Paleoproterozoic geology of SW Montana: Implications for the paleogeography of the Wyoming craton and for the consolidation of Laurentia

The Wyoming craton is conventionally understood to have been an Archean protocontinent, one of several that make up North America today (Fig. 1). It carries evidence for some of the oldest rocks in the core of Laurentia (Chamberlain and Mueller, 2019; Mogk et al., 2020; Mueller and Frost, 2006). In the northwestern Wyoming Province¹ of southwest Montana, however, the Montana metasedimentary terrane (Mogk et al., 1992a, 1992b) has a significant Paleoproterozoic tectonic history that spans the time period 2.5–1.6 Ga and includes lithotectonic suites that are likely Paleoproterozoic in origin (Harms and Baldwin, 2021). This Paleoproterozoic history helps to frame the paleogeography of the Wyoming craton prior to and during its amalgamation with the Kenorland supercontinent and following breakup of that supercontinent, and it sets the stage for final consolidation of the stable Laurentian basement of today’s North America.

The Montana metasedimentary terrane (MMT) constitutes the northwestern division of the Wyoming Province in its present coordinates (Fig. 2). By area, the MMT is dominated by diverse quartzofeldspathic gneisses; some of these are likely paragneisses, but voluminous orthogneisses that are not metasedimentary in origin characterize the MMT (Mogk et al., 2004; Krogh et al., 2011). These quartzofeldspathic gneisses have yielded very old (>3.0 Ga) detrital and magmatic zircons (Mueller et al., 1993, 1998, 2004). The MMT is nevertheless distinguished by the presence of interlayered metasupracrustal rock sequences that include prominent marbles and amphibolites (Mogk et al., 1992a, 1992b), which have been well mapped in the Tobacco Root Mountains (Vitaliano et al., 1979), the Ruby Range (James, 1990), and the Madison Range (Kellogg and Williams, 2000) (Fig. 3).

Along its southeastern flank, the MMT is juxtaposed with the Beartooth-Bighorn magmatic zone (Fig. 2 herein; Mogk et al., 1992a, 1992b). The Beartooth-Bighorn magmatic zone constitutes the ancient core of the Wyoming craton; it was integrated and stabilized during a prominent magmatic phase from ca. 2.9 to 2.8 Ga (Mogk et al., 2020). The MMT did not share in this event; magmatic zircons of that age are conspicuously absent in the MMT (Mogk et al., 2020). There are several discontinuous, and possibly polygenetic, shear zones northwest of and parallel to the Montana metasedimentary terrane–Beartooth-Bighorn magmatic zone contact, including the Madison mylonite zone (Erslev and Sutter, 1990) and the Snowy shear zone (Fig. 3 herein; Erslev, 1992; Mogk et al., 1988; Reid et al., 1975), some parts of which record Paleoproterozoic displacement. Nevertheless, the MMT has been genetically linked to the Beartooth-Bighorn magmatic zone because both share a distinctive Pb isotopic fingerprint that is thought to indicate derivation from the same lithospheric domain and to be a unique characteristic of the Wyoming Province (Mueller et al., 1993; Wooden and Mueller, 1988). Whereas some components of the MMT were likely accreted to the Wyoming craton in the Archean and Paleoproterozoic, the quartzofeldspathic basement gneiss of the MMT in particular can be associated with the Wyoming Province before the earliest Paleoproterozoic. Thus, the MMT can provide evidence of Paleoproterozoic tectonic events affecting the paleogeography of the Wyoming craton as a whole.

The Paleoproterozoic geologic history of the MMT includes extended periods of tectonic quiescence for which no rock suites are preserved. At these times, it appears that the northwest (present coordinates) margin of the Wyoming craton was embedded within a stable continental interior. This stability was punctuated by three phases of tectonic change: (1) an orogenic event ca. 2.45 Ga, which likely saw the addition of allochthonous domains to the Wyoming craton; (2) intrusion of mafic dikes and sills ca. 2.06 Ga, associated with continental extension, rifting, and the reestablishment of an open continental margin; and (3) conversion of that margin to an upper-plate subduction system with a fringing arc and associated back-arc basin by ca. 1.9 Ga. This subduction system produced the climactic event along the Wyoming craton’s northwestern margin—the ca. 1.8–1.7 Ga Big Sky orogeny, during which the Medicine Hat block and Selway terrane were added to the northern, northwestern, and western (present coordinates) margins of the craton (Fig. 2).

The Big Sky orogeny imposed upper-amphibolite– to granulite-grade metamorphic conditions and penetrative ductile deformation across all but the southernmost MMT (Alcock et al., 2013; Baldwin et al., 2017; Cheney et al., 2004a; Condit et al., 2015; Harms et al., 2004b); the Big Sky orogen is defined as the domain that preserves evidence of this tectonism (Fig. 2). K-Ar, ⁴⁰Ar/³⁹Ar, and monazite geochronologic analyses have identified a significant geochronologic front, known as “Giletti’s line” (Fig. 3), northwest of which rocks of the MMT experienced growth of monazite and complete resetting of lower-temperature chronometers during Big Sky orogenesis (Giletti, 1966; Brady et al., 2004a; Cheney et al., 2004b; Roberts et al., 2002). North of Giletti’s line, most metasupracrustal suites record not only Big Sky orogeny recrystallization, but also preserve relict evidence of an older metamorphic event ca. 2.45 Ga. However, several distinctive suites yield only ca. 1.8–1.7 Ga chronometers, and are likely to have been juvenile suites at the time of the Big Sky orogeny (Cheney et al., 2004b; Harms and Baldwin, 2021; Roberts et al., 2002). In the MMT that lies south of Giletti’s line, both monazite and ⁴⁰Ar/³⁹Ar systems record equilibration ca. 2.55 or ca. 2.45 Ga (Hames and Harms, 2013; Lloyd et al., 2012). These rocks avoided the effects of the Big Sky orogeny, except in discrete shear zones such as the Madison mylonite zone (Erslev and Sutter, 1990). The metasupracrustal suites that characterize the MMT can thus be divided into those that must predate the ca. 2.55 or 2.45 Ga orogenic events and those that only experienced the ca. 1.8–1.7 Ga Big Sky orogeny. These contrasting histories across the MMT preserve and expose the polyphase Paleoproterozoic evolution of the northern margin of the Wyoming craton.

The past several decades have seen the steady accumulation of a range of geochronological analyses across the MMT that provide evidence for magmatism and metamorphism in the period ca. 2.55–2.45 Ga (Baldwin et al., 2017; Cheney et al., 2004b; Condit et al., 2015; Foster et al., 2012; Hames and Harms, 2013; Kellogg et al., 2003; Lloyd et al., 2012; Mogk et al., 1988; Mueller et al., 2011, 2012; Roberts et al., 2002). Dates in the range ca. 2.55–2.45 Ga have been attributed to a Tendoy orogeny (Mueller et al., 2012; Baker et al., 2017; Baldwin et al., 2014) or a Beaverhead (and Beaverhead–Tobacco Root) orogeny (Jones, 2008; Krogh et al., 2011). From the body of data now available, however, there has emerged a pattern of ages either at ca. 2.55 Ga, generally clustered south of Giletti’s line, or at ca. 2.45 Ga, mostly north of Giletti’s line and along the western margin of the MMT. The significant temporal separation of these two dates warrants the recognition of two distinct events. We elect to use the term “Beaverhead” orogeny for the ca. 2.45 Ga dates, as that name was first applied in areas north of Giletti’s line, where dates of that age are concentrated; we reserve the term “Tendoy” orogeny for ca. 2.55 Ga dates. We have few constraints on the ca. 2.55 Ga, end-Archean, Tendoy event and will focus here only on what can be known about the ca. 2.45 Ga, early Paleoproterozoic, Beaverhead orogeny.

North of Giletti’s line, overprinting of the MMT during the ca. 1.8–1.7 Ga Big Sky orogeny was sufficiently severe that older, ca. 2.45 Ga orogenic effects can only be understood in broad strokes. Nevertheless, there is widespread evidence for metamorphism accompanied by crustal-melt plutonism in the MMT at that time. North of Giletti’s line, most—but notably not all—metasupracrustal sequences preserve ca. 2.45 Ga monazites as small matrix grains and as relict cores within 1.8–1.7 Ga monazite grains included in porphyroblasts (Cheney et al., 2004b; Pearson and Cheney, 2008). The intensity of the Beaverhead orogenic period is well expressed in the Ruby Range, where the Dillon Gneiss and Lower Christensen Ranch metasupracrustal rocks host a widespread and visually distinctive garnet leucogneiss, presumably the result of partial melting of continental crust. This garnet leucogneiss is characteristically mylonitic and locally includes highly attenuated plagioclase feldspar, attesting to mylonitization at very high, subsolidus temperatures (Fig. 4). The garnet leucogneiss has magmatic zircon, garnet, and elongated monazite, all dated at ca. 2.45 Ga (Alcock et al., 2013; Baker et al., 2017; Baldwin et al., 2017; Harms and Baldwin, 2019; Mukunda et al., 2015), confirming a syntectonic origin for the rock.

In the Highland and Tobacco Root Mountains, quartzofeldspathic gneisses are crosscut by a suite of mafic dikes and sills containing zircon that has been U-Pb dated at 2.06 ± 0.006 Ga (Mueller et al., 2004, 2005). Although subject to metamorphism and recrystallization during the younger Big Sky orogeny, this relationship demonstrates that the original gneissic fabric must be older than 2.06 Ga. Furthermore, the metamorphosed mafic dikes and sills cut surrounding gneisses with sharp contacts, suggesting a cold host at the time of intrusion (Brady et al., 2004b). Although not directly dated, the gneissic fabric, therefore, could also be attributable to the ca. 2.45 Ga Beaverhead event (Harms et al., 2004a).

Metasupracrustal rocks that are interlayered with quartzofeldspathic gneisses cut by 2.06 Ga dikes in the Tobacco Root and Highland Mountains, and that are intruded by 2.45 Ga garnet leucogneiss in the Ruby Range, include thick marble, prominent amphibolite, and subordinate iron formation, pelitic schist, and calc-silicate schist. The protoliths of these metasupracrustal rock types are interpreted to reflect the presence of an active (basalt-producing), shallow-water (carbonate-sustaining) continental margin along this part of the Wyoming craton leading up to the Beaverhead orogeny.

Metamorphosed mafic dikes and sills (Fig. 5) that cut the quartzofeldspathic gneiss complexes in the Tobacco Root and Highland Mountains (Brady et al., 2004b; Hanley and Vitaliano, 1983) have a ²⁰⁷Pb/²⁰⁶Pb zircon intrusive age of 2.06 ± 0.006 Ga (Mueller et al., 2004, 2005). Having been subject to metamorphism and deformation during the younger Big Sky orogeny, these mafic rocks contain metamorphic mineral parageneses, and, in many locations, they have been transposed to near parallelism with the host gneissic fabric (pseudo-sills). Nevertheless, these mafic rocks retain relict igneous textures, including chilled margins; in areas of low strain, the intrusive bodies crosscut gneissic banding at a high angle (true dikes; Brady et al., 2004b; Harms et al., 2004a).

Despite high-grade metamorphism, the mafic intrusive rocks retain an original tholeiitic bulk composition and have modestly enriched rare earth element (REE) patterns that are 10–100 times chondrite (Brady et al., 2004b; Castro et al., 2014; Krogh et al., 2011), consistent with a continental rift setting. Because the MMT was subsequently the site of a subduction-related volcanic arc and was profoundly affected by a collisional orogeny at ca. 1.8–1.7 Ga, the most compatible interpretation is that 2.06 Ga mafic magmatism progressed to complete rifting and resulted in the post–Beaverhead orogeny reestablishment of an open continental margin along this part of the Wyoming craton (Harms et al., 2004a).

Several spatially limited but geologically significant lithotectonic suites in the northern MMT constrain paleogeographic models for the period following 2.06 Ga rifting (Fig. 3). These suites are metasupracrustal in character, but they are distinct from all other metasupracrustal suites of the MMT because they yield only monazites that are younger than 1.9 Ga in age (Cheney et al., 2004b; Pearson and Cheney, 2008; Baldwin et al., 2017), and they are not cut by either 2.06 Ga metamorphic dikes and sills or the 2.45 Ga garnet leucogneiss, despite close proximity today to quartzofeldspathic gneisses that do host these intrusions. These suites are interpreted (1) to have been tectonically juxtaposed with older quartzofeldspathic gneiss suites; (2) to have accumulated in a period of subduction-related arc activity; and (3) to be certainly younger than 2.06 Ga and likely ca. 1.9–1.8 Ga in age and origin (Harms and Baldwin, 2021; Harms et al., 2004a).

These critical lithotectonic suites are:

• (1) Arc rocks of the Little Belt Mountains: The small area of exposed Precambrian crystalline basement rocks in the Little Belt Mountains of north-central Montana (Fig. 2) is dominated by calc-alkaline, dioritic orthogneisses (Vogl et al., 2004). These gneisses have trace-element patterns and primitive initial Nd isotopic ratios consistent with genesis in a subduction setting, and concordant and upper-intercept U-Pb zircon ages ranging from ca. 1.87 to 1.79 Ga (Mueller et al., 2002; Vogl et al., 2004). Dioritic and quartzofeldspathic basement orthogneiss in the Pioneer Mountains (Fig. 2) and Biltmore anticline in southwesternmost Montana yielded similar concordant and upper-intercept zircon crystallization ages of ca. 1.89–1.86 Ga (Foster et al., 2006). From this, we interpret the northwesternmost margin of the Wyoming craton to have been an active subduction zone with a correlated volcanic arc in at least the period 1.89–1.79 Ga, with the younger age marking the demise of this subduction system.

• (2) Back-arc basin oceanic crust of the Spuhler Peak metamorphic suite in the Tobacco Root Mountains: The distinctive Spuhler Peak metamorphic suite of the Tobacco Root Mountains is dominated by mafic schists, which constitute as much as 90% of the suite, but also includes minor aluminous schist, quartzite, and meta-ultramafic bodies (Burger et al., 2004). Marble is absent. The mafic schists are (a) hornblende amphibolites from tholeiitic basalt protoliths (Fig. 6), and (b) orthoamphibolites, representing hydrothermally altered mafic volcanic protoliths (Burger et al., 2004), suggesting they originated in a setting of active ocean crust formation. Trace-element concentrations (Fig. 6) are consistent with an arc-related origin (Burger et al., 2004). We interpret the Spuhler Peak metamorphic suite to represent former ocean crust from a suprasubduction, intra-arc or back-arc basin with active spreading centers.

• (3) Back-arc basin sediments and flows in the Highland Mountains: Biotite-sillimanite-garnet schists and gneisses and their texturally mylonitic equivalents, associated with abundant concordant layers of amphibolite, are widespread and voluminous in the Precambrian crystalline core of the Highland Mountains (O’Neill et al., 1996). Major- and trace-element analyses of the amphibolites show they were subalkaline tholeiites, with mid-ocean-ridge–, back-arc–, and arc-related affinities (Rioseco et al., 2016). The most likely basin of accumulation for the sedimentary and volcanic protoliths of these metasupracrustal rocks would be a suprasubduction back-arc basin (Harms and Baldwin, 2021).

• (4) Active continental margin sediments and flows in the upper Christensen Ranch suite of the Ruby Range: Metasupracrustal rocks of the Christensen Ranch suite in the western Ruby Range include massive marble layers, calc-silicate, quartzite, iron formation, pelitic schist and gneiss, hornblende amphibolite, orthoamphibole schist, and minor quartzofeldspathic gneiss (James, 1990; Dahl, 1979). Above a structural discontinuity in the upper part of the suite, only ca. 1.8–1.7 Ga monazite can be found, and the unit is not intruded by the ca. 2.45 Ga garnet leucogneiss that crosscuts the structurally lower half of the suite, where both ca. 1.8–1.7 Ga and ca. 2.45 Ga monazites occur (Baldwin et al., 2017; Harms and Baldwin, 2019). We conclude that, following the Beaverhead orogeny in this part of the Wyoming craton, and subsequent to continental extension at 2.06 Ga, a carbonate + basalt–bearing, active continental margin was reestablished along the northern Wyoming craton, giving rise to the protoliths of the upper Christensen Ranch suite. This margin would have faced the back-arc basin associated with the Little Belt arc and may have been intermittently affected by back-arc, extension-related basaltic magmatism.

The protolith characteristics and present geographic distribution of these four subduction-related Paleoproterozoic suites are best integrated in a model of south-directed subduction (present coordinates) beneath the northern Wyoming craton (Fig. 7) in which, by ca. 1.9 Ga, south-dipping subduction had created a fringing arc on the Wyoming craton plate and had opened a back-arc basin between the arc and the continental margin of the craton. This is a significant departure from previous interpretations in which the Wyoming craton occupied a lower-plate position relative to a north-dipping subduction zone (Erslev and Sutter, 1990; O’Neill, 1998; Harms et al., 2004a; Mueller et al., 2004, 2005; Vogl et al., 2004; Whitmeyer and Karlstrom, 2007). A more detailed understanding of the age and character of juvenile suites of the MMT, however, requires this revised model. The absence of a purely metasedimentary sequence without significant interlayered metabasalts, as would typify a passive continental margin developed over the ~200 m.y. between rifting at 2.06 Ga and the ca. 1.8 Ga onset of the Big Sky orogeny, also argues for an upper-plate position for the Wyoming craton during this time period. At 1.79 Ga, as collision occurred along this convergent plate boundary, the Little Belt arc ceased activity, the arc and its back-arc basin were obducted onto the Wyoming craton, and the metamorphism and deformation that characterize the Big Sky orogeny during the period ca. 1.78–1.72 Ga began (Cheney et al., 2004a, 2004b; Harms et al., 2004a, 2004b).

To the north and northwest of the MMT, in western Montana, Idaho, and eastern Washington, there are only very limited, isolated exposures of pre–Belt Supergroup, Precambrian basement rocks that would have been situated outboard of the Little Belt arc in the Paleoproterozoic. Paleoproterozoic rocks in this region have been assigned to the Selway terrane (Fig. 2); they share a characteristic geochronologic profile of zircon crystallization ages at both Beaverhead and Big Sky orogeny time (Foster et al., 2006). Notably, zircons older than ca. 2.45 Ga are not in evidence; the Selway terrane likely represents additions to the Wyoming Province that were juvenile at or after the time of the Beaverhead orogeny (Foster et al., 2006). The correlation of Paleoproterozoic age profiles between the MMT and the Selway terrane suggests some shared Paleoproterozoic history and may indicate the pre–Big Sky orogeny existence of a broad pericratonic domain of crustal fragments that had rifted or partially rifted away at 2.06 Ga and then became reamalgamated with the Wyoming craton during the Big Sky orogeny.

In northern Montana and in southern Canada, north of the Little Belt arc, the Medicine Hat block (Fig. 2) is known only from drill core (Ross, 2002), but it is presumed to have been part of the converging plate geometry outboard of the Little Belt arc (Fig. 7).

Both ⁴⁰Ar/³⁹Ar chronology from metasupracrustal rocks in the MMT (Hames and Harms, 2013) and apatite chronology from xenoliths that represent the deep crust in the Big Sky orogen suture zone (O’Sullivan et al., 2021) point to a long residence time in the middle crust and slow thermal reequilibration in the MMT following the end of active tectonism during the Big Sky orogeny. We interpret this to indicate that the northern Wyoming craton was once again embedded within a stable continental body—Laurentia. The northern Wyoming craton would remain so until initiation of Belt basin extension some 300 m.y. later.

Because of its notable antiquity, the Wyoming craton must figure into reconstructions of supercontinents or supercratons at any and all times in the Precambrian. The geologic history of the MMT provides three spatial and temporal constraints on Wyoming craton paleogeography in reconstructions for latest Archean through Paleoproterozoic time (Fig. 8). (Different reconstructions have been proposed for the late Archean, including the Kenorland supercontinent [Williams et al., 1991] and the Superia [Bleeker, 2003] or Arctica [Rogers and Santosh, 2003] supercratons [Evans, 2013]. Insofar as these all link the Wyoming and Superior cratons, we will not distinguish between them here but will refer to them collectively under the umbrella of Kenorland.)

While general comparisons have been used to associate the Superior and Wyoming cratons by at least 2.6 Ga (Bleeker, 2003; Roscoe and Card, 1993; Williams et al., 1991), the well-established, specific correlation of the early Paleoproterozoic Huronian Supergroup of the southern Superior Province with the homologous Snowy Pass Supergroup in the Medicine Bow Mountains of the southern Wyoming Province links the Wyoming craton with the Kenorland supercraton in an “inverted” position relative to present coordinates at around 2.45 Ga (Heaman, 1997; Roscoe and Card, 1993). This reconstruction has been substantiated by paleomagnetic data from the central Wyoming Province (Figs. 8A and 8B; Kilian et al., 2016a). (The 2.45 Ga mafic intrusions and flows have also linked Karelia to the Wyoming and Superior cratons, with Karelia adjacent to today’s western edge of the Wyoming Province [Heaman, 1997]. The geologic character of the MMT offers no independent constraint on this reconstruction.)

In its “inverted” position within Kenorland, today’s northern margin of the Wyoming craton—the MMT—would have constituted the supercontinent’s southern margin from >2.6 to ca. 2.45 Ga. This margin is likely to have been the site of accumulation of at least some of the >2.45 Ga metasupracrustal suites that characterize the MMT, developing in a shallow-marine setting permissive of carbonate accumulation, intermittently interrupted by mafic volcanism. A semipassive continental margin facing a back-arc basin, with periodic rifting, seems fitting (Fig. 8A).

At 2.45 Ga, the geologic setting of the MMT differed significantly from that of the Snowy Pass Supergroup on the opposite side of the Wyoming craton (Fig. 8B). The Snowy Pass and Huronian Supergroups accumulated in epicratonic, extensional basins and are tectonically linked to ca. 2.45 Ga mafic dike swarms and extensive flood basalt flows in the southern Superior Province in the context of an episode of failed supercontinent rifting (Heaman, 1997; Roscoe and Card, 1993). The ca. 2.45 Ga record in the MMT is one of deformation and metamorphism in an orogenic setting (not reasonably attributable to regional heating in a plume setting). Its “inverted” position within Kenorland would have placed the MMT along a free face conducive to collisional tectonics both at 2.55 Ga (Tendoy event) and again at 2.45 Ga (Beaverhead orogeny). It seems possible that collisions along the southern margin of the Wyoming craton, at the same time as extension and incipient rifting along its northern rim, could have played a far-field role in determining that the Huronian and Snowy Pass Supergroup basins ultimately would be failed rift basins.

This leaves the obvious question of what might have collided with the southern margin of Kenorland along the MMT margin of the Wyoming craton at ca. 2.45 Ga. Tectonism in the age range 2.55–2.45 Ga is shared by what Pehrsson et al. (2013) consider the “Rae family” of cratons, which includes presently far-flung cratonic blocks such as the Gawler, Amazonia, North China, and West Africa cratons, among others. Any one or any set of these cratons is a likely collider; here, we use the Rae craton as a placeholder for whichever blocks in the Rae family formed the other side of the Beaverhead orogen (Fig. 8B). Following the Beaverhead orogeny, the geologic history of the MMT is one of quiescence until 2.06 Ga. Whatever collided with the southern margin of Kenorland at ca. 2.45 Ga stayed well stuck in place for some 400 m.y., keeping the MMT buffered within a larger continental mass.

The period 2.2–2.0 Ga saw intrusion of mafic dike swarms, continental rifting, and the ultimate breakup of Kenorland, a process that formed the outlines of the Archean cratons of Laurentia as we know them today. In particular, based on the ages of the Fort Francis and Marathon dike swarms in the Superior craton (Ernst and Bleeker, 2010), and the age of the Kennedy dike swarm in the southern (present coordinates) Wyoming craton (Chamberlain et al., 2003; Cox et al., 2000), the Wyoming Province rifted from the Superior craton between ca. 2.125 and 2.01 Ga (Fig. 8C).

At approximately the same time, the 2.06 Ga igneous protoliths of the metamorphosed mafic dikes and sills were intruded into basement quartzofeldspathic gneisses of the MMT in the northern (present coordinates) Wyoming Province. We interpret these metamorphosed mafic dikes and sills to be the expression of a period of continental rifting that would result in the removal of whatever Rae family continental body had accreted to the Wyoming craton at ca. 2.45 Ga, having encased it in a stable continental interior (Fig. 8C). Rifting associated with the 2.06 Ga dikes would reestablish a continental margin along the MMT side of the Wyoming craton, recorded as the upper Christensen Ranch metasupracrustal suite.

Continental fragments may have rifted away from the Wyoming craton or from the Rae family landmass, or both, during breakup, creating a penumbra of crustal blocks bearing the imprint of ca. 2.45 Ga metamorphism and plutonism outboard of the MMT rifted margin (Fig. 8C) that today form the basement of the Selway terrane (Foster et al., 2006).

Contemporaneous rifting along what would become both its southern and northern continental margins would leave the Wyoming craton as an island continent. Spreading along the divergent boundary that developed on the former Kenorland side of the Wyoming craton started the rotational motion necessary to bring the Wyoming craton into the “right-side-up” position that it occupies today (Fig. 8C). (The same might be said for the Rae craton, as a consequence of combined ocean plate production at both new divergent boundaries.) Growth of these new ocean basins may have continued at an intra-oceanic divergent plate boundary or boundaries for some period of time following rifting at ca. 2.1 Ga. The ages of the oldest arc-related rocks in the Cheyenne belt south of the southern Wyoming Province (Jones et al., 2010; Whitmeyer and Karlstrom, 2007) and in the Little Belt arc (Mueller et al., 2002; Vogl et al., 2004), however, indicate that subduction and ocean closure were under way by ca. 1.89 Ga. The rifted continental margin of the Wyoming craton along the MMT margin cannot have been passive for very long; there is no evidence for a long-lived, post–2.06 Ga stable margin in the conventional Phanerozoic sense within the MMT. Instead, the interlayered metasedimentary and metamafic rocks of the upper Christensen Ranch suite indicate that the MMT margin was rather quickly converted to convergence, where the Wyoming craton occupied a suprasubduction position, flanked by a back-arc basin (Fig. 8D).

Plate divergence and ocean basin expansion along its former Kenorland margin and convergence along its MMT margin would have rotated and translated the Wyoming craton into the realm of the Manikewan Ocean, the Paleoproterozoic ocean basin between the Superior and Rae cratons, the closure of which was responsible for the metamorphism, deformation, and tectonism that defines the Trans-Hudson orogen in Canada (Fig. 8D; Stauffer, 1984; Corrigan et al., 2009). In this position, ocean crust both north and south of the insular Wyoming craton could have been southern arms of the greater Manikewan Ocean. The Wyoming craton and its surrounding plate boundaries would have been integral components in the closure of the Manikewan Ocean over the period 1.92–1.80 Ga, as bracketed by rock suites in the Trans-Hudson orogen (Corrigan et al., 2009). The ca. 1.8–1.7 Ga Big Sky orogeny represents closure of that part of the Manikewan Ocean that lay between the Wyoming and Hearne cratons (Fig. 8E). This phase of Wyoming craton paleogeography is certainly the least speculative, as Manikewan Ocean closure resulted in the present juxtapositions of Archean cratons in North America’s crystalline basement. Observations from the MMT suggest that along-strike correlations and contrasts can be drawn between the Paleoproterozoic history of the Wyoming Province and that of the Trans-Hudson orogen in Canada, in particular, the Flin Flon and Thompson Nickel belts (Fig. 8D), which could refine our understanding of the process by which Paleoproterozoic consolidation happened. The geologic character of juvenile Paleoproterozoic suites of the MMT and the tectonics of the Big Sky orogen suggest the following conclusions.

• (1) Subduction polarity correlates along strike. The destruction of the Manikewan oceanic plate appears to have been the consequence of SE-dipping subduction along craton margins on the SE side of the ocean, i.e., the Thompson Nickel belt along the western Superior craton (Corrigan et al., 2009) and the Little Belt arc along the MMT margin of the Wyoming craton, and of NW-dipping subduction to the NW beneath the Flin Flon complex along the southernmost exposures of the Trans-Hudson orogen in Canada (Fig. 8D herein; Corrigan et al., 2009). While there are no exposures of Paleoproterozoic rocks north of the Little Belt Mountains to test the possible presence of rocks related to NW-dipping subduction (Fig. 2), a deep seismic refraction profile that ran from the northern Hearne Province to the southern Wyoming Province identified two north-dipping reflectors in the lithospheric mantle, which were interpreted as fossil subduction zones (Gorman et al., 2002). One reflector projects upward toward the southern end of the Medicine Hat block, and one projects toward the southern end of the Hearne Province (Gorman et al., 2002), neither of which are positioned to have been related to subduction at the Little Belt arc. We propose that the Big Sky orogeny in Montana was the consequence of the closure of an arm of the Manikewan Ocean by subduction both to the south below the Little Belt arc and to the north below the Medicine Hat block (Fig. 7), correlative with the positions and polarities of subduction zones to the north in Canada.

• (2) The Wyoming craton was the last to arrive. Subduction initiation in the Little Belt arc was contemporaneous with that in the Flin Flon and Thompson belts at ca. 1.89 Ga (Corrigan et al., 2009; Mueller et al., 2002; Vogl et al., 2004), but it lasted slightly longer at the Little Belt arc. The youngest known dioritic gneisses associated with Little Belt arc activity are 1.791 ± 0.01 Ga (Vogl et al., 2004), but, to the north, collision involving the Superior craton and the periphery of the Hearne craton was already under way by 1.83 Ga (Corrigan et al., 2009). Metamorphism associated with collision along the Big Sky orogen spanned 1.78–1.71 Ga, with a peak between 1.78 and 1.73 Ga (Cheney et al., 2004b; Harms et al., 2004a). In contrast, peak metamorphism occurred at 1.82–1.77 Ga across the Trans-Hudson orogen (Schneider et al., 2007). In so far as the geology of the Black Hills reflects the development of the Dakotan orogen, metamorphism brackets collision there at ca. 1.77–1.71 Ga (Allard and Portis, 2013; Chamberlain and Mueller, 2019; Dahl et al., 1999). It appears that final consolidation of Laurentian Archean cratons came with the docking of the Wyoming craton along both its northern and eastern margins (see also Killian et al., 2016b).

• (3) In comparison to the Trans-Hudson orogen, the Big Sky orogen was narrow, hot, and high. Well-constrained peak metamorphic pressures of ~1.1 GPa and peak temperatures of ~800 °C along a clockwise pressure-temperature path have been documented in both the juvenile Paleoproterozoic Spuhler Peak suite and in metasupracrustal units within basement quartzofeldspathic gneisses of the MMT in the Tobacco Root Mountains, requiring burial of supracrustal suites under at least 25 km of tectonic overburden (Cheney et al., 2004a; Harms et al., 2004a). Elsewhere in the MMT, peak pressures were in the general range of 0.8–1.0 GPa (Harms and Baldwin, 2021). In contrast, metamorphic pressures achieved during the Trans-Hudson orogeny were only 0.3–0.7 GPa (Corrigan et al., 2009).

On the other hand, the Trans-Hudson orogen, encompassing the area of metamorphism attributed to collision between the Superior, Slave, Rae, and Hearne cratons (Corrigan et al., 2009), has been compared to the Himalayan orogen in width and other characteristics (St-Onge et al., 2006). The Wyoming Province is ringed to the west, north, and east by accretionary domains that, for the most part, consist of continental fragments formed during ca. 2.45 Ga, Beaverhead orogeny–age tectonism and of juvenile suites originating in the ca. 1.9–1.8 Ga period of convergence (Foster et al., 2006; Harms and Baldwin, 2021; Dahl et al., 1999). These are the Selway terrane, the area surrounding the Little Belt arc and south of the Medicine Hat block (commonly referred to as the Great Falls tectonic zone [O’Neill and Lopez, 1985]), and the Dakotan orogen, including the Black Hills of South Dakota (Fig. 2). Along with the Big Sky orogen, these accretionary belts constitute the damage zone associated with docking of the Wyoming craton. When compared to the width of the Trans-Hudson orogen, however, this is a much narrower realm of orogenic effects. It appears that the Wyoming craton’s collision resulted in more intense deformation but over a more restricted area.

The Paleoproterozoic geologic history of the MMT, situated on the northern flank of the Wyoming Province today, provides greater paleogeographic detail to reconstructions of Kenorland, its breakup, and the construction of Laurentian basement. The MMT preserves evidence of collisional orogeny at ca. 2.45 Ga, which is unique within Kenorland but which correlates with a period of tectonism shared by the Rae family of cratons (Pehrsson et al., 2013). This provides a possible link between Kenorland and the proposed Nunavutia supercraton (Pehrsson et al., 2013) at that time. If so, this was a strong, stable, and long-lived link, because 2.45 Ga collision in the MMT was followed by 400 m.y. of tectonic quiescence, signaling that the Wyoming craton was positioned within a larger continental interior.

Advances in characterizing the time and tectonic setting of the origin of juvenile suites younger than 2.06 Ga in the MMT permit us to create a refined model for the role of the Wyoming craton in the construction of Laurentia’s Precambrian basement. We now understand that, although it may have been paired with north-dipping subduction beneath the Medicine Hat block and Hearne craton, collision along the Big Sky orogen resulted from south-dipping subduction beneath the Wyoming craton. A similar subduction geometry of opposing polarities existed in the contemporaneous Flin Flon and Thompson Nickel belt subduction zones along strike to the northeast (Corrigan et al., 2009). Subduction persisted longer, and terminal collision came later, however, along the Little Belt arc and Big Sky orogen than in the Trans-Hudson orogen. Consolidation of Archean cratons into Laurentian basement was only complete once the Wyoming craton had docked with the Medicine Hat block–Hearne craton and with the Superior craton along the Big Sky and Dakotan orogens, respectively.

We are grateful to the many colleagues who introduced us to the Montana metasedimentary terrane and have worked with us to study this area: Mike O’Neill, Dave Mogk, Paul Mueller, Bob Burger, John Brady, and Jack Cheney. No advancement in understanding the MMT would have been possible without the many contributions of students from Amherst College, the Keck Geology Consortium, and the University of Montana. The clarity of presentation was improved by the careful reviews of Stephen Allard and Darrell Henry; we thank them for their efforts.