ABSTRACT
The Mesoproterozoic is a controversial time within the Earth’s history, and is characterized by high temperature/pressure ratios in metamorphic rocks, a large volume of extensional plutons, very few economic mineral deposits, and possibly a slowdown in plate tectonic processes. In Laurentia, ca. 1.48–1.35 Ga is well known as a time of voluminous ferroan magmatism, which led to conflicting tectonic interpretations that range from continental extension to convergent margin settings. Recently, a ca. 1.50–1.35 Ga orogenic belt was proposed that spanned Laurentia from present-day eastern Canada to the southwestern United States. Unlike the preceding Paleoproterozoic Yavapai/Mazatzal orogenies and the subsequent late Mesoproterozoic Grenville orogeny, the early–mid-Mesoproterozoic Picuris orogeny in the southwestern United States was relatively unrecognized until about two decades ago, when geochronology data and depositional age constraints became more abundant. In multiple study areas of Arizona and New Mexico, deposition, metamorphism, and deformation previously ascribed to the Yavapai/Mazatzal orogenies proved to be part of the ca. 1.4 Ga Picuris orogeny. In Colorado, the nature and extent of the Picuris orogeny is poorly understood. On this trip, we discuss new evidence for the Picuris orogeny in the central Colorado Front Range, from Black Hawk in the central Colorado Front Range to the Wet Mountains, Colorado. We will discuss how the Picuris orogeny reactivated or overprinted earlier structures, and perhaps controlled the location of structures associated with Cambrian rifting, the Cretaceous–Paleogene Laramide orogeny, and the Rio Grande rift, and associated mineralization. We will also discuss whether and how the Picuris orogeny, and the Mesoproterozoic in general, were unique within the Earth’s history.
INTRODUCTION
Globally, the Mesoproterozoic is characterized by a number of anomalies compared with pre- and post-Mesoproterozoic periods. These include a decrease in the number of passive margins, detrital zircons, eclogites, granulites, carbonatites, and orogenic gold (Bradley et al., 2011); higher temperature/pressure (T/P) ratios of metamorphic rocks; and more juvenile and thinned crust (Brown et al., 2020; Spencer et al., 2021). The reason for these anomalies is not known. It has been interpreted previously as a slowdown of tectonic processes during the Columbia/Nuna-Rodinia supercontinent transition, or as a period of plate tectonics characterized by hot, thin, and low orogens (Spencer et al., 2021, and references therein). Stern (2021) considers the possibility of a single-lid tectonic episode during the Mesoproterozoic, where plate tectonic movement was stagnant. Conversely, based on geological observations and lithospheric-scale geodynamic modeling, others argue that plate tectonics have operated in a manner similar to today since ca. 3.0 Ga (e.g., Hawkesworth et al., 2016, 2017; Dhuime et al., 2018; Palin et al., 2020). The Mesoproterozoic Picuris orogen of the southwestern United States (Daniel et al., 2013, 2022a, 2022b) provides an opportunity to investigate and discuss tectonic processes of the Mesoproterozoic. The focus of this field trip is to examine evidence for, and the nature and extent of, the Picuris orogen in the Colorado Front Range. We will also discuss how the Picuris orogeny reactivates earlier Paleoproterozoic structures and may control later Phanerozoic structures.
Proterozoic rocks of the southwestern United States (Figs. 1, 2) preserve a record of significant Paleoproterozoic crustal growth southward from the Archean Wyoming Province (present-day coordinates), and experienced multiple orogenic events including the Paleoproterozoic Yavapai and Mazatzal orogenies (Whitmeyer and Karlstrom, 2007; Jones et al., 2009; Holland et al., 2020), and the Mesoproterozoic Picuris orogeny (Daniel et al., 2013; Aronoff et al., 2016). Voluminous ca. 1.48–1.35 Ga granitic magmatic rocks are also associated with the Mesoproterozoic, and make up ca. 25%–30% of the exposed Proterozoic rocks in the southwestern United States. The occurrences of significant ferroan, granitic magmatic rocks, typically attributed to an extensional environment (Anderson and Morrison, 2005; Frost and Frost, 2022), overlap in time and space with crustal thickening associated with the contractional Picuris orogeny, presenting an intriguing apparent contradiction with respect to the tectonic setting of Laurentia during the Mesoproterozoic (Daniel et al., 2022a, 2022b; Frost and Frost, 2022).
Simplified geologic map of Proterozoic age provinces for North America and Proterozoic rocks exposed in the western United States (modified from Holland et al., 2020; Daniel et al., 2022a). Dashed black lines indicate the approximate area of the Pinware-Baraboo-Picuris orogen. Abbreviations: AZ—Arizona; CO—Colorado; L—Labrador; NM—New Mexico; P—Picuris Mountains; Q—Quebec; SR—Salt River Canyon.
Simplified geologic map of Proterozoic age provinces for North America and Proterozoic rocks exposed in the western United States (modified from Holland et al., 2020; Daniel et al., 2022a). Dashed black lines indicate the approximate area of the Pinware-Baraboo-Picuris orogen. Abbreviations: AZ—Arizona; CO—Colorado; L—Labrador; NM—New Mexico; P—Picuris Mountains; Q—Quebec; SR—Salt River Canyon.
Simplified map of the southwestern United States showing Proterozoic crustal blocks. Modified after Mahatma (2019) and Jones et al. (2010; cf. Condie, 1986; Bennett and DePaolo, 1987; Karlstrom and Bowring, 1988; Wooden et al., 1988; Wooden and DeWitt, 1991). Colorado Mineral Belt indicated with dashed line.
Simplified map of the southwestern United States showing Proterozoic crustal blocks. Modified after Mahatma (2019) and Jones et al. (2010; cf. Condie, 1986; Bennett and DePaolo, 1987; Karlstrom and Bowring, 1988; Wooden et al., 1988; Wooden and DeWitt, 1991). Colorado Mineral Belt indicated with dashed line.
In Colorado, ca. 1.4 Ga deformation and metamorphism were previously recognized and interpreted as relatively localized events related to pluton emplacement or as reactivation along Paleoproterozoic foliations and shear zones (Fig. 2; Selverstone et al., 2000; Shaw et al., 2001; McCoy et al., 2005; Shaw and Allen, 2007; Allen and Shaw, 2011; Lytle, 2016). This field trip includes visits to exposures in the central Colorado Front Range and in the Wet Mountains (Fig. 2) to examine new evidence for ca. 1.4 Ga metamorphism and deformation, and to discuss how to separate the Mesoproterozoic from older Paleoproterozoic events. We will visit one of the reactivated shear zones in the central Colorado Front Range, and discuss its relationship with map-scale folds in the area. We will visit ca. 1.4 Ga plutons in the central Colorado Front Range and in the Wet Mountains, and discuss possible relationships between these and other plutons in the area, as well as with migmatites. We will discuss the larger-scale nature of deformation across the southern margin of Laurentia, and how the Picuris orogen correlates with the Baraboo and Pinware orogens of the midcontinent and eastern Canada, respectively (Fig. 1; Daniel et al., 2022b). We will also highlight overprinting Cambrian rift structures and mineralized rocks, and Cretaceous–Paleogene structures and mineralized rocks associated with the Laramide orogeny. We will discuss whether and how the Mesoproterozoic Picuris orogeny was different from earlier and later orogenies on Earth and how Proterozoic tectonism may have influenced the location of later structures and mineralization.
The Geological Society of America designated this as a Warren Hamilton field trip, which supports student participation. Warren B. Hamilton (1925–2018) integrated geology and geophysics into large-scale tectonic and planetary-scale syntheses. He was an outside-the-box thinker and always encouraged stimulating and thought-provoking discussion. In that spirit, this field trip guide has “questions for discussion” at each stop. We encourage broad discussion of those questions and others that may arise.
PROTEROZOIC OROGENIC EVENTS IN COLORADO
The Yavapai and Mazatzal Orogenies
The Paleoproterozoic Yavapai and Mazatzal orogenies were first defined in Arizona (Wilson, 1939; Condie, 1982; Karlstrom and Bowring, 1988; Bowring and Karlstrom, 1990) and have subsequently been extended to the NE and across the midcontinent (e.g., Bowring and Karlstrom, 1990; Shaw and Karlstrom, 1999; Jessup et al., 2006; Whitmeyer and Karlstrom, 2007). Paleoproterozoic amphibolite facies metasedimentary and metaigneous rocks in the Colorado Front Range (e.g., Condie, 1982; Gable, 2000; Widmann et al., 2000; Kellogg et al., 2008) were interpreted as juvenile arc terranes with associated basins that amalgamated and accreted to the Wyoming Craton, which was part of Laurentia, between ca. 1.8 Ga and ca. 1.6 Ga (Whitmeyer and Karlstrom, 2007). Alternatively, the Paleoproterozoic rocks of the southwestern United States may have formed largely within a continental arc setting that experienced alternating periods of slab roll and extension to create back-arc basins that were subsequently closed and inverted by advancement of the subduction zone and compression (e.g., Condie, 1982; Jones et al., 2009; Holland et al., 2020). In general, two periods of orogenesis including the ca. 1.75–1.68 Ga Yavapai and ca. 1.65–1.60 Ga Mazatzal orogenies have been recognized across the southwestern United States for more than 35 years, which may or may not have been a continuous period of deformation (Jones and Connelly, 2006; Mahan et al., 2013). The resultant Yavapai crustal province (Fig. 2) is now a NE-trending zone of predominantly juvenile crust that extends from Arizona to Michigan (Holm et al., 2007; Whitmeyer and Karlstrom, 2007), and forms most of the Proterozoic basement of Colorado. Calc-alkaline plutons intruded the Yavapai basement of Colorado between about ca. 1.77 Ga and 1.67 Ga (Anderson and Cullers, 1999). The Mazatzal crustal province possibly extends from Arizona to Quebec and Labrador in Canada (Fig. 1; Whitmeyer and Karlstrom, 2007), and probably includes older Paleoproterozoic crustal material (Holland et al., 2020). Between ca. 1.60 Ga and ca. 1.50 Ga Laurentia underwent a period of tectonic quiescence (Whitmeyer and Karlstrom, 2007; Doe et al., 2012; Duebendorfer et al., 2015; Aronoff, et al., 2016). Interestingly, deformation in the Mazatzal Mountains of Arizona is now known to have occurred between ca. 1.47 Ga and 1.43 Ga (Doe et al., 2012; Doe and Daniel, 2019), the general age of the Picuris orogeny (see below), making the term “Mazatzal orogeny” somewhat controversial (Doe and Daniel, 2019).
The Picuris Orogeny
Mesoproterozoic ca. 1.5–1.45 Ga deposition of siliciclastic sediments in the southwestern United States was first reported in the Picuris Mountains, New Mexico (Jones et al., 2011; Fig. 1) and the upper Salt River Canyon, Arizona (Doe et al., 2012; Fig. 1). Additional rocks of this age are known from the Defiance uplift, Four Peaks, and the northern Mazatzal Mountains of Arizona (Doe et al., 2013, Mako et al., 2015; Doe and Daniel, 2019; Figs. 1, 2). These rocks and the underlying Yavapai and Mazatzal basement across much of central and northern New Mexico experienced greenschist to uppermost amphibolite facies metamorphism and deformation between ca. 1.44 Ga and ca. 1.37 Ga, including the southern Tusas and Taos mountains of New Mexico (Pedrick et al., 1998; Kopera, 2003; Fig. 2). Several areas also experienced emplacement of ca. 1.45–1.40 Ga granitic plutons and preserve contact metamorphic aureoles (Four Peaks, upper Salt River Canyon, southern Picuris Mountains; Fig. 2). The recognition of ca. 1.50–1.35 Ga deposition, regional metamorphism, deformation, and magmatism led to the proposed Picuris orogeny (Daniel and Pyle, 2006; Jones et al., 2011; Daniel et al., 2013) and is supported by subsequent work (Mako et al., 2015; Aronoff et al., 2016; Bollen et al., 2022).
Mesoproterozoic, ca. 1.43–1.36 Ga metamorphism and deformation has also been observed in the Wet Mountains, the Needle Mountains, and the region around the Black Canyon of the Gunnison, southern Colorado, and the Big Thompson Canyon of northern Colorado (Fig. 2; Shaw et al., 2001; McCoy et al., 2005; Jessup et al., 2006; Jones et al., 2010; Shah and Bell, 2012; Mahan et al., 2013; Lytle, 2016). Recent and ongoing research from the central Front Range of Colorado also indicates Mesoproterozoic deformation, not only reactivation along shear zones, but also pervasive folding (Lytle, 2016; Mahatma, 2019; Powell, 2020; Shockley, 2021). This is further discussed below.
The Picuris orogen is part of a larger, trans-Laurentian orogen that include the Baraboo and Pinware orogens of the midcontinent and eastern Canada, respectively (Daniel et al., 2022a, 2022b; Fig. 1). This Pinware-Baraboo-Picuris orogenic belt is attributed to a convergent or accretionary plate margin across the southern margin of Laurentia (present-day coordinates). In eastern Canada, the ca. 1.51–1.46 Ga Pinware orogeny involved convergence and subduction within the proto–Grenville province in Labrador and eastern Quebec (Tucker and Gower, 1994; Gower and Krogh, 2002; Groulier et al., 2020). Similarly, rocks of the midcontinent Baraboo orogeny yielded ca. 1.49–1.47 Ma muscovite 40Ar/39Ar dates and ca. 1.49 Ga U–Th–total Pb dates of neoblastic overgrowths on detrital monazite, and a ca. 1.47 Ga granite porphyry age (Medaris et al., 2021). The ca. 1.48–1.35 Ga Picuris orogeny in northern New Mexico involved convergence and possible collision between Laurentia and juvenile crust along the southern margin of Laurentia (Aronoff et al., 2016). This created shortening and thickening of Paleo- and Mesoproterozoic rocks (Daniel and Pyle, 2006; Daniel et al., 2013, 2022b; Bickford et al., 2015; Aronoff et al., 2016; Holland et al., 2020).
Between ca. 1.48 Ga and 1.36 Ga, granitoids were emplaced in a broad belt spanning from the southwestern United States through the Baltic shield (Fig. 2; Windley, 1993; Karlstrom and Humphreys, 1998; du Bray et al., 2018). These granitoids may be divided into ca. 1.49–1.41 Ga and ca. 1.41–1.34 Ga pulses (Whitmeyer and Karlstrom, 2007). They were interpreted as anorogenic (Anderson and Morrison, 2005; Goodge and Vervoort, 2006). However, episodes of significant deformation and metamorphism are now recognized in parts of New Mexico, Arizona and Colorado that are coeval with pluton emplacement (Nyman et al., 1994; Kirby et al., 1995; McCoy, 2001; Daniel and Pyle, 2006; Jones et al., 2010; Shah and Bell., 2012; Doe et al., 2013; Mahan et al., 2013; Mako et al., 2015; Aronoff et al., 2016; Lytle, 2016, Bollen et al., 2022; Daniel et al., 2022b), consistent with a convergent setting.
From the Grenville Orogeny to the Rio Grande Rift
The ca. 1.1 Ga Pikes Peak batholith in Colorado (Smith et al., 1999; Guitreau et al., 2016) is coeval with the ca. 1.2–1.0 Ga Grenville orogeny along the SE margin of Laurentia (e.g., Whitmeyer and Karlstrom, 2007), but no significant deformation occurred as a result of the Grenville orogeny in Colorado. In the Wet Mountains, Cambrian extension or rifting resulted in REE (rare earth element)-rich alkaline intrusive rocks and veins (Armbrustmacher, 1988). Elsewhere, Proterozoic deformation was overprinted by the Late Cretaceous–Eocene Laramide orogeny (e.g., English et al., 2003; Kellogg et al., 2008) and associated mineralization. Paleogene mineralization was concentrated within the Colorado Mineral Belt (Fig. 2) and may or may not have been controlled by Proterozoic structures such as shear zones (Tweto and Sims, 1963; Caine et al., 2010; Chapin, 2012). The latest deformation was localized extension associated with the Oligocene and younger Rio Grande rift, which trends northward from Socorro, New Mexico, to Leadville, Colorado (Olsen et al., 1987; Chapin and Cather, 1994; Caine and Minor, 2009; Minor et al., 2013).
GEOLOGY OF THE CENTRAL COLORADO FRONT RANGE
Prominent Proterozoic structures of the central Colorado Front Range at the latitude of Denver include at least three generations of folds and the Idaho Springs–Ralston shear zone (IRSZ; Figs. 2, 3). These structures were initially all thought to be Paleoproterozoic. This field trip will focus on the increasing evidence for Mesoproterozoic reactivation of the IRSZ and for Mesoproterozoic pervasive folds, as outlined below. We will also discuss the extent and nature of the IRSZ. The IRSZ is a NE-trending Proterozoic shear zone that was previously interpreted to extend from the Mount Evans batholith, ~10 km SE of Idaho Springs, Colorado, to the eastern margin of the Front Range, ~10 km NNW of Golden, Colorado (e.g., Kellogg et al., 2008). Madison Bzdok (née Lytle; Lytle, 2016) conducted detailed mapping along the shear zone and concluded that it does not extend farther SW than Virginia Canyon Road, on the north side of Idaho Springs (Figs. 2, 3). NE-trending shear zones of the central Colorado Front Range including the IRSZ were initially interpreted as part of a suture zone that formed during ca. 1.8–1.6 Ga accretion of small terranes and island arcs to the Proterozoic Wyoming Craton (Bowring and Karlstrom, 1990; McCoy et al., 2005; Abbott and Cook, 2012). This interpretation was largely based on the presence of tectonic mélange at the St. Louis Lake shear zone, ~40 km NNW of the IRSZ (McCoy, 2001). The Proterozoic shear zones were reactivated between ca. 1.45 Ga and ca. 1.38 Ga (Shaw et al., 2001; McCoy, 2001). These shear zones may have controlled the location of Paleogene mineralization that generally, but not exclusively, occurred along the Colorado Mineral Belt (Tweto and Sims, 1963; McCoy, 2001). The interpretation of the IRSZ as a Paleoproterozoic suture zone and its interpreted control on the location of the Colorado Mineral Belt have been questioned by Caine et al. (2010).
Geologic map of the Idaho Springs–Ralston shear zone (IRSZ) area from Lytle (2016), after Gable (2000; cf. Sims, 1964; Sims and Gable, 1964; Wells, 1967; Wrucke and Wilson, 1967; Sheridan et al., 1972; Bryant et al., 1973; Sheridan and Marsh, 1976; Taylor, 1976; Young, 1991). Automated Mineralogy (AM) samples and map areas as described in detail by Lytle (2016) indicated. Bold numbers in the boxes are sample numbers referred to in the text, followed by U–Pb LA-ICP-MS (laser ablation-inductively coupled plasma–mass spectrometry) ages from Lytle (2016).
Geologic map of the Idaho Springs–Ralston shear zone (IRSZ) area from Lytle (2016), after Gable (2000; cf. Sims, 1964; Sims and Gable, 1964; Wells, 1967; Wrucke and Wilson, 1967; Sheridan et al., 1972; Bryant et al., 1973; Sheridan and Marsh, 1976; Taylor, 1976; Young, 1991). Automated Mineralogy (AM) samples and map areas as described in detail by Lytle (2016) indicated. Bold numbers in the boxes are sample numbers referred to in the text, followed by U–Pb LA-ICP-MS (laser ablation-inductively coupled plasma–mass spectrometry) ages from Lytle (2016).
Based on detailed structural mapping, Lytle (2016; Fig. 3) demonstrated that the IRSZ is not as extensive as previously interpreted (e.g., Kellogg et al., 2008). The mylonite zone ends ~700 m NE of the Edgar Mine (see Stop 3), and to the SW, structures are dominated by folds and the foliation is not consistently parallel to the IRSZ. Additionally, a lack of pinch-outs and offset of major units, as well as similar deformation histories and metamorphic conditions on either side suggest that the IRSZ did not form as a continental suture zone. Along the IRSZ, isoclinal F1 folds are overprinted by F2 folds (Lytle, 2016; Figs. 3, 4). F2 folds NW of the IRSZ have subvertical NE-trending axial planes and plunge shallowly NE. SE of the IRSZ they also plunge shallowly NE, but axial planes dip shallowly ENE. Lytle (2016) developed a model where isoclinal F1 folds are folded by asymmetric NW-side-up meter-scale F2 folds, followed by several km-scale NE-plunging, NW-dipping F3 folds (Fig. 4), including a map-scale synform with a hinge zone in the Idaho Springs area. NW-side-down movement along the IRSZ may have been a result of flexural slip on the steeply dipping NW limb of the NE-plunging, NW-dipping F3 fold.
Cross-section sketch showing proposed deformation and timing of deformation from Lytle (2016). (A) Regional metamorphism at ca. 1.68 Ga (D1). (B) Picuris orogeny at ca. 1.45–1.40 Ga (D2). (C) Picuris orogeny at ca. 1.45–1.40 Ga (D3). IRSZ—Idaho Springs–Ralston shear zone; GGC—Golden Gate Canyon State Park; CCP—Central City Parkway.
Cross-section sketch showing proposed deformation and timing of deformation from Lytle (2016). (A) Regional metamorphism at ca. 1.68 Ga (D1). (B) Picuris orogeny at ca. 1.45–1.40 Ga (D2). (C) Picuris orogeny at ca. 1.45–1.40 Ga (D3). IRSZ—Idaho Springs–Ralston shear zone; GGC—Golden Gate Canyon State Park; CCP—Central City Parkway.
U–Pb laser ablation inductively coupled mass spectrometry monazite dates revealed ca. 1.68 Ga and ca. 1.43 Ga events (Lytle, 2016), both within and adjacent to the IRSZ. Samples 123 and 402C (Fig. 3) are from refolded folds NW of the IRSZ. Sample 305A (Fig. 3) was taken from the moderately north-dipping SE limb of the map-scale NE-plunging, NW-dipping F3 fold (see above), within domain 2. Samples 311A and 312B.2 are from the IRSZ on State Highway (SH) 119 (Stop 1; Fig. 3) and show predominately NW-side-down S-C-C′ fabrics and local minor NW-side-up fabrics. Relationships between microstructures and monazite grains suggest that F1 folds formed at ca. 1.68 Ga, and F2 and F3 folding and associated shearing along the IRSZ occurred at ca. 1.43 Ga. Shear along the IRSZ occurred during both those periods.
In summary, Lytle (2016) concluded that the IRSZ may not be as extensive as previously thought. It was active both at ca. 1.68 Ga and ca. 1.43 Ga. Pervasive folding occurred both NW and SE of the IRSZ at those times. Folding at ca. 1.43 Ga suggests that Mesoproterozoic deformation is not solely restricted to shear zones as previously interpreted (Shaw et al., 2001; McCoy et al., 2005).
Mapping and geochronology by Mahatma (2019), Powell (2020), and Shockley (2021) in the Mount Evans and Montezuma 7.5′ quadrangles SW of the Idaho Springs 7.5′ quadrangle indicated that folding in those areas, too, was partly Paleoproterozoic and partly ca. 1.42 Ga. Therefore, ca. 1.43–1.42 Ga pervasive folding apparently affected a large part of the central Colorado Front Range. This deformation may be related to the Picuris orogeny to the south, and the effects of the Picuris orogeny may, therefore, be much more extensive than previously thought.
In addition, U–Pb detrital geochronology of a quartzite south of Mount Evans (Fig. 5) yielded a ca. 1.43 Ga detrital zircon population (Fig. 5; Mahatma, 2019), indicating that some of the sedimentary rocks were deposited during the Picuris orogeny. The quartzite is exposed in a ~300-m-wide outcrop in a package with amphibolite and calc-silicate gneiss, within biotite-sillimanite gneiss. The ca. 1.43 Ga zircon population is interpreted as detrital, and not metamorphic or hydrothermal, because grains show concentric and oscillatory zoning. In addition, some ca. 1.43 Ga grains show zoned cores with unzoned metamorphic overgrowths that were too narrow to date, but that suggest igneous zircon growth prior to deposition and metamorphism of the quartzite. This is discussed in detail by Mahatma (2019). The quartzite and other stratified rocks were metamorphosed shortly after deposition of the quartzite as indicated by ca. 1.42 Ga and ca. 1.39–1.33 Ga metamorphic monazite in pelitic schist in the area (Mahatma, 2019).
Field photographs of quartzite looking NE. (A) Outcrop picture at sample location. U–Pb detrital zircon data show a ca. 1.43 Ga youngest age population (Mahatma, 2019). (B) Close-up of quartzite, showing bedding.
Field photographs of quartzite looking NE. (A) Outcrop picture at sample location. U–Pb detrital zircon data show a ca. 1.43 Ga youngest age population (Mahatma, 2019). (B) Close-up of quartzite, showing bedding.
GEOLOGY OF THE WET MOUNTAINS
The Wet Mountains comprise a NW-trending, fault-bounded block of primarily Proterozoic rocks (Fig. 6). Metavolcanic and metasedimentary rocks throughout the range display foliations that are shallowly to steeply north-dipping, with lineations commonly trending to the north (Siddoway et al., 2000; Jones et al., 2010; Levine et al., 2013). Metamorphic host rocks are intruded by Paleo- and Mesoproterozoic igneous rocks (Noblett et al., 1987; Bickford et al., 1989; Siddoway et al., 2000; Jones et al., 2010; Levine et al., 2013; Hernández-Montenegro et al., 2019). Paleoproterozoic intrusive bodies include the foliated tonalite and granodiorite of the ca. 1705 Ma Twin Mountain and Crampton Mountain plutons and the weakly foliated to undeformed granodiorite of the ca. 1663 Ma Garell Peak pluton (Fig. 6; Bickford et al., 1989). Mesoproterozoic intrusive rocks include the foliated ca. 1442–1439 Ma Oak Creek pluton (Fig. 6; Stop 9), which ranges from quartz monzonite and monzogranite to leucogranite (Bickford et al., 1989; Siddoway et al., 2000; Hernández-Montenegro et al., 2019); the unfoliated ca. 1460 Ma West McCoy Gulch leucogranite pluton (Fig. 6; Cullers et al., 1993); and the ca. 1371–1362 Ma San Isabel pluton (Fig. 6; Bickford et al., 1989), which is a monzogranite to syenogranite that is variably deformed at exposed margins and generally undeformed within the main body of the pluton (Cullers et al., 1992; Jones et al., 2010).
The central and southern Wet Mountains host an extensive network of Mesoproterozoic sill and dike intrusions (Fig. 6; Stop 8; Jones et al., 2010; Levine et al., 2013). The southern Wet Mountains also contain extensive exposures of migmatite gneiss (Fig. 6, Stop 8; Levine et al., 2013). In the Greenhorn Mountain area south and west of the main body of the San Isabel pluton, host rock-intrusion relationships are commonly unclear, because of the extent of migmatization of host rock gneisses and because of the extensive network of centimeter- to meter-scale felsic intrusions (Jones et al., 2010; Levine et al., 2013). The San Isabel pluton contains magmatic epidote (Cullers et al., 1992), suggesting midcrustal emplacement depths. Migmatitic gneisses at the contact with the Oak Creek pluton record peak metamorphic conditions of ~750 °C and ~7 kbar (Hernández-Montenegro et al., 2019).
Granitic sills in the southern Wet Mountains yield U–Pb zircon crystallization ages between 1435 and 1390 Ma (Jones et al., 2010), and are interpreted as syntectonic with respect to an episode of regional NW-directed shortening (Nyman et al., 1994). Metamorphic cooling ages demonstrate that the central and southern Wet Mountains reached temperatures above 500 °C at ca. 1.4 Ga (Shaw et al., 2005). Preliminary Lu–Hf garnet age data from migmatites of the central and southern Wet Mountains suggest that Mesoproterozoic anatexis may have occurred between ca. 1.49 and 1.47 Ga (Aronoff, 2016).
In the northern Wet Mountains, Mesoproterozoic deformation overprinted two generations of Paleoproterozoic fabrics, and strain was concentrated in shear zones (Siddoway et al., 2000). In migmatitic rocks of the southern Wet Mountains (and beyond), extensive partial melt generation may have led to the formation of melt networks that accommodated mid- to lower crustal flow in response to regional compressive stresses (Jones et al., 2010; Levine et al., 2013; Searle, 2013; Levine and Rahl, 2021). Together with the age constraints described above, these structural data suggest that Mesoproterozoic partial melt generation and mid- to lower crustal flow in the Wet Mountains were contemporaneous with the tectonic activity that defined the Picuris orogeny in New Mexico and Arizona (Daniel and Pyle, 2006; Doe et al., 2012; Daniel et al., 2013; Mako et al., 2015; Aronoff et al., 2016; Doe and Daniel, 2019; Bollen et al., 2022).
The northern part of the Wet Mountains is well known for thorium and other REE mineralization associated with Cambrian–Ordovician alkaline intrusions, including the McClure Mountain, Gem Park, and Democrat Creek complexes (Fig. 6; Armbrustmacher, 1988). These complexes are sequentially cross-cut by lamprophyre, syenite, and carbonatite dikes, and mineralized quartz-barite-thorite veins (Armbrustmacher, 1988; Magnin et al., 2021). REE mineralization occurs predominately within carbonatite dikes, quartz-barite-thorite veins, and hydrothermally altered red syenite dikes (Armbrustmacher, 1988). Alkaline magmatic rocks in the Wet Mountains may have been derived from a mantle melt, likely related to failed intracontinental rifting during the Cambrian–Ordovician (Olson et al., 1977; Larson et al., 1985; McMillan and McLemore, 2004; Magnin et al., 2021).
FIELD-TRIP DESCRIPTION AND STOPS
On Day 1, we will first discuss how the Picuris orogeny reactivated or overprinted earlier structures, and perhaps controlled the location of structures associated with Cambrian rifting, the Cretaceous–Paleogene Laramide orogeny and subsequent Rio Grande rift, and associated mineralization. Stops 1–7 will give an overview of the geology of the central Colorado Front Range, including rocks and deformation associated with the Paleoproterozoic Yavapai/Mazatzal orogenies and the Mesoproterozoic Picuris orogeny, as well as some of the Cretaceous–Paleogene structures and mineralization associated with the Laramide orogeny. On Day 2, Stops 8 and 9 will focus on deformation and metamorphism associated with the Paleoproterozoic Yavapai/Mazatzal orogenies and the Mesoproterozoic Picuris orogeny. We will also discuss Cambrian REE-bearing magmatic rocks in the northern Wet Mountains as a result of Cambrian rifting.
DAY 1. Central Colorado Front Range—Introduction to the Proterozoic Yavapai/Mazatzal and Picuris Orogenies
Leave Colorado Conference Center (700 14th Street, Denver, Colorado 80202) at 7:00 a.m.
Drive 32 miles to the outcrop via I-70, U.S.-6 and SH 119 (50 min). Park on the SW side of the road at Stop 1, across the road from the Coyote Motel and small strip mall across the road. Restrooms are available inside the mall. Be very aware of traffic.
Stop 1. Idaho Springs–Ralston Shear Zone (8:00–9:30 a.m.; Includes Introductory Talk)
Location: 39°47′11″ N, 105°27′57″ W; 2363 m elevation; Blackhawk 7.5′ quadrangle (coordinates at parking location along the SW side of SH 119; outcrops are across the road to the NW and the SE of the strip mall).
The rocks are primarily Paleoproterozoic felsic gneiss (Tweto, 1987, Taylor, 1975). They can be generally described as fine- to medium-grained, white to light-gray or tan, moderately well-foliated microcline-quartz-plagioclase-biotite gneiss, with various amounts of biotite and local occurrences of magnetite, hornblende, or garnet (Taylor, 1975; cf. Widmann et al., 2000), and have been interpreted as being volcanic in origin (Tweto, 1987), but metamorphism renders the origin difficult to interpret. Locally, lenses of amphibolite and calc-silicate gneiss occur in the felsic gneiss (Taylor, 1975). On this outcrop, the rocks are a mix of schist, gneiss, and pegmatite. All are deformed to various amounts. The rocks have a subvertical NE-trending foliation (Fig. 7) and steep lineation. S-C-C′ fabrics indicate predominately SE-side-up and locally NW-side-up (Fig. 7). One biotite-muscovite schist yielded ca. 1.67 Ga, ca. 1.63 Ga, and ca. 1.48 Ga in situ U–Pb LA-ICP-MS (laser ablation-inductively coupled plasma–mass spectrometry) monazite dates, and another yielded ca. 1.62 Ga and ca. 1.44 Ga dates (Lytle, 2016). The dates are consistent with ages of deformation along the IRSZ, and of folds away from the shear zone (Lytle, 2016). The relationships between dates and particular shear senses are not clear.
Idaho Springs–Ralston shear zone outcrop images in various Paleoproterozoic rocks. (A) SE-side-up shear bands in schist with quartz-rich pseudomorphs of unknown porphyroblast. (B) Felsic gneiss showing both SE- and NW-side-up S-C-C′ fabrics. (C) Pelitic gneiss with both SE- and NW-side-up S-C-C′ fabrics. (D) Sheared pegmatite. (E) Sketch of S-C and C-C′ fabrics to show general geometry. C is the main foliation. S is deflected by C. C′ is also called ‘shear bands’ or Riedel shears. C′ deflects C and sometimes also S; all three can occur together.
Idaho Springs–Ralston shear zone outcrop images in various Paleoproterozoic rocks. (A) SE-side-up shear bands in schist with quartz-rich pseudomorphs of unknown porphyroblast. (B) Felsic gneiss showing both SE- and NW-side-up S-C-C′ fabrics. (C) Pelitic gneiss with both SE- and NW-side-up S-C-C′ fabrics. (D) Sheared pegmatite. (E) Sketch of S-C and C-C′ fabrics to show general geometry. C is the main foliation. S is deflected by C. C′ is also called ‘shear bands’ or Riedel shears. C′ deflects C and sometimes also S; all three can occur together.
Questions for discussion: What are the shear direction and shear sense? Was this shear zone a suture zone or a smaller-scale shear zone? How is it related to folding in the area? How would one find out?
Drive for 5 miles (8 min.) back toward the SE on SH 119. After passing Douglas Mountain Drive on the left (east) and a small inaccessible tunnel to the right (west), pull off on the left (east) side of the road, before the intersection with U.S.-6.
Stop 2. Mesoproterozoic Pegmatite along Top-to-the-South Shear (9:45–10:15 a.m.)
Location: 39°44′47″ N, 105°23′51″ W; 2108 m elevation; Squaw Pass 7.5′ quadrangle (coordinates at outcrop; park ~50 m to the north along the east side of the road in the pullout and walk south).
Southeast of the IRSZ there are tens of meter-scale shear zones along the C planes of several meter-scale S-C fabrics, displaying top-to-the-S ductile-brittle shear along shallowly N-dipping shear planes. On this outcrop, multiple sheared biotite-bearing pegmatites exist along shallowly northerly dipping top-to-the-S shear planes (Figs. 8A–8D) in Paleoproterozoic (?) heterogeneous biotite gneiss and felsic gneiss (cf. Stop 1; Taylor, 1975; cf. Widmann et al., 2000). The gneissic rocks show shallowly to moderately northerly dipping foliations and shallowly northerly to northwesterly plunging (intersection?) lineations. The pegmatites are deformed internally and boudinaged by the top-to-the-S shear. Zircon in a pegmatite near the traffic light yielded a preliminary and imprecise LA-ICP-MS zircon crystallization age of ca. 1.44 Ga, with inheritance of primarily ca. 1.77 Ga grains and others as old as ca. 1.91 Ga (Kuiper and Chang, 2021, personal observ.). While the ca. 1.44 Ga age is imprecise, it is consistent with the Picuris orogeny and inconsistent with the earlier Yavapai-Mazatzal and later Grenville orogenies. The S-directed shear zone here and other top-to-the-S structures in the area (e.g., Stop 4) are therefore interpreted as associated with the Picuris orogeny. They may represent late flexural slip on the moderately N-dipping SE limb of a NW-dipping syncline (cf. Fig. 4C) during folding associated with the Picuris orogeny. Alternatively, they may have formed as a result of separate S-directed transport after folding, late during the Picuris orogeny.
Outcrop pictures, showing ca. 1.44 Ga pegmatite along top-to-the-S shear. (A) Shear indicated by white dashed line. (B) Part of the pegmatite, boudinaged by the top-to-the-S shear. (C) Google Earth version of A. (D) Close-up of pegmatite. Contacts with the host felsic gneiss are at the hammer head and just above the handle. Hammer for scale in B and D is 28 cm long.
Outcrop pictures, showing ca. 1.44 Ga pegmatite along top-to-the-S shear. (A) Shear indicated by white dashed line. (B) Part of the pegmatite, boudinaged by the top-to-the-S shear. (C) Google Earth version of A. (D) Close-up of pegmatite. Contacts with the host felsic gneiss are at the hammer head and just above the handle. Hammer for scale in B and D is 28 cm long.
Questions for discussion: What is the significance of these shears? Are they related to folding in the area, and to the IRSZ on the moderately N-dipping limb that forms the SE part of the large-scale NW-dipping, NE-plunging synform? How would one find out?
Drive 8.5 miles (15 min.) via U.S.-6, I-70 to Idaho Springs. Take exit 240, which brings you downtown. Go north on 13th Avenue, west on Colorado Boulevard, and north on 8th Avenue to the Edgar Mine. This is a short but steep dirt road. Drive all the way up to the entrance of the mine, if possible, or walk up if needed (depending on road conditions). There is a restroom on-site at the Edgar Mine.
Stop 3. Edgar Experimental Mine (10.30 a.m.–12.00 p.m.)
Location: 39°44′50″ N, 105°31′31″ W; 2399 m elevation; Idaho Springs 7.5′ quadrangle. (For further info and visits, see https://mining.mines.edu/edgar-experimental-mine/.)
The Edgar Mine, the Colorado School of Mines Experimental Mine (Fig. 9), stems from the “Rush to the Rockies” mining period. In the 1870s, it produced high-grade silver, gold, lead, and copper. It is now part of the Colorado School of Mines and an educational laboratory for those learning to find, develop, and process natural resources. In addition to educating Colorado School of Mines mining engineering students, it provides educational tours for the public and school groups. The Edgar Mine has nearly two miles of underground passageways. The Miami tunnel, one of the two main tunnels, was started in the 1890s and was driven to access the Edgar’s silver vein. The back of the Edgar Tunnel (1200 ft [366 m] from the Miami portal) has long since caved. The Army tunnel, the other main tunnel, was drilled less than 20 years ago. The mine has a constant temperature of 54 °F or 12 °C.
The Edgar Mine, the Colorado School of Mines Experimental Mine. (A) Geology graduate students in front of the entrance. (B) The USGS underground classroom. (C) NW-side-up fold in Paleoproterozoic migmatitic biotite gneiss. (D) Pyrite vein. (E) Location map of the Edgar Mine (red lines). Quartz-plagioclase gneiss and quartz-plagioclase-biotite gneiss in blue and quartz-biotite-hornblende gneiss and biotite-microcline-pegmatite in brown (from C. Dattel, Edgar Mine, 2022, personal commun.).
The Edgar Mine, the Colorado School of Mines Experimental Mine. (A) Geology graduate students in front of the entrance. (B) The USGS underground classroom. (C) NW-side-up fold in Paleoproterozoic migmatitic biotite gneiss. (D) Pyrite vein. (E) Location map of the Edgar Mine (red lines). Quartz-plagioclase gneiss and quartz-plagioclase-biotite gneiss in blue and quartz-biotite-hornblende gneiss and biotite-microcline-pegmatite in brown (from C. Dattel, Edgar Mine, 2022, personal commun.).
The oldest rocks include Paleoproterozoic migmatitic biotite gneiss and schist (Fig. 9C), with local sillimanite, magnetite, garnet or cordierite, layers or pods of granite, felsic gneiss, hornblende gneiss, and calc-silicate gneiss (Widmann et al., 2000). The Edgar Mine is located along-strike with (and is arguably part of) the IRSZ. However, the mylonite zone that is characteristic of the IRSZ ends less than 700 m to the NE, where the last outcrops can be seen on the Virginia Canyon “Oh-My-God” Road (nicknamed for its twists and turns, which are much less scary than in the past). In Idaho Springs, a zone of NE-trending foliation exists associated with predominately NW-side-up asymmetric folds that formed at upper amphibolite facies, but no mylonite zone has been identified. The NW-side-up asymmetry of the folds may be consistent with one of the movements on the IRSZ. The termination of the IRSZ led to the conclusion that it is not a suture zone, but an effect of regional folding (Lytle, 2016; Fig. 4).
Cretaceous–Paleogene quartz veins and mineralization (Fig. 9D) are close to parallel to the Proterozoic foliation in and around the Edgar Mine, but on careful inspection, they can be seen to crosscut at very low angle, or locally at high angle (e.g., a pyrite-galena–bearing vein in the Orica drift of the mine).
Questions for discussion: Is Cretaceous–Paleogene mineralization, magmatism, faulting, and veining controlled by Proterozoic structures? What, if any, evidence is there for structural control on ore deposits at any scale in the area?
Lunch (12:00–1:00 p.m.)
Drive 2.5 miles (7 min.) east on Colorado Boulevard, across the roundabout on the east side of town, and across I-70. Turn right, then left on East Idaho Springs Road (and don’t go straight onto I-70 here!). After ~1000 ft, pull off at the driveway entrance without obstructing the driveway. If there are too many vehicles, park another ~1000 ft down the road and walk back up.
Stop 4. Mega Sigma-Clast and Other Shear Structures (1:15–1:30 p.m.)
Location: 39°44′39″ N, 105°29′16″ W; 2281 m elevation; Squaw Pass 7.5′ quadrangle.
Look E on the cliff across the road to find an ~15-m-high sigma-clast (Fig. 10A) and other large-scale shear structures (Fig. 10B) in Paleoproterozoic (?) felsic gneiss (cf. Stop 1; Sheridan and Marsh, 1976). The shear sense is top-to-the-south, as in Stop 2. Similar structures exist throughout this area SE of the IRSZ, and possibly beyond, and provide something to look at when stuck in ski or other mountain traffic.
Mega-shear structures in Paleoproterozoic (?) felsic gneiss indicating top-to-the-south movement. (A) An ~15-m-high sigma-clast in felsic gneiss between arrows indicating shear sense. (B) Part of an ~15-m-high sigma-clast between arrows indicating shear sense.
Mega-shear structures in Paleoproterozoic (?) felsic gneiss indicating top-to-the-south movement. (A) An ~15-m-high sigma-clast in felsic gneiss between arrows indicating shear sense. (B) Part of an ~15-m-high sigma-clast between arrows indicating shear sense.
Along the road at the stop itself, there is another sheared pegmatite. Unlike the one at Stop 2, this one is reactivated and altered during later brittle movement, presumably associated with the Laramide orogeny.
Questions for discussion: Continue discussion of Stop 2. Also, what causes such large-scale structures to form? Is it because of rheology contrast, or because of brittle shear localization, or something else? Where else have you seen such large-scale shear structures?
Drive 7 miles (12 min.) to the next stop. Backtrack toward Colorado Boulevard, across the highway and then immediately on the highway (I-70 westbound) from the roundabout. Take the next exit (240) to the SW on SH 103 (Squaw Pass Road). Park on the large pulloff at the intersection with W Chicago Creek Road. The outcrops are along the cliffs to the NE. Watch out for unstable slopes and rockfall.
Stop 5. Sheared Mount Evans Batholith (1:45–2:30 p.m.)
Location: 39°41′38″ N, 105°37′02″ W; 2689 m elevation; Idaho Springs 7.5′ quadrangle.
The Mount Evans batholith can be described as a weakly to moderately foliated, medium- to coarse-grained granodiorite. It is composed of plagioclase, microcline, quartz, biotite, and accessory hornblende, magnetite, and sphene (Widmann et al., 2000). U–Pb zircon data yielded ages of 1422 ± 2 Ma and 1445 ± 8 Ma (Powell, 2020). While the 1424 ± 6 Ma (cf. Kellogg et al., 2008) Silver Plume batholith to the NW is ferroan, strongly peraluminous, and alkali-calcic, suggesting an extensional setting, the Mount Evans batholith is slightly magnesian, metaluminous, and calc-alkalic, which is more typical for a subduction zone setting (Frost and Frost, 2011; du Bray et al., 2018; Powell, 2020).
In this outcrop, the batholith and a large xenolith of biotite gneiss are deformed by a steeply WNW-dipping local shear zone with NW-side-up movement (Fig. 11). It has been interpreted as the IRSZ in the past (e.g., Kellogg et al., 2008). However, it is too northerly trending for that, and may be a local shear zone. Similarly, foliations and shear zones within ~500 m to the SW, S, and SE, and on a separate outcrop ~4 km to the SW along W Chicago Creek Road, have various orientations. Shear zones are only local here and their significance is not clear and probably minor.
Steeply WNW-dipping local shear zone in ca. 1.44 Ga Mount Evans batholith. (A) Outcrop photo. (B) West-side-up S-C fabric.
Steeply WNW-dipping local shear zone in ca. 1.44 Ga Mount Evans batholith. (A) Outcrop photo. (B) West-side-up S-C fabric.
Questions for discussion: Why does the IRSZ exist to the NE, but not here? Does it reactivate earlier NE-trending structures? How does it relate to other NE-trending shear zones of Colorado, some of which are interpreted as suture zones? What is the significance of the smaller shear zones at and around this stop? What is the significance of the contrasting compositions of the Mount Evans and Silver Plume batholiths?
Drive 0.9 miles (2 min.) to the NE on SH 103, and park on the right (SE) side of the road by a small park with a bridge crossing over to a meadow. Outcrops are across the road.
Stop 6. Weakly Foliated Mount Evans Batholith (2:40–3:10 p.m.)
Location: 39°42′13″ N, 105°36′29″ W; 2612 m elevation; Idaho Springs 7.5′ quadrangle.
Here, the Mount Evans batholith (Fig. 12A) looks much less deformed than the previous outcrop, and has a weak foliation that is more typical for the Mount Evans batholith. However, the rocks contain much more biotite than most of the Mount Evans batholith. While the Mount Evans batholith contains dioritic parts (Aleinikoff, 1993), it is possible that this is actually part of the ca. 1.7 Ga Routt plutonic suite instead.
Weakly foliated ca. 1.44 Ga Mount Evans batholith. (A) Outcrop photo, looking SW. (B) Poles to foliation in the Mount Evans batholith in the Mount Evans 7.5′ quadrangle to the south (from Powell, 2020; used with permission).
Weakly foliated ca. 1.44 Ga Mount Evans batholith. (A) Outcrop photo, looking SW. (B) Poles to foliation in the Mount Evans batholith in the Mount Evans 7.5′ quadrangle to the south (from Powell, 2020; used with permission).
The Mount Evans batholith in general was initially interpreted as having primarily flow foliations (Aleinikoff et al., 1993), but Powell (2020) noted that most foliations dip moderately to the NW (Fig. 12B) and are probably tectonic. This may imply that the shortening direction of the Picuris orogeny was NW-directed, while the top-to-the-south ductile-brittle structures of Stops 3 and 4 imply that it was S-directed. The reason for this variation is unclear. Other ca. 1.4 Ga plutons and other deformed rocks also show evidence for NW-directed shortening in Colorado (Gonzales et al., 1996; Jones et al., 2010; Shah and Bell, 2012), New Mexico (Grambling and Codding, 1982), and Arizona (Doe and Daniel, 2019), while in the Picuris mountains of New Mexico (Daniel et al., 2013) and the Montezuma mining district of Colorado (Shockley, 2021), the shortening direction is north-directed.
Questions for discussion: What was or were the shortening direction(s) during the Picuris orogeny? Did it change through time? How does it fit with shortening in New Mexico and Arizona and with shear zone reactivation? How does one distinguish rocks from the ca. 1.7 Ga Routt plutonic suite from those of the ca. 1.4 Ga Berthoud suite (that includes the Mount Evans and Silver Plume batholiths) and/or the equally heterogeneous ca. 1.1 Ga Pikes Peak batholith?
Drive 1.8 miles (3 min.) to the NE on SH 103, and park on the right (SE) side of the road by or just past four colorful cabins. Outcrops are across the road.
Stop 7. Folded Migmatitic Biotite Gneiss (3:20–3:50 p.m.)
Location: 39°42′50″ N, 105°34′48″ W; 2483 m elevation; Idaho Springs 7.5′ quadrangle.
Paleoproterozoic migmatitic biotite gneiss with local sillimanite as in the Edgar Mine (Stop 3; Widmann et al., 2000) is isoclinally folded, and subsequently refolded by moderately NE-plunging open to close (NW-side-up?) F2 folds (Fig. 13). The lineation is steep and pre-or syn- F2 folding. This location is along strike with the IRSZ of Stop 1, but here (and in the Edgar Mine to the NE) it has died out.
(A) F1 folds in Paleoproterozoic migmatitic biotite gneiss looking down to the NE along an F2 fold hinge. (B) F1 folds in migmatitic biotite gneiss looking down to the NW.
(A) F1 folds in Paleoproterozoic migmatitic biotite gneiss looking down to the NE along an F2 fold hinge. (B) F1 folds in migmatitic biotite gneiss looking down to the NW.
Questions for discussion: Why does the IRSZ die out toward the SW, where structures are dominated by regional folds and no NE-trending mylonite zone exists? Is the shear zone related to folding (flexural flow)?
Drive 150 miles to Walsenburg: Continue north on SH 103 and get on I-70 E. Follow I-70 E and CO-470 E to US-85 S/S Santa Fe Drive. Take exit 17 from CO-470 E. Get on I-25 S from CO-105 E and W Wolfensberger Road. Follow I-25 S to I-25BL S. Take exit 52 from I-25 S toward Walsenburg. The Best Western Rambler is immediately after the exit. 2.5–3.0 h.
Stay overnight in Walsenburg, Colorado.
DAY 2. Wet Mountains—Proterozoic Deformation and Metamorphism and Cambrian Rifting
Depart from Best Western Rambler in Walsenburg at 7:15 a.m.
Drive ~38 miles, 55–75 min. Follow CO-69 N west for 24 miles and then turn right (north) onto County Road 630. After 9.3 miles, turn right (east) onto Turkey Creek Road. Park after 2.4 miles.
Stop 8. Garnet Biotite Migmatite Gneiss
Parking location: 37°53′30.73″N, 105°6′23.85″W; 2670 m elevation; San Isabel 7.5′ quadrangle. Walk ~0.9 miles (heading up in elevation) along the Cisneros Trail.
This section of the Cisneros Trail contains semi-continuous outcrop exposure of biotite migmatite gneiss, garnet biotite ± sillimanite migmatite gneiss, and amphibolite (Fig. 14). Melt in these migmatites may have formed by the biotite dehydration melting reaction: biotite + quartz + sillimanite + plagioclase = melt + K-feldspar + garnet (Levine et al., 2013). Recent work shows that partial melting was focused along subgrain boundaries in the southern Wet Mountains migmatites and that subgrain melt may have contributed to a weak middle and lower crust (Levine and Rahl, 2021).
Outcrop photos showing typical textures of migmatites in the Cisneros Trail field area of the southern Wet Mountains, including quartz-feldspar migmatite, biotite migmatite, and garnet-biotite migmatite. Note isoclinal, ptygmatic, and refolded folds, shear bands, and anatectic garnet.
Outcrop photos showing typical textures of migmatites in the Cisneros Trail field area of the southern Wet Mountains, including quartz-feldspar migmatite, biotite migmatite, and garnet-biotite migmatite. Note isoclinal, ptygmatic, and refolded folds, shear bands, and anatectic garnet.
Hornblende 40Ar/39Ar cooling ages in the southern Wet Mountains demonstrate that the area reached regional metamorphic temperatures above 500 °C at ca. 1.4 Ga (Shaw et al., 2005). Preliminary Lu–Hf garnet age data from embayed and skeletal garnet crystals within a garnet-biotite-sillimanite-migmatite gneiss sample from this stop location indicate anatectic garnet growth at ca. 1.49 Ga (Aronoff, 2016), which is older than U–Pb zircon crystallization ages of 1371–1362 Ma for the San Isabel pluton (Bickford et al., 1989; Cullers et al., 1992) and of 1435 ± 4 Ma and 1390 ± 10 Ma U–Pb for granitic sills of the southern Wet Mountains (Jones et al., 2010). Because host rock-intrusion relationships are cryptic or obscured in the southern Wet Mountains, geochronology data aid in gaining a more comprehensive view of the interplay between metamorphism and pluton emplacement than structural and field relationships alone. This stop is ideally suited for examining the interplay between metamorphism and igneous activity. The preliminary constraints on the timing of metamorphism at this stop suggest that migmatization and partial melt retention versus removal may have set the stage for younger episodes of igneous emplacement.
We will walk the first mile of the trail. The Cisneros Trail is an ~10-mile (~16-km) path from this stop to the San Isabel recreation area. The trail continues past this field-trip stop through exposures of Mesoproterozoic granitic sill and dike intrusions (Jones et al., 2010; Levine et al., 2013). The northern section of the trail (beyond this stop’s location) passes through the main body of the San Isabel pluton, but exposure is poor. Stop 8 will highlight migmatite textures and mineralogy (Fig. 14). The migmatite outcrops of the Cisneros Trail are characteristic of the central and southern Wet Mountains.
Questions for discussion: What can we interpret from migmatite textures in outcrop about the structural or temporal relationship between partial melt formation and felsic intrusions in the southern Wet Mountains? What might different structural interpretations imply about tectonic models of southwestern North America in the Proterozoic?
Drive (48 miles, ~1 h, 15 min.) from Cisneros trailhead to downtown Westcliffe. Head west on Turkey Creek Road until County Road 630, go south on County Road 630 until CO-69 N, and then drive north to Westcliffe.
Lunch in Westcliffe
Drive from Westcliffe 18 miles northeast along Oak Creek Grade (County Road 143; 20 miles, ~40 min.) to East Bear Trailhead pull-out.
Stop 9. Oak Creek Pluton
Parking location: 38° 18′ 19.764″ N, 105° 15′ 16.83″ W; 2255 m elevation; Curley Peak 7.5′ quadrangle.
We will examine a section of the Oak Creek pluton exposed along the road (Fig. 15). The Oak Creek pluton is a foliated Mesoproterozoic pluton with a U–Pb zircon crystallization age of 1442–1439 Ma (Bickford et al., 1989; Cullers et al., 1993). Its fabrics are generally concordant with host rock fabrics (Cullers et al., 1993; Siddoway et al., 2000). Adjacent host rocks include felsic gneiss, amphibolite, and migmatite gneiss (Siddoway et al., 2000; Jones et al., 2010; Levine et al., 2013; Hernández-Montenegro et al., 2019).
Outcrop photo of the foliated Oak Creek pluton (left) and a photo of Stop 9 (right), with roadside exposure of the Oak Creek pluton.
Outcrop photo of the foliated Oak Creek pluton (left) and a photo of Stop 9 (right), with roadside exposure of the Oak Creek pluton.
The Oak Creek pluton ranges in composition from quartz-monzonite to monzogranite to leucogranite (Bickford et al., 1989; Cullers et al., 1993) and is also ferroan in composition (Cullers et al., 1993; Frost et al., 2001; Hernández-Montenegro et al., 2019). This composition is similar to that of the Silver Plume batholith, but not the Mount Evans batholith as described above. Several thermobarometers indicate that the Oak Creek pluton was emplaced at 3–4 kbar, with solidus temperatures of ~650 ºC and liquidus temperatures of ~880 ºC (Cullers et al., 1993). West- to northwest-striking foliations in the Oak Creek pluton are interpreted as magmatic (Bickford et al., 1989), and their concordance with host rock fabrics is interpreted to reflect syntectonic emplacement, perhaps as a series of sill injections, in alignment with a regional stress field (Siddoway et al., 2000).
Migmatites adjacent to the Oak Creek pluton have calculated peak temperatures and pressures of 743 ºC and 7.6 kbar (Hernández-Montenegro et al., 2019). These migmatites preserve a complex history of dehydration partial melting, back-reaction, melt recrystallization, and melt removal (Hernández-Montenegro et al., 2019; cf. Kriegsman, 2001). Preliminary Lu–Hf garnet age data from the same field location indicate that partial melting may have occurred at ca. 1.47 Ga (Aronoff, 2016). The Oak Creek pluton crystallization ages are younger than this preliminary estimate of the timing of migmatization of the host rock. Like Stop 8, metamorphic data suggest that regional metamorphism may have already been under way prior to pluton emplacement. Together, structural, metamorphic, and geochronologic data for the Oak Creek pluton potentially challenge the interpretation that ferroan magmas are markers of extensional tectonic settings.
Questions for discussion: How can we reconcile the structural evidence of syntectonic emplacement of this Mesoproterozoic pluton with the ferroan geochemical character of the pluton?
Continue northeast on from Oak Creek Grade, and follow US-50 E and CO-115 N. Get on I-25 N in Colorado Springs to return to Denver.
DISCUSSION
We conclude this field guide with a discussion about how the Picuris orogen and the regional tectonic questions associated with its creation may fit into the evolution of global tectonics during the Precambrian. Much of the current research on North American Mesoproterozoic tectonics is predicated on the assumption that active plate margins operated in much the same manner as Phanerozoic active margins (Condie, 1982; Bowring and Karlstrom, 1990; Karlstrom et al., 2001; Whitmeyer and Karlstrom, 2007; Jessup et al., 2006; Jones et al., 2009; Daniel et al., 2013, 2022a, 2022b; Holland et al., 2020). However, research into the secular evolution of Earth systems is currently divided on the question of whether the Mesoproterozoic was similar to or different from the Phanerozoic. Competing ideas are discussed below.
Recent work supports the interpretation that global tectonics began to operate on Earth by ca. 3 Ga (e.g., Hawkesworth et al., 2016, 2017; Palin et al., 2020). Between 3.0 and 0.75 Ga, tectonics evolved from systems featuring warm, shallow subduction to the cold, deep subduction systems of the Phanerozoic (Hawkesworth et al., 2016; Palin et al., 2020). Within this evolutionary framework, the ca. 1.45–1.35 Ga Picuris orogeny (Daniel et al., 2013, 2022a, 2022b) occurred during an intermediate stage of Earth tectonics evolution. This “middle age” of plate tectonics was marked by thick continental lithosphere and a warm mantle, but did not yet feature cold subducting slabs (Hawkesworth et al., 2016).
Alternatively, Mesoproterozoic orogenic and other Earth processes may have been different from those in the Paleoproterozoic, and Neoproterozoic/Phanerozoic (e.g., Bradley et al., 2011; Brown et al., 2020; Spencer et al., 2021; Stern, 2021; Liu et al., 2022). Bradley et al. (2011) compiled various data sets and demonstrated various minima, including the number of passive margins, detrital zircons, greenstone-belt collisions, eclogites, granulites, carbonatites, and orogenic gold. Perhaps the quiescence explains the “Boring Billion” of Holland (2006), which was based on a period of little variation in atmospheric oxygen levels. Brown et al. (2020) observed that T/P ratios of metamorphic rocks were higher in the Mesoproterozoic than at any other time in the Earth’s history. High T/P conditions were accompanied by high 176Hf/177Hf ratios and low δ18O values in zircon, indicating more juvenile crust in the Mesoproterozoic (Brown et al., 2020). A high volume of massif-type anorthosites also indicates high T/P conditions and zircon Eu/Eu* ratios indicate thin crust (Spencer et al., 2021). Low T/P ratios in metamorphic rocks are characteristic of plate boundaries (Liu et al., 2022), suggesting that there may have been fewer active plate boundaries in the Mesoproterozoic than before and after that time. In addition to such a proposed “slowdown” in plate tectonic processes, both “stagnant lid tectonics” and “hot, thin, and low orogens” have been proposed for the Mesoproterozoic (Stern, 2021; Spencer et al., 2021). Low T/P rocks are especially characteristic of blueschist and eclogite facies rocks that are associated with subduction. While present in the Paleoproterozoic, these became especially abundant in the Neoproterozoic when there was an apparent transition to modern plate tectonic processes (Brown et al., 2020; Liu et al., 2022).
Paleogeographic reconstructions suggest that the Paleoproterozoic supercontinent may never have broken up fully in the early to mid-Mesoproterozoic and that the transition to Rodinia was primarily during the late Mesoproterozoic (e.g., Pisarevsky et al., 2014, Martin et al. 2020). Along the SE margin of Laurentia, there may have been subduction during most of the Mesoproterozoic, and perhaps slab rollback, associated juvenile crust formation, and/or accretion of juvenile crust (cf. Bickford et al., 2015; Brown et al., 2020; Liu et al., 2022; Daniel et al., 2022a).
At present, the tectonic nature, areal extent, and architecture of the Picuris orogen, and how it relates to a global-scale tectonic regime remain speculative. We encourage further detailed research across the southwestern United States to better separate regional deformation and metamorphism associated with the Picuris orogeny from that associated with older, Paleoproterozoic orogenies. Additional research is also needed to better characterize the nature and origin of ferroan magmatism across the southwest and midcontinent regions. The relationship between the interpreted convergent tectonic setting for the Picuris-Baraboo-Pinware orogen (Daniel et al., 2022a), and the possibly extensional ferroan magmatic belt across southern Laurentia, in part at the same time and location remains enigmatic.
ACKNOWLEDGMENTS
We thank current and past Colorado School of Mines graduate students Ben Magnin, Asha Mahatma, Logan Powell, and Dustin Shockley for discussion and/or contribution. We benefited much from discussion with Jonathan Caine, Mike Doe, and John Ridley. Chris Holm-Denoma and Zhaoshan Chang carried out U–Pb LA-ICP-MS geochronology. We appreciate feedback from participants on a Colorado Scientific Society field trip held on 7 June 2022 to the stops of Day 1 of this field trip. Constructive comments from Mike Williams and Jamey Jones are greatly appreciated.
