Laramide Sedimentary Basins and Sediment-Dispersal Systems

Chapter 13

Laramide Sedimentary Basins and Sediment-Dispersal Systems Timothy F. Lawton Centro de Geociencias, Universidad Nacional Autónoma de México, Querétaro, Mexico

Chapter Outline Introduction Laramide Orogeny Basin Distribution and Classification Timing of Laramide Deformation Basin Structure

529 531 531 533 536

Depositional Systems Laramide Sediment-Dispersal Network Tectonics of the Laramide Orogeny Acknowledgments References

538 544 548 550 550

INTRODUCTION Intermontane broken-foreland basins that developed during the Laramide orogeny in western North America succeeded a widespread Cretaceous foreland basin system linked to the Cordilleran fold and thrust belt (Dickinson et  al., 1988; DeCelles, 2004). The Laramide basins were closely yoked to basement-cored uplifts that partitioned the continental crust beneath the former foreland basin and changed the paleogeography of the Rocky Mountain region from an extensive epicontinental seaway (Chapter 9) to an assemblage of basins occupied by continental depositional environments that received detritus from adjacent uplifted blocks and from rivers that extended beyond the limits of the broken foreland. Depositional environments included alluvial fans, meandering and braided rivers with extensive floodplains, and lacustrine systems, which varied in time and space from lacustrine deltas to lacustrine-margin shoals, intermittently exposed carbonate mudflats, open lacustrine environments, and saline playa settings. Basin development and basement-involved uplift began late in the Cretaceous (Campanian) and continued into the middle to late Eocene. Basement deformation began earlier south of the Colorado Plateau than in the northern part of the broken foreland, and the end of deformation was also spatially diachronous (e.g., Copeland et al., 2017). The deformed province of Laramide basins and uplifts lay east of the former Sevier orogenic belt, which consisted primarily of thin-skinned thrust sheets composed mainly of Proterozoic, Paleozoic, and Mesozoic sedimentary strata, although basement was involved in the hinterland of the Sevier belt (Fig. 1; Camilleri et al., 1997; Chapter 11) and at a basement ramp at the outer limit of the Neoproterozoic-Paleozoic passive margin in northern Utah (Yonkee and Weil, 2015). West of the thrust belt, a region of high, altiplano-like topography evolved during the late development of the Sevier orogenic belt (McQuarrie and Chase, 2000; DeCelles, 2004). An elongate, margin-parallel magmatic arc along the western Cordilleran margin in California became quiescent in the Late Cretaceous, ~85 Ma, probably as a result of low-angle subduction (Dickinson and Snyder, 1978; Bird, 1998; Ducea, 2001; DeCelles et al., 2009). In contrast, arc magmatism continued into the Campanian adjacent to the northern part of the Laramide orogen, as recorded by the Idaho batholith (Fig. 1; e.g., Chetel et al., 2011; May et al., 2013b). In the southernmost part of the Laramide deformed province along the United States–Mexican border between California, Arizona and New Mexico, and Sonora, uplift and basin development were accompanied by arc magmatism that ultimately spread across the southern Laramide province (Coney and Reynolds, 1977; Damon et al., 1981; McDowell et al., 2001; Seager, 2004; González-León et al., 2011), where the magmatic region has been termed the Laramide copper porphyry province (Pecha et al., 2018 and sources therein). Campanian-Maastrichtian intra-orogen magmatism took place in the Colorado mineral belt (Chapin, 2012) and the southwest New Mexico volcanic field, which lay adjacent to the Laramide copper porphyry province, apparently lacked a migratory character, and took place coeval with crustal shortening and basin formation (Fig. 1; McMillan, 2004; Seager, 2004; The Sedimentary Basins of the United States and Canada. https://doi.org/10.1016/B978-0-444-63895-3.00013-9 © 2019 Elsevier B.V. All rights reserved.

529

FIG. 1  Distribution of Laramide sedimentary basins and uplifts formed between Late Cretaceous and late Eocene time. Classification of sedimentary basins follows that of Dickinson et al. (1988) except as noted in text. Sedimentary basin abbreviations: BM, Bull Mountain; Ca, Cabullona; C-LJ, Carthage-La Joya; CM, Crazy Mountains; E, El Rito; EC, El Chanate; EP, Echo Park; ET, El Tuli; FC, Fort Crittenden; G, Galisteo; GR, greater Green River, consisting of four subbasins (GR-B, Bridger; GR-Gd, Great Divide; GR-Sw, Sand Wash; GR-W, Washakie); H, Hanna; Ka, Kaiparowits; Kl, Klondike; La, Laramie; LHT, Little Hat Top; LR, Love Ranch; MP, Middle Park; MV, Monte Vista; NP, North Park; P, Potrillo; Pic, Piceance Creek; R, Ringbone; Ru, Rucker; Sh, Shirley; SB, Sierra Blanca; SP, South Park; Wind R, Wind River. Basement uplift abbreviations (italic font): BH, Black Hills; Bi, Bighorn; B, Beartooth; CC, Circle Cliffs; CT, Coconino terrace; D, Defiance; DC, Douglas Creek arch; FR, Front Range; GM, Granite Mountains; K, Kingman; Kb, Kaibab; Hi, Hidalgo; La, Laramie; M, Monument; Mz, Montezuma; N, Nacimiento; OC, Owl Creek; RS, Rock Springs; Saw, Sawatch; SdC, Sangre de Cristo; SJ, San Juan; SL, San Luis; SM, Sierra Madre; SR, San Rafael; Un, Uncompahgre; Wh, White River; WM, Wet Mountains; WR, Wind River; Z, Zuni. Sources of data: Dickinson et al. (1988); Fouch (1976); Goldstrand (1994); Jacques-Ayala (1999); Cather (2004); Seager (2004); Rasmussen and Foreman (2017). Batholiths of Cordilleran magmatic arc: BB, Boulder batholith; IB, Idaho batholith; PRB, Peninsular Ranges batholith; SNB, Sierra Nevada batholith. LPCP, Laramide porphyry copper province segment of magmatic arc. Intraforeland volcanic fields: AVF, Absaroka volcanic field; CMB, Colorado mineral belt; SWNM, southwestern New Mexico volcanic field (named here). Hinterland volcanic field: CVF, Challis volcanic field. Thick dashed lines bound Precambrian crustal provinces, which in combination with local intrusive bodies, indicate potential sources of detrital zircons in basin fill (adapted from Anderson, 1989; Karlstrom et al., 2004; Thomas et al., 2016). Locations of structure sections in Fig. 3 indicated by white circles. See text for further references.



Laramide Sedimentary Basins and Sediment-Dispersal Systems Chapter | 13  531

Clinkscales and Lawton, 2015; Amato et al., 2017). Therefore, although the California part of the arc was extinguished during the Laramide orogeny, Laramide deformation in the central Rocky Mountain region of the United States was not entirely amagmatic, an important consideration for detrital-zircon provenance studies of the Laramide orogen.

LARAMIDE OROGENY The Laramide orogeny has been defined from both temporal and kinematic criteria, which has led to some confusion as to the meaning of the term. The event is commonly defined in a temporal context to encompass a basement-involved style of deformation that took place from about 75 to 40 Ma (Coney, 1976). Because deformation and uplift swept inboard from the southwest (Heller and Liu, 2016), and then migrated back toward the continental margin (e.g., Fan and Carrapa, 2014; Smith et al., 2014), Copeland et al. (2017) pointed out that the temporal duration of orogeny was significantly less (as little as 10 m.y.) at the distal northeastern edge of the Laramide province than elsewhere in the orogen. A temporal definition is further rendered imprecise by complete temporal overlap of thin-skinned and thick-skinned deformation and by examples of basement deformation that preceded the specified time frame. Basement-involved deformation in the foreland began in Montana as early as the Early Cretaceous, prior to Late Cretaceous thin-skinned deformation (DeCelles, 1986; Perry et al., 1988; Schmidt et al., 1988). In central Utah, thin-skinned deformation near the front of the thrust belt continued until middle to late Eocene and thus temporally overlapped the main phase of basement deformation that occurred farther to the east (Lawton and Trexler, 1991; Lawton et al., 1993b). Moreover, basement-involved disruption of the orogenic foreland did not occur in the Canadian Cordillera; therefore, the temporal definition of the Laramide orogeny is an artificial construct in Canada. The term Laramide is widely used for time-equivalent deformation in Mexico; nevertheless, the Mexican orogen, accompanied by continental-arc magmatism throughout its history, encompasses a wide variety of structural styles (Fitz-Díaz et al., 2018). The term Laramide in Mexico is thus a purely temporal construct. For purposes of this chapter, the Laramide orogeny is defined on the basis of a specific basement-involved deformation style. The orogen encompasses a region of basement-cored uplifts and adjacent intermontane basins that formed in Late Cretaceous to Paleogene time (Fig. 1). This province of uplifts and basins encircles the less-deformed Colorado Plateau on its south, east, and north sides and includes basins formed on the western edge of the plateau between the thrust belt and the monocline-flanked uplifts of the plateau. The Laramide orogen extends northward from northern Sonora, Mexico, about latitude 31°10′ North, to central Montana, near latitude 46°30′ North. It does not extend into Canada, nor does it continue into the Sierra Madre Oriental of Mexico, where variable deformation style was determined by preexisting basement structure and physical heterogeneity of the overlying Mesozoic stratigraphic section (Fitz-Díaz et  al., 2018). The most familiar part of the Laramide province lies in the central Rocky Mountains of the United States, north and east of the Colorado Plateau (Fig. 1), but roughly coeval basins of similar structural style also lie south of the Colorado Plateau, in the Basin-Range Province, which was created by Neogene extension. The southern basins are not as well understood as their northern counterparts as a result of superposed younger deformation, but their presence and distribution are important to a full understanding of the geodynamic setting of Laramide deformation and basins.

BASIN DISTRIBUTION AND CLASSIFICATION Laramide basins are widely distributed on Archean crust of Wyoming and southern Montana, and they flank the relatively rigid block of the Colorado Plateau (Fig. 1). Basins on Archean basement north of the Colorado Plateau include the greater Green River, Hanna, Wind River, Laramie, Big Horn, Powder River, Crazy Mountains, Shirley, and Bull Mountain basins (Fig. 1). Several large basins lie on the high-standing thin lithosphere of the Colorado Plateau: (1) the Uinta and Piceance Creek basins, separated by the north-trending Douglas Creek arch, lie at the northern edge of the Colorado Plateau, where they are flanked by a major reverse fault system on the southern edge of the Uinta uplift; (2) the San Juan basin, flanked by a system of monoclines on its northern and western edges and an oblique-slip reverse fault on its eastern edge (Woodward et al., 1972; Cather, 2004); and (3) the Baca basin, an Eocene depocenter formed late in the history of Laramide deformation, at the southern extreme of the plateau. Two basins, partly dismembered by Cenozoic extension, lie along the western edge of the Colorado Plateau. Termed the Flagstaff and Table Cliff basins, they lie between the thrust belt and the San Rafael and Circle Cliffs uplifts, respectively (e.g., Stanley and Collinson, 1979; Goldstrand, 1994; Larsen et al., 2010). Although depicted in Fig. 1 as two discrete basins, they are separated by the Oligocene-Miocene Marysvale volcanic field and may represent a single structural basin (Hintze, 1988). At least eight basins lie south of the Colorado Plateau on Paleoproterozoic to Mesoproterozoic basement of the Yavapai-Mazatzal basement province (~1.8–1.6 Ga), which is intruded by ~1.4 Ga granitoids (Anderson, 1989; Karlstrom et al., 2004). These elongate basins typically trend west-northwest, with lengths of several tens of kilometers and strike-perpendicular widths of only a few kilometers to a few tens of kilometers. These include the Potrillo, Klondike, Ringbone-Skunk Ranch, Little Hat Top, Rucker, Cabullona-Fort Crittenden, El Tuli, and

532  The Sedimentary Basins of the United States and Canada

El Chanate basins. The Rucker and Little Hat Top basins may have once been contiguous, as suggested by their alongstrike positions (e.g., Seager and Mack, 1986; Mann, 1995). Two major basins, the Raton and Denver basins, lie east of the Colorado Plateau, also on Yavapai-Mazatzal basement. A string of small, narrow basins lies along the eastern edge of the Colorado Plateau (Chapin and Cather, 1981; Cather, 2004). These include the Carthage-La Joya, Galisteo, El Rito, Monte Vista, Huerfano Park, Echo Park, South Park, and North Park basins. Laramide basins of the US Cordillera are classified in general as “broken-retroforeland basins” (Ingersoll, 2012), but classification of basins within the array remains an unsettled topic (Table 1). Chapin and Cather (1981) classified the basins according to a system that emphasized basin structure and morphology as well as patterns of deposition within the basin. Dickinson et al. (1988) utilized a classification that combined paleogeographic distribution and facies characteristics of

TABLE 1  Laramide Basin Classifications Basin

Dickinson et al.

Chapin & Cather

This Chapter

Baca

NA

Green River-type

Ponded

Bighorn

Ponded

Green River-type

Ponded

Bull Mountain

Perimeter

NA

Perimeter

Cabullona

NA

NA

Southern Perimeter

Carthage-La Joya

NA

Echo Park-type

Axial

Crazy Mountains

Perimeter

NA

Perimeter

Cutter Sag

NA

Echo Park-type

Perimeter

Denver

Perimeter

Denver-type

Perimeter

Echo Park

NA

Echo Park-type

Axial

El Chanate

NA

NA

Southern Perimeter

El Tule

NA

NA

Southern Perimeter

Flagstaff

NA

NA

Ponded

Galisteo

Axial

Echo Park-type

Axial

Green River (greater)

Ponded

Green River-type

Ponded

Hanna

Axial

Green River-type

Axial

Kaiparowits

NA

NA

Ponded

Huerfano Park

Axial

Echo Park-type

Axial

Klondike

NA

NA

Southern Perimeter

Laramie

Perimeter

Green River-type

Perimeter

Little Hat Top

NA

NA

Southern Perimeter

Love Ranch

NA

NA

Perimeter

Montezuma

NA

NA

Southern Perimeter

North Park-Middle Park

Axial

Echo Park-type

Axial

Piceance Creek

Ponded

Green River-type

Ponded

Potrillo

NA

NA

Southern Perimeter

Powder River

Perimeter

Green River-type

Perimeter

Raton

Perimeter

Denver-type

Perimeter

Ringbone-Skunk Ranch

NA

NA

Southern Perimeter

Rucker

NA

NA

Southern Perimeter

Laramide Sedimentary Basins and Sediment-Dispersal Systems Chapter | 13  533



TABLE 1  Laramide Basin Classifications—cont’d Basin

Dickinson et al.

Chapin & Cather

This Chapter

San Juan

Perimeter

Green River-type

Perimeter

San Luis

NA

Echo Park-type

Axial

Shirley

Axial

Green River-type

Axial

Sierra Blanca

NA

Denver-type

Perimeter

South Park

Axial

Echo Park-type

Axial

Table Cliff

NA

NA

Ponded

Uinta

Ponded

Green River-type

Ponded

Upper McCoy

NA

NA

Southern Perimeter?

Wind River

Ponded

Green River-type

Ponded

basins within the Rocky Mountain region. In that scheme, the basin array consists of perimeter basins situated around the edge of the broken foreland and typically with persistent fluvial outlets to the exterior of the orogen, ponded basins that contained long-lived lake systems in the central part of the orogen, and axial basins that are narrow basins with significant structural relief that generally lie between the perimeter and ponded basins. The basins illustrated in Fig. 1 are classified according to the scheme of Dickinson et al. (1988), in part because that classification conforms well with their roles in Laramide sediment dispersal, as described later. Laramide sedimentary basins south of the Colorado Plateau were not considered in either classification scheme. They are classified here as southern perimeter basins because, although of smaller scale, they have structural characteristics similar to those of the classic perimeter basins (Lawton, 2000; Bayona and Lawton, 2003; Clinkscales and Lawton, 2015, 2017). The scale of the uplift-basin pairs in the southern Laramide province, although similar in structural style to analogous pairs in the Rocky Mountains, was controlled by spacing of antecedent normal faults, with the result that they have a strike-perpendicular dimension roughly half that of the northern Rocky Mountain basins (Clinkscales and Lawton, 2017).

TIMING OF LARAMIDE DEFORMATION Foreland partitioning by the rise of basement uplifts generally spanned latest Cretaceous and Paleogene time, roughly 76–40 Ma. Criteria for recognition of Laramide deformation vary and include localized increased subsidence rates in the Late Cretaceous, onset of differential subsidence within the former Western Interior basin, reorganization of paleodispersal systems and appearance of lakes within the foreland basin system, syndepositional deformation of growth strata, apatite ­fission-track cooling ages, and paleoelevation inferences from stable isotopes in the uplifted blocks. Because the initial rocks eroded from nascent foreland uplifts included poorly lithified shale and sandstone of the foreland basin system that was being structurally partitioned, incipient uplift is particularly difficult to recognize from changes in basin-fill petrography. Recycling of Mesozoic sedimentary sandstone from incipient uplifts also complicates interpretation of detrital-zircon data sets. Moreover, subsidence mechanisms that accommodated strata near the end of the Cretaceous likely included dynamic mantle effects related to emplacement of a shallow slab with oceanic plateau beneath the Laramide foreland (Liu et al., 2010; Heller and Liu, 2016). Although some disagreement exists as to the initial age of Laramide deformation, integrated studies of basin architecture and growth strata, geochronology, sandstone and conglomerate provenance, fluvial paleocurrents, and thermochronology combine to indicate a general picture of northeastward migration of the Laramide deformation front (e.g., Heller and Liu, 2016), followed by southwestward migration of crustal uplift and basin subsidence patterns (Fan and Carrapa, 2014; Smith et al., 2014; Copeland et al., 2017). Earliest deformation took place in the southern basins, where synorogenic strata derived from adjacent uplifts are ­generally Campanian in age and may locally range back into the Santonian. The Fort Crittenden Formation in southern Arizona contains Santonian-Campanian dinosaur fossils and is overlain by volcanic rocks dated near 75 Ma (Dickinson et  al., 1989); the Cabullona Group of the Cabullona basin in northernmost Sonora is Campanian-Maastrichtian (~82–70 Ma; González-León and Lawton, 1995; González-León et al., 2017). The synorogenic Ringbone and Skunk Ranch formations of southwestern New Mexico range in age from late Campanian to early Maastrichtian (~75–70 Ma) on the basis of dinosaur fossils and interbedded tuffs (Lucas et al., 1990; Lawton et al., 1993a; Jennings et al., 2013; Clinkscales and

534  The Sedimentary Basins of the United States and Canada

Lawton, 2015) and interfinger with syndeformational andesites with latest Campanian-early Maastrichtian 40Ar/39Ar ages (72.4–71.5 Ma, Young et al., 2000; legacy ages recalculated in Clinkscales and Lawton, 2017). Similarly, andesitic breccias of the Tarahumara Formation, dated at 71 Ma, overlie the El Chanate Group in northeastern Sonora (Jacques-Ayala, 1999). In the former two cases, the andesites overlie synorogenic strata that record early Laramide basin development and are themselves folded, indicating that arc volcanism both postdated and accompanied crustal deformation. Ages from the upper part of the thick McCoy Mountains Formation in southern California and western Arizona indicate broadly coeval deposition adjacent to basement uplifts (Tosdal and Stone, 1994; Dickinson and Lawton, 2001; Barth et al., 2004). Bracketing stratigraphic ages of the basin fill of the McCoy basin are based on young detrital zircon grains that yield maximum depositional ages of ~93–84 Ma in the upper part of the section and postdepositional intrusion by a 73.5 Ma granodiorite (Barth et al., 2004). A tuff age in the upper part of the McCoy Mountains Formation is 79 ± 2 Ma, which provides the only direct age on the formation, whose type section is about 7.5 km thick (Tosdal and Stone, 1994). Jacques-Ayala et al. (2009) suggested a correlation of the McCoy Mountains Formation with the El Chanate Group of Sonora. Ages of deformation in the Colorado Plateau indicate onset of foreland partitioning near the end of the Campanian. In southwestern Utah, movement on the Cottonwood monocline flanking the northern part of the Kaibab uplift began in middle Campanian time and affected thickness patterns of the uppermost part of the Wahweap Formation (Fig. 2; Tindall et  al., 2011). The Circle Cliffs uplift began to pond drainage systems of the Canaan Peak Formation at the end of the Campanian (Goldstrand, 1994). The uplifted flanks of the Table Cliff basin, a small depocenter north of the Kaibab uplift, became local sources for sediment of the Maastrichtian Canaan Peak and Pine Hollow formations (Larsen et al., 2010), after which the basin expanded to host extensive fluvial-palustrine-lacustrine facies recorded by the Claron Formation in the late Paleocene-middle Eocene (Goldstrand, 1992, 1994; Biek et al., 2014; Eaton et al., 2018). The San Rafael uplift blocked long, northeast-draining rivers, created growth strata in the upper part of the Mesaverde Group in the late Campanian, and ponded Maastrichtian lacustrine strata east of the thrust belt (Lawton, 1983, 1986). This lacustrine system, which grew to become Paleocene Lake Uinta, occupied a southwestward extension of the Uinta basin, termed the Flagstaff basin (Fig. 1). Lacustrine ponding of the Uinta basin proper, signaled by deposition of the Flagstaff Member of the Green River Formation (Fig. 2; Fouch, 1976; Ryder et al., 1976), indicates significant growth of the Uinta uplift by the middle to late Paleocene, although early growth is recorded by deposition of the upper Campanian Ericson Formation in the Green River basin (Leary et al., 2015). Cather (2004) proposed a three-stage subsidence history for the San Juan and Raton basins in northern New Mexico and southern Colorado. His early phase, ~80–75 Ma, is based on increased subsidence rates in stratal successions long considered as part of the Western Interior basin. The regions of increased subsidence rates, particularly in the San Juan basin, do not conform well with the long-term Laramide architecture of the basin and require changing depocenters within the basins through time (Cather, 2004). Coeval increased rates of subsidence in the Western Interior basin are recorded by eastward expansion of isopachs in New Mexico, Utah, and Colorado, and have been attributed to the isostatic effects of a flat Farallon slab (Cross, 1986) or dynamic effects created by decreased slab angle that preceded Laramide deformation (Lawton, 1994; Nummedal, 2004; Heller and Liu, 2016). More recent modeling of the dynamic effects of a subducted oceanic plateau beneath North America indicates broad subsidence 90–80 Ma, coincident with development of the Western Interior seaway, followed by northeastward migration of a wave of uplift that tracked the onset of foreland crustal deformation (Heller and Liu, 2016). Therefore, subsidence in the Colorado Plateau region at 80 Ma is best considered pre-Laramide by the definition of the orogen proposed here. Fission-track cooling ages and apatite (UTh)/He thermochronology suggest somewhat older basement uplift ages, at about 65–60 Ma in the northern Rocky Mountains, than are indicated by local stratigraphic patterns. Assessment of uplift onset by thermochronologic methods is rendered difficult by a low Phanerozoic geothermal gradient in the Rocky Mountain region and possible refrigeration effects of a low-angle subducted slab, both of which could result in failure of both zircon and apatite tracks to fully anneal, or failure of basement rocks to exceed the He closure temperature, prior to Laramide uplift (e.g., Dumitru et al., 1991; Cerveny and Steidtmann, 1993; Kelley and Chapin, 1995; Crowley et al., 2002). Apatite fission-track ages suggest cooling in the Wind River Range as early as 75 Ma, with most rapid cooling between 60 and 57 Ma (Cerveny and Steidtmann, 1993). Stratigraphic arguments suggest initial uplift of the range at about 70 Ma (Dickinson et al., 1988). Facies patterns and conglomeratic strata in the Wind River basin indicate rapid unroofing of the Washakie and Owl Creek uplifts somewhat later, in the interval 55.5–54.5 Ma (Fan et al., 2011). Apatite fissiontrack ages from the Front Range and Wet Mountains in Colorado indicate cooling in the period 67–57 Ma (Bryant and Naeser, 1980; Kelley and Chapin, 1995). Similarly, apatite (UTh)/He thermochronology suggests cooling of the Bighorn Mountains at 65 ± 5 Ma (Crowley et al., 2002), whereas the Beartooth Conglomerate, equivalent to part of the Paleocene Fort Union Formation, suggests a slightly younger age for uplift of the Beartooth uplift on the west flank of the Bighorn basin (DeCelles et al., 1991a, b). Thus, thermochronometers may record arrival of the shallow slab beneath the foreland

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FIG. 2  Correlation chart of stratigraphic units in selected Laramide basins. Formations designated as foreland basin strata underlie strata of Laramide basins but predate foreland partitioning and formerly extended beyond bounds of Laramide structural basins. Dominant facies of a particular formation indicated; most coarsen toward basin margins into conglomeratic strata or contain intervals of other facies types. Sources: Baca basin: Cather (2004), Prothero et al. (2004); Bighorn basin: Keefer (1965), May et al. (2013b); Galisteo basin: Gorham and Ingersoll (1979), Lucas et al. (1997); Greater Green River basin: Surdam and Stanley (1980), Machlus et al. (2002), Smith et al. (2003, 2008), Leary et al., 2015); McCoy basin: Tosdal and Stone (1994), Barth et al. (2004); Piceance Creek basin: Smith et al. (2008), Heller et al. (2013), Foreman and Rasmussen (2016); Powder River basin: Keefer (1965), Smith et al. (2008), Fan et al. (2011); Raton and San Juan ­basins: Cather (2004), Cather et al. (2008), Wegert and Parker (2011), Bush et al. (2017); southern perimeter basins (Arizona, New Mexico and Sonora): Seager (1983), Lucas et al. (1990), Dickinson et al. (1989), Lawton et al. (1993b), Seager et al. (1997), Jacques-Ayala (1999), Dickinson and Lawton (2001), Jennings et al. (2013), Clinkscales and Lawton (2015), Amato et al. (2017); Table Cliff basin: Christensen (2005), Larsen et al. (2010), Roberts et al. (2013), Lawton et al. (2014), Biek et al. (2014), Albright and Titus (2016), Eaton et al. (2018); Uinta basin: Ryder et al. (1976), Johnson (1985), Bryant et al. (1989), Machlus et al. (2002); Balls et al. (2004), Lawton and Bradford (2011). Timescale: ICS, 2016 (Cohen et al., 2013, updated); North American Land Mammal Age (NALMA) age boundaries from Smith et al. (2008).

536  The Sedimentary Basins of the United States and Canada

prior to crustal ­deformation related to the slab (e.g., Dumitru et al., 1991) and may therefore be less confident indicators of early deformation than sedimentation patterns in the adjacent basins. The end of Laramide deformation is signaled by unconformities cutting deformed strata and generally overlain by upper Eocene-Oligocene volcanic and volcaniclastic strata (Dickinson et  al., 1988; Cather, 2004). Following Laramide deformation, a low-relief erosion surface that truncates Laramide structures was established in Colorado and northeastern Utah. In Colorado, this surface is overlain by a 37–36 Ma ignimbrite (Epis and Chapin, 1975). North and south of the Uinta Mountains, pediments, termed the Gilbert Peak erosion surface, slope away from the Uinta uplift to north and south (Hansen, 1984). The surface is overlain by the Bishop Conglomerate, with a rhyolite tuff near its base that has yielded single-crystal 40Ar/39Ar ages of 34.46 ± 0.26 to 33.73 ± 0.11 Ma (late Eocene-earliest Oligocene; Balls et al., 2004). In central New Mexico, the Laramide orogeny ended in late Eocene time (~38 Ma) when deposition of the Baca Formation, which is middle Duchesnean in its upper part, was succeeded by emplacement of intermediate volcanic and volcaniclastic rocks of the lower Datil Group (Fig. 2; Cather et al., 1987; Cather, 2004; Prothero et al., 2004). Deformation in southern New Mexico was postdated by deposition of middle Eocene andesite of the Rubio Peak Formation with a fission-track age of 41.7 Ma (Thorman and Drewes, 1980). East-west normal faulting in southwesternmost New Mexico, dated by syndeformation intrusions ranging ~34–27 Ma, records extension that postdated Laramide shortening (Clinkscales and Lawton, 2017). A complicated pattern of latest deformation and basin formation therefore suggests that youngest Laramide deformation was diachronous and underwent retrograde southwestward or westward migration (Dickinson et al., 1988; Fan and Carrapa, 2014; Smith et al., 2014). Copeland et al. (2017) concluded that terminal Laramide deformation migrated to the southwest from ~55 to 45 Ma, possibly to end as late as ~30 Ma in the southwestern part of the orogen, after reaching its maximum inboard extent at the Black Hills uplift at c.60 Ma (Lisenbee and DeWitt, 1993).

BASIN STRUCTURE Laramide basins are typically paired with basement-cored uplifts. Some basins, axial basins in particular, are flanked by more than one uplift, but typically thicken toward the overthrust margin, although exceptions to this structure exist in central New Mexico in the Rio Grande rift area (Yin and Ingersoll, 1997). The mechanism of basement uplift, controversial for many years, is now generally considered to be crustal shortening and displacement of uplifted blocks along steep to moderately dipping faults that root into basement (Fig. 3A and B; e.g., Berg, 1962; Smithson et al. 1978; Erslev and Koenig, 2009). In the Rocky Mountain region, most basement faults dip moderately and have displacements exceeding 10 km. For example, the well-imaged Wind River thrust, which emplaces crystalline basement of the Wind River Mountains over the Green River basin to the south, has 13 km of vertical separation, at least 26 km of horizontal separation, and penetrates 25–30 km into the crust as a discrete fault zone (Allmendinger et al., 1985). The thrust dips 35–40 degrees throughout its imaged depth and is planar, with no sign of steepening or flattening. Faults with similar thrust geometries have been widely drilled and imaged seismically in the Rocky Mountain region and provide abundant evidence for crustal shortening (Gries, 1983a). The location and vergence of the Laramide thrusts in the Rocky Mountain region have been attributed to reactivation of Proterozoic and late Paleozoic extensional faults (Marshak et al., 2000); if this is the case, the Laramide displacement on the faults greatly exceeded older normal displacement in order to emplace basement over the Phanerozoic section. Multiphase Proterozoic and late Paleozoic faults reactivated by Laramide shortening have been documented in northern New Mexico (Yin and Ingersoll, 1997). Basement-fault dips adjacent to southern basins range from moderate to steep and thus have similarities with faults that separate Rocky Mountain uplifts and basins. Basin-bounding faults generally trend west-northwest, dip steeply beneath the uplifts, and verge to both northeast and southwest (Keith, 1982; Seager, 1983; Seager and Mack, 1986; Lawton, 2000; Clinkscales and Lawton, 2017). Although long considered to have a major component of sinistral offset (Seager, 1983; Hodgson, 2000), some steep faults having apparent left steps are simply displaced by younger normal faults, and thus transcurrent movement is not necessary to explain Laramide fault kinematics (Clinkscales and Lawton, 2015, 2017). Basement-involved faults in the southern Laramide province commonly separate Jurassic-Lower Cretaceous strata of markedly different thickness, indicating that they reactivated extensional faults of mid-Mesozoic age (Lawton, 2000; Bayona and Lawton, 2003; Clinkscales and Lawton, 2015). In some cases, Laramide basins lie adjacent to the thickest parts of former rift basins (Fig. 3C and D), offset from the rift basins across reactivated faults, which controlled trends and locations of Laramide uplifts and basins. In other cases, there is significant inferred overlap of Laramide sections and precursor rift basins (Clinkscales and Lawton, 2017). Laramide basins are asymmetric and thicken toward active, uplifted basin margins (Fig. 4A and B). This pattern of asymmetry is present in the Rocky Mountain and southern Laramide basins, and to a lesser degree, in axial basins on the east flank of the Colorado Plateau (Chapin and Cather, 1981). Maximum structural relief on basement of 5–10 km is usually expressed

Laramide Sedimentary Basins and Sediment-Dispersal Systems Chapter | 13  537



West Elevation (m)

East

Beartooth uplift

5000

Bighorn basin

J-K C

Kl Km Kmv Kc

Tb

0

J-K

pC

Tr

J-K O-P pC

C

–5000

(A) Blue Mountain Anticline North Jn Elevation Kd Kf (ft) Tr Ju 5000 Pw Pp Pm MD C 0

Uinta basin

Kmc

Jn

Pp

Pm

Thunderbolt ridge

SSW Elevation (ft) 7000 Cave Creek

Kmv

pC

–5000

South

6000

Ju

7000 Lava Flow Qal

Qal

Tr Pw

NNE

Lava flow Jg

Jg

Jce Jcs

Jcs Km

5000

MD C

Pz

Jct

Jct

Jcu?

4000

(B)

5000

Pz

Jo

pC

6000

4000

(C) Late Cretaceous Laramide basin Ku

Late Cretaceous Laramide basin Ku

Kl

Kl

Kl

basement horse

Ju-Kl Pz

Ju-Kl

Pz

Pz

null point pC

Extensional basin formation (Late Jurassic-Early Cretaceous)

Ju-Kl

Ju-Kl Pz

Pz

pC

T A

Shortening by dip-slip (Late Cretaceous)

pC

Shortening by oblique-slip (Late Cretaceous)

(D) FIG.  3  Models for Laramide basin-margin structure. (A) Structure of east flank of Beartooth uplift, developed during deposition of Beartooth Conglomerate (Tb). (B) Fold-thrust model, northern flank of Piceance Creek basin, Colorado. (C) South flank of inverted Jurassic basin, Chiricahua Mountains, Arizona. This is the northern margin of the Rucker basin (Fig. 1). Jurassic strata are not demonstrated to lie south of the reverse fault (Lawton and Olmstead, 1995). (D) Model for active margin of some southern perimeter basins (Rucker and Fort Crittenden basins) in southern New Mexico and Arizona. Laramide basins overlie thin stratigraphic sections developed during a Late Jurassic-Early Cretaceous phase of extensional basin formation, but are generally absent over thick sections of Upper Jurassic–Lower Cretaceous strata. ((A) Redrawn from DeCelles et al. (1991a,b). (B) Redrawn from Berg (1962). (D) After Lawton (2000).)

at major basin-bounding faults. The lateral dimension of basin-uplift pairs of southern perimeter basins is about one-half that of basins in the northern Rocky Mountains, but structural relief is similar (Fig. 4C; Clinkscales and Lawton, 2017). Some axial basins of northern New Mexico and southern Colorado have contrasting structure, in which Laramide strata onlap basinmargin uplifts, which are bounded by reverse faults that dip beneath the basins (Yin and Ingersoll, 1997). Two-dimensional modeling suggests that the dominant subsidence mechanism in Laramide basins was a flexural response to loading by basinmargin uplifts and syndeformational sedimentary strata (Hagen et al., 1985). The modeled flexural profiles are consistent with bending of strong, cold lithosphere with a flexural rigidity between 1021 and 1023 Nm (Hagen et al., 1985).

538  The Sedimentary Basins of the United States and Canada

FIG. 4  Laramide basin cross-sections, showing asymmetry of fill. (A) Reconstructed cross-section of Hidalgo uplift and Ringbone basin of southwestern New Mexico. Pre-Laramide rocks: pC, Proterozoic granite; Pzl, lower Paleozoic (Ordovician-Mississippian) strata; Pzu, upper Paleozoic (PennsylvanianPermian) strata; Ju, Upper Jurassic strata; Kl, Lower Cretaceous strata. Laramide deposits: Kr, Ringbone Formation; Ks, Skunk Ranch Formation; Kh, Hidalgo Volcanic Formation (equivalent to Skunk Ranch Formation, see Fig.  2). (B) North-south section across northern part of Bighorn basin and Beartooth uplift. Tf: Fort Union Formation. Note vertical exaggeration. (C) North-south section from southern part of Bighorn basin across Owl Creek uplift to Wind River basin and Granite Mountains uplift. Explanation: pC, Precambrian rocks; Pz, Paleozoic strata; Tr-K, pre-Meeteetse Mesozoic strata; Kml, Meeteetse Formation and Lewis Shale; Kl, Lance Formation; Tf, Fort Union Formation; Ti, Indian Meadows Formation; Twi, Willwood Formation; Twr, Wind River Formation. ((A) Adapted from Clinkscales and Lawton (2017). (B) From Hagen et al. (1985). (C) Redrawn from Keefer (1965).)

DEPOSITIONAL SYSTEMS Perhaps the singular characteristic of Laramide basins that distinguishes them from the older retroarc foreland basin is their assemblage of wholly continental depositional environments. These range from alluvial-fan, fluvial, deltaic, and lacustrine settings distributed according to distance from primary sediment sources of flanking uplifts. The distribution of the facies types was controlled by location of flanking uplifts and rates of deformation (Fig. 5; Beck et al., 1988) as well as positions of basins within the Laramide broken foreland (Dickinson et al., 1988). Perimeter basins along the eastern flank of the Laramide province contained local lacustrine facies and fluvial facies deposited by river systems that exited the basins to the east. Conglomeratic deposits dominated the axial basins located near the east edge of the Colorado Plateau due to their narrow widths and great structural relief of basin-bounding faults (Chapin and Cather, 1981). Nevertheless, recent provenance analysis has demonstrated that axial-fluvial systems were important in sediment transport in the axial basins (Rasmussen and Foreman, 2017; Sharman et al., 2017). South of the Colorado Plateau, the arc-proximal southern perimeter basins, with comparatively small lateral dimensions, likewise have basin-bounding faults with great displacement that likely created significant topographic relief, suggested by coarse conglomerate and local rock-avalanche deposits that interfinger with siliciclastic lacustrine strata and alternate with axial, basin-parallel fluvial systems (Hayes, 1987; Inman, 1987; González-León and Lawton, 1995; Basabilvazo, 2000; Clinkscales and Lawton, 2015). Ponded basins lying in the interior of the orogen collected runoff from large drainage areas, were connected with adjacent basins via structurally elevated divides, and thus contained large freshwater to saline lakes during middle Paleocene to middle Eocene time when flanking



Laramide Sedimentary Basins and Sediment-Dispersal Systems Chapter | 13  539

FIG. 5  Generalized distribution of depositional environments with respect to basement uplift (after Beck et al., 1988). Numerals correspond to strata deposited during four phases of thrust development proposed by Beck et al. (1988): (1) Early thrusting; (2) rapid thrusting; (3) slow thrusting; (4) postthrusting. Rapid subsidence during the first two phases results in strong asymmetric subsidence and well-defined depositional axis of basin nearer the thrust front than in later two phases. Numerals indicate approximate position of basin depositional axis.

uplifts attained their greatest structural relief (Bradley, 1964; Ryder et al., 1976; Surdam and Stanley, 1979; Johnson, 1985; Fan et al., 2011). As discussed in greater detail later, it is likely that the ponded basins adjacent to the Uinta uplift were the ultimate sinks for most or all sediment transported into them. Depositional systems of Laramide basins fall naturally into four categories: alluvial-fan, fluvial, lacustrine-margin, and open-lacustrine. These systems are recorded by corresponding facies associations arranged more or less concentrically about the structural basins, but strongly influenced by basin asymmetry, paleoslope orientation and direction and distance to source uplifts. The alluvial-fan depositional system constitutes the most source-proximal deposits of Laramide intermontane basins. Strata consist of massive, crudely laminated pebble to boulder conglomerate and subordinate sandstone deposited by ­debris-flow and sheet-flood processes (Andersen and Picard, 1974; Ryder et al., 1976). Rock-avalanche deposits interfinger with lacustrine strata on the active flank of the Ringbone basin of southwestern New Mexico (Clinkscales and Lawton, 2015). Alluvial-fan deposits typically are restricted to regions that lay within a few kilometers of active uplifts (Steidtmann, 1971; Surdam and Stanley, 1980), and commonly preserve internal angular unconformities that record uplift of the adjacent range (Bryant et al., 1989; DeCelles et al., 1991b) or developing intrabasinal structures (Clinkscales and Lawton, 2015). Clast compositions, which match the range of rock types in nearby uplifts, have been used to discriminate individual local uplifts flanking Laramide basins (e.g., Steidtmann, 1971; Fan et al., 2011; Clinkscales and Lawton, 2015). Conglomeratic successions commonly contain inverted-clast stratigraphies that record progressive erosion through the sedimentary cover and into the basement of the uplifts (Hansen, 1984; Graham et al., 1986; Ingersoll et al., 1987; Clemons and Mack, 1988; DeCelles et al., 1991b; Crews and Ethridge, 1993; Cather, 2004). The fluvial depositional system consists of channel-belt and floodplain environments (e.g., Ryder et al., 1976; Bown and Kraus, 1987; Kraus, 1987). Fluvial systems of Laramide basins lay downstream of and adjacent to alluvial-fan environments and upstream of lacustrine margins, or, if lakes were not present, formed axial drainages parallel to the structural axis of the basin (e.g., Steidtmann, 1971; Beck et al., 1988; Rasmussen and Foreman, 2017). Fluvial systems record alluvial plains traversed by braided or meandering channels; overbank settings with extensive paleosols are common in basins with hydrologic outlets (e.g., Willwood Formation of Bighorn Basin; Bown and Kraus, 1987; Kraus, 1987). In basins with well-developed lacustrine systems, alluvial plains varied tremendously in length. Adjacent to active uplifts, alluvial plains formed short links between alluvial fans and lacustrine margins. Short fluvial systems are recorded in the lower Eocene Wasatch Formation of the Sand Wash subbasin, which was derived from the north flank of the Uinta uplift

540  The Sedimentary Basins of the United States and Canada

(Surdam and Stanley, 1980) and the coarse-grained Duchesne River Formation, which was derived from sedimentary rocks of the Uinta uplift directly to the north of the Uinta basin (Andersen and Picard, 1974; Dickinson et al., 1986). During early basin development, the thrust belt to the west continued to be an important sediment source. For example, the North Horn Formation, which includes local lacustrine strata that record initial ponding of eastbound thrust belt-derived drainage systems by the San Rafael uplift, also contains detritus derived from continued uplift of the Sevier orogenic belt (Lawton et al., 1993b). Distributive fluvial systems (sensu Hartley et al., 2010 and Weissmann et al. 2010) were important components of some Laramide basins. Alluvial plains within and adjacent to some ponded basins stretched for hundreds of kilometers and have been reconstructed from their facies distributions or distinctive detrital compositions. The lower Eocene Colton Formation was deposited by a fluvial megafan or distributive fluvial system, only the downstream reach of which is preserved (Jones and Plink-Bjorklund, 2015). Consisting of arkosic detritus, the Colton fluvial fan interfingers with lacustrine facies of the Green River Formation on the south flank of the Uinta basin (Fouch, 1976; Ryder et al., 1976; Dickinson et al., 2012). Formerly considered as derived from basement uplifts in south-central Colorado (Dickinson et al., 1986), detrital-zircon data indicate that the Colton detritus was likely transported from magmatic-arc sources as far away as southern California (Dickinson et al., 2012). The middle Eocene Crazy Hollow Formation, which overlies the Green River Formation of the Flagstaff basin, records delivery of chert-lithic detritus to the southwestern arm of Lake Uinta from southwestern Utah, northwestern Arizona, and southern Nevada (Weiss and Warner, 2001). An almost archetypical distributive fluvial system occupied the Baca basin, which did not begin to form until middle Eocene time (Fig. 1; Cather and Johnson, 1986; Potochnik, 1989). The Mogollon Rim and Eagar formations of Arizona and the correlative Baca Formation of western New Mexico record an extensive alluvial plain that generally fined eastward from braided to meandering rivers with channels that diminished in depth and width downstream. The fluvial fan discharged into a lake of moderate size via deltas recorded by upward-coarsening successions that average 9 m thick (Cather and Johnson, 1986) and likely spilled eastward at times into the axial Carthage-La Joya basin. In New Mexico, alluvial strata overlie a lateritic paleosol developed on Campanian fluvial strata, indicating that a long hiatus preceded basin development. In Arizona, the informal Mogollon Rim formation overlies beveled northeast-vergent monoclines expressed in Paleozoic strata (Potochnik, 1989). Fluvial-sandstone compositions are highly variable, ranging from arkosic to lithic, depending on rock types in nearby uplifts or distant headwaters. Quartz-lithic sandstone with abundant sedimentary grains is common in basins adjacent to uplifts, such as the Uinta uplift, which contains a core of Proterozoic sedimentary strata (Isby and Picard, 1983; Dickinson et al., 1986). Arkosic fluvial sandstone of the Wagon Bed Formation of the Wind River basin indicates a source in exposed basement of the Granite Mountains (Chetel et al., 2011). Volcanic-lithic grains are present in Eocene strata of the Green River, Piceance Creek, and Uinta basins (Surdam and Stanley, 1980; Johnson, 1985). Volcanic-lithic detritus is also present in deposits of axial drainages of southern perimeter basins (Inman, 1987; Basabilvazo, 2000; Dickinson and Lawton, 2001; Barth et al., 2004). There, volcanic grains were derived from Jurassic volcanic rocks in the adjoining uplifts and from time-equivalent volcanic rocks in the Laramide porphyry copper segment of the Cordilleran arc. Ages of detrital zircon in sandstone of local fluvial systems, as with alluvial-fan conglomerate, record unroofing of adjacent source areas (Bush et al., 2017; Pecha et al., 2018). Detrital-zircon data from the McCoy Mountains Formation augment compositional data and offer insight into the nature of basin-margin sources of the McCoy basin. Detrital-zircon ages from the Upper Cretaceous part of the formation are distinctively bimodal (Fig. 6; Barth et al., 2004), with Jurassic-Late Cretaceous grain ages representing detritus from the volcanic arc to the west and southwest, and grains near 1.7 Ga derived from uplifted Proterozoic basement of the basin margins (Fig. 1).

FIG. 6  Detrital-zircon age probability plot for sample from uppermost part of McCoy Mountains Formation (Barth et al., 2004). Abscissa is age in millions of years; peaks represent probability of encountering a particular grain age in the zircon population. Numbers are peak ages in millions of years. Youngest ages in this sample of 84 Ma indicate maximum depositional age of upper part of McCoy Mountains Formation.



Laramide Sedimentary Basins and Sediment-Dispersal Systems Chapter | 13  541

The lacustrine-margin depositional system consists of siliciclastic and carbonate strata that record deltaic, interdeltaic, and littoral environments. These environments shifted laterally tens of kilometers as lakes within these basins expanded and contracted, resulting in complexly interfingering lithofacies (Ryder et al., 1976; Surdam and Stanley, 1979; Johnson, 1985; Smith et al., 2003, 2008). Siliciclastic strata include gray-green calcareous claystone with calcareous nodules and channelform sandstone that accumulated in deltaic and interdeltaic settings. Carbonate strata include grain- and mudsupported rocks deposited on lake-margin strandlines and carbonate flats, respectively. Dolomicrite strata in the Green River Formation represent deposits of exposed carbonate mudflats that were subject to evaporative pumping during lake lowstands (Surdam and Stanley, 1979). Upward-coarsening siliciclastic cycles interpreted as deltaic successions indicate that lake-margin water depths in the greater Green River basin were on the order of 2 m, but uncommon Gilbert delta foresets indicate local water depths of 25 m on the north side of the basin (Surdam and Stanley, 1979). The open-lacustrine depositional system is represented by carbonate mudstone, commonly varved and kerogenous, fossiliferous carbonate wackestone, and claystone with minor sandstone and siltstone. Strata of the open-lacustrine system are generally assigned to the Green River Formation in the Flagstaff, Uinta, Piceance Creek, and greater Green River basins. Varved kerogenous carbonate mudstone, also termed oil shale, was deposited in lake-center environments removed from sources of coarse detrital influx. Evaporitic strata, including bedded trona and halite with interbedded oil shale, are present at several levels of the Green River Formation in the greater Green River basin (Eugster and Hardie, 1975; Surdam and Stanley, 1979; Smoot, 1983; Sullivan, 1985; Pietras et al., 2003a). Nahcolite and halite are interbedded with oil shale in the Uinta basin (Johnson, 1985). Saline conditions are also indicated in carbonate mudstone by the presence of evaporite mineral casts and molds in some horizons and lack of freshwater mollusks (Surdam and Stanley, 1979; Johnson, 1985). Nearshore facies contain abundant mollusks and ostracodes (Ryder et al., 1976; Johnson, 1985). Cycles of kerogen-rich laminated carbonate mudstone bearing evidence for exposure and shallow-water conditions, such as desiccation polygons and flat-pebble intraclast conglomerate, and structureless dolomicrite beds averaging 2 m thick are present in the greater Green River basin (Surdam and Stanley, 1980). The open-lacustrine system has been interpreted as representing perennially deep, salinity stratified lakes (Bradley and Eugster, 1969) or playas (Eugster and Surdam, 1973; Eugster and Hardie, 1975). The former explains the varved carbonate mudstone, whereas the latter explains intervals of evaporitic strata with desiccation features. To reconcile these end-­member hypotheses, Surdam and Stanley (1979) invoked water-depth changes above featureless lake bottoms over which slight changes in water depth resulted in widespread transgressions and regressions. Kerogenous laminated carbonate accumulated as a result of lake-bottom algal productivity (Eugster and Surdam, 1973) during maximum lake highstands, although water depths may not have exceeded 2 m. During lowstands, wide exposed carbonate mudflats were altered to dolomicrite by evaporative pumping. Current models of deposition in the greater Green River basin favor alternations from fluvial-lacustrine conditions deposited during times when sediment and water influx exceeded accommodation (overfilled-lake model) through a fluctuating profundal facies association of kerogenous carbonate mudstone when influx roughly balanced accommodation (balanced-fill-lake model) to an evaporative facies association of linked mudflat-playa settings recorded by evaporite-bearing strata when accommodation exceeded influx (underfilled-lake model; Carroll and Bohacs, 1999). Interaction of Laramide depositional systems and the influence upon them of basin hydrology, tectonics, and longterm climate are illustrated by the lower-middle Eocene Green River Formation and coeval flanking fluvial wedges of the interior ponded basins (Fig. 7). These basins include the greater Green River basin, divided into the Bridger, Great Divide, Washakie, and Sand Wash subbasins, and the Piceance Creek and Uinta basins. Strata of the greater Green River basin thicken westward toward the Wyoming salient of the thrust belt and southward toward the Uinta uplift, both of which were important sediment sources (Surdam and Stanley, 1979; Sullivan, 1985). The north-trending Rock Springs uplift of latest Cretaceous age (Gries, 1983b) partitioned the greater Green River basin. Other Laramide sediment sources included the Granite Mountains and Sierra Madre uplifts on the east side of the basin, the Wind River uplift on the north flank, and beginning in middle Eocene time, the Absaroka and Challis volcanic fields to the north and northwest. The Uinta and Piceance Creek basins thicken northward toward the Uinta uplift and are separated by the north-trending Douglas Creek arch, which influenced thickness trends into the early Eocene (Johnson, 1985). In the Paleocene, the Uinta basin received arc-derived sediment from the south (Davis et al., 2010; Dickinson et al., 2012) and sedimentary detritus from the thrust belt (e.g., Isby and Picard, 1983), whereas the Piceance Creek basin received mainly recycled sedimentary detritus from the Sawatch uplift to the southeast (Foreman and Rasmussen, 2016). In the Eocene, the Uinta Mountains became the primary source for clastic detritus (Andersen and Picard, 1974; Ryder et al., 1976; Dickinson et al., 1986). The Green River Formation represents deposits of Lake Gosiute and Lake Uinta north and south of the Uinta uplift, respectively. In the greater Green River basin, the Green River Formation is divided into four members: the Luman Tongue and Tipton, Wilkins Peak, and Laney members (Fig. 7; Sullivan, 1985; Carroll and Bohacs, 1999; Smith et al., 2003, 2008). These members are lacustrine-margin and open-lacustrine deposits; the Luman and Tipton members record freshwater

542  The Sedimentary Basins of the United States and Canada

FIG. 7  East-west distribution of facies in greater Green River basin. Modified from Smith et al. (2003) with 40Ar/39Ar ages reported in Smith et al. (2008). Other sources of data: Sullivan (1985); Prothero (1996); Murphey et al. (1999).

fluvio-lacustrine conditions that changed near the termination of Tipton deposition to fluctuating profundal conditions, whereas the Wilkins Peak records evaporative mudflat and playa conditions. The Luman and Tipton members correlate with syn- and post-Colton freshwater lacustrine-margin and open-lacustrine deposits of the Uinta basin; the evaporitic Wilkins Peak Member correlates with nahcolite- and halite-bearing hypersaline strata in the lower part of the Parachute Creek Member of the Green River Formation in Piceance Creek basin, but with fresh open-lacustrine facies of the Green River Formation in the Uinta basin (Johnson, 1985; Smith et al., 2008). Single-grain 40Ar/39Ar laser-fusion ages indicate that playa and mudflat evaporative lake conditions persisted for roughly 1.5 m.y., between 51.3 and 49.8 Ma, in the Green River basin (Fig. 7; Smith et al., 2003, 2008). Stratigraphic correlations on the basis of 40Ar/39Ar ages suggest that the transition to evaporative conditions took place later, and persisted longer, in the Uinta basin than in the greater Green River and Piceance Creek basins (Machlus et al., 2002; Smith et al., 2008). An abrupt return to fluctuating profundal conditions north of the Uinta uplift is recorded by the Laney Member, whereas evaporitic conditions continued in Lake Uinta (Smith et al., 2008). The ponded basins north and south of the Uinta uplift thus consistently experienced out-of-phase lacustrine conditions (Carroll et al. 2008; Smith et al., 2008). Individual clastic sedimentary wedges of the Wasatch Formation that impinged upon the margins of Lake Gosiute were derived from all flanks of the basin (Surdam and Stanley, 1979; Sullivan, 1985). The New Fork Tongue, which interfingers with the Tipton Member from the north and northwest, was derived from the thrust-belt reentrant at the northwest corner of the Bridger subbasin, where the Wind River uplift intersects the thrust front, and from the Wind River uplift itself (Sullivan, 1985). Lake Gosiute may have spilled to the northeast during this freshwater overfilled to balanced-fill phase of the lake (Sklenar and Andersen, 1985). Fluvial systems that deposited the Desertion Point Tongue of the Wasatch Formation entered the basin from the southwestern reentrant of the thrust belt; the Desertion Point Tongue interfingers northeastward with the Wilkins Peak Member, whereas the arkosic Cathedral Bluffs Tongue, which interfingers with the Wilkins Peak Member from the east, was derived from basement rocks of the Sierra Madre and Granite Mountains uplifts (Surdam and Stanley, 1980; Sullivan, 1985). A thick wedge of middle Eocene volcanic-lithic detritus prograded southward into the greater Green River basin (Fig. 8; Surdam and Stanley, 1979, 1980). This volcanic-lithic wedge of the Bridger and Washakie formations advanced southward and crossed the east end of the Uinta uplift into the Piceance Creek basin, where it is recorded by the Uinta Formation, which directly overlies the Parachute Creek Member of the Green River Formation. There is some



Laramide Sedimentary Basins and Sediment-Dispersal Systems Chapter | 13  543

FIG. 8  Map of Bighorn, Greater Green River, Piceance Creek, and Uinta basins, showing uplifts that supplied sediment to the basins as discussed in the text. Progradation direction of volcanic-lithic detritus as originally inferred by Surdam and Stanley (1980), with ages of sediment arrival in the basin in bold numbers based upon tuff ages in Fig. 7 and correlations of Fig. 2. Bold numbers in Absaroka volcanic field indicate age range of volcanism (Feeley and Cosca, 2003; Harlan, 2006). Age range of Challis volcanic field from Chetel et al. (2011). Solid arrows are paleocurrent vectors of Surdam and Stanley (1980), open arrows are based on general considerations of possible progradation directions (Surdam and Stanley, 1980; Johnson, 1985). Alternate route of the Idaho River from Challis volcanic field after Carroll et al. (2008) and Chetel et al. (2011). Exposed basement in uplifts depicted in brown. Sedimentary basins: BB, Bridger subbasin; BHB, Bighorn basin; GDB, Great Divide subbasin; HB, Hannah basin; PB, Middle and North Park basins; PCB, Piceance Creek basin; PRB, Powder River basin; SB, Shirley basin; SWB, Sand Wash subbasin.

evidence that the volcanic-lithic wedge subsequently prograded west across the Douglas Creek arch, and thence, along the northern basin axis of the Uinta basin (Surdam and Stanley, 1980; Johnson, 1985), although the amount of volcanogenic detritus that actually reached the Uinta basin may have been negligible (Chetel et al., 2011). Surdam and Stanley (1980), who first documented progradation of the volcanogenic wedge, inferred transport of sediment across the Wind River basin from the time-equivalent Absaroka volcanic field via a gap between the Wind River and Granite Mountains uplifts (Fig. 8). On the basis of 40Ar/39Ar ages of detrital feldspar and bulk lead isotopes of sandstone from the Bridge and Washakie formations, more recent workers have suggested the Challis volcanic field as a more likely source for the volcanic-lithic sediment (Carroll et al., 2008; Chetel et al., 2011). This history of the nuclear ponded basins suggests a complex interplay of climate, tectonism, and sediment supply with a similarly complex mosaic of depositional environments and history of basin evolution. The Green River Formation was deposited during a pronounced warming trend of the Cenozoic extending from 59 to 52 Ma, with a peak at 52–50 Ma, termed the early Eocene climatic optimum (EECO). The EECO was followed by a return to cooler conditions that began abruptly in the interval 50–48 Ma (Zachos et al., 2001). The evaporative facies of the Wilkins Peak and Parachute Creek members thus were deposited during the second half of the EEOC; return to fluctuating profundal conditions took place at the end of the EEOC. It is tempting to infer that drying and subsequent freshening events of Lake Gosiute were slightly

544  The Sedimentary Basins of the United States and Canada

delayed responses to global climatic trends (e.g., Roehler, 1993; Matthews and Perlmutter, 1994). Nevertheless, the diachronous facies shifts in Lakes Gosiute and Uinta argue against regional climate change as the primary driver of lake states, as discussed further in following text. Abundant additional evidence exists for active tectonism and uplift of basin-flanking blocks during changing conditions in Lake Gosiute. Coeval tectonism during Green River deposition is supported by: (1) asymmetric, southward-increasing subsidence in the middle part of the Wilkins Peak Member, documented by detailed correlations along the west side of the Rock Springs uplift (Pietras et al., 2003a); (2) a nearly fourfold increase in sediment accumulation rates between the Tipton and Wilkins Peak members (Smith et al., 2003), either as a result of increased accommodation by structural elevation of the outflow sill of the lake or through increased loading by the adjacent Uinta uplift; (3) inferred coeval uplift of the north flank of the Uinta uplift during deposition of the Tipton Member (Hansen, 1984) and the south flank of the Wind River uplift during deposition of the Wilkins Peak Member (Pietras et al., 2003b); and (4) large sediment influx from the east flank and southwest corner of the basin (Sullivan, 1985). Uplift of eastern source areas may have cut off an inferred early Eocene outlet on the northeast (e.g., Sklenar and Andersen, 1985). The prodigious supply of easily eroded volcaniclastic debris from the coeval Challis volcanic field created a distributive fluvial system that overwhelmed and filled the remaining tectonically generated accommodation space in the greater Green River basin. Because explosive volcanism can strongly impact fluvial sedimentation by providing huge volumes of mobile sediment to coeval fluvial systems (Smith, 1987), basin filling could have taken place in the absence of a climatic driving mechanism. The flood of volcanic-lithic debris containing both neovolcanic and paleovolcanic grains (e.g., Surdam and Stanley, 1980) provides unambiguous evidence for hydrologic integration of the greater Green River and Piceance Creek basins, with potential limited spilling into the Uinta basin. Increased fluvial inflow prior to arrival of the Bridger-Washakie clastic wedge is recorded by a return to fluctuating profundal lake conditions, abrupt freshening and change in water chemistry, and a shift to lighter oxygen-isotopic composition of Lake Gosiute at c.50 Ma (Fig. 7)(Surdam and Stanley, 1980; Johnson, 1985; Machlus et al., 2002; Smith et al., 2003, 2008; Carroll et al., 2008). The shift to open lacustrine conditions in the Green River and Piceance Creek basins, but not in the Uinta basin, disfavors a climatic forcing mechanism. Although runoff may have increased as a result of long-term changes in precipitation, tectonically driven changes in locations and elevations of drainage divides could also have influenced hydrology of the basins (Pietras et al., 2003b; Carroll et al., 2008). Observed diachronic lacustrine conditions in the greater Green River and Piceance Creek basins suggest that integration of a long river system with high-elevation headwaters in the newly developed Challis volcanic field may have increased discharge to the greater Green River basin, providing a mechanism for transition from evaporative to fluctuating profundal conditions and lighter isotopic composition (Carroll et al., 2008; Chetel et al., 2011). No strong sediment input to the south flank of Lake Uinta is evident after the earliest Eocene (Johnson, 1985), suggesting that fluvial input had shifted away from the Uinta basin approximately concomitantly, by ~50 Ma, thereby instigating evaporative conditions there. This facies shift was likely caused by capture of the California River system by stream piracy, as discussed later. In summary, numerous observations indicate that tectonism, climate, and sediment supply were important factors in determining the fill of Laramide basins, in terms of both age and distribution of facies. The relative importance of these factors remains to be elucidated through integration of detailed regional studies and combined stratigraphic, geochronologic, and geochemical approaches. Recent studies suggest that abrupt changes in drainage systems, perhaps caused by local uplift and drainage-basin capture, were the dominant factors that affected hydrology of lacustrine systems in the ponded basins.

LARAMIDE SEDIMENT-DISPERSAL NETWORK Application of UPb geochronological techniques to detrital zircon grains in sandstone (e.g., Gehrels, 2012) since the latter part of the 20th century has revolutionized understanding of continental sediment routing systems in general, and, more specifically, of how sediment was likely transported into and between Laramide basins, as well as beyond the Laramide province to the Gulf of Mexico (e.g., Blum and Pecha, 2014; Sharman et al., 2017). Detrital-zircon geochronology has confirmed sources of sediment in local Laramide uplifts inferred from petrographic and paleocurrent studies and also resulted in new inferences regarding source areas by demonstrating that some sediment transfer took place across hundreds of kilometers, from as far away as the Late Cretaceous magmatic arc. A combination of sandstone petrology and detrital-zircon analysis from stratigraphic sections throughout the Laramide orogen indicates two general categories of fluvial systems, intra-orogen and extra-orogen, transported sediment to and through Laramide basins (Fig.  9). Deposits of intra-orogen river systems record local sources of sediment in adjacent ranges, whereas extra-orogen rivers transported sediment from source areas that lay west or south of the orogen, with mixing of detritus from local source areas during transit through the broken foreland. Intra-orogen sediment dispersal has been documented in several basins. In many cases, information from detritalzircon analysis confirms earlier provenance interpretations from sandstone petrology and paleocurrent analysis, whereas



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FIG. 9  Representative detrital zircon age distributions of Laramide fluvial sandstones. Colour age intervals represent Laurentian age provinces adapted from Gehrels et al. (2011) and Pecha et al. (2018); see Fig. 1 for distribution of age provinces in Laramide province. Sources of data: Deposits of the California River in Kaiparowits basin (Kaiparowits Formation; Lawton and Bradford, 2011) and Uinta basin (Colton Formation; Davis et  al., 2010; Dickinson et al., 2012); Ringbone Formation, Clinkscales and Lawton (2015); Baca Formation, Sawtooth Mountains northwest of Datil, New Mexico, A. Licht and J. Quade, written communication (2012); Wasatch Formation of Piceance Creek (PC) basin, Dickinson et  al. (2012) and Foreman and Rasmussen (2016); San Juan basin, Pecha et al. (2018); Galisteo basin, Sharman et al. (2017); Raton basin, Bush et al. (2017); Bighorn basin, May et al. (2013b); Wind River basin, Fan et al. (2011). Age range plotted 0–3000 Ma for scale purposes, causing uncommon older grains to fall off plots: Fort Union sample contains 7 grains in the range 3430–3007 Ma (May et al. 2013b). Wind River Formation contains 3 grains in the range 3317–3224 Ma (Fan et al., 2011).

in other cases, new interpretations have arisen. Local sediment sources are best identified by arkosic sandstone compositions and large proportions of zircon grains with prominent age peaks that correspond to local basement ages (Fig. 9; e.g., Fan et al., 2011; May et al., 2013b; Bush et al., 2017). Although there are exceptions, analysis of stratigraphic sections within individual basins demonstrates progressive unroofing of Mesozoic and Paleozoic strata from adjacent uplifts. For example, in the Raton and Galisteo-El Rito basins, progressive upsection decreases in the proportion of Paleozoic and Mesozoic grains and concomitent increases in Yavapai-Mazatzal basement ages record erosional stripping of Mesozoic strata from the adjacent Sangre de Cristo and Nacimiento uplifts to expose subjacent granitic and metamorphic basement (Bush et al., 2017; Sharman et al., 2017). Detrital-zircon data from the nearby Huerfano Park basin differ in demonstrating that basement c­ onstituted the primary source for sediment in strata at the base of the Laramide section, represented by the middle Paleocene Poison Canyon Formation, with no evidence of contributions from a sedimentary carapace; moreover,

546  The Sedimentary Basins of the United States and Canada

Cambrian zircon grains were derived from distinctive alkalic intrusions in the Wet Mountain uplift to the north, corroborating south-directed, axial sediment transport in the basin and conflicting with prior interpretations that the adjacent San Luis uplift was the primary sediment source for the basin (Rasmussen and Foreman, 2017). In the Bighorn basin, detrital-zircon ages change from the Maastrichtian Lance Formation to the Paleocene Fort Union Formation (Fig. 9) (May et al. 2013b). The Lance contains abundant 1.7 Ga grains that do not reflect local basement, but instead were transported from Mesoproterozoic sedimentary rocks of the Belt Supergroup in western Montana (May et al., 2013b). The Lance also contains Cretaceous grains likely derived from the Idaho batholith and less common Paleozoic and 1.0 Ga grains recycled from Mesozoic strata, yielding a mix of grains from local and extra-orogen sources during earliest Laramide deposition in the Bighorn basin (e.g., Steidtmann, 1993). Upsection increase in Archean basement grains in the Fort Union Formation likely records progressive erosional removal of the Phanerozoic sedimentary section from the basement cores of adjacent uplifts, isolation of the basin from external drainages, and a drainage reversal within the basin to north-flowing rivers, as documented by paleocurrent studies (Lillegraven and Ostresh, 1988; Fan et al., 2015). Because of their distance from contemporary magmatic provinces, uplifts drained by intra-orogen fluvial systems typically failed to provide near-depositional zircon ages to strata in adjacent basins. This is especially evident in the Bighorn and Wind River basins, where arc-derived zircon grains in Paleogene strata are generally of Cretaceous age and likely recycled from Mesozoic strata on adjacent uplifts (Fig. 9; Fan et al., 2011; May et al., 2013b). In contrast, young, even syndepositional, grains occur near intraforeland volcanic fields, such as the Colorado mineral belt and the southwest New Mexico volcanic field. In the Raton basin, syndepositional grain ages are present in the Vermejo Formation and lower part of the Raton Formation, both of Maastrichtian age (Bush et al., 2017), approximately coeval with deposition in the San Juan basin of the volcaniclastic McDermott Formation, derived directly from time-equivalent volcanic rocks on the north flank of the basin (Cather, 2004; Wegert and Parker, 2011; Pecha et al., 2018). Syndepositional Campanian zircon in the Ringbone basin of southwestern New Mexico could have been derived from coeval volcanic rocks of the Laramide copper porphyry province in Sonora (González-León et al., 2011; Clinkscales and Lawton, 2015) or from time-equivalent volcanic rocks of the adjacent southwest New Mexico volcanic field. Large extra-orogen river systems drained into the Laramide orogen from segments of the largely quiescent magmatic arc that lay west of the broken foreland. These river systems, which have taken the names of the modern states in which they had their headwaters, have been delineated not only by UPb zircon ages, but also by bulk-rock Pb isotope compositions of sandstone and 40Ar/39Ar ages of potassium feldspar grains (Davis et al., 2010; Chetel et al., 2011; Dickinson et al., 2012). These rivers had ancestries that predated development of the Laramide broken foreland, but foreland structure and magmatic history adjacent to the orogen influenced and altered their courses. The rivers played an influential role in the evolution of facies associations in ponded basins into which they drained. Deposits of extra-orogen drainages tend to be enriched in predepositional arc grains (Dickinson et al., 2012; May et al. 2013a, b). These grains of arc derivation do not serve to discriminate between California and Idaho segments of the Late Cretaceous arc because the magmatic histories of the two arc segments are quite similar (Ducea, 2001; May et al. 2013a). This similarity poses a problem for analyses of sediment routing from the Laramide orogen to the Gulf of Mexico proposed on the basis of abundant arc-derived Cretaceous and Paleogene grains in Paleocene-Eocene Wilcox sandstones (e.g., Mackey et  al., 2012; Blum and Pecha, 2014; Sharman et al., 2017; Blum et al., 2017). A large group of associated grain ages near 1.8 Ga, probably derived from the Belt Supergroup (May et al., 2013a, b), and Archean grains from the Wyoming basement province in northern drainages may serve to discriminate northern from southern drainage systems. Two major extra-orogen fluvial systems constitute the California and Idaho rivers. The California River drained northward from the magmatic arc terrane of southern California and Arizona. The river’s ancestry extended back to axial fluvial systems that flowed northward along the foredeep of the Cordilleran foreland basin beginning as early as the Coniacian (e.g., Lawton et al., 2003, 2014; Szwarc et al., 2015). The initial manifestation of the Laramide California River is the Kaiparowits Formation of southern Utah, which contains arc-derived Jurassic and Cretaceous zircon grains, many of the latter syndepositional Campanian grains (Fig.  9; Lawton and Bradford, 2011), and modest numbers of grains derived from Yavapai-Mazatzal basement of the Mogollon highlands (Fig. 10A). Fluvial systems and thickness of the Kaiparowits Formation were influenced by local uplifts, including the Kaibab uplift (Goldstrand, 1994; Biek et al., 2014), deformation of which was underway by the end of Wahweap deposition (Tindall et al., 2011). An alternate route for the Kaiparowits fluvial system to the San Juan basin was advocated by Roberts (2007), but that proposed route was likely blocked by the incipient Circle Cliffs uplift (Fig. 10A). Immediately following Kaiparowits deposition, exhumation of the Coconino terrace (Fig. 1) or analogous Kingman uplift (Beard et al., 2010) and incision of upstream reaches of the early California River drainage directly south of the western Grand Canyon took place at the end of the Campanian (72 Ma; Wernicke, 2010; Flowers and Farley, 2012), probably during deposition of the coarse-grained Canaan Peak Formation. By the end of Paleocene time, a mature California River delivered sediment from southern basement sources (Fig. 10B), indicated by



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FIG. 10  Inferred routes of main drainages in Laramide orogen for four time periods defined by changing spatial distribution of uplift and volcanism in and around the orogen. Inferred northeast-transient deformation primarily after Dickinson et al. (1988), Heller and Liu (2016), Copeland et al. (2017), and references cited in text. Distribution of silicic to intermediate magmatic rocks (rhyolite, granite, granodiorite) for each time period indicated by active magmatic front and intraforeland magmatism generalized from NAVDAT database (NAVDAT, 2017). Uplift abbreviations as in Fig. 1. (A) Northeast and east drainage during development of southern perimeter basins and Colorado Plateau uplifts. Dominant sediment flux from the orogen was to the Western Interior seaway. This interval marks uplift of the Coconino terrace or Kingman uplift (Fig. 1) and incision of upstream reaches of early California River (Wernicke, 2010; Flowers and Farley, 2012). Southeast sediment transport in southern perimeter basins was directed toward northeastern Mexico along the Sierra Madre front (Lawton et al., 2009). Extant plateau uplifts after Beard et al. (2010), Pecha et al. (2018). Incipient rise of Uinta uplift indicated by Ericson Formation (Fig. 2; Leary et al., 2015). Other sources of data: Hayes (1987); Goldstrand (1994); González-León and Lawton (1995); Continued

548  The Sedimentary Basins of the United States and Canada

detrital-zircon data and feldspathic sandstone composition, and terminated at the distributive fluvial system of the Colton Formation on the south margin of Lake Uinta, which at the time was a fresh, fluctuating profundal lacustrine system (Fig. 2; Fouch, 1976; Ryder et al., 1976; Smith et al., 2008; Davis et al., 2010; Dickinson et al., 2012). Sediment from the Uinta basin did not spill east across the Douglas Creek arch, as indicated by local Sawatch and Uncompahgre uplift sources for zircon grains in the Wasatch Formation of the Piceance Creek basin (Fig. 9; Foreman and Rasmussen, 2016). Large-volume discharge to the Uinta basin by the California River likely established open-lacustrine conditions in the basin by late Paleocene time, immediately prior to arrival of the Colton distributive fluvial wedge and before a major lake developed in the Green River basin north of the Uinta uplift (Fig. 2; Smith et al., 2008). After the early Paleocene, the California River sediment load could not continue northeastward to exit the Laramide orogen because it was blocked by the Uinta uplift (Fig. 10B). The California River probably continued to discharge into Lake Uinta through most of early Eocene time, as the lake maintained its profundal fluctuating state even as the Green River and Piceance Creek basins experienced closed, evaporative conditions (Fig. 10C) (Smith et al., 2008). The Idaho River, as defined by Chetel et al. (2011), began to deliver volcanic-lithic detritus to the Green River basin from headwaters in the Challis volcanic field of western Montana and eastern Idaho at ~50 Ma (Figs. 1 and 10D). The arc region of Idaho had contributed zircon of syndepositional age to Upper Cretaceous strata of the Western Interior seaway from at least Cenomanian time (May et al., 2013a, b) until the seaway withdrew from Wyoming and Montana near the end of the Campanian (Robinson Roberts and Kirschbaum, 1995). As noted previously, the Maastrichtian Lance Formation contains Late Cretaceous zircon derived from the batholith, but those grains may have been partly recycled from Cretaceous strata on nearby growing uplifts. The Idaho River drainage was diverted or captured ~50 Ma to deliver the massive influx of volcaniclastic detritus to the Green River basin from the coeval Challis volcanic field, which overlies the Idaho Batholith and Belt Supergroup. When the Green River and Piceance Creek basins returned to profundal conditions upon capture of the Idaho River drainage (Smith et al., 2003, 2008; Carroll et al., 2008), Lake Uinta was evaporative, receiving little or none of the inflow that crossed the Uinta uplift into the Piceance Creek basin and evidently no longer the recipient of a significant extra-orogen drainage. The detrital-zircon population of the Baca Formation is similar to that of the Colton population (Fig. 9), permitting the possibility of a common headwater region for the Colton and Baca formations in the Laramide porphyry copper segment of the Cordilleran arc (e.g., Pecha et al., 2018). The development of east-flowing axial drainage in the Baca basin, even as the California River appears to have dried up, suggests the possibility that the southwestern headwaters were captured by an east-flowing drainage, termed here the Mogollon River (Fig. 10D). Two extra-orogen drainages that headed in the Laramide copper porphyry province and its eastern counterpart, the southwest New Mexico volcanic field, developed as early as middle Paleocene time (Fig. 10B). One of these, which may have developed into the middle Eocene Mogollon River, likely transported arc grains from southern Arizona and Sonora to the Gulf of Mexico by way of the Tornillo perimeter basin (Fig. 1) to provide Jurassic to Eocene zircon grains to the Paleocene-Eocene Wilcox Formation of the Texas Gulf Coast (Fig. 9B–D) (Mackey et al., 2012; Kortyna et al., 2017). The other drainage transported Cretaceous and Paleogene grains along the front of the Sierra Madre orogen to a foreland basin in northeastern Mexico (Fig. 10B and C) (Lawton et al., 2009).

TECTONICS OF THE LARAMIDE OROGENY Spatial and temporal patterns of Laramide deformation and basin development, in combination with plate reconstructions and arc magmatic history at the continental margin, comprise essential data for interpreting underlying mechanisms of FIG.  10 CONT’D  Basabilvazo (2000); Lawton and Bradford (2011); Wegert and Parker (2011); Heller et  al. (2013); May et  al. (2013b). WIS, ­approximate mid-Campanian (~75–74 Ma) shoreline of Western Interior seaway (Robinson Roberts and Kirschbaum, 1995), which migrated rapidly to east in Maastrichtian (hence overlap of some Maastrichtian uplifts with seaway). (B) Development of intra-orogen drainage and external drainage toward the Gulf of Mexico during widespread foreland partitioning. Northern rivers drained around the periphery of the orogen, either to the northern remnant of the Western Interior Seaway (Sharman et al., 2017) or toward the Gulf of Mexico to generate the dominant middle Paleocene sediment influx to the Wilcox Group (Galloway et al., 2011; Mackey et al., 2012). Data sources: San Juan basin paleocurrent direction and fluvial source, Klute (1986); Galisteo basin, Sharman et al. (2017); Raton basin, Bush et al. (2017); Huerfano Park basin, Rasmussen and Foreman (2017); Denver basin, Sharman et al. (2018); Black Hills, Lisenbee and DeWitt (1993), Finzel (2017). Incipient uplifts in Wyoming from Yonkee and Weil (2015). (C) Development of mature California River, which delivered sediment to Colton distributive fluvial system on south edge of Lake Uinta, well integrated intra-orogen drainages, and outflow from perimeter basins. Data sources: Fouch (1976), Ryder et al. (1976), Surdam and Stanley (1980), Dickinson et al. (1988), Cather (2004), Davis et al. (2010), Dickinson et al. (2012), May et al. (2013b), Foreman and Rasmussen (2016), Sharman et al. (2017); Pecha et al. (2018). (D) Drainage network during southwestward migration of late intra-foreland deformation. Late middle Eocene diversion of Idaho River into Green River basin restored fluctuating profundal conditions north of the Uinta uplift and to the Piceance Creek basin (Chetel et al., 2011). Mogollon River headwaters captured California River, resulting in evaporative conditions in Uinta basin, recorded by Uinta Formation, and sediment delivery from active magmatic centers of Laramide Porphyry copper province to axial distributive fluvial system in Baca basin (Cather, 1982; Cather and Johnson, 1986, Potochnik, 1989).



Laramide Sedimentary Basins and Sediment-Dispersal Systems Chapter | 13  549

Laramide orogenesis. The general amagmatic character of Laramide deformation, with some important exceptions, and distribution of basins and uplifts in a cratonic setting inboard of a recently active fold-and-thrust belt suggest analogy with the Sierras Pampeanas province of Argentina, where the Nazca plate is being subducted subhorizontally beneath South America (Allmendinger et al., 1983; Jordan and Allmendinger, 1986). Indeed, the longest-running plate-tectonic model for Laramide deformation posits a flat Farallon slab beneath western North America to create far-field stresses within the foreland capable of effecting the observed crustal shortening (Dickinson and Snyder, 1978; Chapter 11). Foreland shortening was likely due to sublithospheric shear caused by coupling with the subducted slab (Bird, 1988, 1998; Saleeby, 2003; Heller and Liu, 2016). The flat slab has been explained in terms of increased convergence rates between the Farallon and North American plates (Engebretson et al., 1984) or subduction of a buoyant aseismic ridge or oceanic plateau (Livaccari et al., 1981; Henderson et al., 1984). Analysis of paleostress directions assembled from dikes, veins, and fault slip vectors indicates that the dominant shortening direction was ENE (068 degrees) at 75 Ma, with possible slight clockwise rotation of 15–30 degrees at 45–40 Ma (Bird, 2002). Early contraction that initiated basement-involved crustal deformation may have begun as early as 85 Ma through lithospheric delamination beneath southwesternmost North America as the buoyant slab segment encountered the North American plate (Saleeby, 2003). Lithospheric shearing or delamination created an end load to effect early shortening in the southwestern part of the Laramide foreland, which was followed by shortening attendant upon basal shearing of the flat slab in the distal part of the Laramide province. A model of the path and extent of the subducted oceanic plateau beneath North America provides a reasonable match for diachronous deformation observed in the Laramide orogen (Fig. 11). Successive positions of the subducted Shatsky Rise conjugate beneath North America are predicted on the basis of plate reconstructions (e.g., Liu et al., 2010). Calculation of

FIG. 11  Approximate track of Shatsky Rise conjugate beneath North America before and during the early part of the Laramide orogeny. Red dots indicate the reconstructed locations of thickest part of the flat slab, or conjugate Shatsky Rise, at four-m.y. intervals, with ages in Ma indicated near each dot, from Heller and Liu (2016). Position of the rise thick at 60 Ma, during northeasternmost Laramide deformation in the Black Hills uplift, lay 250 km NE of the 64 Ma position. Area with diagonal rule is distribution of Shatsky Rise conjugate at 88 Ma, also from Heller and Liu (2016). The 800 km wide gray band approximates the progressive footprint of the Shatsky Rise conjugate during its prograde migration beneath North America, assuming no rotation of the subducted plateau.

550  The Sedimentary Basins of the United States and Canada

dynamic topography associated with transit of the Shatsky conjugate predicts initial subsidence followed by uplift, the latter coincident with onset of deformation in the foreland and probably due to enhanced coupling between the two plates (Heller and Liu, 2016). Although the onset of basement uplift and basin subsidence across the Laramide foreland is not precisely constrained for reasons outlined already, there is a general northeastward progression of ages of earliest Laramide deposition from the southern perimeter basins to the Powder River basin apparent in the correlation chart of Fig. 2. An advance of approximately 1200 km from the McCoy basin to the Bighorn uplift between 75 and 65 Ma yields a structural-front migration of 120 km/m.y., within the range of the 90–150 km/m.y. Farallon-North America convergence rates calculated for that time interval (Engebretson et al., 1984). The migration rate of the conjugate plateau during the same time interval is in the range of 75–80 km/m.y. (Fig. 11) (Heller and Liu, 2016). Although not exactly equivalent, the velocity ranges are not widely divergent. In conclusion, a flat-slab model for Laramide deformation satisfies a list of diverse observations: (1) kinematic criteria derived from an actualistic plate-tectonic setting in South America; (2) calculated plate-convergence rates between North America and the subducted Farallon plate linked to the Late Cretaceous-Paleogene shortening history in the foreland; (3) temporal and spatial distribution of Paleogene magmatism and calculated stress directions; and (4) predictions of dynamic topography and deformation migration derived from plate reconstructions of a subducted conjugate of a modern oceanic plateau beneath North America. Flat subduction, caused by the buoyancy of the thick oceanic lithosphere of the plateau, therefore constitutes a viable paradigm for the plate-tectonic origin of Laramide orogenesis.

ACKNOWLEDGMENTS Discussions over the years with the following persons influenced my thinking and greatly advanced my understanding of things Laramide: Beth Bradford, Chris Clinkscales, Peter Copeland, Bill Dickinson, Elisa Fitz-Díaz, Carlos González-León, Paul Heller, Bill Seager, and Daniel Stockli. Andrew Barth, Ronald Blakey, William Dickinson, and Beth Welle reviewed an earlier version of this chapter. Reviews of the revised chapter by Ron Blakey and Ray Ingersoll improved content and writing. Although the reviewers do not agree with all of the content, their comments helped clarify and amplify some of my thinking. Above all, I thank W.D. Darton, who first navigated me through the Laramide of the Colorado Plateau nearly 40 years ago.

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