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Late Paleozoic Deformation of Interior North America: The Greater Ancestral Rocky Mountains1

Hongzhuan Ye,2 Leigh Royden,3 Clark Burchfiel,3 and Martin Schuepbach4

ABSTRACT Late Paleozoic deformation within interior North America has produced a series of north-northwest­ to northwest-trending elongate basins that cover much of Oklahoma, Texas, New Mexico, Colorado, and Utah. Each basin thickens asymmetrically toward an adjacent region of coeval basement uplift from which it is separated by synsedimentary faults with great vertical relief. The remarkable coincidence in timing, geometry, and apparent structural style throughout the region of late Paleozoic deformation strongly suggests that these paired regions of basin subsidence and basement uplift form a unified system of regional deformation, the greater Ancestral Rocky Mountains. Over this region, basin subsidence and basement uplift were approximately synchronous, beginning in the Chesterian­Morrowan, continuing through the Pennsylvanian, and ending in the Wolfcampian (although minor post-Wolfcampian deformation occurs locally). The basement uplifts show evidence for folding and faulting in the Pennsylvanian

©Copyright 1996. The American Association of Petroleum Geologists. All rights reserved. 1 Manuscript received August 18, 1994; revised manuscript received January 30, 1996; final acceptance April 8, 1996. 2 Massachusetts Institute of Technology, Cambridge, Massachusetts 02139. Present address: University of Texas at Dallas, P.O. Box 688, Richardson, Texas 75083-0688. 3 Massachusetts Institute of Technology, Cambridge, Massachusetts 02139. 4MAXUS Energy Company, 717 North Harwood Street, Dallas, Texas 75201. Present address: Danube International Petroleum Company, 2651 North Harwood Street, Dallas, Texas 75201. This paper grew out of a 1989 workshop at Maxus Energy Company. Subsequent interpretation of seismic reflection data was completed by Vinnie Rigatti. Additional studies were completed by Hongzhuan Ye during a research fellowship at the Massachusetts Institute of Technology (MIT) under the supervision of Leigh Royden and funded by NSF grant EAR-8721401 as part of a joint project between MIT and the Chengdu Institute of Geology and Mineral Resources, and by NSF grant EAR-9024320 to Leigh Royden as part of the NSF Faculty Awards for Women in Science and Engineering. We thank T. Ewing for helpful suggestions, and P. Kluth, J. Peterson, and R. Stanley for helpful and thorough reviews. We thank J. Lowell for permission to publish Figure 5, and Maxus Energy Company for drafting Figure 2. Hongzhuan Ye thanks Liu Baojun for making it possible for him to study at MIT, and gratefully acknowledges receipt of a scholarship from the Southwestern Section of AAPG.

and Early Permian. Reverse faults and thrust faults have been drilled over many of the uplifts, but only in the Anadarko region has thrusting of the basement uplifts over the adjacent basin been clearly documented. Extensive basement-involved thrusting also occurs along the margins of the Delaware and Midland basins, and suggests that the entire greater Ancestral Rocky Mountains region probably formed as the result of northeast-southwest­ directed-intraplate shortening. Deformation within the greater Ancestral Rocky Mountains was coeval with late Paleozoic subduction along much of the North American plate margin, and has traditionally been related to emplacement of thrust sheets within the Ouachita-Marathon orogenic belt. The nature, timing, and orientation of events along the Ouachita-Marathon belt make it difficult to drive the deformation of the greater Ancestral Rocky Mountains by emplacement of the Ouachita-Marathon belt along the southern margin of North America. We speculate the deformation was driven instead by events within a late Paleozoic Andean margin along the southwestern margin of North America. Evidence for the existence of this previously unrecognized Andean margin comes from east-central Mexico, where a Pennsylvanian and Permian volcanic arc indicates that a northeastdipping subduction boundary lay to the south and west. The interpretation that deformation throughout the greater Ancestral Rocky Mountains occurs by basement-involved overthrusting on gently to moderately dipping thrust faults suggests that potential hydrocarbon reserves beneath crystalline thrust sheets may be much greater than is generally supposed. INTRODUCTION The relationship between subduction at plate margins and deformation within plate interiors is important because it bears on the way in which stress and strain are transmitted through the lithosphere and on the driving mechanisms for subduction and deformation at the plate boundary. In

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some instances, causal relationships between subduction at active continental margins and deformation within the continental interior seem generally established; for example, Laramide deformation in the Rocky Mountain region is thought to be the result of flat-slab subduction beneath the adjacent continental margin (Lipman et al., 1971; Burchfiel and Davis, 1975; Coney, 1976; Dickinson and Snyder, 1978). Another apparent example of deformation within a plate interior resulting from orogeny along the plate margin is the late Paleozoic deformation of southern North America. Here, Late Mississippian­Early Permian deformation within interior North America has produced a series of north-northwest­ to northwest-trending basins flanked by coeval zones of basement uplift (Figure 1). The late Paleozoic deformation of interior North America was approximately coeval with deformation along active margins that ringed much of North America, including subduction and thrusting along the Appalachian orogene, the Ouachita-Marathon orogene, an oceanic subduction boundary along the western margin of the United States, and a poorly known region of late Paleozoic deformation in east-central Mexico. Historically, most workers have regarded the late Paleozoic deformation of interior North America as the result of continental or arc collision along the Ouachita-Marathon orogenic belt (see, for example, Walper, 1977; Burchfiel, 1978; Coney, 1978; Casey, 1980; Kluth and Coney, 1980, 1981; Goldstein, 1981, 1984; Dewey and Pitman, 1982; Thomas, 1983; Viele, 1983, 1986; Warner, 1983; Kluth, 1986). Because the nature of the late Paleozoic deformation within interior North America has not been well documented [and has been variously interpreted as extensional, compressional, and strike slip (Rankin, 1975; Walper, 1977; Goldstein, 1981, 1984; Kluth and Coney, 1981; Budnik, 1986; Kluth, 1986; Soegaard, 1990; Soegaard and Caldwell, 1990; G. Viele, 1992, personal communication)], this causal relationship has been mainly inferred from the broad spatial and temporal coincidence of events within the two regions. In this paper, we summarize some data and recent interpretations of data that call this hypothesis into question. We examine the timing, distribution, and, in so far as possible, the nature of late Paleozoic deformation within interior North America, and relate that deformation to coeval plate boundary activity along the margins of North America. This exercise is not only important for our academic understanding of the relationship between the plate boundary processes and coeval deformation within the continental interior, but is also of great practical importance because part of

this region is one of the richest areas for oil and gas exploration in the continental United States. LATE PALEOZOIC DEFORMATION OF INTERIOR NORTH AMERICA Because the late Paleozoic deformation of interior North America is usually interpreted as the result of compressional stresses transmitted northward from the Ouachita-Marathon belt, it is commonly referred to as the late Paleozoic "foreland" deformation. In this paper, we refer to this either as the zone of late Paleozoic deformation of interior North America, which does not imply a causal link between events within the continental interior and deformation along its periphery or as the greater Ancestral Rocky Mountains (see, for example, Goldstein, 1984; Kluth, 1986). Of the numerous late Paleozoic basins and basement uplifts within this zone of deformation, we selected four major systems for discussion: the Wichita-Anadarko system, the Central basin platform-Delaware-Midland system, the basins and uplifts of the Ancestral Rocky Mountains, and the Pedernal-Orogrande system. A discussion of all of the individual basins and subbasins within this area is clearly beyond the scope of this paper, but those basins mentioned in the text are shown in Figure 1. These zones of late Paleozoic foreland deformation appear to be the reactivation of older structures, and their precise locations are thus determined by preexisting zones of weakness within the crust. For example, the Anadarko basin is superimposed on the Southern Oklahoma aulacogen, a zone of Cambrian rifting (Hoffman et al., 1974; Budnick, 1984; Hoffman, 1989). The Delaware basin is superimposed on the Tobosa basin, which also appears to have been a zone of rifting in the Cambrian (Keller et al., 1980; Arbenz, 1989a; Hoffman, 1989). Precambrian mafic and ultramafic rocks have been found along the southwestern margin of the Uncompahgre uplift, suggesting that deformation within the Ancestral Rockies was also localized by a preexisting zone of weakness within the crust (Baars, 1966; Stevenson and Baars, 1977; Stone, 1977; Budnick, 1986). With a few exceptions, these late Paleozoic basins have been buried by younger sediments or, in the western part of the region, redeformed during the Tertiary, thereby obscuring the original structural relationships between the basins and the basement highs that lie adjacent to them. Recent advances in our knowledge of these basins have come largely through the increased quality and quantity of subsurface data. However, in many places the nature of the deformation remains controversial, and the timing of deformation in these

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(a)

(b)

Figure 1--(a) Locations and names used in the text. Barbed lines are late Paleozoic thrust faults and inferred thrust faults; the barbs are on the upper plate. Hachured area southeast of the Mojave-Sonora megashear was not in its present location in the late Paleozoic. A. = arch, B. = basin, U. = uplift. (b) Locations of cross sections show L1­L4 in Figure 3, L5 in Figure 8, and the detailed map of the Central basin platform (CBP) in Figure 6. Inset shows late Paleozoic (Pennsylvanian) orogens of central North America and the regions of coeval deformation in interior North America. Solid heavy lines on the inset map show locations of Ouachita and Appalachian orogens; heavy dashed lines show hypothetical locations of offshore subduction systems in the Pacific Ocean and along the margin of northern Mexico. Note that the modern shoreline did not exist in the late Paleozoic and is shown for reference only.

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(a)

(b)

Figure 2--Depositional environments for the Ouachita-Marathon region and interior North America (United States) from the Chesterian­ Guadalupian (a­i, respectively), and for interior North America (Mexico) for the Pennsylvanian and Permian (j and k, respectively). Depositional environments were inferred from facies and stratigraphic descriptions cited throughout the text. Contour lines show stage isopachs, where available, mainly from McKee and Crosby (1975) and Sloss (1988), supplemented by other sources cited throughout the text. Where contours are not labeled, the contour interval is 250 m. (When environments changed during an interval, we selected the environment that was most representative of that interval.) Dark-brown line on (j) and (k) shows the limit of available data for Pennsylvanian­ Permian facies in Mexico.

places must be largely constrained by the age, facies, and thickness of sediments deposited within the basins; therefore, much of the following discussion focuses on the sedimentation history of the basins. Wichita-Anadarko System The Wichita-Anadarko system includes, in addition to the Wichita uplift and Anadarko basin proper,

the Amarillo, Criner Hills, and Arbuckle uplifts, the Ardmore and Marietta basins, and the Palo Duro and Hardeman basins on the south side of the Wichita uplift (Figure 1). Amarillo-Wichita Uplift and Anadarko Basin The Anadarko basin trends west-northwest and on its southwest margin is separated from the Wichita-Amarillo uplift by a fault zone with up to

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(c)

Figure 2--Continued.

(d)

12 km of structural relief. However, along most of its length the northeastern margin is mainly unfaulted and sedimentary horizons thin and pinch out against the Cimarron arch to the northeast (Figures 2, 3) (Evans, 1979; Luza et al., 1987). Subsidence of the Anadarko basin began in middle or late Chesterian, when the stable passive margin began to subside rapidly beneath the Anadarko basin and more than 500 m of Chesterian sediments

were deposited (Figure 2a). Under much of the Anadarko basin is an angular unconformity between Chesterian and post-Chesterian strata, beneath which Chesterian and pre-Chesterian strata are folded, thrusted, and eroded (Petersen, 1983; Hill, 1984; Rascoe and Johnson, 1988). The thick Morrowan (Springer) clastic sediments record increased subsidence of the basin and uplift of the Amarillo-Wichita-Criner belt. Morrowan strata are

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Figure 2--Continued.

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predominantly marine shales, sandstones, and prograding fan-delta chert conglomerates. The provenance of the chert conglomerates is probably Mississippian cherty limestones and dolomites from the Amarillo-Wichita uplift (Evans, 1979; Shelby, 1980). Morrowan strata reach a maximum thickness of more than 2 km along the southwest margin of the Anadarko basin and thin toward the northeast, where they are partially eroded.

Atokan sedimentary rocks in the Anadarko basin are predominantly marine mudstones, but the amount of limestone in the section increases northwest of the Anadarko basin (Johnson et al., 1988). Atokan rocks reach a maximum thickness of about 1500 m along the southwestern side of the basin (Frezon and Dixon, 1975). Along the southwestern margin of the basin these marine deposits grade abruptly into the granite wash, alluvial-fan,

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(g)

Figure 2--Continued.

(h)

and fan-deltaic conglomerates that include granite, limestone, and dolomite fragments and coarse sandstone. This thick sequence of fan and fan-delta conglomerate is restricted to the margin adjacent to the Amarillo-Wichita belt, and suggests that rapid uplift and erosion of the Amarillo-Wichita belt resulted in unroofing of crystalline rocks in the core of the uplift by the Atokan. On the northern margin of the Anadarko basin, Atokan sediments

overstep the limit of the Morrowan strata and unconformably onlap the Mississippian rocks toward the north on the Cimarron arch. Along the southern margin of the Anadarko basin, pre-Atokan rocks underwent intense folding and faulting, and Atokan rocks are commonly absent on local positive structures (Rascoe and Johnson, 1988). Desmoinesian sedimentary rocks of the Anadarko basin are similar to those of the Atokan sequence,

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(i)

Figure 2--Continued.

and reach a maximum thickness of more than 1500 m in the deepest part of the basin. The progressive onlap of Morrowan­Desmoinesian sedimentary units toward the northeast indicates a progressive deepening and perhaps slightly northeastward migration of the basin through time. Missourian and Virgilian sedimentary rocks in the Anadarko basin overlie the Desmoinesian rocks unconformably and are predominantly prodeltaic and basinal shale and sandstone with minor carbonates (Frezon and Dixon, 1975). To the north and northwest, shallow-marine shelf carbonates dominate the section (Rascoe and Adler, 1983; Johnson et al., 1988; Rascoe and Johnson, 1988). The granite wash formation is limited to narrow bands along the northern edge of the AmarilloWichita uplift. The combined thickness of these bands reaches a maximum of more than 2000 m (Frezon and Dixon, 1975). In the Anadarko basin, Wolfcampian and Virgilian regional lithofacies patterns are quite similar to the Desmoinesian section (Figure 2d­g). Wolfcampian fan or fan-delta conglomerate and coarse sandstone occur on the north flank of the Amarillo-Wichita belt, and clasts of Paleozoic limestone and Precambrian granite can also be found in these sediments, suggesting that uplift of the Amarillo-Wichita belt probably continued into the early Wolfcampian. However, the Amarillo uplift was finally covered by sediments in the Wolfcampian, indicating that uplift of the Amarillo high ceased in the middle or late Wolfcampian.

Southwest of the Anadarko basin, the WichitaAmarillo uplift is cored by faulted Precambrian and Cambrian basement rocks, with folded and thrusted sedimentary rocks of Paleozoic age near the margin of the Anadarko basin. Most of the Paleozoic section has been removed from the Wichita-Amarillo uplift by erosion, and Wolfcampian rocks were deposited directly on the Cambrian or Precambrian basement (Ham et al., 1964; Rascoe and Johnson, 1988; Perry, 1989). Erosion of the Wichita-Amarillo uplift probably began as early as the Chesterian when cherty carbonates are presumed to have been eroded from the uplifted region(?). Uplift and erosion had clearly begun by the Morrowan, as indicated by deposition of the granite wash conglomerates along the southeast side of the Anadarko basin. The fault zone between the Wichita-Amarillo uplift and the Anadarko basin system consists of northeast-vergent asymmetric folds and overturned folds and basement-involved thrust faults, with thrust faults and blind thrusts propagating into the basin (McConnell, 1986, 1989). Numerous anticlines within the Anadarko basin probably represent hanging-wall structures to southwest-dipping thrust faults (such as the Cordell and Elk City anticlines). Thrust faults and folds are well exposed in some places (e.g., within the Slick Hills; McConnell, 1989), but most of the structures are subsurface and are known only from seismic and drilling information (e.g., Takken, 1968). Deep seismic reflection data collected by COCORP show

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(j)

Figure 2--Continued.

(k)

prominent southwesterly dipping reflections that, near the surface, are coincident with the thrust faults on the south side of the Anadarko basin (Brewer et al., 1983). These ref lections can be

traced to depths of 20­25 km under the WichitaAmarillo uplift, with a moderate dip of 30 to 40°, suggesting that thrusting has involved rocks to middle or lower crustal depths. Brewer et al. (1983)

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Figure 3--Cross sections showing comparative basin geometries plotted at the same scale; see Figure 1 for locations. Note that all basins are strongly asymmetric (and are fault-bounded on the deepest side by uplifted basement blocks). (The basement of the Taos trough is interrupted by a small fault of unknown age; reconstruction of offset restores the basement under the internal part of the basin to the position indicated by the dotted line.) The upper and lower boundaries of the basin sections shown are as follows. L1 = upper boundary is the top Permian (erosional top of Maroon, Webe, and Statebridge formations); lower boundary is the base Morrowan­ Atokan (Bolden, Coffman, and Kerber formations); L2 = upper boundary is the top Permian­Virgilian (Cutler Formation); lower boundary is the base Morrowan­Atokan (Molas Formation); L3 = upper boundary is base Wolfcampian­Missourian (Sangre de Cristo Formation), lower boundary is the base Morrowan­Desmoinesian (Sandia and Madera formations); L4 = upper boundary is the top Permian (erosional), lower boundary is the base Pennsylvanian. Profiles from DeVoto (1980), Soegaard (1990), and Luza et al. (1987).

Wolfcampian sediments probably signifies the end of tectonic activity along this zone, and postWolfcampian infilling of the Anadarko basin by Permian red beds and evaporites probably represents posttectonic filling of a preexisting depression. The asymmetric, southwest-thickening geometry of Chesterian­Wolfcampian sediments is consistent with formation of the basin as a flexural depression in front of the advancing thrust sheets of the Wichita-Amarillo uplift (Figure 3). Flexural modeling indicates considerable loading (during thrust-sheet emplacement) in post­late Desmoinesian (Garner and Turcotte, 1984). Taken together, these data strongly indicate that the Anadarko basin and the Wichita-Amarillo uplift formed as a result of northeast-directed, basement-involved thrust faulting beginning in the Chesterian­Morrowan, and probably terminating in the Wolfcampian (Figure 4). We did not find evidence for any significant extensional deformation in the basin or the uplifted regions and, although we cannot rule out the possibility that some strike-slip displacement occurred during deformation, we found little evidence for it (see Brown, 1984). Therefore, we believe that the late Paleozoic defor mation that for med the Anadarko basin and the Wichita-Amarillo uplift was predominantly basement-involved thrusting within an otherwise stable continental interior, similar to the Late Cretaceous­early Tertiar y deformation that has been so well documented within the Rocky Mountains (Figure 4). Although this interpretation is not particularly new, we present the supporting data in some detail because the thickness, distribution, age, and lithofacies in the Anadarko basin and the nature and timing of deformation observed on the crest of the Amarillo-Wichita uplift are similar to those developed throughout the region of late Paleozoic deformation within interior North America. Thus, these data may be important for comparison when we examine other regions of late Paleozoic deformation where the nature of deformation bounding the basins and basement highs is less clearly established. Ardmore and Marietta Basins (Arbuckle Uplift and Criner Hills) Toward the southeast, the Anadarko basin divides into two subbasins, the Ardmore and Marietta basins, separated by the Criner Hills (Figures 1, 2). These basins appear to have formed in a manner similar to the Anadarko basin, with overthrusting along the southwest margin of the Ardmore basin (in the Criner Hills) beginning in the Chesterian­Morrowan and continuing into

estimated about 10­20 km of thrust displacement along this fault system. The age of thrust faulting can be directly constrained as Morrowan and Atokan by the age of of fset sedimentar y strata within the basin, although less intense deformation of post-Atokan rocks can also be seen on seismic profiles (e.g., Brewer et al., 1983). Continuing uplift of the Wichita-Amarillo high and rapid subsidence of the Anadarko basin into the Wolfcampian suggest that thr usting continued into the Wolfcampian. Overlap of the Wichita-Amarillo high by late

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Figure 4--Time-slice maps showing the approximate location of thrust faulting during the Pennsylvanian and Early Permian. (Where tectonic activity changed during an interval, we selected the position of faulting that was most representative of that interval.) On some panels, a heavy dotted line shows the present location of the frontal thrust faults of the Ouachita­Marathon system.

(b)

at least the Desmoinesian. In the late Desmoinesian, thrusting affected the northeastern margin of the Arbuckle uplift, producing Desmoinesian and post-Desmoinesian uplift as indicated by orogenic conglomerates shed from the uplift and by deformation of the Ardmore and

Marietta basins. The data in support of the interpretation are given by Dott (1934), Tomlinson (1952), Tomlinson and McBee (1959), Bradfield (1968), Ham (1969), Frezon and Dixon (1975), Tennant (1981), Brown (1984), and Sutherland (1984, 1988b).

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Figure 4--Continued.

(d)

Palo Duro and Hardeman Basins (Matador-Red River Uplifts) The Palo Duro basin (including the Whittenburg trough, its northeast extension) and the Hardeman basin are shallow basins located on the south side of the Amarillo-Wichita uplift (Figures 1, 2), and are bounded by the Matador-Red River uplift on the

south. The Palo Duro and Hardeman basins subsided in the Late Mississippian­Late Pennsylvanian and are asymmetric southward-deepening basins. South of the basins, the east-west­trending Matador-Red River uplift is a basement-involved faulted and folded belt that became emergent in approximately the Morrowan, was eroded during

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(e)

Figure 4--Continued.

the Atokan­Desmoinesian, and was completely submerged by the Missourian. The origin of the uplift is not clear due to limited data from the region, but Crosby and Mapel (1975) described folded and faulted early Paleozoic and Precambrian rocks unconformable beneath Pennsylvanian strata on the crest of the Red River and Muenster uplifts. The major source area for sediments in the Palo Duro basin is the Amarillo-Wichita uplift to the north; however, except for the Whittenburg trough, the southward-deepening asymmetry of the basin (Oliver, 1976; Handford, 1979; Handford et al., 1981; Dutton and Goldstein, 1988) suggests that the main structural control on basin subsidence was to the south along the margin of the Matador and Red River uplifts. Lacking more definitive data, we tentatively conclude that the Palo Duro and Hardeman basins were overthrust during basement-involved thrusting and folding in the Matador and Red River areas, and represent a small-scale ver sion of the Wichita-Amar illoAnadarko system. In contrast, the narrow Whittenberg trough northeast of the Palo Duro basin is bounded by a reverse fault with more than 1 km of structural separation along its north side; reverse faults may occur within the trough as well (Brewer et al., 1983; Budnick, 1984). Although the Whittenberg trough could be interpreted as a strike-slip­dominated feature, it can also be interpreted as subsidence related to backthrusting of the Amarillo-Wichita uplift

toward the south, with normal faulting occurring on the south side of the trough in response to flexural bending of the basement during thrusting. Central Basin Platform-Delaware-Midland System The north-northwest­trending Central basin platform is bounded to the west by the Delaware basin and to the east by the Midland basin (Figures 1, 2). The Delaware basin deepens toward the Central basin platform in the east, although the eastward dip of the basin was accentuated during Laramide and Basin and Range deformation (Adams, 1965). Prior to the Chesterian, the Delaware basin was part of a region of slow subsidence (Tobosa basin) dominated by deposition of shallow-water limestones (Hills and Galley, 1988). In the Chesterian­Morrowan, the Delaware basin began to deepen rapidly, as recorded by the deposition of thick black organic shales (Figure 2a, b) (Meyer, 1966). The presence of black shales, the lack of proximal deposits along the eastern margin of the Delaware basin, and the close spacing of the isopachs suggest that the distribution of Morrowan sediments was formerly more extensive than at present (e.g., Meyer, 1966). The Pennsylvanian evolution of the Delaware basin is not well understood largely because the preserved or available Pennsylvanian and Early

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Permian strata were deposited in a deep-water starved basin environment (Vertrees et al., 1964; Cys and Gibson, 1988). However, during this period the basin appears to have deepened asymmetrically toward the east and Early Permian sedimentary strata thicken toward the Central basin platform (Cys and Gibson, 1988). The oldest rocks to overlap the contact between the Delaware basin and the Central basin platform are Wolfcampian, with a maximum thickness of up to 4300 m in the area adjacent to the Central basin platform (Feldman et al., 1962; Cys and Gibson, 1988), indicating that active tectonism along the eastern margin of the Delaware basin had ceased by the late Wolfcampian. However, posttectonic deep-water clastic sediments up to 1500 to 2000 m thick were deposited in the Delaware basin during the Leonardian and Guadalupian (Permian), and the Delaware basin was finally filled in the Ochoan with deposition of evaporites (Castile Formation). East of the Delaware basin, the Central basin platform is made of strongly folded and faulted Paleozoic and Precambrian rocks overlain by Early Permian (middle Wolfcampian­Leonardian) rocks above a regionally developed unconformity (Scobey et al., 1951; Davis et al., 1953; Jones and Matchus, 1984). The Precambrian basement is exposed in a narrow interior part of the Central basin platform with folded and faulted sedimentary cover along the margins of the platform. Northnorthwest­trending asymmetrical anticlines within the sedimentary cover are commonly faulted along their steepest limb (Galley, 1958, 1968). Two unconformities are on the Central basin platform, one of Desmoinesian age and one of Wolfcampian age. Pennsylvanian strata within the Delaware basin contain eastward-derived chert conglomerates (Adams, 1965), and the early Wolfcampian section adjacent to the Central basin platform contains large exotic blocks of Mississippian limestone (Guinan, 1971), indicating uplift and erosion of the Central basin platform. Data from 113 boreholes from the western and eastern parts of the Central basin platform, as well as from adjacent regions, show that stratigraphic inter vals from the Mississippian to the Wolfcampian are repeated and thickened, as well as missing in many places (Galley, 1958, 1968; Guinan, 1971). In our opinion, these data are best interpreted as thickening and repetition of units due to Late Pennsylvanian thrusting and erosion over the tops of anticlines formed during thrusting (discussed in more detail in a following section). The Central basin platform is flanked to the east by the Midland basin, an asymmetrical westwardthickening basin in which sedimentary strata pinch out northward onto the Matador arch and eastward onto the Bend arch (Figure 2). Subsidence of the

Midland basin appear s to have begun in the Desmoinesian when Desmoinesian platform carbonates were drowned and the basin became starved during the Missourian and Virgilian (Adams et al., 1951; Davis et al., 1953; Ewing and Garret, 1984; Jones and Matchus, 1984). The basin was gradually filled by shallow-water platform carbonates and evaporites in the Permian, although tectonically driven subsidence of the basin probably ceased in the Wolfcampian, long before the basin was filled. The nature of the deformation that produced the Delaware basin, Central basin platform, and the Midland basin remains controversial, and these structures have been variously interpreted as normal faults, strike-slip faults, upthrusts (near-vertical reverse faults), and basement-involved overthrusts, and have been inferred to have formed within a wide variety of stress fields and tectonic settings (e.g., Galley, 1958, 1968; Elam, 1969, 1984; Hills, 1970; Walper, 1977; Ewing and Garrett, 1984; Shumaker, 1992; Ewing, 1993). However, some published data and some new seismic data presented in this paper suggest that the dominant mode of deformation was basement-involved overthrusting of the Central basin platform over the Delaware and Midland basins. Along the southwestern margin of the Central basin platform, several wells, especially the Phillips 1-EE University, have shown evidence for large reverse or thrust faults, which put Precambrian rocks over early Paleozoic strata. Although Galley (1958, 1968) interpreted this scenario to result from near-vertical upthrusts, the sections encountered in these wells can also be interpreted to have resulted from basement-involved thrusting (James Lowell, 1975, personal communication) (Figure 5). Within the easternmost part of the central Delaware basin (Coyanosa zone near the Fort Stockton uplift, located at the southern end of Figure 6), folded and faulted Pennsylvanian sedimentary rocks form a series of northwest- to west-trending anticlines. These anticlines are cut by northwest-trending faults of unknown sense along their southern margins. Ewing and Garrett (1984) reported substantial repetition of the Mississippian­Wolfcampian section in drill holes in this and nearby regions (Rojo Caballos), suggesting that these anticlines lie in the hanging walls of thin-skinned, southwest-vergent thrust faults within the sedimentary section (see alternate interpretation by Guinan, 1971). These data, together with evidence for Pennsylvanian­earliest Permian folding and faulting within the Central basin platform, the lack of proximal sedimentary facies along the western margin of the Central basin platform, and the presence of coeval northeast-directed thrusting of the WichitaAmarillo uplift over the Anadarko basin, suggest

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Figure 5--Structural section through the Phillips 1-EE University well in the southwestern Central basin platform showing overthrust on southwest side of the Central basin platform (modified after J. Lowell, 1975, unpublished figure based on interpretation of dipmeter data). Southwest-thrusted and overturned Paleozoic units (Ellenburger­Pennsylvanian) are unconformably overlain by Permian strata (Wolfcampian) (see Shumaker, 1992, for similar interpretation). Dots indicate basal part of the Simpson Formation, and thin solid black lines show beds within the Simpson Formation that could be correlated from well data. The well is located along the southwestern margin of the Central basin platform several kilometers south of the area shown in Figure 6. No vertical exaggeration.

that the Delaware basin formed during basementinvolved southwest-directed overthrusting of the Central basin platform over the Delaware basin. In detail, the structural boundary between the Central basin platform and the Delaware basin is quite complicated. Locally, thrust faults verge northeast and southwest, although west-directed thrust faults dominate the western margin of the Central basin platform, and east-directed thrust faults dominate the eastern margin. In this interpretation, the Delaware basin is the result of down-to-the-east f lexure of its basement in response to loads imposed by emplacement of thrust sheets along its eastern margin. The juxtaposition of the most deeply eroded and shallowest part of the Central basin platform (Fort Stockton uplift) against the deepest part of the Delaware basin also suggests a genetic relationship between the amount of structural relief on the basin margin and subsidence in the basin.

Although the depth of the Midland basin is considerably less than that of the Delaware basin, we have considerable data to support east-vergent overthrusting of the Central basin platform over the Midland basin. Drilling data show repeated and thickened sections of Mississippian­Wolfcampian rocks present along the western margin of the basin adjacent to the Central basin platform and within interior parts of the basin (Galley, 1958, 1968; Guinan, 1971). Figure 7 shows two reflection seismic profiles that cross the eastern edge of the Central basin platform and the western edge of the Midland basin. On these profiles the Paleozoic section (Ordovician through the top Pennsylvanian) is clearly thrust toward the east. On line 2 (Figure 7b) the main thrust fault has formed a ramp anticline in its hanging wall and thrusting is synsedimentary. On line 1 (Figure 7a) the projection of the main thrust fault has been eroded and is truncated by a basal Permian (Wolfcampian) unconformity, and ramp or frontal anticlines in the hanging wall of the main thrust fault have been eroded. On both lines, smaller folds within the sedimentary section west of the main thrust fault (within the Midland basin) appear to be anticlines formed in the hanging wall of smaller thrust faults, probably splays of the main thrust fault system, and are coeval with the main thrust fault. Based on these data, we interpret the Midland basin to have formed as a result of eastward-directed basementinvolved thrusting of the Central basin platform, with some thrusting, perhaps thin-skinned, also occurring in interior parts of the basin. To test whether the geometry of the Midland basin is consistent with flexure of the lithosphere due to loading by thrust sheets along its eastern margin, we used the method of Kruse and Royden (1994) to examine the f lexural behavior of the basement. We used the thickness of the Atokan­ Wolfcampian sedimentary section and Bouguer gravity data across the basin to constrain the Late Pennsylvanian­earliest Permian subsidence of the basement. (The base of the Atokan and the top of the Wolfcampian are both shallow-marine facies, thus representing approximate sea level datum planes.) The profile crossed not only the Midland basin, but extended east across the Bend arch and the Fort Worth basin, which forms the foredeep for the Ouachita thrust belt (Figure 1). Figure 8 shows that the subsidence of this region in the Atokan­Wolfcampian is consistent with flexural bending of an elastic plate with an effective thickness of 30 km subjected to loading along its western margin (the Central basin platform) and along its eastern margin (the Ouachita thrust belt). In this interpretation, the Bend arch represents an intermediate f lexural bulge (or f lexural uplift) that formed in response to downward flexing of the

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lithosphere beneath the Central basin platform and the Ouachita thrust belt. Thus, uplift of the Bend arch is genetically linked to the subsidence in the Midland and Fort Worth basins and to the tectonic activity along the margins of the Midland and Fort Worth basins. (Note that subsidence of the Midland and Fort Worth basins was not exactly coeval, with subsidence of the Fort Worth basin occurring almost entirely in the Desmoinesian, and subsidence of the Midland basin occurring throughout the Desmoinesian­Wolfcampian; see Figure 2 and the section on the Ouachita thrust belt.) The timing of thrusting is best constrained by seismic and drilling data, such as that shown in Figure 7, and by the timing of subsidence within the Delaware and Midland basins and the age of the unconformities developed on top of the Central basin platform. Chesterian­Morrowan subsidence of the Delaware basin indicates that thrusting was initiated at this time, approximately coeval with the initiation of thrusting along the north side of the Wichita-Amarillo uplift (Figure 4). Initiation of rapid subsidence within the Midland basin in the middle to late Desmoinesian indicates that significant basement-involved overthrusting on the east side of the Central basin platform probably did not begin until this time. Thrusting along both sides of the Central basin platform had ceased by the end of the Wolfcampian, as indicated by the middle Wolfcampian Hueco Formation, which unconformably covers rocks ranging in age from Precambrian to Pennsylvanian, and by the infilling of the basins during the Wolfcampian and postWolfcampian. The Desmoinesian and Wolfcampian unconformities that developed on the top of the Central basin platform are also consistent with a reorganization in the orientation or locus of thrusting in the Desmoinesian. This is the same period in which a minor reorganization of basement-involved overthrusting occurred within the southeastern part of the Anadarko basin system (Ardmore and Marietta basins and Arbuckle uplift). Ancestral Rocky Mountain System

Figure 6--Structure contour map of the Central basin platform for the top of the Ordovician Ellenburger Formation (location shown by box in Figure 1b). Isolines and depths are given in thousands of feet. Thrust faults were determined from seismic profiles, drilling data, and isopach patterns. Large numbers indicate location of seismic sections 1 and 2 in Figure 7. Interpretation by V. Rigatti, Maxus Energy Company, 1990.

The Ancestral Rocky Mountain system proper is made up of several basement-involved uplifts and flanking basins, including the Frontrange-Sierra Grande uplift, Denver and Dalhart basins, Uncompahgre uplift, Paradox basin, and, between the two zones of basement uplift, the Central Colorado and Taos troughs. Uncompahgre-Paradox System The northwest-trending Uncompahgre uplift separates the Paradox basin to the west and the

(a)

(b)

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Figure 7--Seismic reflection profiles across the northeastern margin of the Central basin platform showing overthrust nature of the platform edge. Locations given in Figure 6. Time is two-way traveltime in seconds.

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Figure 8--Bouguer gravity data (squares) and deflection data, inferred from the thickness of Atokan­ Wolfcampian sediments, for the Fort Worth and Midland basins (line L5 on Figure 1b). Solid lines show best- fitting (two-sided) flexural model with an elastic plate thickness (Te) of 30 km, treating the Midland-Fort Worth region as an intermediate foreland flexed downward beneath the thrust sheets of the Ouachita orogene and thrust sheets on the east side of the Central basin platform.

Central Colorado trough and Taos trough to the east (Figures 1, 2). The Paradox basin is of latest Mississippian to earliest Permian age and extends from northwestern New Mexico to east-central Utah (DeVoto, 1980). This basin is strongly asymmetrical and deepens northeastward toward the Uncompahgre uplift, where the thickness of Pennsylvanian strata exceeds 2000 m and up to 7880 m of structural relief exists between the deepest part of the Paradox basin and the Uncompahgre uplift (Figure 3) (Stevenson and Baars, 1986). Basin subsidence began in the Morrowan following a brief period of regional erosion in the Late Mississippian. The initiation of subsidence is recorded by the deposition of late Morrowan(?) to early Atokan nonmarine shales, sandstones, and rare conglomerates. The basin appears to have become progressively deeper, with marine deposits appearing by the late Atokan. Rapid subsidence in the Desmoinesian produced more than 1500 m of sediments near the northeastern margin of the basin (Figure 2). Coarse clastic

debris, presumably shed from the rising Uncompahgre uplift, was deposited within a narrow zone along the northeastern margin of the basin (Stevenson and Baars, 1986; Baars et al., 1988). Alluvial-fan and fan-delta deposits argue for active tectonism at this time, and the deposition of arkosic sandstones suggests that crystalline rocks within the Uncompahgre uplift became subject to erosion in the Desmoinesian. There is a gradual facies change within the Paradox basin from these coarse clastic deposits in the northeast through evaporites in the central part of the basin to shallow-marine carbonates and clastics in the southwestern part of the basin, indicating no significant faulting within the basin interior during the Desmoinesian (see, for example, Stevenson and Baars, 1986). Sediments deposited in the Missourian contain no salt or anhydrite, but are otherwise similar in facies and distribution to Desmoinesian deposits. Virgilian alluvial deposits prograded southwestward across the basin, suggesting that deposition became faster than subsidence at about this time. Sedimentation continued until the earliest Permian (Wolfcampian), when the basin was filled by fluvial red beds (Baars, 1962; Baars et al., 1988). Northeast of the Paradox basin, the Uncompahgre uplift consists of a faulted Precambrian metamorphic and plutonic core bounded on its southwestern side by a southwest-vergent fold and thrust belt that involves Precambrian crystalline and Paleozoic sedimentary rocks (see White and Jacobson, 1983). Along the southwestern margin of the Uncompahgre uplift, one drill hole penetrated overturned Paleozoic rocks beneath more than 4000 m of Precambrian crystalline basement (Frahme and Vaughn, 1983) and several other wells have encountered repeated sections of early Paleozoic strata (Stevenson and Baars, 1986). The Uncompahgre belt is thought to continue to the southeast into New Mexico, where Early Pennsylvanian northwest-trending anticlines have been recognized (Figures 2, 4) (Baltz and Bachman, 1956). Uplift of the Uncompahgre belt probably began in the Morrowan when the first coarse terrigenous clasts were shed into the Paradox basin. Continued uplift and erosion probably occurred throughout the Pennsylvanian, as indicated by alluvial deposits in the Paradox basin and by locally developed Desmoinesian unconformities within the Uncompahgre uplift (Spoelhof, 1976; Baars and Stevenson, 1984). Based on these data, we believe that the Uncompahgre uplift is best interpreted as a basement-involved thrust belt, thrust southwest over the Paradox basin. In this interpretation, the Uncompahgre uplift and the Paradox basin are analogous to the Wichita-Amarillo-Anadarko system, where better resolution by subsurface data shows

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the presence of basement-involved thrust faults to great depth. Although the detailed timing of late Paleozoic thrusting is not well known in the Uncompahgre-Paradox system, the initiation of uplift and subsidence in the Late Mississippian­ earliest Pennsylvanian and the cessation by earliest Permian are nearly identical with those processes observed in the Wichita-Amarillo-Anadarko system. Central Colorado and Taos Troughs Along the northeast margin of the Uncompahgre uplift, which is bordered by the Taos and Central Colorado troughs, late Paleozoic structures have been obscured by Laramide deformation and Tertiary volcanism. Sediments within the Taos and Central Colorado troughs are generally similar and the troughs are thought to have been periodically connected (Lindsey et al., 1986), although Soegaard (1990) and Soegaard and Caldwell (1990) suggested that they were separated by an arch in the postmiddle Desmoinesian. Morrowan­Wolfcampian sediments within the basins thicken southwest toward the Uncompahgre uplift, reaching a maximum thickness of more than 5000 m in the Central Colorado trough (DeVoto et al., 1988). In the Central Colorado trough these sediments also thicken to the northeast, toward the Frontrange uplift. Subsidence on top of an eroded Mississippian carbonate platform began in the Morrowan with deposition of deepwater marine shales and turbidites in the basin center and coarse clastic rocks in more marginal parts of the basin (DeVoto et al., 1988). In the Central Colorado trough coarse clastic sediments, shed from the Uncompahgre uplift and including fan and fandeltaic feldspathic conglomerate and arkosic sandstone derived from Precambrian basement, were deposited along the southwestern edge of the basin beginning sometime between the Morrowan and the Desmoinesian. Coarse clastic sediments were also shed into the northeastern margin of the Central Colorado trough from the Frontrange uplift in the Morrowan­Wolfcampian. In the Taos trough, similar deposits appear in the Atokan­Desmoinesian. The Central Colorado and Taos troughs were filled and dominantly occupied by alluvial and fluvial sediments by the Early Permian (Wolfcampian). Although definitive structural data are generally lacking along the margins of the Taos and Central Colorado troughs, Soegaard (1990) interpreted the Taos trough to have formed by flexural subsidence in response to thrusting along the northeastern margin of the Uncompahgre uplift. Given the similarity in facies and subsidence between the Central Colorado and Taos troughs and the Paradox basin, and the observation that subsidence in these basins was coeval with basement-involved thrusting on the southwestern side of the Uncompahgre uplift,

we agree with his interpretation. Because the Central Colorado trough most likely had the same origin as the Taos trough, we infer that a zone of northeast-directed basement-involved thrusting along the northeastern margin of the Uncompahgre uplift was active at the same time as thrusting along its southwestern margin. Thrusting probably began in the Morrowan­Atokan, coeval with the initiation of basement uplift and basin subsidence. Deposition of undeformed Guadalupian terrestrial sediments over folded and faulted Desmoinesian­Wolfcampian strata on the east side of the Uncompahgre uplift indicates that deformation continued throughout the Late Pennsylvanian and had terminated by the Early Permian, consistent with subsidence and facies data from the basins (Sharps, 1955, 1962; Freeman, 1971; Freeman and Bryant, 1977; DeVoto et al., 1988). Unconformities that occur locally near the eastern margin of the Uncompahgre uplift also indicate tectonic activity between the Desmoinesian and Wolfcampian (Pierce, 1969; DeVoto et al., 1971; DeVoto and Peel, 1972). Southwest-directed thrusting of the Frontrange uplift over the northeastern part of the Central Colorado trough appears to have been active from the Morrowan to Wolfcampian, but the eastern side of the Taos trough is mainly unfaulted. Frontrange-Sierra Grande-Denver Basin System The Frontrange­Sierra Grande uplift forms a north- to northwest-trending region of late Paleozoic uplift (Figures 1, 2). This uplift is bordered on its western side by the Taos and Central Colorado troughs and on its eastern side by the Denver and Dalhart basins. However, sediments within the Taos trough onlap and pinch out against the Frontrange­Sierra Grande uplift, indicating that subsidence within the Taos trough was unrelated to active faulting along its eastern boundary (Figure 3) (Soegaard, 1990). In contrast, late Paleozoic sediments within the Denver basin pinch out eastward against the Cambridge and Las Animas arches and their northern extension (the Chadron arch), and thicken westward toward the Frontrange, forming an asymmetric basin with a maximum thickness of more than 1300 m. The thicknesses and geometries of sedimentary rocks filling the Denver basin indicate a histor y of subsidence similar to that described for the Paradox basin, with subsidence beginning in the Morrowan (or late Chesterian?) and ending in the Early Permian (Figure 2) (Mudge, 1967; DeVoto et al., 1971; DeVoto and Peel, 1972; Mallory, 1972a, b, 1975; Wilson, 1975; DeVoto, 1980). Coarse clastic debris, partially derived from exposed basement, likewise indicates uplift and

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erosion of the Frontrange uplift throughout the Pennsylvanian. South of the Denver basin, the Dalhart basin also formed a westward-thickening asymmetric basin in the Pennsylvanian. The Dalhart contains more than 700 m of Pennsylvanian sediments in its deepest part, and is partially continuous with the Anadarko basin. Influx of granite wash debris from the west by the Atokan indicates that uplift and erosion to basement of the Sierra Grande uplift had begun by this time. Deposition of granite wash continued into the Wolfcampian, when it prograded across the entire basin and onlapped the basement on the east side of the basin (Bravo dome). The basement uplifts were finally covered by shallow-marine deposits at the end of the Wolfcampian (Johnson et al., 1988). Data to constrain the nature of late Paleozoic structures along the east side of the Frontrange and Sierra Grande uplifts are generally absent because these structures have been largely overprinted by Laramide thrusting; however, the similarity in timing of uplift and subsidence, the asymmetric nature of the basins, and the facies present within the basins suggest that the Denver and Dalhart basins formed in a manner similar to the other paired zones of basement uplift and subsidence, and were bounded along their western margins by east-directed basement-involved Pennsylvanian thrust faults. Pedernal-Diablo-Orogrande System The Pedernal and Diablo uplifts are north- and north-northwest­striking basement highs that lie along the trend of the Uncompahgre uplift to the north (Figures 3, 4). The uplifts are bordered on the east by the Delaware basin and on the west by the Orogrande basin. The Orogrande basin developed on top of a Mississippian carbonate shelf that underwent regional down-to-the-south tilting prior to development of the north-northwest­trending Orogrande basin. In this asymmetr ic basin, Pennsylvanian and earliest Permian strata thicken to the east, toward the Pedernal-Diablo uplift (Figure 2). The earliest sedimentation in the Orogrande basin proper is comprised of easterly derived Early Pennsylvanian sandstones and chertcobble conglomerates, suggesting erosion of the Paleozoic sedimentary cover in the Pedernal and Diablo uplifts at this time (King and Harder, 1985). Desmoinesian­Missourian rocks in the Orogrande basin are mainly shallow-marine limestones with local lenses of arkosic sandstone and shale, suggesting initial unroofing of basement in the Pedernal uplift (King and Harder, 1985). Increased deposition of arkose continued through the late Virgilian­early Wolfcampian, indicating progressive

unroofing of basement. The maximum thickness of the Virgilian basinal sediments is about 750 m along the east side of the basin; the early Wolfcampian sequence is also thick and the Orogrande basin was filled during the late Wolfcampian (Greenwood et al., 1977; King and Harder, 1985). Late Wolfcampian shallow-marine limestone and nonmarine red beds (Hueco and Abo formations) were deposited over the Orogrande basin and part of the Pedernal uplift, indicating the cessation of uplift and subsidence by the earliest Permian. The facies boundary between marine limestones and the clastic wedge composed of the erosional debris from the Pedernal uplift migrated westward across the Orogrande basin in the Late Pennsylvanian. In addition, the facies boundar y between a shallow-marine shelf in the east and a more basinal area in the west also migrated westward during the Late Pennsylvanian, and in places these shallow-marine rocks were eventually folded, faulted, and uplifted along what is now the western margin of the Pedernal uplift (e.g., in the Sacramento Mountains) (Pray, 1961; Greenwood et al., 1977). East of the Orogrande basin, pre-Permian, Pennsylvanian, and older rocks form tight folds and faulted folds on the Pedernal uplift (and its southward extension into the Otero uplift). The folds are asymmetric toward the west (King and Harder, 1985), with eroded Precambrian basement contained locally in the cores of the folds. Drilling data along the western margin of the Pedernal uplift indicate that Pennsylvanian strata were folded into northwest-trending anticlines (e.g., the Sacramento Mountains) (Otte, 1959a; Pray, 1959). Chert conglomerates and quartz sandstone were shed from the Pedernal uplift into the Orogrande basin from the Morrowan to the Virgilian, followed by arkosic detritus from the late Virgilian to the early Wolfcampian (King and Harder, 1985). This indicates erosion, deformation, and uplift of the Pedernal uplift throughout the Pennsylvanian. On the Diablo uplift, Cambrian­early Wolfcampian sedimentary rocks were strongly folded and faulted and Precambrian basement was exposed and eroded before deposition of undeformed middle Wolfcampian strata (Hueco Formation). Taken together, the presence of Pennsylvanian folding and faulting on top of the Pedernal-Diablo uplift, the eastward-deepening asymmetry of the Orogrande basin, the basinward (westward) migration of all facies boundaries in the Late Pennsylvanian, and the progressive overprinting of the depositional system by folding and faulting along the east side of the basin (now the west side of the Pedernal uplift) strongly suggest westvergent basement-involved thrusting along the

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west side of the Pedernal-Diablo uplift. On the east side of the Pedernal uplift, the Humble 1 Huapache well penetrated the Huapache fault and encountered a repeated stratigraphic section of unknown age across a fault plane (Hayes, 1964; King and Harder, 1985). This finding suggests that east-vergent thrust or reverse faulting of the Pedernal uplift over the Delaware basin occurred along the east side of the uplift (see also Meyer, 1966). However, the down-to-the-east asymmetry of the Delaware basin indicates that the major structural controls on its subsidence lie along its eastern boundary, and we infer that overthrusting along its west side is probably of minor importance. The shedding of coarse clastic debris derived from the Pedernal uplift into both the Orogrande and Delaware basins during the Morrowan indicates that uplift, erosion, and probably deformation of the Pedernal uplift had begun by the Morrowan. Clearly, deformation was well underway by the Desmoinesian, when the basement of the PedernalDiablo uplift was exposed and eroded, and had finished by the middle Wolfcampian with the deposition of the shallow-marine Hueco Formation unconformably across faulted and folded Pennsylvanian and older rocks. Basins and Uplifts of Northeastern Mexico The late Paleozoic deformation of northeastern Mexico is poorly known, partly because almost none of the rocks deformed in the late Paleozoic are now exposed at the surface, and partly because the southwestern region has probably been modified by several tectonic events. In particular, late Paleozoic structures that truncated the southern part of the Cordilleran orogene may have extended into northwestern Mexico. Additionally, some workers have suggested that the postulated Jurassic­Cretaceous Mojave-Sonora megashear traverses northern Mexico and extends into the Gulf of Mexico, thereby truncating the Paleozoic structures of North America (Figure 1) (e.g., see Anderson and Schmidt, 1983). Nevertheless, the available geologic record is sufficient to show that late Paleozoic deformation in this region produced structures that have some similarities with the northwest-trending systems of basins and basement uplifts within interior North America. The Magdalena­Sierra del Nido uplift and the Pedregosa-Chihuahua basin probably represent the southwesternmost part of the paired systems of northwest-trending basins and basement uplifts that developed within interior North America in the late Paleozoic (Figures 1, 2). In the late Paleozoic the Pedregosa basin was probably connected southeastward with the Chihuahua trough

and even extended farther southeast into eastcentral Mexico (Lopez-Ramos, 1969; Thompson et al., 1978; Mellor and Breyer, 1981; Handschy, 1986; Torres-Roldan and Wilson, 1986; Handschy et al., 1987). Data for the Pedregosa basin are much better than data for the Chihuahua basin because almost none of the rocks in the Chihuahua basin are exposed at the surface; data come from several drill holes and scattered outcrops in east-central Mexico (for example, in the Las Delicias and Cuidad Victoria areas). Subsidence of the Pedregosa basin began in the Morrowan with regional southwestward tilting and erosion of the previously stable carbonate platform. Deltaic to shallow-marine conglomerate, sandstone, shale, and dark oolitic limestone were deposited in the Pedregosa basin in the Morrowan and the Atokan. Along the central part of the basin, deepmarine carbonate turbidites and debris flows interstratified with basinal mudstone, and limestones, shales, and sandstones were deposited during the Desmoinesian­early Wolfcampian (Schuepbach, 1973; Greenwood et al., 1977; Kottlowski et al., 1988). Late Wolfcampian­Leonardian red bed facies and limestone facies (Earp Formation) unconformably overlie the Morrowan­early Wolfcampian strata (Horquilla formation), and probably represent filling and the end of subsidence in the Pedregosa basin. Thinning and pinch-out of the flysch section eastward onto the Diablo uplift and regional northeast onlap of Early Pennsylvanian strata indicate that the Pedregosa basin was asymmetrical, deepening to the southwest. Within the Chihuahua basin and its continuation southeast into east-central Mexico, thick flysch or flysch-like beds and local volcaniclastic sediments are considered to be middle and late Paleozoic in age, but the oldest sediments with fossil control yield Morrowan conodonts (see Lopes-Ramos, 1969; Wilson et al., 1969; Ramirez-Ramirez, 1984; McKee et al., 1988). Sedimentation in the Chihuahua basin continued until the middle Permian (Guadalupian) with deposition of chaotic breccias and flysch. A northwesterly trend for the Chihuahua basin is inferred only from the local trend of depositional features in the Las Delicias area (McKee et al., 1988). The southwestern boundary of the Pedregosa basin is well defined by the northwest-trending Magdalena­Sierra del Nido uplift, but there is little information about the southwestern boundary of the Chihuahua basin. The Magdalena uplift was an active area of erosion during the late Paleozoic, and not buried until the Leonardian (King, 1939; Cooper and Arellano, 1946; Navarro and Tovar, 1974). To the south, uplift of the Sierra del Nido region in the late Paleozoic is suggested by northeast-directed sediment transport and by northeast-vergent axial planes of synsedimentary slump structures in the

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Figure 9--Comparative timing of events in the Ouachita-Marathon region and in interior North America and western Mexico in the late Paleozoic. Thinner black bars show age of thrusting along the basin margins (for the Ouachita orogene, only thrusting proximal to the margin is shown); heavier black bars show timing of subsidence in the interior basins and the foredeep of the Ouachita orogene. V = volcanic activity in western Mexico, as described in the text.

late Paleozoic sediments adjacent to the uplift (Handschy, 1986; Handschy et al., 1987). The southern part of the uplift is poorly understood and was probably dismembered and displaced by the Mojave-Sonora megashear. In the Sierra del Nido region, the uplift can be clearly shown to be a northeast-directed basementinvolved thrust belt accompanied by east-vergent folding of Paleozoic sedimentary rocks. Several major thrust faults involve Early Permian flysch and Precambrian crystalline rocks (Bridges, 1971; Handschy, 1986; Handschy et al., 1987). In the northernmost Chihuahua basin, the age of uplift and deformation must be inferred from the sedimentary record. The thick, late Paleozoic flysch sequence immediately east of the Sierra del Nido uplift changes upward from quartz-rich to feldsparrich composition with granite debris, indicating uplift and unroofing of basement occurred in the Pennsylvanian (Mellor and Breyer, 1981). Later, during the Laramide orogeny in the Cretaceous, only Precambrian and Early Permian clasts are found within the Cretaceous synorogenic sediments, indicating that the entire pre-Permian section was eroded prior to the Cretaceous. This absence of prePermian Paleozoic clasts also indicates major uplift and erosion of the Sierra del Nido uplift in the preCretaceous, probably by the Early Permian. Farther southeast, the late Paleozoic flysch sequence in the Chihuahua basin contains fragments of crystalline basement (de Cserna et al., 1968). Lacking direct data to date the initiation of uplift, we infer that uplift and deformation began in the Morrowan, synchronous with subsidence of the basins. Uplift of the Magdalena­Sierra del Nido

region probably ceased in the late Wolfcampian­ Leonardian, as is suggested by deposition of late Wolfcampian­Leonardian red beds and shallowwater limestones over the Pedregosa basin and part of the Magdalena uplift (Kottlowski et al., 1988). Deformation, uplift, and subsidence in the Chihuahua basin ceased later than subsidence of the Pedregosa basin, possibly as late as the Guadalupian or the Ochoan (Flawn and Diaz G., 1959). The asymmetric geometry of the PedregosaChihuahua basin system, the presence of folds and basement-involved thrust faults along the east side of the Sierra del Nido uplift, and the late Paleozoic (Pennsylvanian) age of basement uplift and basin subsidence suggest that late Paleozoic deformation in this region involved northeast-directed thrusting of the Magdalena­Sierra del Nido uplift over the Pedregosa-Chihuahua basin. Age constraints, although not always definitive, suggest that deformation began at the same time as thrusting within interior North America (Morrowan), and in the Pedregosa basin ended at the same time (middle Wolfcampian). However, deformation along the margin of the Chihuahua basin appears to have lasted slightly longer than the deformation to the northeast, until the middle Permian (Figure 9). INTERPRETATION An overview of the late Paleozoic interior basins and uplifts reveals a remarkable coincidence in timing, geometry, and apparent structural style, and strongly suggests that these paired regions of basin subsidence and basement uplift form part of

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a unified system of regional deformation extending across much of Oklahoma, Texas, New Mexico, Colorado, and Utah. The north-northwest­ to northwest-trending elongate basins thicken asymmetrically toward the adjacent region of basement uplift (Figure 3) from which they are separated by synsedimentary faults with large vertical separation (usually several kilometers or more). On the opposite side of the basins, subsidence was generally accompanied by arching, mostly without faulting, of basement, and strata commonly pinch out onto these basement arches (except in the Delaware basin, which has faults on both sides of the basin). Over the entire region of late Paleozoic interior deformation, the initiation and cessation of basin subsidence and basement uplift were synchronous, beginning in the Chesterian­Morrowan, continuing through the Pennsylvanian, and ending in the Wolfcampian, although some minor postWolfcampian deformation has been found in a few places (e.g., along the southern extension of the Uncompahgre uplift) (Figure 9). Significant exceptions occur mainly in regions where there was a local reorganization of uplift and subsidence in the Desmoinesian; for example when subsidence was transferred from the Ardmore basin to the Marietta basin and when some of the faulting on the west side of the Central basin platform was transferred to the east side and subsidence began in the Midland basin. In addition, dur ing the Chesterian­Wolfcampian, the most rapid phases of basin subsidence and probably the most active periods of tectonism appear to have been diachronous throughout the region of late Paleozoic interior basins and uplifts (P. Kluth, 1995, personal communication). Documentation of the most active periods of tectonism along the margins of the late Paleozoic basins is complicated by the recognition that tectonism and basin subsidence do not correlate directly with the thickness of basin sediments, particularly in regions where the basins remained sediment starved for extended periods of time. Although the documentation of the periods of most rapid deformation is beyond the scope of this paper, we note that similar diachroneity can be observed in the Laramide deformation of the Rocky Mountains, because compressional strain is partitioned nonuniformly in space and time. All of the basement-involved uplifts show evidence for folding and faulting in the Pennsylvanian and Early Permian, coeval with the time of uplift; reverse faults and thrust faults have been drilled over the edges of many of these uplifts. In some of the basin systems, thrusting of the basement uplifts over the adjacent basins has been clearly documented by seismic reflection and drilling data; for example in the Amarillo-Wichita-Anadarko system, the Central basin platform-Delaware-Midland system, and the

Uncompahgre-Paradox system. As yet, we have little or no direct evidence for basement-involved overthrusting in the other basin systems, but the orientation, geometry, timing, and observable style of deformation are so similar throughout the entire region of late Paleozoic interior deformation that it seems highly likely that all of these basins have a common origin related to thrusting of basement over the margins of the basins. In addition, the orientation of the basins and uplifts is such that thrusting appears to have occurred in response to northeast-southwest­directed compressional stresses, although it is clear that inhomogeneities and preexisting zones of weakness in the crust are important in controlling the exact distribution of deformation. In map view and in cross section, the late Paleozoic interior deformation has produced structures and geometries similar to those produced during Laramide deformation of the Rocky Mountains (Figures 3, 10). In the Rocky Mountain region, basement-involved Laramide thrusting is clearly established, although prior to our acquisition of good subsurface data for this region, arguments in the literature about the origin of these paired basins and adjacent basement uplifts were similar to arguments that per sist for the late Paleozoic deformation of interior North America. Because the tectonic styles of the systems are so similar, and because a part of the region of late Paleozoic deformation in interior North America includes the Ancestral Rocky Mountains proper, we use the term "greater Ancestral Rocky Mountains" to describe the entire system of late Paleozoic basins and adjacent basement uplifts within interior North America. Laramide deformation of the Rocky Mountains occurred during eastward subduction of oceanic lithosphere under the North American continent, and is generally thought to be related to the geometry of the subducted slab in Laramide time (Lipman et al., 1971; Burchfiel and Davis, 1975; Coney, 1976; Dickinson and Snyder, 1978). Similarly, in the late Paleozoic, the western, eastern, and southern margins of North America were ringed by active margins along which deformation was approximately coeval with basement-involved shortening within the continental interior. The points that we address here are how the late Paleozoic deformation within the greater Ancestral Rocky Mountains is related to coeval events along the plate margins, what these events are, and their locations along the plate margin. LATE PALEOZOIC TECTONIC EVENTS OF PERIPHERAL NORTH AMERICA Late Paleozoic events along the eastern, western, and southern margins of North America have been

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Figure 10--Map view comparing the geometry of the modern basement-involved deformation in the Sierra Pampeanas, South America (left), to the early Eocene deformation of interior North America (Rocky Mountains) (center), and to the Pennsylvanian deformation of interior North America (greater Ancestral Rocky Mountains) (right). Black areas show regions of uplifted basement. Heavy dashed lines with barbs indicate trench position or inferred position, and light lines with open barbs show position of thin-skinned thrust belts. Area inside the rectangle shows region of flat-slab subduction or possible flat-slab subduction. Parts of the figure are modified after Jordan et al. (1983).

studied extensively, and the timing of these events is relatively well known (for discussion and references, see Sloss, 1988; Hatcher et al., 1989a, b; Burchfiel et al., 1992a, b). Along the eastern margin of North America, collision with the combined African/South American (Gondwana) continent occurred in Mississippian­Pennsylvanian, at least along the southern part of this margin, producing the northeast-trending Appalachian orogenic system. Along the southern part of the Appalachian orogene, postcollisional convergence probably continued into the Wolfcampian, producing extensive shortening and basementinvolved over thr usting within the Nor th Amer ican cr ust (Hatcher et al., 1989a, b). Although some authors have tried to relate deformation of the greater Ancestral Rocky Mountains to events within the Appalachian orogene (e.g., Budnick, 1986), we find major problems with this interpretation. Fir st, shor tening within the Appalachians, which clearly did involve shortening and deformation of cr ystalline basement, occurred in a northwest-southeast direction. This shortening direction is perpendicular to the northeast-southwest direction inferred for shortening within the greater Ancestral Rocky Mountains. Second, the Appalachian orogene is located a considerable distance (more than 3000 km) from the western part of the greater Ancestral Rocky Mountains. For these reasons, we consider it

unlikely that thrusting along the Appalachian orogene caused the defor mation of the greater Ancestral Rocky Mountains. Late Paleozoic tectonic activity along the western margin of North America offers little possibility as a cause for Ancestral Rocky Mountains deformation. Convergence between oceanic plates did occur throughout most of the late Paleozoic, but it was mostly in offshore island arcs and did not cause deformation within the North American craton or its passive margin. In fact, Miller et al. (1992) presented considerable evidence that the western margin of North America was subject to extensional deformation during most of the period Ancestral Rocky Mountain deformation was active. Along the southwestern part of the Cordilleran margin in southern California, extending into western Mexico, left-lateral strike-slip deformation with important extensional and minor convergent components was active during the late Paleozoic (Stone and Stevens, 1988; Walker, 1988), a style of tectonism difficult to relate to the Ancestral Rockies. Only two major convergent events are recorded within the North American craton and continental margin of the Cordilleran orogene: the Antler orogeny during the Mississippian and the Sonoma orogeny during the Late Permian­Early Triassic. The former is too old and the latter too young to be related to most of the deformation within the Ancestral Rocky Mountains (see Miller et al., l992). Within

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the Cordilleran miogeocline, and locally extending into offshelf regions, mild folds and faults and major subsidence took place in the late Paleozoic, but the location of these features suggests they are the westernmost part of the Ancestral Rocky Mountain deformation rather than related to the cause of the deformation (Burchfiel et al., 1992). Historically, most attempts to relate the deformation of the greater Ancestral Rocky Mountains to events around the periphery of North America have focused on events along the southern margin of North America, where thrust sheets of the Ouachita-Marathon orogenic belt were emplaced onto the North American craton in the late Paleozoic (see, for example, Burchfiel, 1978; Coney, 1978; Casey, 1980; Kluth and Coney, 1980, 1981; Goldstein, 1981, 1984; Thomas, 1983; Viele, 1983, 1986; Kluth, 1986; Soegaard and Caldwell, 1990). For the most part, supporting arguments have been based on the broad temporal coincidence of events within the two regions and on their spatial proximity. However, evidence from northeastern and east-central Mexico suggests that a late Paleozoic active margin also existed along the southwestern margin of North America, adjacent to and broadly contemporaneous with shortening in the plate interior. Examining the timing and nature of deformation within each of these zones of peripheral deformation, and attempting to identify modern tectonic analogs for each of these plate boundary systems, helps explain our conclusions. Late Paleozoic Deformation in the Ouachita-Marathon Orogenic System The Ouachita-Marathon orogenic belt is a thinskinned foreland fold and thrust belt on the late Paleozoic passive margin of southern North America (Figure 1). With the exception of a few local features that involve the crystalline basement (such as the Devil's River uplift), deformation within the Ouachita-Marathon belt affects predominantly sedimentary rocks from the distal part of the North American passive margin and allochthonous synorogenic, mainly deep-water, sedimentar y sequences that have been thrust onto North America. Although the Ouachita-Marathon belt commonly has been regarded as the result of a continent-continent or arc-continent collision along a south-dipping subduction boundary, the geologic features of the Ouachita-Marathon orogenic belt are perhaps more consistent with those of an accretionary prism obducted on the edge of the North American continent (Flawn et al., 1961; Vernon, 1971; Nicholas and Rozendal, 1975; King, 1980; Moore et al., 1981; Lillie, 1985; Leander and Legg, 1988; Thomas, 1988, 1989; Arbenz, 1989a, b; Keller

et al., 1989a, b; Nicholas and Waddell, 1989; Viele, 1989; Viele and Thomas, 1989). Other distinctive features of the Ouachita-Marathon orogenic belt include the absence of extensive involvement of crystalline basement in deformation, lack of highgrade metamorphism, lack of significant erosion, low topographic elevation (as evidenced by its present elevation of around 1 km and adjacent sedimentary sequences that show no evidence of high topographic relief during orogeny), and lack of protracted shortening of continental crust following the initial emplacement of thrust sheets onto the craton (see timing constraints section). The Ouachita-Marathon orogenic belt mainly consists of north-northwest­verging folds and thrust faults, although some south-vergent structures occur (e.g., the Broken Bow and the Benton uplifts) (Nielson et al., 1989). Along the Ouachita-Marathon orogenic belt, thrusting was toward the north and northwest, over the continental margin of North America and perpendicular to the trend of the OuachitaMarathon belt (King, 1977; Arbenz, 1989a, b; McBride, 1989; Nicholas and Waddell, 1989; Viele, 1989). Unless cited otherwise, data summarized below can be found in Sloss (1988), Hatcher et al. (1989a, b), and Bally and Palmer (1989). The timing of thrusting within the OuachitaMarathon belt and emplacement of thrust sheets onto the passive margin of North America are best constrained by the subsidence and sedimentation histories of the foredeep basin that formed immediately in front of the advancing thrust sheets, as well as by the age of synorogenic f lysch within the allochthonous thrust sheets. Locally, where the age of folding and thrusting within the belt can be constrained directly, it is in good agreement with ages inferred from the synorogenic and foredeep sediments. The initiation of convergence along the Ouachita-Marathon belt is somewhat uncertain, but was underway by the Mississippian, as evidenced by the presence of synorogenic, deep-water flysch of Mississippian age (e.g., the Stanley Formation and the Tesnus Formation) (Ross, 1986; McBride, 1989; Morris, 1989). This early thrusting presumably ref lects south-dipping subduction within an offshore, deep-water environment, and probably occurred at some distance from the North American continental margin. Except for the eastern part of the Ouachita belt, the advance of the orogenic system to a location near or at the edge of the continental margin is marked by the beginning of rapid subsidence of the previously stable passive margin, as indicated by an abrupt transition from dominantly shallow-water facies with slow deposition rates to a deep-water or starved environment. Within the easternmost part of the Ouachitas, deep-water conditions never developed within the Black Warrior basin, and the

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approaching thrust belt is marked by progradation of deltaic clastic sediments over the carbonate shelf in the late Meramecian (Thomas, 1988, 1989). To the west in the Arkoma basin, the passive continental margin began to subside in the Morrowan, accompanied by deposition of erratic limestone boulders indicating creation of surface relief (Shideler, 1970). Down-to-the-south normal faulting of the margin also began by the Morrowan (Arbenz, 1989b), indicating flexural bending of the crust in front of the advancing thrust belt. This initial subsidence of the Arkoma basin was succeeded by deposition of thick sequences of deep-water flysch during progressive subsidence and normal faulting of the continental margin. The earliest events in the Fort Worth basin region are somewhat uncertain because a large part of the basin has been overridden by thrust sheets (Keller et al., 1989b), but drowning of the shelf, down-to-basin normal faulting, and forelandward progradation of the clastic sequence are clearly recorded in the Atokan deposits (Thompson, 1982, 1988), indicating proximity of the advancing thrust belt by the Atokan. Toward the south and west in the Marathon region, deposition of synorogenic flysch continued through the Atokan. Subsidence of a continental margin carbonate platform in front of the advancing thrust belt is dated as late Desmoinesian in the Val Verde basin when the shelf was drowned and a late Desmoinesian orogenic clastic wedge prograded forelandward over the subsiding platform (Crosby and Mapel, 1975; Wuellner et al., 1986). These data show that advance of the thrust belt to a location proximal to the continental margin occurred diachronously from east to west (Figures 4, 9). Emplacement of thrust sheets onto the continental margin and cessation of thrusting are also diachronous from east to west, with thrusting terminating almost immediately after emplacement along most of the margin. Onshore thrusting within the Black Warrior basin probably ended in the late Atokan because thrust sheets of the Appalachian orogene override the Ouachita belt in the Middle Pennsylvanian and because the youngest foreland basin deposits known to be preserved in the Black Warrior basin are of Atokan age (Butts, 1926; Thomas, 1989). To the west in the Arkoma basin, emplacement of Ouachita thrust sheets onto the continental margin probably occurred in the late Atokan­Desmoinesian as indicated by the progradation of deltaic and fluvial deposits across the foredeep basin (Morris, 1974; Houseknecht and Kacena, 1983; Sutherland, 1988a). Thrusting in this region was probably finished in the Desmoinesian because the most external thrust faults in the belt cover middle Desmoinesian sediments. Additionally, asymmetrical, southward-deepening

subsidence of the foredeep basin had finished by the middle Desmoinesian, indicating that loading and flexural subsidence of the continental margin had finished by this time. Onshore thrusting in the Ouachita belt adjacent to the Fort Worth basin probably began in the Atokan as suggested by Atokan alluvial-coastal plain deposits preserved in the eastern part of the basin (Thompson, 1988). Thrusting here probably terminated in the Desmoinesian because the asymmetric foredeep basin was filled by coastal-plain deposits and platform carbonates. Within the Marathon region to the west and south, the tectonic evolution of the thrust belt is slightly more complex. Emplacement of thrust sheets onto the continental margin probably occurred in the late Desmoinesian (or Missourian?), because a late Desmoinesian orogenic clastic wedge prograded over the drowned carbonate platform in the Val Verde basin (Crosby and Mapel, 1975). Unlike the other segments of the Ouachita-Marathon orogenic belt, deformation within the Marathon region continued (or was renewed?) for a considerable time after the emplacement of thrust sheets onto the continental margin. From the Virgilian to the Wolfcampian, deformation of the foredeep region continued with northeast-directed overthrusting that involved crystalline basement in the Devil's River uplift (King, 1977; Wuellner et al., 1986; McBride, 1989). Whether this more protracted period of postemplacement thrusting in the Marathon region should properly be considered part of the Ouachita-Marathon orogeny is problematic. This period also can be interpreted as reactivation and overprinting of the Ouachita-Marathon orogenic belt proper by basement-involved thrusting along the margins of the Delaware basin and the Central basin platform, and the continuation of thrusting to the southwest into the Marathon region. A comparison of the Ouachita-Marathon orogenic system to young thrust belts developed within the Mediterranean region, where collision between Eurasia and a combined African-Arabian continent is underway, indicates that the Ouachita-Marathon area displays geological features nearly identical to some of the thrust belts in the Mediterranean region. Therefore, it may prove instructive to examine the tectonic settings within which these different Mediterranean thrust belt systems have evolved. (For a more detailed discussion of these Mediterranean systems, see Royden, 1993a, b.) Where parts of the Eurasian and African plates that contain thick continental crust are in contact, approximately east-west­trending thrust belts with high topographic mountains have been formed (e.g., the European Alps, Figure 11). These belts are typified by large amounts of erosion and denudation, exposure of high-grade metamorphic rocks at

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Figure 11--Configuration of Miocene­Holocene thrust belts within the Mediterranean region. Barbed lines show the approximate external limit of thrust belts, double-barbed lines indicate thrust belts formed at advancing plate boundaries, single-barbed lines indicate thrust belts formed at retreating plate boundaries, and vertical lines show approximate areas of upper plate extension above retreating plate boundaries. Large arrows indicate inferred direction of slab retreat at retreating plate boundaries.

the surface, development of antithetic thrusting, extensive deformation of crystalline basement to mid-crustal depths, a protracted history of postcollisional convergence, and molasse deposition in front of the advancing orogenic belt. The extensive postcollisional shortening across these belts (which in the Alps began in the Eocene­Paleocene and has lasted more than 40 m.y.) reflects the continuing convergence of Europe and Africa, and deformation appears to be driven by large compressional stresses transmitted across the orogene. Between these zones, where the thick continental crust of Eurasia and Africa has already collided, are zones where the thick continental crust of the two continents has not yet come into contact. These regions are floored by thin continental or oceanic crust and are covered by relatively deep water. In these areas Miocene­Holocene deformation has been dominated by thrusting and subduction along arcuate belts with low topographic relief and by regional extension within the upper plate of the subduction system (e.g., the Carpathian, Apennine, and Hellenic thrust belt systems). These belts are typified by little erosion or denudation, low-grade to no metamorphism, little to no involvement of crystalline basement in shortening, and flysch deposition in front of the advancing orogenic belt. These thrust belts are active only so long as they involve subduction of a deep-water foreland, and thrusting generally stops when the

thrust belt is emplaced onto a shallow-water or cratonic foreland. This observation is also consistent with gravity profiles across these subduction boundaries; these profiles show that subduction is driven primarily by the high density of the subducted slab (Royden, 1993a, b). Subduction in these belts does not occur within a regime of large horizontal compressive stresses; rather, horizontal shortening occurs only in a narrow zone at the toe of the advancing accretionary wedge, whereas deformation within the upper plate is dominated by extension parallel to the direction of subduction and convergence. This explains why basement-involved shortening is seldom observed in these belts, either beneath the thrust belt or in its foreland. In summar y, comparison of the OuachitaMarathon orogenic system to young thrust belts within the Mediterranean region shows that the geologic style of the Ouachita-Marathon thrust belt is distinctly different from that of thrust belts that develop along zones of ongoing collision between converging continents (or at least collision between the parts of the continents that have thick crust). The Ouachita-Marathon geological features are characteristic of those displayed by the accretionarytype thrust belts developed in zones where collision of thick crust has not yet occurred and where subduction is driven by slab pull and accompanied by upper-plate extension. Although upper-plate extension behind the Ouachita-Marathon belt has

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not been documented, several workers have suggested recently that the Desmoinesian subsidence within the upper plate of the Ouachita-Marathon belt (in east Texas) may be the direct result of upper-plate extension coeval with thrusting and subduction (Bally, 1989; Royden et al., 1990). If this correlation between the Ouachita-Marathon belt and accretionar y-type thrust belts of the Mediterranean is correct, then it is unlikely that emplacement of thrust sheets along the OuachitaMarathon belt was associated with the transmission of large compressional stresses into its foreland. Late Paleozoic Subduction in Northeastern Mexico Late Paleozoic plate boundary activity in northeastern Mexico is largely unknown, mainly because this part of Mexico was disrupted by motion along the Mojave-Sonora megashear and possibly by other events of late Paleozoic age. However, arcrelated volcanic rocks and volcaniclastic sediments of late Paleozoic age occur extensively in eastcentral Mexico (Bridges, 1962, 1971; de Cserna et al., 1968; Denison et al., 1971; Cunningham, 1975; Rowett and Hawkins, 1975; Ramirez-Ramirez, 1984; Handschy, 1986; Handschy et al., 1987; McKee et al., 1988). Volcanism in this area has been interpreted as Permian (King, 1944; Bridges, 1971), but fusulinids in contemporaneous limestone debris interbedded with volcaniclastic beds indicate that volcanism was active in the Desmoinesian (or pre-Desmoinesian) (McKee et al., 1988). Volcanism also occurred in the middle Permian (Guadalupian), as indicated by chaotic deposits containing andesite, dacite, welded tuff, laharic breccia, and red volcaniclastic siltstone and sandstone, and by a granitic pluton dated at 225 Ma (McKee et al., 1988). Volcanic clasts consist of andesitic, rhyolitic, and dacitic volcanic debris ranging in size from sand to great blocks 10­100 m long (McKee et al., 1988). The calc-alkalic chemistry of these volcanic rocks suggests that volcanism developed on continental crust (Jones et al., 1986; Jones and McKee, 1987) or on top of an older volcanic arc, as do inclusions of granite and schist in volcaniclastic debris flows. The large size of some of the volcanic clasts deposited in the Chihuahua basin indicates that eruption must have occurred immediately adjacent to the basin. This late Paleozoic volcanic arc testifies to the presence of a nearby subduction boundary of late Paleozoic age. In our interpretation, this arc is probably an Andean-type volcanic arc erupted through the continental crust, and implies northeast-dipping subduction of oceanic lithosphere

along a subduction boundary southwest of the Magdalena­Sierra del Nido uplift and the Chihuahua basin. Because much of dating on the volcanic arc comes from rock clasts within Pennsylvanian and Permian strata, we can state only that volcanism was active during or prior to the Desmoinesian and was also active in the Permian. WHICH SUBDUCTION BOUNDARY CAUSED INTRAPLATE SHORTENING? Arguments that bear most directly on the relationship between the late Paleozoic deformation of interior North America and deformation around the plate margins include the relative timing of deformation in each system, the spatial proximity and orientation of regional strain fields, structural relationships present where two of the tectonic systems overlap, and the existence of modern tectonic analogs in which intraplate deformation can be causally related to coeval events at the plate margins. Because deformation within the greater Ancestral Rocky Mountains begins and ends everywhere at about the same time and because the style of deformation and the direction of shortening are so similar across the entire region, we believe it unlikely that deformation within the greater Ancestral Rocky Mountains occurred in response to different tectonic events in different parts of the region, or that it occurred in response to different tectonic events operating at different times. If the deformation in the greater Ancestral Rocky Mountains was indeed caused by plate boundary processes along the margin of North America, then, in our opinion, it most probably was caused by processes that did not change very much in space or time during the deformation of the greater Ancestral Rocky Mountains. Although we cannot rule out the possibility that diachronous events along the Ouachita-Marathon belt exerted some control on the late Paleozoic deformation within its foreland, the nature, timing, and orientation of events along the Ouachita-Marathon belt make it very difficult to relate its emplacement to most of the deformation within the greater Ancestral Rocky Mountains. First, the geological features of the Ouachita-Marathon thrust belt are typical of thrust belts developed in regions of incomplete continental collision where the parts of the continents containing thick crust have not yet come in to contact. Subduction in these belts usually does not occur within a regime of large horizontal compressive stress and, except at the toe of the thrust belt, deformation within the upper plate is dominated by regional extension and subsidence. Such thrust belts are generally not associated with deformation of crystalline basement or with the development of

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compressional structures within the foreland. We know of no instances in which the emplacement of such a thrust belt can be shown to have caused extensive deformation within its foreland region. Second, even if the emplacement of thrust sheets in the Ouachita-Marathon belt were accompanied by horizontal compressive stresses large enough to have caused extensive deformation within their foreland, there is no part of the Ouachita-Marathon belt in which both the orientation of the thrust belt and the direction of thrusting can be simultaneously related to northeast-southwest shortening in the greater Ancestral Rocky Mountains (Figures 4, 9). Within the easternmost (Black Warrior) and westernmost (Val Verde) parts of the Ouachita-Marathon orogene, thrust sheet transport appears to have been toward the north­northeast. However, most of the shortening within the greater Ancestral Rocky Mountains did not occur directly in front of these parts of the thrust belt, but rather occurred in parts of the foreland that lie laterally at a significant distance from the thrust belt. Thus, any horizontal compressional stresses generated across the Black Warrior region would not produce northeast-southwest compression within the greater Ancestral Rocky Mountains, and horizontal compressional stresses generated across the Val Verde region would not produce northeast-southwest compression within the wester n par t of the greater Ancestral Rocky Mountains (for example, in the Paradox and Orogrande regions). In the central part of the Ouachita-Marathon orogene, the direction of thrust sheet transport is nearly perpendicular to the northeast-southwest direction of shortening within the greater Ancestral Rocky Mountains. Third, in no part of the Ouachita-Marathon belt does emplacement occur throughout the entire period of deformation in the greater Ancestral Rocky Mountains. Diachronous emplacement along the entire belt does span the period of deformation within the greater Ancestral Rocky Mountains, but because emplacement is occurring in different directions at different times, it seems unlikely that this could have generated any sort of consistent stress field throughout the emplacement history of the Ouachita-Marathon belt. Fourth, where there is direct contact between the greater Ancestral Rocky Mountains and the Ouachita-Marathon belt at the southeast end of the Anadarko basin system, northeast-trending structures within the thin-skinned Ouachita belt do not interfere with or modify northwesttrending basement-involved structures within the Anadarko basin system (Arbuckle region) (Granath, 1989). Likewise, west of the Val Verde basin southwest-trending thin-skinned structures in the Marathon thrust belt are also at right angles

to northwest-trending basement-involved structures that pass beneath the Marathon belt (Reed and Strickler, 1990). (Only in the Val Verde region does the Marathon thrust belt contain northwesttrending structures that involve crystalline basement. These structures, which are anomalous within the Ouachita-Marathon belt, may be part of the greater Ancestral Rocky Mountains system and thus represent the southeastward continuation of structures within the Delaware basin and Central basin platform rather than primary structures of the Marathon thrust belt.) Taken together, these considerations make it difficult to drive the late Paleozoic deformation of the greater Ancestral Rocky Mountains by emplacement of the Ouachita-Marathon belt along the southern margin of North America. Another alternative is that the late Paleozoic deformation of interior North America is the result of plate boundary processes operating southwest of the greater Ancestral Rocky Mountains. Here, in east-central Mexico, a volcanic arc of late Paleozoic age indicates the presence of a coeval subduction boundar y to its southwest. Because the subduction boundary itself is not known from the geologic record, the tectonic style of this late Paleozoic margin is essentially unknown and must be inferred from its volcanic arc and from coeval deformation within its hinterland. In this paper, we speculate that this late Paleozoic subduction boundary formed a northwest-trending Andean margin, and that the greater Ancestral Rocky Mountains formed as the result of intraplate deformation within the overriding plate (hinterland) of the subduction boundary. Although the exact orientation of the late Paleozoic arc within east-central Mexico is not known, the arc appears to have been produced by subduction of oceanic lithosphere beneath North America, indicating approximately northeastdipping subduction. This subduction probably was not the direct continuation of the OuachitaMarathon belt because the vergence of subduction in the Ouachita-Marathon belt is opposite. The eastdipping subduction boundary in northern Mexico may have been continuous with the subduction boundary known to have persisted along the western margin of North America throughout the late Paleozoic. This subduction boundary also dipped to the east, but was associated with an offshore island arc separated from western North America by a significant, but unknown, width of oceanic crust (see Miller et al., l992). The back-arc oceanic region east of this subduction boundary appears to have narrowed southward into southeastern California, but we found no clear evidence of how the arc or its associated subduction zone would have connected with a late Paleozoic arc subduction boundary in northern Mexico.

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Horizontal compressional stresses generated across a northwest-trending, northeast-dipping subduction boundary located along this part of the North American continental margin would be consistent with generation of northeast-southwest shortening within the greater Ancestral Rocky Mountains. The duration of subduction is not known, but minimum age constraints on the timing of subduction are provided by the ages of the arc volcanic rocks. These data show that arc volcanism was active in the Desmoinesian or pre-Desmoinesian and in the Early Permian. Because the initiation of arc volcanism has not been dated, volcanism and subduction were possibly active throughout the Pennsylvanian and were coeval with the entire period of deformation in the greater Ancestral Rocky Mountains. Alternatively, it is possible that volcanism was active in the pre-Pennsylvanian and in the Early Permian, with a volcanic gap throughout the Pennsylvanian. Obviously, more extensive dating of late Paleozoic arc rocks in eastcentral Mexico is needed to better constrain the duration of arc volcanism and subduction. At least two well-documented examples of basement-involved intraplate shortening have been related to plate boundary activity along an adjacent Andean margin. One instance of shor tening occurred within the Rocky Mountain region in the Late Cretaceous to early Cenozoic (Laramide), during eastward subduction of oceanic lithosphere beneath North America (Armstrong, 1974; Burchfiel and Davis, 1975). The second example occurs along the east side of the Argentine Andes (Sierra Pampeanas), where east-west shortening of the South American craton has been attributed to eastward subduction of oceanic lithosphere beneath South America (Jordan et al., 1983). In both cases, the overall direction of intraplate shortening is perpendicular to the trend of the subduction boundary and shortening involves parts of the overriding plate across a region more than 500 km wide and at distances in excess of 500 km from the subduction boundary (Figure 10). It is striking that the general geometry and configuration of basement uplifts in these younger systems is very similar to those produced during the late Paleozoic deformation of the greater Ancestral Rocky Mountains. Thus, at least two well-documented modern analogs exist to support the hypothesis that the greater Ancestral Rocky Mountains formed as the result of basement-involved shortening within the hinterland of an active Andean margin in east-central Mexico. Because late Paleozoic events along the southwestern margin of North America are so poorly known, detailed speculations about the exact relationship between events along this plate boundary and events within the greater Ancestral Rocky

Mountains are probably not warranted. However, it is worth noting that both of the well-documented examples of intraplate shortening behind a coeval Andean margin are associated with subduction of the downgoing plate at a very shallow angle and with a temporal and spatial gap in the associated volcanic arc in the region of flat-slab subduction (Jordan et al., 1983). The late Paleozoic age of the greater Ancestral Rocky Mountains and the timing of the volcanic arc in east-central Mexico likewise may have been related to a period of flat-slab subduction along the southwest margin of North America, and may have been preceded or followed by periods of subduction at steeper dips. Thus, the volcanic arc observed in east-central Mexico may represent only the easternmost extent of a longerlived arc system. Alternatively, the arc may record the entire duration of subduction along this part of the North American continental margin, perhaps terminating with the initiation of strike-slip motion on the Mojave-Sonora megashear. Although the morphologies of the greater Ancestral Rocky Mountains, the Laramide Rocky Mountains, and the Sierra Pampeanas were similar during the times at which they were tectonically active, there is a striking difference between the posttectonic morphology and elevation of the greater Ancestral Rocky Mountains and the Laramide Rocky Mountains (the Sierra Pampeanas are still active, so they cannot be used for comparison). The topographic elevation of the Laramide Rocky Mountain region is far above sea level, and uplift of the region from approximately sea level to its present elevation is thought to have occurred during the Cenozoic, during or after Laramide deformation. In contrast, the greater Ancestral Rocky Mountains subsided and were buried by thick posttectonic sedimentary strata shortly after the cessation of deformation in the earliest Permian. The significance of the differing posttectonic development of the Laramide Rocky Mountains and the greater Ancestral Rocky Mountains is unclear. It may be that the posttectonic uplift of the Laramide Rocky Mountains is unrelated to Laramide shortening and instead is due primarily to the hot-spot activity that has affected much of the west-central United States in the Miocene­Holocene (e.g., Morgan, 1972). Nevertheless, the regional posttectonic subsidence of the greater Ancestral Rocky Mountains is puzzling because one would normally expect regional uplift during periods of thrusting and crustal shortening. Thus, even if the basin regions subsided during shortening, one would expect a corresponding uplift of the basement highs to give a regionally averaged increase in topographic elevation. Without further data, we can only speculate that perhaps processes connected with cessation of flat-slab subduction beneath the

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greater Ancestral Rocky Mountains may have produced regional subsidence. In summary, we interpret the greater Ancestral Rocky Mountains to have formed by late Paleozoic intraplate shortening within the hinterland (upper plate) of a coeval Andean margin along the southwestern periphery of North America. In this interpretation, the greater Ancestral Rocky Mountains would be directly analogous to the younger system of intraplate shortening developed in the Laramide Rocky Mountains. In this sense, the greater Ancestral Rocky Mountains may truly be considered to be ancestral to the Rocky Mountains because they share a common tectonic origin and structural style, as well as a common geographic location. IMPLICATIONS FOR OIL AND GAS EXPLORATION The traditional interpretation of uplifted regions within the greater Ancestral Rocky Mountains as basement platforms bounded by steep reverse or normal faults has discouraged extensive drilling through the crystalline basement. Although a number of wells have drilled into and through basement rocks at the edges of some of the uplifted regions, the interpretation of the basement-bounding faults as steep structures provides little room for source or reservoir rocks beneath the edges of the basement blocks. Our interpretation that deformation throughout the greater Ancestral Rocky Mountains occurred by basement-involved overthrusting on gently to moderately dipping thrust faults suggests that the potential for subthrust exploration for oil and gas beneath the basement may be very much greater than is generally supposed. In particular, the reinterpretation of the basement-bounding faults as gently to moderately dipping thrust faults implies that there may be a much greater width of source and reservoir rocks present beneath the basement-involved thrust sheets than would be possible beneath steeply dipping reverse faults, and that reservoir rocks may be present at a sufficiently shallow depth to allow penetration of reservoir rocks by drilling. We therefore suggest that future exploration for new oil and gas fields along the margins of these uplifted basement blocks may benefit from the exploration approach used in fold and thrust belts around the world. For example, in the region of the Central basin platform, the Paleozoic section contains excellent source beds, reservoir rocks, and seal, and the thermal history of the source beds has been favorable for maturation of hydrocarbons. Hydrocarbon exploration in this region has focused primarily on the location of structural (and stratigraphic) traps within the sedimentary section above and adjacent

to the Central basin platform. In our interpretation, the Central basin platform was uplifted during the late Paleozoic basement-involved thrusting on moderately to gently dipping faults (see seismic sections in Figure 7). Thus, the Midland basin is interpreted as a foredeep basin formed in front of west-dipping basement-bounding thrust faults on the east side of the Central basin platform. Sediments are involved in thrusting in the frontal parts of the thrust system, where the main thrust fault cuts upsection from basement rocks into the overlying sediments. In our interpretation, folds within the internal parts of the Midland basin are part of the same deformation system and formed as frontal and ramp anticlines above thrust faults within the sedimentary section (these faults are probably splays off of the main basement-bounding thrust fault). Folds on top of the Central basin platform probably occur above a series of secondary thrust faults, many of them antithetic to the main west-vergent thrust fault. As in many other overthrust regions, significant quantities of hydrocarbons may be trapped within the sedimentary rocks that have been overthrust by crystalline basement in the frontal parts of the thrust system. Because source, seal, reservoir, and maturation are known to be regionally present within the Paleozoic section of the Central basin platform, any structures in Paleozoic rocks that have four-way closure are likely to be productive. We suggest that closed structures in Paleozoic rocks beneath the crystalline thrust sheets are similarly likely to contain hydrocarbons. If our interpretation that the basement-bounding thrust faults have a gentle to moderate dip is correct, then there would be considerable room beneath the crystalline thrust sheets for closed structures at depths sufficiently shallow to be accessible by drilling, as well as for a large region of available source material. Thus, exploration for such structures beneath the frontal parts of the crystalline thrust sheets may be more likely to locate new and larger oil and gas fields than has generally been supposed. In addition, basement-involved thrust faults also may occur within and in front of the main basinbounding thrusts faults, and structural traps also may be present beneath basement rocks in the internal parts of the uplifted regions, as well as beneath the adjacent basins. REFERENCES CITED

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