Syntectonic sediment loading and fold-thrust belt structural architecture: An example from the central Appalachians (USA)

The structural geometry of a growing thrust wedge can be influenced by many variables (see reviews in Buiter, 2012, and Graveleau et al., 2012), such as rock strength and pore fluid pressure (e.g., Mourgues and Cobbold, 2006), detachment strength (e.g., Costa and Vendeville, 2002; Smit et al., 2003; Malavieille, 2010; Simpson, 2010), preexisting stratigraphy and sedimentary basin taper (e.g., Boyer, 1995; Soto et al., 2003), and sedimentation and erosion (e.g., Mugnier et al., 1997; Malavieille, 2010). In particular, numerous studies have examined the effect of syntectonic wedgetop depositional load on the development and structural architecture of a growing thrust wedge (e.g., Marshak and Wilkerson, 1992; Storti and McClay, 1995; Mugnier et al., 1997; Simpson, 2006; Bonnet et al., 2007, 2008; Stockmal et al., 2007; Wu and McClay, 2011; Buiter, 2012; Graveleau et al., 2012; Fillon et al., 2013a, 2013b; Erdõs et al., 2015; Sun et al., 2019). Both analog and mechanical (numerical) models predict that syntectonic sedimentation controls the structural geometry of the fold-thrust belt. In general, with high sedimentation, thrust length increases and hence thrust spacing increases, while with low to no syntectonic load, thrust length and spacing are short (Marshak and Wilkerson, 1992; Storti and McClay, 1995; Stockmal et al., 2007; Bonnet et al., 2008; Fillon et al., 2013a, 2013b) or concentrated toward the hinterland (Sun et al., 2019). Basically, high sediment loads deposited in the foredeep proximal to the thrust wedge stabilize the wedge (negative-alpha basin of Fuller et al., 2006; Willett and Schlunegger, 2010), resulting in a jump of the thrust front creating a piggyback basin. The new active thrust then develops into a critical wedge, resulting in further sedimentation. With smaller loads, there is less wedge stabilization, and the thrust wedge develops closely spaced imbricate thrusts. These model observations are borne out in fold-thrust belts where syntectonic sedimentary rocks are preserved (e.g., Willett and Schlunegger, 2010; Fillon et al., 2013a; Sun et al., 2019).

However, in many ancient fold-thrust belts, erosion has removed most, if not all, syntectonic deposits and evaluating the effect of load on thrust wedge development is difficult without a burial-depth proxy such as apatite and/or zircon fission-track data (e.g., McQuarrie and Ehlers, 2017), vitrinite reflectance (e.g., Schegg, 1992; Hardebol et al., 2007), conodont alteration index (Epstein et al., 1977), or clay mineral alteration (e.g., Frey, 1987). Fluid inclusion micro-thermometry can also provide an estimate of the depth of fluid trapping, which can be used to determine the extent of eroded overburden (e.g., Evans, 2010; Evans et al., 2014). In the central Appalachians (northeastern United States), a large data set of maximum burial depth data derived from fluid inclusion microthermometry (Evans, 2010; Evans et al., 2012, 2014; this study) is used to provide a picture of the maximum syntectonic load in the central Appalachians during the Alleghenian orogeny. The analysis integrates the structural architecture and retrodeformation of the central Appalachians using a series of regional cross sections and retrodeformed cross sections. The goal is to investigate the relationship between the distribution of syntectonic load and the structural development and final architecture of the thrust wedge.

The central Appalachian fold- and- thrust belt (Fig. 1A) was formed during the late Carboniferous–Permian Alleghanian orogeny. It consists of a sublinear segment in the south with a structural trend of ~030° and an ~060°-trending segment in the northeast (Fig. 1A). This curve is the Pennsylvania salient, the origin of which is a matter of debate (e.g., Gray and Stamatakos, 1997; Macedo and Marshak, 1999; Wise, 2004; Wise and Werner, 2004; Ong et al., 2007). The structural style is controlled by a 2–3-km-thick Cambrian–Ordovician carbonate rock lithotectonic unit (Fig. 1B) that underlies 3–4 km of an Ordovician–Carboniferous mixed-lithology cover rock sequence (Fig. 1B) (Hatcher et al., 1989; Evans et al., 2016). The structural geometry of the fold-thrust belt is that of an active roof duplex of the Cambrian–Ordovician section and is characterized by locally significant forethrusting and coupled deformation of the cover rock sequence (Dunne and Ferrill, 1988; Ferrill and Dunne, 1989; Couzens-Schultz et al., 2003; Evans, 2010; Sak et al., 2012). The significant detachment folding and layer-parallel shortening that occurs above the roof thrust is similar to those of the Brooks Range in Alaska (Wallace and Hanks, 1990; Wallace, 1992, 1993; Wallace et al., 1997). The floor thrust of the duplex is in the Cambrian Waynesboro Formation, while the roof thrust is in the Ordovician Martinsburg (Reedsville) Formation (Gwinn, 1964, 1970). In the hinterland, two crystalline massifs characterized by Proterozoic metamorphic rocks dominate: the Blue Ridge thrust system in the southern segment, and the Reading Prong–Lebanon Valley nappe system in the eastern segment (e.g., Hatcher et al., 1989). These antiformal stacks are similar in style and structural setting to those shown as crystalline massifs in the Alps (Bonnet et al., 2007, 2008) and other fold-thrust belts (Malavieille, 2010).

The Blue Ridge massif is thrust over an imbricated section of the Cambrian–Ordovician lithotectonic unit (Harris et al., 1982; Evans, 1989), while the nature of the Reading Prong–Lebanon Valley nappe system is uncertain due to lack of subsurface control. Forelandward of the massifs is the Valley and Ridge province, which is dominated by imbrication of the Cambrian–Ordovician lithotectonic unit into a duplex and by detachment folding in the cover rock sequence (e.g., Gwinn, 1964, 1970; Evans, 2010; Sak et al., 2012). Large-displacement emergent thrust faults separate the massifs from the Valley and Ridge: the North Mountain and Blue Ridge thrusts in the south, and the Yellow Breeches thrust to the east. Emergent thrusts are rare in the Valley and Ridge and typically have less than a kilometer of displacement.

Seventeen regional cross sections (Fig. 1A; Plate 1) show that the structural geometry varies significantly from the 030°-striking southwestern segment to the 060°-striking northeastern segment. Each cross section is restored to its pre-thrusting configuration using a line-length and area balance of the Cambro-Ordovician carbonate lithotectonic unit. The cover rock sequence is assumed to have deformed by passive folding above the Cambro-Ordovician section and forethrusting along the Martinsburg detachment ahead of the imbricate thrusts.

In the east, in the Anthracite belt, low-amplitude, low-displacement fault-bend-style folds dominate the Cambrian–Ordovician section (Plate 1, sections 1–3). In the central part of the salient, the structural geometry is defined by a duplex with imbricate horses of Cambro-Ordovician carbonates that gradually pass into an antiformal stack of two to three carbonate thrust sheets near the Appalachian structural front and define the Nittany anticlinorium (Plate 1, sections 4–8). The duplex is characterized by closely spaced thrust imbricates and is referred to as the Juniata culmination (Fig. 1A; Gwinn, 1970; Pennsylvania culmination of Nickelsen, 1963; Rodgers, 1970).

In southern Pennsylvania, the eastern part of the belt is defined by a series of imbricated Cambro-Ordovician carbonate horses with leading-edge fault-propagation-style folds. These gradually pass into the Broadtop synclinorium to the west, and then two additional carbonate horses with similar leading-edge folds that compose the Wills Mountain anticlinorium toward the Appalachian structural front (Plate 1, sections 8–12). In the south, in West Virginia and Virginia, the eastern part of the belt is dominated by a complex series of foldthrust structures that die out toward the south and are gone by section 16 (Plate 1). To the west, the Wills Mountain anticlinorium gradually passes into a flat-on-flat structure (Wilson, 1989; Wilson and Shumaker, 1992; Smart et al., 1997; Evans, 2010) that has >20 km of duplicated Cambro-Ordovician lithotectonic unit (Plate 1, sections 14–17).

The Valley and Ridge province lacks significant syntectonic sedimentary rocks that could be used to evaluate the deformation history (e.g., Wiltschko and Dorr, 1983; DeCelles, 1994; DeCelles et al., 1995). Pennsylvanian-age sedimentary rocks are present within the Valley and Ridge in the Broadtop synclinorium and in the Anthracite belt. These rocks were deposited distal to the advancing thrust front, which was >100 km toward the hinterland at the time of deposition, and consist of deltaic deposits (Pottsville Formation through Conemaugh Group) in the Broadtop synclinorium and braided stream deposits in the Anthracite belt (Pottsville Formation through Lewellyn Formation). Permian and younger(?) syntectonic sediments are not preserved in the Valley and Ridge. However, 76 m of rocks interpreted to be Permian are found adjacent to the Appalachian structural front in Maryland (O’Harra, 1900; Berryhill et al., 1956; Brezinski and Conkwright, 2013). Forelandward of the Valley and Ridge is the Appalachian Plateau province, which preserves Pennsylvanian sedimentary sequences throughout and Permian rocks in southwestern Pennsylvania, eastern Ohio, and northern West Virginia. Prior to erosion, the thickness of these syntectonic rocks was ~1.5–4.5 km (Evans, 2010; Evans et al., 2014). These rocks are dominated by deltaic sequences with occasional marine incursions in the Pennsylvanian and early Permian (e.g., Donaldson and Shumaker, 1981; Donaldson et al., 1985; Ettensohn, 2004, 2008; Fedorko and Skema, 2013).

A total of 296 quartz, calcite, and more rarely barite fracture-filling cements (veins) were collected as part of a regional study from 236 sites throughout the Valley and Ridge province from all exposed stratigraphic levels (Fig. 2). In addition, data are included from 209 vein samples from 151 sites from Evans (2010), Evans et al. (2012, 2014), and other workers (Srivastava and Engelder, 1990, 1991; Lacazette, 1991; O’Kane, 2005; Castles, 2010; Lacek, 2015). The veins are all interpreted to be Alleghanian in age, are pre-folding to very early synfolding, and therefore record deformation conditions prior to and during folding and exhumation (e.g., Evans, 2010; Evans et al., 2012, 2014). The locations of the samples are shown on the cross sections in Plate 1 and on the map in Figure 2. The pre-deformation stratigraphic positions of the samples are shown in the restored cross sections in Plate 1.

The sample preparation and analytical methods are summarized in Evans and Battles (1999), Evans (2010), Evans et al. (2012, 2014), and Ferrill et al. (2020). Three kinds of fluid inclusions were found in the samples: two-phase aqueous inclusions containing a brine and a vapor bubble, single-phase inclusions containing CH₄ ± CO₂, and two-phase liquid hydrocarbon inclusions. Standard fluid inclusion microthermometry was performed on all of the vein samples in order to derive the fluid trapping conditions. Heating and freezing of the inclusions was conducted on a Linkham THMSG600 heating-freezing stage on a Leica DM2500P microscope. The accuracy of measurements is estimated at ±0.2 °C for T_hH (single-phase hydrocarbon inclusion homogenization), T_mIce (water ice melting), and T_mCO2 (CO₂ ice melting used to determine CO₂ content) and ±0.5 °C for T_hA (aqueous inclusion homogenization) and T_hLH (liquid hydrocarbon inclusion homogenization). All samples have data from multiple fluid inclusion assemblages. However, for this study, only data from the fluid inclusion assemblage that provides the greatest trapping depth in each sample are included here (fluid inclusion data and estimated burial depths can be found in Table S1 in the Supplemental Material¹).

The trapping depth of the fluid inclusions was determined by one of the three methods described in Evans (2010) and Evans et al. (2014). Two methods assume trapping at or near lithostatic conditions, while the third provides actual trapping conditions. As shown by Evans and Battles (1999) and Evans (2010), the probable range for late Paleozoic geothermal gradient for most of the Valley and Ridge province in the central Appalachians is 20–25 °C km⁻¹, while in the Anthracite belt, gradients are lower (16–22 °C km⁻¹). Except in the Anthracite belt, the fluid inclusion trapping pressures determined here use an average paleogeothermal gradient of 22.5 °C km⁻¹. Once the trapping pressure has been calculated for a sample, the trapping depth is determined using a lithostatic gradient of 26 MPa km⁻¹.

The geothermal gradient method uses the maximum homogenization temperature of aqueous fluid inclusions in an assemblage, and a trapping pressure is determined from the geothermal gradient range (Fig. 3). A surface temperature of 20 °C is assumed for the late Paleozoic in the central Appalachians. The aqueous inclusions are assumed to be methane saturated because all stratigraphic units contain methane fluid inclusions. Therefore, no pressure correction for the aqueous fluid inclusions is needed (Hanor, 1980) and the homogenization temperature is the trapping temperature (Table S1). If trapping conditions are actually less than lithostatic, then trapping would occur at lower pressure (Fig. 3). Therefore, this method provides a maximum estimate of the trapping depth. Where only liquid hydrocarbon inclusions are present, they are used to determine the trapping depth with the geothermal gradient method.

For single-phase CH₄ ± CO₂ fluid inclusions, the methane isochore method (Evans, 2010; Evans et al., 2014) is used to determine the trapping pressure (Fig. 3). The homogenization temperature of the inclusions (T_hH) in addition to the melting temperature of the CO₂ (T_mCO2, if present) is used to determine the density of the trapped fluids (van den Kerkhof, 1988, 1990; Kisch and van den Kerkhof, 1991; Thiéry et al., 1994), which is used to calculate an isochore (Duan et al., 1992a, 1992b; Bakker, 2003; Bakker and Brown, 2003). The range of trapping pressures is determined where this isochore meets the lithostatic gradient lines for 20–25 °C km⁻¹ (Fig. 3). However, the isochore extends beyond the 25 °C km⁻¹ lithostatic gradient line, and if conditions are less than lithostatic, trapping conditions would be higher (Fig. 3). Therefore, this method provides a minimum estimate of the maximum trapping depth. Because the CH₄ ± CO₂ inclusion homogenization values are a function of the density of the trapped fluid and hence the trapping pressure, only the highest-density (lowest T_hH) inclusion within an assemblage is assumed to represent the densest inclusion and hence the highest-pressure inclusion (Vityk and Bodnar, 1995, 1998). These inclusions provide the greatest trapping depth estimate (Table S1).

If both aqueous and CH₄ ± CO₂ fluid inclusions can be shown to be coeval in the same fluid inclusion assemblage, they would represent two immiscible fluids trapped at the same time. For these inclusions, the aqueous inclusion homogenization temperature–CH₄-CO₂ inclusion isochore intersection method may be used (Mullis, 1987), and a unique trapping temperature and pressure can be determined (Fig. 3). Such mutual trapping is uncommon and difficult to demonstrate (e.g., Goldstein, 2001; Chi et al., 2021). Only several fluid inclusion assemblages used in this study meet the criteria of coeval trapping (Table S1).

Stratigraphic thicknesses from published reports were used to estimate the stratigraphic thickness from the sampled unit to the top of the Mississippian. The top of the Mississippian is used because it is an easily identifiable boundary, even though some workers consider the beginning of Alleghanian sedimentation to have commenced during the deposition of the Mississippian Mauch Chunk Formation (e.g., Slingerland and Beaumont, 1989).

The stratigraphic distance from the stratigraphic level of a sample location to the top of the Mississippian was subtracted from the trapping depth determined from fluid inclusion microthermometry, providing an estimate of the thickness of the late Carboniferous–Permian (syntectonic) strata (Table S1 [see ^{footnote 1}]; Plate 1). The calculated load values are shown above the restored cross sections as they are determined from pre-folding veins (Plate 1) and are assumed to have been initially emplaced on the undeformed wedge top. The data are projected from ~15 km on either side of each cross section.

The syntectonic thicknesses determined are variable throughout the region and reflect the timing of fluid trapping and may represent trapping before, during, or possibly after maximum load during syn-folding syntectonic erosion (Evans and Fischer, 2012; Evans et al., 2012) or at pressures less than lithostatic, so not all values represent the maximum load at a particular location. In addition, a single location may have multiple maximum load values (Plate 1). For example, a single sample site may contain mineral veins from multiple sets that formed at different times during the Alleghanian orogeny (Table S1). One vein set may be early and have trapped fluids at low burial conditions, while a later set may have trapped fluids at higher burial conditions. Alternatively, one vein set may record burial conditions prior to folding, while a later vein set or a later mineral stage may record burial conditions during uplift and exhumation that are lower than the earlier vein set. See Evans et al. (2012), Evans and Fischer (2012), and Evans et al. (2014) for examples.

In some cases, the syntectonic load was determined to be zero or slightly less than zero (Plate 1; Table S1). Many of these values are typically at the level of the pre-Alleghenian sedimentary thickness through the Mississippian and are interpreted to indicate fluid trapping prior to or early during syntectonic load accumulation (Plate 1).

Importantly, the calculated syntectonic load was not emplaced instantly across the fold-thrust belt but instead is time transgressive, with erosion toward the hinterland resulting in forelandward deposition. Therefore, the thicknesses illustrated in Plate 1 represent not a static load pile at any particular point in time but a summary of maximum loads over the region during fold-thrust development.

In general, the maximum calculated load values are consistent over a large area, for example, across restored cross section 2 and above thrust sheets 4 through 7 in restored cross section 11 (Plate 1). Similarly, loads are consistent even though samples are from widely separated (as much as 5 km) stratigraphic intervals, for example, samples 15, 25, 40, 42, and 44 on restored cross sections 1–3 (Plate 1). In other areas, only a few sample locations define the maximum load trend, such as above thrust sheets 1 and 2 in restored cross section 3 and above thrust sheets 1 through 3 in restored cross section 13 (Plate 1).

One noticeable trend is that samples in the predominantly carbonate sections of the Siluro-Devonian and Cambro-Ordovician commonly have lower calculated loads than nearby samples from clastic units. This may be due to early fracturing and rapid sealing of veins in the relatively closed fluid systems in the carbonate sections (e.g., Evans and Battles, 1999; Evans et al., 2012). The veins in these rocks may be recording a pre-maximum overburden.

Along with the calculated syntectonic load shown on the restored cross sections on Plate 1, the load values are plotted on a regional map in Figure 2, which depicts the present-day configuration of the fold-thrust belt. Only the highest syntectonic load values are reflected by the contours on the map. Two large areas of high syntectonic load are interpreted to be major Alleghanian depositional centers (Fig. 2) that may have developed as negative-alpha piggyback basins (Fuller et al., 2006; Willett and Schlunegger, 2010) from local foredeeps foreland-ward of the hinterland massifs. The depositional center to the northeast is above the current Anthracite belt synclinorium and forelandward of the Reading Prong massif. It has the largest syntectonic load values of as much as 16 km closest to the hinterland. The southwestern depositional center is generally thickest above the Broadtop synclinorium with 5 to >7 km of load, but the high thickness distribution broadens to the south forelandward of the North Mountain thrust and Blue Ridge massif. The intervening Juniata culmination has the lowest load distribution. In general, the load thickness decreases to less than 1–3 km toward the Appalachian structural front (Fig. 2; Evans et al., 2014). Several locations just forelandward of the structural front, however, show increasing loads of locally as much as 3–5 km in the eastern Appalachian Plateau province.

The cross sections in Plate 1 provide a more detailed view of the load distribution. In the Anthracite belt area, maximum post-Mississippian sedimentary strata thickness is greatest toward the hinterland and is as much as 14–16 km thick in sections 1–3 (Plate 1), decreases to ~10 km in sections 4 and 5 (Plate 1), and is <2 km by section 6 (Plate 1). Toward the foreland in sections 1–3, the maximum load is 6–8 km but decreases to ~5 km in sections 4 and 5. A single location forelandward of the Appalachian structural front in section 5 (Plate 1) has a load of 6 km. In section 3 (Plate 1), there is a reduced load over thrust sheet 3. This thrust sheet develops into the fault-related fold underlying the Berwick anticline, and the low load may reflect fold growth during syntectonic sedimentation.

As the Cambro-Ordovician carbonate duplex develops to the southwest into the Juniata culmination with closely spaced imbricate thrusts, the syntectonic loads decrease markedly such that in sections 5–9 (Plate 1), maximum loads are typically <3 km, only locally exceeding 5 km. The load distribution is not consistent across the sections but instead is characterized by highs separated by lows. It is unknown whether this reflects the actual load distribution or is a manifestation of the widely spaced data. Unlike in sections 1–5, there is no significant buildup of sediment toward the hinterland, which is at the juncture of the two massifs.

Low hinterlandward loads continue southward through sections 10 and 11 (Plate 1). In these two sections, however, there is a buildup of as much as 7–9 km of load in the area that is above the Broadtop synclinorium. The high load is bounded on both sides by antiformal stacks. This large load continues to the south (Plate 1, sections 12–17) and merges with a thick (as much as 9 km) load accumulation that increases toward the hinterland, reflecting the increasing displacement on the North Mountain and Blue Ridge thrust sheets (Evans, 1989, 2010).

In order to understand the relationship of the syntectonic loads determined in this study to the structural development of the fold-thrust belt, it is necessary to spatially restore the regional load distribution to a pre-deformation configuration. Geiser (1988) and Hatcher and Geiser (2010) discussed the difficulties of constructing cross sections in and restoring curved orogens in general, and the Pennsylvania salient specifically, due to the problem of lateral compatibility of fault number, spacing, and displacement. Due to the fact that a series of cross sections constructed normal to the structural grain will all converge (Fig. 1), a three-dimensional approach is necessary to approach a valid regional reconstruction (Hatcher and Geiser, 2010; Watkins et al., 2017).

In this study, the spatial restoration was done on the Cambrian–Ordovician lithotectonic unit that defines the structure of the region. For the Pennsylvania salient area, the initial restoration used 33 regional and 15 smaller line- and area-balanced cross sections. For the southern part of the study area, the restoration by Evans (2010) was used with modification. The 17 cross section lines in Figure 1 were spatially corrected for their convergence by allowing for structures between the cross-section lines to be the same distance apart after the cross sections were restored. The resulting restored section lines are shown in Figure 4. The final restoration lines are not absolute and rely on numerous assumptions; however, the general pattern is necessary. The lines are used for regional balancing and are not meant to reflect transport or displacement trajectories.

The cover rock sequence is restored separately by taking into consideration macroscale folding and faulting as well as finite-strain layer-parallel shortening in the Valley and Ridge and Appalachian Plateau derived from published reports (e.g., Nickelsen, 1966; Engelder and Engelder, 1977; Faill et al., 1989; Ferrill and Dunne, 1989; Hogan and Dunne, 2001; Sak et al., 2012). By incorporating the shortening in the Appalachian Plateau into the retrodeformation, cover rocks currently above the Valley and Ridge are displaced toward the hinterland. The restored positions of the sample locations (Fig. 3) that serve as passive markers in the cover rock sequence were determined by assuming a homogeneous finite strain and by using the Puppet Warp function in Adobe Photoshop. In the process of the restoration, the North Mountain thrust and Blue Ridge massif are pulled back to their early Alleghanian positions (Evans, 1989, 2010) as is the Reading Prong massif (Fig. 4).

The restored load distribution (Fig. 4) shows the two distinct depositional centers. The one forelandward of the Reading Prong massif thins markedly toward the west and northwest. The depositional center forelandward of the North Mountain thrust and Blue Ridge massif to the south shows 3 to >7 km of load over much of the future Valley and Ridge, with a thick tongue extending into the future Broadtop synclinorium. The area extending from the juncture of the two massifs generally has <3 km of load, and this is the area of the future Juniata culmination. Again, this is the summation of loads throughout the orogeny and does not represent a static load pile at any one point in time.

Several studies have used various proxies to estimate the syntectonic load in the central Appalachian orogen. Geodynamic modeling of lithospheric flexure by Beaumont et al. (1987, 1988) predicted 7 km of load during the Early Pennsylvanian in the area that is now eastern Virginia and North Carolina. This 7 km load expanded and moved to central Virginia and North Carolina during the Late Pennsylvanian. These Pennsylvanian loads resulted in 300–600 m of sediment in southern and eastern Pennsylvania. For the Permian, the model by Beaumont et al. (1988) predicted 9–12 km of load in southern Pennsylvania, eastern Maryland, and northern Virginia, resulting in 6.1–7.6 km of sediment in central and southern Pennsylvania and as much as 9.1 km of sediment in eastern Pennsylvania. These estimates are broadly within the range of the results of this study and illustrate that syntectonic loads are highly variable within an orogen and the thickness is transient and dependent upon structural setting.

Conodont alteration index (CAI) values for Devonian rocks are summarized in Repetski et al. (2014) and in Figure 5A. The distribution of CAI values reflects the distribution of syntectonic loads determined in this study. Based on the calibration in Epstein et al. (1977), the CAI values of 4.0–5.0 in the area of the Anthracite belt are equivalent to ~190–300 °C based on ~1 m.y. heating at maximum temperature (Epstein et al., 1977). Maximum measured fluid inclusion T_hA values (Table S1 [see ^{footnote 1}]) fall within this range. Similarly, in the southwestern Valley and Ridge, the CAI values of 3.5–4.5 are equivalent to ~150–230 °C, while in the area of the Juniata culmination, CAI values are 3.0–3.5, indicating temperatures of ~120–160 °C (Epstein et al., 1977). Again, these values are similar to the maximum measured fluid inclusion T_hA values (Table S1).

Coal vitrinite reflectance values for the region (Ruppert et al., 2010) are summarized in Figure 5B. Although the data are only from the Pennsylvanian and Permian coal-bearing rocks, the overall pattern of high thermal maturity is similar to the distribution of large syntectonic loads presented here. In the Anthracite belt, R_o vitrinite reflectance values range from 2.4 in the western part of the belt to >5.5 in the southeastern part (Fig. 5B). Based on the basin%R_o model of Nielsen et al. (2017), these values indicate temperatures of 205–270 °C, which are similar to the maximum T_hA values determined from fluid inclusion microthermometry. Both Levine (1986) and Hulver (1997) suggested deep syntectonic burial of as much as 6–9 km in the Anthracite belt based on vitrinite reflectance, while Paxton (1983) called for 7–9 km based on rock density and porosity.

In contrast, several authors have called on low burial depths and migrating hot fluids to account for the high vitrinite reflectance in the Anthracite belt as opposed to deep burial (Daniels et al., 1990, 1994; Kisch and van den Kerkhof, 1991; Harrison et al., 2004). However, the extremely dense CH₄-CO₂ inclusions reported here (Table S1) support very high burial depths (up to 16 km).

In the Catskill Mountains of southeastern New York, ~125 km along strike from the Anthracite belt, Carboniferous sediment thicknesses are estimated to be as much as 6.4 km based on Upper Devonian vitrinite reflectance (Friedman and Sanders, 1982; Friedman, 1987; Gurney and Friedman, 1987). Lakatos and Miller (1983) and Johnsson (1986) used apatite fission-track analysis to estimate at least 4–7 km of post-Devonian sedimentary rocks in eastern and central New York. In contrast, Gerlach and Cercone (1993) estimated only 0.7–3.4 km of Carboniferous strata in central New York based on Upper Devonian vitrinite reflectance, but they used a much higher geothermal gradient (50 °C km⁻¹) compared to Johnsson (1986) (20 °C km⁻¹) and Lakatos and Miller (1983) (25 °C km⁻¹).

Zhang and Davis (1993) used vitrinite reflectance data to estimate Permian loads to be 4.1 km in the Broadtop synclinorium based on maximum R_o of ~2.0, which their coalification model determined to represent a 155 °C maximum temperature. In the Appalachian Plateau province, their model determined R_o values of ~0.7–1.5 to represent 89–139 °C, which corresponds to a 2.5–3.5 km burial depth. The Zhang and Davis (1993) model used a higher geothermal gradient (~33 °C km⁻¹) than that used here (22.5 °C km⁻¹), resulting in lower estimated burial depths. If the basin%R_o model of Nielsen et al. (2017) is used to estimate temperature, an R_o of ~2.0 indicates a temperature of ~200 °C in the Broadtop synclinorium and 150–180 °C in the Appalachian Plateau, in agreement with the fluid inclusion T_hA values presented here.

Reed (2003) and Reed et al. (2005) used vitrinite reflectance and fluid inclusion microthermometry to estimate a burial depth of ~4.4 km for Upper Pennsylvanian rocks at the Appalachian structural front in West Virginia, which is consistent with samples from both outcrop and core in Evans (2010).

Although they did not calculate overburden directly, Orkan and Voight (1985) used fluid inclusion microthermometry of methane-rich and water-rich inclusion pairs from mineral veins at 11 sites in the Pennsylvania Valley and Ridge to estimate trapping pressures of 25–160 MPa, which are generally significantly lower than in this study, although it is not clear whether the lowest T_hH values were used. However, their aqueous inclusion T_hA data are comparable to those of this study.

Throughout most of the study area, the restored syntectonic load values in the central Appalachians are comparable to those in active fold-thrust belts. For example, in a restored model of the Alps, Burkhard and Sommaruga (1998) used vitrinite reflectance data to infer that prior to erosion, 3–5 km of Tertiary syntectonic sediment filled the Swiss Molasse Basin during Alpine deformation in front of the Penninic thrust and the Subalpine Molasse thrust. This number is similar to that shown by Schegg et al. (1997) of 4.3 km and Mraz et al. (2019) of ~4–5 km. Moss (1992) used vitrinite reflectance to infer that as much as 5 km of load had been eroded from thrust sheets within the French Subalpine Chains. Similar (4–5 km) and possibly higher (as much as 7 km) values of syntectonic overburden within foredeep deposits of the Macigno Costiero Formation have recently been estimated for the neighboring northern Apennines fold-thrust belt of peninsular Italy, which shares features and a deformation history analogous to those of the French Subalpine Chains and Swiss Jura Mountains (Tavarnelli et al., 2021).

In the Sub-Andean belt, Echavarria et al. (2003) showed 6–7 km of Tertiary molasse sediments deposited in northwestern Argentina. In the El Simbolar syncline in northern Argentina, at least 5 km sediment is found in more forelandward synclines (Echavarria et al., 2003). Baby et al. (1995) document 6.5–7.0 km of synorogenic basin fill in the Alto Beni syncline, a piggyback basin in the Sub-Andean zone of northern and central Bolivia. Louterbach et al. (2018) also showed 5–7 km of syntectonic sediments in the Beni Basin of the northern Bolivia sub-Andean thrust belt.

Webb (2013) and DeCelles et al. (1998) showed 4–6 km of syntectonic sediment in the Siwalik Formation in the Himalayan foreland, ahead of the Sub-Himalayan thrust zone, while Le Garzic et al. (2019) showed ~3.5 km of syntectonic sediments in a large syncline in the northwestern Zagros fold-thrust belt in the Kurdistan Region of Iraq, 75 km forelandward of the last major emergent thrust fault (High Zagros fault) in an area of a blind duplex. In the Ebro Basin of Spain, Jones et al. (2003) documented >1 km of syntectonic sediment just forelandward of the Cavalls-Pándols thrust sheet complex, an antiformal stack in a similar structural position as the Nittany anticlinorium and Wills Mountain anticlinorium in the central Appalachians. Piggyback basins are common features in fold-thrust belts (e.g., Ori and Friend, 1984; Lawton and Trexler, 1991; Coogan, 1992; Hippolyte et al., 1994; Martín-Martín and Martín-Algarra, 2002; Thompson et al., 2015) and may play a role in the final structural architecture of the belt (e.g., Leturmy et al., 2000).

The large 10–16 km syntectonic loads in the Anthracite belt due to the emplacement of the Reading Prong massif have few analogues in the literature. For example, Heermance et al. (2008) showed at least 10 km of syntectonic load in the Kashi foreland of northwestern China.

Sequential deformational models for each of the summary cross sections (Fig. 2) show the Pennsylvanian–Permian syntectonic load thicknesses projected onto each section from ~20 km on either side of the section, and the maximum restored load is shown in stage 1 of each sequence (Plate 2). A sequential hinterland-to-foreland thrust progression is assumed. There are few data on the regional lithology of the basal Waynesboro detachment and therefore on the actual strength of the detachment. The cross sections assume that the syntectonic load was distributed across a depositional wedge top (e.g., Ford, 2004; Kendall et al., 2020) above the deforming wedge. The wedge top is the top of the Upper Ordovician through Mississippian cover rock sequence, which is underlain by the Cambro-Ordovician lithotectonic unit that controls the structural deformation.

The dip of the upper surface of the thrust wedge (α) and the dip of the base of the wedge (β) define the taper angle (α + β). The initial taper angle of each model is defined by the thickness of the restored syntectonic load that sits on the wedge top and the dip of the basal detachment in the Cambrian Waynesboro Formation. The dip of the basal detachment in the central Appalachians is likely not the same as it was during the late Paleozoic due to isostatic adjustments and may have been higher than today, however this would not affect the modeled taper angle. The deformation sequences are created from the regional restored balanced cross sections and then drawn as a forward model in stages. The wedge taper and syn-deformational load in each stage are derived iteratively to match the observed final maximum load distribution. Because actual α and β values are not known, the syntectonic load is simply placed on a wedge top and there is no compensation for isostatic load adjustment.

In each model, the rear of the wedge is dominated by an antiformal stack of metamorphic and/or igneous basement rocks and is likely strong and riding above a weak mylonitic décollement; therefore, it may have had a higher taper angle. The thrust stack includes thrust sheets that have been transported along reactivated Late Ordovician Taconic-age faults (e.g., Valentino et al., 1994; Krol et al., 1999; Kunk and Burton, 1999; Valentino et al., 2004; Kunk et al., 2005; Southworth et al., 2006; Wintsch et al., 2010). These thrust sheets may have provided the elevation in the rear of the wedge to allow for the observed syntectonic sediment thicknesses. The last stage shown in each model is the present-day cross section.

Cross section A-A′ from the eastern salient (Fig. 2; Plate 2A) shows the imbricate stack of the Reading Prong–Lebanon Valley nappe system that developed early during the Alleghanian orogeny (Faill, 1998). Above this structure is the Martic thrust sheet, which has been shown to have had activity into the Pennsylvanian (Valentino et al., 2004). With forelandward progression of the massif, syntectonic erosion develops a thick sediment wedge as much as 16 km thick over the future Anthracite belt (Plate 2A, stages 1 and 2) and 5–7 km toward the foreland. To achieve this thickness distribution, a taper angle of ~6.5° is required to fit the restored load thickness (Fig. 4A).

With continued thrusting and the growth of the Reading Prong massif, early-deposited syntectonic sediments are eroded and recycled into the proximal foredeep (Plate 2A). Due to the stabilizing effect of the sediment load, thrusting jumps ahead to the Berwick anticline structure, which renews the critical wedge state. The Berwick anticline may have been growing during the development of the sediment wedge, with concurrent erosion of prior-deposited syntectonic sediment, given that the observed load is thinner over the structure (Plate 2A, stage 3). The final structural geometry of the Cambro-Ordovician lithotectonic unit is dominated by long thrust sheets (average length ~22 km) with low-amplitude fold-thrust structures (Plate 2A, stages 4 and 5).

In a study of detrital zircon ages from the Tumbling Run Member of the Pottsville Formation at the base of the syntectonic pile, Becker et al. (2005) concluded that the source of sediment in the Anthracite belt was the metamorphic rocks in the Piedmont (Fig. 1), with some Laurentian contribution. The source may have been the reactivated Taconic thrust sheets. Similarly, Gray and Zeitler (1997) suggested that the source of zircon-bearing clasts in the stratigraphically higher Sharp Mountain Member of the Pottsville Formation was primarily the Piedmont with a secondary Grenville-age source. A southern and possible longitudinal source of sediment of the Pottsville Formation is also supported by detrital muscovite ages (Monami et al., 2022).

In order to accommodate the crustal loading associated with deep basin burial, Levine (1986) called on crustal downwarp associated with the Scranton gravity high. Multiple workers have suggested that at least part of the overburden is tectonically emplaced as a thrust sheet (Orkan and Voight, 1985; Levine, 1986; Faill, 1998). However, there is no structural evidence for a thrust sheet to have been emplaced this far toward the foreland (Plate 2A, stage 1). The thick syntectonic sequence above the Anthracite belt likely reduced the strength of the basal detachment due to the high burial temperatures and may have contributed to lowering the wedge taper.

Cross section B-B′ across the Juniata culmination (Fig. 2; Plate 2B) shows the Cambro-Ordovician carbonate lithotectonic unit deformed into a series of thrust imbricates with an average thrust spacing of ~15 km. The rear of the wedge is dominated by an antiformal stack of thrust sheets overlying the South Mountain–Blue Ridge thrust sheet. The restored syntectonic loads are low, generally <2 km, but locally >3 km, with intermittent small peaks. These peaks of sediment load presumably developed at points of sediment buildup ahead of the incremental development of the duplex structure associated with concurrent erosion of syntectonic sediments above the newly formed thrust sheets (Plate 2B, stages 2–5). The thickness of the peaks requires a taper angle of ~5.5°. The imbricate nature of this part of the fold-thrust belt did not allow the development of any major synclinoria as sediment traps. The cross section extends from the terminus of both the Blue Ridge and Reading Prong massifs where there was no significant foredeep developed and also less topographic relief to supply sediment to the wedge top.

Cross sections C-C′ and D-D′ are located the southern limb of the salient (Fig. 2; Plates 2C and 2D). Fold-thrust structures of the Cambrian–Ordovician lithotectonic unit have an average thrust spacing of ~17 km in section C-C′ and ~22 km in section D-D′. In both sections, the maximum load is above the area that develops into the Broadtop synclinorium, while in section D-D′, there is an additional thick load that is forelandward of the far-traveled North Mountain thrust sheet (Evans, 1989, 2010). The thickness of the peaks requires a taper angle of ~6.0°.

Some of the syntectonic load in these sections may be due to the emplacement of the North Mountain thrust sheet, which places Cambro-Ordovician carbonates above a similar section of carbonates for as much as 60 km (Plates 2D–2F, stage 1), displacing the entire cover rock sequence. It is likely that part of this cover rock sequence was transported across the North Mountain thrust and eroded into the foreland (Evans, 2010). However, the earliest syntectonic sediments of the Upper Mississippian Mauch Chunk Formation and Pennsylvanian Pottsville Formation in the Broadtop synclinorium have detrital zircons that indicate a Piedmont source (Becker et al., 2005; Park et al., 2010; Thomas et al., 2017), again possibly from reactivated Taconic thrust sheets.

With the growth of each fold structure, sedimentation across the wedge top increases and the taper angle is approached that is necessary to create new thrusts and to reactivate preexisting ones to drive the wedge forward. Rapid denudation would have kept topography subdued (e.g., De Paor and Anastasio, 1987; Burbank and Beck, 1991), therefore it is possible that little topography was developed across growing anticlinal structures. However, the Wills Mountain anticlinorium (Plates 2C and 2D) is interpreted to have been positive topographic features that “dammed” syntectonic sediment dispersal, creating a piggyback basin in the Broadtop synclinorium that allowed 5 to >7 km of syntectonic sediment to accumulate, stabilizing this part of the wedge (Plates 2C and 2D, stages 4 and 5). This would have reduced the wedge taper, possibly leading to intermittent out-of-sequence thrusting in the Blue Ridge massif would have increased thrust wedge taper angle in the hinterland, driving forward the next thrusts. Loads forelandward of the Wills Mountain anticlinorium are likely related to the growth and erosion of these structures (Evans, 2010).

Cross sections E-E′ and F-F′ are located in the linear segment of the Valley and Ridge that extends through eastern West Virginia and western Virginia (Fig. 2; Plates 2E and 2F). The structure is defined by thrust sheets in the Cambro-Ordovician lithotectonic units with leading-edge folds in eastern section E-E′ and by more closely spaced imbricates throughout section F-F′. Both sections are characterized by the extensive flat-on-flat of the Wills Mountain duplex in the western Valley and Ridge. As in section C-C′, growth of the Wills Mountain duplex resulted in the formation of a smaller piggyback basin in the Broadtop synclinorium. The syntectonic loads are thicker toward the hinterland in section E-E′ than in F-F′, although that may be sampling bias. The thickness of the peaks requires a taper angle of ~6.0°.

Again, much of the syntectonic load in these two sections may be due to the emplacement of the North Mountain thrust sheet, which displaced the entire cover rock sequence. Both sections also show a decrease in load toward the foreland, with a very low load over the Wills Mountain anticline, which may indicate that it was a growing fold with synkinematic erosion of prior-deposited syntectonic sediments and part of the cover rock sequence.

Apatite fission-track (AFT) ages and apatite (U-Th)/He (AHe) ages for the study area are summarized in Figures 6 and 7. In the central Valley and Ridge province of Pennsylvania, Blackmer et al. (1994) interpreted the AFT ages in the Juniata culmination (125–246 Ma) as being synorogenic and coeval with unroofing with growth of the underlying duplex. These dates are slightly younger than the age of folding (270–250 Ma) determined by paleomagnetic data (Stamatakos et al., 1996). The rapid unroofing of the culmination is consistent with the low syntectonic loads over the area as reported in this study. Similarly, several older AFT ages (137–246 Ma) are found along the southern Nittany anticline and along large folds in the eastern Appalachian Plateau province, possibly indicating early unroofing with fold growth. This is also an area of low syntectonic load. In contrast, the younger AFT ages in the Anthracite belt (111–152 Ma) are consistent with long-term unroofing of a thick rock sequence. In the Broadtop synclinorium, where as much as 7–9 km of syntectonic load is interpreted, sparse AFT ages (125–189 Ma) also reflect long-term unroofing of a thick rock sequence. A single AFT age of 95 Ma (Roden, 1991) is found in Upper Devonian rocks in the footwall of the North Mountain thrust where as much as 9 km of syntectonic load is interpreted (Fig. 4). AFT ages in the Piedmont (Gates and Glover, 1989; Kunk et al., 2005) are generally young (198–124 Ma) and indicate cooling through the AFT closure temperature (100 ± 10 °C; Ketcham et al., 1999) from the Early Jurassic through the Early Cretaceous. In the Reading Prong massif, AFT ages (144–168) are older than those in the Anthracite belt, suggesting earlier uplift and exhumation. Pooled AHe ages (McKeon et al., 2014) from the Reading Prong massif (Figs. 6 and 7) vary from 72 to 155 Ma and are younger than AFT ages in the area, reflecting the lower closure temperature (~60 °C; Flowers, 2009). Similarly, AHe ages from Pennsylvanian rocks in the eastern Appalachian Plateau province of West Virginia (Reed et al., 2005) are also generally younger than the nearby AFT ages.

Basler et al. (2021) examined zircon (U-Th)/He (ZHe) from multiples sites across the Valley and Ridge province of West Virginia and into the western Piedmont province (Figs. 6 and 7F). They concluded that the Valley and Ridge province underwent rapid Alleghanian exhumation during the middle Permian. In particular, their inverse thermal model suggests that the Wills Mountain anticline and related duplex structure (WMA on Fig. 7F and Plate 2F) were formed by 280–270 Ma and being exhumed by 270 Ma. This rapid duplex growth and exhumation may be reflected in the lower restored syntectonic loads above the Wills Mountain duplex (Plate 1, sections 14−17). Relatedly, Basler et al. (2021) suggested that the Blue Ridge massif was rapidly cooling by the late Permian (Figs. 6 and 7F). The closure temperature for ZHe is ~200 °C (Reiners et al., 2002), although it may be lower (Gérard et al., 2022).

Zircon fission-track (ZFT) ages in the central Appalachians (Gates and Glover, 1989; Kohn et al., 1993; Kunk et al., 2005; Naeser et al., 2016) are shown in Figures 6 and 7. ZFT ages from the Appalachian Plateau and Valley and Ridge provinces are all older than depositional age, indicating that they have not experienced the temperatures having been reset (Naeser et al., 2016). Although there are some older ages present in the western Blue Ridge massif, most ZFT ages in the core of the Blue Ridge massif are Alleghanian (ca. 340–283 Ma), reflecting active thrusting and uplift through the ZFT closure temperature (~235 ± 25 °C; Brandon et al., 1998). Muscovite cooling ages in the Blue Ridge massif are also Late Mississippian to Early Pennsylvanian (Kunk and Burton, 1999; Kunk et al., 2005; Southworth et al., 2006; Bailey et al., 2017). Younger ages (Figs. 6 and 7) in the western Piedmont (298–241 Ma) and eastern Piedmont (176–142 Ma) may be due to the erosion of thick crystalline thrust sheets. In the southwestern Reading Prong massif in the southeastern Piedmont province of Pennsylvania, ZFT ages are interpreted to have been reset by a Mesozoic thermal event (Kohn et al., 1993) so the ZFT ages do not record a thermal history between the end of the Alleghanian orogeny (ca. 260 Ma) and Mesozoic rifting (ca. 230 Ma).

In general, the thermochronologic data are consistent with the syntectonic overburdens determined in this study. In areas of low restored load, exhumation through the AFT closure temperature was during or shortly after the Alleghanian orogeny, while in areas of thick restored load, AFT ages are Mesozoic. Both AFT and ZHe ages suggesting exhumation during fold growth are consistent with the thinner restored loads above major anticlinoria.

The spatial distribution of syntectonic load determined for the central Appalachians shows a strong relationship to the structural geometry of the lithotectonic unit that defines the architecture of the fold-thrust belt. In areas of thick loads (5 to >10 km), the Cambro-Ordovician carbonates are deformed into broad thrust sheets with low-amplitude thrust-related folds, while in areas of low load (<3–5 km), the structural style is dominated by closely spaced imbricate thrust sheets. These patterns reflect both analog and mechanical model scenarios that show where syntectonic sedimentation is high, thrust length increases and hence thrust spacing increases, while where syntectonic sedimentation is low, thrust length and spacing are short (Marshak and Wilkerson, 1992; Storti and McClay, 1995; Simpson, 2006; Stockmal et al., 2007; Bonnet et al., 2007, 2008; Wu and McClay, 2011; Buiter, 2012; Graveleau et al., 2012; Fillon et al., 2013a, 2013b; Erdõs et al., 2015; Sun et al., 2019). This suggests that the observed syntectonic load, although not the only controlling factor, was a primary influence on the structural development.

According to Ettensohn (2008), the central Appalachian orogen was probably a high-standing Himalayan-type mountain range by the end of the Alleghanian orogeny. Rodgers (1987) suggested that a broad altiplano similar to the modern central Andes may have extended across much of the southern and central Appalachians that would have been a 300–400-km-wide range with an average relief of ~4 km (Beaumont et al., 1987; Slingerland and Furlong, 1989; Faill, 1998).

Although syntectonic erosion of a growing thrust wedge can play a role in controlling the evolution of the wedge (e.g., Persson and Sokoutis, 2002; Bonnet et al., 2007, 2008; Whipple, 2009; Konstantinovskaya and Malavieille, 2011; Dal Zilio et al., 2020), few data exist on the extent of erosion of the growing central Appalachian fold-thrust belt. However, based on plate reconstructions (e.g., Scotese, 2001), during the late Carboniferous and Permian, the central Appalachian fold-thrust belt was located in the tropical, equatorial belt, which is characterized by humid to perhumid climate (Cecil et al., 2003, 2004; Ettensohn, 2008). Therefore, the region would likely have had a high-precipitation climate. This could have led to relatively rapid syntectonic erosion rates. The apparent low loads over certain anticlines and anticlinoria suggest the synkinematic erosion of prior-deposited syntectonic sediments and part of the cover rock sequence exposed over growing folds.

Several studies have evaluated the co-effects of syntectonic erosion and sedimentation on thrust wedge development (e.g., Mugnier et al., 1997; Bonnet et al., 2008; Malavieille, 2010; Wu and McClay, 2011). Based on the analysis by Wu and McClay (2011), fold-thrust belts exposed to high sedimentation and low moderate erosion would have reactivated to highly active hinterland thrusts. Although there are few data on thrust reactivation in the Blue Ridge massif, there is evidence for Permian motion in the eastern Reading Prong massif (Ratcliffe, 1980) and for Pennsylvanian and Permian motion on transpressional faults in the Piedmont (e.g., Krol et al., 1999; Kunk and Burton, 1999; Kunk et al., 2005; Southworth et al., 2006; Wintsch et al., 2010) that may have served to elevate the back of the growing wedge.

By using fluid inclusion microthermometry of synkinematic veins, a syntectonic load can be estimated in highly eroded orogenic belts. In the central Appalachians, restored loads indicate two major depocenters as piggyback basins during the Alleghanian orogeny, one above the Anthracite belt with as much as 16 km of syntectonic load probably sourced by the erosion of the growing Reading Prong massif, and the other over the Broadtop synclinorium with as much as 7 km of load likely sourced by the erosion of the growing Blue Ridge massif and emplacement of the North Mountain thrust sheet. In the intervening Juniata culmination, the loads were generally <3 km. The structural architecture of the fold-thrust belt reflects this load distribution with widely spaced thrusts in areas of high load and closely spaced thrusts in areas of low load. The maximum syntectonic burial depths presented here will aid in future work to constrain the thermochronologic history of the central Appalachians.

The author recognizes the time and effort by undergraduate students at Central Connecticut State University who contributed to this study: S. Braddock, A. Delisle, K. Duffy, C. Lafonte, K. Landry, J. Leo, E. Lincoln, C. MacDonald, J. Marino, R. McMahon, I. Murphy, D. Pietkevich, V. Swenton, and N. Zygmont. In addition, the author acknowledges the inspiration for deciphering the structure of the Pennsylvania salient from Don Wise and the late Nick Nickelsen and Rodger Faill. The author also thanks D. Ferrill, E. Tavarnelli, and F. Pazzaglia for constructive comments that greatly improved earlier versions of this paper. Partial funding for this project was provided by American Chemical Society Petroleum Research Grant 51793-UR8.