pb-Zn-Ag Deposit, 1071 British Columbia. Geological Association of Canada, Mineral Deposits Division, 1072 Special Publication 1: 127-135.

1073 Sears, J.W., Chamberlain, K.R., and Buckley, S.N., 1998, Structural and U-Pb 1074 geochronological evidence for 1.47 Ga rifting in the Belt Basin, western Montana. 1075 Canadian Journal of Earth Sciences, 35: 467-475.

1076 Stewart, E.D., Link, P.K., Fanning, C.M., Frost, C.D., and McCurry, M., 2010, 1077 Paleogeographic implications of non–North American sediment in the 1078 Mesoproterozoic upper Belt Supergroup and Lemhi Group, Idaho and Montana, 1079 USA. Geology, 38: 927–930, doi:10.1130/G31194.1.

1080 Stewart, J.H., 1972, Initial deposits inDraft the Cordilleran geosyncline: Evidence of a late 1081 Precambrian (<850 m.y.) continental separation. Geological Society of America 1082 Bulletin, 83, 1345-1360.

1083 Stewart, J.H., Gehrels, G.E., Barth, A.P., Link, P.K., Christie-Blick, N., Wrucke, C.T., 1084 2001, Detrital zircon provenance of Mesoproterozoic to Cambrian arenites in the 1085 western United States and northwestern Mexico. Geological Society of America 1086 Bulletin, 113: 1343-1356 doi:10.1130/0016- 1087 7606(2001)113<1343:DZPOMT>2.0.CO;2.

1088 Thompson, R.N., Morrison, M.A., Hendry, G.L., and Parry, S.J., 1984, An assessment 1089 of the relative roles of crust and mantle in magmagenesis: An elemental approach. 1090 Philosophical Transactions of the Royal Society London, Series A, 310: 675-692.

1091 Thorkelson, D., Furlanetto, F., Medig, K.P.R., Verbaas, J., and Rainbird, R., 2016, 1092 Australia-Laurentia relations in the late Paleoproterozoic to Mesoproterozoic: A 1093 Wilson cycle followed by sinistral translation. Geological Society of America 1094 Abstracts with Programs, 48: No. 7. doi: 10.1130/abs/2016AM-278590

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1095 Thorkelson, D., Verbaas, J, Medig, K.P.R., and Furlanetto, F., 2017, Proterozoic Wilson 1096 cyclicity and its application to Laurentia-Australia interactions. Geological Society of 1097 America Abstracts with Programs, 49, No. 5. doi: 10.1130/abs/2017RM-293108

1098 Verbaas, J., Thorkelson, D.J., Milidragovic, D., Crowley, J.L., Foster, D., Gibson, H.D., 1099 and Marshall, D.D., 2018, Rifting of western Laurentia at 1.38 Ga: The Hart River 1100 sills of Yukon, Canada. Lithos 316–317: 243–260. 1101 https://doi.org/10.1016/j.lithos.2018.06.018

1102 Watkinson, A.J., and Ellis, M.A., 1987, Recent structural analysis of the Kootenay arc in 1103 northeastern Washington, in Schuster, J.E., ed., Selected papers on the geology of 1104 Washington. Washington Division of Geology and Earth Resources Bulletin 77: 41- 1105 53.

1106 Whipple, J.W., and Johnson, S.N., 1988, Stratigraphy and Lithocorrelation of the 1107 Snowslip Formation (Middle ProterozoicDraft Belt Supergroup), Glacier National Park, 1108 Montana. U.S. Geological Survey Bulletin 1833: 30 p.

1109 Winston, D., Woods, M.O., and Byer, G.B., 1984, The case for an intracratonic Middle 1110 Proterozoic Belt-Purcell Basin: Tectonic, stratigraphic and stable isotopic 1111 considerations, in McBane, J.D., and Garrison, P.B., eds, Northwest Montana and 1112 adjacent Canada. Montana Geological Society 1984 Field Conference Guidebook, 1113 103-118.

1114 Winston, D., 2007, Revised stratigraphy and depositional history of the Helena and 1115 Wallace Formations, Mid-Proterozoic Piegan Group, Belt Supergroup, Montana and 1116 Idaho, in Link, PK. And Lewis, R.S., eds., Proterozoic Geology of western North 1117 American and Siberia. SEPM Special Publication 86: 65-190.

1118 Wood, D.A., Joron, J.L., and Treuil, M., 1979, A re-appraisal of the use of trace 1119 elements to classify and discriminate between magma series erupted in different 1120 tectonic settings. Earth and Planetary Science Letters, 45: 326-336.

1121 Xu, Y.J., Cawood, P.A., Zhang, H.C., Zi, J.W., Zhou, J.B., Li, L.X., and Du, Y.S., 2019, 1122 The Mesoproterozoic Baoban Complex, South China: A missing fragment of western

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1123 Laurentian lithosphere. Geological Society of America, (published online Nov. 5, 1124 2019). https://doi.org/10.1130/B35380.1

1125 Yonkee, W. A., Dehler, C. D., Link, P. K., Balgord, E. A., Keeley, J. A., Hayes, D. S., 1126 Wells, M. L., Fanning, C. M., and Johnston, S. M., 2014, Tectono-stratigraphic 1127 framework of Neoproterozoic to Cambrian strata, west-central US: Protracted rifting, 1128 glaciation, and evolution of the North American Cordilleran margin. Earth-Science 1129 Reviews, 136: 59–95, https://doi.org/10.1016/j.earscirev.2014.05.004

1130 Zhang, S.H., Ernst, R.E., Pei, J.L., Zhao, Y., Zhou, M.F., and Hu, G.H., 2018, A 1131 temporal and causal link between ca. 1380 Ma large igneous provinces and black 1132 shales: Implications for the Mesoproterozoic time scale and paleoenvironment. 1133 Geology, 46: 1–4, https://doi.org/10 .1130 /G45210 .1

1134 Zhao, G., Sun, M., Wilde, S.A., and Li, S., 2004, A Paleo-Mesoproterozoic 1135 supercontinent: Assembly, growthDraft and breakup. Earth-Science Reviews, 67: 91– 1136 123, doi:10.1016/j.earscirev.2004.02.003.

1137 Zirakparvar, N.A., Vervoort, J.D., McClelland, W., Lewis, R.S., 2010. Insights into the 1138 metamorphic evolution of the Belt-Purcell basin; evidence from Lu–Hf garnet 1139 geochronology. Canadian Journal of Earth Sciences 47: 161–179.

1140 Tables

1141 Table 1. Tabulation of statistically significant detrital zircon age populations (peak 1142 ages), calculated using the AGEPICK macro for Microsoft Excel (Gehrels, 2009).

1143 Table 2. Whole rock geochemistry of mafic igneous samples (latitude/longitude datum: 1144 WGS84).

1145

1146 Figures

1147 Fig. 1. Map of the Laurentian craton showing the distribution of Precambrian tectonic 1148 elements (adapted from Hoffman, 1988; Ross and Villeneuve, 2003; Jones et al., 2015).

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1149 Greenland is restored to its position prior to Cenozoic rifting from North America.

1150 Dashed line labelled “Sri=0.706” in western North America (Armstrong, 1988) indicates 1151 the approximate (buried) rifted western margin of the craton. The Belt-Purcell basin and 1152 other mid-Proterozoic basins are labelled (references in text). Areas of Neoproterozoic 1153 strata in western North America taken from Moynihan et al., 2019. Abbreviations: BH-- 1154 Buffalo Head, CB--Cumberland batholith, FS—Fort Simpson magmatic arc, GB—Great 1155 Bear magmatic arc, GF—Great Fall tectonic zone, H—Hottah terrane, W—Wopmay 1156 orogen. The area of Fig. 2 is indicated by small box on the west side of the Belt-Purcell 1157 basin.

1158 Fig. 2. Geologic map of Chewelah, WA area (adapted from Miller, 2001) showing 1159 locations of detrital zircon samples (red stars) and igneous geochemical samples 1160 (yellow stars). Line Z-Z’ gives the location of the generalized cross section shown in Fig. 1161 3. Locations mentioned in text are labelled.Draft 1162 Fig. 3. Generalized cross section from Z-Z’ (2.12 x vertical exaggeration). Unit colors 1163 and abbreviations are the same as in Fig. 2.

1164 Fig. 4. Lithologic photos of rock units sampled for detrital zircon analyses; rock hammer 1165 with 35 cm long handle for scale (except where noted) : A. Massive, medium-grained 1166 sandstone of the Revett Formation at location of sample 07SB306A; B. Alternating thin 1167 layers of fine quartzitic sandstone, tan siltstone and black argillite of Snowslip Formation 1168 at location of sample 08SB203A (open 8 cm diameter Brunton compass in red circle); 1169 C. Fine- to medium-grained quartzite ledges with the argillite of Half Moon Lake at 1170 location of sample 07SB305A (daypack for scale); D. Cut slab of thinly layered black 1171 slate, tan siltstone and white, graded fine sandstone of the Togo Formation at location 1172 of sample Togo3 (with cm scale); E. 15 cm medium-grained quartzite (above orange 1173 tape on rock hammer) of sample 07SB302B in thinner siltite-slate sequence of lower 1174 Wabash-Detroit Formation; F. Thick-bedded dolomite grading up to thinner bedded fine 1175 quartzite (at left margin) in upper Wabash-Detroit Formation at location of sample 1176 07SB303A (rock hammer for scale); G. Bold outcrops of medium-grained thick-bedded 1177 quartzite of Buffalo Hump Formation (closed 8 cm diameter Brunton compass in red 1178 circle) just uphill from roadcut location of sample BH1; H. Cleaved quartzitic pebbly

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1179 sandstone (12 cm long hammerhead for scale) at location of sample HB1; I. Cut slab of 1180 quartzitic pebbly mudstone of Monk Formation (scale labelled in cm) from site of sample 1181 DT3.

1182 Fig. 5. Probability density plots (Ludwig, 2008) for samples from this study (parentheses 1183 after sample number gives number of grains accepted for use in plot over the total 1184 number of grains analyzed). Two published samples are also plotted: Prichard 1185 Formation, Member C, from lowest Belt Supergroup (Lewis et al., 2007) and Cambrian 1186 Addy Quartzite (GLST-1: Linde et al., 2014; 2018). Arrows point to stratigraphic position 1187 of samples in generalized stratigraphic column on the left. Colored intervals behind 1188 probability plots are labelled below x-axis to identify igneous age intervals found in 1189 Precambrian crustal elements shown in Fig. 1.

1190 Fig. 6. Geochemical plots of Huckleberry lava and underlying mafic dikes cutting Deer 1191 Trail Group (data in Table 2): A. Chondrite-normalizedDraft multi-element spidergrams of 1192 basalt from the Huckleberry Formation (HB2), two underlying dikes (RM100 cuts 1193 Stensgar dolomite; Togo2 cuts Togo Formation), and Hawaiian lava (Laboratory 1194 standard BHVO-1, glassy lava from Kilauea caldera) for comparison. Chondrite- 1195 normalization values from Thompson et al.(1984). The lack of relative depletion of Nb 1196 and Ta (vertical pale blue stripe) indicates these are not subduction-related arc 1197 magmatic rocks but are akin to enriched mid-ocean ridge basalts (E-MORB), ocean- 1198 island basalts (OIB) and continental rift basalts. B. Three mafic igneous samples from 1199 this study (solid squares) plotted on triangular Hf-Th-Ta diagram of Wood et al. (1979) 1200 also suggest an E-MORB/continental rift origin. Shown for comparison are larger field 1201 of Huckleberry-Irene lavas and coeval Gunbarrel intrusive rocks in northwestern US and 1202 western Canada (Mackinder et al., 2019).

1203 Fig. 7. Middle Mesoproterozoic evolution of western Laurentia and its evolving ties to 1204 eastern Australia-eastern Antarctica as part of supercontinent Columbia (starting figure 1205 modified from Medig et al., 2014). A. At 1.45 Ga, eastern Australia-eastern Antarctica 1206 serves as the main detrital source to incipient rift basins in western North America from 1207 northern Yukon (PR1 basin) to the US-Canada border (Belt-Purcell basin) and nearly to 1208 the US-Mexico border (Yankee Joe and Trampas basins). The Yukon basin sediment

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1209 was primarily derived from 1.5 Ga plutonic rocks in the Mt. Isa region (Medig et al., 1210 2014), while the basins to the south were dominated by (Belt-Purcell) or had smaller 1211 contributions from (Yankee Joe-Trampas) 1.60-1.55 Ga plutonic rocks of the Mawson 1212 Continent and eastern Northern Australia cratons. B. From 1.45-1.39 Ga, northern 1213 Australia begins to rotate counterclockwise away from western Laurentia, while eastern 1214 Antarctica remains attached to Laurentia in the present-day southwestern US (Stewart 1215 et al., 2010). Magmatism, uplift and exhumation in the Yavapai-Mazatzal province 1216 (Jones et al., 2015), as well as in eastern Antarctica (Goodge et al., 2008), produce a 1217 flood of detritus in north-flowing rivers that deliver sediment to the Missoula and Lemhi 1218 Groups in the Belt-Purcell basin. At the end of that period, a last gasp of distinctive 1.58 1219 Ga zircon from the northern Mawson craton flows eastward into the western edge of the 1220 Belt-Purcell basin (this paper). C. From 1.37-1.30 Ga, the northern and southern 1221 Australia continue to rotate counterclockwise away from Laurentia, beginning to open a 1222 wedge-shaped ocean basin betweenDraft them, as suggested by pervasive intrusion of 1223 partial melts of upwelling asthenosphere in northern Yukon. The Deer Trail basin 1224 overlaps the western edge of the Belt-Purcell basin and continues to receive sediment 1225 from the south. Preliminary detrital zircon Hf isotopic data (Brennan et al., 2019) 1226 suggest that eastern Antarctica may still be contributing sediment to the basal strata 1227 (Togo Formation) of the Deer Trail Group and that Antarctica remains attached to 1228 Laurentia in the southwestern US at least until ca. 1.36 Ga. The presence of a single 1229 pluton of 1.58 Ma about 100 km east of Chewelah in the Eocene Priest River 1230 metamorphic core complex (Evans and Zartman, 1990) may indicate that a fragment of 1231 the western sediment-source craton was left behind with western Laurentia during the 1232 pre-Deer Trail rifting.

1233

1234 Supplemental File 1 – Data tables from LA-ICP-MS analyses of detrital zircons

1235 Supplemental File 2 – Metadata for sample preparation, LA-ICP-MS instrumentation 1236 and settings (after Horstwood et al., 2016)

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Draft

Fig. 1. Map of the Laurentian craton showing the distribution of Precambrian tectonic elements (adapted from Hoffman, 1988; Ross and Villeneuve, 2003; Jones et al., 2015). Greenland is restored to its position prior to Cenozoic rifting from North America. Dashed line labelled “Sri=0.706” in western North America (Armstrong, 1988) indicates the approximate (buried) rifted western margin of the craton. The Belt-Purcell basin and other mid-Proterozoic basins are labelled (references in text). Areas of Neoproterozoic strata in western North America taken from Moynihan et al., 2019. Abbreviations: BH--Buffalo Head, CB--Cumberland batholith, FS—Fort Simpson magmatic arc, GB—Great Bear magmatic arc, GF—Great Fall tectonic zone, H— Hottah terrane, W—Wopmay orogen. The area of Fig. 2 is indicated by small box on the west side of the Belt-Purcell basin.

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Draft

Fig. 2. Geologic map of Chewelah, WA area (adapted from Miller, 2001) showing locations of detrital zircon samples (red stars) and igneous geochemical samples (yellow stars). Line Z-Z’ gives the location of the generalized cross section shown in Fig. 3. Locations mentioned in text are labelled.

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Fig. 3. Generalized cross section from Z-Z’ (2.12 x vertical exaggeration). Unit colors and abbreviations are the same as in Fig. 2.

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Draft

Fig. 4. Lithologic photos of rock units sampled for detrital zircon analyses; rock hammer with 35 cm long handle for scale (except where noted) : A. Massive, medium-grained sandstone of the Revett Formation at location of sample 07SB306A; B. Alternating thin layers of fine quartzitic sandstone, tan siltstone and black argillite of Snowslip Formation at location of sample 08SB203A (open 8 cm diameter Brunton compass in red circle); C. Fine- to medium-grained quartzite ledges with the argillite of Half Moon Lake at location of sample 07SB305A (daypack for scale); D. Cut slab of thinly layered black slate, tan siltstone and white, graded fine sandstone of the Togo Formation at location of sample Togo3 (with cm scale); E. 15 cm medium-grained quartzite (above orange tape on rock hammer) of sample 07SB302B in thinner siltite-slate sequence of lower Wabash-Detroit Formation; F. Thick-bedded dolomite grading up to thinner bedded fine quartzite (at left margin) in upper Wabash-Detroit Formation at location of sample 07SB303A (rock hammer for scale); G. Bold outcrops of medium-grained thick-bedded quartzite of Buffalo Hump Formation (closed 8 cm diameter Brunton compass in red circle) just uphill from roadcut location of sample BH1; H. Cleaved quartzitic pebbly sandstone (12 cm long hammerhead for scale) at location of sample HB1; I. Cut slab of quartzitic pebbly mudstone of Monk Formation (scale labelled in cm) from site of sample DT3.

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Draft

Fig. 5. Probability density plots (Ludwig, 2008) for samples from this study (parentheses after sample number gives number of grains accepted for use in plot over the total number of grains analyzed). Two published samples are also plotted: Prichard Formation, Member C, from lowest Belt Supergroup (Lewis et al., 2007) and Cambrian Addy Quartzite (GLST-1: Linde et al., 2014; 2018). Arrows point to stratigraphic position of samples in generalized stratigraphic column on the left. Colored intervals behind probability plots are labelled below x-axis to identify igneous age intervals found in Precambrian crustal elements shown in Fig. 1.

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Fig. 6. Geochemical plots of Huckleberry lava and underlying mafic dikes cutting Deer Trail Group (data in Table 2): A. Chondrite-normalized multi-element spidergrams of basalt from the Huckleberry Formation (HB2), two underlying dikes (RM100 cuts Stensgar dolomite; Togo2 cuts Togo Formation), and Hawaiian lava (Laboratory standard BHVO-1, glassy lava from Kilauea caldera) for comparison. Chondrite- normalization values from Thompson et al.(1984). The lack of relative depletion of Nb and Ta (vertical pale blue stripe) indicates these are not subduction-related arc magmatic rocks but are akin to enriched mid- ocean ridge basalts (E-MORB), ocean-island basalts (OIB) and continental rift basalts. B. Three mafic igneous samples from this study (solid squares) plotted on triangular Hf-Th-Ta diagram of Wood et al. (1979) also suggest an E-MORB/continental rift origin. Shown for comparison are larger field of Huckleberry-Irene lavas and coeval Gunbarrel intrusive rocks in northwestern US and western Canada (Mackinder et al., 2019).

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Draft

Fig. 7. Middle Mesoproterozoic evolution of western Laurentia and its evolving ties to eastern Australia- eastern Antarctica as part of supercontinent Columbia. A. At 1.45 Ga, eastern Australia-eastern Antarctica serves as the main detrital source to incipient rift basins in western North America from northern Yukon (PR1 basin) to the US-Canada border (Belt-Purcell basin) and nearly to the US-Mexico border (Yankee Joe and Trampas basins). The Yukon basin sediment was primarily derived from 1.5 Ga plutonic rocks in the Mt. Isa region (Medig et al., 2015), while the basins to the south were dominated by (Belt-Purcell) or had smaller contributions from (Yankee Joe-Trampas) 1.60-1.55 Ga plutonic rocks of the Mawson Continent and eastern Northern Australia cratons. B. From 1.45-1.39 Ga, northern Australia begins to rotate counterclockwise away from western Laurentia, while eastern Antarctica remains attached to Laurentia in the present-day southwestern US (Stewart et al., 2010). Magmatism, uplift and exhumation in the Yavapai- Mazatzal province (Jones et al., 2015), as well as in eastern Antarctica (Goodge et al., 2008), produce a flood of detritus in north-flowing rivers that deliver sediment to the Missoula and Lemhi Groups in the Belt- Purcell basin. At the end of that period, a last gasp of distinctive 1.58 Ga zircon from the northern Mawson craton flows eastward into the western edge of the Belt-Purcell basin (this paper). C. From 1.37-1.30 Ga,

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the northern and southern Australia continue to rotate counterclockwise away from Laurentia, beginning to open a wedge-shaped ocean basin between them, as suggested by pervasive intrusion of partial melts of upwelling asthenosphere in northern Yukon. The Deer Trail basin overlaps the western edge of the Belt- Purcell basin and continues to receive sediment from the south. Preliminary detrital zircon Hf isotopic data (Brennan et al., 2019) suggest that eastern Antarctica may still be contributing sediment to the basal strata (Togo Formation) of the Deer Trail Group and that Antarctica remains attached to Laurentia in the southwestern US at least until ca. 1.36 Ga. The presence of a single pluton of 1.58 Ma about 100 km east of Chewelah in the Eocene Priest River metamorphic core complex (Evans and Zartman, 1990) may indicate that a fragment of the western sediment-source craton was left behind with western Laurentia during the pre-Deer Trail rifting.

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Table 1: Tabulation of statistically significant detrital zircon age populations (peak ages), calculated using the AGEPICK macro for Microsoft Excel (Gehrels, 2009).

Age range Age range Age range Age range Age range Age range Age range Age range Age range Age range Age range Estimated Stratigraphic Sample Formation n Ma Ma Ma Ma Ma Ma Ma Ma Ma Ma Ma Unit age unit 1100-1199 n 1250-1299 n 1300-1399 n 1400-1499 n 1490-1610 n 1611-1699 n 1700-1799 n 1800-1899 n 1900-1999 n 2400-2499 n 2600-2699 n GLST_1 Addy Cambrian 70 1100 4 1703 5 1814 6 2455 3 Cambrian 1742 5 1842 8 Addy Qtzite 1782 39 1867 9 DT_3 Monk Neoproterozoic 92 Windermere 1252 + 16a 46 1728 21 2653 4 HB_1 Huckleberry Neoproterozoic 101 Group 1143 + 19b 19 1434 32 1740 12 1851 10 1972 3 BH_1 Buffalo Hump Meso/Neoprot. 101 ?uncertain ? 1112 + 20c 30 1427 29 1820 14 2452 4 2695 3 07SB303A Wabash-Detroit upper Mesoproterozoic 11 1359 + 38d 3 1670 4 1742 4 1810 3 07SB302A Wabash-Detroit lower Mesoproterozoic 95 Deer Trail 1300 + 24e 5 1403 15 1684 19 1743 32 2412 3 2629 5 Togo3 Togo Mesoproterozoic 86 Group 1362 + 35f 4 1726 56 07SB305A Argillite of Half Moon Lk Mesoproterozoic 73 1376 + 27g 5 1666 46 08SB203A Snowslip Mesoproterozoic 28 1391 + 25h 6 1578 13 1653 10 Belt 07SB306A Revett Mesoproterozoic 90 1579 + 18i 46 1723 20 1828 4 Supergroup 04RL175 Prichard C Mesoproterozoic 107 1467 9 1607 34 1683 4 1725 12 1817 4

Early 1.34-1.48 Ma NA gap with Mojave- Main Grenville Mazatzal Trans-Hudson Earliest Paleo- Grenville arc granite-rhyolite belt within western Yavapai Archean Probable source of zircon grain population Province Province or 2.0-1.8 Ga arcs Proterozoic magmatism?DraftYavapai-Mazatzal Province source province Youngest age given as weighted mean age; MSWD values: a1.5, b0.9, c0.9, d1.2, e0.5, f0.9, g0.3, h1.0, i1.0; n = number of grains

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Table 2: Whole rock geochemistry of mafic igneous samples (latitude/longitude datum: WGS84). Togo2 RM100 HB2 Latitude 48.14993 48.16564 48.17204 Longitude -117.90683 -117.95802 -117.98010

Major Elements (anhydrous wt %): SiO2 48.7 48.9 50.2 TiO2 3.03 5.01 3.00 Al2O3 14.1 13.6 14.2 FeO* 14.5 18.0 13.4 MnO 0.21 0.07 0.20 MgO 6.60 7.60 6.00 CaO 8.58 1.62 8.57 Na2O 3.54 2.77 3.38 K2O 0.38 1.89 0.71 P2O5 0.36 0.56 0.31 LOI % 2.62 4.57 2.34

Trace Elements by XRF (ppm): Ni 61 45 77 Cr 70 73 138 Sc 32 37 35 V 347 462 374 Ba 83 303 192 Rb 27 35 15 Sr 292 35 161 Zr 171 298 176 Y 28 52 34 Nb 20.6 33.9 20 Ga 23 24 20 Draft Cu 59 45 184 Zn 119 101 108 Pb 3 6 2 La 20 28 17 Ce 48 73 43 Th 1 4 3 Nd 29 45 26 U 2 2 2

Trace Elements by ICP (ppm): La 20.0 31.5 17.2 Ce 46.4 74.1 42.6 Pr 6.51 10.2 6.03 Nd 29.0 45.2 27.2 Sm 7.24 11.4 7.07 Eu 2.62 3.6 2.17 Gd 7.05 11.9 7.45 Tb 1.10 1.9 1.23 Dy 6.29 11.2 7.28 Ho 1.16 2.3 1.41 Er 2.90 5.8 3.64 Tm 0.38 0.80 0.49 Yb 2.28 4.9 2.96 Lu 0.32 0.7 0.42 Ba 84.0 292 182 Th 1.7 3.63 1.59 Nb 21.5 34.0 19.9 Y 29.1 53.1 34.2 Hf 4.56 8.45 4.94 Ta 1.42 2.31 1.37 U 0.52 1.21 0.47 Pb 1.97 3.97 3.04 Rb 27.6 34.6 13.7 Cs 6.21 4.41 0.60 Sr 297 36.0 164 Sc 33.7 38.6 37.0 Zr 182 323 187

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