The Rocky Mountain Front, southwestern USA Open Access

Why the Southern Rocky Mountains trend north-south while the Pacific–North American convergent margin and its related tectonic and magmatic features mostly trend northwest (Fig. 1) has been one of the enduring geologic mysteries of the southwestern United States. The imposing topographic escarpment along the eastern flank of the Southern Rocky Mountains between Las Vegas, New Mexico, and southern Wyoming (Fig. 1) is often referred to as the Rocky Mountain front (RMF). However, as a tectonic feature the RMF approximately coincides with long 105°W for ∼1500 km from near the U.S.-Mexico border to southern Wyoming, where it is deflected northeastward along the boundary of the Wyoming Archean province to the Black Hills of South Dakota (Karlstrom and Humphreys, 1998; Marshak et al., 2000). Since its origin in the Mesoproterozoic, the RMF has been reactivated several times and has been a significant influence in development of the Ancestral and Laramide Rocky Mountains, as well as Cenozoic magmatic and rifting events. Here we utilize a variety of geological and geophysical data to establish the timing, extent, and tectonic character of the RMF from the Mesoproterozoic onward.

This paper began out of curiosity as to how and why the RMF makes progressive steps to the right as one moves northward along it in Colorado and Wyoming, and what that could tell us about Laramide tectonics. The expectations were modest. However, as the investigation progressed, we realized that two more important subjects were involved, the Precambrian ancestry of the RMF and the nature of the lithospheric structure that underlies it. The ancestry is known, but the underlying structure is yet a mystery. So, the paper evolves from the question of the steps to the underlying lithospheric structure and ends with a list of constraints and descriptions of three geophysical studies that provide some insight into possible lithospheric controls.

Karlstrom et al. (2004, p. 23) pointed out, “An important but incompletely investigated Proterozoic north-striking boundary exists along the Rocky Mountain front…East of this boundary, 1.4 Ga rocks include shallow level plutons and volcanic rocks, whereas west of the boundary, rocks of the same age were emplaced at ∼10 km depths…Thus, a Proterozoic fault (post 1.4 Ga) with up to 10 km of east-down dip slip seems to be required to explain the rock distribution.” Yet, geological and geophysical cross sections typically cross the RMF without sensing anything unusual in the subsurface. The same is true of the Precambrian basement map of Colorado compiled by Sims et al. (2001) from interpretations of aeromagnetic anomalies.

The Neoproterozoic extensional events that accompanied breakup of the Rodinia supercontinent between 1.1 Ga and 0.6 Ga (Karlstrom and Humphreys, 1998; Marshak et al., 2000; Timmons et al., 2001; Keller et al., 2005; Luther et al., 2012) are also part of the tectonic ancestry of the RMF. This continental breakup established the structural framework of Laurentia (now the Precambrian core of North America), including the Cordilleran passive margin and a series of northwest- and north-trending fault zones. Karlstrom and Humphreys (1998, fig. 3 therein) interpreted a north-trending generalized fault zone extending from southern New Mexico to northern Colorado at 1.1 Ga as the newly created Rocky Mountain trend. Similarly, Marshak et al. (2000) interpreted a north-trending eastern edge of the Rocky Mountain–Colorado Plateau province extending from the Mexico-U.S. border region to South Dakota ca. 0.9 to 0.7 Ga. The incipient RMF transected the northeast-trending Yavapai (1.8–1.7 Ga) and Mazatzal (1.7–1.6 Ga) provinces that were progressively added to the southern boundary of the Wyoming Archean craton during the Proterozoic, as well as the 1.4 Ga Granite-Rhyolite Province (Karlstrom et al., 2004).

Sanders et al. (2006) estimated from ⁴⁰Ar/³⁹Ar thermochronometry that ∼12 km of exhumation of Mesoproterozoic rocks occurred after 1.0 Ga west of the RMF near Las Vegas, New Mexico, compared to ∼3–5 km of exhumation between 700 and 600 Ma east of the RMF. Sedimentary and volcanic rocks of the Mesoproterozoic Las Animas Group (ca. 1.1 Ga) in southeastern Colorado (Tweto, 1980, 1987) and the Debaca Group (ca. 1.26 Ga) in southeastern New Mexico (Karlstrom et al., 2004; Amarante et al., 2005) have been preserved in a stable cratonic setting east of the RMF, in contrast to the apparent uplift and denudation west of the RMF. The alkaline Pikes Peak granitic batholith (Barker et al., 1975; Wobus, 1976) in central Colorado was emplaced at shallow depths on the RMF ca. 1.09 Ga (Smith et al., 1999; Karlstrom et al., 2004) and is close to the present western margin of the continental interior craton. Thus, there seems ample evidence to conclude that the RMF occupies a linear lithospheric boundary that originated between 1.4 and 1.1 Ga during the Mesoproterozoic.

Further evidence comes from the North American Cordilleran belt of alkaline igneous rocks that follows the RMF, and is known as the Rocky Mountain alkalic province (Lindgren, 1933; Wooley, 1987; Allen and Foord, 1991; Mutschler et al., 1991; McLemore, 1996; Cappa, 1998; Kelley and Ludington, 2002; Jensen and Barton, 2000, 2007). Figure 2 summarizes the alkaline igneous rocks along the RMF in Colorado (modified from Cappa, 1998). The igneous rocks and their associated mineral deposits range in age from 1090 Ma (Mesoproterozoic) for the Pikes Peak granitic batholith to 27–20 Ma (Oligocene–Miocene) at Spanish Peaks and the numerous synrift intrusions of the Sangre de Cristo Range (Miggins, 2002). Important mineral deposits include diamond-bearing kimberlite diatremes in the State Line district (other kimberlite intrusives occur near Estes Park and Boulder), major gold deposits of the northeastern Colorado Mineral Belt (CMB) at Central City and in Boulder County, the largest vein-type uranium deposit in the U.S. near Ralston Buttes, a world-class gold deposit in an Oligocene diatreme complex at Cripple Creek, rare earth and thorium deposits associated with Cambrian alkaline intrusives in the Wet Mountains, and precious metal deposits in Eocene–Oligocene volcanic centers at Silver Cliff and Rosita (see Cappa, 1998; Abbott and Cook, 2012, for well-organized summaries and references). Similar alkaline rocks and gold deposits are present along the RMF in New Mexico (McLemore, 1996; Kelley and Ludington, 2002).

Along the convergent-margin volcanic arcs of both North and South America, igneous rocks closest to the thicker, colder lithospheres of the continental interiors tend to be strongly alkaline (Mutschler et al., 1987, 1991; Allen and Foord, 1991; Déruelle, 1991; Kay and Gordillo, 1994; Kelley and Ludington, 2002). The RMF parallels the western margin of the thick, cold cratonic interior of North America (Lee and Grand, 1996; Lerner-Lam et al., 1998; West et al., 2004; Yuan and Romanowicz, 2010) with its ≥200-km-thick lithosphere, low heat flow, and fast P-wave velocities (Gao et al., 2004; Eaton, 2008). The CMB provides a good example of the effects of thick lithosphere and low heat flow on the compositions of igneous rocks and associated mineral deposits. The northeastern end of the CMB is on the RMF near Boulder (Fig. 2), and the southwestern end is on the Colorado Plateau near the Four Corners intersection of contiguous states. Both ends have relatively thick lithospheres, low heat flow, and alkaline intrusions with associated gold deposits (e.g., Boulder County and Central City at the northeastern end; Allard stock in the La Plata Mountains near the southwestern end). Between, the CMB is characterized by mainly granodiorite–quartz monzonite stocks with associated base metal and silver deposits (Mutschler et al., 1987; Bookstrom, 1990; Chapin, 2012). The alkaline compositions are thought to be due to smaller degrees of partial melting at greater depth, probably within the lithospheric mantle and/or lower crust, and possibly with enrichment of source rocks by metasomatic activity (Mutschler et al., 1987; Stein and Crock, 1990; Kelley and Ludington, 2002; Pilet et al., 2008; Chapin, 2012).

Progressive westward collision and suturing of Gondwana with Laurentia (Fig. 3B) from Late Mississippian to early Permian time formed the Ouachita-Marathon orogenic belt in Oklahoma and Texas (Kluth, 1986; Dickinson and Lawton, 2003; Kues and Giles, 2004; Keller and Stephenson, 2007; Nance and Linnemann, 2008; Soreghan et al., 2012). Far-field stresses imposed on the Laurentian foreland (now southwestern North America) resulted in widely distributed intracratonic basement-cored uplifts and adjoining basins. The overall pattern of Ancestral Rocky Mountains (ARM) uplifts and basins trends northwestward ∼2000 km from the Llano uplift in southern Texas to southern Idaho (Fig. 3A) with a width of 750–1000 km (Kluth, 1986; Dickinson and Lawton, 2003; Soreghan et al., 2012). The individual uplifts trend mainly west-northwest to north-northwest except along the RMF (105°W), where northward trends dominate (Kluth, 1986; Dickinson and Lawton, 2003; Soreghan et al., 2012). The northward alignment along the RMF is also visible in Figure 4, the residual gravity map of Soreghan et al. (2012), and on the aeromagnetic map (not shown) of Karlstrom et al. (2004). A series of little-known ARM structures mapped by Kelley (1972) extends for more than 300 km south of the Sangre de Cristo Range, along 105°W in southeastern New Mexico.

The origin of ARM deformation remains enigmatic in several respects in spite of abundant stratigraphic and structural data gathered during intense petroleum exploration and many decades of geologic and geophysical research. The great areal extent of the intracratonic deformation, the overall pattern of uplifts and basins, the type of faulting, and the lack of magmatism are all problematic. The tectonic setting of the ARM deformation is controversial; interpretations range from a continental-scale collisional system associated with suturing of Laurentia and Gondwana (Moores, 1991; Nance and Linnemann, 2008) to a Laramide-style foreland contractional system related to subduction along the southwest margin of Laurentia (Ye et al., 1996). The contemporaneity in timing of the intracratonic deformation with the progressive southwestward collision and suturing of Laurentia and Gondwana and the lack of magmatism makes this the generally accepted model. For a detailed space-time analysis of the time-transgressive, westward-younging history of block uplifts and basin filling, see Kluth (1986), Dickinson and Lawton (2003), and Kues and Giles (2004).

The orientation of ARM uplifts and basins with northwest and northward trends crossing each other in southern Colorado and northern New Mexico is puzzling. But when the effect of reactivation of older structures is considered, the causes of the pattern become clear. The failed rift of the Cambrian southern Oklahoma aulacogen, with related mafic and alkaline intrusions that extend intermittently for ∼1500 km northwestward to the Uncompahgre uplift in southwestern Colorado (Figs. 3 and 4), apparently controlled the major northwest-trending ARM structures of Pennsylvanian and early Permian age (Kluth, 1986; Keller and Stephenson, 2007; Soreghan et al., 2012). The north-trending ARM structures reflect basement control by the RMF. Figure 5 illustrates how these competing northwest and northward trends have influenced ARM, Laramide, and Rio Grande Rift structures in southern Colorado and northern New Mexico from Pennsylvanian to Neogene time (Chapin and Seager, 1975). Transmission of compressive stresses from collision and suturing along the Ouachita-Marathon orogenic belt (Fig. 3B), and/or from Andean-type subduction in northeastern Mexico (Ye et al., 1996), activated preexisting structures with many uplifts overlapping in time and some in space. Deformation tended to be along high-angle thrusts or reverse faults, some accompanied by strike-slip movement, but normal faults are also present. The lack of volcanism indicates far-field stresses regardless of origin.

The Late Cretaceous–Paleogene (ca. 75–45 Ma) partitioning of the Cordilleran foreland from the Sevier fold and thrust belt to the RMF (Fig. 6) reflects northeastward-directed compressive stresses from the Cordilleran convergent margin and the northeastward subhorizontal subduction of the Farallon plate (Coney, 1972, 1978; Coney and Reynolds, 1977; Cross and Pilger, 1978b; Hamilton, 1981; Gries, 1983; Bird, 1984; Cross, 1986; Dickinson et al., 1988). Deformation of the Rocky Mountain foreland coincided temporally with an increase in the rate of Farallon–North American convergence from ∼100 km/m.y. to as much as 150 km/m.y. during the interval ca. 75–45 Ma (Coney, 1978; Jurdy, 1984; Engebretson et al., 1985; Chapin, 2012, fig. 2 therein). Contraction in the Sevier fold and thrust belt west of the foreland continued intermittently until ca. 50 Ma (Jordan, 1981; DeCelles, 1994; DeCelles and Mitra, 1995). Thus, the Rocky Mountain–style basement-cored arches (Erslev, 1993) and block uplifts, including those along the RMF, overlapped significantly in time with contraction in the Sevier fold and thrust belt (Dickinson et al., 1988; Kulik and Schmidt, 1988). This raises the interesting possibility that Laramide deformation was driven concurrently at two levels: (1) end loading of the middle and upper crust by far-field forces from the Cordilleran convergent margin, and (2) northeastward drag imparted to the basal lithosphere through viscous coupling with the subducting Farallon plate.

The beginning of Laramide deformation and initial reactivation of ARM uplifts are well constrained by the preservation of synorogenic strata in the adjacent basins ( Appendix 1). The change is often manifest in the transition from marine shales to delta front sands overlain by coastal plain terrigenous deposits with interbedded coal seams and the first signs of detritus shed from the Precambrian basement (Weimer, 1976; Raynolds, 2002, 2003). The tectonic interpretations of Paleocene to middle Eocene terrestrial deposits are not so straightforward; they are often separated by significant lacunae from the upper Cretaceous–lower Paleocene transitional sequence, and are more difficult to date due to scarcity of fossils and datable volcanic ashes.

We selected four key localities along the RMF for their potential to help elucidate the history and tectonic development of the RMF. The first three are localities where the RMF steps 40–50 km to the right while maintaining the overall trend paralleling 105°W. At the fourth locality, the RMF was deflected ∼250 km to the northeast along the southeastern margin of the Archean Wyoming province to the Black Hills of South Dakota. The locations of the steps are shown in Figure 6.

In southern Colorado, the RMF steps to the right ∼40 km from the Sangre de Cristo Range to the Wet Mountains (Figs. 6 and 7); between are the northern end of the Laramide Raton Basin and its northern extension, the Paleogene Huerfano Park basin (Fig. 7). Kleinkopf et al. (1970) showed a 15 mgal, north-northwest–trending, closed Bouguer gravity low over Huerfano Park. The syntectonic sedimentary fill of Huerfano Park consists of the Cuchara and Huerfano Formations and the Farasita facies, with an aggregate thickness of more than 1500 m (Cather, 2004). The Huerfano and Farasita were considered by Briggs and Goddard (1956), Robinson (1963, 1966), and Scott and Taylor (1975) to be laterally equivalent facies. The Huerfano Formation contains red sandstone and mudstone derived mainly from sedimentary redbeds of Pennsylvanian and Permian age in the Sangre de Cristo Range. The Farasita conglomerate contains yellowish-gray coarse conglomerate and sandstone derived mainly from granitic crystalline rocks in the Wet Mountains, and the Cuchara Formation is a mixture of Huerfano and Farasita lithologies (Scott and Taylor, 1975). Early-middle Eocene vertebrate faunas have been found in the Huerfano Formation (Robinson, 1966; Lindsey, 1998).

The Huerfano Formation unconformably overlies the Paleocene Poison Canyon Formation of the Raton Basin and is overlain by volcaniclastic fluvial and debris-flow deposits of the lower Oligocene Devils Hole Formation (Johnson and Wood, 1956; Scott and Taylor, 1975). The volcanic constituents of the Devils Hole Formation were derived from the late Eocene–early Oligocene Silver Cliff and Rosita Hills volcanic centers, ∼20 km to the north-northwest (Fig. 7) in the Wet Mountain Valley, and the Deer Peak volcanic center ∼10 km to the northeast in the Wet Mountains (Scott and Taylor, 1975; McIntosh and Chapin, 2004). The Ar⁴⁰/Ar³⁹ ages (reported in McIntosh and Chapin, 2004) range from ca. 35 to 32 Ma. (For a perspective of how the Huerfano Park basin relates to other Laramide basins in northern New Mexico and southern Colorado, see Cather, 2004, fig. 2 therein.)

Huerfano Park merges to the northwest with the Wet Mountain Valley graben, which separates the northeast-vergent thrusts of the east flank of the Sangre de Cristo Range from the normal-faulted western margin of the Wet Mountains. The thrust-faulted Pennsylvanian and Permian formations of the Sangre de Cristo Range were deposited in the central Colorado trough, a long, narrow Ancestral Rocky Mountain basin between the huge San Luis–Uncompahgre uplift on the southwest and the Apishapa and Frontrangia uplifts (“Frontrangia” is a feature distinguished from the Laramide Front Range) on the northeast (Figs. 5 and 6). The Laramide Wet Mountains and Wet Mountain Valley were superimposed on the Apishapa uplift, which was denuded to the Precambrian basement during the Ancestral Rocky Mountain orogeny (Tweto, 1975; Lindsey et al., 1983). Three alkalic intrusive complexes of Cambrian to Early Ordovician age occur in the northern Wet Mountain Valley and were exposed during Pennsylvanian–Permian uplift and erosion of the Wet Mountain block (Naeser, 1967; Taylor et al., 1975; Olson et al., 1977). The complex thrust faulting of the northern Sangre de Cristo Range is prominent on geologic maps of Colorado as the most highly deformed thrust-faulted terrain in the state. During northeast-oriented Laramide compression, the Pennsylvanian–Permian sedimentary fill of the central Colorado trough was shortened between the San Luis–Uncompahgre uplift on the southwest and the Wet Mountain block on the northeast. Lindsey et al. (1983) estimated the magnitude of shortening as ∼8–14 km (or 40%–70%) across an original width of 20 km. Burbank and Goddard (1937) estimated 15–17 km shortening for Paleozoic and Mesozoic rocks in Huerfano Park.

Northeast-directed Laramide compression was transmitted across the Wet Mountain block, resulting in the eastern margin of the Wet Mountains being thrust over upper Cretaceous rocks of the Canon City embayment (Scott et al., 1978; Weimer, 1980; Jacob and Albertus, 1985). Several anticlines and synclines involving Precambrian to upper Cretaceous rocks closely parallel the thrusted eastern margin of the Wet Mountains. The eroded core (∼22 × 6 km) of the Red anticline (Fig. 7) reveals patches of Ordovician rocks overlying Precambrian basement and overlain by arkosic redbeds of the Pennsylvanian–Permian Fountain Formation and the Jurassic Morrison Formation (Scott et al., 1978; Jacob and Albertus, 1985). Interpretations of the Wet Mountain thrust range from a low-angle thrust near the surface, with several kilometers of overhang (Jacob and Albertus, 1985), to a high-angle reverse fault (Weimer, 1980). Reactivation of the Wet Mountain block occurred in the early-middle Eocene, recorded by the southwestward-prograding alluvial fans of Precambrian detritus that formed the Farasita facies of the Huerfano Formation (Scott and Taylor, 1975). Reactivation occurred again in Miocene–Pliocene time during development of the Wet Mountain Valley graben as a subsidiary basin of the Rio Grande Rift containing alluvial fill correlated with the Santa Fe Group (Scott and Taylor, 1975).

Tweto (1975, p. 1) stated that “… eroded stumps of late Paleozoic mountain ranges made up of Precambrian rocks” were present at the beginning of the Laramide orogeny beneath a blanket of Cretaceous sedimentary rocks. The apatite fission track (AFT) cooling ages in Kelley and Chapin (2004) show that the stump of the Wet Mountain block (∼2–3 km in the subsurface) had cooled below ∼120–100 °C before burial by upper Cretaceous sedimentary deposits, but was not buried deeply enough to reset the AFT ages (Fig. 7). The Wet Mountain block was not uplifted sufficiently for the rocks with Paleozoic and Mesozoic AFT ages to be removed by erosion. At Greenhorn Mountain, the high peak at the south end of the Wet Mountains (Fig. 7), ∼1200 m of Precambrian rocks are still present with pre-Laramide, partially reset AFT ages between 228 and 166 Ma (Kelley and Chapin, 2004).

A recent study of the erosion and uplift history of the central and southern Rocky Mountains (Cather et al., 2012) indicates that the Wet Mountains and adjacent high country of the Colorado Rocky Mountains were not strongly uplifted and dissected by erosion until late Miocene to Holocene time (ca. 6–0 Ma), which helps explain preservation of the old AFT ages. Today, the middle Eocene erosion surface (Epis and Chapin, 1975) that MacGinitie (1953) estimated was carved at elevations <900 m (see also Cather et al., 2012) rises southward along the crest of the Wet Mountains until it reaches modern elevations as high as 3500 m at Greenhorn Peak. The apparent step to the northeast of the RMF from the Sangre de Cristo Range to the Wet Mountains was simply reactivation of the Apishapa uplift, beveled during and after the Pennsylvanian–Permian ARM orogeny.

The RMF steps northeastward ∼50 km from the eastern edge of the Wet Mountains to the eastern edge of the southern Front Range (Figs. 6 and 7); between are the Cañon City embayment, its northern extension, the Garden Park syncline, and the Cañon City–Florence basin (Fig. 7;  Appendix 1). A broad plateau composed mostly of Precambrian rocks is to the northwest; it is thinly covered by upper Eocene to lower Oligocene volcanic rocks of the Thirtynine Mile volcanic field, now part of the Central Colorado volcanic field (McIntosh and Chapin, 2004). The area discussed here is mostly within the Pueblo (Scott et al., 1978) and the Denver 1° × 2° geologic quadrangles (Bryant et al., 1981b). The Central Colorado volcanic field has been deeply dissected by erosion; widely scattered remnants are present in the northern Wet Mountain Valley, South Park, the southern Front Range, and the High Plains of eastern Colorado. The Central Colorado volcanic field is mainly in an elongate belt that extends north-northwest from the Raton Basin to southern Wyoming and includes the Laramide intermontane basins of Huerfano Park, Echo Park, South Park, and Middle Park–North Park (called the Colorado Headwaters Basin by Cole et al., 2010). The basins were classified as axial basins by Dickinson et al. (1988). The belt containing the basins is essentially the lower elevation terrain between the high frontal ranges on the east and high interior ranges like the Sangre de Cristo, Sawatch, and Sierra Madre on the west (Figs. 6 and 7).

The old AFT cooling ages that characterize the Precambrian rocks of the Wet Mountain block extend northward across the Arkansas River gorge (Fig. 7). A sample traverse along the Arkansas River revealed that AFT ages between 92 and 174 Ma extend ∼30 km westward from the range front before giving way to Laramide cooling ages (75 Ma or younger) near the Texas Creek fault (Fig. 7). The old AFT ages are near the southern edge of a broad area in the Front Range, south of the CMB, with AFT ages older than 100 Ma. Near the center of the southern Front Range, near Bailey, Colorado, we (Kelley and Chapin, 2004) found ∼1200 m of Proterozoic rocks preserved above the base of the Late Cretaceous partial annealing zone (PAZ) of fission tracks in apatite. North of the Arkansas River gorge, on the plateau-like surface beneath the Central Colorado volcanic field, an inlier of Ordovician rocks and patches of Jurassic Morrison Formation and Cretaceous Dakota Sandstone directly overlie Proterozoic rocks. Thus, the southern Front Range and its western flank compose the reactivated core of the ARM Frontrangia uplift, as noted by Mallory (1972), Tweto (1975), and Kluth (1997). The geographic extent and thickness (as much as 1200 m) of Proterozoic rocks that cooled during uplift and denudation in Paleozoic and early to middle Mesozoic time, but are still preserved in the Laramide Rocky Mountains, are surprising.

The broad plateau of mostly Proterozoic rocks that underlies the Central Colorado volcanic field north of the Arkansas River is cut by a series of north-northwest–trending major faults (Fig. 7), some of which bound narrow strike-slip basins that contain as much as 600 m of prevolcanic arkosic alluvium (Chapin and Cather, 1981; Chapin et al., 1982). The best exposed basin, named after Echo Park canyon on the north side of the Arkansas River gorge, extends ∼65 km northward to near Hartsel in the South Park basin. The Echo Park basin is ∼1.5–5 km wide and as deep as 600 m. The fault zone bounding the east side contains lenses of Morrison and Dakota sandstones and exhibits dextral kinematic indicators (Chapin and Cather, 1981). Large uranium deposits discovered in the Echo Park basin in 1977 (Chapin et al., 1982) led to extensive drilling programs between the Arkansas River and South Park. The drilling documented the existence of other narrow, fault-bound basins containing arkosic alluvial fill, but much of the data remain proprietary. A late-early to early-middle Eocene age of the Echo Park Alluvium ( Appendix 1) was estimated based on pollen assemblages in core samples (R.H. Tschudy, U.S. Geological Survey, 1983, written commun.). The Echo Park Alluvium is overlain by the Wall Mountain Tuff, ³⁹Ar/⁴⁰Ar dated as 36.7 ± 0.07 Ma (McIntosh and Chapin, 2004).

Northeast-directed Laramide compressive stresses were transmitted across the remnants of Frontrangia uplift just as they were across the Wet Mountains block. The east flank of the southern Front Range was crowded against Phanerozoic sedimentary formations underlying the southern Denver Basin along an echelon series of moderately steep thrusts (Tweto, 1979; Trimble and Machette, 1979; Kluth, 1997; Weimer and Ray, 1997). The scarcity of seismic lines and drill holes has resulted in a wide variety of conflicting structural interpretations of the range-bounding faults. In a detailed, model-based structural analysis, Sterne (2006) interpreted the southeastern flank of the Laramide Front Range as a complex series of stacked triangle zones. Nesse (2006) included a foldout sheet containing 21 cross sections of the Front Range from 38°30′N to 41°00′N. The cross sections reveal the structural asymmetry of the Front Range, characterized by low-angle, west-vergent thrusts along the west flank and relatively short, moderate-angle, east-vergent thrusts and high-angle reverse faults along the east flank (see  Appendix 1). The Laramide Front Range does not coincide completely with the Ancestral Rocky Mountain Frontrangia uplift, as shown in Figure 8; the Frontrangia uplift trends more to the northwest, so that the two uplifts diverge in northern Colorado (Kluth, 1997; Erslev, 2005). The east flank of the Laramide Front Range changes character north of its intersection with the Colorado Mineral Belt (discussed in the following).

A very interesting and important change in structural style of the eastern margin of the Laramide Front Range occurs between Golden and Boulder, Colorado (Fig. 9). The transition in AFT ages is surprisingly abrupt, occurring within a 3.5 km interval between Golden Gate and Van Bibber Canyons northwest of Golden (Kelley and Chapin, 2004). No major faults cross the margin of the Denver Basin to account for the abruptness; however, the transition occurs across the trend of the northern CMB. As pointed out by Tweto and Sims (1963), the CMB cuts indiscriminately across the geologic grain of Colorado with remarkable continuity, seemingly independent of the tectonic elements it crosses. In Chapin (2012) the origin of the CMB was interpreted as due to an underlying segment boundary in the subhorizontally subducting Farallon plate that became dilated as the Farallon slab was overridden by the thicker lithospheres of the Wyoming and continental interior cratons. Thus, the control of the CMB was not in the North American plate and no faults need cross the Proterozoic-Phanerozoic boundary of the range front. For alternative hypotheses see Tweto and Sims (1963), McCoy et al. (2005), and Jones et al. (2011).

The Front Range–Denver Basin margin changes from a north-northwest trend south of the CMB intersection to a northward trend (Figs. 8 and 9) within ∼10 km of the intersection. The Golden thrust is no longer mappable at the surface north of this zone. The structural style of the range front also changes from northeast-vergent thrusts with triangle zones characteristic of the southern margin to northwest-trending, northeast-dipping, high-angle reverse faults (Erslev 1993; Nesse, 2006; Cole and Braddock, 2009) north of the CMB intersection. The axis of the Denver Basin is close to the range front south of the CMB due to loading of the basin by overhanging northeast-vergent thrusts. Starting at the CMB intersection, the basin axis veers to the northeast to near Greeley, Colorado (Fig. 9); between Greeley and the Wyoming border the axis is 35–40 km east of the range front. This is the area in which the structural style of the Front Range changes from northeast-vergent thrusts to northeast-dipping, high-angle reverse faults, also called southwest-vergent backthrusts by Erslev (1993) and Erslev and Selvig (1997), and possibly caused by basement wedging and backthrusting above a blind northeast-directed thrust in the basement (Erslev 1993; Erslev and Holdaway, 1999). The Denver and Cheyenne Basins contain asymmetric thrust-loaded deeps in the Denver and Cheyenne areas (Fig. 10) separated by a relatively shallow saddle in the Greeley area known as the Greeley arch (Weimer, 1996).

Figure 9 displays histograms of AFT cooling ages for areas of the Front Range north and south of the CMB intersection (Kelley and Chapin, 2004). South of the CMB, the AFT ages range from ca. 30 Ma to older than 300 Ma, and many samples show relatively short track lengths indicative of long residence within the AFT partial annealing zone. In contrast, north of the CMB the AFT ages span a narrow range from ca. 80 to 40 Ma and characteristically have relatively long track lengths indicative of fast cooling. In Kelley and Chapin (2004), we interpreted the narrow range of AFT ages and their correlation with the Laramide orogeny north of the CMB as being due primarily to the PAZ of AFT ages being up within the thicker and relatively low thermal conductivity Pierre Shale that was subsequently eroded away, revealing the Laramide cooling ages in the Proterozoic basement beneath the PAZ. In contrast, the end-Cretaceous PAZ south of the CMB was within the resistant basement rocks and preserved from erosion. However, the abruptness of the transition (∼3.5 km) across the CMB, the northeastward deflection of the axis of the Denver Basin, the shallowing of the basin, and the change in structural style suggest that uplift of the Front Range north of the CMB may also have been a factor.

There is additional evidence of early Laramide uplift across the CMB. The Boulder-Weld coal field (Fig. 9) is a 40-km-long, 10–15-km-wide northeast-trending zone of horst and graben faulting from which more than 100 × 10⁶ t of coal were produced between 1859 and 1979 from the Maastrichtian Laramie Formation (Carroll, 2009). The faults have been variously interpreted as normal, reverse, strike slip, growth, or décollements. Kittleson (2009) appears to have solved the problem via correlation of sedimentary units and faults in ∼1450 geophysical well logs that resulted in recognition of down-to-the-basin gravity sliding away from the Longmont fault (Fig. 9) above a bedding-plane detachment ∼76 m below the top of the Pierre Shale. The detached blocks included the regressive, delta front Fox Hills Sandstone (30–67 m) and the overlying delta plain coal-bearing Laramie Formation (213–262 m). The coal beds were generally thicker on the graben blocks (Davis and Weimer, 1976), indicating that the detachment occurred in a terrestrial near-surface environment with peat bogs accumulating organic matter contemporaneously with faulting. Davis and Weimer (1976) described the faulting as growth faulting and also reported unusually thick sections of Fox Hills Sandstone. The Golden–Boulder area has the greatest thickness of Pierre Shale (2295–2448 m) in the Denver Basin (Davis and Weimer, 1976). High sedimentation rates may have been a contributing factor in the gravity sliding. The importance to this study is the evidence supporting early Laramide uplift (Maastrichtian) of the Front Range north of the CMB. The gravity sliding may also have been aided by magmatic inflation and seismic activity in the northern CMB. A cluster of five intrusive centers (Fig. 9) at the northeast end of the CMB are within 20–30 km of the Boulder-Weld coal field and range in age from ca. 75 to 45 Ma (Gable, 1984; Mutschler et al., 1987).

Figure 9 also shows the outlines of the huge Wattenberg gas field and its associated geothermal anomaly, shown by contours of vitrinite reflectance values of Ro 1.0 and 1.4 (Higley et al., 2003). The Wattenberg field is a continuous-type gas accumulation located along the axis of the Denver Basin adjacent to its intersection with the CMB. Production is from the lower Cretaceous Muddy Sandstone at depths of 2070–2830 m and from overlying upper Cretaceous units, Codell Sandstone, Niobrara Formation, and the Hygiene and Terry Sandstone Members of the Pierre Shale. Stratigraphic traps, tight sands, and a paleostructural high provided the trapping mechanisms (Weimer, 1996; Higley et al., 2003; Nelson and Santus, 2011). The Wattenberg field is pertinent to this study of the RMF because of (1) its location at the intersection of the CMB with the Denver Basin; (2) its geothermal anomaly and residual magnetic intensity anomaly that may indicate a magmatic intrusion at depth (T. Grauch, inHigley et al., 2003); (3) its location where the axis of the Denver Basin is deflected northeastward to a position 35–40 km east of the basin margin where the axis parallels the north-trending, reverse-faulted segment of the Front Range (Fig. 9); and (4) the structural and stratigraphic control provided by thousands of drill-hole logs. The five small-displacement, oblique-slip fault zones mapped across the Wattenberg field by Weimer (1996) were not included in Figure 9 because they appear to have little tectonic significance, other than causing compartmentalization effects within the gas field. No northeast-trending faults cut the Front Range–Denver Basin margin; this is another example of the CMB cutting indiscriminately across the geologic grain of Colorado, independent of the tectonic elements it crosses, as pointed out by Tweto and Sims (1963).

The RMF continues northward along the east edge of the Front Range, which merges almost imperceptively with the Laramie Range of southern Wyoming. In like manner, the Denver Basin merges northward with the Cheyenne Basin across a broad saddle known as the Greeley arch (Weimer, 1996). The Denver and Cheyenne Basins are two halves of a 500-km-long basin (if measured at the –900 m contour level in Fig. 10) with basin deeps near Denver and Cheyenne (Matuszczak, 1976). The axis of the combined basins is close to the range front (the RMF) at the Denver and Cheyenne deeps due to loading of the basin by east-vergent thrusts (Fig. 10). The relatively shallow middle of the Denver and Cheyenne Basins is opposite the section of the Front Range dominated by northeast-dipping, high-angle reverse faults (or backthrusts) that bring the basin up relative to the range (Erslev, 1993; Erslev and Selvig, 1997). This section of the Front Range contains the Indian Peaks Wilderness, Rocky Mountain National Park, and the Continental Divide, with many peaks in the 3600–4000 m range. The area is largely underlain by the massive Mesoproterozoic granite of the Longs Peak–St. Vrain batholith (Cole and Braddock, 2009), the southern margin of which approximately coincides with the CMB and the changes in structural style of the Front Range and Denver Basin as outlined here.

The Laramide AFT cooling ages (ca. 80–40 Ma), highlighted by the histogram in Figure 9 for the northern Front Range, are similar to those in the Laramie Range except for ∼15 km on either side of the boundary with the Archean Wyoming province (Fig. 11). Old AFT ages ranging from 378 to 172 Ma (Kelley, 2005) are present near the boundary, generally referred to as the Cheyenne belt (Karlstrom and Houston, 1984; Karlstrom et al., 2005). One of us (Kelley, 2005) concluded that the distribution of AFT ages in the Laramie Range is indicative of long-wavelength folding of the basement, similar to other Laramide uplifts in southern Wyoming. Two range-parallel doubly plunging anticlines with Laramide AFT ages are separated by a low area along the Cheyenne belt (Fig. 11), where old AFT ages and short track lengths characterize rocks that cooled within the Late Cretaceous PAZ. Reflection seismic lines recorded by the Consortium for Continental Reflection Profiling (COCORP) across the Laramie Range–Cheyenne Basin margin (Fig. 11) revealed a series of westward-dipping reflections traceable as deep as 10–12 km and arranged en echelon with dips of 20°–50° (Brewer et al., 1982). Reflectors inferred to be sedimentary formations could be traced as far as 3 km west of the mountain front under the overhang of Precambrian rocks (Brewer et al., 1982). Thus, the RMF at the latitude of the Laramie Range is similar to that of the Front Range south of the CMB.

Northeast-directed compressional stresses transmitted across the Wyoming Archean province during the Laramide orogeny generated the large (320 × 110 km) north-northwest–trending Black Hills uplift (Figs. 1 and 6). The Laramide evolution of the Black Hills uplift is covered in the map by Lisenbee (1985) and in Shurr et al. (1988) and Lisenbee and DeWitt (1993). However, the Cheyenne belt between the Laramie Range and the Black Hills has received scant attention in the geologic literature. For discussions of the Precambrian tectonics and metallogeny of the Hartville uplift, Wyoming, see Sims et al. (1996) and Sims and Stein (2003). The Wheatland-Whalen fault zone (Fig. 11) trends northeast from the Laramie uplift toward the Black Hills, but the surface is largely covered by the upper Eocene White River Group and Miocene Ogallala Formation (Love and Christiansen, 1985) and there is little agreement on the nature of the faulting. The Precambrian basement is at shallow depths along a subsurface arch connecting the Laramie and Black Hills uplifts (Blackstone, 1993; Erslev, 1993), so petroleum exploration has been minimal. For these reasons, we end our coverage of the RMF at the Laramie Range.

Orogenic activity along the RMF occurred in several episodes of differing intensity and causes during the Late Cretaceous and the Cenozoic. Evidence for four episodes of Cenozoic erosion and uplift in southwestern North America was presented in Cather et al. (2012). In episode one, the Laramide orogeny (ca. 75–45 Ma) reactivated the RMF through regional northeast-directed compression. Intermontane basins remained at low elevations with the RMF marking the eastward limit to basement-cored uplifts, inferred here to be due to its proximity to the thick (≥200 km) continental interior lithosphere (Lee and Grand, 1996; Lerner-Lam et al., 1998; West et al., 2004; Gao et al., 2004; Yuan and Romanowicz, 2010). In episode two, late-middle Eocene erosion (ca. 43–37 Ma) was centered in Wyoming and northern Colorado and is inferred to be in response to lithospheric rebound following cessation of Laramide dynamic subsidence. Carving of the Rocky Mountain erosion surface and shallowing of the Denver Basin north of the CMB were two effects along the RMF.

The third episode of erosion and uplift (discussed in Cather et al., 2012) occurred in late Oligocene–early Miocene time (ca. 27–15 Ma) in response to increased mantle buoyancy related to major volcanism in the southern Rocky Mountains and Sierra Madre Occidental of Mexico. Early development of the Rio Grande Rift overlapped with this episode and rifting increased markedly during the interval ca. 16–6 Ma (Chapin and Cather, 1994). Uplift and outward tilting of the east flank of the Rio Grande Rift provided source areas for aggradation of the Ogallala Formation on the western Great Plains (Chapin and Cather, 1994).

Accelerated uplift and erosion during the late Miocene–Holocene (ca. 6–0 Ma) episode four of Cather et al. (2012) resulted in major enhancement of the RMF along the Front Range, Wet Mountains, and Sangre de Cristo Range. The erosion was increased by opening of the Gulf of California ca. 6 Ma (Oskin and Stock, 2003; Chapin, 2008) which, together with the effect of higher elevation, intensified the North American monsoon. The east-facing, ∼500–1200 m topographic escarpment along the east flank of the southern Rocky Mountains owes its existence to a combination of late Miocene–Holocene uplift and erosional exhumation (Leonard and Langford, 1994; Chapin and Cather, 1994; Leonard et al., 2002). The intensified erosion removed the alluvial fans along the mountain fronts, eroded hogbacks of tilted Paleozoic and Mesozoic formations, and beheaded the Ogallala Formation from its source terrains in the mountains. Strike valleys paralleling the mountains are home to the larger cities of the Front Range urban corridor with the Rocky Mountain Front as their scenic backdrop.

The Front Range dominates the RMF in Colorado with its ∼280 km length, ∼60–100 km width, and ∼2 km of topographic relief; however, it is not a monolithic feature. The northern half differs significantly from the southern half, the CMB being the dividing line. As illustrated in Figure 9, north of the CMB the AFT cooling ages are nearly all Laramide in the narrow range of 80–40 Ma; the Front Range trends north and is dominated structurally by northeast-dipping, high-angle reverse faults. South of the CMB, the AFT ages cover a wide range, from ca. 458 Ma (partially reset) at the summit (4300 m) of Pikes Peak (Naeser et al., 2002) to ca. 27 Ma (Oligocene) near Cripple Creek (Kelley and Chapin, 2004). The pre-Laramide AFT samples exhibit short track lengths indicative of slow cooling and lengthy presence within the end-Cretaceous PAZ. At Greenhorn Peak in the Wet Mountains and near the town of Bailey in the central Front Range, ∼1200 m of Proterozoic rocks with pre-Laramide cooling ages still reside above the base of the end-Cretaceous PAZ. The inescapable conclusion that the southern Front Range has undergone modest exhumation compared to the northern third is also suggested by the preservation of several inliers of rocks ranging in age from Cambrian to Cretaceous in the southern part of the range.

There are also conspicuous changes in the adjacent Denver Basin that relate to structural changes along the Front Range. South of the CMB, the Front Range trends north-northwest and the axis of the Denver Basin is close to the range front due to gravitational loading of the basin by northeast-vergent thrust faults. The Laramide synorogenic sedimentary deposits of the Laramie Formation and the D1 and D2 sequences of Raynolds (2002, 2003) occupy the resultant asymmetric syncline. However, as the structure of the Front Range margin changes to northeast-dipping, high-angle reverse faults north of the CMB, the axis of the Denver Basin is deflected ∼60 km to the northeast (Fig. 9), the basin becomes shallower and more symmetrical, and the synorogenic sedimentary sequences D1 and D2 of Raynolds (2002, 2003) are no longer present. This leaves the Laramie Formation (Maastrichtian) as the only Laramide synorogenic sedimentary unit in the northern Denver Basin. When and where did the other synorogenic sediments go? The Laramie Formation in the northern Denver Basin is much thicker than to the south (∼540 m vs. ∼260 m) and is dominantly fluvial with several relatively thick channel sandstones (Kirkham et al., 1980). Unfortunately, the middle Eocene (ca. 43–37 Ma) episode of uplift and erosion (Cather et al., 2012) removed much of the synorogenic sedimentary record. The record was further obscured by the upper Eocene–lower Oligocene (ca. 35.5–30 Ma) blanket of eolian and fluvial tuffaceous sediments of the White River Group (Larson and Evanoff, 1998). The Front Range and Denver Basin have also undergone changes in regional tilting, as reflected by changes in direction of stream flow, stream capture, and tilting of the Rocky Mountain erosion surface from southeastward in the late Eocene to northeastward and then eastward in the Neogene to Holocene (Steven et al., 1997).

Laramide synorogenic sedimentation was similar in character and timing (dominantly Paleocene–early Eocene) on opposite sides of the Front Range ( Appendix 1), but the structural style was quite different. The eastern margin is crowded against the Denver Basin by east-vergent moderate- to high-angle thrusts (∼35°–70°) and triangle zones south of the CMB, and high-angle, northeast-dipping reverse faults north of the CMB. However, along the western margin, contraction culminated in low-angle, west-vergent thrusts that advanced over the Middle Park and South Park basins by as much as 10–15 km (Erslev et al., 1999; Kellogg et al., 2004; Nesse, 2006; Cole et al., 2010). In contrast, documented overhangs of the Denver Basin by thrusts along the east margin of the Front Range are only in the range of 2–3 km (Weimer and Ray, 1997; Nesse, 2006). The short overhangs were apparently adequate to cause subsidence and create accommodation space that was rapidly infilled by proximal sedimentary facies, leading to further subsidence and an asymmetric basin.

The Middle Park and South Park basins along the west side of the Front Range may have initially developed in a similar manner, with moderately steep, west-vergent thrusts and asymmetric basin fills. Volumetrically, the bulk of Laramide synorogenic sedimentation occurred in the Paleocene ( Appendix 1). However, in latest Paleocene–early Eocene time, a final episode of horizontal compression drove the west-vergent basin-margin thrusts to override the Middle Park and South Park basins by 10–15 km (Erslev et al., 1999; Kellogg et al., 2004; Cole et al., 2010). This intense interval of horizontal compression coincided temporally with the third major episode of shortening in the Sevier fold and thrust belt, as the Hogsback thrust (Fig. 6) added ∼21 km of additional shortening at a rate of ∼3 mm/yr during late Paleocene–early Eocene time (ca. 56–50 Ma; DeCelles, 1994). The east flank of the Front Range was little affected except for modest uplift and deposition of the D2 sequence (Raynolds, 2002, 2003) in the Denver Basin. The Middle Park and South Park basins, however, are internally complex, with intrabasin, west- to southwest-vergent thrusts and tight folding (Izett, 1968; Izett and Barclay, 1973; Beggs, 1977; Wellborn, 1977).

A remarkable attribute of the RMF is its ability to maintain its position through multiple orogenies and changes in orientation and strength of tectonic stresses. During the final paroxysm of Laramide contraction in latest Paleocene–early Eocene time, the RMF remained in place as thrusts advanced from east and west on opposite sides of a foreland partitioned by basement-cored uplifts, sedimentary basins, and the Colorado Plateau. That the RMF is a deeply penetrating, fundamental flaw in the North American lithosphere is also indicated by the wide range in age of alkaline igneous rocks and mineral deposits (summarized in Fig. 2). The Colorado Front Range is a composite structure with significant differences from north to south and east to west, and with geologic time. It represents a microcosm of the overall RMF.

An important question remains. What is the nature of the lithospheric flaw that underlies the RMF? It may not be possible to answer this question given the present lack of definitive subsurface data, but we can outline some constraints and list some features that suggest a solution. For example, the RMF has three features that are found in many suture zones: (1) a long, narrow tectonic belt that follows a cratonic boundary, (2) a long-term existence measured in hundreds of millions of years, and (3) the presence of a variety of alkaline igneous intrusions. However, exposures of Proterozoic rocks along the RMF lack the complex structural deformation, high-grade metamorphic rocks, and indications of major strike-slip faulting that are found in many suture zones. Several geologic features strongly indicate that the RMF is a major lithospheric boundary: (1) during the Laramide orogeny (ca. 75–45 Ma) contractional basement-cored uplifts ended at the RMF; (2) igneous intrusions and volcanism also ended at the RMF; (3) the flatly subducted Farallon slab broke at, or near, the RMF by ca. 37 Ma and rolled back toward the southwest as it sank into the mantle, causing widespread late Eocene–Oligocene ignimbrite volcanism; (4) the Rio Grande Rift formed in close proximity and approximately parallel to the RMF; (5) the RMF has been, and remains, a major heat flow boundary; and (6) the RMF separates regions with different patterns of anomalies on the residual gravity map of Figure 4.

The results of three types of geophysical studies are especially pertinent to the question of what underlies the RMF. The first concerns the use of magnetometer arrays to map anomalies in magnetovariation fields in the western United States and Canada (see review by Gough, 1989). The natural magnetovariation fields are produced by primary currents in the ionosphere and magnetosphere and by secondary induced currents in the solid Earth. The induced currents flow preferentially in the more conductive rocks. High conductivity can be generated by either saline hot waters in fractures or by a few percent of silicate melt in interconnected spaces in the crust or lithospheric mantle. Alabi et al. (l975) and Camfield and Gough (1977) identified a narrow zone of high conductivity, which was named the North American Central Plains (NACP) conductor, and mapped it over a total length of ∼1800 km from ∼57°N in Canada to the Laramie Range in southern Wyoming. Gough (1989, Fig. 2 therein) extended the NACP conductor southward along the RMF to near the U.S.-Mexico border with a similar conductor mapped along the Wasatch Range that borders the Colorado Plateau on the west. Gough (1989, p. 148) stated “…there is little doubt that the NACP marks a major fracture zone.” That the NACP extends over such a long distance, including areas that have not undergone magmatism since the Proterozoic, suggests that saline waters in fractures may be the dominant cause of the enhanced conductivity. That the geomagnetic conductivity anomalies tracked the eastern and western boundaries of the Rocky Mountain–Colorado Plateau province rather than the zone of partial melting beneath the southern Rocky Mountains and Rio Grande Rift (Gao et al., 2004; West et al., 2004; Karlstrom et al., 2012) also suggests the presence of saline waters in fracture zones.

The second series of geophysical studies began with the PASSCAL (Program for Array Seismic Studies of the Continental Lithosphere) Rocky Mountain Front experiment in 1991–1992, in which 36 seismographs were deployed from eastern Utah to western Kansas and north-south within the state borders of Colorado. The two-dimensional array recorded broadband teleseismic data from 446 earthquakes; the data were processed using receiver-function techniques to estimate crustal thickness and seismic velocities in the crust and upper mantle (see Lerner-Lam et al., 1998, for details). Regional body wave and shear wave tomography (Lee and Grand,1996; Lerner-Lam et al., 1998) imaged the transition from seismically slow upper mantle beneath the Colorado Rocky Mountains to the seismically fast cratonic structure beneath eastern Colorado and western Kansas. The transition was seen as gradational east of the RMF with an abrupt increase in velocity near the Colorado-Kansas border, interpreted as the western edge of the continental interior craton (Lee and Grand, 1996; Lerner-Lam et al., 1998). A temperature contrast of at least 350 °C, as well as partial melt beneath the Rocky Mountains, was estimated by Lee and Grand (1996) to be required to match the contrast in shear wave velocity between the mantle beneath Kansas and the Colorado Rocky Mountains.

The third series of geophysical studies was based on the La Ristra (Colorado Plateau/Rio Grande Rift Seismic Transect Experiment) deep-imaging seismic profile, generated using naturally occurring earthquake sources, and extending from the Four Corners area to southeastern New Mexico and west Texas along a S45°E line (Gao et al., 2004; West et al., 2004; Reiter and Chamberlin, 2011). The La Ristra profile crosses the RMF at an oblique angle, but projections constructed approximately normal to the RMF (Reiter and Chamberlin, 2011) show an abrupt transition over ∼50 km from low heat flow (∼41–50 mW/m²) and relatively fast shear wave velocity of the southern High Plains to the high heat flow (∼60 to >100 mW/m²) and relatively slow shear wave velocity of the RMF and eastern shoulder of the Rio Grande Rift. The La Ristra experiment documented present-day partial melting and mantle convection beneath the southern Rocky Mountains and Rio Grande Rift with a remarkably sharp seismic velocity and heat flow boundary at the RMF.

In summary, the lithospheric flaw underlying the RMF has provided a tectonic boundary and a pathway to the surface for alkaline magmas, kimberlite intrusions, and hydrothermal fluids under varying tectonic conditions over more than a billion years of geologic time. However, the nature of the lithospheric structure underlying the RMF remains elusive.

We thank Virginia McLemore and Mike Timmons for informal reviews of an early version of the manuscript. Chapin benefited from discussions with Virginia McLemore concerning her work on the Great Plains margin gold deposits. James Cole and Jeremiah Workman kindly provided copies of the Estes Park quadrangle and a preliminary copy of the Fort Collins 1:100,000 quadrangle. The patience and skills of Lynne Hemenway in word processing and Leo Gabaldon in computer graphics were extremely helpful. We also thank the reviewers E.A. Erslev, G.R. Keller, and associate editor R.M. Flowers for helpful comments and suggestions that resulted in a better paper.

Location

Basin

Structural style

Laramide sedimentary units

Age, thickness, paleontologic, paleomagnetic, and/or isotopic dating

Key references

East-vergent, low-angle thrusts from Sangre de Cristo Range

Huerfano Formation (Fm.), lower-middle Eocene, 0–1500 m, vertebrate faunas

Poison Canyon Fm., Paleocene, 0–590 m, overlies upper Cretaceous formations

Johnson (1959); Scott and Taylor (1975); Lindsey (1998) 

Dextral strike-slip faults, north-northwest trending

Echo Park Fm., early-middle Eocene, 0–600 m, pollen, narrow elongate basins

capped by 37 Ma Wall Mountain Tuff

Chapin and Cather (1981); Chapin et al. (1982) 

West-vergent, low-angle Elkhorn thrust

South Park Fm., Paleocene–early Eocene(?), 0–3050 m, K-Ar 65.5–56.3 Ma

Laramie Fm., Maastrichtian, 90 m

Sawatzky (1964, 1967); Barker and Wyant (1976); Wyant and Barker (1976);

Beggs (1977); Bryant et al. (1981a); Raynolds (2003) 

West-vergent, low-angle Williams Range thrust

Middle Park Fm., Paleocene–lower Eocene, ∼1000 m, pollen zones P3 and P5

60–58 Ma Windy Gap Volcanic Member (Mbr.), ∼200 m, ⁴⁰Ar/³⁹Ar 60.5 Ma.

Pierre Shale, Hygiene Sandstone Mbr., ammonoid fossils, 75.8 Ma

Izett (1968); Izett and Barclay (1973); Erslev et al. (1999); Kellog et al. (2004); Cole et al. (2010) 

Several west-vergent, short, en echelon steep reverse faults (Cole et al., 2010)

Coalmont Fm., middle and upper Paleocene–lower Eocene, ∼2700 m, pollen

zones P5 and 6, locally P3 and 4

Hail (1965, 1968); Wellborn (1977); Flores (2003); Cole et al. (2010) 

East-vergent, moderate-angle Wet Mountain thrust

Poison Canyon Fm., Paleocene, 150–300 m

Raton Fm., Maastrichtian–Paleocene, 73–152 m, pollen

Vermejo Fm., Maastrichtian, 45–230 m, marine invertebrates

Scott and Taylor (1974); Scott (1977); Weimer (1980) 

East-vergent, moderate-angle thrusts, asymmetric basin subsidence;

major lacuna ca. 64–56 Ma

D2 sequence, upper Paleocene–lower-middle Eocene, 0–360 m, pollen, paleomagnetic analyses (pmag.),

K-Ar, ca. 56–53 Ma

D1 sequence, Maastrichtian–lower Paleocene, 0–600 m, pollen, pmag.,

K-Ar, ca. 68–64 Ma

Laramie Fm., Maastrichtian, 60–150 m, pollen, pmag., ca. 69–68 Ma

Soister and Tschudy (1978); Trimble and Machette (1979); Bryant et al. (1981b);

Weimer (1996); Weimer and Ray (1997); Kelley (2002); Nichols and Fleming (2002); Farnham and Kraus (2002); Raynolds (2002, 2003); Raynolds and Johnson (2003); Hicks et al. (2003) 

Northeast-dipping, northwest-trending, high-angle reverse faults (also called backthrusts; Erslev, 1993)

Laramie Fm., Maastrichtian, 60–450 m

Denver Basin shallows to north over transverse Greeley arch and merges with Cheyenne Basin

Colton (1978); Erslev and Selvig (1997); Erslev et al. (1999); Cole and Braddock (2009) 

East-vergent, moderate-angle thrusts; asymmetric basin subsidence

Laramie or Lance Fm., Maastrichtian, 100–450 m, vertebrate fossils

Matuszczak (1976); Kirkham et al. (1980); Brewer et al. (1982)