The Mid-Atlantic shore to the Appalachian highlands: field trip guidebook for the 2010 joint meeting of the Northeastern and Southeastern GSA Sections (GSA Field Guide 16)

The Mid-Atlantic Shore to the Appalachian Highlands: Field Trip Guidebook for the 2010 Joint Meeting of the Northeastern and Southeastern GSA Sections

edited by Gary M. Fleeger Pennsylvania Geological Survey 3240 Schoolhouse Road Middletown, Pennsylvania 17057-3534 USA and Steven J. Whitmeyer Department of Geology and Environmental Science James Madison University 800 S. Main Street, MSC 6903 Harrisonburg, Virginia 22807 USA

Field Guide 16 3300 Penrose Place, P.O. Box 9140

Boulder, Colorado 80301-9140 USA

2010

Copyright © 2010, The Geological Society of America (GSA), Inc. All rights reserved. GSA grants permission to individual scientists to make unlimited photocopies of one or more items from this volume for noncommercial purposes advancing science or education, including classroom use. For permission to make photocopies of any item in this volume for other noncommercial, nonprofit purposes, contact The Geological Society of America. Written permission is required from GSA for all other forms of capture or reproduction of any item in the volume including, but not limited to, all types of electronic or digital scanning or other digital or manual transformation of articles or any portion thereof, such as abstracts, into computer-readable and/or transmittable form for personal or corporate use, either noncommercial or commercial, for-profit or otherwise. Send permission requests to GSA Copyright Permissions, 3300 Penrose Place, P.O. Box 9140, Boulder, Colorado 80301-9140, USA. GSA provides this and other forums for the presentation of diverse opinions and positions by scientists worldwide, regardless of their race, citizenship, gender, religion, or political viewpoint. Opinions presented in this publication do not reflect official positions of the Society. Copyright is not claimed on any material prepared wholly by government employees within the scope of their employment. Published by The Geological Society of America, Inc. 3300 Penrose Place, P.O. Box 9140, Boulder, Colorado 80301-9140, USA www.geosociety.org Printed in U.S.A. Library of Congress Cataloging-in-Publication Data The Mid-Atlantic shore to the Appalachian highlands : field trip guidebook for the 2010 joint meeting of the Northeastern and Southeastern GSA Sections / edited by Gary M. Fleeger and Steven J. Whitmeyer. p. cm. -- (Field guide ; 16) Includes bibliographical references. ISBN 978-0-8137-0016-8 (pbk.) 1. Geology--Piedmont (U.S. : Region)--Fieldwork. 2. Geology--Blue Ridge Mountains Region-Fieldwork. 3. Geology--Middle Atlantic States--Fieldwork. 4. Geology--Appalachian Mountains-Fieldwork. I. Fleeger, Gary M. (Gary Mark) II. Whitmeyer, Steven J. QE78.3.M53 2010 557.5--dc22 2010000940 Cover, front: View to the northeast of Germany Valley, West Virginia, from an overlook on Route 33. The exposed ledges on the left and right margins and the distant center of the picture are outcrops of Silurian Tuscarora sandstone; the floor of the valley is composed of Ordovician carbonate rocks. Thus, Germany Valley is a several kilometer-scale breached anticline that plunges shallowly to the northeast. Photo by Steve Whitmeyer. Back: View from Bearfence Mountain in Shenandoah National Park, looking north toward the ridges of Massanutten Mountain, with outcrop of Catoctin greenstone in the foreground. Photo by Chuck Bailey.

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Contents

Preface . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . v 1. The Peach Bottom area in the Pennsylvania-Maryland Piedmont . . . . . . . . . . . . . . . . . . . . . . . . . 1 R.T. Faill and R.C. Smith II 2. Soils, geomorphology, landscape evolution, and land use in the Virginia Piedmont and Blue Ridge . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 31 W.C. Sherwood, A.S. Hartshorn, and L.S. Eaton 3. Magmatic layering and intrusive plumbing in the Jurassic Morgantown Sheet, Central Atlantic Magmatic Province . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 51 L. Srogi, T. Lutz, L.D. Dickson, M. Pollock, K. Gimson, and N. Lynde 4. The early through late Pleistocene record in the Susquehanna River Basin . . . . . . . . . . . . . . . . 69 D.D. Braun 5. Stratigraphy, structure, and tectonics: An east-to-west transect of the Blue Ridge and Valley and Ridge provinces of northern Virginia and West Virginia . . . . . . . . . . . . . . . . . . . . . . . . . . . . 103 L.S. Fichter, S.J. Whitmeyer, C.M. Bailey, and W. Burton 6. Teachers guide to geologic trails in Delaware Water Gap National Recreation Area, Pennsylvania–New Jersey J.B. Epstein This guide is available at http://fieldguides.gsapubs.org/ (open access) or as GSA Data Repository item 2010097 posted at www.geosociety.org/pubs/ft2010.htm or on request from editing@ geosociety.org, Documents Secretary, GSA, P.O. Box 9140, Boulder, CO 80301, USA.

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Preface This guidebook features field trips offered during the joint meeting of the Northeastern and Southeastern Sections of the Geological Society of America (GSA) held in Baltimore, Maryland, in March 2010. Chapters in this guide reflect the meeting’s theme (“Linking North and South: Exploring the Connections between Continent and Sea,”) in that they span the lowlands of eastern Pennsylvania to the highlands of northeastern West Virginia (Fig. 1). Four physiographic provinces are covered: Piedmont (Piedmont Upland and Gettysburg-Newark Lowland Sections), Blue Ridge, Valley and Ridge, and Appalachian Plateau. The geologic foci are likewise variable, ranging from Precambrian basement rocks to Pleistocene sediments. The chapters are organized alphabetically, with premeeting trips listed first (Faill and Smith, Sherwood et al., Srogi et al.), followed by postmeeting trips (Braun, Fichter et al.) Topics range from surficial materials and landscape evolution (Sherwood et al., Braun) to magmatism and igneous processes (Srogi et al.) to stratigraphy, structure and tectonics (Faill and Smith, Fichter et al.) In addition, at least two of the field guides are specifically targeted at teachers and instructional pedagogy of field-oriented education (Fichter et al., Epstein). Field Trip 6, by Epstein, is based on a field guide originally published in conjunction with the 2006 GSA Annual Meeting in Philadelphia. As a result, the field guide for this trip is not being published in this volume. However, an updated guide, which includes additional material not included in the 2006 version, is available (open access) at http://fieldguides.gsapubs.org/, or as GSA Data Repository item 2010097 at www.geosociety.org/pubs/ft2010.htm. The editors would like to thank all of the authors, field trip organizers, and leaders for the countless hours that went into producing this volume. Special thanks to the reviewers who helped improve this volume: Alan Benimoff, Duane Braun, Lee Daniels, Rick Diecchio, David Eggler, John Haynes, Louis Heidel, Dan Richter, Bill Sevon, Steve Shank, Scott Southworth, Aaron Thompson, and Gil Wiswall. Many thanks to the GSA Publications Department staff for help in preparing, formatting, and generally shepherding this guide through the production process. We hope that geoscientists and educators alike will find the contents of this field guide thought provoking in discussions of regional geology and pedagogy, as well as useful for planning future field excursions. Gary M. Fleeger and Steven J. Whitmeyer

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Field Trip 5

Field Trip 2

Baltimore

Field Trip 4

Field Trip 1

Field Trip 3

Field Trip 6

Figure 1. Google Maps relief map of the Mid-Atlantic region showing the approximate locations for the field trips covered by the chapters in this volume. Field Trip 1—Faill and Smith; Field Trip 2—Sherwood et al.; Field Trip 3—Srogi et al., Field Trip 4—Braun; Field Trip 5—Fichter et al.; Field Trip 6—Epstein.

The Geological Society of America Field Guide 16 2010

The Peach Bottom area in the Pennsylvania-Maryland Piedmont Rodger T. Faill* 3407 Rutherford Street, Harrisburg, Pennsylvania 17111-1850, USA Robert C. Smith II* 22 Longview Drive, Mechanicsburg, Pennsylvania 17050, USA

ABSTRACT The Appalachian Piedmont in south-central Pennsylvania and north-central Maryland contains metasedimentary siliciclastic rocks (phyllites to quartzites) that were deposited largely offshore of Laurentia, prior to and during the early history of the Iapetan Ocean. The Peach Bottom area is centered on the belt of Peach Bottom Slate and overlying Cardiff Quartzite, which is surrounded by the late Neoproterozoic and early Paleozoic rocks of the Peters Creek and Scott Creek (new name) Formations. Their provenance was the Brandywine and Baltimore microcontinents that lay farther offshore of the Laurentian coast. This area also includes an ophiolitic mélange that formed in front of an advancing island arc in Iapetus. All these rocks lay largely undisturbed throughout much of the Paleozoic, experiencing only chloritegrade greenschist facies metamorphism through deep burial. Alleghanian thrusting associated with the growth of the Tucquan anticline imparted their present widespread, monocline, steep southeast dip of the bed-parallel foliation.

The rocks within the Peach Bottom area have long been considered part of a regional syncline (Knopf and Jonas, 1923, 1929), with the Peach Bottom Slate as the youngest stratigraphic unit. Traditionally, this core was presumed to be underlain by the Cardiff Quartzite and Peters Creek Formation, which lay adjacent on both the northwest and the southeast limbs. None of the rocks in the Peach Bottom area support this traditional interpretation.

INTRODUCTION The Peach Bottom study covers an ~40 km × 13 km quadrilateral-shaped area located in Lancaster and York Counties, Pennsylvania and Harford and Cecil Counties, Maryland (Fig. 1). The south-southeast flowing Susquehanna River crosses midway along its length. The boroughs of Delta, Pennsylvania, and Cardiff, Maryland, lay some 10 km southwest of the river. The distinctive Peach Bottom Slate underlies a narrow outcrop belt (0.25–2 km wide) for 30 km along the middle of the Peach Bottom area. The Cardiff Quartzite, which conformably overlies the slate, surrounds the slate along much of its perimeter. Together, they constitute the Peach Bottom Slate Belt.

GEOLOGIC SETTING The dominant structure in the south-central Pennsylvania Piedmont siliciclastics is the Tucquan anticline (Frazer, 1880), a very large, west-southwest–trending, gently plunging, upright

*[email protected]; [email protected] Faill, R.T., and Smith, R.C., II, 2010, The Peach Bottom area in the Pennsylvania-Maryland Piedmont, in Fleeger, G.M., and Whitmeyer, S.J., eds., The MidAtlantic Shore to the Appalachian Highlands: Field Trip Guidebook for the 2010 Joint Meeting of the Northeastern and Southeastern GSA Sections: Geological Society of America Field Guide 16, p. 1–30, doi: 10.1130/2010.0016(01). For permission to copy, contact [email protected]. ©2010 The Geological Society of America. All rights reserved.

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fold (Fig. 2). Its east end lies at or near the southwest terminus of Mine Ridge; its west end in the vicinity of the MarylandPennsylvania border is not well defined; it either dies out or is truncated. The dominant foliation that pervades the metasedimentary schists parallels bedding and best displays Tucquan anticline form: subhorizontal along the hinge, northwest dipping north of the hinge, and southeast dipping to the south (Freedman et al., 1964). The schists of the Tucquan anticline are bounded on the northwest by the Martic Line, a fairly sharp transition between the Tucquan schists and the slope carbonates to the southeast, and schists to the northwest that accumulated on the Laurentian continental margin. The nature of the Martic Line lies at the center of a long-standing, contentious issue (Miller, 1935; Cloos and Hietanen, 1941; Rodgers, 1970; Higgins, 1972; Wyckoff, 1990; and Wise and Ganis, 2009), as to whether it is a conformal stratigraphic contact or a thrust fault. The nature of the Martic Line bears on the Peach Bottom area. If the Line is a conformal contact, the age of the Peach Bottom area rocks may be younger than

middle Ordovician (barring some intervening structure(s)). If a thrust, the Peach Bottom area rocks could be much older. The Laurentian carbonate shelf consists of a thick sequence of largely carbonate rocks ranging in age from lower Cambrian to middle Ordovician. This sequence overlies a thinner sequence of siliciclastic rocks (quartzites to mudstones) that in turn overlies the metabasalts and metarhyolites of the Catoctin Formation. The Catoctin Formation itself overlies an even thinner sequence of the Neoproterozoic Chilhowee siliciclastics to the west. The Catoctin Formation is Neoproterozoic to lower Cambrian in age (564 ± 9 Ma in the north-central Appalachians, Aleinikoff et al., 1995; Tollo et al., 2004). The Baltimore Mafic Complex and equivalent rocks to the southwest in Maryland bound the southeast side of the Tucquan schists (Fig. 2). The Baltimore Mafic Complex consists in large part of peridotites and pyroxenites that intruded the lower reaches of an island arc (herein called Cecil Island Arc, following Faill, 1997) in Iapetus at 489 ± 7 Ma (Sinha et al., 1997). The peridotites and pyroxenites have since been altered largely to serpentinites,

Figure 13

Figure 9

Wakefield Conowingo Dam

Holtwood Delta

McGuigan Quarries Bald Friar Hill

Figure 10

Figure 1. Geologic map of the Peach Bottom area in south-central Pennsylvania and north-central Maryland, within portions of the Conowingo Dam, Delta, Holtwood, and Wakefield 7½-minute quadrangles. The Delta Duplex is the fault-bounded zone that contains the Peach Bottom Slate, Cardiff Quartzite, and Sykesville (north) Formation.

Peach Bottom area, Pennsylvania-Maryland Piedmont talcs, and carbonates (mostly magnesite). During the advance of the Baltimore Mafic Complex, sediments, basalts, and ultramafic fragments accumulated as a precursory mélange, the Sykesville Formation (Muller et al., 1989) in front of Baltimore Mafic Complex. The sharp contact of the Sykesville schists, metabasalts, and steatized ultramafites with the Tucquan schists, the striking lithic contrast, and the very different provenances support the probability of a thrust fault contact between the two at the Martic Line. The Tucquan schists have been metamorphosed to greenschist facies. Almandine garnets are common in the hinge of the fold. The rocks on either side are at biotite grade, and authigenic albite is common. Farther away from the hinge, chlorite is the dominant mineral in the pelitic fractions. This pattern of higher-grade metamorphism in the presumably older rocks in the hinge implies a greater depth of burial for those rocks, and suggests that the metamorphism occurred prior to the development of the anticline. Structurally, the Tucquan anticline appears to be simple (Fig. 2). Throughout much of the anticline, the dominant foliation tends to parallel the compositional layering. This suggests that the foliation (defined by the phyllosilicate minerals) developed concurrently with the burial metamorphism, while the beds were horizontal. The growth of the Tucquan anticline thus postdated the burial metamorphism. A steeply dipping to subvertical, east-northeast–trending foliation transects much of the Tucquan anticline (Freedman et al., 1964). This foliation postdates the dominant foliation but is not as penetrative. Its age, character, and attitude suggest it may be axial-planar, and thus coeval, to the upright Tucquan anticline. Smaller folds (from decimeter to meters in size) are present

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locally, with subvertical axial surfaces and steeply southeastplunging axes. Local faults of various orientations are also present, with displacements of a few meters at most. Crenulations are common, many of them plunging moderately to the northeast, but gently southwest-plunging crenulations are present as well. THE OCTORARO BASIN The Rodinian continent was amalgamated by the Grenvillian orogeny (from ca. 1200 to 1000 Ma) at the end of the Mesoproterozoic. What was to become the Peach Bottom area lay well within Rodinia during the subsequent Neoproterozoic, underlain by rocks of the Grenville orogen. It appears that significant rifting activity occurred within Rodinia during the middle of the Neoproterozoic, from 760 to 700 Ma (Tollo and Aleinikoff, 1996; Tollo et al., 2004), as evidenced by the development of the Robertson River Igneous Suite in Virginia (Tollo and Aleinikoff, 1996) and possibly the Ocoee basin. Whether true oceanic crust was formed in the rifts, or continental crust was simply attenuated, one or more basins seemed to have developed, which accumulated siliciclastic sediment derived largely from the surrounding Rodinian highlands. The pre-Catoctin Octoraro Formation, if it is indeed Neoproterozoic in age, may have accumulated in one of the rift basins developed at that time. Additionally, it was at ca. 735 Ma that the igneous magmas that would become the Baltimore Mafic Complex (Smith and Barnes, 2008) and the 433 ± 2 Ma Sword Mountain Olivine Melilitite (Smith et al., 2004) were separated from the mantle. Late in the Neoproterozoic, additional crustal attenuation by mantle forces led to a second rifting that split Rodinia largely

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Figure 2. Regional geologic map of the Piedmont surrounding the Peach Bottom area, illustrating the locations of the Tucquan anticline, the microcontinental fragments (Brandywine and Baltimore), and the Cecil Island Arc that contains the Baltimore Mafic Complex, the Sykesville Formation, and the James Run Formation.

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along the Grenville orogen. The separation of these main continental fragments (Laurentia, Baltica, and West Gondwana) created ocean basins between them: Iapetan Ocean between Laurentia and Baltica; and the Rheic Ocean between Baltica-Laurentia and West Gondwana1. This separation was accompanied by a voluminous outpouring of largely basaltic magma along the continental margins of rifting phase, subaerial Catoctin Metabasalt along eastern Laurentia, and, mainly in Pennsylvania, Catoctin Metarhyolite from remelted Grenvillian lower crust. More chemically evolved drifting phase metabasalts were extruded within the Iapetan Ocean. These evolved, drifting stage metabasalts are represented by the Sams Creek Formation in north-central Maryland and the geochemically identical Fishing Creek Metabasalt (see Stop 4 below) in the Peach Bottom area. Iapetus (Rheic Ocean) widened during the Cambrian, perhaps to a width of more than 1000 or even 2000 km. During this time, continental fragments in the form of microcontinents (two of them presently represented by the Brandywine and Baltimore domes) moved about within the Iapetan basin. The Brandywine and the Baltimore microcontinents (Faill, 1997) drifted into position off this part of the Laurentian coast, probably in the Cambrian. They appear to be exotic to this part of the Grenville orogen for two reasons. First, although both microcontinents are composed of high-grade Grenville rocks, their schists and gneisses do not resemble lithically (lesser mafic gneiss) any of the Laurentian Grenville rocks to the northwest, not of the Blue Ridge, the Trenton Prong, Mine Ridge or the Honey Brook Upland, the Reading Prong, nor even of the Adirondacks (Rankin, 1975; Rankin et al., 1993). Second, the absence of metadiabase of Catoctin affinity (Smith and Barnes, 2004, their p. 42 and table 1—the dikes are the “older diabase” of Bascom and Stose, 1932) suggest that, ca. 570 Ma, they were not near this part of Laurentia during Catoctin magma activity. Their later positioning off Laurentia augmented the Octoraro basin by isolating the attenuated continental (Octoraro) and adjacent oceanic crust from the main part of Iapetus (Fig. 3). The arrival of the microcontinents offshore of Laurentia greatly changed the pattern of sedimentation. As with any subaerially exposed landmass, the Brandywine and Baltimore microcontinents were eroding and shedding sediment onto their surrounding margins in Iapetus. With their arrival to form the southeastern barrier of the Octoraro basin (current geographic directions), their northwestward-flowing sediment began adding to the basin (Fig. 3). In the absence of any other major sediment source (carbonate shelf on northwest side), the microcontinents became the principal supplier of siliciclastic sediment for the Octoraro basin. The sediment accumulation was probably characteristic of any continental passive margin, thickest at the continental edge and tapering into the basin. Deltas would have formed in front of the major rivers, and long-shore currents would have filled the 1

intervening areas with sediment to form a continuous bajada in the southeastern part of the Octoraro basin. The coarsest-grained sediment accumulated near the landmasses, whereas the finergrained fraction was carried out to the basin center or farther. The size and lateral extent of the bajadas and deltas (subaerial portions), and the shoreline positions to the northwest, would depend on the amount of subsidence of the basin floor and the volume of the entering sediment. Considering the tens of cubic kilometers volume of the present remnant of the Octoraro basin, the microcontinents must have been considerably larger than their present exposure would suggest. Laurentia (at least this part of it) apparently became less of a source of sediment during the Early Cambrian because the continental margin underwent a profound change at that time and became a widespread carbonate shelf that persisted well into the middle Ordovician. Some sands (Chilhowee and Gatesburg) from Laurentia were intermingled and incorporated within the carbonates during the Early Cambrian, but little siliciclastic material entered or crossed the carbonate shelf subsequently. The duration of the Octoraro basin is unknown. As argued below, it may have initiated as one of several intracontinental basins during the mid-Neoproterozoic rifting, and evolved into a continental margin basin with the opening of Iapetus late in the Neoproterozoic. When it ceased to be a sediment-accumulating feature through uplift or deformation is not known. We would suggest that it was unaffected by the Taconic orogeny because it contains none of the nappe structures that are so characteristic of Taconic deformation in southeastern Pennsylvania and western New Jersey. Also, considering the large dextral translation along the east-northeast–trending Cream Valley fault (in southeastern Pennsylvania) that presumably was active during the late Paleozoic Alleghany orogeny, the Octoraro basin may have been far removed from its present location during much of the Paleozoic. The Cecil Island Arc, comprising the Baltimore Mafic Complex, the James Run Formation, the Port Deposit, and intervening metasediments, developed during the middle Ordovician (Horton et al., 1998) and was probably thrust over the Octoraro basin during the Alleghany orogeny. However, it is not known if it impinged on the Octoraro basin before that time. THE PEACH BOTTOM AREA Nearly all of the rocks in the Peach Bottom area are siliciclastic metasedimentary rocks—the other rocks include metabasalts, diabases, and ultramafic rocks altered principally to serpentinite, talc, and listwaenite (mainly dolomite) (Table 1; see Appendix 1 for more detailed descriptions of each unit). The two principal stratigraphic units in the Peach Bottom area are the Peters Creek Formation and the Scott Creek Formation (south and north of the Delta Duplex, including the Peach Bottom Slate Belt), which together underlie 80% of the map area

Iapetus was the western part of the Rheic Ocean. Iapetus was closed by the docking of Avalonian (periGondwanan) microcontinents against Laurentia during the middle Paleozoic (Faill, 1997).

Peach Bottom area, Pennsylvania-Maryland Piedmont (Fig. 1). They represent two distinct portions of the continental passive margin sediments deposited on the northwest side of the Brandywine and Baltimore microcontinents. The Peters Creek Formation contains a number of upward-fining cycles (particularly in the lower, Puddle Duck Member) indicative of subaerial deposition that is characteristic of upper delta plains. Furthermore, the Peters Creek sediments are generally coarser grained, contain more quartz grains, and are more feldspathic than the Scott Creek sediments. This would suggest that the Peters Creek represents a more proximal part of the passive margin. The Scott Creek Formation is generally finer grained, with a greater proportion of phyllosilicate minerals, and it contains several intervals of dark gray to grayish-black phyllites that are meters- to tens of meters-thick. In addition, millimeter-scale laminations pervade much of the formation, suggestive of a shallow water depositional environment found in lakes or tidal flats. Neither of these formations appears to have accumulated much of their sediment in deep-water environments. First, turbidites are rare to absent for both units. Second, the low energy levels inferred from the more phyllitic intervals, especially the dark gray beds, would suggest deeper water (sub-wave-base), but their alternation with high-energy, quartzose beds would require large, frequent changes in water depths. Third, beach deposits have not been identified, which argues against frequent subaerial to deep-water transitions. On the other hand, the shoreline may have been more muddy than high energy. Fourth, the finer grained beds could have accumulated in flood plains and lakes within the interfluves. In other words, the vertical alteration of coarse and fine-grained deposits may simply reflect lateral shifting of river courses through time. Overall, it appears that the sediments in this part of the Octoraro basin were largely fluvial and thus rather

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high on the continental margin, mostly in the subaerial part of the bajadas and delta. The Sykesville Formation (including the horse(s) in the Delta Duplex) and the northern edge of the Baltimore Mafic Complex, underlie 15% of the Peach Bottom area. The Sykesville of the present Peach Bottom study area and the presumably equivalent Morgan Run Formation at the Liberty Reservoir of Carroll County, Maryland (Muller et al., 1989; Smith, 2006) is an ophiolitic mélange comprising poorly bedded sediment, wellbedded sediment, dismembered fragments of Bald Friar Metabasalt (Smith and Barnes, 1994, 2004; Smith, 2006), and bodies of steatized and carbonated ultramafites. This formation probably formed as an accretionary deposit in front of the advancing Baltimore Mafic Complex, but probably at sufficient distance from the Octoraro basin to have had no (little?) effect on its sedimentation or tectonic development. It impacted the Octoraro basin when the Baltimore Mafic Complex was accreted later in its history. The Delta Duplex (Fig. 1), comprising the Peach Bottom Slate, the Sykesville (northern horse), a large elongate body of serpentinite extending west from Delta, and various well-bedded schists, underlies 5% of the Peach Bottom area. The Peach Bottom Slate is unique in this part of the Pennsylvania and Maryland Piedmont, and its provenance is unknown. It is black, tough, extensively fractured (hence the large wastage in quarries), and contains significant acid insoluble carbon (>1%) and chloritoid (5%–15%). Bedding is usually not discernable either because of its original lithic uniformity or because it was obscured by subsequent deformation. The high carbon content suggests deep-water deposition (to preserve the organic matter), but its conformable and rather sharp contact with the overlying, high-energy Cardiff Quartzite would require a rapid shallowing of water depth,

Figure 3. Schematic cross section of the Octoraro basin early in the Ordovician, showing the inferred relations among the various sedimentary units, the microcontinental source of sediment, and a possible underlying attenuated crustal structure. The Brandywine (and Baltimore, not shown) microcontinent(s) shed sediment into the Octoraro basin during the early Paleozoic, creating a broad bajada (coalesced deltas) comprising the proximal Peters Creek Formation and the more distal Scott Creek Formation. The Octoraro Formation occupied the basin center. Mid-Neoproterozoic rifting created the attenuated continental crust over which the Octoraro basin formed, in which late Neoproterozoic sediments (Octoraro Formation) accumulated. The Baltimore Mafic Complex and Sykesville Formation were later obduced onto the Octoraro basin, possibly as late as the Permian Alleghanian orogeny. The position of the Neoproterozoic-Cambrian boundary is based primarily on the Catoctin-related Fishing Creek Metabasalt.

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TABLE 1. PEACH BOTTOM AREA STRATIGRAPHY JURASSIC/TRIASSIC Rossville Diabase (Jrd)—Homogeneous mafic igneous rocks occurring in subvertical dikes. The Rossville Diabase has age of 201.0 Ma (Dunning and Hodych, 1990). It crops out along the Norfolk & Southern railroad at the Williamsons Point tunnel south of Peach Bottom Station, and at the north end of the parking lot for Peach Bottom nuclear power plant. Quarryville Diabase (Trd)—Homogeneous, olivine mafic igneous rocks occurring in subvertical dikes. The Quarryville Diabase is in the swarm around and east of the Peach Bottom Marina and, based on 2nd Survey literature, lies under the flooded mouth of Peters Creek. The ages of the remaining stratigraphic units are very uncertain, probably within the span from late Neoproterozoic to Ordovician. The units are listed in sequence across the Peach Bottom Area map, from south to north. The age of the following two units is probably Cambrian to Ordovician. Baltimore Mafic Complex (bmc)—Serpentinite, with subordinate talc and magnesite (thickness not determined, >1000 m). Sinha et al. (1997) dated the Baltimore Mafic Complex at 489 ± 7 Ma. Sykesville Formation (sy)—A mélange of quartzose schist and metasandstone containing dismembered metabasalts, and boulders to large bodies of steatized ultramafics derived from Baltimore Mafic Complex (900–3000+ m). The age of the following unit may extend from late Neoproterozoic to Ordovician. Peters Creek Formation (pc)—A metasedimentary sequence of schists, quartzites, and phyllites (6400 m). Cooks Landing Member (new name) (pcc)—Thin- to medium-bedded quartzose schists and silty phyllites occurring in cyclic alternation (3100 m). Puddle Duck Member (new name) (pcp)—Medium- to thick-bedded metasandstones, schists, and phyllites, often arranged in upward fining cycles (3300 m thick). The age of the following two units may be late Neoproterozoic. Cardiff Quartzite (cd)—Thin- to medium-bedded sequence of interbedded quartzite and quartzose schist (30–100? m). Peach Bottom Slate (pb)—Very fine-grained homogeneous, chloritoid-bearing, black slate, with indecipherable bedding (1000+ m). The age of the following unit may be Ordovician Sykesville Formation (northern belt) (syn)—Metasediments, Bald Friar Metabasalt fragments [metabasalts], and listwaenite (carbonated and steatized ultramafite) (100 to 200 m). The Sykesville, derived in part from the Baltimore Mafic Complex, must be at least slightly younger than 489 ± 7 Ma. The age of the following unit may extend from late Neoproterozoic to Ordovician. Scott Creek Formation (new name) (sc)—Laminated schists, and silty phyllites, and dark gray phyllites, interbedded at various scales (4700 m). Whitaker Member (new name) (scw)—Interbedded thin- to thick-bedded phyllitic schist, laminated schist, and dark gray phyllite (800+ m). Coyne Lock Member (new name) (scc)—Variably laminated, thin- to medium-bedded, very fine- to finegrained schists with subordinate medium dark gray phyllitic schist and silty phyllite (2300 m). Bryansville Member (new name) (scb)—Complex sequence of quartzose schists, schists, and phyllites, and a metabasalt (Fishing Creek) (1600 m). Fishing Creek Metabasalt (scbf)—Medium dark green, coarsely laminated, epidote-bearing granular schist, occurring ~150 m from the top of the Bryansville Member (20 m). The Fishing Creek Metabasalt, geochemically evolved from ca. 570 Ma Catoctin Metabasalt, is perhaps on the order of 10 million years younger. The age of the following unit may be late Neoproterozoic. Octoraro Formation (oct)—Thin- to thick-bedded, albitic, quartzose schist with intervals chloritic schist (>1000 m thick).

Peach Bottom area, Pennsylvania-Maryland Piedmont possibly in conjunction with uplift of the source area. The Sykesville (north) horse contains fragments of the Bald Friar Metabasalt, a 9-m-thick body of listwaenite (an ultramafite steatized by carbonate-bearing fluids), and a large serpentinite body. This serpentinite resembles serpentinites of the Baltimore Mafic Complex and is probably a separated fragment of it. The provenance of the well-bedded schists is unknown. The southeastern edge of the Octoraro Formation underlies the northwestern edge of the Peach Bottom area. This probable Neoproterozoic deposit predates the younger (presumably, as argued below) Peters Creek and Scott Creek Formations in the Octoraro basin. STRUCTURE AND METAMORPHISM The most pervasive structure in the Peach Bottom area is the metamorphic foliation (S1) that parallels the compositional layering (bedding, S0) in the metasedimentary rocks. This thoroughly penetrative S1 fabric is a consequence of the very strong parallelism of the phyllosilicate minerals (primarily muscovite and chlorite, and biotite where present), microlaminations, and millimeter-scale laminations. We maintain that this widespread parallelism of fabric and bedding resulted from deep burial (10–15? km) of the undeformed sediments that induced extensive mineral recrystallization at the chlorite grade of greenschist facies metamorphism. In the absence of lateral tectonic stresses, the maximum principal stress during recrystallization was vertical, thereby aligning the phyllosilicate minerals with the still horizontal compositional layering, the bedding. Reviewers of this field guide, and other persons, raised the possibility that the parallelism of bedding (S0) with the primary S1 foliation may have occurred by some unspecified transposition of bedding. We reject this suggestion for several reasons. • First, the presence of compositional layering in these rocks is most simply explained as a sequential, episodic accumulation on a basin floor of sediment with contrasting properties, producing what is commonly called bedding. This interpretation is supported by the presence of numerous cross beds and troughs in the metasandstones, by upward-fining fluvial cycles, and by pillows in the metabasalts. • Second, what microscopic process (cataclasis or ductile flow [pure shear]? slip on parallel surfaces [simple shear]?) would have internally rotated bedding? We observed no structures in these rocks that reflect such processes. • Third, the foliation is mineralogically an integral fabric of the rock. How does one internally rotate the bedding but not the foliation? • Fourth, the foliation postdates the bedding. If the transposing preceded the foliation, the question of transposition is moot because what spatial frame of reference can be used to document and measure the internal rotation? It would be more appropriate to ask why the foliation parallels the bedding.

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• Fifth, the foliation parallels bedding across the entire ~30 km width of the Tucquan anticline. What vast geologic event or process could have internally rotated bedding so uniformly in such a huge volume of rock over such a large area? • Sixth, introducing transposition to explain the parallelism creates difficult questions (1 through 5 above) and unnecessary complications. It is not our intent to “prove” any specific “fact,” idea, or concept. Instead, we offer a story for the Peach Bottom area that is (1) simple and (2) consistent with features in the rocks. All the other structures (folds, faults, crenulations, and cleavages) postdate the foliation because they affect both the bedding and foliation equally. We believe that burial metamorphism supplies the simplest and most plausible explanation for the widespread parallelism of the foliation to bedding. The second most pervasive structure is the southeastward dip of bedding (and primary foliation), averaging 65° across the entire width of the Peach Bottom area. We suggest that this “monoclinal” structure was created by the development and growth of the Tucquan anticline, with the Peach Bottom area lying within the southeast limb (Fig. 4). We presume that the Tucquan is an Alleghanian structure, and that, consistent with Alleghanian tectonism in the foreland to the northwest, it grew above a basal décollement, with a thrust in the fold core that rose from the décollement. The trend of the Tucquan anticline (its hinge) is 065° azimuth, and its axis plunges gently to the southwest. Dynamically, the simplest model is movement perpendicular to the fold trend, driven by a maximum principal stress in the same direction, 335° azimuth (and subhorizontally). If the Scott Creek and Peters Creek rocks formed a single stratigraphic sequence, it would measure some 10 km thick. Burial metamorphism of such a section would produce a higher metamorphic grade in the deepest, oldest rocks relative to the shallowest, youngest ones (as indeed is seen farther northwest, in the structurally lower (and older) rocks near the Tucquan hinge). That virtually all the rocks across the Peach Bottom area are at chlorite grade, greenschist facies suggests that a structural break must exist that divides the “single section” into two or more parts that were assembled from separate, chlorite-grade areas. Indeed, the Peters Creek Formation was thrust over the Scott Creek Formation on the Delta Duplex, bringing together two approximately coeval blocks from different parts of the Brandywine and Baltimore microcontinental margins (Fig. 3). Similarly, the Sykesville Formation and Baltimore Mafic Complex from an oceanic area was obduced (thrust) over the Peters Creek Formation on the McGuigan thrust (Fig. 1). The third, rather widespread structure and most prevalent in and near the Delta Duplex, is the crenulation of the primary foliation, S1. This deformation is a spaced, nonpenetrative structure consisting of numerous parallel, closely spaced, narrow (submillimeter) zones at moderate to large angles to the primary foliation. In cross-section views, it appears as a second cleavage, S2, which is steeply dipping to subvertical, trending eastward at a small angle to the S1 foliation strike. The cleavage manifests itself on the S1 foliation surfaces as sets of numerous, minute

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(submillimeter), parallel folds (F2). The majority of these crenulations plunge ~40° (within a range of 10° to 60°) to the east; a smaller number of them plunge gently (0°–20°) to the southwest. The variation in the crenulation plunge (and therefore the trends of the crenulation cleavage) is presumed to result from local variation in stress direction and/or the S1 foliation during their development. The predominance of the easterly plunges suggests that the maximum principal stress producing the crenulations trended more northerly (~000° azimuth) than the earlier stress responsible for the southeast dip of bedding and foliation (335° azimuth). This sequence of stress regimes is consistent with that observed (northwesterly to northerly) in the Great Valley and Valley and Ridge (Anthracite) provinces to the north (MacLachlan et al., 1975; Faill, 1998). Similarly, the southwesterly plunges would indicate a more westerly directed maximum principal stress. Locally, within zones some 100–300 m wide, the S1 foliation trends more southwesterly, at some 20 degree angle to main structural grain. Small (20 cm to 2 m), moderately open folds are present sporadically throughout the Peach Bottom area. They generally occur in clusters of six or less folds. The fold plunges vary from cluster to cluster, ranging from subhorizontal to steeply southeast (~60°). Their sporadic occurrence, and diverse plunges, suggests that they represent local strain adjustments within

the more widespread thrusting and bed rotation. A very few larger folds are associated with small complex zones of faulting (wedges and duplexes). The traditional view of the rocks and structures in this part of the Pennsylvania-Maryland Piedmont is, as mentioned above, a regional syncline with the Peach Bottom Slate in its core being the youngest of the stratigraphic units. This interpretation is untenable in view of the geographic distribution of the various lithologies, the regional and local structure, and the metamorphism of these rocks. These various aspects include the following: (1) The metasediments southeast (Peters Creek Formation) and northwest of the slate belt are sufficiently different lithically to require a new stratigraphic formation name (Scott Creek Formation) for those northwest of the slate belt. (2) An ophiolitic mélange (Sykesville Formation) lies between the Baltimore Mafic Complex and the Peters Creek Formation, containing various arc-related metabasalts and numerous ultramafic fragments spalled from the Baltimore Mafic Complex. This same unit is present in the Delta Duplex, adjacent to the northwest side of the Peach Bottom Slate Belt. (3) Two thrust faults bound the Delta Duplex that comprises the Peach Bottom Slate Belt (the Peach Bottom Slate and

Figure 4. Schematic cross section illustrating the effects of Permian Alleghanian décollement tectonism on the Octoraro basin. The Octoraro and Scott Creek Formations occupy the southeast limb of the Tucquan anticline. The Peters Creek Formation was thrust over the Scott Creek Formation, riding on the Delta Duplex. The Cecil Island Arc originated in Iapetus during the Early Cambrian and was thrust (obduced) over the Octoraro basin during the Permian Allegany orogeny.

Peach Bottom area, Pennsylvania-Maryland Piedmont overlying Cardiff Quartzite, which together form a local anticline), a structural fragment (horse) of Sykesville Formation, and schists of unknown provenance. (4) The dominant metamorphic foliation (S1) parallels bedding in nearly all of the stratigraphic units. (5) The S1 foliation dips to the southeast at 60°–65° on average, across the entire 13-km-wide Peach Bottom area, as part of the southeast limb of the Tucquan anticline. (6) The presence of only lower greenschist facies (chlorite zone) metamorphism across the entire area imposes restrictions on the tectonic development. A simple regional syncline is not consistent with these aspects. A revised stratigraphy (as described above), with a few new units, is necessary to account for the lithic complexities revealed by the recent mapping. In addition, the areal distribution of these lithologies requires a new structural model, one dominated by thrust faults. A tectonic synthesis follows. TECTONIC SYNTHESIS Tectonic assembly of these various bodies of rock by thrusting seems to be the simplest model to account for the present distribution of the various lithologies and stratigraphic units in the Peach Bottom area (Fig. 4). The Octoraro basin continued to receive sediment primarily from the Brandywine and Baltimore microcontinents throughout the Ordovician. During the Late Ordovician, landmasses to the east that emerged during the Taconic orogeny produced sediment that spread over the submerging carbonate shelf north and west of the present location of the Peach Bottom area, commencing the development of the Appalachian basin. The Cecil Island Arc containing the Late Cambrian Baltimore Mafic Complex, and the somewhat younger Sykesville ophiolite mélange in front of it, advanced toward Laurentia throughout much of the early Paleozoic, and closed Iapetus against the Octoraro basin, perhaps as late as the early Silurian. The principal sedimentary units, the Peters Creek and the Scott Creek Formations, were deposited in the Octoraro Sea adjacent to the ocean, and may have been roughly coeval, being lateral facies of one another. The relation of the Peach Bottom Slate and Cardiff Quartzite to the other units is not known. That they lay structurally below the Peters Creek Formation suggests that they may have been early deposits in the Octoraro basin. The hinge of the gently southwest-plunging Tucquan anticline lays 20 km northwest of (and parallels) the Peach Bottom Slate belt; the Octoraro Formation extends from that hinge southeastward to the Peach Bottom area (Fig. 2). These pelitic rocks contain almandine garnets near and in the hinge, indicating an upper greenschist metamorphic facies for these oldest metasedimentary rocks. Also, some kyanite, probably with chloritoid, occurs near Rawlinsville on or near the crest of the Tucquan anticline (Hietanen, 1951). Biotite is present for a few kilometers from the hinge, but, for 20 km farther southeastward, the pelitic rocks are dominated by chlorite. The primary foliation, S1, is sub-

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horizontal in the hinge of the Tucquan anticline. Southeastward, the foliation dip gradually increases to 60°. This attitude persists throughout the remainder of the southeast limb, across the Peach Bottom Slate Belt, to the Baltimore Mafic Complex in Maryland. Where discernable, this foliation parallels bedding. Thus, the Peach Bottom Slate Belt, indeed the entire Peach Bottom area, lies within this chlorite domain of the Tucquan southeast limb. Traditionally, closure of Iapetus has been included in the Taconic orogeny, typically dated at ca. 458 Ma at the latitude of Pennsylvania (e.g., Rodgers, 1970). This would seem to be an appropriate time for the obduction of the Baltimore Mafic Complex and the Sykesville Formation into and/or onto the Octoraro basin. However, the dominance of a single, initially horizontal bed-parallel foliation resulting from burial metamorphism, and the absence of nappes and associated axial planar cleavage structures characteristic of Taconic deformation in the Great Valley in eastern Pennsylvania and New Jersey, implies that the Taconic orogeny probably did not affect the rocks now in the Peach Bottom area. Additionally, obduction over the Peters Creek sediments would have loaded those sediments to greater depths, thereby augmenting the burial metamorphism of those rocks to grades higher than lowest greenschist facies. These two aspects suggest that perhaps these rocks were distant from the area of Taconic activity, and that the obduction of the Baltimore Mafic Complex and Sykesville occurred later in the Paleozoic (postOrdovician). We cannot rule out that this corresponds to a period of baddeleyite rim overgrowth at ca. 310 Ma. We propose that the deformation that resulted in the steep dips of foliation and bedding, and the overprints of a second subvertical foliation and small folds, probably occurred during the late Paleozoic Alleghany orogeny. It was at this time that the African continent (western Gondwana) advanced toward Laurentia during the closure of the Rheic Ocean. This continental collision produced the décollement and thrust and fold tectonism that dominated the Alleghany orogeny in the central Appalachian foreland. The Piedmont, being hindward of the foreland, is similarly allochthonous, above the very same, subhorizontal décollement. We suggest that the Tucquan anticline grew over thrusts from this décollement (Fig. 4), creating throughout the southeast limb the steep southeast dips in the bedding and foliation in the Octoraro and Scott Creek Formations. Immediately after, the Peters Creek Formation and underlying Cardiff and Peach Bottom beds ramped up the southeast limb on the Delta Duplex. It is not known, nor suggested by us, when the Cecil Island Arc was obduced over the Peters Creek Formation. Figure 4 attempts to illustrate this tectonism, but somewhat unsuccessfully. The severest problem lies with the Delta Duplex. In general, horses (fragments) in a duplex can only be derived from the rocks of the adjacent hanging and footwalls. With the Delta Duplex, the hanging and foot walls consists of the Peters Creek and Scott Creek Formations, respectively, with only two to three kilometers of offset. But the horses are composed of the Peach Bottom Slate and Cardiff Quartzite and the Sykesville Formation, neither of which occur within the Scott Creek or

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Peters Creek rocks. How did they become incorporated within the duplex? An additional problem in Figure 4 lies with the thickness of the Peach Bottom Slate. The slate outcrop is ~1 km in width, overlain by the Cardiff Quartzite, yet the slate is portrayed as being more than 4 km thick in Figure 4. Much of that thickness could have been drawn as Octoraro Formation, but the relations at depth are completely unknown. The slate is not exposed elsewhere in this part of the Piedmont, the slate (and Cardiff) is in fault contact with the adjacent Peters Creek and Scott Creek formations, and so its original location in the Octoraro basin and its thickness are unknown. With these and other caveats, Figure 4 is at best an imperfect working hypothesis. ITINERARY The trip begins at the Sheraton Baltimore City Center Hotel with travel to Wakefield, Pennsylvania and nearby areas in southern Lancaster County, south-central Pennsylvania (Fig. 5). There, we will examine the nature of the Neoproterozoic and lower Paleozoic siliciclastic rocks of the Pennsylvania Piedmont (including the Peach Bottom Slate), and present a geologic history of the area. The field trip comprises visits to five outcrops on the east side of the Susquehanna River, from Pleasant Grove in southern-

Wakefield SQ

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most Lancaster County, Pennsylvania (very near the Maryland boundary) to Drumore some 9 km (5.5 miles) to the north (Fig. 6). Two of the stops are on the Norfolk and Southern Railroad property (Port Road Branch) along the east shore of the river2. The other three stops are along roadsides, which usually require no prior permissions. Only one stop (Stop 2) requires a fair amount of walking along the tracks (1.5 km each way), which is not difficult but does take time. The other stop along the railroad requires less walking. The three roadside stops are easily accessible. The Port Road Branch (originally the Columbia and Port Deposit Branch of the Pennsylvania Railroad) was electrified in the 1930s, which involved placing overhead “power takeoff” wires on catenaries that were supported by steel I-beam poles. Electric motive engines were discontinued in the 1970s (replaced by diesel-electric engines), and the catenaries were removed. The poles, however, remain, which has proved useful because they were individually numbered, from 1 to 800+, beginning at the south end of the branch at Havre de Grace, Maryland. The poles are spaced from 30 to 100 m apart (depending on the sharpness of the curves) and specific outcrops can be located as lying between, e.g., catenary poles 357 and 358. With a measuring tape, individual beds, contacts, structures, samples, and other features can be located within 1 m along the track. Therefore, to assist in finding the described features in the two railroad stops, the catenary pole number and number of meters are provided, in the format 423/45 m (i.e., 45 m north of pole 423). The meter distances are measured from the southern pole of the interval, increasing in the same direction as the ascending pole numbers, from south to north. Mileage Incremental Cumulative 0.0

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Figure 5. Road map showing the route from Baltimore to the Peach Bottom area (Wakefield, Pennsylvania).

Start at convention hotel. N to E FAYETTE STREET. Turn LEFT onto E FAYETTE STREET. Turn LEFT onto ST PAULS STREET. Turn RIGHT onto E LOMBARD STREET. Turn LEFT onto S HOWARD STREET. S HOWARD STREET becomes I-395 S. Merge onto I-95 N toward New York. Keep LEFT to take I-95 N. Toll after tunnel. Merge onto MD-24 N (Exit 77B) toward BEL AIR.

Some railroads in the United States have become quite restrictive with respect to access, even to the extent of arresting “trespassers.” The Norfolk-Southern has generally been quite tolerant of geologists visiting outcrops along the railroad, requesting only that care be taken about oncoming trains. Usually, trains do not run on these tracks during daylight hours unless the previous night’s scheduled runs were delayed. However, policies and practices can change between now (November 2009) and the day of the field trip.

Peach Bottom area, Pennsylvania-Maryland Piedmont 6.4 1.8 11.6

32.9 34.7 46.3

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Stay STRAIGHT to go onto MD 24. MD 24 becomes U.S.-1 N. Turn LEFT onto U.S.-222, follow U.S.-222. Turn RIGHT onto GOAT HILL ROAD. Proceed east on GOAT HILL ROAD. Park in small pull-off on the right side of road just beyond the small bridge over creek.

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Stop 1. Goat Hill Road: Sykesville Formation, Including the “Conowingo Creek metabasalt,” Serpentinite, and Steatized Ultramafite The Sykesville Formation (Muller et al., 1989) is herein extended into Lancaster County, Pennsylvania, from Carroll and Howard Counties, Maryland, where it was defined (Higgins and Conant, 1986; Muller et al., 1989). It lays in the southeastern part of the Peach Bottom area (Fig. 1), between the Baltimore Mafic

Figure 6. Local road map showing the route to each of the five field trip stops. bmc—Baltimore Mafic Complex; oct—Octoraro Formation; pb—Peach Bottom Slate; pc—Peters Creek Formation; pcc—Cooks Landing Member; pcp—Puddle Duck Member; scb—Bryansville Member; scc—Coyne Lock Member of the Scott Creek Formation; scw—Whitaker Member; syn—Sykesville (north) Formation.

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Complex and the Peters Creek Formation. The Sykesville Formation comprises the southernmost part of what had previously been mapped in the Susquehanna River area as the Peters Creek Formation (Knopf and Jonas, 1923), and as the Wissahickon Formation (Knopf and Jonas, 1929; Southwick and Owens, 1968). The Sykesville Formation in the area of Bald Friar Hill (Fig. 1) just NW of the Baltimore Mafic Complex, Maryland, includes at least four fragments (large clasts) of Bald Friar Metabasalt (Smith and Barnes, 1994, 2004; Smith, 2006), one of which includes relatively undeformed pillows. Their preservation here is similar to the situation at Gilpins Falls, Cecil County, Maryland, where pillows are well preserved in the James Run volcanics (Higgins, 1977). However, Bald Friar Hill is best known for its talc quarries in steatized fragments of ultramafite. On this trip, the Bald Friar Metabasalt will be seen in Sykesville Formation just north of the Peach Bottom Slate. The 489 ± 7 Ma (Sinha et al., 1997) Baltimore Mafic Complex lies just 3.5 km south of Stop 1. The magma that formed the Baltimore Mafic Complex was isolated from the mantle at 735 Ma (based on 187Os/188Os by Ryan Mathur, Juniata College, on an Os-Ir-Ru micronugget supplied by Smith and Barnes (2008). This same 735 Ma age was obtained by Ken Foland as a TNd mantle separation age for the 433 ± 2 Ma Sword Mountain Olivine Melilitite in the Valley and Ridge province of northern Washington County, Maryland (Smith et al., 2004). The 735 Ma age also happens to be the approximate igneous age of the Robertson River Igneous Suite in Virginia (Tollo and Aleinikoff, 1996; Smith, 2003). All three units are generally regarded to have initiated because of extension, i.e., rifting within Rodinia. It is not known if Rodinia was separated by this 735 Ma extension to form an intervening ocean. The extension may have proceeded only to forming several intracratonic basins over attenuated continental crust, analogous to the Late Triassic basins in eastern North America. The Sykesville Formation formed as an accretionary ophiolite mélange in front of the advancing Cecil Island Arc and Baltimore Mafic Complex probably during the Early Ordovician as the Iapetan Ocean was closed against eastern Laurentia. The provenances of the Sykesville include: siliciclastic sediment from nearby microcontinents, cannibalized blocks from itself, both forearc and backarc basalt flows, and various spalled Baltimore Mafic Complex ultramafic bodies as the Cecil island arc advanced by thrusting over the Sykesville mélange. The resulting wedge-shaped sedimentary body contained both wellbedded parts, and chaotic parts churned by the advancing island arc. Davis M. Lapham of the Pennsylvania Geological Survey mapped a large number of these mafic and ultramafic bodies during the 1950s and 1960s. Lapham’s map of this area has been incorporated in the Peach Bottom area geologic map for the Pennsylvania Geological Survey, recently completed by Rodger T. Faill (unpublished). The ultramafites generally have not preserved their original mineralogy in the central Appalachian Piedmont. Their chemistry is so different from that of the enclosing country rock (mostly

siliciclastic schists) that, even under the fairly low-grade metamorphism, contact metasomatism between them creates a blackwall reaction. The blackwall spans both the ultramafite and country rock, and consists of a series of mineralogical zones (Sanford, 1982). Blackwalls in different regions have many common features, regardless of different lithologies and different metamorphic grades. The blackwall zones in this part of the Piedmont (from ultramafite to country rock) can be summarized as: ultramafite (usually antigorite in this area); talc + ferroan magnesite; talc; tremolite-actinolite + clinochlore; clinochlore (the “black” in blackwall); and country rock. Because chromium remains immobile in the ultramafite, and titanium in the country rock, the location of the original contact can be determined, as was done by Smith (1993, 1994) on the south side of the Peach Bottom slate belt. We will see that blackwall at Stop 2B. Typically during blackwall formation, Si, Fe, and CO2 are introduced into the ultramafite, and H2O and Mg are lost from it. Where CO2 is readily available, a carbonation process yields a listwaenite, typically composed of dolomite or magnesite, talc, and quartz. These can resemble marbles to the uninitiated, but the original disseminated chromite proves their origin. Where CO2 is somewhat less available, a steatization process produces more talc and less dolomite or magnesite. The alteration of an ultramafite, in many situations, increases its volume at the expense of the country rock. This, together with the formation of steatite (talc), increases the ductility of the ultramafite to a level only slightly greater than bubble gum. Under conditions of differential stress, the ultramafite acts as a lubricant that facilitates movement within the rock. For this reason, altered ultramafites are often associated with major faults. As an ophiolitic mélange, the Sykesville Formation comprises five main lithologies: massive sandstone; bedded schist; diamictite, metabasalt blocks of two distinct compositions; and steatized ultramafic blocks. One metabasalt and the ultramafic rocks can be examined in Stop 1. The rocks at Stop 1 lie on the northern edge of Sykesville outcrop belt, presumably very close to the (local) base of the unit, just above the McGuigan thrust (Fig. 1) that carried the Sykesville Formation over the Peters Creek Formation. Stop 1 consists of three outcrops on the south side of Goat Hill Road, just east of a small bridge over Conowingo Creek (its upper reaches), 0.4 km east of U.S. Route 222, southeast of Wakefield. The first exposure is just east of the parking area, 55 m southeast of the bridge and 3–4 m above road level. It is one of several discontinuous exposures of the “Conowingo Creek metabasalt” (informal name, Smith and Barnes, 2004; Smith, 2006). The metabasalt is medium dark greenish to greenish-gray, containing mostly fine-grained epidote (zoisite?), along with chlorite and actinolite. The rock is very thin to medium bedded, and laminated. The metabasalt is dense and has been variably sheared. Probable vesicles suggest that these blocks are fragments of basalt flows. Pillow structure can be discerned in some of the other exposures to the west, indicating tops to the south. The dominant foliation (bedding) here dips moderately eastward,

Peach Bottom area, Pennsylvania-Maryland Piedmont

Rock /

Chondites

at 351–47 (strike azimuth-dip°), with a pronounced down dip lineation. This attitude is at variance with most other bedding and S1 foliation orientations throughout the Peach Bottom area. This divergent attitude may reflect rotation of this metabasalt fragment during deposition within the mélange. On the other hand, the position of this outcrop in a low mound on this north-facing slope suggests another possibility—recent slumping. Two other foliations are present as well. The significance of the “Conowingo Creek metabasalt” is that its chondrite-normalized rare-earth element (REE) plots (lower two plots in Fig. 7) indicate an intermediate REE depletion, a “steer-horn” pattern that is generally considered characteristic of boninites. Modern boninites are typically found only in association with the forearc portion of island arcs above subduction zones. The “Conowingo Creek metabasalt” is one of two metabasalt compositions in the Sykesville Formation, the other being the Bald Friar Metabasalt (Smith and Barnes, 2004; Smith, 2006). The Bald Friar Metabasalt, exposed southwest of here on the east shore of the Susquehanna River (as well as in Chester and York Counties, Pennsylvania, and Carroll County, Maryland, and also on the northwest side of the Peach Bottom Slate, Lancaster County), is probably a backarc basin basalt (BABB) (Smith, 2006) that typically forms above linear spreading cen-

Figure 7. Chondrite normalized rare-earth elements for three Conowingo Creek metabasalt samples from the area of Stop 1 to Goat Hill Road. The uppermost plot is for sample CONJSEIII from a 0.25 m × 0.4 m pillow. The high TiO2 content of 5.1% suggests that it has suffered severely from “pillow enrichment” (major elements in glass lost, but not immobile trace elements), but it seems to have been a normal ocean floor basalt). It was collected from the pasture on the north side of Conowingo Creek on the south side of Goat Hill Road, ~275 m E of U.S. Route 222. The middle plot is for sample CONJSEII from a mass of possible 0.2 m × 0.8 m sheared pillows, 80 m S15 W of the bridge. The lowermost plot is for sample CONJSE from 55 m SE of the bridge. The middle and lower plots exhibit classic steer-horn patterns, considered characteristic of boninitic metabasalts.

13

ters, probably (in this case) from the backarc (southeast) side of the island arc system. The second outcrop, 52 m to the northeast along the road from the first, exposes sheared serpentinite overlying talcmagnesite (ferroan) containing trace octahedral chromite. The presence of euhedral chromite rules out a carbonate protolith. The serpentinite and talc are probably altered from a peridotite-dunite ultramafite (originally probably 95% forsterite and 1% chromite). The present lithic composition varies considerably from place to place on this hillside, including serpentinite, talc, magnesite, and trace chromite, reflecting significant carbonation as well as steatization. The absence of an intense thermal aureole (and it would have been immense, given the very hot temperature [~1200 °C for a forsterite melt]) indicates that this body was emplaced cold, either as a horse within a fault zone, or as block within the mélange. The primary foliation dips moderately to the southeast, commensurate to S1 dips throughout the Peach Bottom area. A low rib running up the hill leads to a small quarry in mostly serpentinites, with some forsterite. Much of this hillside here and to the west is underlain by slightly greenish dark gray, very fine grained serpentinites as evidenced by the numerous boulders and cobbles. A rib of serpentinite parallels the slope contour some 30 m up from the road, and profuse amounts of small to large boulders of serpentinite lie across the slope. Intermingled with the serpentinite blocks are cobbles and boulders of very dark gray rock, a blackwall product comprised mainly of chlorite. The third “exposure” consists of four “Conowingo Creek metabasalt” outcrops, beginning ~80 m SSW of the bridge and extending discontinuously westward for another ~140 m along the south side of Conowingo Creek. The very good outcrops at the west end (220 m from the bridge) appear to be fresh exposures uncovered by stream erosion. The metabasalt here is different from the metabasalt near the bridge in that it contains considerably more zoisite (white grains) epidote that is apparently filling amygdules. The rock appears massive, but a pervasive (if faint) foliation does dip steeply to the southeast, ~067–74, with a lineation trending 122° azimuth probably produced by the intersection with an even fainter subvertical foliation trending 167° azimuth. Bands of zoisite concentrations bend around this lineation, perhaps representing pillow axes. Seventy meters to the northeast is a smooth surface of metabasalt sloping toward the creek. Little can be ascertained from this outcrop. Nearby, on the flood plain, is a large (60–70 cm) boulder of brecciated quartz and interlayered medium dark gray, fine-grained schist. This boulder may have come from the McGuigan Thrust, some distance upstream to the northeast. Seventy meters farther to the northeast (80 m from the bridge) are a pair of outcrops partway up the slope, beyond an electric fence (sometimes activated). The 6-m-wide by 3.5-m-high outcrop on the west appears to consist of a mass of 0.2 by 0.8 m pillows, the long axes of which trend 160° azimuth, plunging 53 SE. Locally at least, apparent S0 trends 073–67, parallel to the primary S1 foliation. A second planar fabric, dipping variably to the northeast, ~165–61, may be in response to the pillow structure.

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The intersection of these two planar fabrics has produced the prominent southeast-plunging lineation on S1. Summary Stop 1 represents the northernmost edge of the Cecil Island Arc system that includes metasediments, metabasalts, and steatized ultramafites in the Sykesville Formation (ophiolitic mélange), the Baltimore Mafic Complex (ultramafic intrusions), and perhaps the James Run Formation (volcanics), the Port Deposit Tonalite, and intervening siliciclastic metasedimentary rocks. Attenuated fragments within the Sykesville (ultramafites and the Bald Friar Metabasalt) appear on the north side of the Peach Bottom slate belt (Fig. 8, Stop 2E), and an ultramafite is present on the south side (Stop 2B), presumably emplaced by a major thrust (Smith, 1993).

The Cecil Island Arc was active from the (Late?) Cambrian, through the Ordovician, and possibly into the Silurian. There is no evidence that the Taconic Orogeny affected this island arc system (other than intrusion by Ordovician plutons). The recumbent nappes that characterize the Taconic deformation in the Laurentian margin deposits to the north and east of the Peach Bottom area are not present here. No clear-cut Taconic radiometric dates have yet been reported for the Peach Bottom area, nor is there evidence of Taconic activity to the north, in the foreland Great Valley province west of the Susquehanna River. Most likely (or most simply), the Cecil island arc system was not near the Laurentian margin affected by the Taconic orogeny, but was transported here at a later time, probably during the Permian Alleghanian orogeny. Mileage Incremental Cumulative 0.0 0.4

54.4 54.8

1.0

55.8

2.0

57.8

0.9

58.7

0.6

59.3

0.3

59.6

Return to U.S. ROUTE 222. Turn RIGHT onto U.S. ROUTE 222, proceed into WAKEFIELD. Turn LEFT onto PEACH BOTTOM ROAD. Intersection with CHERRY HILL ROAD. Continue west on PEACH BOTTOM ROAD. Intersection with RIVERVIEW ROAD. Continue west on PEACH BOTTOM ROAD. Intersection with SLATE HILL ROAD. Turn left onto bridge over PETERS CREEK. Continue on south side of creek, around curve, and onto flat area with railroad siding. Park on grassy area between road and railroad. Walk onto tracks, turn RIGHT, and cross railroad bridge over PETERS CREEK.

Stop 2. Peach Bottom Station: Peters Creek Formation (Puddle Duck Member), Cardiff Quartzite, Peach Bottom Slate, and Sykesville Formation (North)

Figure 8. Topographic map of Stop 2 showing locations of the five sub-stops, A through E, along the Norfolk and Southern Railroad. The parking area is just off the southeast corner of the map. The numbers along the shoreline, from 422 to 446, are the catenary pole numbers along the track. The topographic contours are labeled in feet above sea level.

Stop 2 crosses the most complete exposure of the rocks in the Delta Duplex (Fig. 9). It begins near the base of the Puddle Duck Member (new name, Table 1, also see Appendix) of the Peters Creek Formation (Stop 2A), which is in contact with the top of the Cardiff Quartzite. The presence of a thin layer of talc with an attendant blackwall reaction zone (Stop 2B) suggests that this contact is faulted. It is the southern edge of the Delta Duplex. Northward along this railroad exposure, the duplex contains the Cardiff Quartzite, the Peach Bottom Slate, and a large fragment of the Sykesville (north) Formation. This northern edge (fault contact) of the Peach Bottom Slate is probably along a talc (steatite) which is mineralogically and geochemically identical to the one exposed at Stop 2B (Smith, 1993).

Peach Bottom area, Pennsylvania-Maryland Piedmont The Cardiff Quartzite stratigraphically overlies the Peach Bottom Slate (Stop 2C). This contact appears to be conformable, but much of the Peach Bottom schist (the upper, transitional part of the Peach Bottom Slate; Table 1; Appendix) may have been eroded prior to the deposition of the high-energy Cardiff quartzites. The Peach Bottom Slate (below the upper schist) is fairly uniform in character across the 500-m-wide belt. It is a tough, high-carbon, black slate containing significant chloritoid. A small quarry is at Stop 2D, and a much larger quarry lays a short distance to the north. The northwest contact of the slate is covered, but 5 ± 2 m beyond the slate (and in fault contact with) is a high-chromium dolomite (listwaenite) (Stop 2E) that is probably a steatized remnant of a small ultramafite body through which hot carbonaterich solutions altered. It lies in the same zone as the talc (steatite) farther uphill at an approximation elevation of 300 feet. To the north are a few analyzed Bald Friar Metabasalt fragments (Smith, 2006, his appendix II and table 1) interspersed within bedded quartzose schist. The presence of the ultramafite and the Bald Friar Metabasalt suggests correlation with the Sykesville Formation (hence the term “Sykesville (north) Formation”). The north contact of the Delta Duplex lies in the moderately large valley beyond (Whitaker Station). The Whitaker Member of the Scott Creek Formation lies on the other side of the valley.

Figure 9. Geologic map of the central portion of the Delta Duplex. The duplex comprises the Peach Bottom Slate (pb) and the overlying Cardiff Quartzite (cd), and the Sykesville (north) Formation (syn). The duplex is in fault (probably thrust) contact with the Peters Creek Formation, Puddle Duck Member (pcp) on the southeast, and the Scott Creek Formation, Whitaker Member (scw) on the northwest. The Coyne Lock Member of the Scott Creek Formation (scc) underlies the northwest part of the map. The now-inactive slate quarries are: 1—Gorsuch open-pit; 2—Faulk-Jones; 3—Kell; 4—Funkhauser (joined from original Johnson, R.L. Jones, and McLaughlin quarries; 5—John Humphrey; 6—Edward Evans.

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To reach the outcrop Stop 2A, ascend the track, turn right (north) and cross the railroad bridge over Peters Creek. The first catenary pole north of the creek is number 421. Walk past the first outcrop, which is very sandy and deformed, and continue to the next large outcrop, which extends from pole 424/13 m to pole 426/10 m. (Pole interval 424–425 is ~80 m, both 425–426 and 426–427 are ~70 m.) Pole 425 has a large, silver-colored, ticking signal box mounted on it; a wooden pole just to the south holds a smaller silver box. Stop 2A. Peach Bottom Station: Puddle Duck Member, Peters Creek Formation Three different lithic assemblages are present in this outcrop, each showing different sedimentary structures, reflecting three contrasting depositional environments. From the north, the oldest beds are very thick bedded, quartzose sandstones; the medial portion displays upward-fining cycles; and the youngest, in the southern part of the outcrop, exhibits alternating parallel-bedded silty and argillaceous layers. All three of these assemblages, and one or more others, reappear frequently throughout the Peters Creek Formation. We suggest that these repetitions resulted from a continual shifting of depositional environments on the delta and continental margin. The oldest beds in this part of the outcrop, from pole 426/ 10 m to 425/25 m, are quartzose sandstones in thick and very thick beds, many of which exhibit cross beds and troughs, indicating a rather high energy environment which has winnowed out the finer grained sediment. Intervals of parallel bedded, wellfoliated sandstone, with flat but somewhat undulose bedding (cm-scale vertically, over 1 m lateral) are interspersed. Possible environments could be a point bar in a river, or a beach. The beds are inferred to be right-side-up (tops to the southeast) based on the cross beds and trough cutouts. Within this stretch, two intervals of sandstone are sufficiently resistant to form protruding “spines” on the slope above the tracks. The medial stretch of this outcrop, from pole 425/25 m to 424/29 m, exhibits an upward-fining pattern of sandstones overlain by siltstones and thence mudstones (phyllites). Generally, the basal sandstone has a sharp base, overlying (and even cutting into) underlying phyllite and silty phyllite. The cycles range in thickness from 1 to 5 m; some smaller ones appear incomplete, as if interrupted in their development. Some small cycles are nested within larger cycles, and some cycles lack sand in their basal layers. The bases of two particularly prominent cycles are at 424/ 29 m and 424/32 m. The youngest beds, in the southern part of the outcrop from pole 424/29 m to 424/13 m, present a quite different pattern— alternating beds of siltstone and mudstone (phyllite). The beds are thin to medium bedded, planar bedded, parallel bedded, with sharp tops and bottoms. Truncations suggest stratigraphic tops are to the southeast. The beds are rarely laminated. Channels are absent, but the beds thin and thicken laterally only gradually. Overall, this part is quite phyllitic, and contains scattered

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pyrite crystals. This pattern is suggestive of turbidites, but the upward fining within beds is not present. The lateral continuity of the beds suggests a subaqueous environment, below wave base. Possible environments could include a deltaic lagoon, or bayou, or perhaps a pond or ephemeral lake higher on the delta. Any one of these would provide water depths too shallow for much wave action, and yet could be contiguous with the other, higher energy environment As is true for the entire Peach Bottom area, the dominant structure at Stop 2A is the moderately southeast dip of bedding (S0) and primary foliation (S1), at 039–64. A steeper, less well developed foliation (S2) dips more steeply, also to the southeast, at 043–81. A more discontinuous (spaced) and infrequently present planar structure is a moderately planar fracture set (S3) that trends more northerly than the other foliations, at ~019–60. Crenulations appear only on the primary foliation, S1. The crenulations are asymmetric microfolds on a submillimeter scale, verging northwest. They occur as sets of numerous parallel lineations on a single foliation surface, covering areas as small as 10 cm2 to more than a square meter. They are very long relative to their wavelengths, a ratio of 100:1 being common. In general, they plunge moderately, gently, or subhorizontally to the northeast and, to a lesser extent, to the southwest. Mesoscale folds (centimeter to decimeter in size) are sporadically present, especially from pole 424/17–424/30 m. They are moderately tight (interlimb angles of 60°–120°) with axial planes subparallel to S2. The S2 fans in the phyllitic hinges of some folds but not the more sandy layers. The fold axes tend to plunge moderately to the northeast, similar to that of the crenulations. Faults are not common. Wedges, contraction faults of single, or a few layers, appear occasionally, as at pole 424/64 m. A minor fault zone is present at pole 424/17 m. Proceed northward along track from pole 425 to pole 429. Northward from pole 428 is very silvery phyllitic schist and dark phyllite, somewhat slippery on foliation, as if containing paragonite. Stop 2B. Peach Bottom Station: Cardiff Quartzite and Ultramafite The lithology changes markedly over a span of 6–8 m at pole 429, from the phyllites of the Peters Creek Formation on the south to the quartzites of the Cardiff Quartzite on the north. What is significant here is the nature of the contact. Traditionally, it has been considered a conformable contact, in part because there is no angular discordance, and in part there was no reason to question its being conformable. We argue otherwise. At this contact (centered on pole 429/0.0 m) is a 1.9-m-thick talc-magnesite-chlorite schist zone that is not immediately evident without some digging. The talc formed by steatization of an ultramafic rock, probably a peridotite composed of forsterite, pyroxene, and chromite during burial metamorphism. Blackwall reaction, especially with the more pelitic Peters Creek on the south side, resulted in 3 cm dark, chloritic blackwall,

17 cm of slightly altered Peters Creek, 3 cm moderately altered, magnetite-rich Peters Creek, 20 cm of moderately altered Peters Creek, and then the 1.9 m of talc-magnesite-chlorite schist. On the north side, the talc lies in contact with a chlorite schist at pole 429/1.0 m that changes northward to a chlorite-muscovite schist (including a 20-cm-thick very chloritic schist) and then a quartzmuscovite-chlorite schist to a quartzite (Cardiff), at pole 429/ 3.7 m. Interestingly, the dominant foliation in this zone is subvertical, 051–85, subparallel to the S2 foliation farther south at Stop 2A in the Peters Creek Formation. However, by pole 429/3.0, the moderately dipping attitude (043–64) has returned to the primary foliation, which continues through the Cardiff exposure. We suggest that this zone represents a major fault between the Peters Creek Formation and the Cardiff Quartzite. An ultramafic sliver is not a natural component of a sedimentary sequence (excepting mélanges), even at a sedimentary contact. The absence of any thermal aureole indicates that the ultramafite was not intruded as magma. We suggest that the ultramafite slice was caught up in a fault zone at depth, and that its talcy nature facilitated movement on the fault surface. This is not a local occurrence. Talc has been found at this horizon at two other locations, one as far from here as Whiteford, Maryland, 8 km to the west, on the north side of the slate belt. The presence of talc between the Cardiff Quartzite and the Peters Creek Formation is apparently widespread. The development of this fault may well have been early, as a separate event preceding the Delta Duplex movements. The Cardiff Quartzite is distinctive, consisting predominantly of clean quartzites (>95% quartz), alternating with subordinate amounts of quartz-muscovite-chlorite schists, especially in the lower part. Many of the quartzites are thickly laminated with dark partings, both near the top and near the base. Bedding in the quartzites is mostly thin to medium (from 5 to 10 cm, some as much as 20 cm). The thickness of many quartzite beds is not constant—they thin and thicken by as much as 20%–50% over a meter laterally. Cross bedding is present in a few of the beds. Bed surfaces tend to be smooth and undulose, or even flat, but a few bedding surfaces are regular. Very thin (<1 mm) muscovite partings are between some quartzite beds. Pebbles up to 2 cm in size occur in quartzite beds near the top, between 429/4.0 and 429/ 8.0 m. A few pebbles may also be present lower in the section, at 429/~23 m. The pebbles appear to be flattened in the foliation, and elongated subhorizontally, in the strike direction. The Cardiff Quartzite becomes progressively more schistose down section, beginning ~429/18 m. In this lower part, intervals of quartzite (40–200 cm) are interspersed with generally thinner intervals (10–50 cm) of very fine grained quartz-muscovitechlorite schist. Contacts between the contrasting intervals are generally sharp, smooth, and undulose. Very thin laminae of white mica are common at the contacts between the quartzites and overlying schists. The bedding S0 (= primary foliation S1) dips variably (55°– 72°) to the southeast, striking ~040. The second foliation, S2, is present. Crenulations, mostly moderately northeast plunging,

Peach Bottom area, Pennsylvania-Maryland Piedmont are present on many bedding surfaces, especially in the coarsely laminated quartzite at pole 429/21 m. The Cardiff Quartzite surrounds much of the slate belt. It continues uninterrupted (with one exception) on the southeast side of the belt, from Octoraro Creek some 8 km to the east to Broad Creek 12 km to the southwest. The Cardiff can be traced around both ends of the belt, and partway, on the northwest side, back to the Susquehanna River. On the western end, float or outcrop can be found as far as Cardiff; on the eastern end, evidence of it dies out under the colluvium and saprolite on the uplands around U.S. Route 272. Its apparent absence on the northwest side between U.S. Route 272 and Cardiff is attributed to its being faulted out by the Delta Duplex. The Cardiff Quartzite is also absent for a 1 km stretch along the southeast outcrop, ~5 km southwest of here. The Peters Creek Formation lies in contact with the Peach Bottom Slate, and the Cardiff is unusually thin just to the east and west. Either faulting cut out the Cardiff Quartzite, or it was not deposited in this locality. Proceed northward along the track to catenary pole 430. Stop 2C. Peach Bottom Station: Cardiff Quartzite and the Peach Bottom “Schist” The fault between the Peters Creek Formation and the Cardiff Quartzite inferred from the presence of ultramafites points to a tectonic assembly from different original settings. In contrast, the apparent conformal contact between the Peach Bottom Slate and the Cardiff Quartzite indicate that those two form a coherent sedimentary package, one distinct in provenance and possibly history from the surrounding Peters Creek and Scott Creek sedimentary package. The contact between the Peach Bottom Slate (Peach Bottom “schist”) and the Cardiff Quartzite is placed at pole 429/ 28.8 m. The contact appears conformal because small angular fragments of black slate are imbedded in quartzite beds, the Cardiff beds parallel the Peach Bottom beds, and no structural break is discernable at or near the change in lithology. The contact is not sharp, but consists of a transition from silty slate to quartzite over a zone some 4 m thick. The lowest 3 m of the Cardiff (to the south) consists of alternating thin to medium bedded quartzite and very thin to medium bedded schist. To the north, the uppermost part of the Peach Bottom Slate (“schist”) is a 20-cm-thick dark schist overlying a 30-cm-thick interval of dark schist and quartzite, all rather contorted and intruded by vein quartz. Below is 30 cm of dark, sparkly (chloritoid-rich) schist. The slate over the next 20 m is an interlayered sequence of thin to medium layers of slate and silty slate. These rocks have been called the Peach Bottom “schist” (Behre, 1933) even though they bear no resemblance to the coarse-grained quartzmuscovite-chlorite fabric usually associated with pelitic schists. It is the presence of the “silty” slate beds near the Cardiff Quartzite that distinguishes the “schist” from the slates throughout most of the belt, especially those worked in the numerous quarries.

17

Another characteristic of the “schist” is the common presence of small (1–5 mm) brown spots, and angular vugs, which probably held iron carbonate crystals. A few layers of “graphitic” slate occur sporadically through the “schist.” These layers are softer with a less pronounced cleavage than the more common slate. Chloritoid is present (up to 25%) in nearly all the “schist,” as well as in the rest of the Peach Bottom Slate. Several layers containing numerous pyrite cubes occur throughout the “schist.” Some quartz veins in the “schist” contain lazulite, a blue phosphate mineral. It is worth speculating about the implications of this large change from Peach Bottom to Cardiff lithology over such a thin interval. The predominance of phyllosilicate minerals over detrital quartz suggests a low-energy depositional environment lacking currents, as in a standing water body, below wave base. In contrast, the near absence of phyllosilicate minerals in the Cardiff quartzites indicates a high-energy depositional environment typical of a major fluvial channel or a beach. The thinness of the transition interval implies either a very rapid emergence of the subaqueous depositional site, or that some of the original, much thicker transition was removed by erosion. The marked lithic contrast between the silty slates below the contact, and the quartz-muscovite-chlorite phyllitic schist interbedded with the quartzites points to a greater change in sediment provenance and deposition than can be explained by simple emergence. We suggest that a significant amount (several tens of meters) of sediment was removed by the arrival of the higher energy environment of Cardiff sedimentation, and that the contact is a disconformity. The transition from “schist” to slate is not easily drawn because the frequency of “silty” layers progressively lessens to the north. Pole 432/00 m is probably as good an arbitrary contact as any (Smith, 1993). A 6-m-wide fault zone occurs just north of pole 432, at approximately pole 432/06 m. It is a mylonite, lacking the well-developed foliation so common in the slate, and it contains numerous “balls” (5–10 cm) of vein quartz. It trends 022–85, distinctly across the regional grain, subparallel to the S3 foliation. A 276 ± 6 Ma K-Ar date was obtained from a sample of this mylonite zone (Smith, 1994). Proceed northward along the track to catenary pole 436. Stop 2D: The Peach Bottom Station: Peach Bottom Slate at the Bonsell and Yard Quarry The small Bonsell and Yard slate quarry lies between poles 436 and 437. The quarry is ~6–8 m wide and extends into the slope some 30 m. A 40Ar/40K whole-rock analysis by Kruger Enterprises, Inc., (16 August 1994) yielded an age estimate of 279 ± 6 Ma (Smith, 1994) for a sample collected from this quarry (10 m northeast of the north edge of the culvert at 436/18.8m, 1.6 m above water level). The Peach Bottom Slate is remarkably uniform in lithic character across the belt, which is some 600 m wide here along the Susquehanna River. Along its length, the slate belt ranges in width from 200 to 1000 m, a variability due to structural

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complications within the Delta Duplex rather than real thickness changes. The slate is dark to very dark gray, very fine grained, generally massive (bedding usually cannot be discerned), with a pronounced cleavage, sometimes with a second cleavage, and often crossed by large, moderately spaced fractures of various orientations. It is considered the hardest, toughest slate to have been produced in this country—production all but ceased by the 1920s (Behre, 1933; Berkheiser, 1994). The rock here at the Bonsell and Yard pit is characteristic of the better commercial grade slate. The slate is characteristically very dark gray, homogeneous, argillaceous, quite heavy, and with a very planar cleavage. However, a few layers here are siltier than others, even this far below the formation top; they may represent original bedding. Load casts may also be present. Valentino (1994, p. 98–99) attributes the dark color of the Peach Bottom Slate to extremely fine-grained micas produced by ultramylonitization in a shear zone rather than the presence of carbon. Analyses (1995, by Activation Laboratories, Ltd.) of 13 samples collected across most of the slate belt along the railroad indicate an enhanced level of carbon (median value 1.53%, ranging from 1.22% to 2.22%). Interestingly, carbon levels are half as much as that along both sides of the slate belt—from 0.6% to 0.9% in the 20 m of slate adjacent to the Cardiff, and from 0.36% to 0.95% over 60 m on the northwest side. Throughout the slate belt, four minerals comprise 90% of the slate: muscovite, quartz, chlorite, and chloritoid. Accessory minerals include opaques (magnetite, hematite, or ilmenite), goethite, and dark coatings (carbon?) on mineral grains. Trace minerals may include rutile, zoisite, biotite, leucoxene, and/or zircon. Valentino (1994, p. 88) describes the Peach Bottom as containing up to 90% of very fine grained muscovite and sericite, with accessory quartz, chlorite, chloritoid, and ilmenite. X-ray diffraction measurements (J.H. Barnes, Pennsylvania Geological Survey, 2006, personal commun.) of 17 slate samples collected along the railroad over the 60 m interval adjacent to the Cardiff Quartzite (including both slate and the “schist”) yields median values for: quartz—25% (3%–43%); muscovite—30% (10%– 55%); chlorite—20% (0%–40%); chloritoid—15% (2%–25%); and other—10% (1%–37%). Chemical analyses of four slate channel and/or composite samples presented by Smith (1994, p. 185, his table 13-C) and 16 slate samples summarized by Chiarenzelli and Valentino (2006) indicate that the protolith of the slate had a composition similar to the post-Archean average shale (PAAS) of Taylor and McClennan (1985). The slightly elevated levels of Fe, U, V, and Zn suggest a black shale protolith. The slightly higher Al2O3 values may be in response to the lower SiO2 values, which reflect loss of silica, probably during the deformation that produced the cleavage. The low CaO, MgO, and Na2O could represent intense weathering of the sediment source, or loss during metamorphism. The lithic uniformity of the slate, its phyllitic (fine-grained) character, the absence of larger clastic grains, and the absence of sedimentary tractive features (channels and cross beds) point

to accumulation of mud (the protolith material for the slate) in a quiet subaqueous depositional environment below wave base. This environment would readily preserve organic material in the sediment that led to the present enhanced carbon content (Agron, 1950). Small, silt-size quartz grains may have been part of the original sediment, but the angularity and flattening parallel to cleavage of those remaining suggest that many were largely dissolved during subsequent metamorphism and deformation. The fairly uniform distribution of chloritoid prisms throughout the slate suggests that their growth was coeval with the metamorphism, incorporation material freed by the recrystallizing of the original clays. They occur in short, stubby to long prisms, with no preferred orientation. Prisms at moderate to large angle to the aligned phyllosilicates (the cleavage) often have quartz or chlorite “beards” along their sides, extending in the cleavage direction. In addition, the phyllosilicate minerals tend to bend around the end of these prisms, suggesting either growth of the prisms against the phyllosilicates, or compaction of the phyllosilicates around the prisms. What is apparent is that this geometric arrangement would not persist where significant slip has occurred parallel to the cleavage. Such slip would have rotated the more rigid chloritoid prisms into greater parallelism with the cleavage. The Peach Bottom Slate has long been considered to lie in the core of a regional syncline (Knopf and Jonas, 1923, 1929; Behre, 1933; Agron, 1950; Freedman et al., 1964; Wise, 1970). The persistent southeast dip of bedding (and primary foliation, S1) across the Peach Bottom area described in the Geologic Setting above invalidates this view. However, the slate does lie in the core of a fold, an anticline, surrounded by the overlying Cardiff Quartzite. The absence of discernable bedding throughout the slate obscures this anticlinal structure along most of the belt, but the relations at the southwest end of the slate belt at Pylesville, Maryland (Fig. 10) clearly demonstrate its anticlinal nature. The Cardiff Quartzite underlies the northeastward rising slope east of Broad Creek, with the slate underlying the crest of the ridge farther east. The quartzite outcrop extends northeastward on both flanks of ridge, as one would expect in a gently plunging fold. The quartzite also lies at the foot of the steep slope on the west side of Broad Creek, and is overlain by silty phyllites (Fig. 11) similar to those of the Peters Creek Formation just south of the Cardiff along the east side of the Susquehanna River. If this fold were an eastward plunging syncline, then the quartzite would underlie most of the hill west of Broad Creek. The presence of the phyllite in most of this hill indicates the local structure (within the Delta Duplex) is a gently southwestward plunging anticline. Much of the slate exhibits crenulations on the cleavage surfaces. Each crenulation is a parallel-sided (more or less) zone of microfolding (rather kink-like) in which the phyllosilicate minerals are rotated from 30° to more than 90° from their attitude in the adjacent microlithons (the unrotated rock between the crenulation zones). The zones lay at large angle (generally more than 60°) to the main cleavage orientation. The similarity of the

Peach Bottom area, Pennsylvania-Maryland Piedmont phyllosilicates within the zones to those in the microlithons indicates that the crenulations developed after the metamorphism. Interestingly, virtually no quartz exists within the zones, even where angular quartz exists in the adjacent microlithons. Apparently the high strain within the crenulation zones completed the dissolution of the quartz. The subvertical S2 cleavage pervades much of the Peach Bottom Slate, and is dominant in the many of the slate quarries within the belt. Although Valentino (1994) cites this fabric as evidence of regional (and horizontally directed) dextral shear, no horizontally oriented structures (slickensides and/or mullions) are present in the slate (nor elsewhere within the Peach Bottom area) that support such a strain. The moderately east-plunging (and gently southwest-plunging) crenulations (minute folds) on S1 that are so prevalent throughout the Peach Bottom area are kinematically inconsistent with a horizontal shear deformation. We suspect that lithology is the primary factor in the dominance of the S2 cleavage in the slate. Throughout the rest of Peach Bot-

19

tom area, this cleavage is often apparent in the phyllites but is not mesoscopically visible in the more quartzose beds. We attribute the S2 fabric as an axial plane cleavage developed in conjunction with the growth of the Tucquan anticline, which would explain its widespread occurrence. It is not evidence of local shear zones. Cathodoluminescence measurements of fluid inclusions across the Peach Bottom area (Lynn, 2009) demonstrate no systemic variation as a function of location or lithology. Proceed northward along track to pole 442. We will pass the Gorsuch Open Pit Quarry that extends approximately from pole 438 to pole 441. It is a large quarry extending up the slope on several levels. We will not enter the quarry on this trip because: of lack of time; parts of it are difficult to walk through; and the exposures in it, some of which are very good, add little to what we have already seen. Slate exposures extend by the track from approximately pole 440/50 m to pole 442, but we will not stop to examine these because the more interesting rocks lie north of pole 442.

A’ pcp

pb

N cd

EXPLANATION

A

pcp

pcp

Peters Creek Fm Puddle Duck mbr

pb

Peach Bottom Fm

cd

Cardiff Fm

Contact 0

200

400

600

800

meters

Figure 10. Map of the Pylesville area showing SSW-plunging anticlinal nose of Peach Bottom Slate overlain by Cardiff Quartzite. Contact between Cardiff Quartzite and Peters Creek Formation is interpreted as a fault. See Stop 2D text for discussion.

20

Faill and Smith

Stop 2E. The Peach Bottom Station: Sykesville Formation, Bald Friar Metabasalt, and Ultramafites The several different lithologies exposed between pole 442 and 444 are similar to or identical to those in the Sykesville Formation mélange between the Baltimore Mafic Complex and the Peters Creek Formation (e.g., Stop 1). This Sykesville “north” body of rock is also a mélange, and is the other principal component of (along with the slate) the Delta Duplex (Fig. 9). The Sykesville “north” body is bounded by faults, against the Scott Creek Formation on the northwest side, and against the Peach Bottom Slate on the southeast side. This body also includes the large (100 m × 1000 m) lens of serpentinite that extends westward from the north side of Delta and Cardiff Boroughs. Along the Norfolk and Southern railroad, a 9.5-m-thick dolomite-talc body (including, 65 m up the slope above track level, ~2 m of talc-magnesite schist containing 1500 ppm each of Cr and Ni, on its northwest side) lies in the 10 m interval north of pole 442. The large amount of chromium (1800 ppm Cr) and nickel (1200 ppm Ni) in this carbonate rock indicates that it was originally an ultramafite that was metasomatized by fluids rich in Ca and CO2. The absence of a thermal aureole indicates that it was not intruded, but was a sedimentary fragment within the Sykesville mélange, which itself was tectonically emplaced as part of Delta Duplex. Metasediments, mostly laminated quartz-muscovite-chlorite schists with numerous vein quartz stringers, lay along the next ~40 m. The foliation, averaging 036–62, trends more northerly than the regional 060 trend, possibly a result of the late, D3 deformation. The foliation is not planar, but undulose, with relief of a few centimeters over distances of tens of decimeters. Surfaces against which laminae are truncated are probably troughs, or possibly wedge faults. A thin (~30 cm) fragment of the Bald Friar Metabasalt is present at pole 442/50 m, with a very thin (7 cm) sliver of carbonate-talc metasomatized ultramafite containing 1000 ppm Cr and 880 ppm Ni. The geochemistry of the two larger

Bald Friar Metabasalt fragments here on the NW side of the slate is nearly identical to that for samples from the type locality, Bald Friar, Maryland (on the east side of the Susquehanna River just NW of the Baltimore Mafic Complex, Cecil County, Maryland); just NW of the slate on the west side of the Susquehanna River in York County, Pennsylvania; in Chester County, Pennsylvania; and at the Liberty Reservoir, Carroll County, Maryland. This geochemical identity indicates that they were all derived from the same magma (Smith, 2006; see also Table 2). This correspondence, along with the other lithologies and the mélange character, justifies designating this body as a northern exposure of the Sykesville Formation. Metasediments (again, mostly quartz-muscovite-chlorite schists) continue northward for ~22 m, to pole 443/~50 m, where a 20-cm-thick “banded iron formation” and a 7-cm-thick carbonate layer are exposed. The 0.5- to 1-cm-thick solid chlorite layers on both sides of the carbonate, and the chloritic schist up to 50 cm beyond, represent a blackwall reaction zone. Just to the north is a 3.2-m-thick, very green laminated metabasalt (a fragment of Bald Friar Metabasalt) locally containing magnetite. The concentration of magnetite may reflect alteration on the sea floor (near a vent?) in which part of the basalt was replaced. The tightly folded (nearly isoclinally) and faulted lower contact of the metabasalt, nearly enclosing some of the underlying schist, was probably formed during emplacement of the metabasalt on soft sediments. A strong, widely spaced lineation (almost rodding) in the metabasalt plunges northeast at 33°. A finer lineation (crenulation) plunges northeast at 54°, and is an intersection lineation of S1 with S2. Two other bodies of Bald Friar Metabasalt and steatized ultramafites are present on the slope above the railroad track. A large gully lies just north of these exposures, through which the northern bounding fault of the Delta Duplex passes. The outcrop on the other side of the gully exposes the phyllites and schist of the Whitaker Member of the Scott Creek Formation. Return southward along track to the vehicles.

PEACH BOTTOM ANTICLINE -longitudinal section at Pylesville

600

Pylesville

Slate Ridge

pb

400 300

200

pcp

200

100

cd

pb

100 vertical exaggeration 2 : 1

0 0

200

400

EXPLANATION pcp pb

Peters Creek Fm Puddle Duck mbr Peach Bottom Fm

cd

Cardiff Fm Contact

0 600

800 meters

meters

feet

500

ENE Broad Creek

Ridge Road

WSW 700

Figure 11. Longitudinal cross section of Slate Ridge near Pylesville. If the structure were an ENE-plunging syncline, the hill west of Broad Creek would consist of Cardiff Quartzite rather than Peters Creek Formation.

Peach Bottom area, Pennsylvania-Maryland Piedmont

21

TABLE 2. SELECTED TRACE ELEMENTS IN TWO SAMPLES OF THE BALD FRIAR METABASALT FROM THE PRESENTLY KNOWN ENDS OF THE STRIKE BELT Name and county TiO2 Zr Hf Nb Ta Th Y La Ce LRMRA6, Carroll 1.35 91 2.1 2 <0.1 0.2 35 2.5 11 BRANDYSE, Chester 1.35 92 1.8 <2 <0.1 <0.1 31 2.5 9 Note: The two rows of analyses of Bald Friar Metabasalt are from samples collected from a strike length of 143 km. For practical purposes, we would consider being from the “ same” magma.

Retrace steps along railroad, across the railroad bridge over Peters Creek, to the vehicles. Mileage Incremental Cumulative 0.0

59.6

0.3

59.9

1.3

61.2

0.9

62.1

0.1

62.2

0.9

63.1

0.1

63.2

Board vehicles, return across road bridge over Peters Creek. Turn LEFT onto SLATE HILL ROAD, go up long hill. Intersection with BALD EAGLE ROAD on Left, continue straight ahead on SLATE HILL ROAD. Stop sign, turn LEFT onto CHERRY HILL ROAD. Yield sign, bear LEFT, proceeding down hollow on BENTON. HOLLOW ROAD. Bear RIGHT onto BALD EAGLE ROAD. Park at pull-off area on left side of road.

Stop 3. Benton: Scott Creek Formation, Coyne Lock Member The Scott Creek Formation The Scott Creek Formation is interpreted to be a more distal part of the deltaic and continental margin deposits that spread northwestward (current coordinates) from the Brandywine and Baltimore microcontinents offshore of Laurentia (see “The Octoraro basin” above). Being farther from the source, the Scott Creek sediments naturally are overall less sandy, in volume and grain size, than those of the more proximal Peters Creek Formation. However, the Scott Creek is not without sand—it just has less of it, deposited in different, more subaqueous environments. It also exhibits a distinctive fine-scale lamination not common in the Peters Creek, not only in the sandy and silty beds, but also in the more argillaceous beds. The subdivision of the Scott Creek Formation into three members is based primarily on the amount of sand and sandstones. The lower member, the Bryansville, and the upper member, the Whitaker tend to be more argillaceous than the medial Coyne Lock Member. However, none of the three is uniform in character—all three exhibit numerous intervals, from ~1 to >10 m thick, of contrasting lithologies and beds forms. In addition, most such intervals can be found in each member, but comprising dif-

ferent proportions of the entire member. These contrasting lithic intervals probably represent contrasting depositional environments that shifted repeatedly on the delta and continental margin in response to migrating distribution systems. One other lithology distinguishes the Scott Creek Formation from the Peters Creek Formation—intervals of medium dark to very dark gray silty phyllites and phyllitic siltstones. Three major intervals of ~10± meters thickness or dark gray phyllitic rocks are present in the Scott Creek, plus some thinner intervals dispersed throughout the formation. They occur at the base of each member, or perhaps it better said that the member bases were placed in these dark gray intervals. Regardless, they apparently represent periods of quiet, low-energy deposition, possibly with accumulation of substantial carbon. The Coyne Lock Member The Coyne Lock Member of the Scott Creek Formation consists primarily of laminated quartzose schist. The beds are light to medium gray, thin to medium bedded, very fine to finegrained, quartz-muscovite-chlorite schists. The rocks are variably laminated, in which the laminations consist of alternating medium-light gray quartz-rich (80%) laminae and very dark gray phyllitic laminae. Generally, the lamination thickness varies with grain size, from ~1 mm to 3 mm. Some of the quartzose laminae are lensoid, suggesting ripples and microtroughs, often providing a flaser-like aspect. Concentrations of heavy minerals (zircon, monazite, xenotime, rutile, and titanite) at several horizons throughout the Coyne Lock Member are sedimentary lag deposits. Rather profuse magnetite-rich layers also occur in several intervals in this member. Exposures of the Coyne Lock Member of the Scott Creek Formation Rather continuous exposures of the upper part of the Coyne Lock Member lie along the Norfolk and Southern railroad along the Susquehanna River. However, the time required to reach the railroad and examine the outcrops in any sort of detail greatly exceeds the available time. Fortunately, three outcrops along Bald Eagle Road, just north of the parking area, provides exposure of the sandier lithologies present in this member, and of the dark gray lithology. The first outcrop is just north of the (unnamed) creek, on the east side of the road. Medium to thick bedded quartzose sandstone interval of 4–6 m thickness alternates with laminated, rather quartzose schist, in intervals from 20 to 25 m. As

22

Faill and Smith

we have seen through much of this day, the bedding S0, which parallels the primary foliation, S1, dip moderately to the southeast, at 048–60. The second outcrop displays mostly quartzose, laminated schist with a few near quartzite beds. Imbedded in these rocks is one <1-m-thick interval of very dark gray quartzose schist. Again, the bedding S0 = foliation S1 dips moderately to the southeast, 049/61. A faint spaced cleavage appears in the more phyllitic beds, dipping steeply (~80°) to the southeast. The third outcrop is a bit different in that it lacks quartzite beds, and not much schist is present. The rock is finely laminated (~1 mm) with alternating sucrosic quartz and very dark gray to black phyllite. At the north end of the outcrop is a 2-m-thick interval containing abundant intruded vein quartz masses. Return to vehicles and proceed north on Bald Eagle Road.

63.2

0.7

63.9

64.7

1.5 1.2

66.2 67.4

Turn RIGHT onto SUSQUEHANNOCK DRIVE. Turn LEFT onto STATE PARK ROAD. Park in parking area in SUSQUEHANNOCK STATE PARK.

(Overlook above Susquehanna River.) 0.0 67.4 Board vehicles, return to SUSQUEHANNOCK DRIVE. 1.2 68.6 Cross SUSQUEHANNOCK DRIVE onto PARK DRIVE. 0.2 68.8 Turn RIGHT (south) onto FERNGLEN ROAD. 1.0 69.8 Stop just north of small bridge over creek; park by side of road.

Mileage Incremental Cumulative 0.0

0.8

Stop 4. Fishing Creek: Scott Creek Formation, Fishing Creek Metabasalt (in Bryansville Member)

Return to vehicles. Continue north on BALD EAGLE ROAD. Turn LEFT onto HARMONY RIDGE ROAD.

The Fishing Creek Metabasalt (Smith and Barnes, 1994, 2004) is an important part of the Peach Bottom area story. The metabasalt lies near the top of the Bryansville Member of the

Scott Creek Formation Bryansville Member M

s

Drumore “tectonite” ” zone

STOP 4 Scott Creek Formation Coyne Lock Member

STOP 5

Fishing Creek Metabasalt Bryansville / Coyne Lock contact

0

1

2 km

Limits of Drumore “tectonite” zone

Figure 12. Topographic and geologic map of the Fishing Creek Metabasalt and the Drumore “tectonite” zone (Valentino, 1994). See Stop 4 and Stop 5 texts for discussions. Modified from Smith and Barnes (1994).

Population (N) TiO2 Zr Hf Nb Ta Th Ni V Y La Ce Reference Eocene basalt, West Virginia and Virginia (7) 2.28 151 3.8 34 1.9 2.8 68 252 26 32.4 60 R.C. Smith unpublished Rossville Diabase type locality (1) 0.88 77 1.3 7 0.2 1.0 70 254 24 5.2 12 Smith and Barnes (2004) York Haven Diabase type locality (1) 1.12 109 2.1 11 0.3 1.6 73 220 24 8.7 18 Smith and Barnes (2004) Quarryville Diabase type locality (1) 0.41 59 0.9 4 <0.1 0.9 320 160 20 4 9 Smith and Barnes (2004) Sword Mountain Olivine Melilitite (8) 4.06 307 4.6 106 7.8 8.4 335 297 22 72 134 Smith et al. (2004) Catoctin Metadiabase dikes (23) 2.78 210 4.3 19 0.9 1.2 40 330 36 17 38 Smith and Barnes (2004) Accomac Metabasalt (5) 2.72 1 . 90 4 .6 19 1 .0 1. 3 60 2 44 32 16 37 Smith and Barnes (2004) White Clay Creek Amphibolite (21) 2.64 150 4.1 14 0.9 1.6 66 393 30 15.4 37 Smith and Barnes (2004) Catoctin Metabasalt* (48) 2.35 157 3.6 12 0.6 0. 7 80 32 9 36 11.5 29 Smith and Barnes (2004) Tunnel Mine Metadiabase (5) 3.70 355 7.6 54 3.4 5.9 191 217 49 62 123 Smith (2003) plus 2 unpublished Williams Quarry metadiabase informal (3) 3.51 295 5.4 91 5.4 6.6 102 189 44 67 132 Smith (2003) Unnamed near Mount Rogers Formation (1) 3.04 185 4.6 15 1.5 1.3 67 281 32 20 40 Smith (2003) Note: Some may justifiably wonder how geochemically similar are basalts from various extension events. This table provides some idea of the geochemical distinctiveness of “different” rift-related basalts in the Pennsylvania, Maryland, West Virginia, and Virginia area, approximately from youngest to oldest. Also, consider examining table 3 of Smith and Barnes (2004) for an idea of how related basalts in the Catoctin sensu lato group vary as rifting and extension evolved into drifting. TiO2 in percent. Others in ppm. *Catoctin Metabasalt of South Mountain Anticlinorium NW of Tunnel Hill–Jacks Mountain fault system, Pennsylvania.

Scott Creek Formation (Table 1); it extends from the Susquehanna River (just north of Drumore) northeastward for 4 km (Fig. 12). Its geochemistry indicates that it is related [slightly younger, offshore equivalent] to the Catoctin Metabasalts that were extruded on the Laurentian continental margin late in the Neoproterozoic (Table 3). This age has significant implications: (1) Because the Fishing Creek Metabasalt is a stratabound, extrusive flow that lies in the upper part of the Bryansville Member of the Scott Creek Formation, it suggests that the metasediments below the Fishing Creek Metabasalt are Neoproterozoic in age, and the overlying metasediments are younger than 564 Ma (mostly Paleozoic). (2) Geochemically, the Fishing Creek Metabasalt is an extremely close equivalent to the Sams Creek Metabasalt flows at the type locality and at the de facto principal reference section in Maryland (Table 4). This equivalence suggests that the two were extruded at the same time, and thus they provide a basis for correlating across Tucquan anticline, from the southeast limb (Fishing Creek Metabasalt) to the northwest limb (Sams Creek Metabasalt). Because the Sams Creek extends into some of the metabasalt localities in southern York County, correlation across the Tucquan anticline is possible to that area as well. The Fishing Creek Metabasalt consists of greenschistfacies, epidote-chlorite-actinolite-albite metabasalt containing some calcite. It also commonly contains accessory titanite and 1-mm octahedral magnetite, and locally biotite. Trace specular hematite occurs at a few localities. The Fishing Creek Metabasalt is relatively uniform chemically along strike, including an apparent intrusive phase at the northeast end. It exhibits geochemical features transitional from within-plate basalts to ocean-floor basalt. When compared to 48 samples of Catoctin Metabasalt from the South Mountain anticlinorium NW of the Tunnel Hill–Jacks Mountain fault system, the Fishing Creek Metabasalt and Sams Creek Metabasalt are both slightly more oceanic. This suggests a slight shift from a rifting environment (with remobilization of deep Laurentian crust as metarhyolite) for the Catoctin to slightly more drifting environment for the Fishing Creek and Sams Creek, which lack known associated metarhyolite. Thus, the Sams Creek and Fishing Creek are likely slightly younger than the Catoctin and, likely more outboard than the Catoctin Metabasalt and Catoctin Metarhyolite of the South Mountain anticlinorium (Smith and Barnes, 2004). Uncompleted work by Ken A. Foland, Ohio State University, and Robert C. Smith, II established a linear plot for 143Nd/144Nd versus 147Sm/144Nd for the Catoctin Metabasalt sensu lato. The relative positions on this Sm-Nd isochron largely confirmed the groupings, relative ages, and the rift to drift concept for the 12 populations within that overall group proposed by Smith and Barnes (2004, their table 3). The Fishing Creek Metabasalt is greenish-gray, silty, and laminated. The rock color is lighter (medium light to medium dark greenish gray) in the siltier beds, and darker (medium dark to dark greenish gray) in the more phyllitic beds, probably reflecting larger amounts of chlorite. Knots and strings of blocky, fractured epidote are numerous, ranging in size up to 4 cm across,

TABLE 3. FISHING CREEK METABASALT INSERT

Peach Bottom area, Pennsylvania-Maryland Piedmont

23

24

Faill and Smith TABLE 4. MEDIANS OF SELECTED TRACE ELEMENTS IN THE SAMS CREEK METABASALT TYPE AND DE FACTO PRINICPAL REFERENCE SECTION, MARYLAND; THREE SIMILAR SAMPLES FROM SOUTHERN YORK COUNTY, PENNSYLVANIA; AND THE FISHING CREEK METABASALT OF THE BRYANSVILLE MEMBER, SCOTT CREEK FORMATION, LANCASTER COUNTY, PENNSYLVANIA N TiO2 Zr Hf Nb Ta Th Ni V Y La Sams Creek Metabasalt, York County 3 2.02 160 3.5 21 1.1 1.6 160 328 38 16.5 Fishing Creek Metabasalt Member of Scott Creek 7 1.94 154 3.6 21 1.1 1.7 96 223 30 14.8 Formation, Lancaster County Sams Creek Metabasalt, Maryland, type and 6 1.88 146 3.0 22 1.0 1.4 126 306 32 14.2 reference section Note: N—number of samples analyzed.

and 20 cm or more in length. Magnetite and/or specular hematite, and calcite are accessory minerals. Bed thickness ranges from thin to thick, generally thicker bedded in the lower part, and very thin to medium bedded in the upper part. Bedding is quite parallel, but some beds exhibit lateral thickness changes of 10%–20% (a few 30%–40%) over a lateral distance of 2–3 m. The metabasalt is generally flat bedded, but some shallow troughs are present, especially in the lower part. A coarse lamination (2–3 mm thick) pervades the beds, which may have been enhanced by the subsequent burial metamorphism. A number of strings (very thin [<6 mm] layers) of vein quartz parallel the layer partings. Rough surfaces (Fe-stained brown) parallel to bedding and foliation may have been chill zones of glass that were replaced by carbonate by the action of deuteric hot fluids within a few thousand years of extrusion. Calcite is a common filler of amygdules (up to 5 mm in size). No obvious erosional cutouts are evident. No definitive original flow structures, pahoehoe toes, horizontal feeder tubes, or pillows are discernable at most outcrops. Such primary features that were present may have been obliterated by the strongly developed primary foliation. The flow was probably short-lived because no quartz-muscovite-chlorite schists were observed within the metabasalt sequence. The top of the metabasalt is overlain by a medium gray, very fissile muscovite-chlorite-quartz phyllite. Dissolved vugs indicate that the metabasalt did have considerable calcite and/ or dolomite, produced by deuteric alteration from fluids (meteoric fluids (mostly) and magmatic fluids), which entered the flow early (while still hot). The Fishing Creek Metabasalt is pervaded by a welldeveloped foliation, S1, which parallels the layering and laminations. This foliation is quite planar and flat, only gently and broadly undulose on a 2–4 m scale. However, the foliation does bend (refract) around epidote and quartz knots, probably in response to the growth of the knots after S1 developed. The foliation is so well developed that the metabasalt frequently tends to split into thin (1–8 mm) slabs on weathered surfaces. This foliation dips moderately to the southeast, 063–56. A second foliation (S2), not as strongly developed, trends more northerly and steeply (034–71) than the primary foliation. A third spaced foliation (072–78) has a strike similar to the primary foliation, but dips more steeply to the southeast. A few of these partings seem to have been open,

Ce 38 33 31

being once filled with carbonate. One of these S3 surfaces passes through several layers with down-on-the-south displacement of 0.5–1 cm. This movement may be as late as Jurassic, as part of the deformation of the early Mesozoic Gettysburg basin to the north. A prominent lineation pervades much of the Fishing Creek Metabasalt outcrop, here at Fernglen Road, and farther east along the metabasalt outcrop belt. It appears on the primary foliation (S1) as a crenulation or even rodding. It plunges moderately to the southwest (192–55), the direction of the intersection of S1 and S2, L S1 × S2. Here at Fernglen Road it is represented primarily as a crenulation, but farther east it appears as rods, pencils, or “cigars” up to 2–3 cm across, and ~20 cm long (down the S1 × S2 intersection). In the metabasalt, they consist mostly of crystalline epidote, sometimes with calcite. This SW-lineation is also present in the schists above and below the metabasalt. Crenulation lineations with other plunges are fairly common though less prominent; they are almost always on the primary foliation, S1. The subhorizontal ones are probably related to the S3 foliation; others may reflect local undulations in the foliation surfaces, which result in divergent plunges. A few foliation surfaces display two or more crenulations (usually over small areas [<20–30 cm square]), in response to multiple movements along the surface. Small folds are locally present, with fold axes plunging parallel to the crenulation plunge. It is worth noting that subhorizontal shear indicators are not present in this outcrop, suggesting that there is no significant strike-slip component to the deformation here. Keep in mind that this outcrop lays within the Drumore tectonite zone of Valentino (1994). The Fishing Creek Metabasalt has been traced for 4.5 km from the Susquehanna River northeastward across several entrenched meanders of Fishing Creek, underlying saddles in the intervening low hills. It has not yet been found on the west side of the river. At its northeast end, the presence of coarse-grained metabasalt (some porphyritic) and even gabbroic float (no outcrop) suggests proximity to the source vent. The absence of any metabasalt farther northeastward suggests a southwestward paleoslope at the time of extrusion. The thickness ranges from 6 to 37 m; the greater thicknesses here and immediately to the northeast suggest the flow may have filled a topographic low. Along Fernglen Road (Stop 4), the base of the metabasalt is ~50 m north of the small bridge over the unnamed creek; the top is just

Peach Bottom area, Pennsylvania-Maryland Piedmont north of the muscovite-chlorite schist outcrop immediately north of the bridge. A paragonite-bearing sequence of phyllites lies above the Fishing Creek Metabasalt, exposed along Fernglen Road, but is presently mostly seen only as float some 100–150 m south of the small bridge over the creek. The sequence is a dark gray phyllite, containing muscovite, paragonite, and some quartz. Low-grade metamorphism (300–350 °C, chlorite-grade, greenschist facies) of igneous rocks can produce different minerals, depending on the abundance of potassium and/or sodium. For example, paragonite (or albite) would occur if sufficient Na were abundant. On the other hand, muscovite or K-feldspar (and not chlorite or garnet) would develop if sufficient K were present. In the absence of sufficient Na or K relative to aluminum, pyrophyllite will appear, which at higher metamorphic grade would become kyanite, andalusite, or sillimanite. The presence of paragonite here suggests that sea water was readily available during metamorphism. In contrast, the profusion of the anhydrous albite in the upper part of the underlying Octoraro Formation would suggest a more restricted access to seawater. This occurrence of paragonite-bearing phyllite may have a more regional significance. The Ijamsville Formation to the southwest in Maryland, which is interpreted to be a tuffaceous deposit (Edwards, 1986, 1988) overlies paragonite-bearing rocks. Correlation between these two paragonite-bearing sequences would greatly aid regional interpretations of basin development and tectonism. It cannot presently be ruled out that this set of paragonite-bearing rocks is reworked volcaniclastic detritus from the Catoctin Metarhyolite. Mileage Incremental Cumulative 0.0

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Board vehicles, continue SOUTH on FERNGLEN ROAD. Bear RIGHT onto FISHING CREEK ROAD. Turn LEFT, proceed on SUSQUEHANNOCK DRIVE to river. Park at far end of gravel area to left of railroad.

Stop 5. Drumore: Scott Creek Formation, Coyne Lock Member, and the Drumore Tectonite Zone Drumore consists of a cluster of houses and summer cottages, many of them lined up along the top of a low ridge overlooking the Susquehanna River. The large parking area and the (concrete floor of the now dismantled) substation were presumably built on artificial fill between the track and low ridge. A grade crossing over the track lies just south of a very tall, unnumbered steel pole that stands ~25 m north of pole 495, providing access to outcrops north and south of Drumore. The oldest beds in the Coyne Lock Member of the Scott Creek Formation are displayed in the exposures along the low

25

ridge on the east side of the parking area and substation, extending from pole 497/~35 m to pole 494. Black phyllite is exposed at the northern end of the parking area (pole 497/~35 m), containing numerous small to medium size (5–10 cm) lenses and irregular masses of vein quartz. These rocks grade upsection (southward) into medium gray silty schists and laminated schists with many fewer vein quartz masses. From pole 497 southward for 30+ m, the rock is dark gray silty phyllite and very dark gray phyllite, containing numerous medium to large lenses and irregular masses (5–40 cm) of vein quartz. Throughout this interval, the foliation is quite planar (except around the quartz masses), dipping moderately steeply to the southeast, at an average of 046–63. Crenulations on the foliation plunge moderately steeply to the east (091–59) and southwest (204–49). Southward (up section), the rocks grade into medium gray phyllitic schist, silty schist, and laminated schist that continue very nearly to pole 495. Medium dark gray phyllitic schists dominate farther south to end of outcrop. The foliation in this part is planar as well (except in proximity to the numerous vein quartz masses and lenses), dipping moderately steeply to the southeast, at an average of 052–65. Crenulations on the foliation similarly plunge moderately steeply to the east (109–59) and less steeply to the southwest (214–30). Wedging (thrust faulting at low angle to bedding) is suggested by occasional cutoffs and truncations of the foliation. A few narrow (2–3 cm) bands kink the foliation, apparently being late structures. The rocks at this stop are characteristic of the finer grained parts of the Scott Creek Formation that form the basal parts of each member. David Valentino et al. (1994) have argued that these rocks are part of a regional shear zone, which he named the Drumore Tectonite Zone for these exposures. We disagree with his interpretation for a number of reasons: (1) Dark gray phyllites interspersed with lighter gray phyllitic schists are not necessarily phyllonites. Mudstone and shale protoliths would naturally transform into wellfoliated phyllites under lower greenschist-grade burial metamorphism. The presence of carbon would produce their dark gray color. This is supported by the presence of similarly dark phyllites at other stratigraphic levels in the Scott Creek Formation (for example, at Stop 4), none of which are in the Drumore Zone or in any other regional shear zone. (2) The primary foliation dips moderately steeply to the southeast throughout this exposure, at an angle identical to that elsewhere along this stretch of the Susquehanna River. The later, steeply southeast-dipping foliation, S2, is rarely present. (3) Crenulations and folds in the Drumore Zone appear no different than elsewhere across the area—their geometry exhibits no evidence of subsequent distortion or shearing. No subhorizontal structures (slickensides and/ or mullions) are present that would characterize a subvertical shear zone. Either the shearing occurred before the folding and crenulations, or no shearing occurred.

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(4) The Fishing Creek Metabasalt is included within the Drumore Shear Zone (Valentino, 1994, his figures 27 and 28, p. 86–87), but the metabasalt shows no evidence of having been deformed by shear. No structures are present (such as folds with subvertical fold axes, pronounced subhorizontal smear lineations, or slickensides that would indicate a subhorizontal shear deformation). Cross the railroad track at the grade crossing and walk southward along the gravel road. The first outcrop along the track extends from pole 491/ 35 m to 491/00 m. The rocks appear quite homogeneous with little indication of bedding. The rocks vary from medium gray silty phyllitic schist (some laminated) to dark gray phyllitic schist, very similar to those by the parking area. Apparently, the dark gray basal interval of the Coyne Lock Member extends this far south. Foliation is quite planar (except near the numerous vein quartz masses), though undulose, yielding average dips of 056–56 and 045–69. The two clusters in this data suggest the presence of two different foliations, the primary bed-parallel foliation (S1), and the later, not as well developed, Tucquan axial plane foliation (S2). Crenulations on the foliation plunge subhorizontally to the northeast (044–06), moderately steeply to the east (097–55), and moderately to the southwest (206–29). Wedging may also be present. The next outcrop to the south, from pole 488/30 m to 487/ 30 m, is badly covered with soot, thereby obscuring much of the lithology and structure. At the northern end, the laminated chlorite-quartz-muscovite schist is folded (kink-like) on a centimeter and decimeter scale with axial surfaces dipping moderately steeply to the southeast, spaced at 14–25 cm. Overlying these deformed schists is a thick interval of very dark gray chloritemuscovite phyllitic schists and dark gray phyllitic schists, locally containing large lenses of white vein quartz. The southern portion contains quartz-chlorite-muscovite schist and quartz-muscovitechlorite schist with thick laminae parallel to the primary foliation. The gradual increase from dark phyllites to quartzose schists suggests a possible upward-coarsening cycle. The foliation at the northern end has the usual moderate dip to the southeast (046– 61), but southward the dip is a more northerly trend, 017–68, not unlike the S3 attitude seen at Stop 2D. The next outcrop to the south, from poles 486/27 m to 486/15 m, exposes medium dark greenish-gray, laminated quartz-chlorite-muscovite silty schist and dark greenish-gray chlorite-muscovite-quartz phyllitic schist. Two foliations are present: the dominant one dips steeply to the southeast (044–80, S3), the other less steeply to the southeast (S1). The laminae in the schists parallel the steeper foliation. The intersection of these foliations has produced a coarse lineation on S1 that plunges moderately to the southwest. The intermittent outcrop farther south, from 485/10 m to 484/20 m, exposes a very laminated rock and many folds. Bedding is not readily discerned, except for a pyritiferous layer and a few phyllitic beds. The rock is chloritic, being mostly a quartzchlorite-muscovite schist and laminated schist and subordinate

quartz-muscovite-chlorite schist. Many very thin lenses of sucrosic vein quartz parallel the laminae. Overall, the foliation is very steeply southeast dipping, suggesting that it is either the Tucquan axial plane cleavage (S2) or the more northerly trending S3. The axial surfaces of the folds parallel the very steep foliation. The next outcrop to the south, from 483/35 to 493 10 m, is even more deformed, particularly on the south end. Bedding is not obvious, but laminations appear to be subhorizontal, and profusely folded. At the north end, bedding dips moderately to the southeast, at 046–57. The dominant foliation is 046–78. We have seen at this Stop 5 two very different rock assemblages and structures. The northern part, at the parking area and the first outcrop along the railroad, the rocks were phyllitic, often with numerous lenses and masses of vein quartz. A single structure pervaded these rocks, the regionally extensive, moderately steeply southeast dipping foliation, S1, which parallels bedding. Southward from there, the rocks have become more schistose and laminated, and extensively folded. The S1 foliation has given way to, or overprinted by, the steeply southeast dipping Tucquan axial plane foliation, S2. Return to vehicles. Mileage Incremental Cumulative 0.0

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End of field trip.

Return to and board vehicles. Proceed north on SUSQUEHANNOCK DRIVE to HARMONY RIDGE DRIVE. Turn RIGHT onto HARMONY RIDGE DRIVE, proceed to FURNISS ROAD. Turn RIGHT onto FURNISS ROAD, proceed to U.S. ROUTE 222 in WAKEFIELD. Turn RIGHT onto U.S.-222, follow U.S.-222 to U.S. ROUTE 1. Turn RIGHT onto U.S. ROUTE 1, follow U.S. 1 S. West end Conowingo Dam. Turn SLIGHT LEFT to take MD-24 S ramp to BEL AIR. Turn LEFT onto MD-24. Merge onto I-95 S toward BALTIMORE (toll). Merge onto I-395 N (EXIT 53) toward DOWNTOWN. I-395 N becomes S HOWARD STREET. Turn RIGHT onto W BALTIMORE STREET. Turn LEFT onto N HOWARD STREET. Turn LEFT onto W FAYETTE STREET. Arrive at 101 WEST FAYETTE STREET (Sheraton Baltimore City Center Hotel).

Peach Bottom area, Pennsylvania-Maryland Piedmont APPENDIX By Rodger T. Faill The Appendix includes detailed descriptions of stratigraphic units that are herein defined, or extended into the Peach Bottom area (PBA). They include the Sykesville Formation (extended into the PBA), the Peters Creek Formation (redefined and subdivided into two new members), and the Scott Creek Formation (newly defined along with its three members). See Stop 2 for descriptions of the Peach Bottom Slate and the Cardiff Quartzite. Sykesville Formation (900–3000+ m Thick) The lower part of the Sykesville Formation at Stop 1 is rather quartzose, consisting primarily of very thick-bedded metasandstones and metagraywackes, with subordinate amounts of thinner bedded schists. A few of the sandstones contain quartz pebbles, but pebbles and cobbles of graywacke and quartz are not common in this lower part of the Sykesville. Some of the sandstones are quite micaceous; others contain pseudo-flaser structures and very thin (<1 mm) black phyllitic laminae. A few dark gray sandstones are unusually heavy. Bald Friar Metabasalt is present as dismembered flows in the form of locally well-preserved pillows, large boulders (2.5-m-thick blocks), and isolated layers, probably fragments derived from formerly more widespread flows. Although poorly laminated, these rocks appear granular. Petrographically, they are epidote-rich, with subordinate chlorite. Minor minerals include quartz, zoisite, and carbonate. Accessory minerals are titanite, rutile, and zircon. These epidote-chlorite schists may include a volcaniclastic component. Some of them are very contorted as if by soft-sediment folding; others are deeply scoured by overlying sandstone. The most prominent metabasalt, the Bald Friar Metabasalt (Smith and Barnes, 2004; Smith, 2006), is exposed on the east side of the Susquehanna River, near Bald Friar, Maryland. Fragments of this backarc metabasalt are present along strike to the northeast into Chester County, Pennsylvania, and to the southwest into York County, Pennsylvania, and Carroll County, Maryland (Smith, 2006). The latter notes that the Bald Friar Metabasalt is an excellent geochemical match for the Caldwell 1 b basalts of southeastern Quebec (Bédard and Stevenson, 1999). This lower part of the Sykesville Formation also contains several medium (a few meters) to very large bodies of ultramafic rocks (originally peridotite-dunite) that have been variably steatized to serpentinite, magnesite, talc, and the common blackwall minerals. One such body of serpentinite, magnesite, and talc is well exposed in the McGuigan quarries that lie on the northeast side of Bald Friar Hill (Fig. 1), in Maryland, ~400 m up a small tributary on the east side of the Susquehanna River. These minerals are steatized products of an ultramafic (95% forsterite, 1% chromite, 4% other mafic minerals) body (McKague, 1964). The generally medium gray rock that dominates the southern (younger) part of the Sykesville Formation is quartzose metasandstone, coarse-grained, and quite homogeneous, with subordinately more schistose rock; phyllites are absent. This lithology is also present in the Sykesville Formation of Cecil County to the south (Higgins and Conant, 1986). Bedding commonly is absent or ill defined; elsewhere the rock is very thick bedded, and locally trough cross bedded (suggesting a high-energy depositional environment). Quartz pebbles and graywacke cobbles are present in some layers. Some metasandstones are medium-dark to dark gray, containing sucrosic quartz pebbles. A few of the metasandstones are feldspathic, with alkaline feldspars constituting as much as 25% and plagioclase up to 10%. Petrographically, the more massive rock is granular, with large equant alkali feldspars surrounded by a matrix of quartz, muscovite, and chlorite. Epidote and/or zoisite, carbonate, magnetite, and ilmenite are minor minerals.

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Accessory minerals include rutile, tourmaline, apatite, hornblende, and zircon. In the more schistose beds, alternating quartz-alkali-chlorite and muscovite-chlorite-biotite-epidote laminae define the layering; some of the schists are laminated. Petrographically, the schists have quartz, alkali feldspar, muscovite, chlorite, and epidote as major minerals. Minor minerals include magnetite, zoisite, and biotite. Accessory minerals include hornblende, perthite, plagioclase, tourmaline, apatite, titanite, monazite, and zircon. Fragments of the informally named “Conowingo Creek metabasalt [informal],” a forearc basalt having the boninitic affinity (Smith and Barnes, 1994), are also present in the Sykesville Formation. The Sykesville is equivalent to the Morgan Run Formation, a component of the Liberty Complex (Muller et al., 1989), a trench-fill mélange in the Liberty Reservoir area, Carroll County, Maryland, that has been both thrust and folded. Smith (2006) analyzed identical samples of Bald Friar Metabasalt associated with ultramafites in the Morgan Run Formation. The northern edge (“base”) of the Sykesville Formation is interpreted to be a thrust fault (Smith, 1993) because the lithic contrast Sykesville and adjacent Peters Creek Formations. This thrust is herein named the McGuigan thrust for the McGuigan quarries just east of the Susquehanna River, northeast of Bald Friar Hill, that contain talc, magnesite, and serpentinite blocks, which facilitated obduction of the Sykesville and Baltimore Mafic Complex terrane over the Octoraro basin (Peters Creek Formation) metasedimentary rocks. The Sykesville Formation is well bedded in part, but the paucity of good compositional layering in the other parts, the coarse grained of those deposits, and the presence of isolated small clasts (<1 cm) to large blocks (>>20 m) suggest that the Sykesville Formation is a poorly sorted, ophiolitic mélange deposit. As the Baltimore Mafic Complex advanced across the Octoraro Sea, converging toward the Laurentian continental margin, its weight depressed the crust before it, forming an oceanic (subduction) trench before it. Into this trench were dumped fragments of the Baltimore Mafic Complex itself, a large volume of siliciclastic sediment, and the nearly concurrent basalt flows (Bald Friar Metabasalt and the “Conowingo Creek metabasalt”), all mixed to form an ophiolitic mélange, the Sykesville Formation (Muller et al., 1989). pc—Peters Creek Formation (6400 m Thick) The Peters Creek Formation (Knopf and Jonas, 1923) is a metasedimentary sequence of schists, quartzites, and phyllites. It is divided into two members based on sedimentary structures: the lower, fluvial Puddle Duck Member and the upper, deltaic Cooks Landing Member. pcc—Cooks Landing Member (New Name) (3100 m Thick) The Cooks Landing Member of the Peters Creek Formation consists of thin- to medium-bedded quartzose schists and silty phyllites in cyclic alternation. The coarser grained beds are laminated metasandstones to metasiltstones with sharp tops and bottoms. Mineralogically, the rocks are quartz-muscovite-chlorite schists. A very few of the metasandstones are feldspathic, with alkaline feldspars constituting as much as 25% and plagioclase up to 10%. Interstitial carbonate is present (up to 35%) in a very few of the metasandstones. Accessory minerals include magnetite, carbonate, and/or ilmenite, epidote, zoisite, and clinozoisite. Trace amounts of titanium are present in the form of rutile, brookite, and titanite; tourmaline and detrital zircons are also occasionally present. The absence of troughs, scours, cross bedding, and ripples suggest deposition in an aqueous environment below wave base. The laminations in the sand beds suggest strong traction (upper flow regime), as is characteristic of clinoforms in the lower reaches of deltas).

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The contact between these two members is gradational over ~200 m. The absence of a beach facies between them suggests either a fairly quiet shoreline or a rapid transition from fluvial to marine conditions. pcp—Puddle Duck Member (New Name) (3300 m Thick) The Puddle Duck Member of the Peters Creek Formation contains medium- to thick-bedded metasandstones, metasiltstones, and silty phyllites. The lower Puddle Duck Member (2600 m thick) contains numerous upward-fining cycles ranging from 150 to 200 m thickness. The basal 30–60 m of medium- to thick-bedded metasandstones that contain troughs and cross bedding. The quartz-bearing sandstones grade upward into siltstones and mudstones, imparting the appearance of upward-fining cycles characteristic of a fluvial environment in the subaerial part of a delta. Mineralogically, the rocks are quartz-muscovite-chlorite schists. A very few of the metasandstones are feldspathic, with alkaline feldspars constituting as much as 25% and plagioclase up to 10%. Interstitial carbonate is present (up to 35%) in a very few of the metasandstones. Accessory minerals include magnetite, carbonate, and/or epidote, ilmenite, and pyrite. Trace amounts of titanium are present in the form of rutile, brookite, and titanite; tourmaline and detrital zircons are also occasionally present. sc—Scott Creek Formation (New Name) (4700 m Thick) The Scott Creek Formation is herein defined for the exposures along Scott Creek north of Delta and the Peach Bottom Slate Belt northward to Muddy Creek. The members are (from the bottom): Bryansville, Coyne Lock, and Whitaker members. Important reference sections exist along the Port Road (Norfolk and Southern Railway) on the east side of the Susquehanna River and, on the west side of the river, south and west of the Peach Bottom power station and in the parking lot to the north. The Scott Creek Formation is a sedimentary sequence in which the bedding dips at a fairly uniformly steep attitude (55°–75°) across the entire 5.4 km width of the belt. Limited sedimentary data (channels in sandstones) indicates that stratigraphic tops are toward the southeast, suggesting that, in the absence of any duplication by faulting, the present stratigraphic section is ~4900 m thick. Its lower contact with the Octoraro Formation appears to be conformable, even interbedded. Its uppermost beds are in fault contact (Whitaker Station thrust) with the Sykesville Formation ophiolitic mélange on the north side of the Peach Bottom Slate. The Scott Creek Formation consists of schists and silty phyllites interbedded at various scales. The schists contain primarily quartz, muscovite, and chlorite along with a variety of accessory minerals. The finer grained phyllites contain less quartz. Some dark gray phyllites are also present at specific horizons. A metabasalt, the Fishing Creek Metabasalt, occurs in the lower portion of the Bryansville Member. scw—Whitaker Member (New Name) (800+ m Thick) The Whitaker Member of the Scott Creek Formation consists primarily of interbedded thin to thick-bedded medium gray phyllitic schist, medium light gray laminated quartz-chlorite-muscovite schist, and light gray quartzose schist. The alternating schistose sandstone and siltstone beds are thin to very thin bedded, with some truncated cross laminations. Some schist layers in the middle part contain scattered pyrite cubes. Two- to 5-m-thick intervals of dark- and pale-green to mediumgray phyllite and laminated phyllite are present within the member. The base of the Whitaker Member consists of a 30-m-thick alternation

of finely to coarsely laminated quartzose schist and dark gray laminated schist, chloritic phyllite, and slaty phyllite. This interval is variably magnetic, containing few to a profusion of magnetite octahedra, some up to a few millimeters in size. The foliation is in general quite planar and parallels compositional layering (bedding). The main exceptions are in many of the phyllitic layers, where irregular masses of vein quartz disrupt the planarity of the foliation. In addition, the foliation is quite contorted in the 50– 100 m north of the fault contact with the Sykesville Formation. scc—Coyne Lock Member (New Name) (2300 m Thick) The Coyne Lock Member of the Scott Creek Formation consists primarily of laminated quartzose schist. The beds are light to medium gray, thin to medium bedded, very fine to fine-grained, quartzmuscovite-chlorite schists. The rocks are variably laminated, in which the laminations consist of alternating medium-light gray quartz-rich (80%) laminae and very dark gray phyllitic laminae. Generally, the lamination thickness varies with grain size, from ~1 mm to 5 mm. Some of the quartzose laminae are lensoid, suggesting ripples and microtroughs, often providing a flaser-like aspect. Concentrations of heavy minerals (zircon, monazite, xenotime, rutile, and titanite) at several horizons throughout the Coyne Lock Member are suggestive of sedimentary lag deposits. Rather profuse magnetite-rich layers also occur in several intervals in this member. Medium dark gray phyllitic schist and silty phyllite comprise a subordinate portion of the Coyne Lock Member. These more phyllitic layers occupy intervals from 1 to 10 m thick, interspersed among the more quartzose schists. Laminations are not prevalent in these beds, although they are present in some of the siltier layers. The basal part (300–350 m) of the Coyne Lock Member is well exposed at the mouth of Fishing Creek (Drumore) on the east side of the Susquehanna River. This ~300-m-thick interval consists primarily of medium dark gray to very dark gray phyllite, and subordinately medium gray silty phyllite and dark greenish-gray phyllite. Some of the siltier intervals are laminated. Quartz segregations are common (up to 40%–50% in places), generally in one of three forms: very thin (1– 3 mm) laminae (stringers); longer ribbons and lenses (5–20 cm thick); and large, irregularly shaped equidimensional masses (30–50 cm). The thinner quartz bodies tend to parallel the foliation, whereas the foliation generally wraps around the larger quartz masses. The rocks between the basal portion and the remainder of the Coyne Lock Member are transitional. The dark phyllites in the base grade upward through laminated silty phyllite and laminated chloritequartz-muscovite schist to the main part of the member with an increase in bed thickness, lamination thickness, grain size, and quartz content. The laminations consist of quartz-rich silt and sand grains interlayered with dark gray to black phyllitic. These rocks in turn grade upward into the medium to thickly laminated quartz-muscovite-chlorite schist and quartz-chlorite-muscovite schist characteristic of the main part of the member. On the west side of Susquehanna River, just south of Coyne Lock, a metabasalt lies exposed near the entrance to the north parking lot of the Peach Bottom power plant. It occurs ~1500 m above the base of the member. scb—Bryansville Member (New Name) (1600 m Thick) The Bryansville Member of the Scott Creek Formation contains a complex sequence of quartzose schists, schists, and phyllites, and a metabasalt. The basal part of the member consists of alternating quartzose schists and phyllitic schists. The quartzose schists are light to medium gray, flat bedded, laminated to medium bedded, and exhibit shallow troughs and bed pinchouts that suggest a high-energy

Peach Bottom area, Pennsylvania-Maryland Piedmont depositional environment. The intervening layers are medium dark greenish-gray phyllitic schists. Upsection, the quartzose schists become medium to thick-bedded quartzites, with thicker troughs. Overlying the quartzites are thin to medium bedded phyllitic and granular schists, some containing carbonate and/or albite. The medium gray quartzites, quartzose schists, and medium dark gray phyllitic schists above exhibit a 5- to 8-m-thick cyclicity suggesting upward-fining cycles. From 400 m above the base, most of the remaining member contains sequences consisting of repeated intervals of very dark gray phyllites, laminated to thin-bedded quartzose muscovite-chlorite schists, and thin to medium bedded quartzites and quartzose schists. The 25-m-thick Fishing Creek Metabasalt lies near the top of the member, overlain by dark gray (slightly greenish) phyllitic schist and chloritemuscovite-quartz schist. A few layers with considerable magnetite are present in the lower part. scbf—Fishing Creek Metabasalt (25 m Thick) The Fishing Creek Metabasalt (Smith and Barnes, 2004) occurs in the upper part of the Bryansville Member, ~150 m below the top. It is a moderately thick flow (~25 m) that is autochthonous with respect to the enclosing metasediments. The metabasalt is medium dark green, coarsely laminated, epidote-chlorite-albite rock with numerous magnetite octahedrons. The epidote also occurs in knots. Calcite is present in vugs, possibly replacing basalt glass during a hot fluid phase. The Fishing Creek Metabasalt is geochemically and stratigraphically equivalent to the Sams Creek Metabasalt of Maryland, itself a more “drifting” facies of the “initial rifting” facies of the Latest Precambrian Catoctin Formation (Smith and Barnes, 2004). The Fishing Creek Metabasalt also provides cross-Tucquan anticline correlation to the Glenrock, York County area, where it is more broken up. oct—Octoraro Formation (>1000 m Thick) The Octoraro Formation is distinguished by the presence of large metamorphic albite grains throughout many beds. Lyttle and Epstein (1987) redefined (reinstated) the Octoraro as the albite-chlorite schist facies portion of, and separated it from, what was formerly the Wissahickon Formation (Knopf and Jonas, 1923). This report describes only the uppermost beds exposed within the Peach Bottom area (lying just northwest of the overlying Scotts Creek Formation). Hill (2007) recently mapped and discussed the main portion of the Octoraro Formation lying farther to the northwest. The uppermost beds of the Octoraro Formation are thin to thick bedded, very parallel bedded, with very flat cross-bed troughs. Fine- to medium-grained beds of medium-gray quartzose (quartz-muscovitechlorite-albite) schist dominate, alternating with intervals (2–5 m thick) of greener, more chloritic schist. The former beds are distinctive in that white, medium-size grains of authigenic albite are common to dense in many of the layers. The Octoraro Formation is likely to be either lowermost Cambrian or late Neoproterozoic in age, based on its position below the Fishing Creek Metabasalt of the overlying Bryansville Member of the Scott Creek Formation.

REFERENCES CITED Agron, S.L., 1950, Structure and petrology of the Peach Bottom slate, Pennsylvania and Maryland, and its environment: Geological Society of America Bulletin, v. 61, p. 1265–1306, doi: 10.1130/0016-7606(1950)61[1265: SAPOTP]2.0.CO;2. Aleinikoff, J.N., Zartman, R.E., Walter, M., Rankin, D.W., Lyttle, P.T., and Burton, W.C., 1995, U-Pb ages of metarhyolites of the Catoctin and Mount

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Rogers Formations, central and southern Appalachians: Evidence for two pulses of Iapetan rifting: American Journal of Science, v. 295, p. 428–454. Bascom, F., and Stose, G.W., 1932, Description of the Coatesville and West Chester quadrangles, Pennsylvania-Delaware: U.S. Geological Survey Folio 223, 15 p. Bédard, J.H., and Stevenson, R., 1999, The Caldwell Group lavas of southeastern Quebec: MORB-like tholeiites associated with the opening of Iapetus Ocean: Canadian Journal of Earth Sciences, v. 36, p. 999–1019. Behre, C.H., Jr., 1933, Slate in Pennsylvania: Pennsylvania Geological Survey Mineral Report 16, 400 p. Berkheiser, S.W., 1994, Some commercial aspects of he Peach Bottom Slate: The problem of being too good, in Faill, R.T., and Sevon, W.D., eds., Various Aspects of Piedmont Geology in Lancaster and Chester Counties, Pennsylvania: Lancaster, Pennsylvania, 59th Field Conference of Pennsylvania Geologists Guidebook, p. 143–145. Chiarenzelli, J.R., and Valentino, D.W., 2006, Chemical homogenization during retrograde slate formation: Geological Society of America Abstracts with Programs, v. 38, no. 2, p. 86–87. Cloos, E., and Hietanen, A., 1941, Geology of the “Martic overthrust” and the Glenarm Series in Pennsylvania and Maryland: Geological Society of America Special Paper 35, 207 p. Dunning, G.R., and Hodych, J.P., 1990, U/Pb zircon and baddeleyite ages for the Palisades and Gettysburg sills of the northeastern United States: Implications for the age of the Triassic/Jurassic boundary: Geology, v. 18, p. 795– 798, doi: 10.1130/0091-7613(1990)018<0795:UPZABA>2.3.CO;2. Edwards, J., Jr., 1986, Geologic map of the Union Bridge quadrangle, Carroll and Frederick Counties, Maryland: Maryland Geological Survey Atlas. Edwards, J., Jr., 1988, Geologic map of the Woodsboro Quadrangle, Carroll and Frederick Counties, Maryland: Maryland Geological Survey Atlas. Faill, R.T., 1997, A geologic history of the north-central Appalachians, Part 1. Orogenesis from the Mesoproterozoic to the Taconic orogeny: American Journal of Science, v. 297, p. 551–619. Faill, R.T., 1998, A geologic history of the north-central Appalachians, Part 3. The Alleghany orogeny: American Journal of Science, v. 298, p. 131–179. Faill, R.T., 2009, Geologic map of the Conowingo Dam, Delta, Holtwood, and Wakefield 7-½ minute Quadrangles, Lancaster and York Counties, Pennsylvania: Pennsylvania Geological Survey, unpublished map. Frazer, P., Jr., 1880, The Geology of Lancaster County: Pennsylvania Geological Survey, 2nd series, v. CCC, 350 p. Freedman, J., Wise, D.U., and Bentley, R.D., 1964, Pattern of folded folds in the Appalachian Piedmont along Susquehanna River: Geological Society of America Bulletin, v. 75, p. 621–638, doi: 10.1130/0016-7606(1964)75[621:POFFIT]2.0.CO;2. Hietanen, A., 1951, Chloritoid from Rawlinsville, Lancaster County, Pennsylvania: American Mineralogist, v. 36, p. 859–868. Higgins, M.W., 1972, Age, origin, regional relations, and nomenclature of the Glenarm Series, central Appalachian Piedmont: A reinterpretation: Geological Society of America Bulletin, v. 83, p. 989–1026, doi: 10.1130/0016-7606(1972)83[989:AORRAN]2.0.CO;2. Higgins, M.W., 1977, The Baltimore Complex, Maryland and Pennsylvania, in Sohl, N.F., and Wright, W.B., eds., Changes in stratigraphic nomenclature by the U.S. Geological Survey, 1976: U.S. Geological Survey Bulletin 1435-A, p. A127–A128. Higgins, M.W., and Conant, L.B., 1986, Geologic map of Cecil County, Maryland: Maryland Geological Survey, 1 sheet, 1:62,500 scale. Hill, J.C., 2007, Bedrock geologic map of the northern portion of the Holtwood and Wakefield quadrangles, Lancaster and York Counties, Pennsylvania: Pennsylvania Geological Survey, 4th ser., Open-File Report OFBM 07-04.0, 16 p. (PDF). Horton, J.W., Jr., Aleinikoff, J.N., Drake, A.A., Jr., and Fanning, C.M., 1998, Significance of middle to late Ordovician volcanic-arc rocks in the central Appalachian Piedmont, Maryland and Virginia: Geological Society of America Abstracts with Programs, v. 30, no. 7, p. A-125. Knopf, E.B., and Jonas, A.I., 1923, Stratigraphy of the crystalline schists of Pennsylvania and Maryland: American Journal of Science, 5th series, v. 5, p. 40–62. Knopf, E.B., and Jonas, A.I., 1929, Stratigraphy of the crystalline schists of Pennsylvania and Maryland: U.S. Geological Survey Bulletin 799, p. 24–41. Lyttle, P.T., and Epstein, J.B., 1987, Geologic map of the Newark 1 × 2 degree Quadrangle, New Jersey, Pennsylvania, and New York: U.S. Geological Survey Report of Investigations 1715.

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MacLachlan, D.B., Buckwalter, T.V., and McLaughlin, D.B., 1975, Geology and mineral resources of the Sinking Spring quadrangle, Berks and Lancaster Counties, Pennsylvania: Pennsylvania Geological Survey, 4th ser., Geologic Atlas A177d, 228 p. Markham, J.L., 2009, Fluid history of the Peach Bottom Slate and adjacent units, southeastern Pennsylvania [M.S. thesis]: Bowling Green, Ohio, Bowling Green State University, 65 p. McKague, H.L., 1964, The geology, mineralogy, petrology, and geochemistry of the State Line serpentinite and associated chromite deposits [Ph.D. thesis]: University Park, Pennsylvania, Pennsylvania State University, 166 p. Miller, B.J., 1935, Age of the schists of the South Valley Hills, Pennsylvania: Geological Society of America Bulletin, v. 46, p. 715–756. Muller, P.D., Candela, P.A., and Wylie, A.G., 1989, Liberty Complex: Polygenetic mélange in the central Maryland Piedmont, in Horton, J.W., Jr. and Rast, N., Mélanges and Olistostromes of the U.S. Appalachians: Geological Society of America Special Paper 228, p. 113–134. Rankin, D.W., 1975, The continental margin of eastern North America in the southern Appalachians: The opening and closing of the proto–Atlantic ocean: American Journal of Science, v. 275A, p. 298–336. Rankin, D.W., Drake, A.A., Jr., and Ratcliffe, N.M., 1993, Proterozoic North American (Laurentian) rocks of the Appalachian orogen, in Rankin, D.W., Chiarenzelli, J.R., Drake, A.A., Jr., Goldsmith, R., Hall, L.M., Hinze, W.J., Isachsen, Y.W., Lidiak, E.G., McLelland, J., Mosher, S., Ratcliffe, N.M., Secor, D.T., Jr., and Whitney, P.R., Proterozoic rocks east and southeast of the Grenville front, in Reed, J.C., Jr., Bickford, M.E., Houston, R.S., Link, P.K., Rankin, D.W., Sims, P.K., and Van Schmus, W.R., eds., Precambrian: Conterminous U.S.: Geological Society of America, The Geology of North America, v. C-2, p. 378–422. Rodgers, J., 1970, The Tectonics of the Appalachians: New York, Wiley Interscience, 271 p. Sanford, R.F., 1982, Growth of ultramafic reaction zones in greenschist to amphibolite facies metamorphism: American Journal of Science, v. 282, p. 543–616. Sinha, A.K., Hanan, B.B., and Wayne, D.M., 1997, Igneous and metamorphic U-Pb zircon ages from the Baltimore Mafic Complex, Maryland Piedmont, in Sinha, A.K., Whalen, J.B., and Hogan, J.P., eds., The Nature of Magmatism in the Appalachian Orogen: Geological Society of America Memoir 191, p. 275–286 Smith, R.C., II, 1993, Tell-tale talcs—Chemical clues to unravel the Earth’s secrets: Pacific Geology, v. 24, no. 1, p. 2–6. Smith, R.C., II, 1994, Bald Friar Metabasalt and geochemistry of ultramafic fragments, in Stop 14. Peach Bottom section at the Susquehanna River, in Faill, R.T., and Sevon, W.D., eds., Various Aspects of Piedmont Geology in Lancaster and Chester Counties, Pennsylvania: Lancaster, Pennsylvania, 59th Field Conference of Pennsylvania Geologists Guidebook, p. 180–184. Smith, R.C., II, 2003, Late Neoproterozoic felsite (602.3 ± 2 Ma) and associated metadiabase dikes in the Reading Prong, Pennsylvania, and rifting of Laurentia: Northeastern Geology and Environmental Science, v. 25, p. 175–185. Smith, R.C., II, 2006, Bald Friar Metabasalt and Kennett Square Amphibolite: Two Iapetan ocean floor basalts: Northeastern Geology and Environmental Science, v. 28, p. 238–253. Smith, R.C., II, and Barnes, J.H., 1994, Geochemistry and geology of metabasalt in southeastern Pennsylvania and adjacent Maryland, in Faill, R.T., and Sevon, W.D., eds., Various aspects of Piedmont Geology in Lancaster

and Chester Counties, Pennsylvania: Lancaster, Pennsylvania, 59th Field Conference of Pennsylvania Geologists Guidebook, p. 45-72A. Smith, R.C., II, and Barnes, J.H., 2004, White Clay Creek amphibolite: A Piedmont analog of the Catoctin metabasalt, in Blackmer, G.C., and Srogi, L., eds., Marginalia—Magmatic arcs and continental margins in Delaware and southeastern Pennsylvania: West Chester, Pennsylvania, 69th Field Conference of Pennsylvania Geologists, p. 28–45. Smith, R.C., II, and Barnes, J.H., 2008, Geology of the Goat Hill Serpentine Barrens, Baltimore Mafic Complex, Pennsylvania: Journal of the Pennsylvania Academy of Science, v. 82, no. 1, p. 19–30. Smith, R.C., II, Foland, K.A., and Nickelsen, R.P., 2004, The Lower Silurian Clear Spring Volcanic Suite, Sword Mountain Olivine Melilitite (433 ± 3 Ma) and Hanging Rock Tuff/Diatreme, Washington County, Maryland: Geological Society of America Abstracts with Programs, v. 36, no. 2, p. 71. Southwick, D.L., and Owens, J.P., 1968, Geologic map of Harford County, Maryland: Maryland Geological Survey, 1 sheet, 1:62,500 scale. Taylor, S.R., and McClennan, S.M., 1985, The continental crust: Its composition and evolution: An examination of the geochemical record preserved in sedimentary rocks: Oxford, Blackwell Scientific Publications, 312 p. Tollo, R.P., and Aleinikoff, J.N., 1996, Petrology and U-Pb geochronology of the Robertson River igneous suite, Blue Ridge province, Virginia— Evidence for multistage magmatism associated with an early episode of Laurentian rifting: American Journal of Science, v. 296, p. 1045–1090. Tollo, R.P., Aleinikoff, J.N., Bartholomew, M.J., and Rankin, D.W., 2004, Neoproterozoic A-type granitoids of the central and southern Appalachians— Intraplate magmatism associated with episodic rifting of the Rodinian supercontinent: Precambrian Research, v. 128, p. 3–38, doi: 10.1016/j .precamres.2003.08.007. Valentino, D.W., 1994, Stop 9, Drumore Tectonite along the East Branch of Octoraro Creek, and Stop 12, The Drumore Tectonite and the Peters Creek Formation, in Faill, R.T., and Sevon, W.D., eds., Various Aspects of Piedmont Geology in Lancaster and Chester Counties, Pennsylvania: Lancaster, Pennsylvania, 59th Field Conference of Pennsylvania Geologists Guidebook, p. 16 and 175–176. Valentino, D.W., Gates, A.E., and Glover, L., III, 1994, Late Paleozoic transcurrent tectonic assembly of the central Appalachian Piedmont: Tectonics, v. 13, p. 110–126, doi: 10.1029/93TC02313. Wise, D.U., 1970, Multiple deformation, geosynclinal transitions and the Martic problem in Pennsylvania, in Fisher, G.W., Pettijohn, F.J., Reed, J.C., Jr., and Weaver, K.N., eds., Studies of Appalachian geology: central and southern: New York, Interscience, Publishers, p. 317–333. Wise, D.U., and Ganis, G.R., 2009, Taconic Orogeny in Pennsylvania: A 15–20 m.y. Apennine-style Ordovician event viewed from its Martic hinterland: Journal of Structural Geology, v. 31, p.887–899. Wyckoff, D., edited by Crawford, W.A., 1990, A history of the Martic line controversy: The pre-WWI years, in Scharnberger, C.K., ed., Carbonates, Schists, and Geomorphology in the Lower Reaches of the Susquehanna River Valley: Lancaster, Pennsylvania, 55th Field Conference of Pennsylvania Geologists Guidebook, p. 12–39.

MANUSCRIPT ACCEPTED BY THE SOCIETY 2 DECEMBER 2009

Printed in the USA

The Geological Society of America Field Guide 16 2010

Soils, geomorphology, landscape evolution, and land use in the Virginia Piedmont and Blue Ridge W. Cullen Sherwood Anthony S. Hartshorn L. Scott Eaton Department of Geology and Environmental Science, James Madison University, 800 S. Main Street, Harrisonburg, Virginia 22807, USA

ABSTRACT The object of this field trip is to examine the geology, landforms, soils, and land use in the eastern Blue Ridge and western Piedmont geologic provinces in Orange County in central Virginia. A complex mix of igneous, sedimentary, and metamorphic bedrocks, ranging in age from Mesoproterozoic to Triassic (possibly some Jurassic) underlie the area. Soils are equally varied with a total of 62 series mapped in Orange County alone. The area being relatively stable tectonically, landforms generally reflect the resistance to weathering of the bedrock. Area landforms range from a low ridge over Catoctin greenstone to a gently rolling Triassic basin. Soils examined on the trip represent three orders: Ultisols, Alfisols, and Inceptisols. Residual soils clearly reflect the compositions of the parent rocks and saprolites are common. Map patterns of forested versus nonforested lands bear a striking resemblance to the distribution patterns of the different soil and bedrock types. Our work has shown that the vast majority of the land in central Virginia, even that forested today, shows evidence of past clearing and cultivation. However, the harsh demands of growing tobacco wore out the less fertile and more erodible soils by the mid-nineteenth century resulting in their abandonment and the subsequent regeneration of the vast tracts of hardwood forests we see today. Only the most productive soils remain in agriculture. REGIONAL GEOLOGY

is composed of Mesoproterozoic crystalline rocks, some >1 Gy (Southworth et al., 2009). During the Grenvillian orogeny, granitic intrusions and associated metamorphism affected the core. Moving outward from the core, rocks of Neoproterozoic age form the flanks of the massif. Here a series of sedimentary rocks derived from the weathering of the core—ranging from alluvial conglomerates to marine shales—were deposited along with volcanic ash. These rocks make up the Swift Run Formation on the west side and the Lynchburg Group on the east side of the eroded core (Fig. 2). Subsequently, basalt flows associated with

Virginia can be divided into five geologic provinces (Fig. 1). Our trip will include portions of the Blue Ridge and Piedmont provinces in central Virginia. Blue Ridge Province The oldest rocks in Virginia are found in the Blue Ridge Province, a complex basement massif. The core of the massif

Sherwood, W.C., Hartshorn, A.S., and Eaton, L.S., 2010, Soils, geomorphology, landscape evolution, and land use in the Virginia Piedmont and Blue Ridge, in Fleeger, G.M., and Whitmeyer, S.J., eds., The Mid-Atlantic Shore to the Appalachian Highlands: Field Trip Guidebook for the 2010 Joint Meeting of the Northeastern and Southeastern GSA Sections: Geological Society of America Field Guide 16, p. 31–50, doi: 10.1130/2010.0016(02). For permission to copy, contact [email protected]. ©2010 The Geological Society of America. All rights reserved.

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the opening of the proto–Atlantic Ocean flowed out over the sedimentary units. Later, these basalts were metamorphosed to form the greenstones of the Catoctin Formation. Following the rifting and associated basalt flows, clastic sediments associated with the proto–Atlantic Ocean margin covered portions of the flows. These make up the Chilhowie Group, which occupy the western flank of the Blue Ridge Mountains and the Candler Formation east of Southwest Mountain. Metamorphism that altered the Catoctin basalts to greenstone affected the siliceous clastic rocks producing mainly quartzites, phyllites, and schists (Rader and Evans, 1993). Topographically, the Blue Ridge Mountains with elevations of up to 1200 m (4000 ft) dominate the western portion of the Blue Ridge Province. These mountains are composed of a mix of basement rocks, Catoctin greenstone, and Chilhowie metasediments. Today, the Shenandoah National Park occupies much of the Blue Ridge Mountain crest in northwestern Virginia. Moving eastward from the Blue Ridge Mountains toward the core of the Blue Ridge Province, a number of monadnocks or inselbergs are evident in the western portion of the province. Farther east, the land is gently rolling until it reaches a low ridge composed of Catoctin greenstone that forms the eastern edge of the Blue Ridge Province.

83°45'W 40°37'N

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Figure 1. Generalized geologic provinces (Appalachian Plateau, Valley and Ridge, Blue Ridge, Piedmont, and Coastal Plain) of Mid-Atlantic States. Inset rectangle shows location of this field trip.

Figure 2. Simplified cross section of the Blue Ridge anticlinorium (modified from Sherwood and Eaton, 1993, p. 57).

Structurally, the Blue Ridge has been variously described as an anticlinorium with the flanks held up by the relatively resistant Catoctin greenstone and the oldest formations in the center (Fig. 2). Other interpretations involve a series of thrust faults to account for the surface pattern of rock types (Fig. 3). It is generally agreed that the rocks of the Blue Ridge Province are allochthonous, having been thrust northwestward over the Paleozoic rocks of the Valley and Ridge. Piedmont Province The Piedmont Geologic Province borders the Blue Ridge Province to the southeast. Today’s complex bedrock geology reflects an equally complex geologic past. Basically, the rocks of the Virginia Piedmont are composed of a series of allochthonous terranes that have accreted to the North American plate by the closing of the ancient proto–Atlantic Ocean. Figure 4 is a generalized geologic map showing the extent of the individual terranes. The western Piedmont terrane is composed of early Paleozoic igneous and metasedimentary rocks associated with the suture zone with the Blue Ridge Province. Rocks of the western Piedmont are thought to be composed of fragments of the divergent continental margin created during the opening of the protoAtlantic and later metamorphosed to greenschist and amphibolite facies (Rankin, 1975). To the southeast of the western Piedmont terrane lies the Chopawamsic volcanic belt (Fig. 4). This belt is composed of a series of volcanic and plutonic rocks that appear to have originated as an island arc offshore in the proto–Atlantic Ocean (Pavlides, 1980). The Lahore pluton (Figs. 2 and 5; Pavlides, 1982) at Stop 6 is one of the several mafic intrusions occurring within the Chopawamsic Belt in Virginia. Converging forces caused the arc to collide with the North American plate. Both the western Piedmont and Chopawamsic volcanic terranes have been intruded by late Paleozoic granites, and undergone several episodes of deformation and metamorphism, producing an exceedingly complex geology. Today the Piedmont is characterized by a gently rolling landscape incised by numerous streams and a few large rivers. Deep weathering and saprolites are common, and bedrock outcrops are scarce, occurring mainly in the beds and banks of streams.

Figure 3. Cross section of the regional geological (modified from Bailey, 2006).

Virginia Piedmont and Blue Ridge Culpeper Mesozoic Basin Subsequent to the closing of the proto-Atlantic, tensional stresses began to affect Earth’s crust in what is now eastern North America. During the Triassic, rifting at several points created a series of grabens and half-grabens. The largest of these features in Virginia is the Culpeper Basin. The Culpeper Basin extends from just south of Orange County northward into Maryland, a distance of more than 148 km (Lindholm, 1979). The rocks of the basin have traditionally been associated with the Triassic Newark Group (Fig. 5). For many years the Culpeper and related basins were referred to as Triassic basins. Pollen studies (Cornet, 1977), however, determined the rocks to range from Upper Triassic to Lower Jurassic in age. As rifting occurred, sediment eroding from the adjoining highlands was rapidly transported into the basins. Thick beds of sand, silt, and clay were deposited throughout the basins, and coarse conglomerates formed near the border faults. Red beds deposited in shallow fresh water and on broad alluvial plains

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are the most common rock types found in the Culpeper Basin (Lindholm, 1979). In general, the sedimentary sequence exhibits coarse clastic rocks at the base fining upward to the Bull Run shale. In Virginia, most of the conglomerates are found contiguous to the normal faults bordering the west side of the basin. However, near the south end of the basin, north of Barboursville, Lindholm (1979) reported conglomerates to be more extensive along the southeast side. One of the authors (Sherwood) also described well-developed conglomerate along the southeast border of the basin north of the Rappahannock River in Fauquier County. Lithologies of the clasts making up the conglomerates change along the faults reflecting the source rocks being weathered and transported into the basin. Examination of the western border conglomerate, ~10 km north of our Stop 2 at Barboursville, found it to contain a variety of clasts, including creamcolored shale, vein quartz, phyllite, greenstone, and granite (Sherwood, 2003). In general, the beds within the basin dip to the west indicating periodic or subsequent movement along the

Figure 4. Terrane map of the Piedmont and Blue Ridge provinces, Virginia (modified from Bailey, 2006).

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western faulted edge. Early dinosaur tracks have been observed at several exposures within the basin (Roberts, 1928). Farther north across the Rapidan River, the Culpeper Basin contains numerous mafic igneous intrusions and several basaltic lava flows. The mafic intrusions are predominately diabase or “trap rock.” These intrusions serve as the major sources of crushed stone in northern Virginia. In some cases, contact metamorphism caused by the intrusions result in zones of hornfels extending as much as hundreds of meters into the surrounding sedimentary rocks. REGIONAL SOILS The “diversity and complexity of the geology in central Virginia are accompanied by an equally diverse and complex mantle of residual soils” (Plaster and Sherwood, 1971, p. 2813). On today’s field trip, we will focus on the uppermost two meters of the regolith. This depth will allow observations of the critical A and B horizons or solum on which modern soil classification is based. As we know, soils reflect the variable influences of soilforming state factors (Jenny, 1941), including climate, organisms, relief, parent material, and time. A major focus of this field trip will be on parent material and time (age of the soils). Pavich

(1986) framed this field trip well, when he wrote “… how climate affects the rock weathering is complexly dependent upon soil and rock structure.” In the Virginia Blue Ridge and Piedmont, soil properties show great fidelity to the underlying lithology. The most common soil order in the Piedmont is the Ultisol (Buol, 1973; Fig. 6), highly weathered soils with low base saturation. Base saturation is a measurement of the percentage of a soil’s cation exchange capacity (negatively charged edges) that is occupied by the base cations Ca2+, Mg2+, Na+, and K+. Ultisols predominate over metasedimentary rocks and felsic intrusions. A second well-developed soil order, the Alfisols, consists of intermediately weathered soils with high base saturation; Alfisols are typically mapped over mafic and ultra-mafic parent materials, such as gabbro and diabase (Genthner, 1990), as will be seen at Stop 6. Other common orders mapped in this area are Inceptisols and Entisols; these are both relatively unweathered soils showing little or no morphologic profile differentiation, and are commonly mapped on steep slopes, highly dissected uplands, or floodplains. While the current landscape appears stable, and the soil properties we observe and use to classify the soils (texture, color, quantity and types of clays) are consistent with long residence times, portions of the Piedmont landscape may have been less

Figure 5. Geological units in the field trip area, with individual stops (1–6) indicated (modified from Dicken et al., 2008).

0.14 0.18 0.11 0.08 0.10

stable in the past. For example, ~60% of soil parent material along a 22-km pipeline trench in Culpeper and Orange counties was interpreted to be colluvium, while 10% was alluvium (Whittecar, 1985). However, the regional fidelity between lithology and soil properties (and land use), implies large areas of the Piedmont landscape are residual soils, formed in place.

35 Plant available water 3 –3 (cm water cm soil)

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33 12 14 5 43 Clay loam Silt loam Sandy loam Loamy sand Sandy clay

Figure 6. Simplified soils map of the field trip area, showing broad soil series zones: Stop 1 will feature an example of a soil mapped as the Watt series (Inceptisol soil order), Stop 2 the Bucks-Penn series (Alfisol, Ultisol), Stop 3 the Davidson series (Ultisol), Stop 4 the Nason-Tatum series (Ultisols), Stop 5 the Wehadkee series (Inceptisol), and Stop 6 the Iredell series (Alfisol).

Textural polygon

Midpoint % clay (<2 μm)

On this field trip, a number of soil properties will be highlighted, including Munsell colors, textures, B horizon thicknesses, clay mineralogy, as well as base saturation. As a quick guide to the overarching importance of soil texture, Table 1 shows a number of properties associated with five of the 12 polygons that are commonly found on texture triangles, calculated using an online hydraulic properties calculator (http://www.pedosphere.com/resources/texture/triangle_us .cfm?192,220). Most of the field trip sites we visit today exhibit very old soils. As Pavich et al. (1989, p. 48) have noted, Piedmont soils “show relatively the same amount of development” as postMiocene soils (Markewich et al., 1987; Owens et al., 1983). For example, quartz dissolves and illite transforms to vermiculite in soils >100 k.y., and >1 m thick sola can be 1–2 Ma (Markewich et al., 1987). While a comparison with Coastal Plain soils suggests Piedmont soils are “no older than Pliocene (5 Ma) and probably no older than Pleistocene (2 Ma),” a number of studies suggest that specific geomorphic features such as the oldest James River terraces are 10 Ma or older (e.g., Howard et al., 1996). Residence times are to some extent a reflection of the balance between soil

TABLE 1. SELECT HYDRAULIC PROPERTIES ASSOCIATED WITH A SUBSET OF SOIL TEXTURAL CLASSES Saturated Field capacity Wilting point Midpoint Typical bulk Saturated volumetric –3 3 –3 3 hydraulic (cm water cm soil) (cm water cm soil) % silt density water content 3 –3 –3 conductivity (g cm ) (cm water cm soil) (2–50 μm) –1 (cm h ) 34 1.32 0.31 0.50 0.32 0.18 66 1.44 2.37 0.46 0.28 0.10 23 1.51 1.60 0.43 0.21 0.10 11 1.68 6.40 0.37 0.15 0.06 7 1.31 0.13 0.50 0.33 0.23

Soil Properties and Nomenclature

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production rates and total denudation (physical erosion + chemical weathering) rates. However, the quantification of these rates is not by any means straightforward. Saprolite The first part of this field guide described the geological setting for the soils we will see today. In addition to the underlying bedrock, however, this part of Virginia is underlain by meters of saprolite (Pavich, 1985; Genthner, 1990), in some cases exceeding 50 m (Pavich et al., 1989). Importantly, however, this saprolite thickness shows great spatial variability (Fig. 7), with greater thicknesses reported over felsic lithologies and thinner saprolites over mafic lithologies, such as the amphibole monzonite that underlies our last Stop (Stop 6: Iredell soil series).

B A Figure 7. Variation in regolith thickness developed over more felsic rock such as foliated metasedimentary and granitic rocks (left, regolith A; ~19 m) and more mafic rock such as diabase (right, regolith B; ~4 m) in northern Virginia (modified from Pavich et al., 1989, p. 5). Note the shift in vertical scale between regolith A and B.

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Figure 8 (continued on following page). Piedmont regolith cross sections: (A) Catawba Valley, North Carolina (modified from Kerr, 1881, p. 347), (B) northern South Carolina (modified from Eargle, 1940, p. 337), (C) Fairfax County, Virginia (modified from Pavich et al., 1989, p. 44).

Virginia Piedmont and Blue Ridge The extreme thickness of the regolith in the southern Piedmont has been a matter of historical interest to both geologists and soil scientists alike. W.C. Kerr (1881) remarked: “To a foreign geologist, entering the Middle and South Atlantic States for the first time, a hundred miles or more from the coast, the most striking and novel feature of the geology is the great deal of earth which almost everywhere mantles and conceals the rocks” (p. 345). Kerr provided some of the earliest cross sections of Piedmont soils (Fig. 8A), and called attention to what he termed “three kinds of earthy layers, each having a different structure and origin” (p. 345). Roughly translated, these are the soils, saprolite, and unweathered rock. Kerr used these strata, and their relationships with the geomorphic and soil surface, to suggest that “the present topography of the surface is the result of extensive erosion, subsequent to the accumulation of these deposits” (p. 347, italics his). He alludes to inverted topography, and also goes on to invoke “frost drift” as the mechanism most clearly responsible for the locations of colluvial material Eargle (1940, 1977) further explored the question of the origin of the soil mantle. He suggested that lateral soil transport was

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responsible for what he termed “episodic accumulation.” This included burial of organic-rich paleosols, not just in hollows, but occasionally on topographically high points (Fig. 8B). Some variations within individual soil types were attributed to “the partial assortment of materials during soil migration” (Eargle, 1940, p. 337–338), a description similar to that used by Milne (“differential reassortment of mass”; 1935a, 1935b, 1936) to describe the genesis of African catenas. These translocation concepts have been expanded upon (e.g., Morison, 1948; Sommer et al., 2000) to argue that, just as eluvial and illuvial processes (such as the transfers of clays) can be invoked to explain vertical differentiation of soil profiles, so too can these processes be invoked to explain lateral differentiation of catenas. More recently, Pavich et al. (1989, Fig. 8C) have observed similar discrepancies between the upper boundaries of soils and saprolite. These observations by Kerr and Eargle, and those of Pavich et al. (1989; Fig. 8C), together point to a fundamental topographic mismatch between the top of the regolith and the top of the saprolite. This mismatch suggests residual interpretations of the origin of Piedmont soils might be incomplete, and are

C

Figure 8 (continued).

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consistent with the ideas that parent material has been translocated and that topographic inversion has occurred in many areas of the Piedmont. Systematic differences in saprolite thickness across the Piedmont can be traced to the underlying lithologies, because “despite the mass loss due to dissolution of less resistant minerals (such as plagioclase and biotite), the quartz and muscovite remain relatively unaltered, and grain-to-grain contacts of muscovite retain structural integrity” (Pavich et al., 1989). Thus, the prolonged and intense chemical weathering of felsic materials leads to the formation of a quartz- and/or muscovite-rich skeleton that allows for continued solutional (“plasma” sensu Pédro, 1983) losses, leading to great saprolite thicknesses. Weathering of mafic materials, by contrast, produces collapse of the profile, yielding much thinner saprolites. In a study of the Maryland Piedmont, Costa and Cleaves (1984) suggested saprolite thickness reflected parent material mineralogy and degree of metamorphism. They also suggested the mineralogy of the saprolite reflected landscape setting: kaolinite and quartz dominated drier, upland portions of catenas, whereas kaolinite, quartz, and smectites dominated wetter, lower portions of catenas. The Piedmont, while upstream of a passive margin, is not tectonically quiescent. Tertiary marine formations such as the Miocene Calvert Formation now lie as much as 150 m above sea level, implying uplift of 20 m/m.y. (Darton, 1951). Furthermore, fall line compressional faults have thrust crystalline rocks over younger sedimentary rocks, and these faults have been active during the Cenozoic (Mixon and Newell, 1977; Prowell, 1976). Long-term lowering of the Potomac Valley (~15 m/ m.y.) through the Piedmont has been suggested to reflect a combination of “slow flexural uplift of the Atlantic margin from offshore sediment unloading, isostatic response to denudation, and protracted late Cenozoic sea-level fall” (Reusser et al., 2004, p. 499). Genthner (1990) has suggested that the eastern Blue Ridge and Piedmont’s rolling hillslopes led Hack (1960) and Pavich (1986) to believe they represent an equilibrium landscape: erosion rates appear to be matching long, slow uplift rates.

FIELD TRIP ITINERARY This field guide will emphasize U.S. soil taxonomic classifications instead of, for example, World Reference Base classifications (the WRB is successor to the Food and Agriculture Organization’s classification system), because the U.S. system names have, in our opinion, a higher information density. And it is worth remembering Lewis Carroll’s (1871) thinly veiled explanation for why we classify soils: “What’s the use of their having names,” the Gnat said, “if they won’t answer to them?” “No use to them,” said Alice, “but it’s useful to the people that name them, I suppose. If not, why do things have names at all?”

Introduction Today’s trip will cover portions of the eastern Blue Ridge and western Piedmont geologic provinces, and the Culpeper Mesozoic basin in Orange County, Virginia (Figs. 1 and 9). The object will be to examine relationships between bedrock geology, soils, and historic uses of the land. Over the years, the authors have studied the historic patterns of land use and their relationships to the bedrock geology and soils of Virginia. Central Virginia presents an ideal outdoor laboratory for studies of this type because of the sharply juxtaposed and contrasting bedrock geology, and soils and the relatively long (human) historical context. The long residence times of the soils, in some cases >1 m.y., have allowed for the clear expression of normally nuanced climate, biotic, relief, and parent material effects on soillandscape relationships. Finally, the general area remains relatively free of intensive urbanization so historic land use patterns largely reflect long-term agricultural and silvicultural land uses. Figure 10 shows the land use patterns that have developed over the past four hundred years. Basically, areas cleared of forest cover are assumed to be underlain by soils worthy of the effort involved in clearing and maintaining the land in nonforest

Figure 9. Road map of field trip area, with the six stops indicated.

Virginia Piedmont and Blue Ridge

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Figure 10. Satellite map showing land use in the vicinity of each of our stops.

(mainly agricultural) use. Forested areas, on the other hand, could either be underlain by soils with significant limitations for other uses or be the product of decades of afforestation following nineteenth and/or twentieth century clearing and abandonment. Having said this, we are acutely aware that modern views value forests for many different reasons. Consequently, in the future the old systems of land values may change, and future patterns of land use in this region may well change with them. In a nutshell, Stops 3 and 4 are underlain by Ultisols, Stop 6 by Alfisols, and Stops 1 and 5 are graphite schist- and floodplain-associated Inceptisols, respectively; Stop 2 includes soils mapped as Ultisols (Bucks soil series) and Alfisols (Penn soil series) (Table 2). Miles 0.0 3.5 4.6

Ruckersville. From Ruckersville proceed east on Rt. 33 to Eheart. Turn left on Rt. 644, proceed to intersection with Rt. 657.

Stop

Soil order

1 2 3 4

Inceptisol Ultisol/Alfisol Ultisol Ultisol/Ultisol

5 6

Inceptisol Alfisol

Stop 1. Eheart Local Geology Stop 1 is located over the Lynchburg Formation named for exposures of schist and gneiss near Lynchburg, Virginia, by Jonas (1927). Rocks of the Lynchburg are Neoproterozoic in age, existing between the Grenville basement and the overlying Catoctin Formation (Fig. 4). The predominant rock types making up the Lynchburg in the vicinity of Stop 1 are biotite-muscovite schists and phyllites, interfingered with bands of graphite schist. Stratigraphic treatment of the graphite-rich rocks at Stop 1 has differed among the various investigators over the years. Jonas (1927) referred to these rocks as the Johnson Mill member of the Lynchburg Formation. Nelson (1962) gives them formational status as the Johnson Mill Graphite Slate Formation. Allen (1963) refers to a graphite schist facies of the Lynchburg Formation. While it is generally agreed that the schists and phyllites of the Lynchburg Formation are sedimentary in origin, the origin of the carbon concentrated in the graphite schist presents an

TABLE 2. SALIENT FEATURES OF THE SIX STOPS Soil series Piedmont acres Orange County acres (000) (%) Graphite schist Watt 2.7 660 (0.3) Triassic siltstone Bucks/Penn 65/367 1800 (0.8)/3700 (1.7) Greenstone Davidson 553 21,900 (10.0) Schist and phyllite Tatum/Nason 720/430 24,100 (11.0)/43,800 (20.0) Parent rock

Alluvium Wehadkee 655 1300 (0.6) Amphibole Iredell 231 440 (0.2) monzonite 2 Note: Orange County has an area of 875 km , or ~219,000 acres. U.S Department of Agriculture (1971).

Notes 85% forested Productive agricultural soil All-purpose soil Highly erosive and acidic; ubiquitous Redoximorphic features Shrink-swell clays, high base saturation

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intriguing question. A search of the existing literature resulted in no evidence that this topic has been addressed in a systematic way. Rumors have circulated for many years to the effect that the rocks contained ~8% graphite and mining had been considered at one time. Soils Here we will observe a soil mapped as part of the Watt soil series, also formally known as (“afka”) a loamy-skeletal, mixed, semiactive, mesic Typic Dystrudept, meaning (U.S. soil taxonomic classifications are read from right to left, and the formative elements of the subgroup are in italics). Order→Inceptisol (weakly developed, “infantile” soil) Suborder→Udept (udic soil moisture regime: wetter than ustic [semiarid grasslands or savannas], drier than perudic [precipitation exceeds evapotranspiration in every month of the year]) Great group→Dystrudept (low base saturation) Subgroup→Typic Dystrupdept (insufficiently distinguished [morphologically or chemically] to merit classification as another subgroup). Watt (A) horizons are typically dark-gray (5Y 2/1), channery silt loams, dominated by graphitic schist fragments, and can be extremely acid with very low base saturation. Watt soils are commonly mapped over rocks with traces of sulfides, which results in very low pH values. Watt B horizons are thin, show an increase in rock fragments with depth, and bedrock is typically within 0.5–1 m of the surface. Because of the dark parent material, “in most places, the C horizon is darker colored than the solum” (Official Series Description, Natural Resources Conservation Service [NRCS]). The areal extent of Watt soils was derived with an online Soil Extent Mapping Tool (link [e.g., http://www.cei. psu.edu/soiltool/semtool.html?seriesname = WATT] at the official series description web site): Watt soils cover 2700 acres across the Piedmont, and 745 acres (0.3%) across Orange County. The Watt soil we will examine is interfingered at a very fine spatial scale with the Glenelg soil series, afka fine-loamy,

mixed, semiactive, mesic Typic Hapludults (Table 3). From the official Glenelg series description: “Depth to bedrock is 2–3 m or more. Rock fragments range from 0 to 35% throughout the solum and 5–55% in the C horizon. Fragments are mostly hard white quartz or schist and range from gravel or channers to stones in size. Stone content ranges from 0 to 5%. Mica content increases sharply in the lower part of the solum and substratum. Unlimited reaction ranges from very strongly acid to slightly acid.” Land Use Due to its shallow depth to bedrock, high acidity (pH <5), and low organic matter content, the Watt series exhibits severe limitations for most land uses. In Orange County, over 80% of the Watt remains in forest cover (see Fig. 10). Even the quality of the forests is generally poor. Acid-tolerant tree species such as chestnut oak, hickory, and maple predominate, but growth rates are slow. The understory contains dogwood and wild blueberries. The Orange County Soil Survey notes the best trees to plant on Watt soils for pulp wood or timber production are loblolly and short leaf pines. Use of the Watt soil for agriculture in Virginia is almost nonexistent and is generally not recommended. All of the characteristics noted above plus excessive drainage create poor conditions for crops and even hay or pasture. U.S. Department of Agriculture (USDA) tables (Carter et al., 1971) for the Watt do not even list expected yields for the principal grains and hay, and rate its use for pasture as lowest of the 63 soils mapped in Orange County (U.S. Department of Agriculture, 1971). Potential for development also has the Watt series receiving low ratings. The principal limitation is the shallow depth which makes it unsuitable for common on-site sewage disposal systems that use septic tanks and drain fields. Alternative systems can be used but are expensive and require a high level of maintenance. Watt soils can also present significant engineering problems, such as the necessity for blasting during excavations for foundations, deep ditches, or road cuts. They also provide inadequate fill material and exhibit poor compactibility.

TABLE 3. SELECTED PROPERTIES OF THE WATT SOIL SERIES Horizon

1

Lower depth Colors Textures Other (cm) 2 3 A1 5 5Y 2/1 Channery silt loam 20% graphitic schist gravels (>2 mm); very strongly acid A2 23 5Y 3/2 Channery silt loam 25% graphitic schist gravels; very strongly acid Bw 35 5Y 4/2 Channery silt loam 40% graphitic schist gravels; very strongly acid C 65 5Y 2/1 Very channery silt loam 55% graphitic schist gravels; very strongly acid R 65+ (Rock) Note: See http://ortho.ftw.nrcs.usda.gov/cgi-bin/osd/osdlist.cgi. 1 All colors are moist and follow Munsell notation: e.g., 5Y 2/1 is hue 5Y, value 2, chroma 1. 2 Textures are modified according to the contribution (by volume) and type of rock fragments >2 mm: for <15% gravels or channers (prismoidal or subprismoidal, flat) rock fragments, there is no modifier; 35% > volume > 15%, “gravelly” or “channery”; 60% > volume > 35%, “very gravelly” or “very channery” (Schoeneberger et al., 2002, p. 2–31). 3 Reaction classes or field pH values are as follows: very strongly acid: 5.0 > pH > 4.5; strongly acid: 5.5 > pH > 5.1; moderately acid: 6.0 > pH > 5.6; slightly acid: 6.5 > pH > 6.1; neutral: 7.3 > pH > 6.6; slightly alkaline: 7.8 > pH > 7.4; moderately alkaline: 8.4 > pH > 7.9 (Schoeneberger et al., 2002, p. 2–70).

Virginia Piedmont and Blue Ridge Miles 6.2 8.0

From Stop 1, proceed east on Rt. 657 to left on Rt. 33. Barboursville intersection of Rt. 33 and Rt. 20.

Stop 2. Barboursville Local Geology Stop 2 places us at the far south end of the Culpeper Mesozoic basin described earlier under Regional Geology. For many years the basin was thought to be continuous to a point just south of Barboursville, where it ends at a transverse fault. More recent investigations, however, have uncovered a distinct break in the basin ~13 km north of Stop 2, just west of Orange Court House. Whether this relatively small area, separate from the main basin, will continue to be considered as a part of the Culpeper Basin or will assume a new name in the future has not been determined at this writing. The rocks underlying Stop 2 are known as the Bull Run shale (Roberts, 1928) or Bull Run Formation (Lindholm, 1979). They exhibit the characteristic color of most sedimentary rocks found in the Mesozoic basins throughout the eastern states. It has been variously described as red, pinkish-red, or pinkish purple and, once seen, the color can be recognized instantly. Lithologically, the rocks at Stop 2 are predominately shales and siltstones. Fresh and even slightly weathered exposures exhibit white specks of plagioclase feldspar, and desiccation cracks are common. Early dinosaur tracks have been found at a number of exposures farther north within the basin. Soils Just north of the intersection of Rt. 20 and Rt. 33, we will meet soils that have been mapped as part of the Penn and Bucks soil series (Table 4). Penn soils are afka fine-loamy, mixed, superactive, mesic Ultic Hapludalfs.

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Order→Alfisol (clay-rich, high base saturation) Suborder→Udalf (udic soil moisture regime) Great group→Hapludalf (insufficiently distinguished [morphologically or chemically] to merit classification as another great group) Subgroup→Ultic Hapludalf (sufficiently leached that base saturation is lower than for a Typic Hapludalf). Penn surface horizons are dark reddish brown (5YR 3/3), silt loams, and have relatively low base saturation. Subsurface horizons show increases in gravels (>2 mm) or channers, as well as a pronounced color shift to redder hues (2.5YR 4/4) that reflect the underlying parent material, often reddish shale, siltstone, or finegrained sandstone of Triassic–Jurassic age (145–250 m.y. B.P.). Penn soils are closely related to the Bucks soil series, differing primarily in terms of depth to bedrock (Penn <1 m) and base saturation (Penn > Bucks). (Bucks soils are afka as fine-loamy, mixed, active, mesic Typic Hapludults.) Across the Piedmont, Penn soils cover 367,000 acres, while the Bucks soils cover 65,000 acres; across Orange County, Penn soils comprise 1800 acres (0.8%), while Bucks soils comprise 3700 acres (1.7%) (U.S. Department of Agriculture, 1971). Land Use Soils of the Bucks and Penn series are the most commonly mapped units within the Mesozoic basins of Virginia. They are also among the most productive soils and often occur together in close proximity on the landscape. Although the Orange County Soil Survey maps Stop 2 as Bucks, our experience has shown that we may well encounter Penn. Both Bucks and Penn soils are intensively farmed throughout the rural portions of the Culpeper Basin. The land use pattern in Figure 9 clearly shows the preponderance of clear land in a band extending to the northeast from Stop 2. Several soil attributes account for this intensive utilization.

TABLE 4. SELECTED PROPERTIES OF THE PENN AND BUCKS SOIL SERIES Horizon

Lower depth (cm)

Colors

Textures

Other

Penn soil series Ap 25 Bt1 45 Bt2 57 Bt3 66 Cr 84 R 84+

5YR 3/3 5YR 5/4 2.5YR 4/4 2.5YR 3/4 10R 4/3 10R 3/3

Silt loam Silt loam Silt loam Channery loam Very channery loam (Rock)

5% channers (>2 mm); slightly acid 8% channers; slightly acid; few clay films 10% channers; moderately acid; few clay films 20% channers; moderately acid; few clay films 40% channers; moderately acid; very few clay films 95% angular flagstones

Bucks soil series Ap 20 BA 45 Bt1 75 Bt2 105 2C 126 2R 126+

10YR 4/4 7.5YR 4/4 5YR 4/4 2.5YR 3/4 2.5YR 3/4 2.5YR 3/2

Silt loam Heavy silt loam Heavy silt loam Silt loam Shaly silt loam (Fractured shale)

Strongly acid Very strongly acid Very strongly acid; discontinuous clay films Very strongly acid; discontinuous clay films 35% shale fragments; strongly acid

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Topographically both series occupy nearly level to gently sloping landscapes where erosion is limited. They are well drained and sufficiently deep to bedrock to allow good root development. The Penn series, being an Alfisol, has, by definition, a relatively high base saturation. Although the Bucks series is classified as an Ultisol, the presence of fresh or only partially weathered plagioclase feldspar in the bedrock and subsoil provide nutrients and acid neutralization. The Orange County Soil Survey states: “These (Bucks Series) are among the best soils for farming in Orange County.” Although timber trees thrive on Bucks and Penn soils, forest acreage tends to be limited because the land is usually cleared for other uses. As often is the case with good agricultural soils, the Bucks and Penn soils are also well suited for engineering and urban development. Consequently, pressures to build on these soils are high, and many highly productive farms throughout the Culpeper Basin are disappearing under concrete and asphalt. Miles 15.6 19.2 21.2

Proceed north on Rt. 20 to right on Rt. 639. Proceed east on Rt. 639 to Rt. 15 (by Montpelier, home of James Madison). Turn left on Rt. 15 and proceed to Agriculture Experiment Station.

Stop 3. Northern Piedmont Agricultural Research and Extension Center (NPAREC) Local Geology Stop 3 is located on the eastern limb of the Blue Ridge Anticlinorium near the eastern edge of the Blue Ridge Province (Fig. 2). Here the bedrock is the Catoctin Formation, comprised of a thick series of basaltic lava flows subsequently metamorphosed to greenstone. Thickness of the formation is reported by Nelson (1962) to exceed 7 km at the Orange County line 12 km to the southwest of Stop 3, while thickness values reported by others working farther south are somewhat less. Greenstone is the dominate rock type within the Catoctin Formation and in the vicinity of Stop 3. Fresh bedrock is mainly a grayish-green to dark-yellowish green, fine grained, somewhat schistose, chlorite and actinolite-bearing lithology with common epidosite segregations. Rocks displaying amygdaloidal features

usually occur at the tops of the individual flows where vesicles are filled with an assortment of minerals. White quartz, jasper, epidote, and pink orthoclase feldspar fillings are common. These rocks are much sought after by rock and mineral collectors who saw and polish them producing striking display specimens. Unfortunately, for any collectors along on this trip, amygdaloidal zones have not been described at this specific locality. Associated and interbedded with the metabasalt are conformable beds of metasedimentary rocks. Impure quartzites and arkosic sandstones appear to be the most common of these. However, metasiltstone and phyllite are also reported (Rader and Evans, 1993). These units range from a few cm to 100 m in thickness but are usually less than 20 m. Most occur as lenses with limited horizontal extent. A conglomerate unit up to 400 m thick and containing greenstone clasts was described by Nelson (1962) as occurring within the Catoctin in Albemarle County, 12 km to the southwest of Stop 3. A number of other geologists (Furcron, 1939; Espenshade, 1986; Kline et al., 1990) have described greenstone breccia within the Catoctin. Finally Lambeth (1901) described “alaskite” dikes intruding the Catoctin in the vicinity of Thomas Jefferson’s Monticello some 40 km to the southwest. These dikes composed of nearly pure microcline have a striking pink color that weathers to orange. One of the authors (Sherwood) determined the foundation and basement walls at Thomas Jefferson’s Monticello to be constructed of this stone. Soils Just south of the town of Orange along Rt. 15 lies the NPAREC. Here, we will meet a soil that has been mapped as part of the Davidson soil series, afka a fine, kaolinitic, thermic Rhodic Kandiudult (Table 5). Order→Ultisol (clay-rich, low base saturation) Suborder→Udult (udic soil moisture regimes) Great group→Kandiudult (distinguishable chemically by low-activity clays, meaning the clay cation exchange capacity is <16 cmol+ kg clay-1) Subgroup→Rhodic Kandiudult (very red). Davidson surface horizons are typically dark red (2.5YR, 5YR), clay loams, dominated by kaolinite, and very acid, and thus, have very low base saturation (<13%). Subsurface horizons are dark red (10R 3/6), and occasionally contain stone lines. Davidson soils form from rocks high in ferromagnesian minerals

TABLE 5. SELECTED PROPERTIES OF THE DAVIDSON SOIL SERIES Lower depth Colors Textures Other (cm) Ap 18 5YR 3/3 Loam Strongly acid 4+ Bt1 30 2.5YR 3/6 Clay loam Strongly acid; few clay films; few black concretions (Mn ) 4+ Bt2 58 10R 3/6 Clay Strongly acid; few clay films; few Mn Bt3 132 10R 3/6 Clay Strongly acid; common clay films Bt4 180 2.5YR 3/6 Clay Strongly acid; many clay films; mottles Note: The term “mottles” was replaced in 1995 by "redox features"; more specifically, the term "depletions" is used for gray areas and "concentrations" for red, yellow, or brown areas. Horizon

Virginia Piedmont and Blue Ridge (e.g., greenstone); bedrock is typically >1.8 m (Genthner, 1990). Davidson soils comprise ~553,000 acres across the Piedmont, and 23,000 acres (10%) in Orange County (U.S. Department of Agriculture, 1971). Land Use The soils formed from the Catoctin greenstone have long been considered to be among the most desirable in the Blue Ridge and Piedmont of Virginia. Many of the earliest land patents from the King of England were underlain by the Catoctin Formation. Figure 9 shows intense use of the land on either side of Southwest Mountain where the Davidson and related greenstone soils are free of forest cover. This is strong evidence that the early settlers recognized the greenstone-derived soils. A list of the antebellum families living over the Catoctin contains many FFVs (First Families of Virginia). Even today some of the most beautiful estates in Virginia are located here. The homes of four United States Presidents—Jefferson, Madison, Monroe, and Taylor—are all located over greenstone soils within 45 km of this stop. Although many of these homes have commanding views of the surrounding countryside, doubtless the principal draw was the productivity of the greenstone-derived soils that underlie the flanks of Southwest Mountain. Only the steepest slopes were unsuitable for crop production. The Davidson, Rabun, and Fauquier series that form from weathered greenstone are all deep, well drained, and productive. A state soil scientist recently measured over 6 m of B horizon at a site underlain by Davidson soil only 6 km south of Stop 3. During the colonial period and even after, much of this acreage was devoted to crops, particularly tobacco, corn, and wheat. The Orange County Soil Survey estimates corn, wheat, alfalfa, and pasture yields from Davidson soils as among the highest in the County. Today pastures and hay production predominate. The same properties that make the Davidson, Rabun, and Fauquier soils excellent for agriculture apply to tree growth. Hardwoods, particularly poplars and oak species grow rapidly in these soils. The Landmark Forest at Montpelier just over the ridge to the west of this stop has been essentially undisturbed for ~100 yr. It contains impressive specimens of the forenamed trees. Tulip poplars with 1.5-m diameters at breast height and oaks with 1.2-m diameters are common there. Greenstone-derived soils also have few limitations for a range of nonagricultural uses. Their deep profiles and welldrained properties are ideal for conventional on-site wastewater systems such as those involving septic tanks and drain fields. Consequently, development pressures for single homes, subdivisions, and commercial enterprises are great. The fact that most of the land over the greenstone is in “strong hands” (i.e., families with a tradition of far-sighted management of the land) may be a factor that limits runaway development over these soils. Davidson soils also exhibit more than adequate engineering properties (Parker et al., 1983). Despite high silt and clay con-

43

tent that classifies them as A-7–5 in the American Association of State Highway and Transportation Officials (AASHTO) system, and as MH in the Unified system, the clay fractions are high in kaolinite. Kaolinite-rich soils are the most stable for engineering uses of the clay-rich soils. They exhibit low shrink-swell potential and moderate optimum moisture, shear strength, liquid limits, plasticity indices, and California bearing ratio (CBR) values. Consequently these soils are commonly used for fill material and other engineering purposes. Miles 23.2 26.3 27.2

Proceed south on Rt. 15 to left (east) on Rt. 639. Turn left on Rt. 643. Proceed to the intersection of Rt. 643 and 638.

Stop 4. The Flat Woods Local Geology Stop 4 is located in the western Piedmont (see Fig. 4). After leaving Stop 3 we proceeded southward parallel to the strike of the Catoctin Formation, then turned eastward at Madison Run. At that point we left the Catoctin and crossed the Mountain Run fault and the Candler Formation. The Candler Formation, of Cambrian age, lies stratigraphically above the Catoctin. Major rock types within the Candler are phyllite, metasiltstone, metatuff, and dolomitic marble. At the top of the formation at this point is the Everona Limestone. Mack (1965) describes the Everona as ~55 m of dark-blue, thin-bedded, slatey limestone containing thin stringers of calcite. Weathering of the limestone has created a distinct, narrow valley that can be traced for several kilometers along the strike of the formation. Note we cross through the valley just east of the railroad. As is usually the case with limestone, the narrow valley has been cleared and farmed since the colonial period (see Fig. 10). The limestone was also mined at several sites during the mid 1800s for the production of hydrated lime. The bedrock at Stop 4 overlies the Candler Formation. The Virginia State Geologic Map (Rader and Evens, 1993) refers to these rocks as “Stratified rocks of the western Piedmont.” Lithologically they are a complex of metagraywacke, quartzose schist, phyllite, and mélange. The metagraywackes have been highly sheared and altered to quartzose chlorite and biotite schist with some blue quartz. According to Pavlides (1980) these rocks grade upward into a sequence of metavolcanic and metasedimentary rocks correlated with the Chopawamsic Formation of northern Virginia. Soils Here we will meet a soil that has been mapped as part of the Tatum soil series, afka a fine, mixed, semiactive, thermic Typic Hapludult (Table 6). Tatum surface horizons are typically brown (7.5YR 4/4), silt loams, contain 25% quartz and sericite schist channers, and have low base saturation. Subsurface horizons are much redder (2.5YR 5/6) clay loams. Tatum soils typically overlie sericite, schist, phyllite, or other fine-grained metamorphic

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rock. Nason soils are very closely related to Tatum soils, with an identical classification, but with the principal difference being subsoil horizons are less red (5YR to 10YR; Table 7). Tatum and Nason soils together cover 1.15 million acres (0.72M and 0.43M acres, respectively) of the Piedmont, and 24,000 (11%) and 45,000 (20%) acres, respectively, of Orange County (U.S. Department of Agriculture, 1971). Land Use Tatum and Nason soils cover vast areas of the western and central Piedmont of Virginia. In Orange County alone, the two series comprise >30% of the total land area with about three quarters of the land area in forests (Fig. 9). Acid tolerant vegetation predominates, with chestnut oaks, Virginia pine, black gum, and blueberries common. During the colonial period and even up to the Civil War most of these soils were cleared and farmed, principally for tobacco. However, due to their acidity, low natural fertility, and highly erodible nature, most of the land quickly became exhausted and was abandoned. The Piedmont, in fact, has been the locus of considerable research on past soil erosion rates (e.g., Ireland et al., 1939; Trimble, 1985). In 1995, it was suggested that current soil erosion rates were unsustainable relative to soil production rates and expensive to remedy (Pimentel et al., 1995). Follow-on letters and studies from economists (e.g., Crosson, 1995) led to an exchange of letters, which triggered yet another response from Trimble and Crosson (2000), titled “U.S. soil erosion rates— Myth or reality?” In this article, Trimble and Crosson (p. 250) wrote “We do not seem to have a truly informed idea of how much soil erosion is occurring in this country, let alone of the processes of sediment movement and deposition.” Recent studies have estimated that farmland denudation is occurring at rates of between ~600 and ~4000 m/m.y. (Wilkinson and McElroy, 2007; Montgomery, 2007), and that these rates are approximately two orders of magnitude larger than soil production rates. It will be important on this field trip to relate current land uses, including abandonment following attempts at agriculture, to soil properties. During the nineteenth century, as the soils were exhausted large numbers of the settlers moved westward in their quest for “new land.” The worn-out and heavily eroded fields were left to nature, where the old field succession of Virginia pine, black locust, blackberries, and broom sedge followed. Within a few years the Virginia pines came to dominate. Even today, fields abandoned during the twentieth century contain thick stands of this species. However, Virginia pines are shade intolerant and relatively shortlived trees. After about seventy years the pines begin to die off and young shade-tolerant hardwoods such as oaks, hickory, poplar, maple, and ash take over. Today, hardwood forests predominate over Tatum and Nason soils in the Virginia Piedmont. In the twenty-first century, with readily available lime and fertilizers, some Tatum and Nason soils can be successfully farmed, and this is the case on a modest scale in Orange County. However, stringent management practices, particularly to control erosion, are required.

Tatum and Nason soils are normally deep and well drained so they usually pass the “perc” test required for conventional septic-tank and drain-field waste water systems. These properties together with lax local zoning regulations and low land prices have resulted in an explosion of strip development along the secondary roads in the “flat woods” area. This condition will be evident as we proceed to our next stop. Thirty years ago these roads traversed virtually unbroken forests. Miles 31.3

Proceed east on Rt. 638 to bridge over Cooks Creek.

Stop 5. Floodplain of Cooks Creek Local Geology The bedrock at Stop 5 lies within the Copowamsic volcanic belt. As noted earlier, this belt is one of the allochthonous terranes making up the western and central portions of the Virginia Piedmont. Pavlides (1990) mapped the rocks here as part of the Malange Zone III of the Mine Run Complex, probably of Ordovician age. Pavlides recognized three distinct malange zones within the Mine Run Complex. While the matrix rocks are predominately schists and phyllites, the zones are differentiated on the basis of the degree of deformation and the compositions of the included blocks of other lithologies. In Zone III many of the matrix rocks are highly deformed and contain abundant euhedral magnetite. Exotic blocks of mafic composition include amphibolite, ultramafics, serpentinite, and talc. Some blocks contain more than one rock type. Nonmafic blocks composed of biotite gneiss are also present. Despite the interesting bedrock in this area, the principal purpose of Stop 5 is to examine the geomorphology (Fig. 9). At this point, Cooks Creek, a tributary to the North Anna River, has abandoned its former channel and moved ~100 m to the south. In so doing, the stream has left a well-defined abandoned channel and a nearly level wetland floodplain that sustains standing water in wet sessions. It also left a stretch of cut bank that resembles a large amphitheater. Soils Where Rt. 638 crosses Cooks Creek, we will visit a floodplain soil that is framed by an amphitheater with ~20 m of relief, reflecting millennia of incision. This Wehadkee soil series, afka a fine-loamy, mixed, active, nonacid, thermic Fluvaquentic Endoaquept, has grayish-brown (10YR 5/2) surface horizons, with some mica flakes, and moderate base saturation (Table 8). Some redoximorphic features are evident in subsurface horizons, and the irregular decrease in organic carbon with depth leads to the “Fluvaquentic” subgroup classification. Wehadkee soils develop from sediments derived from schist, gneiss, granite, phyllite, and other metamorphic and igneous rocks. For this particular soil at this stop, the parent material represents material transported by Cooks Creek to this location. The Wehadkee series comprises

Oi A E Bt1 Bt2 Bt3 Bt4 C Cr

Horizon

Cg

Ap Bg1 Bg2

Horizon

Lower depth (cm) 3–0 3 23 38 51 71 97 127 157

127

Lower depth (cm) 20 43 102 10YR 6/1

10YR 5/2 10YR 4/1 10YR 6/1

Sandy loam

Fine sandy loam Loam Sandy clay loam

Other

(Deciduous forest litter) Strongly acid Very strongly acid 5% quartz gravels; very strongly acid; few clay films 5% quartz gravel; very strongly acid; few clay films 5% schist gravels; very strongly acid; common clay films; mottles 25% schist gravels; very strongly acid; common clay films; mottles 25% schist gravels; very strongly acid; mottles

Few mica flakes; moderately acid 3+ Few mica flakes; moderately acid; common Fe masses 3+ Common mica flakes; moderately acid; common Fe masses 2+ Common mica flakes; moderately acid; common Fe masses; 3+ prominent Fe masses

TABLE 8. SELECTED PROPERTIES OF THE WEHADKEE SOIL SERIES Colors Textures Other

(Peat) Silt loam Silt loam Silty clay loam Silty clay Silty clay Channery silty clay loam Channery silt loam saprolite Weathered, fractured sericite schist

TABLE 7. SELECTED PROPERTIES OF THE NASON SOIL SERIES Textures

(Litter/twigs) 10YR 3/2 10YR 5/4 10YR 5/8 7.5YR 5/8 5YR 4/8 5YR 4/6 5YR 4/6, 2.5YR 5/6, 7.5YR 5/6

Colors

TABLE 6. SELECTED PROPERTIES OF THE TATUM SOIL SERIES Lower depth Colors Textures Other (cm) Oe 5–0 (Litter/twigs) (Mucky peat) (Eve rgreen litter) Gravelly silt loam 25% qpssc; 5% cobbles (75–250 mm); very strongly acid A 10 7.5YR 4/4 Bt1 33 5YR 4/6 Gravelly silty clay loam 18% qpssc; very strongly acid; few clay films Bt2 79 2.5YR 5/6 Silty clay loam 5% qpssc; strongly acid; common clay films BC 107 5YR 5/8 Silty clay loam Very strongly acid; mottles C 137 5YR 6/4, 5YR 5/8 Channery silt loam saprolite 20% ssc; strongly acid; mottles Cr 157 Highly fractured sericite schist Strongly acid R 157+ Unweathered, slightly fractured sericite schist Note: Qpssc—quartz pebbles and sericite schist channers; ssc—sericite schist channers.

Horizon

Virginia Piedmont and Blue Ridge 45

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~655,000 acres of the Piedmont, 62,000 of those in Virginia (1%), and 1300 (0.6%) of those in Orange County (U.S. Department of Agriculture, 1971). Land Use Like most of the land over Wehadkee soils in Orange County, Stop 5 is forested. Wetland tolerant tree species common here are river birch, sycamore, alder, maple, iron wood, and willow oak. In addition, fine specimens of tulip poplar are present. Poplar, a desirable timber species, can grow rapidly and attain immense size under the conditions found here. Noncanopy vegetation includes spice bush, button bush, and ferns. While tree growth can be rapid in Wehadkee soils, harvesting can present problems during wet seasons. Many forms of wildlife such as deer, muskrat, raccoon, wild turkey, and a variety of other birds flourish in areas with Wehadkee soils. Surprisingly, some 40% of the Wehadkee soils in Orange County have been cleared and used for pasture. The carrying capacity of these soils, at 80 cow-acre-days per year ranks about midway when compared to other soils within the County. However, during very dry years, pastures on Wehadkee soils are desirable, providing good forage when the grasses on the upland soils do not. A very small acreage over the Wehadkee soils in Orange County is devoted to growing crops, mainly corn. Again these fields can aid farmers during dry years, but generally cropping on Wehadkee soils is not recommended. Because of the seasonally high water table and susceptibility to flooding, Wehadkee soils are not recommended for development. Miles 32.6 34.7 38.8

Proceed east on Rt. 638 to intersection with Rt. 612, turn right on Rt. 612. Proceed east on Rt. 612 to intersection with Rt. 669, turn left on Rt. 669. Proceed north on Rt. 669 to Stop 6, 0.2 miles north of Lahore.

Stop 6. Lahore Local Geology Stop 6 is over the Lahore pluton (see Fig. 2). The pluton was named by Pavlides (1990) for the small village of Lahore located immediately to the south of our stop. It was intruded into malange zone III during the Ordovician. It is one of a number of mafic intrusions emplaced within the Central Virginia volcanicplutonic belt. Pavlides (1990) recognized three distinct lithologies within the pluton. (1) Amphibole monzonite, mesocratic, medium-grained amphibole monzonite, and amphibole-quartz monzonite. A foliation is defined by the alignment of tabular feldspar crystals. (2) Pyroxene monzonite with color ranging from dark gray to black. The rock is massive to weakly foliated and con-

sists of large augite and plagioclase grains (some zoned) and opaque oxides. (3) Mafic and ultramafic rocks consisting of partially serpentinized pyroxenite and diopside. Stop 6 is located over lithology type 1—amphibole monzonite. The landscape over the Lahore pluton is very gently rolling to nearly level. This topography is in subtle contrast to that in the surrounding countryside, where hill slopes are somewhat steeper and stream valleys are narrower and more pronounced. The average land elevation over the pluton is slightly lower than that surrounding the intrusion. This difference is even more pronounced over the famous Green Springs intrusion, located ~18 km southwest of Lahore. Because of its slightly lower elevation, the Green Springs area is often referred to as the Green Springs basin. Soils Just north of Lahore Road off Rt. 669, we will meet a soil mapped as the Iredell series, afka a fine, mixed, active, thermic, Oxyaquic Vertic Hapludalf (Table 9). Iredell surface horizons are typically dark grayish-brown (2.5Y 4/2), sandy loam, with few gravels, and with moderate base saturation. Subsurface horizons show strong increases in shrink-swell–prone clays (presence of slickensides) as well as base saturation. In the Virginia Piedmont, Iredell soils are most closely associated with mafic plutons, such as diabase, monzonite, diorite, and gabbro. Iredell soils cover ~231,000 acres across the Piedmont, and 427 acres (0.2%) of Orange County (U.S. Department of Agriculture, 1971). Unlike any of the other soils on this field trip, some of the soil pH values are >7.9, an unusual feature for soils in humid climates with long residence times. The taxonomic classification of Iredell soils provided above is specific to a location, per the official series description, 71 m “north of a fire hydrant, across the road from the Southside School, [1.6 km] south of Chester, South Carolina, along U.S. Highway 72.” This classification indicates the mineralogy is mixed, a characterization that is also likely to apply to our specific Iredell soil (Plaster and Sherwood, 1971), since several clay minerals were found in the B horizon (Fig. 11). Note that while montmorillonite increases sharply in the C horizon, this increase is only relative to other clay minerals. In fact, total clay shows a precipitous decline below the argillic B horizon, so it is possible that there could have been more montmorillonite present in the B horizon than in the C horizon, because of the >5-fold difference in total clay content (Fig. 11). Not too surprisingly for an Alfisol, there is a pronounced increase in clay with depth, followed by a sharp decrease with depth (Fig. 11A). From the A2 horizon (sample I-1, 25–37 cm) to the B horizon (samples I-2 and I-3, 37–91 cm) to the C horizon (samples I-4 through I-6, 91–183 cm), clay-sized material (<2 μm) changed from 18% to 66% to 12%. Just as striking as this clay bulge is the shift in clay mineralogy (Fig. 11B), from clays dominated by quartz and vermiculite, to illite in the B horizon, to montmorillonite in the C horizon. The dominance of clay-sized quartz at the surface of the profile was attributed to

Virginia Piedmont and Blue Ridge

47

TABLE 9. SELECTED PROPERTIES OF THE IREDELL SOIL SERIES Horizon Ap1 Ap2 Btss1 Btss2 Btg BC C1 C2 C3

Lower depth (cm) 13 18 28 51 61 69 81 112 157

Colors

Textures

2.5Y 4/2 10YR 4/2 10YR 4/3 10YR 4/3 2.5Y 4/2 5Y 4/3 fmdgg, vpb, and yb (m) fmdgg, vpb, b, and yb (m) fmdgg, yb, b, and vpb

Sandy loam Loam Clay Clay Silty clay Loam Loam Sandy loam Sandy loam

Other 4+

1% fine pebbles; slightly acid; few black concretions (Mn ) 4+ Neutral; few Mn 4+ Slightly acid; common clay films, slickensides; many Mn 4+ Neutral; common clay films, slickensides, Mn ; few weathered feldspar crystals 4+ Neutral; common clay films, slickensides, Mn ; few weathered feldspar crystals Common saprolite; neutral; common clay films; mottles 80% saprolite; neutral; few clay films 90% saprolite; moderately alkaline; few clay films 90% saprolite, 10% hard rock fragments; moderately alkaline

Note: b—black; fmdgg—finely mottled dark greenish gray; m—mottled; vpb—very pale brown; yb—yellowish brown.

intense leaching (eluviation) that is characteristic of a udic soil moisture regime (Plaster and Sherwood, 1971) and very long soil residence times. Plaster and Sherwood (1971), in addition to characterizing the particle size distribution and clay mineralogy, also examined the elemental composition of a representative Iredell soil profile and its hornblende metagabbro parent material (Table 10). The authors then calculated two weathering indices: first, the weathering potential index (WPI; the percentage molar ratio of ΣCaO, MgO, K2O, Na2O [ΣBC] minus water to ΣBC plus SiO2, Al2O3, and Fe2O3); and second, the base:aluminum (BA) ratio of ΣBC to the molar value of Al2O3. The weathering potential index (Rieche, 1950) has been modified (Short, 1961), and is only one of a great number of weathering indices that have been developed (Birkeland, 1999).

A

Cross-comparisons between weathering indices generally produce comparable results; for WPI, specifically, unweathered materials generally have strongly positive values (Birkeland, 1999: table 3.6, p. 70), while weathered material can have negative values. Iredell WPI values were similarly positive between gabbro parent material (+19%) and the C horizon composed of saprolite (+16%). WPI values decreased sharply in the B horizon (−36%) and A horizon (−24%), a result attributed in part to the sharply reduced hydraulic conductivity of the clayey B-horizon, which was dominated by 2:1 clays. In effect, as mafic primary minerals weather to secondary minerals, vertical flow paths are effectively rerouted laterally, in effect drying out the lowermost part of the profile. This hydrologic dependence on lithology played a somewhat analogous role in explaining the difference in saprolite

B

Figure 11. Clay, silt, and sand content (A) and clay mineralogy (B) for a representative Iredell soil profile (Plaster and Sherwood, 1971).

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Sherwood et al. TABLE 10. CHEMICAL COMPOSITION (ALL VALUES %) OF A REPRESENTATIVE IREDELL SOIL, WITH TWO WEATHERING INDICES: WEATHERING POTENTIAL INDEX (WPI) PERCENT AND THE BASE:ALUMINUM (BA) RATIO Lower SiO2 Al2O3 Fe2O3 CaO MgO Na2O K2O LOI Total WPI BA (cm) A2 37 54.40 10.40 23.89 1.43 0.00 2.60 0.67 6.65 100 –23.8 0.80 B 91 49.92 21.33 14.93 1.88 1.34 2.06 0.68 9.31 101 –35.8 0.48 C 183 45.89 17.07 10.67 7.90 10.09 1.61 1.55 3.64 98 16.1 2.75 Rock 46.37 16.67 13.70 9.45 7.47 2.05 1.60 2.41 100 19.3 2.56 Note: From Plaster and Sherwood (1971); Plaster (1968). LOI—loss on ignition.

thickness over felsic versus mafic lithologies. BA ratios showed a comparable pattern. Plaster and Sherwood (1971) raised two questions that we copy here for the purposes of stimulating discussion: (1) Why is there such a dramatic shift in montmorillonite from the B to C horizons? And, (2) Where has the talc identified in the parent material gone? With regards to the depth profile of montmorillonite, Eades (1953) suggested that fine-grained montmorillonite could be eluviated from the B horizon and illuviated into the C horizon. These elemental depth profiles produced at least two unexpected results. First, it was noted that CaO and MgO might typically be expected, as biologically cycled bases, to show a decrease, not increase, with depth. One possible explanation is that the A2, versus the A1, horizon may not have captured the tree-pumping signal. Second, K2O:Na2O ratios decreased sharply with depth, whereas the authors expected to see no increase in Na2O with depth. This result is consistent with a hydrologic rerouting following in situ formation of low-conductivity clays (see Table 1, which does not account for differences in clay mineralogy). For example, illite, a clay that displays a strong affinity for K, decreases with depth. Another factor may be the addition of K-rich fertilizers at the surface over many decades to increase crop yields. The WPI values for the Iredell soil show that the saprolite (sample I-5) has experienced little chemical weathering, which Plaster and Sherwood (1971) attributed to the “high concentration of hydrophyllic clays in the B horizon”; the “impermeable nature of this clay-rich zone has hindered the downward migration of water and stands as the major reason for the lack of chemical weathering” of the saprolite (p. 2824). Land Use Today, most types of land development in areas underlain by Iredell and related “blackjack” soil series—a catch-all term for Jackland, Whitestore, Orange, Zion, and other high shrinkswell clayey soils—is generally discouraged. Several characteristics of these soils contribute to this policy. First, most of the land will not “perc” due to poor drainage and a seasonally high water table. Second, soils over mafic intrusions usually contain significant amounts of high shrink-swell smectitic clays such as montmorillonite. The taxonomy of the closely related Orange series (fine, smectitic, mesic Albaquic Hapludalfs) indicates a 2:1 shrink-swell clay like montmorillonite makes up over 50% of the

clays present. Shrink-swell soils are highly unstable for engineering uses such as foundations, fills and other types of construction. The question then arises, are these soils suitable for agriculture? Tables in the Orange County Soil Survey listing crop yields for the various soil series rate the Orange-Iredell complex in the lowest echelons for corn, wheat, and hay, and low to medium for pasture. Interestingly, the tables give higher yield figures for these same soils on 2%–7% slopes than on 0%–2% slopes. Normally, for most soils the more level slopes are more productive because they are less susceptible to erosion. The reason for these seemingly illogical values can be summed in a single term—drainage! Level landscapes underlain by Iredell and Orange soils exhibit such poor drainage that crops suffer from too much water. Moderate slopes improve drainage and runoff and provide better conditions for crop production. An ancillary problem caused by smectite-rich soils can be harvesting during wet years. When these soils are disturbed while wet, structure is destroyed and they turn to a viscous liquid. It is not unusual for tractors to become mired up to their axles when these conditions occur. Finally, drainage affects chemical weathering processes. As Pavich (1986, p. 587) has noted, “the functional relationships of regolith production and erosion must begin with consideration of the hydrologic processes operating within the soil. The soil acts to partition rainfall into evapotranspiration, runoff, and recharge to the saprolite…. Since water movement is dependent on rock structure, rock weathering rate may be more dependent on soil water balance and rock structure than mineral dissolution kinetics if the rock contains at least one mineral phase that reacts rapidly with dilute, acidic solutions.” Examining the land use patterns in Figure 10, it is evident that despite the latent problems associated with the Iredell soil, the trends are clear. Land underlain by mafic igneous rocks and Iredell or similar soils has been preferentially cleared and farmed over the years. Areas underlain by Tatum and Nason soils are largely forested. The reason can be traced back to the colonial period of Virginia history. As discussed at Stop 4, settlers whose land was over highly acid and erodible soils such as the Tatum and Nason, literally wore out the land in a few short years. A combination of the demands of tobacco production and poor farming methods were largely responsible. A settler to this area, John Craven, wrote in 1833 “… the whole face of the country presented a scene of desolation that baffles description, farm after farm had been worn out, and washed and gullied, so scarcely an acre could be found in

Virginia Piedmont and Blue Ridge place fit for cultivation.…” Even George Washington described the existing farming methods as “ruinous” and visiting English farmer William Strickland noted “...Virginia farms of the area were much worn out … nearly exhausted … and tobacco and maize … a curse” (Strickland, 1801). As noted earlier, these conditions led to a massive outmigration of settlers from the Virginia Piedmont and throughout the southern states and the abandonment of hundreds of thousands of acres of degraded land. Over time, these lands slowly returned to forest. Even today most of these lands remain forested. On the other hand, many of the early settlers over the Iredell, Orange, and related soils formed over the mafic intrusions remained on the land. While the heavy clay soils are difficult to work, they are considerably less erosive and are characterized by naturally high pH and nutrient retention. These properties allowed the land to be successfully farmed for generations, and many families to become relatively prosperous. A number of these families remain on the land today. These areas lack forest cover and appear as light gray in Figure 10. The relationship between the cleared (nonforested) areas and the areal extent of the mafic intrusions is striking. More than one field geologist has noted the contact between the mafic intrusions and the surrounding metasedimentary rocks can be located by following the tree line on the topographic maps of the area. ACKNOWLEDGMENTS We thank the following colleagues for their reviews of early editions of this manuscript: Lee Daniels, Greg Hancock, Louis Heidel, Dan Richter, Dave Starner, Aaron Thompson, and Rich Whittecar. Dot McLaren kindly provided assistance with several of the illustrations. Landowners in this area have been particularly generous over the years with granting us access to their properties. We are appreciative of over thirty years of students’ and colleagues’ critiques of portions of this field trip. REFERENCES CITED Allen, R.M., 1963, Geology and mineral resources of Greene and Madison Counties: Virginia Division of Mineral Resources Bulletin 78, 102 p. Bailey, C.M., 2006, The geology of Virginia: A resource for information, photographs, maps and diagrams of the geology of the Commonwealth, http:// web.wm.edu/geology/virginia/ (accessed December 2009). Birkeland, P.W., 1999, Soils and geomorphology: New York, Oxford University Press, Third edition, 430 p. Buol, S.W., ed., 1973, Soils of the southern states and Puerto Rico: Southern Cooperative Series Bulletin 174, 105 p. Carroll, L., 1871, Through the Looking-Glass, and What Alice Found There: London, Macmillan. Carter, J.B., et al., 1971, Soil survey of Orange County, Virginia: Washington D.C., U.S. Department of Agriculture, Soil Conservation Service. Cornet, B., 1977, The palynostratigraphy and age of the Newark Supergroup [Ph.D. thesis]: University Park, Pennsylvania, Pennsylvania State University, 505 p. Costa, J.E. and Cleaves, E. T.,1984, The Piedmont landscape of Maryland: a new look at an old problem: Earth surface processes and landforms; v. 9, p. 59–74. Craven, J.H., 1833, Letter by J.H. Craven, in Craven, A.O., 1925, Soil exhaustion as a factor in the agricultural history of Virginia and Maryland, 1606– 1860: Urbana, Illinois, University of Illinois, 184 p.

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Crosson, P., 1995, Soil erosion and its on-farm productivity consequences: What do we know?: Science, v. 269, p. 461, doi: 10.1126/science.269.5223.461. Darton, N.H., 1951, Structural relations of Cretaceous and Tertiary formations in parts of Maryland and Virginia: Geological Society of America Bulletin, v. 62, p. 745–780, doi:10.1130/0016-7606(1951)62[745:SROCAT]2.0.CO;2. Dicken, C.L., Nicholson, S.W., Horton, J.D., Kinney, S.A., Gunther, G., Foose, M.P., and Mueller, J.A.L., 2008, Preliminary integrated geologic map databases for the United States: Delaware, Maryland, New York, Pennsylvania, and Virginia: Washington, D.C., U.S. Geological Survey OpenFile Report 2005-1325, version 1.1 (August 2008) (http://pubs.usgs.gov/ of/2005/1325/). Eades, J.L., 1953, A study of the mineralogy of clays in soils as a product of the parent material [M.S. thesis]: Charlottesville, Virginia, University of Virginia, 99 p. Eargle, D.H., 1940, The relations of soils and surface in the South Carolina Piedmont: Science, v. 91, p. 337–338, doi: 10.1126/science.91.2362.337. Eargle, D.H., 1977, Piedmont Pleistocene soils of the Spartanburg area, South Carolina: South Carolina Division of Geology, State Development Board, Geologic Notes, v. 21, p. 57–74. Espenshade, G.H., 1986, Geology of the Marshall Quadrangle, Fauquier County, Virginia: U.S. Geological Survey Bulletin 1560, 60 p. Furcron, A.S., 1939, Geology and mineral resources of the Warrenton quadrangle, Virginia: Virginia Geological Survey Bulletin 54, 94 p. Genthner, M.H., 1990, The variability and geomorphology of Appling, Cecil, and Davidson soils on sideslopes in the Virginia Piedmont [M.S. thesis]: Blacksburg, Virginia, Department of Crop and Soil Environmental Sciences, Virginia Polytechnic Institute and State University, 217 p. Hack, J.T., 1960, Interpretation of erosional topography in humid temperature regions: American Journal of Science, 258-A, p. 80–97. Howard, J.L., Amos, D.F., and Daniels, W.L., 1996, Micromorphology and dissolution of quartz sand in some exceptionally ancient soils: Sedimentary Geology, v. 105, no. 1–2, p. 51–62, doi: 10.1016/0037-0738(95)00133-6. Ireland, H.A., Sharpe, C.F.S., and Eargle, D.H., 1939, Principles of gully erosion in the Piedmont of South Carolina: Washington, D.C., U.S. Department of Agriculture, Technical Bulletin 633, 142 p. Jenny, H., 1941, Factors of soil formation: a system of quantitative pedology: New York, McGraw-Hill Book Company, 281 p. Jonas, A.I., 1927, Geologic reconnaissance in the Piedmont of Virginia: Geological Society of America Bulletin, v. 38, no. 5, p. 837–846. Kerr, W.C., 1881, On the action of frost in the arrangement of superficial earth material: American Journal of Science, Third series, v. 21, no. 125, May, article 44, p. 345–358. Kline, S.W., Conley, J.F., and Evens, N.H., 1990, Hyaloclastic pillow breccia in the Catoctin metabasalt of the eastern limb of the Blue Ridge anticlinorium in Virginia: Southeastern Geology, v. 30, no. 9, p. 241–257. Lambeth, W.A., 1901, Notes on the geology of the Monticello area, Virginia [Ph.D. thesis]: Charlottesville, Virginia, University of Virginia, 22 p. Lindholm, R.C., 1979, Geologic history and stratigraphy of the Triassic–Jurassic Culpeper Basin, Virginia: Geological Society of America Bulletin, v. 90, part 2, p. 1702–1736, doi: 10.1130/0016-7606(1979)90<995:GHASOT> 2.0.CO;2. Mack, T., 1965, Characteristics of the Everona Formation in Virginia: Virginia Division of Mineral Resources Information Circular 10, 16 p. Markewich, H.W., Pavich, M.J., Hall, R.L., Johnson, R.G., and Hern, P.P., 1987, Age relations between soils and geology in the Coastal Plain of Maryland and Virginia: U.S. Geological Survey Professional Paper 1589-A, 34 p. Mixon, R.B., and Newell, W.L., 1977, Stafford fault system: Structures documenting Cretaceous and Tertiary deformation along the Fall Line in northeastern Virginia: Geology, v. 5, p. 437-440. Montgomery, D.R., 2007, Soil erosion and agricultural sustainability: Proceedings of the National Academy of Sciences of the United States of America, v. 104, p. 13,268–13,272, doi: 10.1073/pnas.0611508104. Morison, C.G.T., 1948, The catena and its application to tropical soils: Commonwealth Bureau of Soil Science Technical Bulletin 46, p. 124–130. Nelson, W.A., 1962, Geology and mineral Resources of Albemarle County, Virginia: Division of Mineral Resources Bulletin 77. Owens, J.P., Hess, M.M., Denny, C.S., and Dwornik, E.J., 1983, Postdepositional alteration and near-surface minerals in selected Coastal Plain formations of the Middle Atlantic states: U.S. Geological Survey Professional Paper 1067-F, p. F1–F45. Parker, J.C., Amos, D.F., and Baker, J.C., 1983, Engineering properties of selected soils of the Virginia Piedmont: Blacksburg, Virginia, Department

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of Agronomy, Virginia Polytechnic Institute and State University, Bulletin 83-6, 78 p. Pavich, M.J., 1985, Appalachian Piedmont morphogenesis: Weathering, erosion and Cenozoic uplift, in Morisawa, M., and Hack, J.T., eds., Tectonic Geomorphology: Binghamton, New York, Proceedings 15th Annual Geomorphology Symposium Series, State University of New York, p. 299–319. Pavich, M.J., 1986, Processes and rates of saprolite production and erosion on a foliated granitic rock of the Virginia Piedmont, in Colman, S.M., and Dethier, D.P., eds., Rates of Chemical Weathering of Rocks and Minerals: Orlando, Florida, Academic Press, p. 551–590. Pavich, M.J., Leo, G.W., Obermeier, S.F., and Estabrook, J.R., 1989, Investigations of the characteristics, origin, and residence time of the upland residual mantle of the Piedmont of Fairfax County, Virginia: Washington, D.C., U.S. Geological Survey Professional Paper 1352, 58 p. Pavlides, L., 1980, Revised nomenclature and stratigraphy relationship of the Fredericksburg Complex and Quantico Formation of the Virginia Piedmont: Washington, D.C., U.S. Geological Survey Professional Paper 1146, 29 p. Pavlides, L., 1982, Petrology of the “Lahore” complex and Ellisville pluton, composite granitoid bodies in the Piedmont of Virginia: Washington, D.C., U.S. Geological Survey Professional Paper 1275, 60 p. Pavlides, L., 1990, Geology of part of the Northern Virginia Piedmont: Washington, D.C., U.S. Geological Survey Open-File Report 90-548, map and text. Pédro, G., 1983, Structuring of some basic pedological processes: Geoderma, v. 31, p. 289–299, doi: 10.1016/0016-7061(83)90042-3. Pimentel, D., Harvey, C., Resosudarmo, P., Sinclair, K., Kurz, D., McNair, M., Crist, S., Shpritz, L., Fritton, L., Saffouri, R., and Blair, R., 1995, Environmental and economic costs of soil erosion and conservation benefits: Science, v. 267, p. 1117–1123, doi: 10.1126/science.267.5201.1117. Plaster, R.W., 1968, Geologic, mineralogic, and engineering studies of three Virginia soil profiles [masters thesis]: Charlottesville, Virginia, University of Virginia, 109 p. Plaster, R.W., and Sherwood, W.C., 1971, Bedrock weathering and residual soil formation in Central Virginia: Geological Society of America Bulletin, v. 82, p. 2813–2826, doi: 10.1130/0016-7606(1971)82[2813:BWARSF] 2.0.CO;2. Prowell, D.C., 1976, Implications of Cretaceous and post-Cretaceous faults in the Eastern United States: Geological Society of America Abstracts with Programs, v. 8, no. 2, p. 249–250. Rader, E.K., and Evans, N.H., 1993, editors, Geologic map of VirginiaExpanded explanation: Charlottesville, Virginia Division of Mineral Resources, 80 p. Rankin, D.W., 1975, The continental margin of eastern North America in the southern Appalachians: The opening and closing of the proto–Atlantic Ocean: American Journal of Science, v. 278-A, p. 1–40.

Reusser, L.J., Bierman, P.R., Pavich, M.J., Zen, E., Larsen, J., and Finkel, R., 2004, Rapid late-Pleistocene incision of Atlantic passive-margin river gorges: Science, v. 305, p. 499–502, doi: 10.1126/science.1097780. Rieche, P., 1950, A survey of weathering processes and products: New Mexico University Publications, Geology, 95 p. Roberts, J.K., 1928, The geology of the Virginia Triassic: Charlottesville, Virginia, Virginia Geological Survey Bulletin 39, 205 p. Schoeneberger, P.J., Wysocki, D.A., Benham, E.C., and Broderson, W.D., eds., 2002, Field Book for Describing and Sampling Soils, Version 2.0: Lincoln, Nebraska, Natural Resources Conservation Service, National Soil Survey Center, 174 p. Sherwood, W.C., 2003, Examination of boulders from several anthropogenic features at Montpelier: Unpublished technical report on file with the Head of Archeology, James Madison, Montpelier, 11 p. Sherwood, W.C., and Eaton, L.S., 1993, Soils, geology, and land use in the Virginia Piedmont, p. 51–69, in Field Trip Guidebook: Harrisonburg, Virginia, National Association of Geology Teachers, Southeast Section Meeting, 14–16 May 1993, James Madison University. Short, N.M., 1961, Geochemical variations in four residual soils: The Journal of Geology, v. 69, p. 534–574, doi: 10.1086/626770. Sommer, M., Halm, D., Weller, U., Zarei, M., and Stahr, K., 2000, Lateral podzolization in a granite landscape: Soil Science Society of America Journal, v. 64, p. 1434–1442. Southworth, S., Bailey, C.M., Eaton, L.S., Hancock, G., Lamoreaux, M.H., Litwin, R.J., Burton, W.C., and Whitten, J., 2009, Geology of the Shenandoah National Park Region: Guidebook for 39th Annual Virginia Geological Field Conference (October 2–3, 2009), 40 p. Strickland, W., 1801, Observations on the agriculture of the United States of America: London, Bulmar and Company, 74 p. Trimble, S.W., 1985, Perspectives on the history of soil erosion control in the eastern United States: Agricultural History, v. 59, no. 2, p. 162–180. Trimble, S.W., and Crosson, P., 2000, U.S. soil erosion rates—Myth or reality?: Science, v. 289, p. 248–250, doi: 10.1126/science.289.5477.248. U.S. Department of Agriculture, 1971, Soil Survey of Orange County, Virginia: Washington, D.C., U.S. Department of Agriculture, Soil Conservation Service in cooperation with the Virginia Agricultural Experiment Station, 169 p. Whittecar, G.R., 1985, Stratigraphy and soil development in upland alluvium and colluvium, north-central Virginia Piedmont: Southeastern Geology, v. 26, no. 2, p. 117–129. Wilkinson, B.H., and McElroy, B.J., 2007, The impact of humans on continental erosion and sedimentation: Geological Society of America Bulletin, v. 119, p. 140–156, doi: 10.1130/B25899.1. MANUSCRIPT ACCEPTED BY THE SOCIETY 2 DECEMBER 2009

Printed in the USA

The Geological Society of America Field Guide 16 2010

Magmatic layering and intrusive plumbing in the Jurassic Morgantown Sheet, Central Atlantic Magmatic Province LeeAnn Srogi Tim Lutz Department of Geology and Astronomy, West Chester University, West Chester, Pennsylvania 19383, USA Loretta D. Dickson Department of Geology and Physics, Lock Haven University, Lock Haven, Pennsylvania 17745, USA Meagen Pollock Department of Geology, College of Wooster, Wooster, Ohio 44691, USA Kirby Gimson Nicole Lynde Department of Geology and Astronomy, West Chester University, West Chester, Pennsylvania 19383, USA

ABSTRACT This field trip explores igneous layering in the Morgantown Sheet, southeastern Pennsylvania, a Jurassic diabase intrusion that is part of the Central Atlantic Magmatic Province, formed during rifting of Pangea. The Pennsylvania Granite Quarry (Stop 1) is a dimension stone quarry in the southern side of the sheet, in which the cut walls display intermittent modal layering crosscut by channels of mafic diabase. Plagioclase-rich layers overlie pyroxene-rich layers in packages with slightly concave-up “wok” shapes ~0.3–0.4 m in dimension and ~0.35–0.5 m thick. Mafic diabase—both layers and crosscutting channels—contain 15–25 modal percent orthopyroxene phenocrysts and are interpreted as basaltic magma replenishments. Orientations of layering and channels suggest this part of the sheet was originally a horizontal sill ~400 m thick, at about six kilometers depth, and that the sheet was tilted 20°–25° to the north after crystallization. The Dyer aggregate quarry (Stop 2) is in the northeast side of the sheet that dips ~80° southeast (Birdsboro dike). Here, rhythmic plagioclase-pyroxene layering also dipping ~80° is found in the interior and near the margin of the ~255-m-wide dike. Augite and plagioclase compositions are very similar in samples from different vertical heights in the sheet, suggesting localized rather than sheet-wide fractionation. We compare the Morgantown Sheet layering to similar features in the Palisades sill, New Jersey, and Basement sill, Antarctica, and discuss models for their formation. Srogi, L., Lutz, T., Dickson, L.D., Pollock, M., Gimson, K., and Lynde, N., 2010, Magmatic layering and intrusive plumbing in the Jurassic Morgantown Sheet, Central Atlantic Magmatic Province, in Fleeger, G.M., and Whitmeyer, S.J., eds., The Mid-Atlantic Shore to the Appalachian Highlands: Field Trip Guidebook for the 2010 Joint Meeting of the Northeastern and Southeastern GSA Sections: Geological Society of America Field Guide 16, p. 51–67, doi: 10.1130/2010.0016(03). For permission to copy, contact [email protected]. ©2010 The Geological Society of America. All rights reserved.

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INTRODUCTION AND FIELD TRIP OBJECTIVES This field trip is a brief exploration of magmatic layering within a Jurassic mafic intrusion, the Morgantown Sheet (Fig. 1). The Morgantown Sheet is part of the Central Atlantic Magmatic Province, a large igneous province associated with, and dismembered by, the rifting of Pangea and opening of the Atlantic Ocean. The areal extent of predominantly tholeiitic basaltic magmatism in the Central Atlantic Magmatic Province is vast, perhaps as much as seven million square kilometers (Marzoli et al., 1999) on parts of North and South America, Africa, Europe, and Antarctica. Isotopic ages cluster at 200 ± 4 Ma (Marzoli et al., 1999, and references therein), suggesting that an enormous volume of basaltic magma was erupted and emplaced within a geologically short time (McHone, 2000). One of the best studied

intrusions within the Central Atlantic Magmatic Province is the 300-m-thick Palisades sill in the Newark Basin, New Jersey (Fig. 2), commonly referred to as a classic example of a verticallydifferentiated intrusion of tholeiitic magma (e.g., Lewis, 1908; Walker, 1940; Walker, 1969; Shirley, 1987; Husch, 1990). By contrast, the Morgantown Sheet, located in southeastern Pennsylvania between the Newark and Gettysburg basins (Fig. 2), has not been studied in as much detail (e.g., Smith, 1975; Gottfried et al., 1991a, 1991b). Modal layering and other aspects of magmatic plumbing in the Morgantown Sheet are well exposed and easily accessed in two active quarries, the Pennsylvania Granite Quarry (PAGQ, Elverson, Pennsylvania) and the Dyer Quarry (DQ, Birdsboro, Pennsylvania). In this guidebook paper, we include a brief comparison with magmatic layering in the Palisades sill in order to illuminate the processes that may have been

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Figure 1. Field photos of magmatic layering in the Pennsylvania Granite Quarry. (A) Wall cut parallel to the strike of the sheet (SP wall); 4.5 m wide × 3.7 m high. (B) Traced and digitized plagioclase-rich layers on SP wall. Black lines are Series 1 layers; red lines are Series 2; yellow lines are Series 3 layers. (C) Strike-normal (SN) wall; 4.9 m wide × 3.1 m high. Note mafic layer with cuspate lower boundary parallel to plagioclase-rich layers (ML). (D) Traced and digitized plagioclase-rich layers on SN wall. Black lines are Series 1 layers; orange lines are Series 2; green lines are Series 3 layers. See text for discussion of Series.

Jurassic Morgantown Sheet, Central Atlantic Magmatic Province responsible for the layering and other features we shall see in the Morgantown Sheet. Comparison of Layering in the Morgantown Sheet and Palisades Sill We are working on three extensive exposures in the Morgantown Sheet: the Dyer Quarry (DQ), Birdsboro, Pennsylvania; the Pennsylvania Granite Quarry (PAGQ), Elverson, Pennsylvania; and a roadcut on both sides of Exit 1 from Interstate 176 just north of Morgantown, Pennsylvania (Fig. 2). The northeast side of the sheet is a dike (Birdsboro dike) striking N54°W and dipping ~80° southwest. The entire width of the Birdsboro dike and both contacts with Triassic sediments are exposed in the Dyer Quarry (Fig. 3). The PAGQ is located near the southeast corner of the sheet, and the roadcut goes across ~550 m of the southwestern corner of the sheet (Fig. 2); country-rock contacts are not exposed at either locality but are inferred from existing geologic maps (Berg et al., 1980; Berg and Dodge, 1981). The schematic cross sections in Figure 2 show the Birdsboro dike and PAGQ sill with no vertical exaggeration. The geometry of the north and west sides of the sheet and the subsurface relationships are not known. It is possible that the north side is a separate, higher-level sill, similar to the saucer sills of the Golden Valley sill complex, Karoo Basin, South Africa (Polteau et al., 2008a, 2008b). At all three localities in the Morgantown Sheet, the magmatic layering is generally on the millimeter to decimeter scale

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and defined by color reflecting differences in the modal proportions of pyroxene and plagioclase, but there are important differences. In the DQ (Fig. 3), the layering is rhythmic and steeply dipping, parallel to dike margins. The layers are of approximately equal thickness and consistent in orientation and thickness over several meters. Rhythmic layering has been observed on four levels in the interior of the dike, and within 50 m of the southwest contact on one level. By contrast, in both the roadcut and PAGQ (Figs. 4A and 4B), layering is best defined by concentrations of plagioclase forming light-colored layers that range from a few millimeters to a few centimeters thick. The plagioclase-rich layers are subhorizontal or dip north-northeast at gentle to moderate angles, and appear to be curving and to branch or intersect with other plagioclase-rich layers (Figs. 4A and 4B). Layering in the roadcut is more variable in orientation than in the PAGQ and is more difficult to observe due to numerous weathered and mineralized surfaces. The PAGQ is a dimension stone quarry, so the walls and floor have been cut by diamond-studded cables into fairly smooth, flat surfaces, and the rock has fewer fractures and mineralized surfaces. In the PAGQ, layering is ubiquitous and has greater consistency in orientation: nearly horizontal on walls cut approximately parallel to the map trend (strike) of the sheet (~east-west), and dipping ~20° north on walls cut normal to strike (~north-south) (Fig. 1). Although much of the Palisades sill is composed of massive, homogeneous diabase, cm-scale modal layering is evident on weathered outcrops near 10 m, 70 m (Naslund et al., 1992),

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Figure 2. Maps and cross sections of the Morgantown Sheet. Geologic map simplified from the 1980 Pennsylvania geologic map (Berg et al., 1980). Dark gray shading—diabase intrusions and basalt lava flow (Jacksonwald, J lava). Thin black lines—diabase dikes. Light gray shading—Triassic sedimentary units. No shading—Precambrian and Paleozoic units. A–A′ and B–B′—locations of schematic cross sections; A—location of the Dyer Quarry (DQ) in the Birdsboro Dike, B—location of the Pennsylvania Granite Quarry (PAGQ). Roadcut—location of the roadcut along Exit 1 of I-176 north of Morgantown. Inset: map of the region showing locations of the Morgantown Sheet (rectangle) and Palisades sill, New Jersey (shaded), within the Triassic basin (dotted lines). Schematic cross sections show the inferred subsurface shape of the Morgantown Sheet, with no vertical exaggeration.

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and 100 m above the base of the sill. The layering results from variation in the modal abundance of pyroxene and plagioclase and appears as thin, light and dark colored bands on the weathered rock surface. One roadcut at the 100-m level on Henry Hudson Drive located ~0.4 km south of the Palisades Interstate Park Headquarters in Alpine, New Jersey exhibits layering that varies significantly in spacing and orientation. Rhythmic layers parallel the dip of the sill, are evenly spaced ~3 cm apart, and are remarkably straight, extending as far as ~7 m across the roadcut (Fig. 4C). In contrast, wispy, curved, discontinuous, lens-shaped layers with spacing up to 10 cm apart occur a few meters above the rhythmic layers (Fig. 4D). Some layers curve and sag, truncating underlying layers and giving the appearance of cross-bedding. Overall, the orientation of the layers varies as much as 20°, where some layers dip to the north and others dip to the south. The layering in the Palisades sill is strikingly similar to layering in the Morgantown Sheet (compare Figs. 3B and 4C and Figs. 4B and 4D), suggesting that previous work on the Palisades sill will provide insight into processes in the Morgantown Sheet.

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A Figure 3. Field photographs of the Dyer Quarry (DQ). (A) View looking north-northwest, taken in 2004, level 7 at the top. Contact with lighter-colored Triassic sediments just visible on the extreme right. Wall at each level is 50 ft high (15.24 m). (B) Contact-parallel, nearly vertical rhythmic mafic-felsic layering in the DQ. Sample from level 7 (circle), at the top of the quarry in 2004 (since removed). Geological Society of America scale in upper right is 16.5 cm in height. Brightness and contrast have been digitally enhanced. (C) View looking north-northwest on level 5 showing contact of diabase (left, dark) with Triassic sediments (right, light-colored). Black line shows present dip of ~80ºSW. Wall is 50 ft high (15.24 m).

Magmatic Layering and Magmatic Plumbing Igneous layering seems to be a fundamental part of the crystallization repertoire for basaltic magmas, given its common occurrence in mafic intrusions. There are competing theories to explain how layered structures form, although the many shapes, thicknesses, and types of layers suggest that more than one process can result in layering (e.g., Parsons and Becker, 1987; Naslund and McBirney, 1996). Historically, the study of magmatic layering has shaped petrologists’ larger conceptualizations of magma emplacement, chamber formation, differentiation, and crystallization (compare, for example, Wager and Brown, 1967; McBirney and Hunter, 1995; and Marsh, 2006). Petrologists have largely discarded the model of pluton formation in which a large batch of mafic magma is instantaneously emplaced within a chamber; instead, there is growing evidence that intrusions develop through repeated injections of magma, each having its own composition, crystal “cargo,” and history of transit through the crust (e.g., Coleman et al., 2004; Font et al., 2008). A newer and widely accepted model (e.g., Marsh, 1996; 2006) is that intrusions crystallize from the outside in, and have a solidification front that moves from the country-rock contact inwards. The hot, molten interior is thus encased by a solidification zone (or boundary layer) on the roof, sides, and floor. At any point in time before complete solidification, going from the liquid interior outwards, the crystallinity varies from zero to 100%, and the temperature varies from the liquidus temperature to below the solidus near the margins where the magma has solidified completely. The solidification zone may be further subdivided (e.g., Marsh, 1996), into a crystal suspension (0%–25% crystalline), a crystal mush (25%–~55% crystalline), and a rigid crust (~55%–100% crystalline). Within the context of solidification zone crystallization, there are several plausible models for the formation of magmatic layering. The models we shall emphasize in this paper are: (1) Mechanical sorting of phenocrysts during flow of magma (e.g., Marsh, 1996; Bédard et al., 2007; Petford, 2009); (2) Rafts or plumes of gravitationally-unstable crystal mush that detach from the roof or walls, sink, and flow across the floor (Irvine et al., 1998; Philpotts and Dickson, 2002); (3) Compaction of crystal mush and compaction-driven reactions that redistribute minerals on a millimeter to centimeter scale (Dickson and Philpotts, 2001). Other processes that may be relevant include the physical movement of melt between the liquid interior of the pluton and the porous crystal mush; chemical exchange between the crystal mush and the liquid interior that can lead to oscillatory behavior and feedback processes (e.g., Wang and Merino, 1993); kinetic control of mineral growth resulting from differences in nucleation rates or the dynamic interplay of diffusion rates for heat and chemical components; and subsolidus coarsening and recrystallization (see Naslund and McBirney, 1996, for a complete review). Mafic intrusions cannot form in isolation within the crust, and must be part of larger interconnected networks that transport

Jurassic Morgantown Sheet, Central Atlantic Magmatic Province basaltic magmas from the upper mantle into the site of emplacement. The Morgantown Sheet was part of a plumbing system (magmatic mush column in the parlance of Marsh, 1996, 2004) that ultimately fed continental flood basalt volcanism at the surface. Although no basalts are preserved above the Morgantown Sheet, just to the east is the Jacksonwald lava flow near the northwest border of the basin (Fig. 2), sitting atop another (presumably) interconnected set of intrusions. The Birdsboro dike, which forms the northeast side of the Morgantown Sheet, may have been a feeder for fissure eruptions, but it is not known whether magma flowed through the entire sheet on the way to the surface, as proposed for the Palisades sill (Puffer et al., 2009), or whether the PAGQ sill was just a lateral offshoot of the flood-basalt plumbing. In the discussion below and on the field trip, we will try to emphasize the connections among the parts of the Morgantown

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Sheet in the context of a dynamic and tectonically-active magmatic system. We are in the early stages of a collaboration that will ultimately bring together petrography, geochemical analyses of whole-rock and mineral samples, anisotropy of magnetic susceptibility (AMS), and quantitative textural analysis to evaluate models of magmatic flow, compaction, crystallization, replenishment, and thermal history. In the sections below, we first briefly review work on Palisades sill layering (Dickson, 2006). Then, we present preliminary results including quantitative data on orientation, length, concavity, sinuosity, and sequence, based on digitization of 1452 plagioclase-rich layers (Gimson et al., 2009). To our knowledge, this is the only detailed statistical analysis of modal layering in mafic intrusions using modern computer technology. We conclude with a proposed model for the formation of layering in the Morgantown Sheet.

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Figure 4. Field photographs of magmatic layering in the Morgantown Sheet and Palisades sill. Brightness and contrast have been digitally enhanced. (A) Complex, light-colored, plagioclase-rich layers within medium gray diabase from the Pennsylvania Granite Quarry. Note parallel and crosscutting dark layers and channels. Cut from a wall parallel to the strike of the sheet. Geological Society of America scale in upper right is 16.5 cm in height. (B) Similar plagioclase-rich layers in the northeast side of a roadcut along Exit 1 of I-176 north of Morgantown, Pennsylvania. West Chester University student for scale. (C) Regularly spaced, rhythmic layers located at the 100-m level of the Palisades sill in a roadcut on Henry Hudson Drive, opposite the Palisades Interstate Park Headquarters in Alpine, New Jersey. This section of rhythmic layers is capped by a final prominent plagioclase-rich layer. Gloved hand on upper left for scale. (D) Wispy, curved discontinuous lens-shaped layers located at 100 m above the base of the sill. Notebook for scale is ~19 cm in height.

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A MODEL FOR THE FORMATION OF CM-SCALE RHYTHMIC MODAL LAYERING IN THE PALISADES SILL, NEW JERSEY: INSIGHTS FOR THE MORGANTOWN SHEET In the differentiation model for the Palisades sill proposed by Shirley (1987), crystallizing mush in the roof was gravitationally unstable, detached and sank to the floor in the form of dense plumes. There, the cumulate pile was compacted with upward expulsion of residual liquid. Bergantz and Ni (1999) modeled viscous fluids similar to tholeiitic basaltic magma and showed that density instabilities fall as coherent plumes that deposit on the floor with little mixing or stirring; consequently, the deposited material may retain some textural characteristics from its origin in the roof. Philpotts and Dickson (2000, 2002) proposed that plumes of crystal mush including neutrally buoyant plagioclase crystals were an efficient way of transporting material from the roof to the floor of the 200-m-thick Holyoke flood-basalt flow. Detailed petrographic and chemical analysis of layering in the Palisades sill provides evidence for the compaction and recrystallization of roof-derived plumes that resulted in modal layering (Dickson and Philpotts, 2001; Dickson, 2006). Detailed Analysis of Modal Layering The modal layering in the Palisades sill consists of couplets of a plagioclase-rich layer and an underlying pyroxene-rich layer. A sharp boundary occurs in the center of each couplet with the concentration of pyroxene decreasing gradationally downward and the concentration of plagioclase decreasing gradationally upward. In thin section, the boundary between plagioclase-rich layers and pyroxene-rich layers is sharp but undulates slightly over large pyroxene crystals protruding partly into the overlying plagioclase-rich layer. To investigate the chemical variation among these layers, 44 slices, each ~3 mm thick, were cut parallel to the layers and analyzed chemically. Major and minor element concentrations match the modal variations. The MELTS program (Ghiorso and Sack, 1995) was used to calculate liquidus temperatures of pyroxene and plagioclase for each slice. Results show that the composition of each slice plots closely on either side of the pyroxeneplagioclase cotectic, the phase boundary along which pyroxene and plagioclase crystallized simultaneously from the melt. Each slice of the rock is saturated in either pyroxene or plagioclase by an average temperature difference of 12 °C, and the compositional sum of each plagioclase-pyroxene couplet is equivalent to the cotectic composition of homogeneous rock. The bulk composition of the rock at the 100-m stratigraphic level is unique in the sill in having a cotectic composition. Below the 100-m level the rocks are pyroxene saturated and above the 100-m level the rocks are plagioclase saturated. Evidence suggests that crystal settling was an unlikely cause of the rhythmic modal layering. Density calculations from the MELTS program (Ghiorso and Sack, 1995) show that plagio-

clase would have been neutrally buoyant in the melt composition. Additionally, pyroxene-rich layers are enriched in TiO2 and contain late-crystallizing ilmenite, suggesting that these layers retain significant interstitial residual liquid rather than having separated from the liquid. Layer compositions suggest that they developed from mineral phase redistribution in what was initially homogeneous crystal mush having a near-cotectic composition. Petrographic evidence supports this hypothesis. Texturally, plagioclase-rich layers, contain only ophitic pyroxene, interstitial to and enclosing plagioclase, whereas the homogeneous unlayered diabase and pyroxene-rich layers also contain large pyroxene phenocrysts and granular pyroxene clusters. Loss of some pyroxene from the plagioclase-rich layers would explain this textural difference. Dickson (2006) proposed that pyroxene dissolved in the plagioclase-rich layers and precipitated in the pyroxene-rich layers as granular clusters of small (0.2–0.4 mm) polygonal pyroxene grains, many with ~120° dihedral angles. Similar rocks in the thick Holyoke basalt flow also were interpreted to have formed by recrystallization of roof material that fell to the floor of the sill (Philpotts and Dickson, 2000, 2002). In addition, textural analysis of the plagioclase-rich layers shows that up to 20% of plagioclase crystals have preferential alignments that lie within 20° of the vertical direction. On the other hand, fewer than 4% of plagioclase crystals in the pyroxene-rich layers are aligned. Dickson (2006) and Dickson and Philpotts (2001) assumed that all plagioclase crystals were initially randomly oriented and concluded that the textural anisotropy of the plagioclase-rich layers developed during compaction. Crystals can become preferentially aligned in a compacting crystal mush, if there is room for them to rotate. Gray et al. (2003) and Shirley (1987) demonstrated that crystal mush in the floor zone of the Palisades sill was compacted; however, only that portion with a cotectic composition developed the rhythmic layers (Dickson and Philpotts, 2001). Dickson (2006) proposed that compaction of the crystal mush caused pyroxene to dissolve and the components to diffuse downward a short distance, on the order of millimeters, where it precipitated to form a pyroxene-rich layer. Dissolution of pyroxene required only small degrees of differential stress from compaction because the bulk composition was so close to the cotectic. Removal of pyroxene left the mush enriched in plagioclase with room for the crystals to rotate and become aligned. Results of modeling the low abundance of incompatible elements, such as TiO2, in plagioclaserich layers are consistent with loss of ~20% residual liquid during compaction. Plagioclase crystals within the pyroxene-enriched layers had no room to realign due to increased pyroxene crystallization; essentially they were cemented in place and maintained their random orientation. In summary, chemical and textural evidence suggests that rocks near the 100-m-level began as homogeneous, cotectic crystal mush crystallized in the roof zone that sank to the floor of the sill as dense crystal-rich plumes, spread out, compacted, and recrystallized to form rhythmic modal layers. The shapes and sizes of the modal layering can be used to infer aspects of the development of the Palisades sill magma

Jurassic Morgantown Sheet, Central Atlantic Magmatic Province chamber (Dickson, 2006). If the base of the plagioclase-rich layer in each modal couplet delineates the base of the plume, then the shape of the couplet reflects the topography of the floor (or the boundary between the crystal mush and crystal suspension within the solidification zone). Disturbance and deformation of the mush floor is likely at plume impact sites (G.W. Bergantz, 2005, personal commun.). Thus, curved and wispy layers (Fig. 4D) suggest an uneven or lumpy floor and could have resulted from deposition proximal to plume impact sites. Straight and regularly spaced rhythmic layers (Fig. 4C) suggest that the plume spread out evenly across a relatively flat or gently sloping floor, distal to the impact site. Since curved, wispy layers are located only meters above rhythmic layers in the Palisades sill, it can be concluded that the location of sinking plumes changed over time. Furthermore, a minimum estimate for the volume of roof-derived plumes can be obtained assuming that each rhythmic layer represents a deposited plume with dimensions of a shallow cylinder. For a layer 3 cm thick and 7 m across, the minimum volume of the plume would be 1.15 m3. Stoke’s Law calculations indicate that, if the plume fell as a sphere with this volume and a diameter of ~1.3 m and a very slight density contrast with the host magma (0.005 mg/m3), it would have sunk to the floor of the sill in less than 24 h. MODAL LAYERING IN THE MORGANTOWN SHEET: FIELD CHARACTERISTICS, PRELIMINARY MINERAL DATA AND DISCUSSION OF MODELS The Morgantown Sheet diabase has many similarities to the Palisades sill diabase as well as some significant differences. Generally speaking, mineralogy and texture are similar, consisting primarily of plagioclase, augite, and pigeonite. However, we have not (yet?) found strongly differentiated or iron-enriched diabase, or an olivine layer in the Morgantown Sheet. Furthermore, diabase in the PAGQ contains ~20 modal percent orthopyroxene, typically as almost euhedral phenocrysts. In this section we present field observations and interpretations, analysis of the spatial statistics of the plagioclase layering in the PAGQ, and preliminary data on modes and mineral chemistry. We conclude with a discussion of models for the origin of the layering, drawing on interpretations of the Palisades sill but with some differences to explain specific characteristics of the Morgantown Sheet. Field Characteristics of Mafic Diabase in the PAGQ: Implications for Emplacement and Post-Intrusion History Mafic pyroxene-rich diabase in the PAGQ is observed as layers parallel to and always underneath, not on top of, plagioclase-rich layers. These layers are roughly horizontal on quarry walls oriented parallel to strike (Figs. 1A and 1B), but dip 20°–25° north on quarry walls oriented normal to strike (Figs. 1C and 1D). This intermittent modal layering strongly resembles the Palisades sill (Fig. 4), where the pyroxene-rich layers are interpreted to have formed by compaction-driven recrystallization of roof-derived crystal mush (Dickson, 2006). In the PAGQ, how-

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ever, there is evidence for another possible origin. Some thick (~0.5 m) mafic diabase layers have a cuspate lower margin that shows load-cast and flame structures (Figs. 1C and 1D). Similar structures are interpreted to develop where higher-density basaltic magma is injected into a plutonic magma chamber and flows over a less-dense crystal mush (e.g., Wiebe, 1993). Mafic diabase in the PAGQ is enriched in euhedral phenocrysts of orthopyroxene relative to the unlayered diabase. The orthopyroxene phenocrysts resemble those in the Basement sill, Ferrar dolerites, Antarctica, a magmatic system contemporaneous with the Central Atlantic Magmatic Province. Marsh (1996, 2004) and Bédard et al. (2007) interpreted the orthopyroxene phenocrysts to be primocrysts carried by the magma when it intruded, and showed that phenocryst abundance was related to the flow of magma outwards from a magma feeder. Froelich and Gottfried (1999) note that orthopyroxene phenocrysts tend to be found in the southeastern part of several of the Jurassic diabase sheets in Pennsylvania. Based on these field observations, and by analogy with the Basement sill, we interpret the mafic diabase to be magma replenishments bearing orthopyroxene phenocrysts and suggest that the PAGQ may overlie a feeder dike or pipe into the Morgantown Sheet. This also may explain why the sheet is thicker near the PAGQ (Fig. 2). Diabase in the roadcut outcrop lacks the orthopyroxene phenocrysts because they were concentrated close to the feeder by their greater density. In addition to the layers, zones of mafic diabase also crosscut the plagioclase-rich layers, form narrow channels above places where they crosscut the layers, and those channels terminate at a higher level (Fig. 4A). Data collected from one wall in the PAGQ show that the plagioclase-rich layers are more likely to be deflected upwards (78%) than downwards (22%) at the contact with the channels, suggesting that the mafic diabase was moving upward. We interpret the mafic diabase channels as having been mobile melt replenishments that were rising through and “stalled out” within the solidification zone. In seeming contradiction to this interpretation, the pyroxene-rich channels are more dense than the plagioclase-rich layers they crosscut. We suggest that the channels lost some portion of their melt component which continued to rise upwards, leaving behind the crystal-rich and dense mafic diabase we see today. The scale and configuration of the mafic diabase layers and channels may provide information about the crystallinity of the solidification zone and the style and scale of melt migration. For example, the many small layers and channels in Figure 4A suggest a dispersed mini–dike-sill complex that might have formed if the solidification zone was only rigid enough to form decimeter-scale fractures. The mafic diabase channels are approximately perpendicular to plagioclase-rich layers on both strike-parallel walls (Fig. 4A) and strike-normal walls, suggesting that they had originally a vertical orientation. Undergraduate students in the 2008 petrology class at West Chester University measured the dip of the dark channels on several of the strike-normal walls as 67° ± 12° to the north (n = 38). If originally vertical (90°), these results indicate that ~20°–25° of post-magmatic tilting has occurred, the same as

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the dip of the plagioclase-rich layers on the strike-normal walls (Fig. 1C). The implication is that the magmatic layering was originally approximately horizontal and that the south side of the Morgantown Sheet was a subhorizontal sill at the time of intrusion (Srogi et al., 2008). Post-magmatic tilting of the Birdsboro dike had a very small effect on its strike and dip; rotation by 20° about an east-west axis yields an estimate of N54°W, 84°NE, or approximately vertical at the time of intrusion. Given the orientation of the sheet in the DQ and PAGQ, we can estimate its thickness and deepest level of exposure. The thickness of the Birdsboro dike in the DQ is ~255 m; the south side of the sheet ranges in thickness from ~200 m up to 400–700 m in the vicinity of the PAGQ. With an aerial extent of ~11 × 20 km, this provides a volume estimate of ~50–100 km3 after tilting and erosion—a significantly larger area than estimated by Froelich and Gottfried (1999). The sheet is substantial, but still thin (Fig. 2), with an aspect ratio around 50:1. The distance from the PAGQ to the northwest basin border and a dip of 20°N result in a minimum estimate of ~6 km for the original depth of the Morgantown Sheet exposed in the PAGQ. The PAGQ is located in the lower half of the sheet, ~100–150 m above the inferred lower contact. The roadcut obliquely traverses the sheet beginning ~80 m above the inferred lower contact and ending ~25 m beneath the inferred upper contact, assuming that it, too, dips ~20°. The estimates for post-Jurassic tilting of the Morgantown Sheet are in broad agreement with models for the basin development (e.g., Faill, 1973, 2003; Schlische et al., 2003; Schlische and Withjack, 2005) and its thermal history (Huntoon and Furlong, 1992). The Morgantown Sheet is located in the extreme western end of the Newark Basin (Pennsylvania–New Jersey), where it narrows into the “narrow neck” connection to the Gettysburg basin (Pennsylvania-Maryland), (Froelich and Gottfried, 1999). Most models for the failed-rift basin involve the opening of a half-graben during Triassic time, with a normal fault along the north-northwest border of the basin in Pennsylvania (e.g., Root and MacLachlan, 1999; Schlische et al., 2003; Schlische and Withjack, 2005), although Faill (1973, 2003) considers all basin tectonism to be post-Jurassic. The narrow-neck was not necessarily a narrow part of the basin in the Triassic, but was the location of an extensive alluvial fan draining the upland to the north (Faill, 2003), or possibly developed on top of a bedrock ridge leading to shallower depositional environments than in the Newark or Gettysburg basins (Smoot, 1999). Faill (2003) also proposes that the narrow-neck region experienced the greatest amount of late Mesozoic tilting in the Pennsylvania basins, perhaps as much as 30°–40°N. Whereas the dip of Triassic sediments measured near the contact in the DQ (~31°NW) agrees with Faill’s tilting estimate, our data from the PAGQ suggest that only ~20° of the tilting occurred after ~200 Ma, and that there probably was northward tilting of basin sediments before (and during?) Jurassic magmatism. Post-magmatic tilting also is consistent with the results of thermal modeling, based on apatite and zircon fission-track measurements and vitrinite reflectance, that indicates Jurassic deposition of sediments up to 2.5– 3 km thick (Huntoon and Furlong, 1992).

Spatial Characteristics and Statistical Analysis of Layering We have focused our statistical analysis of the magmatic layering on the excellent exposures in the PAGQ, where two walls were selected for detailed work (Fig. 1), one oriented approximately normal to the strike of the sheet (the SN wall), and the other approximately strike-parallel (the SP wall). The top of the SN wall was ~10 m above the top of the SP wall; the latter has since been removed by quarrying. Digital photographs taken at 2048 × 1536 resolution were imported into Adobe Photoshop, and the brightness and contrast were digitally adjusted to enhance the light-colored plagioclase-rich layers. A total of 1452 plagioclase-rich layers were manually traced and digitized sequentially as three Series (Gimson et al., 2009). Series 1 layers are the most prominent and enriched in plagioclase and show the greatest contrast with adjacent diabase; Series 3 layers show the least contrast. The data were imported into S-PLUS statistical software for analysis. Data for the lengths (Fig. 5), sinuosities and concavities (Fig. 6) suggest that the plagioclase-rich layers are generally planar to

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Figure 5. Frequency distributions (%) of lengths (m) of the plagioclase-rich layers in the Pennsylvania Granite Quarry. (A) Strike-parallel wall. (B) Strike-normal wall.

Jurassic Morgantown Sheet, Central Atlantic Magmatic Province “wok-shaped” in three dimensions. The point-to-point lengths of the layers range from 3 cm to just over 3 m, with the SN wall having more long layers than the SP wall (Fig. 5). Otherwise, most layers are short and average lengths vary with Series 1 > Series 2 > Series 3 for both walls. Sinuosity (Fig. 6) is a measure of the “curviness” of a line. It is the ratio of the sum of the lengths of the individual segments that make up the digitized line relative to the distance between the end-points of the line. The sinuosity of a straight line is 1; a line with sinuosity = 2 is twice as long as the distance between its ends. Our results show that the majority of plagioclase-rich layers in the PAGQ are quite straight; indeed, a

logarithmic scale is used on Figures 6A and 6C to maximize the separation of layers with sinuosity values close to 1. The concavity orientations of the layers are shown on the x-axis in Figures 6A and 6C. On the SP wall (Fig. 6A), the data strongly cluster into two groups centered on concavity orientation 90 (concaveup and horizontal) and −90 (concave-down and horizontal). The majority of layers are concave-up, as shown on Figure 6B where concave-up and concave-down digitized lines are plotted separately on a “map” of the SP wall (note that the x- and y-axes are horizontal and vertical distances, respectively). Similar results are observed for the SN wall, but the orientation of concavity

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Figure 6. Sinuosity and concavity orientation (°) of plagioclase-rich layers in the Pennsylvania Granite Quarry. (A) Strike-parallel (SP) wall, concavity orientation versus log sinuosity. Symbols below x-axis show example concavity orientations corresponding to specific degrees. Note clustering of sinuosities at low values (straight), and clustering of concavity orientations at approximately horizontal concave-up (90º) and concave-down (−90º). (B) Map of SP wall; x- and y-axes are dimensions of SP wall (m). Lines are plotted for Series 1 and 2 layers of sinuosity <3; thicker lines are concave-up, and thinner lines are concave-down. Note majority of lines are concave-up. (C) Concavity orientation versus log sinuosity for strike-normal (SN) wall, as in (A). Note clustering of sinuosities at low values (straight) and clustering of concavity orientations at dips of ~20º from horizontal. (D) Map of SN wall with lines plotted as in (B). Note majority of lines are concave-up.

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(Fig. 6C) is influenced by the ~20° dip of the layers (Fig. 6D). Layer-length distributions are consistent with random sections through many “wok-shaped” layers with median dimensions of 0.3–0.4 m. These lateral dimensions are significantly shorter than the maximum length of the rhythmic layers in the Palisades sill (~7 m) but are consistent with the observed lengths of the “wispy” discontinuous layering in the Palisades sill (Fig. 4D). The sequence of Series 1, 2, and 3 layers was determined from their intersection with transect lines taken 0.5 m apart and oriented perpendicular to layers in the SP and SN walls. The frequencies with which layers of any given Series are followed by layers of the same or different Series were determined from 863 intersections counting from the bottom to the top of the transect. A chi-square test shows that it is unlikely (p ≅ 0) that the frequencies result from a random arrangement, and that there is a reduced tendency for Series 1 to be followed by another Series 1, and for Series 2 and 3 to be adjacent to one another. In other words, Series 2 and 3 layers occur preferentially between Series 1 layers. The maximum distance between Series 1 layers based on the sequence analysis is 35–50 cm (up to 0.5 m), providing an estimate for the thickness of the plagioclase-pyroxene layers. The dips of the layers calculated from the digitized data fall in a fairly narrow range on both walls but with different values (Fig. 7). Figure 7 is similar to a rose diagram, but the dips of the digitized layer segments have been smoothed using a 10° window. On the SN wall, the great majority of plagioclase-rich layers of all Series show a narrow range of dips ~18°–22°N (Fig. 7B). On the SP wall, the dips are more variable but there is a strong peak for all Series at 0° ± 15°, close to horizontal (Fig. 7A). These dip data, along with the orientations of the steep, crosscutting dark channels, are the basis for our inference that the entire Morgantown Sheet was tilted ~20° north sometime after 200 Ma (the approximate age of the magmatism). On the SP wall, a prominent plagioclase-rich layer traverses almost the entire length of the wall in the middle, and the dips are noticeably different above and below this layer (Fig. 6B). Layers in the lower half of the wall tend to dip west, while the layers in the upper half tend to dip east, and the difference is most pronounced for the Series 1 layers. The scale and orientations of the plagioclase-rich layers on the SP wall are similar to the discontinuous style of layering in the Palisades sill that Dickson (2006) proposed to be the result of the impact of plumes of roof-derived crystal mush on the floor of the sill. Mineralogy, Textures, and Mineral Compositions To date, we have examined over 100 thin sections from the three sample localities in the Morgantown Sheet (the PAGQ, the roadcut on Exit 1 off I-176 near Morgantown, and the Dyer Quarry, DQ), and have determined the modal mineralogy by point-counting 12 rock thin sections from the DQ and PAGQ and 12 samples of aggregate from the DQ. Mineral compositions were analyzed in polished thin sections coated with carbon using the FEI Quanta environmental scanning electron microscope (ESEM) with Oxford Inca energy-dispersive detector (EDS) at

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Figure 7. Distributions of dip directions of plagioclaserich layers in the Pennsylvania Granite Quarry shown on diagrams oriented like each wall, with geographic direction along the x-axis and vertical direction along the y-axis. Bold line shows the directions of dip, similar to a rose diagram but smoothed by a 10º window. Concentric circles are percentages of dip values (25%, 50%, 75%, and 100%), normalized to most abundant dip direction at 100% (outer circle). Results are shown on a 360º scale. (A) Strike-parallel wall, all Series. Bold line intersects outer circle near 0º, indicating that most dips are close to horizontal. Smaller peaks on bold line between the 25% circle and close to the 50% circle indicate smaller proportion of lines dipping within ~20°– 25º of horizontal. (B) Strike normal wall, all Series; note strong peak near 20ºN.

Jurassic Morgantown Sheet, Central Atlantic Magmatic Province West Chester University. Accelerating voltages and acquisition times were 20 KeV and 10–20 seconds for plagioclase to minimize loss of Na; and 20 or 30 KeV, 20–30 s for all other minerals. The spot size control was set to optimize the acquisition rate to 3000–3500 counts per second on the sample with dead time between 30%–35% (manufacturer’s recommendation). The probe diameter was on the order of a μm. Quant optimization was on cobalt metal. Element abundances in weight percent were calculated by the software with oxygen by stoichiometry and normalized to 100%. Detection limits for any individual element are ~0.1 wt%. The resulting semiquantitative data are highly reproducible with very good agreement with microprobe analyses for normalized cation proportions, except for systematically slightly higher Si concentrations (variable, but up to ~1 wt%). We have confidence in comparing data from the three localities analyzed under similar operating conditions. Petrographic Observations and Modal Data Diabase in the DQ, PAGQ, and roadcut is phaneritic with grain sizes ranging from fine- to medium-grained. Coarsegrained to pegmatitic diabase is found in the DQ and roadcut, and contains abundant granophyre, quartz, and K-feldspar (~20 modal percent granophyre minerals on average). One aphanitic chill-margin sample, collected about a meter from the southwestern contact, contains microphenocrysts of augite, plagioclase, and olivine (euhedral but entirely replaced by secondary minerals), and one elongated inverted pigeonite phenocryst, in a matrix of intergranular plagioclase and augite. Unlayered diabase ranges from more mafic compositions with roughly equal modal proportions of plagioclase and pyroxenes, to more felsic compositions averaging 52% plagioclase and 24% pyroxenes. Diabase from the PAGQ contains more pyroxene than the diabase in the DQ and, as noted, contains 15%–20% orthopyroxene phenocrysts that are not observed in the roadcut and DQ. Thin section textures are hypidiomorphic-granular and sub-ophitic; plagioclase, augite, and pigeonite comprise the framework of the rock with interstitial pyroxenes, hornblende, biotite, chlorite, quartz, K-feldspar, and granophyre. The hydrous minerals include those formed at higher temperatures (biotite; hornblende and other amphiboles commonly replacing pyroxenes), and at lower temperatures (chlorite; sericite and other minerals in altered plagioclase and K-feldspar). There is no apparent correlation between the proportions of higher-temperature and lower-temperature hydrous minerals among samples. Unlayered diabase from the PAGQ contains significantly fewer hydrous minerals than averaged, unlayered DQ diabase, whether amphiboles (1.5% versus 20%) or altered feldspars (<1% versus 13%). Despite the outcrop-scale differences between intermittent modal layering (PAGQ and the roadcut) and rhythmic layering (DQ), all of the layered diabase samples are similar in thin section and differ primarily in the identity of pyroxene phenocrysts: mainly orthopyroxene with some augite in the PAGQ, and augite and pigeonite in the roadcut and DQ. Mafic layers are dominated by euhedral to subhedral pyroxene (Figs. 8A, 8C, and 8E), while

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the finer-grained plagioclase-rich layers contain ophitic and interstitial pyroxene (Figs. 8B, 8D, and 8F). These microtextures are very similar to those in the rhythmically-layered diabase in the Palisades sill. Plagioclase-rich layers contain about twice as much interstitial quartz, K-feldspar, and granophyre as the pyroxene-rich layers, and more sericitized and altered feldspar. Calculated as a percentage of total feldspar (to compensate for the higher overall amounts of feldspar), plagioclase-rich layers contain 9%–11% altered plagioclase and K-feldspar compared with 1%–3% in mafic layers. The greatest differences between layered diabase samples are found by comparing the PAGQ and DQ. Layered PAGQ diabase contains significantly lower amounts of amphiboles and interstitial granophyre minerals compared with layered DQ diabase (a pattern also observed in the unlayered diabase). In the PAGQ, some plagioclase-rich layers show significant preferred orientation of plagioclase grains, closer packing and smaller patches of interstitial minerals (Fig. 8B) compared with the DQ (Fig. 8F). Finally, PAGQ unlayered diabase is similar to the pyroxene-rich mafic layers, and very different from the plagioclase-rich layers in the PAGQ. By contrast, the average, unlayered felsic diabase from the DQ has modal mineral proportions roughly half-way between the pyroxene-rich and plagioclase-rich layers from the DQ, similar to the differences in the rhythmic layers reported for the Palisades sill. Preliminary Mineral Composition Data Mineral compositions reported here were analyzed in samples from the southern portion of the sheet collected from four different heights above the inferred base of the sill (all heights are approximate): (1) 85 m above the inferred base—unlayered diabase from the roadcut; (2) 125 m—layered diabase from the PAGQ; (3) 215 m—layered diabase from the roadcut; and (4) 260 m—unlayered diabase from the roadcut; 25 m below the inferred upper contact. In addition, we report results from both the chilled margin and layered diabase in the interior of the Birdsboro dike, exposed in the DQ. Because the sheet is dipping ~20° north, the dike in the DQ was originally as much as five kilometers above the sill exposed in the PAGQ and roadcut. Despite few samples, a striking picture emerges from the mineral compositions. The cores of augite and plagioclase phenocrysts have very similar compositions throughout the sheet (filled diamonds, Fig. 9). Within each sample, augite and plagioclase have more Fe-rich and Na-rich rim compositions, respectively (open triangles, Fig. 9), interpreted to be normal zoning to lowertemperature compositions. However, the PAGQ sample shows the smallest compositional range, and the Birdsboro dike shows the most extreme Fe-enrichment in augite and Na-enrichment in plagioclase (Fig. 9). In addition, there are two notable differences in plagioclase compositions among the sample localities. Plagioclase phenocrysts throughout the sill and in the dike chill margin contain Na-rich cores which are not observed in the layered

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Figure 8. Back-scattered electron images of mafic and plagioclase-rich layers sized so that all images are at the same scale; scale bar is 1 mm long. Abbreviations: opx—orthopyroxene, aug—augite, pig—pigeonite, plag—plagioclase, kf—K-feldspar. Darkest gray interstitial grains (unlabeled) are albitic plagioclase, K-feldspar, quartz, or granophyre. In (A), (C), and (E), note zoning in pyroxene from Mg-rich cores (darker gray) to Fe-rich rims (brighter gray). In (B), (D), and (F), note zoning in plagioclase: Ca-rich is brighter gray and Na-rich is darker gray. See text for discussion of zoned plagioclases outlined in dotted lines. (A) Mafic diabase channel in Pennsylvania Granite Quarry (PAGQ). (B) Plagioclase-rich layer in PAGQ sample. (C) Mafic diabase layer in roadcut sample. (D) Plagioclase-rich layer in roadcut sample. (E) Mafic diabase layer in Dyer Quarry (DQ) sample. Note zoning in augite that suggests clusters of small crystals with cementing overgrowth. (F) Plagioclase-rich layer in DQ sample.

Jurassic Morgantown Sheet, Central Atlantic Magmatic Province sample from the dike interior (see Fig. 9 and plagioclase crystals outlined with dotted lines in Figs. 8B, 8D, and 8F). Furthermore, plagioclase phenocrysts in the sill commonly have at least one Ca-rich zone surrounding the Na-rich core (Figs. 8B and 8D), of the same composition as the Ca-rich cores in the dike interior (Fig. 9); however, the Ca-rich zone is lacking in the chilled margin phenocrysts (Fig. 9). At all four, widely separated localities in the sill, we also found oscillatory-zoned plagioclase phenocrysts with compositions that are virtually identical to each other and similar to the compositions of more simply zoned plagioclase. Interpretations of Mineralogical Data The sheet-wide constancy of augite core compositions and similarity of plagioclase compositions are consistent with Smith (1975), who found that magma compositions were almost uniform within the ~600-m-thick York Haven sheet. The results suggest that fractionation was limited to a relatively small scale within the solidification zone, rather than occurring at the scale of the entire sheet (e.g., Marsh, 1996; Bédard et al., 2007). The ubiquity of oscillatory-zoned plagioclase suggests a process operating throughout the chamber that repeatedly brought partly crystalline material in contact with hotter or more primitive magma. Local thermal or compositional variations could explain differences in plagioclase core and rim compositions observed in the chilled margin, dike interior, and sill.

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In the steep rhythmic layers of the Birdsboro dike, the abundant late-stage minerals and strong intracrystalline zoning are consistent with crystallization from trapped liquid, implying less percolation and loss of evolved melt (Hunter, 1996), possibly due to more rapid cooling or the absence of compaction by gravity. PAGQ diabase has more extreme modal mineralogy, with much smaller amounts of late-stage minerals, and more cumulate-like textures than diabase at the roadcut or in the DQ dike. These features are consistent with greater percolation and loss of evolved residual liquids from the solidification zone and explain why the PAGQ minerals are not as strongly zoned. We suggest that the hot magma replenishments in the PAGQ, of greater thickness and density due to the orthopyroxene phenocrysts, led to both greater compaction and slower cooling there compared with the roadcut. While slower cooling might help maintain porosity and permeability, compaction ultimately drove the loss of most interstitial liquid from the PAGQ. These ideas require testing by techniques that provide constraints on cooling rates and thermal history, such as the pioneering work of Holness using measurements of true dihedral angles at mineral grain contacts (Holness et al., 2005, 2007a; Holness, 2006). She has shown that true dihedral angles preserve evidence of magma replenishments (Holness, 2005) and the thermal maturation (Holness et al., 2007b) of crystal mush in mafic intrusions. This technique holds promise for future investigation of the Morgantown Sheet.

B

Figure 9. Mineral compositions in the Morgantown Sheet. Bottom part of graphs show Pennsylvania Granite Quarry and roadcut samples plotted against calculated height above the base of the sill; top graph shows Dyer Quarry samples at the same vertical scale, except for chill-margin sample, which is plotted at top. Horizontal lines connect analyses of grains from the same sample; cores and rims were analyzed on the same grains; then each was averaged separately. (A) Augite phenocryst compositions, cores (filled diamonds) and rims (open triangles). Note similarity of core compositions in dike and sill, and greater Fe-enrichment in augite rims in the dike. (B) Plagioclase phenocrysts without oscillatory zoning. Note similarities of Na-rich cores (filled diamonds) and Ca-rich zones (light gray filled squares) throughout the sill and in the dike. Note lack of Na-rich cores in dike interior and lack of Ca-rich zones in chill-margin phenocrysts. Note greater Na-enrichment in plagioclase rims (filled triangles) in the dike interior compared with the sill.

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Models for the Formation of Layering in the Morgantown Sheet The scale and texture of the intermittent, plagioclasepyroxene modal layering in the sill exposed in the PAGQ and roadcut (Figs. 1A and 4B) strongly resembles the wispy, discontinuous layering in the Palisades sill at the 100-m level outcrop (Fig. 4D). At the thin section scale, the textures of the plagioclase-rich layers in the Morgantown Sheet (PAGQ and roadcut) are similar to those in the Palisades sill, having smaller grain sizes and interstitial, ophitic pyroxenes. Dickson and Philpotts (2001) and Dickson (2006) propose that crystal mush tore loose from the roof solidification zone of the Palisades sill and sank as plumes that spread out over the crystal mush on the floor. Building on the work of Shirley (1987) and Gray et al. (2003), they further propose that the stress of compaction acting on mush of cotectic composition resulted in the dissolution, downward-migration, and precipitation of pyroxene to form a couplet of roof-derived mush depleted in pyroxene (a plagioclase-rich layer) overlying a pyroxene-rich layer. This model would explain the dimensions and morphology of the “wok-shaped” plagioclase layers in the PAGQ as resulting from the impact of roof material on the floor crystal mush. The observation that smaller, less prominent layers (Series 2 and 3) tend to occur between larger, more prominent layers (Series 1) could be explained by this model as smaller pieces of roof material that descended more slowly and piled on top of larger and heavier pieces. The estimated volume of an average PAGQ layer, using the maximum thickness between Series 1 layers of 0.5 m, is ~0.60 m3, smaller than the rhythmic layers, but perhaps comparable to the discontinuous layers in the Palisades sill. Unfortunately, the Dickson-Philpotts model does not explain the layers and crosscutting channels of mafic diabase enriched in orthopyroxene phenocrysts that are found in the PAGQ but not observed in the Palisades sill. The plagioclase-rich layers in the Morgantown Sheet also strongly resemble features described as anorthositic to gabbronoritic schlieren and veins in the Basement sill, Antarctica (Bédard et al., 2007). Both the Basement sill and Morgantown Sheet also contain localized concentrations of orthopyroxene phenocrysts and have nearly uniform mineral compositions throughout the intrusions. In the model proposed by Marsh (2004) and Bédard et al. (2007), the magma that formed the Basement sill was injected as a slurry rich in orthopyroxene and small plagioclase phenocrysts, which were mechanically segregated in response to buoyancy forces. These authors further propose that plagioclase-rich schlieren and veins were channels of the feldspar-charged melt “frozen” in the process of separating from the orthopyroxene-enriched residue, initially by sorting of phases during dynamic flow, or in local dilation zones within a crystal mush under shear forces (e.g., Petford, 2009), and later by coalescence into ascending vertical pipes. The Marsh-Bédard model would thus explain the formation of intermittent modal layering in the PAGQ by lateral flow segregation of phenocryst-bearing magma. However, this model does not account for the concave-up

shapes and relatively small lateral extent of the layers (~0.3–0.4 m), which are better explained by the Dickson-Philpotts model. Also, in the Marsh-Bédard model the plagioclase-rich layers form channels and vertical pipes, whereas in the PAGQ, the plagioclase-rich layers appear rigid and are crosscut by the mafic diabase that forms vertical channels and pipes. Where the plagioclase occurs on the sides of the mafic diabase channels, it appears to be a piece of the adjacent layer, broken and dragged upwards, rather than a channel of ascending melt as described by Bédard et al. (2007). The mafic diabase in the PAGQ provides evidence for the injection and migration of replenishing magmas in the lower solidification zone of the Morgantown Sheet. Magma replenishments are not explicitly considered in either the Marsh-Bédard or Dickson-Philpotts model, despite evidence for their role in the Basement sill (Bédard et al., 2007) and Palisades sill (Husch, 1990; Gorring and Naslund, 1995). If the orthopyroxene-rich layers originated as repeated magma inputs at the floor of the Morgantown Sheet, earthquakes accompanying magma intrusion events may have triggered the roof collapse that resulted in the overlying plagioclase-rich layers. This model has the benefit of a larger perspective that views the intrusive sheet and flood basalts as an interconnected magmatic system within a tectonically-active rift basin. Hoffer (1965) suggested that there may be seismic control of igneous layering, and Shaw (1980) attempted to quantify and relate magma volumes and rates of transport, to crustal failure and stress regimes, and the frequency and magnitude of seismic events. Naslund and McBirney (1996) suggest that seismically-induced layering should be laterally continuous across the chamber, but that is only strictly true if the seismicity triggers nucleation-driven layer formation in the solidification zone. If, on the other hand, the layering is caused by earthquake-triggered roof collapse, then individual layers would be of variable extent, reflecting the shape and rheology of the roof solidification zone, and may be broken, disrupted, and overlapped by material that continued to cascade down onto the floor. This model also predicts that unlayered diabase would form in thinner parts of the sheet, more distal from the magma feeder, where the roof and floor would crystallize more rapidly with a much thinner liquid interior. These might be good places to look for internal chilled margins between magma replenishments (Polteau et al., 2008b). The nearly vertical rhythmic layering in the Birdsboro dike cannot have resulted from any gravity-driven process, by definition, although many other processes are possible. An intriguing model is the spontaneous development of mineral segregations in crystal-rich magma flowing within the dike conduit, as observed for industrial slurries (e.g., Petford, 2009). Segregation is dependent on dike dimensions, flow rate, and crystal abundances, sizes, and shapes; therefore, the absence of layering could be explained by differences in volume, flux, or crystallinity of magma inputs. ROAD LOG AND FIELD TRIP STOP DESCRIPTIONS On the field trip, we will visit the Birdsboro dike in the Dyer Quarry and the layered sill in the Pennsylvania Granite Quarry.

Jurassic Morgantown Sheet, Central Atlantic Magmatic Province Both are active quarries, so we emphasize below the important features to look for but cannot give exact locations for the most part. Permission to visit must be obtained from the companies that own the quarries. Hard hats and appropriate footwear are required. Assemble in the designated area at the Sheraton Baltimore City Center Hotel, 101 West Fayette Street, Baltimore, Maryland 21201. Mileage

Directions

0.0

Head west on W Fayette St. toward N Liberty St., 269 ft.Turn right at Park Ave. Turn right at W Mulberry St. Continue on Orleans St., 200 ft. Turn right at Colvin St. Turn right at E Fayette St. Turn right at I-83 N. Entering Pennsylvania. Take Exit 19A for PA-462 W. Merge onto N Hills Rd. Turn right at U.S.-30 E. Continue straight onto U.S.-222 N. Take the exit toward I-76/PA-272/Denver/Pennsylvania Turnpike. Turn right at Spur Rd./State Route 1040. Take the ramp onto I-76 E. Take Exit 298 to merge onto I-176 N toward PA-10/ Morgantown/Reading. Partial toll road. Take Exit 1B to merge onto PA-10 S/Reading Rd. toward PA-23/Morgantown. Continue to follow PA-10 S. Turn left at E Main St./PA-23 E. Continue to follow PA-23 E. Turn left at Trythall Rd. Turn right into driveway leading to the office of Pennsylvania Granite Corporation, 410 Trythall Rd, Elverson, Pennsylvania 19520-8957.

0.2 0.9 1.2 1.4 54.3 54.5 55.1 77.8 93.6 93.8 94.6 106.6 107.2

108.4 115.8 117.0

Stop 1: Pennsylvania Granite Quarry (PAGQ), Elverson, Pennsylvania The PAGQ dimension stone quarry has been operating for more than 30 years, first as a local business and now owned by Rock of Ages, Inc., Vermont. The diabase is exceptionally dense, tough, and unfractured, so it is cut for countertops, monuments, and memorials, including Dale Earnhardt’s mausoleum and part of the Vietnam Veterans’ Memorial in Washington, D.C. The stone is prized for the patterns of the light gray, plagioclase-rich layers within the black diabase, and is sold under the trade names American Black Granite and Jet Mist. Interestingly, a nearby abandoned quarry has only uniformly dark diabase, with no plagioclase-rich layering. Currently, the stone is shipped mostly to Italy and China, and the company estimates that there is another 1500 years of useable dimension stone in the quarry. Fractures mineralized with calcite and/or chlorite (termed white lines and green lines by the quarrymen, respectively) and other flaws limit

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recovery of useful stone to 5%–15%; i.e., 90% waste on average. A waste pile of about five million tons of stone at the top of the quarry is now being crushed for aggregate by another company. The process of cutting the walls and floor, with limited use of explosives, is very slow compared with an aggregate quarry, which is good for the geologists studying the quarry walls. Sampling the walls in situ is difficult and can only be done with permission of the quarry manager. Samples should be taken from the loose blocks designated as waste material. The cut walls are generally stable, but for safety, trip participants must stay 10 feet from the edge of a wall or a berm, and 50 feet from a working saw. We shall begin the quarry tour in the southwest corner which is not being actively quarried at the present time. Virtually all of the important features can be observed here (Fig. 1C), including: (1) the nature and orientations of the modal layering; (2) parallel and crosscutting dark layers representing magma replenishments; (3) apparent load casts, flame structures, and chaotic layering that suggest different rheological behaviors of the crystal mush; and (4) younger mineralized fractures. We shall tour the available quarry walls and discuss additional features and the implications for the magmatic processes that formed the sheet. If time permits, we shall return to the quarry office and visit the saw plant to learn more about the quarry operation. Lunch Break: Warwick County Park, located ~3 miles southeast of PAGQ, on County Park Rd. south of PA-23. Mileage

Directions (from PAGQ, not including lunch stop)

117.0

Head south on Trythall Rd. toward PA-23 E/ Ridge Rd. Turn right at PA-23 E/Ridge Rd. Turn right at N Chestnut St./PA-82 N. Turn left at Elverson Rd./PA-82 N. Turn right at PA-82 N/Twin Valley Rd. Turn left at Haycreek Rd./PA-82 N, go 259 ft. Make slight left at Rock Hollow Rd. Turn left at Rock Hollow Rd./T348. Turn left into gate for Dyer Quarry.

117.7 122.0 122.5 124.0 128.9 129.8 131.9

Stop 2: Dyer Quarry (DQ), Birdsboro, Pennsylvania The Dyer Quarry also has been in operation for several decades and is currently owned by Anderson Construction, Inc. It is an aggregate quarry, and every size fraction of the crushed stone is sold: for example, the largest pieces are used for rip rap to prevent coastal erosion; fist-sized cobbles make excellent railroad ballast; smaller pieces are used for gravel and road aggregate; and the finest material is used in filtration systems. At the time of this field trip (March 2010) the quarry will likely comprise four or five vertical levels each with a wall 50 ft high. The rock is extensively fractured and unstable; field trip participants

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should stay 10 feet from the walls and be aware of loose rock above their heads. Although we will not know the exact locations available within the quarry until shortly before the trip, the features that we will focus on include (Fig. 3): (1) Triassic sediments and intrusive contacts exposed on both sides of the Birdsboro Dike; (2) chilled margins and grain size variations across the dike; (3) joint and fracture patterns in the diabase, some with slickenlines, as reflections of the cooling and tectonic history; (4) steeply dipping rhythmic modal layering, primarily in the dike interior; (5) internal chill margins and other evidence for multiple dike intrusions; and (6) distribution of pegmatitic diabase and granophyre and migration of evolved liquids within the dike. We shall tour the available quarry walls and discuss magmatic processes within the dike and connections to the rest of the sheet and to probable lava flows at the surface. Mileage

Directions

131.9

Turn left to head north on Rock Hollow Rd. toward Quarry Rd. Take the 2nd right into Old River Rd., go 213 ft. Take the first left onto Main St./PA-724 W. Merge onto I-176 N via the ramp to Reading. Take Exit 11B to merge onto U.S.-422 W toward Reading. Take the exit onto U.S.-222 S/U.S.-422 W toward U.S.-222 S/Lancaster/Lebanon. Continue on U.S.-30W. Turn left at N Hills Rd. Turn right at E Market St./PA-462 W, go 410 ft. Slight right to merge onto I-83 S toward Baltimore. Take Exit 4 for MD-2/St. Paul St. Turn right at E Fayette St. Turn left. Arrive Sheraton Baltimore City Center Hotel.

132.9 136.2 136.6 141.8 170.6 193.1 193.7 245.7 299.1 299.7

ACKNOWLEDGMENTS We acknowledge with profound gratitude Mr. Lance Battersby and the crew of the Pennsylvania Granite Quarry and Mr. Eric Friend and the crew of the Dyer Quarry, for providing access, samples, information, logistical support, and good cheer to the authors and students from West Chester University (WCU) for several years. The first author also acknowledges Dr. Frederick C. Monson, Technical Director, Center for Microanalysis, Imaging, Research, and Training, West Chester University, for training, technical support, and for the intellectual curiosity that has led to innovative uses of the environmental scanning electron microscope–energy dispersive X-ray for this research. Funding from a West Chester University Faculty Development Grant and College of Arts and Sciences Scholarly Development

Award has supported the first author’s research with students. The first author gratefully acknowledges the work of student researchers, Diane Oronzio and Nicole Steiner, and students in the WCU petrology classes, Josh Andrewson, Eileen Capitoli, Jason Daliessio, Eric Dieck, John Fiorello, Sarah Johnson, John LaBold, Leandra Larsen, Russ Losco, Jennie Matkov, Tiffany McClennen, Rob Miller, Tiffany Neumann, Seth Pelepko, Sean Rafferty, Dan Sivco, Danny Wojton, Gabe Antonello, John Antonucci, Shanna Babiak, Andrew Bentley, Paul Girafalco, James Hannagan, Bryan Narwich, Lauren Peterson, Nicole Tornaritis, Justin Turpin, and Russelle Westermann. The authors thank Steve Shank and Dave Eggler for thoughtful reviews that significantly improved the manuscript. REFERENCES CITED Bédard, J.H.J., Marsh, B.D., Hersum, T.B., Naslund, H.R., and Mukasa, S.B., 2007, Large-scale mechanical redistribution of orthopyroxene and plagioclase in the Basement Sill, Ferrar Dolerites, McMurdo Dry Valleys, Antarctica: Petrological, mineral-chemical and field evidence for channelized movement of crystals and melt: Journal of Petrology, v. 48, p. 2289–2326, doi: 10.1093/petrology/egm060. Berg, T.M., and Dodge, C.M., compilers and editors, 1981, Atlas of preliminary geologic quadrangle maps of Pennsylvania: Pennsylvania Geological Survey, Map 61, scale 1:52,500, 636 p., 624 maps. Berg, T.M., Edmunds, W.E., Geyer, A.R., et al., compilers, 1980, Geologic map of Pennsylvania (2nd ed.): Pennsylvania Geological Survey, 4th series, Map 1, scale 1:250,000, 3 sheets [web release]. Bergantz, G.W., and Ni, J., 1999, A numerical study of sedimentation by dripping instabilities in viscous fluids: International Journal of Multiphase Flow, v. 25, p. 307–320, doi: 10.1016/S0301-9322(98)00050-0. Coleman, D.S., Gray, W., and Glazner, A.F., 2004, Rethinking the emplacement and evolution of zoned plutons: Geochronologic evidence for incremental assembly of the Tuolumne Intrusive Suite, California: Geology, v. 32, p. 433–436, doi: 10.1130/G20220.1. Dickson, L.D., 2006, Detailed textural analysis of the Palisades sill, New Jersey [Ph.D. thesis]: Storrs, University of Connecticut, 183 p. Dickson, L.D., and Philpotts, A.R., 2001, Analyzing the complex texture of a simple diabase: Palisades sill, New Jersey: Geological Society of America Annual Meeting Abstracts with Programs, v. 33, no. 6, p. A-137. Faill, R.T., 1973, Tectonic development of the Triassic Newark-Gettysburg basin in Pennsylvania: Geological Society of America Bulletin, v. 84, no. 3, p. 725–740, doi: 10.1130/0016-7606(1973)84<725:TDOTTN> 2.0.CO;2. Faill, R.T., 2003, The early Mesozoic Birdsboro central Atlantic margin basin in the Mid-Atlantic region, eastern United States: Geological Society of America Bulletin, v. 115, no. 4, p. 406–421, doi: 10.1130/0016-7606(2003)115<0406:TEMBCA>2.0.CO;2. Font, L., Davidson, J.P., Pearson, D.G., Nowell, G.M., Jerram, D.A., and Ottley, C.J., 2008, Sr and Pb Isotope micro-analysis of plagioclase crystals from Skye Lavas: An insight into open-system processes in a flood basalt province: Journal of Petrology, v. 49, p. 1449–1471, doi: 10.1093/petrology/ egn032. Froelich, A.J., and Gottfried, D., 1999, Chapter 12B: Early Mesozoic—Igneous and contact metamorphic rocks, in Schultz, C.H., ed., The Geology of Pennsylvania: Harrisburg, Pennsylvania, Pennsylvania Geological Survey Special Publication no. 1, p. 203–209. Ghiorso, M.S., and Sack, R.O., 1995, Chemical mass transfer in magmatic processes. IV. A revised and internally consistent thermodynamic model for the interpolation and extrapolation of liquid-solid equilibria in magmatic systems at elevated temperatures and pressures: Contributions to Mineralogy and Petrology, v. 119, p. 197–212, doi: 10.1007/BF00307281. Gimson, K., Lutz, T., and Srogi, L., 2009, Quantitative measurements and spatial statistics of igneous layering in the Morgantown Sheet, southeast Pennsylvania: Geological Society of America Annual Meeting Abstracts with Programs, v. 41, no. 7, p. 665–666.

Jurassic Morgantown Sheet, Central Atlantic Magmatic Province Gorring, M.L., and Naslund, H.R., 1995, Geochemical reversals within the lower 100 m of the palisades sill, New Jersey: Contributions to Mineralogy and Petrology, v. 119, no. 2–3, p. 263–276, doi: 10.1007/BF00307286. Gottfried, D., Froelich, A.J., and Grossman, J.N., 1991a, Geochemical data for Jurassic diabase associated with early Mesozoic basins in the Eastern United States: Western Newark Basin, Pennsylvania and New Jersey: U.S. Geological Survey Open-File Report 91-0322-D, p. 27. Gottfried, D., Froelich, A.J., and Grossman, J.N., 1991b, Geochemical data for Jurassic diabase associated with early Mesozoic basins in the eastern United States: Gettysburg Basin and vicinity, Pennsylvania and Maryland: U.S. Geological Survey Open-File Report 91-0322-E, p. 56. Gray, N.H., Philpotts, A.R., and Dickson, L.D., 2003, Quantitative measures of textural anisotropy resulting from magmatic compaction illustrated by a sample from the Palisades sill, New Jersey: Journal of Volcanology and Geothermal Research, v. 121, p. 293–312, doi: 10.1016/S0377-0273 (02)00463-8. Hoffer, A., 1965, Seismic control of layering in intrusions: American Mineralogist, v. 50, p. 1125–1128. Holness, M.B., 2005, Spatial constraints on magma chamber replenishment events from textural observations of cumulates: the Rum Layered Intrusion, Scotland: Journal of Petrology, v. 46, no. 8, p. 1585–1601, doi: 10.1093/petrology/egi027. Holness, M.B., 2006, Melt-solid dihedral angles of common minerals in natural rocks: Journal of Petrology, v. 47, no. 4, p. 791–800, doi: 10.1093/petrology/ egi094. Holness, M.B., Cheadle, M.J., and McKenzie, D., 2005, On the use of changes in dihedral angle to decode late-stage textural evolution in cumulates: Journal of Petrology, v. 46, no. 8, p. 1565–1583, doi: 10.1093/petrology/ egi026. Holness, M.G., Anderson, A.T., Martin, V.M., Maclennan, J., Passmore, E., and Schwindinger, K., 2007a, Textures in partially-solidified crystalline nodules: A window into the pore structure of slowly-cooled mafic intrusions: Journal of Petrology, v. 48, no. 7, p. 1243–1264, doi: 10.1093/petrology/ egm016. Holness, M.B., Nielsen, T.F.D., and Tegner, C., 2007b, Textural maturity of cumulates: A record of chamber filling, liquidus assemblage, cooling rate and large-scale convection in mafic layered intrusions: Journal of Petrology, v. 48, no. 1, p. 141–157, doi: 10.1093/petrology/egl057. Hunter, R.H., 1996, Texture Development in Cumulate Rocks, in Cawthorn, R.G., ed., Layered Intrusions: Amsterdam: Elsevier, p. 77–101. Huntoon, J.E., and Furlong, K.P., 1992, Thermal Evolution of the Newark Basin: The Journal of Geology, v. 100, p. 579–591, doi: 10.1086/629609. Husch, J.M., 1990, Palisades Sill: Origin of the olivine zone by separate magmatic injection rather than gravity settling: Geology, v. 18, p. 699–702, doi: 10.1130/0091-7613(1990)018<0699:PSOOTO>2.3.CO;2. Irvine, T.N., Andersen, J.C., and Brooks, C.K., 1998, Included blocks (and blocks within blocks) in the Skaergaard intrusion: Geologic relations and origins of rhythmic modally graded layers: Geological Society of America Bulletin, v. 110, no. 11, p. 1398–1447, doi: 10.1130/0016-7606(1998)110<1398:IBABWB>2.3.CO;2. Lewis, J.V., 1908, The Palisades diabase of New Jersey: American Journal of Science, v. 176, p. 155–162. Marsh, B.D., 1996, Solidification fronts and magmatic evolution: Mineralogical Magazine, v. 60, p. 5–40, doi: 10.1180/minmag.1996.060.398.03. Marsh, B.D., 2004, A magmatic mush column Rosetta Stone: The McMurdo Dry Valleys of Antarctica: Eos (Transactions, American Geophysical Union), v. 85, p. 497 and 502. Marsh, B.D., 2006, Dynamics of magmatic systems: Elements, v. 2, p. 287– 292, doi: 10.2113/gselements.2.5.287. Marzoli, A., Renne, P.R., Piccirillo, E.M., Ernesto, M., Belliene, G., and DeMin, A., 1999, Extensive 200-million-year-old continental flood basalts of the Central Atlantic Magmatic Province: Science, v. 284, no. 5414, p. 616– 618, doi: 10.1126/science.284.5414.616. McBirney, A.R., and Hunter, R.H., 1995, The cumulate paradigm reconsidered: The Journal of Geology, v. 103, p. 114–122, doi: 10.1086/629727. McHone, J.G., 2000, Non-plume magmatism and rifting during the opening of the central Atlantic Ocean: Tectonophysics, v. 316, p. 287–296, doi: 10.1016/S0040-1951(99)00260-7. Naslund, H.R., and McBirney, A.R., 1996, Mechanisms of formation of igneous layering, in Cawthorn, R.G., ed., Layered Intrusions: Amsterdam, Elsevier, p. 1–43.

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Naslund, H.R., Hristov, L.G., and Gorring, M.L., 1992, Cm-scale modal layering in the Palisades Sill, New York–New Jersey, USA: Geological Society of America Abstracts with Programs, v. 24, no. 7, p. 86. Parsons, I., and Becker, S.M., eds., 1987, Origins of Igneous Layering: Dordrecht, D. Reidel Publishing, North Atlantic Treaty Organization Advanced Science Institutes, Series C, Mathematical and Physical Sciences, v. 196. Petford, N., 2009, Which effective viscosity?: Mineralogical Magazine, v. 73, no. 2, p. 167–191, doi: 10.1180/minmag.2009.073.2.167. Philpotts, A.R., and Dickson, L.D., 2000, The formation of plagioclase chains during convective transfer in basaltic magma: Nature, v. 406, p. 59–61, doi: 10.1038/35017542. Philpotts, A.R., and Dickson, L.D., 2002, Millimeter-scale modal layering and the nature of the upper solidification zone in thick flood-basalt flows and other sheets of magma: Journal of Structural Geology, v. 24, p. 1171– 1177, doi: 10.1016/S0191-8141(01)00099-2. Polteau, S.A., Mazzini, A., Galland, O., Planke, S., and Malthe-Sorenssen, A., 2008a, Saucer-shaped intrusions: Occurrences, emplacement and implications: Earth and Planetary Science Letters, v. 266, p. 195–204, doi: 10.1016/j.epsl.2007.11.015. Polteau, S., Ferre´, E.C., Planke, S., Neumann, E.-R., and Chevallier, L., 2008b, How are saucer-shaped sills emplaced?: Constraints from the Golden Valley Sill, South Africa: Journal of Geophysical Research, v. 113, B12104, 13 p. Puffer, J.H., Block, K.A., and Steiner, J.C., 2009, Transmission of flood basalts through a shallow crustal sill and the correlation of sill layers with extrusive flows: The Palisades intrusive system and the basalts of the Newark Basin, New Jersey, U.S.A: The Journal of Geology, v. 117, p. 139–155, doi: 10.1086/595663. Root, S.I., and MacLachlan, D.B., 1999, Chapter 21: Gettysburg-Newark Lowland, in Schultz, C.H., ed., The Geology of Pennsylvania, Harrisburg, Pennsylvania: Pennsylvania Geological Survey Special Publication no. 1, p. 298–305. Schlische, R.W., and Withjack, M.O., 2005, The early Mesozoic Birdsboro central Atlantic margin basin the Mid-Atlantic region, eastern United States: Geological Society of America Bulletin, v. 117, no. 5-6, p. 823–828, doi: 10.1130/B25498.1. Schlische, R.W., Withjack, M. O., and Olsen, P.E. 2003, Relative timing of CAMP rifting, continental breakup, and inversion: Tectonic significance, in Hames, W.E., McHone, G.C., Renne, P.R., and Ruppel, C., eds., The Central Atlantic Magmatic Province: American Geophysical Union Monograph 136, p. 33–59. Shaw, H.R., 1980, The fracture mechanisms of magma transport from the mantle to the surface, in Hargraves, R.B., ed., Physics of Magmatic Processes: Princeton, Princeton University Press, p. 201–264. Shirley, D.N., 1987, Differentiation and compaction in the Palisades Sill, New Jersey: Journal of Petrology, v. 28, no. 5, p. 835–865. Smith, R.C., III, 1975, Geology and geochemistry of Triassic diabase in Pennsylvania: Geological Society of America Bulletin, v. 86, p. 943–955, doi: 10.1130/0016-7606(1975)86<943:GAGOTD>2.0.CO;2. Smoot, J.P., 1999, Chapter 12A: Early Mesozoic—Sedimentary rocks, in Schultz, C.H., ed., The Geology of Pennsylvania, Harrisburg, Pennsylvania: Pennsylvania Geological Survey Special Publication no. 1, p. 180–202. Srogi, L., Lutz, T., Dine, J., Oronzio, D., and Lynde, N., 2008, Igneous modal layering in a Jurassic basaltic dike-sill complex, southeastern Pennsylvania, USA: Geological Society of America Abstracts with Programs, v. 40, no. 6, p. 252. Wager, L.R., and Brown, G.M., 1967, Layered Igneous Complexes: Edinburgh, Oliver & Boyd. Walker, F.R., 1940, Differentiation of the Palisade diabase, New Jersey: Geological Society of America Bulletin, v. 51, p. 1059–1106. Walker, K.R., 1969, The Palisades Sill, New Jersey: A reinvestigation: Geological Society of America Special Paper 111, 178 p. Wang, Y., and Merino, E., 1993, Oscillatory magma crystallization by feedback between the concentrations of the reactant species and mineral growth rates: Journal of Petrology, v. 34, p. 369–382. Wiebe, R.A., 1993, Basaltic injection into floored silicic magma chambers: Eos (Transactions, American Geophysical Union), v. 74, p. 1 and 3.

MANUSCRIPT ACCEPTED BY THE SOCIETY 2 DECEMBER 2009 Printed in the USA

The Geological Society of America Field Guide 16 2010

The early through late Pleistocene record in the Susquehanna River Basin Duane D. Braun Emeritus, Department of Geography and Geosciences, Bloomsburg University, Bloomsburg, Pennsylvania, 17815, USA

ABSTRACT A transect of the Coastal Plain, Piedmont, Ridge and Valley, and Appalachian Plateau physiographic provinces valley will be made by going up the Susquehanna Valley. Early to late Pleistocene features will be examined at eleven sites. The prePleistocene evolution of the Appalachian landscape will also be discussed at two sites. The journey will start on the Coastal Plain and travel to the lower Susquehanna bedrock gorge across the High Piedmont, where strath terraces will be examined. From there the Susquehanna River will be followed upstream to the Low Piedmont where broad outwash terraces, of early to late Pleistocene-age, flank the River. Continuing up-river into Ridge and Valley Province, the water gaps at Harrisburg, Pennsylvania, will be viewed and their origin discussed. The Susquehanna River will be followed across the remainder of the Ridge and Valley with stops to view early and mid-Pleistocene-aged features. Emphasis will be placed on the amount of erosion that has occurred since the early Pleistocene and the development of a pseudo-moraine landscape. At the wind gap through Bald Eagle Mountain, the mid-Miocene to present evolution of the overall Appalachian landscape will be discussed. The evidence for the early Pleistocene-age for Glacial Lake Lesley will be examined in the West Branch Susquehanna valley. In the deep valleys section of the Appalachian Plateau, the mid- and late Pleistocene glacial termini will be examined. Turning south into the Ridge and Valley, the trip will conclude with examination of glaciofluvial deposits of probable mid-Pleistocene-age.

DAY 1. BALTIMORE, MARYLAND, TO WILLIAMSPORT, PENNSYLVANIA

be discussed at two sites. The journey will start on the Coastal Plain and head to the lower Susquehanna bedrock gorge across the High Piedmont, where strath terraces will be examined (Stop 1). From there the Susquehanna River will be followed upstream to the Low Piedmont where broad outwash terraces flank the river (Stop 2). Continuing upriver into Ridge and Valley Province, the water gaps at Harrisburg, Pennsylvania, will be viewed (Stop 3). The Susquehanna River will be followed across the remainder of the Ridge and Valley with stops to view early and

The trip will travel up the Susquehanna Valley (Fig. 1) to examine the early to late Pleistocene features and the overall Appalachian Mountain landscape. The trip will also be a transect of all the Appalachian physiographic provinces: Coastal Plain, Piedmont, Ridge and Valley, and Appalachian Plateau. The prePleistocene evolution of the Appalachian landscape will also

Braun, D.D., 2010, The early through late Pleistocene record in the Susquehanna River Basin, in Fleeger, G., and Whitmeyer, S.J., eds., The Mid-Atlantic Shore to the Appalachian Highlands: Field Trip Guidebook for the 2010 Joint Meeting of the Northeastern and Southeastern GSA Sections: Geological Society of America Field Guide 16, p. 69–101, doi: 10.1130/2010.0016(04). For permission to copy, contact [email protected]. ©2010 The Geological Society of America. All rights reserved.

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Figure 1. Relief map of the field trip region; stops are numbered squares. Dashed lines are physiographic province boundaries. The provinces are: CP—Coastal Plain, PD—Piedmont, MB—Mesozoic Basin portion of the Piedmont, BR—Blue Ridge, RP—Reading Prong, RV—Ridge and Valley, AP—Appalachian Plateau. Solid lines are glacial limits of different ages with the age denoted by its marine isotope stage number (MIS 2, etc.): Late Pleistocene or Late Wisconsinan—MIS 2; Mid-Pleistocene—MIS 6–18; MIS 6—Illinoian; Early Pleistocene—MIS 20+.

Susquehanna River Basin mid-Pleistocene–aged features (Stops 4–7). The pre-Pleistocene landscape evolution will be discussed at Stop 3 and 7. After staying overnight in Williamsport, Pennsylvania, the trip will view early Pleistocene deposits in the West Branch Susquehanna Valley (Stop 8) and then head north into the deep valleys section of the Appalachian Plateau. There the mid- and late-Pleistocene glacial termini will be examined (Stops 9–11). Turning south into the Ridge and Valley, probable mid-Pleistocene deposits will be examined (Stop 12). Note: Approximately one hour travel time to the first stop. From the hotel, 101 W. Fayette Street, Baltimore (Lat. 39° 17′ 24.73″ N, Long. 76° 33′ 01.55″ W), turn left (west) onto W. Fayette Street, turn left (south) onto Hopkins Place, turn right (west) onto W. Lombard Street, turn left (south) onto S. Howard Street, that street becomes I-39. Then bear right onto ramp to I-95 north, take I-95 north to exit 77, bear right onto ramp for MD 24 north (Emmorton Road), right (north) onto combined U.S. 1– MD 24, and left continuing on MD 24. Then take right onto MD 165 to Pennsylvania line (road becomes PA 74), right onto PA 372, go over Norman Wood Bridge, Holtwood hydroelectric dam on left, and park on right just past bridge (Lat. 39° 49′ 21.01″ N, Long. 76° 19′ 06.08″ W). Stop 1A. Bedrock Channel of the Susquehanna River from the Norman Wood Bridge The purpose of this stop is to observe the bedrock channel of the Susquehanna River and discuss the processes of strath genesis. Walk out on bridge to observe Susquehanna channel. The view for the bridge affords an outstanding perspective of the Holtwood Islands and the modern Susquehanna channel. At low water, this is an impressive view of active straths, and it indicates this river’s ability to carve out a wide valley bottom while still vertically incising. The extreme eastern channel bank is not a flat strath, but rather has been revealed as one of six “deeps” (Mathews, 1917) within the Piedmont reach of this river. Deeps are spoon-shaped, unconnected excavations that hug the eastern side of the modern channel. Mathews (1917) shows that the only deep that was exposed is intensely potholed and sculpted. It is assumed the other deeps are comparable. The deep here at Holtwood is over 40 m, which places its base ~6 m below sea level. The deeps are superimposed on a lower Susquehanna long profile that is convex from Harrisburg all the way to tidewater (Fig. 2). The many islands that dot the Holtwood area attest to relatively rapid incision of a paleo-strath through this reach (Thompson, 1990). Note that even though the island tops appear flat, a correlation based on the tops alone suggests an upstream dip (Fig. 3). The passage of a knickpoint at a rate similar to the vertical down-wearing of the channel would result in a timetransgressive, upstream-dipping strath, here represented by the island tops, that are old down stream, but young upstream where it merges with the modern strath just in front of the dam. Strath genesis is probably linked to times when the river has more abrasive tools, such as during glacial outwash. Wide straths

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are favored by both the availability of tools as well as protracted periods of base level stability. In contrast, the strath is destroyed by vertical incision, concentrated in numerous rivulets, during times of low bedload flux and/or during times of base level fall. Upstream migration of knickpoints, born at the fall zone by impulsive glacio-eustatic fall, is an additional mechanism for channel bed lowering. The strath exposed here, presumably carved during the late Pleistocene, is currently being destroyed by vertical incision of numerous small channels and rivulets. If the island tops are a paleostrath, it was abandoned by both vertical incision, as well as the passage of a knickpoint. Turn around and recross the Norman Wood Bridge. Turn right immediately after the bridge into the Lock 12 parking area (Lat. 39° 48′ 47.78″ N, Long. 76° 19′ 43.41″ W). Restrooms are below the parking lot. Stop 1B. Holtwood Islands (River Stage Permitting) The purpose of this stop is to inspect the islands at Holtwood and discuss the mechanisms of bedrock channel erosion, including the formation of potholes and the deeps. The islands along the western channel bank are easily reached by following the blue blazes of the Mason-Dixon trail. Potholes are an important feature on the island tops and along their sides. The potholes, along with the deeps have been used to argue for large, perhaps catastrophic floods in the lower Susquehanna River during the Pleistocene (Thompson, 1985, 1990) in part because these features scale with similar features associated with the scabland floods of the Columbia River. Two potential sources of catastrophic floods have been suggested. The first has an origin along the North Branch of the river as catastrophic subglacial floods responsible for shaping the drumlin fields of New York State (Shaw, 1989). We will review the evidence at Stop 5 for the unlikelihood of

Figure 2. Longitudinal profile of the lower Susquehanna River illustrating the location of the deeps through the Piedmont reach.

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Figure 3. Air photo (top), cross section (middle), and long profile of the islands at Holtwood (bottom). The white line in the air photo represents the location of the cross section, east on left, west on right. The white box in the photo represents a swath profile of 10-m resolution digital elevation model topography that is used to reconstruct the mean elevations of the islands through the Holtwood Gorge. The numbered strath levels in the bottom cross section roughly correspond to the four straths of Reusser et al. (2006) shown below in Figure 9; however, Reusser et al. (2006) do not argue for upstreamdipping straths as depicted in this diagram.

such floods actually coming down the North Branch of the river. The other possible origin of catastrophic floods is the failure of an ice dam along the West Branch and the rapid draining of glacial Lake Leslie. The presence of Lake Leslie has been recently verified (Gardner et al., 1994; Ramage et al., 1998) as an early Pleistocene, ice-dammed lake in the Bald Eagle Valley of central Pennsylvania (Stops 7 and 8). Alternatively, the potholes and the deeps are not relict, but rather modern features actively being carved by the river. The presence of similar features on rivers like the Potomac or tributaries to the Susquehanna River that never experienced glaciation in their headwaters argue against catastrophic floods. Furthermore, the Susquehanna River has generated large historic discharges such as hurricane Agnes (>106 cfs), and spring rain-on-snow events have generated historic icechocked discharges in excess of several hundred thousand cubic feet per second. New cosmogenic dating of four strath surfaces here in the Holtwood Gorge provide the first quantified look at the rates of river incision for this reach of the Susquehanna (Reusser et al., 2006; Fig. 4). Exposure ages modeled from 10Be activities indicate that fluvially eroded bedrock surfaces within Holtwood Gorge increase predictably in age with height above the channel floor, and that all are late Pleistocene features. The highest well-preserved terrace (level 3) yields a mean exposure age of 36.1 + 7.3 ka (n = 14). The middle and lowest terraces, levels 2 and 1, yield mean exposure ages of 19.8 + 2.7 ka (n = 20) and 14.4 + 1.2 ka (n = 10), respectively. One-way analysis of variance (ANOVA) demonstrates that the terrace ages are distinguishable (p<0.0005), confirming that the three levels do indeed represent separable periods of strath formation, and terrace abandonment. Two samples collected from heavily weathered and eroded topographic high points (LR-01 and LR-43), standing >20 m above the channel floor, yield model ages of 97.1 + 10.5 ka and 84.5 + 9.1 ka, respectively. Because the bedrock sampled at these two locations was shattered and no longer preserved water-polished surfaces, we report these ages as lower limiting estimates only; the removal of rock and the associated cosmogenic nuclides by weathering and erosion means that these surfaces could be far older than their model exposure ages suggest. Model ages for samples collected from bedrock surfaces between the prominent terraces (n = 22) range from 45.8 + 4.9 ka to 15.3 + 1.6 ka, and in general increase in age with height above the channel floor. Figure 5 is a combined longitudinal profile of the present Susquehanna and its terraces from Harrisburg to Chesapeake Bay. It suggests the possible correlation of outwash terraces upstream of the fall zone to fragmental terrace remnants within the Piedmont fall zone gorge. Upstream of the gorge in the carbonate areas, the Pleistocene terraces containing crystalline erratics are as high as 40 m or 130 ft above present river level. But the gorge is 200 m or 600 ft deep where it crosses the most resistant Piedmont rocks. This implies that much of the cutting of the present gorge pre-dates the first early Pleistocene glacier that entered the Susquehanna drainage.

Susquehanna River Basin Turn left onto Rt. 372 and recross the bridge, ascend out of the Holtwood Gorge and onto the high Piedmont that is underlain by metamorphic rock, turn left onto River Road. Follow River Road as it wanders across the hills east of the river for the next several miles bearing right or left where necessary. Where River Road crosses Pequea Creek at Martic Forge is the Martic line, a major lineament and fault in the High Piedmont, the significance of which has long been argued by “Piedmont geologists.” Cross Conestoga Creek at Safe Harbor. Turn left after crossing the creek and again follow River Road as it wanders across the hills east of the river, bearing right or left where necessary for the next several miles.

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Descend from the Piedmont Highland onto the Piedmont Lowland that is underlain by Paleozoic carbonates. The Susquehanna River channel and its overall valley are exceptionally wide here. River Road now runs beside the Susquehanna River and becomes PA 441. Follow 441 through Columbia, Pennsylvania, turning right then left. Go over Chickies Ridge (held up by the Cambrian Chickies quartzite) and continue on 441 going by Marietta on the left. The river here is again cutting carbonates of the low Piedmont (Conestoga limestone) and its valley is again exceptionally wide, permitting accommodation space for the development of a broad flight of terraces. Turn left onto Vinegar Ferry Road and drive down to dead end near the river (Lat. 40° 03′ 30.10″ N, Long. 76° 36′ 23.68″ W). If the river stage is high, one will not make it all the way down to the dead end. We will descend a flight of five Pleistocene terraces approaching Stop 2. Stop 2. Pre-Pleistocene to Holocene Susquehanna River Terraces at Marietta, Pennsylvania

Figure 4. Ages (top) and incision rates (middle) calculated from straths in the Holtwood Gorge. Time-transgressive upstream younging of strath level 2 (bottom) (from Reusser et al., 2006).

The purpose of this stop is to observe a flight of Pleistocene terraces and discuss their genesis. Pleistocene terraces are well preserved along this reach of the Susquehanna River underlain by carbonates. Pleistocene terrace stratigraphy along the Susquehanna River was first investigated by Peltier (1949), who demonstrated that there are several alluvial fills underlying terrace treads extending all the way from the head of Chesapeake Bay to heads of outwash at the glacier margin. The precise stratigraphic and genetic relationship between these terraces was more recently investigated by Engel et al. (1996). There are at least six Pleistocene terraces and one possibly pre-Pleistocene terrace remnant preserved here at Marietta (Fig. 6). Water well data and former gravel pits confirm that treads Qt1 through Qt6 are underlain by several meters of stratified sand and gravel. QTg is not a stratified deposit, but rather a scattered lag of rounded clasts of Appalachian provenance, with an occasional diabase boulder. QTg can be traced downstream all the way to the Coastal Plain where it projects both in elevation and in composition to the upper part of the Pliocene Pensauken Formation. If this correlation is correct and QTg is late Pliocene (~2.5 Ma), the incision rate for the river at this point is 20 m/m.y. Terraces Qt1 through Qt6 are distinguished compositionally by containing granite and gneiss, rock types not found in the upper Susquehanna basin that could only have been introduced by glaciation. For this reason, Qt1 through Qt6 are generally held to be genetically related to glacial outwash. Soil chronosequences establish a correlation between the terraces here and those at the heads of outwash 150 km upstream. Qt4 through Qt6 have soils consistent with late Pleistocene (Wisconsinan) age. Qt3 has a significant loess cap and soil development consistent with a preWisconsinan, presumably Illinoian age. Qt2 and Qt1 are thought to be pre-Illinoian age.

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Figure 5. A combined longitudinal profile of the present Susquehanna and its terraces from Harrisburg to Chesapeake Bay (Pazzaglia and Gardner, 1996).

Figure 6. Pleistocene terraces at Marietta (from Engle et al., 1996). Note Qt3 and Qt2 labels are reversed in the legend. W—water well. Stop 2 is located along cross section A–A′.

Susquehanna River Basin There are two competing ideas regarding the preservation of Pleistocene terraces along the river that the strath geometry may help resolve. One possibility is that the Susquehanna River has been incising throughout the Pleistocene, producing the accommodation space to preserve the terraces. Each pulse of glacial outwash provides the tools to carve the strath and then aggrade the terrace alluvium. The alternative explanation is that the Susquehanna River incised to its current elevation in the Pliocene or early Pleistocene and the terrace represent the uneroded, “wings” of thick alluvial fills related to glacial outwash. The latter idea hinges on the fact that more recent glaciations were either less severe or sourced less sediment than earlier glaciations. The terraces at Marietta, on the flank of the broad Susquehanna Valley carved in the Piedmont Lowland, are not present in the deep, narrow gorge visited at Stop 2. Only late Pleistocene strath terraces are present there and remnant gravels on the uplands (Fig. 7). Ascend the terraces and turn left onto 441 (River Road). Go past Bainbridge on the left. At Bainbridge you pass out of the Low Piedmont CambroOrdovician carbonates and onto the Newark Basin redbeds. To the left across the river are the Hellam Hills, held up by the Cambrian Chickies quartzite. After passing through the narrows, cut through a Mesozoic diabase sill and the village of Falmouth. On the left is the infa-

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mous Three Mile Island nuclear-fired steam electric station, site of the 1979 partial core meltdown. The river channel here is cutting across relatively erodable Mesozoic basin red mudstones and sandstones and is again exceptionally wide here with a number of large islands and broad Pleistocene terraces. Enter Middletown, cross Swatara Creek, continuing on 441. At T, turn left onto 230 (W. Harrisburg Pike). Pass Harrisburg airport on the left. W. Harrisburg Pike becomes Martin Ave. Turn left onto Rt. 283 and ascend out of the Susquehanna Valley. Bear right, continuing on Rt. 283 and then bear right onto ramp for I-283 north. Take I-283 north to where it merges with I-83 north, take I-83 north to the next exit (48) and bear right onto ramp for Union Deposit Road. Turn left onto Union Deposit Road and after the road becomes Market Street turn right onto National Civil War Museum Drive at the entrance to Reservoir Park. Follow the drive in a counterclockwise loop around the park, turn left onto Concert Drive, and then bear left onto Lincoln Drive. Park on left at beginning of the parking lot for the Civil War Museum (Lat. 39° 48′ 47.78″ N, Long. 76° 19′ 43.41″ W) and walk up the slope to the top of the hill. Bad weather alternate driving instructions, if Stop 3 deleted: After merging onto I-83, continue north to bear left onto the ramp for I-81 west. Then bear right onto ramp for U.S. 22–322 north. Bear right onto ramp for Linglestown Road, turn left onto Linglestown Road, turn right onto North Front Street and left into Fort Hunter Park (Lat. 40° 20′ 29.79″ N, Long. 76° 54′ 34.94″W).

Figure 7. Representative cross section of the lower Susquehanna River Valley showing the relationship between pre-Pleistocene upland gravels on the High Piedmont and Pleistocene terraces on the Low Piedmont (Pazzaglia and Gardner, 1993).

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Stop 3. Overview of the Susquehanna River Water Gaps at Harrisburg and a Review of the Long-Term Evolution of the Appalachian Landscape and Its Transverse Drainage There is no certainty that William Morris Davis viewed the Susquehanna River water gap from here, but he may have done so. Both Ashley and Fenneman did, and each published a photograph taken from this site (Ashley, 1935, figure 1, plate 123; Fenneman, 1938, figure 55). The fold structure crossed by the river is shown schematically in Figure 8. At this site, the ridge crests of Blue, Second, and Peters mountains were visualized by Davis (1889) and later workers to be accordant summits that represented remnants of the Schooley peneplain of late Cretaceous to early Tertiary age. The rolling hills from Blue Mountain to here were presumed to represent a partial peneplain, the Harrisburg peneplain (Campbell, 1903), that developed on interbedded shales and graywacke sandstones in middle Tertiary times. Reservoir Park is a high point on the Harrisburg surface underlain by thicker beds of sandstone. The fall zone peneplain, the pre-Cretaceous erosion surface under the Coastal Plain sediments, projects far above the present landscape. The supposed Somerville partial peneplain developed on shale and limestone is not present here.

The geomorphic problem to consider here is the development of drainage transverse to the regional structure. How does a river like the Susquehanna become entrenched into resistant rocks and why at this particular location? Davis (1889) had considerable problems explaining these water gaps. He hypothesized a special process of “estuaried” lower reaches of streams during a Cretaceous transgression onto the Schooley peneplain to allow meandering rivers to notch themselves into the resistant ridges. Johnson (1931) hypothesized a Cretaceous transgression across much of a peneplained Pennsylvania with consequent streams developed on the southeastward-dipping Cretaceous sediments that became superimposed on the transverse structure. Such a Cretaceous transgression is not supported by what is now known about Cretaceous-age Coastal Plain stratigraphy. The lower Cretaceous is dominated by fluvial deposits. That the water gaps at Harrisburg are located there due to transverse structural weaknesses such as faults has been proposed by a number of workers (Ashley, 1935; Meyerhoff, 1972; Thiesen, 1983; Hoskins, 1987). The problem has been that there is no prominent offset of the ridges there, and the river makes a dog-leg turn going through the gaps, necessitating separate weakness zones in the northern ridge relative to the southern ridges.

Figure 8. Schematic block diagram of the Susquehanna River water gaps north of Harrisburg, Pennsylvania (Lobeck, 1932, Sheet No. 31). The relative thickness of the resistant Tuscarora and Pocono sandstones is exaggerated relative to the less resistant lowlandforming units. Also the gap through Peters Mountain does not really line up with the gaps through Second and Blue Mountains.

Susquehanna River Basin Thompson (1949) envisioned normal headword erosion being sufficient to cut through the resistant ridges and thought that the river cut through Blue Mountain where it did because the Tuscarora sandstone was particularly thin there. Judson (1975) also thought the gaps were cut by normal headword erosion after Mesozoic rifting started the drainage reversal in the Appalachians. Oberlander (1985) hypothesized superposition of the river from a broad band of Mauch Chunk Formation shale and sandstone when the landscape was at a much higher structural level. Sevon (1986) suggested that the river’s position may have been determined by the front edge of a large thrust sheet that once covered the Anthracite coal region. The late Paleozoic through Mesozoic history of Pennsylvania should be first considered when hypothesizing about the Cenozoic evolution of the Susquehanna drainage and its water gaps (Sevon, 1985, 1986). The Alleghanian orogeny stacked up a 5–9 km thickness of thrust sheets in eastern Pennsylvania (Fig.

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9) (Beaumont et al., 2000). An additional thrust sheet may have been emplaced over the Anthracite coal region (MacLachlan, 1985; Levine, 1986). This continental collision mountain range would have had a relief of 3.5–4.5 km in eastern Pennsylvania (Slingerland and Furlong, 1989). Material eroded from these mountains during the Permian and Triassic was deposited to the northwest as an alluvial plain extending beyond Pennsylvania. Pennsylvanian-age coal rank patterns (Zhang and Davis, 1993) indicate that the thickness of this sediment was ~3.8 km at the Allegheny Front (edge of Appalachian Plateau) and 2.4 km or less at the Ohio-Pennsylvania border. Rifting during the Late Triassic and Jurassic opened basins southeast of and within southeast Pennsylvania, reversing drainage to the southeast on the northwest flank of each basin. In southeastern Pennsylvania, the Gettysburg-Newark rift basin initiated drainage northwestward just south of the Harrisburg area, the ancestral Susquehanna River (Fig. 10). Coarse-grained

Figure 9. Schematic block diagram of the Pennsylvania landscape in the Late Permian at the end of the stacking of thrust sheets during the Alleghanian orogeny continental collision.

Figure 10. Schematic block diagram of the Pennsylvania landscape in the Late Triassic after the Gettysburg-Newark rift basin formed. Reversal of the drainage from northeast to southwest began along the northwest side of the rift basin, the ancestral Susquehanna River (SU) and ancestral Schuylkill River (SC).

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depocenters, with conglomerate units, mark the initiation of the northwest drainage, the Schuylkill first and the Susquehanna later in the Mesozoic. Eroding headword, the Susquehanna would have had some advantages over the Schuylkill that would enable it to become the master stream in the region. First, the Susquehanna is where the entire Blue Ridge Province was faulted out by the Mesozoic rift basin so the Susquehanna was immediately working headword in the Ridge and Valley Province. Second, the resistant coarse-grained sediments in the Paleozoic section thicken markedly to the east of the Susquehanna River’s location. Third, the Schuylkill would have had to work headward directly into the center of the proposed Anthracite thrust sheet while the Susquehanna could have worked around the western and northern edge of the thrust sheet. This would have permitted the Susquehanna to break into the alluvial plain northwest of the mountain range early on and behead the northwest drainage. The details of what happened in the initial and even later stages of the development of Susquehanna River’s course are unknown because it happened several kilometers above the present surface. At Stop 7, it will be emphasized that since the middle Miocene, the past 15 million years, there has been about one kilometer of denudation! The end result of all this erosion into the fold structures of the Ridge and Valley Province is that the drainage is now well adjusted to the structure, and supposedly anomalous situations such as the Susquehanna River water gaps must be explained without peneplanation or Cretaceous transgression. (The material above has been condensed and updated from Sevon, in Gardner et al., 1993.) Continue on Lincoln Drive, bear right onto Civil War Museum drive and again loop around the park. Go past Concert

Drive, turn left onto State St., right onto N17th S., left onto Arsenal Blvd. (U.S. 22), and descend into the Susquehanna Valley. Bear right onto N. Cameron St., continuing on U.S. 22 north. Cross over I-81 on now combined U.S. 22–322. Bear right onto ramp for Linglestown Road, turn left onto Linglestown Road, turn right onto North Front Street and left into Fort Hunter Park (Lat. 40° 20′ 29.79″ N, Long. 76° 54′ 34.94″ W). Stop 3A. Susquehanna Water Gaps and Terraces at Fort Hunter County Park The purpose of this stop is to view the classic Susquehanna River water gaps north of Harrisburg (Fig. 11) and to review the suggested origins for such water gaps. The park buildings occupy a high point on the floodplain that just projected above the largest historic flood on June 1972, Tropical Storm Agnes (Page and Shaw, 1973). The high area is a remnant of a late Wisconsinan recessional outwash terrace level ~8 m (25 ft) above present low water level, Peltier’s (1949) Valley Heads terrace. Much of the city of Harrisburg lies on the 15 m (50 ft) high outwash terrace developed during the late Wisconsinan maximum. A post-glacial aged abandoned channel separates this high point from the eastern bedrock wall of the water gap. A low berm or strath, 2–3 m high, runs along the edge of the present Susquehanna channel. The channel here is on bedrock, as it almost entirely is from the late Wisconsinan terminus to Chesapeake Bay. The geomorphic puzzle here is how the river became entrenched across the ridges of resistant rock and why the water gap developed where it did. The various theories on the origin of this water gap can be place in several groups (modified from Sevon, 1989b): (1) superposition—from an estuary on the

Figure 11. Oblique air photo looking northeast into the water gaps of the Susquehanna River with the northern part of the city of Harrisburg on the right (Capitol dome just off photo on right).

Susquehanna River Basin Schooley peneplain (Davis, 1889) or from overall Cretaceous marine transgression (Johnson, 1931); (2) headward erosion along structural weaknesses (Ashley, 1935; Meyerhoff, 1972; Thiesen, 1983; Hoskins, 1987); (3) resistant sandstone strata were thinner here (Thompson, 1949); and (4) position determined by the front of a large overthrust sheet in the Anthracite region to the northeast (Sevon, 1986). Proof or disproof of the various hypotheses is difficult because on the order of 9 km of rock has been removed since the Alleghanian deformation that produced the folds here and structures at the present depth of erosion do not necessarily reflect those at the level where erosion began (Sevon, 1989b). Leave Fort Hunter, turn left onto N. Front Street, right onto Fishing Creek Valley Road, and left onto U.S. 22–322 north. Go through two water gaps in Second and Peters Mountains, cross the Susquehanna River, and bear right onto U.S. 11–15 north. From Stop 3A to Stop 4 we will cross a syncline, an anticline, a syncline, another anticline, and another syncline (Fig. 12) of the Ridge and Valley Province. The ridges are held up by the Mississippian Pocono sandstone and Pennsylvanian Pottsville conglomerate. Between the ridges we will be crossing areas of rolling hills underlain by the 15,000- to 20,000-ft-thick sequence of mostly red Devonian Catskill and younger interbedded sandstone and shale. In an arid climate this area would look like today’s “painted desert” of Arizona. Bedrock is exposed almost continuously in the bed of the main-stem Susquehanna River from Sunbury to Chesapeake Bay. The gradient steepens slightly across each resistant unit in the Ridge and Valley Province and, as noted before, steepens markedly across the High Piedmont. Approaching Selinsgrove, bear right onto ramp for Rt. 35 and then turn right onto South Market St. Turn left at Sassafras Road (3rd left), continue ahead on Greenridge Road, and turn right onto Clifford Road. Just before railroad tracks, park on left (Lat. 40° 48′ 12.51″ N, Long. 76° 54′ 08.95″ W). Stop 4. Early Pleistocene Head of Outwash and the Pseudo-Moraine at Selinsgrove, Pennsylvania At this site Leverett (1934) observed the knob and kettlelike landscape and assumed that it represented a relict constructional glacial moraine (Fig. 13). This caused him to project a 50 km or 30-mi-long, few-km-wide, ice tongue down the North Branch and main-stem Susquehanna Valley to this point. Long tongues of ice would be a popular way of projecting preWisconsinan glacier borders in the Ridge and Valley Province for later workers up through the 1980s. The problem was that such tongues of ice required nearly zero ice surface profile gradients and impossibly fluid glacial ice. Also the late Wisconsinan border clearly showed that lobation of ice down the strike valleys was only 3–8 km or 2−5 miles, close to what would be expected from a theoretical one bar ice surface profile and from profiles of existing glaciers. The glacial borders on the field trip map (Fig. 1) show Braun’s current understanding of the amount

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of lobation of the glacier’s ice front in the Valley and Ridge of eastern Pennsylvania. Leverett (1934) observed the same knob and kettle-like landscape in the Great Valley in the Allentown area in eastern Pennsylvania and also placed the Illinoian boundary there. Surficial geology mapping in the 1990s indicates that the “moraine” is actually a pseudo-moraine that has a polygenetic origin only initially related to glaciation (Braun, 1996, 1999). The pseudomoraine does mark the southernmost and oldest identifiable glacial border in eastern Pennsylvania, and there are subtle differences in the landscape to either side of that border. Deposits from this most extensive glaciation in the West Branch Susquehanna Valley (Stop 8) and in the Anthracite coal area retain a reversed magnetic polarity and so are of early Pleistocene age (marine isotope stage [MIS] 22+ or pre-Illinoian-G+ age). The glacial deposits under the pseudo-moraine have been undergoing weathering and erosion for at least 880,000 years (MIS 22). The degree of erosion of the deposits is most clearly seen in the slate and shale belt north of Allentown where only the broadest hilltops retain any glacial materials. On a few hilltops where glacial till (diamict) thicker than 6 ft (2 m) remains, numerous depressions form a “patterned” ground effect that suggests a periglacial origin for the depressions (Braun, 1996, 1999). Where only a few feet (<1 m) of glacial deposits remain overlying the slate or shale bedrock, weathering has penetrated and rubified (reddened from oxidation of iron) the bedrock several feet (1– 2 m) below the glacial material. In most of the slate and shale belt the glacial materials have been eroded completely from the hills and redeposited as colluvium in the valleys. Any knob and kettle topography has long since been removed from the landscape. In the carbonate belt around Allentown and here at Selinsgrove, glacial diamicton remnants are more extensive, having been “trapped” by 880,000 years of solutional lowering of the landscape (Fig. 14). Reviews of studies of carbonate denudation rates in Pennsylvania indicate that deposits this old should have been lowered on the order of 100 ft (30 m) or more (Sevon, 1989a; Ciolkosz et al., 1995). There should be at least 3–10 ft (a meter to several meters) of residuum from the dissolution of the carbonates under the “let down” old glacial material (Sevon, 1989a; Ciolkosz et al., 1995, White, 2000) and this is what is observed under the pseudo-moraine (Braun, 1996, 1999). Also there is a subtle but distinct pattern of thinner diamict remnants on the hilltops and thicker remnants in the valleys and especially in solutional depressions (seen in both outcrop and borehole observations) indicating erosional redistribution of the original glacial deposits. The pseudo-moraine does have a relatively thicker and more continuous glacial mantle than elsewhere and probably once had a genuine morainic topography ~100 ft above the present landscape (Fig. 14). Now though the deeply eroded and “let down” colluvium (collapse-uvium?) derived from glacial material is draped over underlying bedrock features (Fig. 14). In places on the present knob tops in the pseudo moraine, colluvium derived from carbonate residuum overlies the colluvium derived from the glacial diamict (Braun, 1996, 1999). This

Figure 12. Schematic block diagram of the central Susquehanna region from Selinsgrove to Harrisburg (Lobeck, 1932, sheet 34). Stops are numbered squares.

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Susquehanna River Basin suggests repeated episodes of topographic reversal as the lowering of the landscape has proceeded. The present swell and sag topography is composed of smaller scale periglacial depressions that developed in the reworked glacial material and that are superimposed on larger scale bedrock solution features. (For more discussion of periglacial activity in this region, see several review papers in Clark and Ciolkosz, 1988; Clark et al., 1992; Braun, 1994b; and Marsh, 1999) The numerous smaller scale depressions commonly contain wetlands and perennial ponds that make the pseudo-moraine landscape distinctly different from the surrounding more “karstic” landscape. Thick sand and gravel deposits in ice marginal kames at the base of South Mountain at Allentown have been partly buried by extensive periglacially derived boulder colluviums and further attest to the reshaping of the landscape by periglacial activity. At this early Pleistocene terminus site in the Susquehanna Valley the original deposit should have been a large head of outwash like that at the late Wisconsinan terminus. The upper few meters or several feet is a roundstone clast-dominated diamicton that does look till-like. A number of excavations in the 1980s and 1990s just to our east for a housing development and for Susque-

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hanna University showed oxidized and weathered but still stratified sand and gravel on the side slopes of valleys cut into this feature. So indeed this is a very old, weathered, and eroded head of outwash. This site is underlain by a broad belt of limestone of the Keyser and Tonoloway Formations on the plunging nose of an anticline. That means that while the pseudo-moraine is presently at ~180 m or 600 ft elevation, the original top surface of the head of outwash was near 210 m or 700 ft elevation. The Susquehanna River today at Selinsgrove is at an elevation of 122 m or 400 ft, 60 m or 200 ft below the pseudo-moraine top. The late Wisconsinan head of outwash on the North Branch Susquehanna River is at most 49 m or 160 ft above the present river bed. If the early Pleistocene head of outwash here had a similar thickness, then the river has incised only ~12 m or 40 ft more than limestone has dissolved, for a total incision of 60 m or 200 ft since glaciation. Continue across the railroad tracks and turn right onto U.S. 522. Bear left onto N. Market St., continuing on U.S. 522 north. Merge with U.S. 11–15 continuing ahead (north) through the strip development of Hummels Wharf and Shamokin Dam. Turn left on U.S. 11 and continue along the west bank of the Susquehanna River. Just before bridge, turn left onto County Line Road

Figure 13. Topographic map of the Selinsgrove area with Leverett’s (1934) Illinoian moraine area outlined by a solid line. The closure on the depressions is too small to show with the 20 ft contour interval. Braun’s current interpretation of the position of the early Pleistocene (marine isotope stage 22+) glacial terminus is shown as a dotted line.

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and then turn sharply right into Shikellamy State Park. Ascend slope and park on right at overlook (Lat. 40° 52′ 54.46″ N, Long. 76° 48′ 06.46″ W). Stop 5. Overview of the Confluence of the West and North Branches Susquehanna River at Sunbury and Discussion of the Use of Relative Degree of Weathering in Mapping Pre-Wisconsinan Glacial Deposits This vantage point at Shikellamy State Park overlooks the confluence of the West Branch Susquehanna River (on the left) with the North Branch Susquehanna River (straight ahead) to form the main-stem Susquehanna River (on the right), the largest river in the eastern United States (Fig. 15). The West Branch has been flowing south across strike to here while the North Branch has been flowing down strike, obliquely cutting through the secondary Trimmers Rock sandstone ridge to reach here. The mainstem Susquehanna River from here will flow transverse to all Appalachian Ridge and Valley and Piedmont structural trends to reach Chesapeake Bay. The surficial geology map of the western-central Susquehanna lowlands was done using relative soil profile development to define and map the units (Fig. 15) (Marchand and Crowl, 1991; from field work primarily done in the 1970s; Marchand, 1978).

The map has a problematic “patch-work quilt” pattern of supposedly younger glacial materials surrounded by supposedly older glacial materials. This “patch-work quilt is not what is observed within the late Wisconsinan terminus (Braun, 1994a). On the central Susquehanna map, supposedly early Wisconsinan till (dt) and late Illinoian White Deer till (wt) patches are surrounded by supposedly older early Illinoian Laurelton drift (Id). The two small early Wisconsinan till patches are on the Bloomsburg redbeds, while the White Deer till is noted as having a limestone component, unlike that of other pre-Wisconsinan drifts. This suggests that there are problems separating parent material effects from time effects in these weathering profiles. Also, the younger material is far less continuous than the older material, the opposite of what would be expected. The undifferentiated pre-Wisconsinan deposit map unit covers the most area and indicates that differentiating the units by weathering profile characteristics is problematic more often than not. It should also be noted that work since the 1970s has shown there is no early Wisconsinan material or early Illinoian material here and even late Illinoian material is questionable (Braun, 1988, 1994a, 2004a; Ridge et al., 1990). The material in this area should be of some pre-Illinoian age (MIS 12+). In addition, the glacial borders drawn on the basis of the mapped till patches produce exceptionally long and low to reverse gradient ice lobes in the strike valleys that require impossibly

Figure 14. Lumpy landscape of the “pseudo-moraine” is from the at least 880,000 years of limestone dissolution since glaciation that has lowered the land surface on the order of 100 ft. The “kettles” are sinkholes with smaller periglacial landforms on their flanks (Braun, 1999).

Susquehanna River Basin fluid ice. The boundaries drawn actually represent lines drawn at the base of the ridges separating where the glacial deposits have been completely eroded away from where there are remnant glacial deposits. The central Susquehanna surficial geology map (Fig. 15) shows that depending on the degree of weathering profile development leads to glacial deposit and glacial boundary map patterns that are totally unlike that of the late Wisconsinan and in some regards are physically impossible. Probably what these map patterns are showing is the complex interplay of different depths of truncation of weathering profiles on different parent materials on different landscape positions. It is highly unlikely that there are any nontruncated weathering profiles in this area of early Pleistocene, older than 880 ka landscape. For instance, how does one reliably differentiate between a deeply truncated

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Penny Hill soil (supposedly the oldest till) and a less deeply truncated Laurelton soil? What this work may best demonstrate is that in the eroding, moderate relief Valley and Ridge landscape, one may accurately differentiate the remnant amount of weathering profile development but that differentiation does not yield a reasonable distribution or age for pre-Wisconsinan glacial material. These various drift deposits can be explained more simply and rationally as a complex of different degrees of truncation of a pre-Illinoian weathering profiles coupled with different lithofacies of pre-Illinoian drift. (The material above was condensed from Braun, 1994a; late Wisconsinan–pre-Illinoian(G?) glacial events in eastern Pennsylvania.) As noted at Stop 1, Shaw (1989) proposed catastrophic floods of channeled-scabland scale formed the drumlin field in upstate New York and came down the North Branch

Figure 15. Surficial deposit map of the central Susquehanna Valley region (Marchand and Crowl, 1991). The map shows a “patch-work quilt” pattern of younger and older glacial deposits. Within the late Wisconsinan limit, no such pattern is observed. MIS—marine isotope stage.

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Susquehanna River during recession of the late Wisconsinan. Shaw’s technique of estimating drumlin-forming subglacial discharges yields a peak discharge of 4.8 × 106 m3/s for the North Branch Susquehanna River. Braun (1990, 2008) noted that evidence of such large-scale flooding along the valley of the North Branch Susquehanna was lacking, particularly in the Bloomsburg area to the east of here. Near Bloomsburg the North Branch Susquehanna River was blocked by pre-Wisconsinan glacial deposits (MIS 6 or 12) and was diverted down a strike valley to a narrow water gap through a strike ridge of one of its tributaries, the Rupert water gap (Braun et al., 1984). Due to the “youthful” age of the Rupert water gap, it is less than one-half the cross-ectional area of any water gap upriver of it. Thus the “undersized” Rupert water gap is the critical “choke point” for hypothesized catastrophic floods and should show dramatic evidence of scabland-like erosion features. Also a low-level bypass for the flood water, the continuation of the strike valley to the west that the river leaves as it enters the Rupert gap, should show significant scabland-like features. No such scabland-like erosion or deposition features are observed there. Instead of flood features, undisturbed, deeply weathered pre-Illinoian glacial deposits (like at Stop 4) mantle the strike valley floor (also noted by

Leverett [1934] and Peltier [1949]), often capped either by Wisconsinan colluvium or Wisconsinan loess. Return to and turn right onto County Line Road. Turn right onto U.S. 15 north, continuing north through Lewisburg, across I-81, and through Allenwood. Turn right onto Rt. 54. In a couple of miles turn left onto Brouse Road and then left onto State Home Road. Pull over on the right side of State Home Road where it starts to descend the flight of terraces stepping down to the river directly ahead (south) (Lat. 41° 12′ 00.37″ N, Long. 76° 50′ 10.73″ W). Stop 6. Mid-Pleistocene to Holocene West Branch Susquehanna River Terraces near Muncy, Pennsylvania The West Branch Susquehanna River takes a sharp southward turn in the vicinity of Muncy (Fig. 19). On the west side of the river bend, a series of bench levels have developed. The Penn State Soil Characterization Laboratory sampled the benches from the river to the footslope area of the adjacent ridge as a part of the national cooperative soil survey of Lycoming County. These data are a part of the Penn State Soil Characterization Lab’s computer database (Ciolkosz, 1999). These bench levels were subsequently studied by Engel et al. (1996). Figure 16 gives an

Figure 16. Topographic map, surficial geology, and cross section of Muncy, Pennsylvania, showing the distribution of terraces and soil pit locations. w— well log data showing depth to bedrock (Pennsylvania Geological Survey). BEM is Bald Eagle Mountain (Engel et al., 1996).

Susquehanna River Basin interpretive drawing of the bench area. The benches below the Qt4a level were flooded by the Agnes Flood of June 1972 (flood of record for this area) and are therefore floodplains. The Qt4a level is extensive and probably Wisconsinan age (Fig. 17). The three levels above Qt4a are all loess covered with what appears to be Wisconsinan loess. Thus, the three upper benches are preWisconsinan in age, and they are underlain by diamictons. Engel et al. (1996) correlated them as progressively older terrace levels with the oldest being early Pleistocene in age (~770 ka to 970 ka) (Fig. 17). Continue on State Home Road and descend the flight of terraces. Turn right onto Rt. 405, in Montgomery continue ahead on Rt. 54, immediately turning right, and stay on Rt. 54 west. Turn

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right onto U.S. 15, ascend Bald Eagle Mountain, and go through a deep, narrow wind gap. Descending the mountain, turn right into overlook beside the road (Lat. 41° 13′ 39.13″ N, Long. 76° 56′ 14.13″ W). Stop 7. Glacial Lake Lesley, the Wind Gap through Bald Eagle Mountain, and How the Three-Tier Landscape Developed: Davis, Hack, and More Recent Hypotheses Glacial Lake Lesley and Glacial Retreat Direction in the Bald Eagle Valley We are now at the east end of Glacial Lake Lesley (Fig. 18) when the lake had its greatest potential extent and depth (~150 m

Figure 17. Correlation chart showing time scales, central Pennsylvania glacial stratigraphy, and Susquehanna River terraces. R—Sediments with reversed magnetic polarity. Dashed lines represent poorly constrained ages. Engel et al. (1996) correlated them as progressively older diamicton terrace levels with the oldest being early Pleistocene in age (~770 ka to 970 ka). References: Gardner et al. (1994); Peltier (1949); Pazzaglia and Gardner (1994).

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or 500 ft). The lake was spilling over the divide to the Juniata River in the Bald Eagle Valley 112 km or 70 mi to our west (Williams, 1895; Leverett, 1934; Ramage et al., 1998). The bedrock floor of that outlet was at an elevation of ~311 m or 1020 ft (borehole data in Ramage et al., 1998). As will be noted at Stop 8, the lake sediments exhibit reversed magnetic direction, so the lake is early Pleistocene in age (MIS 22+). As that glacier receded to this area, the lake may have already started spilling around the plunging anticlinal nose of Bald Eagle Mountain because the glacier should have been retreating to the northeast. Striation directions are all northeast-southwest in the late Wisconsinan and Illinoian (or older) glaciated areas and at a single site so far discovered in the early Pleistocene glaciated area south of Bald Eagle Mountain. The Illinoian (or older) glacial limit is trending southeast-northwest (see Fig. 1) a few km north of Bald Eagle Mountain (ongoing mapping by Braun). So the early Pleistocene glacial ice-front should have also had southeast-northwest trend, rather than the more south-north trend in Figure 18 (from Bucek, 1975) or a southwest-northeast trend (Ramage et al., 1998, Fig.

1). A southeast-northwest ice-front would have left a lesser thickness of ice against for a calving bay re-entrant to cut through to drain the lake. Ongoing mapping by Braun suggests that a middle Pleistocene event (MIS 16? on Fig. 1) should have also dammed a lake in the West Branch Susquehanna Valley but thus far no varves with believable normal polarity have been found above the varves with reversed polarity. Both normal and reversed polarity varves have been discovered in the middle Anthracite Coal Basin beyond the Illinoian or older limit (Sasowsky, 1994). The Wind Gap through Bald Eagle Mountain At this site the plunging anticlinal nose of Bald Eagle Mountain, held up by the Tuscarora sandstone, is cut by a 215 m or 700-ft-deep wind gap whose bottom elevation is presently 367 m 1205 ft (Fig. 19). The floor of the center of the wind gap is covered by boulder colluvium but bedrock is exposed starting at least at 340 m or 1120 ft in a gully on the north side of the floor of the wind gap. Immediately north of the wind gap and Bald Eagle Mountain, the West Branch Susquehanna River is at an elevation

Figure 18. The maximum extent of Glacial Lake Lesley, drawn along the 1100 ft contour. Its outlet at Dix (double-line arrow) drained southwest into Juniata River drainage. Glacial limits of various ages are shown as dashed lines labeled with marine isotope stage (MIS) numbers (modified from Gardner et al, 1993, figure 4.6).

Susquehanna River Basin

Figure 19. Topographic map of the Williamsport area showing the wind gap through Bald Eagle Mountain and the three-tier landscape in the Ridge and Valley and Plateau Provinces. Contour interval = 20 m.

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152 m or 500 ft, 188 m or 620 ft or so below the bedrock floor of the wind gap. The nearness in elevation of the wind-gap floor and the highest level of Glacial Lake Lesley has led some to informally suggest that the gap was an outlet for the lake. But the cutting of the wind gap started at the 580 m or 1900 ft level in the top of the mountain, well above lake level. Also 215 m or 700 ft of incision into the Tuscarora sandstone could not have occurred in the few hundred or thousand year life of the glacial lake. The cutting of the wind gap is from stream capture around a plunging anticlinal nose, probably sometime in the Pliocene. It is a type of stream capture peculiar to fold belts where a larger trunk hung up on resistant bedrock is captured by one of its own tributaries working headword around the nose of the plunging anticline on much weaker bedrock. As the entire landscape was eroding down, the West Branch Susquehanna would have found itself incising into the Tuscarora sandstone at the crest of the plunging anticline. That would have slowed its incision and greatly increased its gradient at the site. This would have kept the river upstream of the gap from incising at the rate of the river downstream of the gap—it would have been “hung up.” A strike valley on much weaker shale and limestone would have existed around the plunging nose of the anticline, as it does today. An east bank or down-plunge direction tributary to the river on the south side of the plunging anticlinal nose would be able to readily work headword around the nose of anticline. Eventually the tributary would capture the river on the north side of the anticline, bringing the river around the nose of the anticlinal ridge, its present position. The capture would have been accelerated by headward spring sapping from subterranean flow through the limestone core of the strike valley. The process is starting again today at the axis of the anticline where the river now has an abruptly steepened gradient (Fig. 22). How the Three-Tier Landscape Developed: Davis, Hack, and More Recent Hypotheses The final purpose of this stop is to view the three-tier landscape in the Ridge and Valley and adjacent Plateau (Fig. 19) and review past and current explanations of how that landscape developed. There have really been only two explanations for the Appalachian landscape, Davis (1889, 1899) and Hack (1960, 1965). Davis explained the three-tier landscape as being remnants of three peneplains (low relief fluvial erosion surfaces) of differing age (Fig. 9). Davis thought the highest peneplain devel-

oped during the Cretaceous and named it the Schooley peneplain (Davis and Wood, 1890). That peneplain was uplifted in the early Tertiary, and a partial peneplain developed during the middle Tertiary (named the Harrisburg peneplain by Campbell, 1903). A second uplift occurred in the late Tertiary and then a second partial peneplain developed locally on the weakest rock (named the Somerville by Davis and Wood, 1890). This explanation requires erosion to carve out only the present relief of ~500 m or 1500 ft since the Cretaceous. Later workers mainly added to the number of peneplains (Bascom, 1921), extended the Coastal Plain cover (Johnson, 1931), or made the peneplains younger (Johnson, 1931; Ashley, 1935). Hack explained the three-tier landscape as being a result of continuous erosion working on rocks of three differing resistances that produces the three different elevations. The hardest sandstone forms the highest ridges, interbedded sandstone and shale form an intermediate elevation rolling-hill landscape, and the shale and limestone units form the gently rolling lowest elevations (Figs. 19 and 20). A concise comparison of various elements of the Davis and Hack theories is given in Table 1. The crucial test of Davis’ theory is to examine the amount of sediment eroded from the Appalachians and deposited in the Coastal Plain and offshore. Long-term rates of post-Triassic erosion have been reconstructed from the offshore sediment load in the Baltimore Canyon Trough (Poag and Sevon, 1989) (Fig. 21). In summary, exhumation was relatively fast in the Late Cretaceous (>30 m/m.y. assuming constant drainage basins size), very

Figure 20. Schematic cross section of the breached Berwick anticline showing the relationship of bedrock units to the three-tier landscape and postulated peneplains.

TABLE 1. DAVIS VERSUS HACK: APPALACHIAN LANDSCAPE EVOLUTION Davis: Geographic cycle Hack: Dynamic equilibrium 1. Time-dependent landscape 1. Time-independent landscape 2. Cyclic uplift and stillstand 2. Continuous uplift—varying rate 3. Episodic or cyclic erosion 3. Continuous erosion—varying rate 4. Sequential landscape—youth, maturity—old age, relief reduction 4. Nonsequential landscape change—perpetual maturity 5. Overall structure truncated by topography 5. Structure and topography closely adjusted to each other 6. Rocks of differing resistance develop differences in elevation and 6. Rocks of differing resistance preserve peneplain remnants of differing ages slope angle 7. Summit accordance—evidence for past peneplain 7. Summit accordance from subequal stream spacing on rocks of similar resistance

Susquehanna River Basin slow in the Tertiary (<30 m/m.y. assuming constant drainage basins size), and then very fast in the Miocene, continuing to the present. The slug of Miocene–Holocene sediment in the Baltimore Canyon Trough requires removal of at least 1–2 km of rock in the past 20 m.y (Braun, 1989b; Pazzaglia and Brandon, 1996; Pazzaglia and Gardner, 2000). This makes it unlikely that a preMiocene peneplain could be preserved in the present landscape. This continuous, rapid mid-Miocene to present erosion is what has formed the three-tier Ridge and Valley landscape (starting 1 km or so above the present ridge crests), and that continuity of erosion is what Hack emphasized in his later work (1975). It is the dynamic equilibrium part of Hack’s model that is most problematic. Uncertainty in measuring basin-wide erosion rates and river-specific incision rates makes it difficult to know whether the Appalachian landscape is maintaining relief in a dynamic equilibrium state (Hack, 1960), or relief is locally increasing or decreasing. Little separation between the straths of successive glacio-fluvial terraces on the Susquehanna suggests that the river is incising slower than periglacial activity is eroding the hilltops, essentially reducing relief in the landscape. However, close to the fall zone, the river profile is clearly convex indicating base level fall and active incision. Process rates are slow in this landscape, and there are significant lag times (Pazzaglia et al., 1998) making it very difficult to determine whether the Appalachian landscape is in dynamic equilibrium or in some disequilibrium state. Looking back further in time at the entire post-Triassic sediment yield curve (Fig. 21) shows that there were two earlier pulses of clastic sediment delivered to the Baltimore Canyon Trough, each followed by an exponential decrease to a very low sediment yield. This fits the Davisian model of landscape evolution and also fits with two major unconformities at the edge of the

Figure 21. Cross section of the Baltimore Canyon Trough (BCT, top) and the flux of eroded rock to the BCT over the past 180 m.y. (bottom). Note the pulses of sediment, followed by exponential decay, as well as the large pulse of sediment in the past 20 m.y. (from Pazzaglia and Brandon, 1996).

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Coastal Plain, one underneath the late Cretaceous sediments and the other under the mid-Miocene sediments. Even before the Late Jurassic sediment pulse shown on Figure 21, there had been deep unroofing of the Appalachians in the Triassic as shown by the erosion surface under the rift basins and by the thermochronologic data. Much of that eroded material was shed in a vast alluvial plain to the west and north, the direction of drainage at the close of the Alleghanian orogeny in the Permian. Triassic continental rifting rejuvenated the Appalachians, opened the Atlantic Ocean, and began the reversal of the drainage to the east. The exponential decline in sediment yield from the Late Jurassic to the close of the Early Cretaceous (160–100 Ma) (Fig. 21) is a result of the Triassic rift topography being reduced and beveled to form the fall zone low-relief erosion surface (unconformity). Davis interpreted that surface to be a peneplain. A Late Cretaceous (~90 Ma) transgression then brought fluvial-deltaic and possibly shallow marine deposits across the fall zone and possibly farther west. The Late Cretaceous sediment pulse (Fig. 21) indicates renewed uplift and erosion in the Appalachians for a yet unknown reason. That uplift was followed by another exponential decline in sediment yield that lasted for 65 million years from the latest Cretaceous to the early Miocene. This decline in sediment yield is related to a second reduction in relief in the Appalachians toward a low-relief erosion surface, the Schooley surface. Davis interpreted that surface to be a peneplain; another possible interpretation is that it is a pediplain or an etchplain. From limited subsurface information, both the earlier fall zone and Schooley unconformities have relief similar to the present fall zone landscape (Gardner et al., 1993), suggesting there were never truly low-relief erosion surfaces in the Appalachians. Recent work supports an early Miocene age for Schooley surface in the New Jersey Highlands west of the Coastal Plain (Stanford, 1997; Stanford et al., 2001). Along the Susquehanna River, the Piedmont upland surface next to the Coastal plain is also considered to be of Miocene age (Pazzaglia and Gardner, 1994). It is expectable that the area next to the hinge line between erosion and deposition retain remnants of an erosion surface dating from before the mid-Miocene sediment pulse. But again, to account for the volume of sediment produced, much greater erosion is necessary upstream, and that erosion has removed all vestiges of a Miocene erosion surface farther inland. The cause of the mid-Miocene sediment pulse, particularly from the Susquehanna basin, is problematic since there is no known tectonic event that could cause the sediment pulse (Sevon, 1989b). Here we present two alternative views: Pazzaglia suggests that the Susquehanna drainage divide remained near the Piedmont– Ridge and Valley boundary until the early Miocene and then rapidly shifted westward to its present position on the Appalachian Plateau in the middle Miocene. That rapid headward erosion would produce the sediment pulse offshore. New U-Th/He thermochronologic data and existing apatite fission-track (AFT) thermochronology reveals a denudation pattern consistent with a systematic, recent shift in the

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drainage divide linked to a similar shift in the depositional basins accumulating unroofed sediments on the Coastal Plain (Naeser et al., 1999, 2001). In the central Appalachians, U-Th/He and AFT closure ages are virtually identical for the Ridge and Valley province and suggest very modest amounts of unroofing in the past ~200 Ma. In contrast, the U-Th/He ages are systematically younger than AFT ages to the east in the Piedmont and Blue Ridge provinces, suggesting more and continuous unroofing during Alleghanian orogenesis and Mesozoic rifting. The problem with these data, particularly in Pennsylvania, lies in the fact that the largest amount of unroofing coincides with the lowest contemporary erosion rates and lowest standing topography, whereas the least amount of unroofing is associated with the highest contemporary erosion rates and the most rugged and highest standing topography. A way to explain this apparent paradox is to envision a drainage divide that forms on the rift flank in the Mesozoic and remains pinned, east of its current position for much of the Mesozoic and early-mid Cenozoic. Only in the late Cenozoic has the divide jumped westward to its current position, leaving a dissected and rugged topography in its wake, but not yet encompassing enough time to lock-in Cenozoic ages for the low-temperature thermochronometer. The rapid, westward shift in the divide may actually be rolling southward along orogen strike as suggested by: (1) the location of the highest rates of Appalachian stream incision, (2) considerable separation of the crest of the range and the crest of the long-wavelength topography (defined as the topography filtered at the flexural wavelength), and (3) a southwest sweep of Miocene to Quaternary depocenters on the Atlantic Coastal Plain (Poag and Sevon, 1989). In contrast, Braun envisions that the Susquehanna basin eroded headward to near its present divide by the beginning of the Tertiary (Judson, 1975, argued for a beginning of Cretaceous date). Then in mid-Miocene times a critical climate threshold was passed that initiated rapid erosion of the entire basin. At that time the landscape was underlain by a very deep weathering profile developed during the previous 40 million years of low sediment yield (Fig. 21). That landscape had a moderate relief equal to or greater than that of today, Garner’s (1974) Humid Climax Topography. That deeply weathered moderate relief landscape permitted the very rapid erosion that produced the mid-Miocene sediment pulse. Much of this scenario was also suggested by Sevon (1989b, 1999) but starting with a low-relief landscape. The Susquehanna River is located where Triassic rifting most deeply cut into the Appalachians and faulted out the Blue Ridge– Reading Prong. During most of the Triassic a southeast draining river entered the north side of the rift basin and formed a major depocenter 50 km east of the present course of the river (Smoot, 1999). That river is the likely ancestor to the Susquehanna and marks the early initiation of the reversal of drainage from the northwest to the southeast as envisioned by Davis (1889). The volume of sediment implies significant headward erosion of the ancestral Susquehanna in the Triassic. Another option is that the depocenter east of the present Susquehanna marks the ancestral

Schuylkill River, and a depocenter west of the present Susquehanna marks the ancestral Susquehanna (Poag and Sevon, 1989; Sevon, 1989b). The Triassic erosion was then followed by two other episodes of erosion in the Jurassic and Cretaceous that produced the sediment pulses shown on Figure 8. The Susquehanna was a primary source of sediment for the Jurassic pulse, a secondary source for the Cretaceous pulse, and a primary source for most of the Tertiary (Poag and Sevon, 1989). There should have been sufficient Mesozoic–early Tertiary erosion to extend the Susquehanna near to its present head. Whereas the entire length of the present Appalachian drainage divide does not coincide with the regional geologic structure, physiographic provinces, or gravity highs, it does roughly parallel the edge of the Coastal Plain and the present shoreline (Judson, 1975). This parallelism implies that the overall length of the river systems from the Susquehanna southward has been similar since the Cretaceous. Judson (1975) also noted that the greatest negative gravity anomaly coincides with the drainage divide along the Blue Ridge front in the southern Appalachians. He argued that the crust is much thicker there and that is what has “pinned” the divide to near that position and that is where the landscape is most in disequilibrium (Hack, 1973). Northward the crust is much thinner and that has permitted a more rapid, earlier migration of the divide across the Ridge and Valley. Compared to other rivers to the south, the Susquehanna had the least geologic work to do, eroded headward the most readily, the earliest on, and achieved a “mature” divide position where the stream gradients immediately to either side of the divide are essentially equal. (This discussion has been condensed from that in Braun et al., 2003 and Pazzaglia et al., 2006.) Turn right onto U.S. 15 and continue along the north flank of Bald Eagle Mountain into South Williamsport. Cross the West Branch Susquehanna River, continuing straight ahead onto Market St. Turn left onto W. 4th St. and then right at the Genetti Hotel at 200 W. 4th St. (Lat. 41° 14′ 26.88″ N, Long. 77° 00′ 18.49″ W). DAY 2. WILLIAMSPORT, PENNSYLVANIA, TO BALTIMORE, MARYLAND Leave hotel turning right onto W. 4th St, left onto Hepburn St., and right onto ramp to I-180 west. Where U.S. 15 turns north, I-180 becomes U.S. 220 west. Continue on U.S. 220 to Jersey Shore. For the next ten miles or so we will be driving on or next to broad, late Wisconsinan outwash terraces and higher remnants of Illinoian or older terraces (Fig. 22). Approaching Jersey Shore, bear right onto ramp to N. Main St. turn left onto Rt. 44 (Allegheny St), and immediately cross the West Branch Susquehanna. Ascend a gentle slope and turn left onto Front St. in the village of Antes Fort. At the firehouse, park on the right next to the railroad tracks (Lat. 41° 11′ 26.36″ N, Long. 77° 13′ 34.28″ W). We will walk west up the tracks to the Stop 8 exposure just east of the Rt. 44 bridge over the railroad. BEWARE: you may have a close encounter with a coal train coasting quietly down the gentle grade from the west.

Susquehanna River Basin Stop 8. Early Pleistocene Till and Varves, Mid- to Late Pleistocene Colluvium, and Late Pleistocene Outwash Terraces at the Antes Fort “Fan” in the West Branch Susquehanna Valley Williams, who named Glacial Lake Lesley in 1895, called the landform at Antes Fort a fan-cone (Fig. 23). He thought that it formed from a torrential discharge of glacial origin through the Antes Creek gap in Bald Eagle Mountain from the Nippenose Valley on the south side of the mountain (Williams, 1917, 1920). Leverett (1934) used both fan-cone and alluvial fan to describe the landform. He noted that in the railroad cut 27 ft of gravelly material was on top of laminated clay that extended down to shale. He also thought that the 660 ft (200 m) elevation of the top of the fan marked the maximum level of the glacial lake. Andreus (1993) did a more detailed description of the railroad cut (Fig. 24) and found that the gravel was only a few feet thick veneer over tens of feet of glacial lake sands and varves, and, at the west end, glacial till under the varves. The apex of the fan-like landform is actually a bedrock hill capped by glacial sediments. The gravelly veneer is alluvium deposited by Antes Creek after the lake drained and before the river and the Creek incised into the glacial deposits. The early Pleistocene till and varves are essentially unweathered at this site thanks to burial by overlying sands and gravelly alluvium. The varves are reddish from clay derived from the thick redbed sequence in the Devonian Catskill Formation (Fig. 25). They retain a strong reversed magnetic polarity signature and place the material in the early Pleistocene (MIS 22+) (Gardner et al., 1993; Braun et al., 1993). It has been argued elsewhere

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(Braun 1989a, 1994a, 2004a; Ramage et al., 1998) that these sediments were deposited in MIS 22 since that was the first of the major cold events of the past million years and that this landscape has been eroding relatively rapidly as was discussed at Stop 7. On the mountain toe-slope immediately south of the fan-like landform are bouldery colluvial deposits from periglacial gelifluction activity in mid- and late Pleistocene times. The ubiquitous boulder colluvium deposits in eastern and central Pennsylvania have been highly stable (not being eroded or deposited) under the present climate and have many morphological features that indicate a gelifluction origin (Clark and Ciolkosz, 1988; Clark et al., 1992; Braun, 1989a, 1994b). Immediately north of the fan-like landform are a series of late Pleistocene outwash terraces on the inside of a large incised meander loop of the West Branch Susquehanna River. At this site, Peltier (1949, table 41) had noted a 1000-ft-wide Binghamton terrace ~25 ft above the river and a 3000-ft-wide Olean terrace, ~45 ft above the river. The Olean terrace represented outwash from the late Wisconsinan terminus, and the Binghamton terrace represented conditions when the late Wisconsinan terminus had retreated to the Pennsylvania–New York border. The West Branch Susquehanna was receiving no meltwater by the time the glacier had retreated to New York so the Binghamton terrace represented a temporary stabilization of the post-glacial downcutting of river. Peltier did not note that at the village of Antes Fort there is an Illinoian or older terrace remnant 60–70 ft above the river, but he noted such remnants at that elevation elsewhere along the West Branch Susquehanna. Braun (1995) corroborated Peltier’s terrace mapping at the site.

Figure 22. Terraces of the West Branch Susquehanna River (Peltier, 1949, figure 35). The Olean or late Wisconsinan maximum terrace typically lies 14 m or 45 ft above the present river. The steepened gradient of the present river between Williamsport and Muncy is where the river is beginning to incise into the sandstone of the plunging nose of Bald Eagle Mountain.

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Return to U.S. 220 bearing right onto U.S. 220 east. Entering Williamsport, bear right onto ramp to U.S. 15 north and drive upstream along the Lycoming Creek Valley. Bear right onto ramp to Beautys Run Road, turn right and then left onto Lycoming Creek Road. Where that road merges with Rt. 110, continue ahead and in a couple of miles enter the narrow gorge of the diverted Lycoming Creek. Then turn left onto Saint Michaels Road and pass under U.S. 15. Ascend the buried valley of Lycoming Creek to the intersection of Bobst Mountain Road on the right. Turn

right onto Bobst Mountain Road and immediately park on the right (Lat. 41° 20′ 37.37″ N, Long. 77° 06′ 23.94″ W). Stop 9. Mid-Pleistocene (Illinoian or Older) Terminus and the Buried Valley of Lycoming Creek We have just entered the Appalachian Plateau and have gone through the relatively narrow, deep diverted course of Lycoming Creek (Fig. 26) at Powys, Pennsylvania. We then drove up onto

Figure 23. Topography and surficial deposits of the Antes fan-like feature at Antes Fort (Andreus, 1993, figure 9).

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Figure 24. Stratigraphic section of the north face of the railroad cut at Antes Fort (Andreus, 1993, figure 10).

Figure 25. Detailed stratigraphic column of the early Pleistocene deposits in the railroad cut at Antes Fort (Gardner et al., 1993, figure 2.11). AF1 and AF2 are samples with reversed magnetic polarity (Gardner et al., 1994).

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Figure 26. Surficial geology map (Faill et al., 1977; surficial geology by Sevon) of the area around the Lycoming Creek diversion site (Stop 9). The large ice-contact stratified drift deposit at Stop 9 was originally interpreted as a late Wisconsinan (Woodfordian) deposit by Sevon. All other glacial deposits in the area were thought by Sevon to be of early Wisconsinan age (Altonian). Gravel pit exposures show the material at Stop 9 and at other sites in the area where the material retains most of its weathering profile to be deeply weathered and thus of Illinoian or older age (Qiic) rather than early or late Wisconsinan age. Holocene–late Wisconsinan material: Qaf—alluvial fan, Qac— alluvium-colluvium, Qc—colluvium, Qbc—Boulder colluvium. Qw denotes supposed Woodfordian-age material: Qwoa—alluvium-outwash, Qwic—ice-contact stratified drift. Qa denotes supposed Altonian-age material: Qaoa— alluvium-outwash, Qaic—ice-contact stratified drift, Qat—till.

Susquehanna River Basin the large mass of ice-contact stratified drift that forms a head of outwash that blocks the pre-glacial course of Lycoming Creek. We are now standing on the local drainage divide, the top of the head of outwash, with the steeper ice-contact edge face to our northeast and the gentler outwash ramp to our southwest (Fig. 26). There are ~200 ft of sand and gravel under our feet. Sevon (in Faill et al., 1977) mapped the surficial geology of the area around Stop 9 (Fig. 26) and recognized that the mass of sand and gravel blocked the pre-glacial course of Lycoming Creek, diverting meltwater and post-glacial drainage through a saddle at Powys, and cutting the present narrow gorge of Lycoming Creek from Powys to Haleeka (Fig. 26). He interpreted the material to be of late Wisconsinan age (Woodfordian) on the basis of the shallowness of the soil profile and the preservation of original constructional topography. He thought that an early Wisconsinan (Altonian) glacier had advanced a few miles farther south across the area. There, till primarily derived from reddish bedrock appeared to be somewhat more weathered and eroded than in the area covered by late Wisconsinan glaciation. He observed no deposits of till or glaciofluvial material that he considered to be of Illinoian-age. Crowl and Sevon (1980) also placed the late Wisconsinan terminus at this site and projected a long, narrow ice tongue 10 km down the Lycoming Creek Valley at an abnormally gentle ice surface gradient of less than 100 ft/mi to the site. The area was examined in more detail in 1993 (Reccelli, 1993a, 1993b) and it was determined that most of the site had had much of its weathering profile either truncated by erosion or the original sand and gravel was buried by colluvium from the valley sides. At and just southwest of the crest of the head of outwash rubified clasts were common at the surface and in an overgrown gravel pit beside Bobst Mountain Road hand excavation showed the sand and gravel to have a depth and intensity of weathering typical of a pre-Wisconsinan (late Illinoian or older) glaciation rather than a late Wisconsinan glaciation. The upper several feet were a roundstone diamict where stratification and many weatherable clasts had been destroyed by weathering, cryoturbation, and bioturbation. Beneath that was a 5 ft or so thick zone of intensely oxidized, high chroma 10 YR color, sand and gravel with some relict bedding, many rubified clasts, and some friable “ghost” clasts. South of the site a new roadcut along U.S. 15 (Fig. 26) showed an oxidation and weathering profile at least 20 ft thick, further indicating that the glacial deposits in the area were of pre-Wisconsinan age. The thickness of the head of outwash deposit at Stop 9 was confirmed by a seismic refraction profile tied to a well point at its southeast end (Fig. 26). Depth to bedrock was 205 ft at the northwest end of the profile and 235 ft at the southeast end of the profile and well point. The bedrock surface is at an elevation of ~600 ft, the same elevation of present day Lycoming Creek through the narrows at Powys. Turn around and then turn left onto Saint Michaels Road. Descend to the Lycoming Creek Valley, go U.S. 15, and turn left onto Rt. 110 to continue up Lycoming Creek Valley. At Trout

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Run, continue ahead where Rt. 14 merges with Rt. 110. Almost immediately turn left onto Trout Run Road, cross Lycoming Creek, and turn left along the creek onto Susque Road. Continue upstream across broad outwash terraces and turn right into gravel parking lot for Pennsylvania game lands (Lat. 41° 24′34.07″ N, Long. 77° 02′ 20.62″ W). Stop 10. Late Wisconsinan Terminal Head of Outwash and Terraces in the Lycoming Valley We have just driven 10 km up the narrow, deep Lycoming Creek Valley and are now standing less than 1 km south of where the late Wisconsinan (MIS 2) terminus is presently mapped (Braun, 2007) (Fig. 27). It has been in the Lycoming Valley where the placement of the late Wisconsinan limit by previous workers had varied by ~20 km north to south, the most of about any place in northeastern Pennsylvania. Lewis (1884) had placed the terminus across the valley ~9 km north of this site. Denny and Lyford (1963) had placed the terminus ~1.5 km south of this site and, as noted at Stop 9, Sevon (in Faill et al., 1977) and Crowl and Sevon (1980) had placed the terminus 10 km south of this site. The problem has been that since the late Wisconsinan glacier started retreating from the terminus here, erosion has removed most of the head of outwash deposits that once should have been on the floor of this narrow, deep valley. During glacial recession this valley acted as a major meltwater sluiceway as the glacier receded for 40 km to the northeast. During that recession it was the outlet for one proglacial lake that probably drained catastrophically through here and another proglacial lake (Glacial Lake Towanda) that spilled through here for the last 20 km of the overall 40 km ice recession from the region. The steepness of the valley sides also caused the deposits on those slopes to have been almost entirely eroded away. Also previous workers were doing reconnaissance work and mostly worked along the roads. The present map was produced by walking the mountain sides and almost continuous tracing of the late Wisconsinan terminus. In places the terminus location was determined solely by the presence or absence of erratic clasts on ledge tops on the steep mountain sides. The terminus is presently placed midway between where Grays Run enters Lycoming Creek and Hagerman Run enters Lycoming Creek, essentially at the location of the stream gauging station on the creek (Fig. 26). Immediately east of that site are the last exposures of what could be identified as late Wisconsinan till or ice-contact stratified drift. At the mouth of Glendenen Run (south edge of Fig. 26) a patch of outwash gravel with a top elevation 220 ft above present stream level shows that a more than 200 ft thick head of outwash occupied the Lycoming Creek Valley at the terminus. Heads of outwash of that thickness are common in major stream valleys all along the terminus in northeastern Pennsylvania (Braun, 1994a, 2004a). The head of outwash acted as a buttress that the glacier ramped up against with ice thickness of at least 200 ft at the terminus (ice surface elevation at least 900 ft). Just 1.8 mi northeast of the terminus,

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glacial erratics on the nose of Shriver Ridge were observed at an elevation of 1700 ft (just north of Fig. 26), giving an average ice surface gradient of more than 400 ft per mi for the snout of the glacier in the Lycoming Creek Valley. That is similar to the ice surface gradient determined for the terminus to the east in the Muncy Valley lowland (Braun, 2004b, 2004c). No late Wisconsinan ice occupied Glendenen Run or Hagerman Run (west side of Fig. 26) except at its head. Turn right onto Susque Road, continuing upstream, cross Lycoming Creek, and turn right onto Rt. 14. Turn right onto Field Station Road, again crossing Lycoming Creek. Ascend valley to divide and turn right onto Rose Valley Road. Bear left onto Lake Road, cross part of Rose Lake, continue along the east shore of the lake, and turn right into the parking lot for the boat launch area (Lat. 41° 23′ 10.99″ N, Long. 76° 58′ 51.43″ W).

Stop 11. Late Wisconsinan Terminal Moraine at Rose Lake At this site, in contrast to the previous narrow, deep valley site at Stop 10, the late Wisconsinan terminus runs across a rollinghill landscape in the breached Barbours anticline where the relief and hillslope angles are low to moderate. Here the deposits at the terminus are little eroded and retain constructional moraine topography (Fig. 28). The deposits are so thick as to cover the entire landscape and partly block the pre-glacial fluvial drainage system. Lewis (1884) traced what he considered to be the terminal moraine (southern limit) of glaciation across the region in 1881(only a single age of glaciation was recognized at that time). Lewis observed the moraine on the east side of Rose Lake and recognized that the Rose Lake area once drained eastward down Murray Run but was blocked by the moraine so a new channel

Figure 27. Surficial geology map (Braun, 2007) of the late Wisconsinan terminus in the Lycoming Creek Valley (Stop 10). It shows the discontinuous nature of late Wisconsinan deposits even at the terminus in this steeply sloping landscape. The line with ticks is Braun’s late Wisconsinan terminus with a steeply sloping ice profile near the terminus. The line with triangles is Crowl and Sevon’s (1980) terminus with a gently sloping ice profile. Qa—alluvium, Qat—alluvial terrace, Qaf—alluvial fan, Qwo—outwash terrace (w—Wisconsinan-age), Qwic—ice-contact stratified drift, Qwt—till, Qwtb—bouldery till, Qbc—boulder colluvium, Qbc/t—boulder colluvium over till, Qsc—stony colluvium, Qcl—colluvium with Illinoian-age lag clasts, Qit—Illinoian-age till, R—bedrock. Lines of ticks—sandstone ledges; isochore line—30 ft of thickness.

Susquehanna River Basin was cut westward through red shale. That channel is the site of the dam that impounds Rose Lake today. Glacial deposits are estimated to be in excess of 150 ft thick under the moraine in the now buried center of the pre-glacial Murray Run Valley (Fig. 28). It should be noted that the late Wisconsinan limit on the south and west sides of Rose Lake (Fig. 28) represents a short-lived, advanced position of the terminus. The long-term equilibrium position of the late Wisconsinan terminus was on the east side of Rose Lake where well-developed knob and kettle topography

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forms both till and kame moraine areas. The presence of a shortlived, advanced position of the terminus was noted previously by Braun (2004b, 2004c) to the east in the Muncy Valley lowland. As previously discussed at Stop 10, the late Wisconsinan terminus drawn on Crowl and Sevon’s (1980) map required the ice profile near the terminus to be far too gently sloped, in places essentially horizontal. That is the case south of Rose Lake, where Crowl and Sevon drew the glacial limit as ice tongues with little or no ice gradient to the south extending south into two hollows

Figure 28. Surficial geology map (Braun, 2006b) of the area around the late Wisconsinan terminal moraine at Rose Lake (Stop 11). The line with ticks is Braun’s late Wisconsinan terminus, and the line with triangles is Crow and Sevon’s (1980) terminus. Holocene–late Wisconsinan-age material: Qa—alluvium, Qc—colluvium, Qsc—stony colluvium. Qw denotes Wisconsinan-age: Qwt—till, Qwtm—till moraine, Qwkm—kame moraine (sand and gravel); heavy lines denote ridges in the moraine. Qi denotes Illinoian or older age material: Qil—colluvium with lag clasts, Qit—till. R—bedrock; lines of ticks—sandstone ledges; isochore lines at 30, 100, and 150 ft of thickness.

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and onto the adjacent quadrangle (Fig. 28). During field work for the current map (Braun, 2006b)(Fig. 28), till with abundant erratics was observed at the south edge of the Bodines quadrangle but no farther south. Placing the limit there produces a steeper and more reasonable ice gradient. Turn right onto Lake Road, left onto Calvert Road, and right onto Southard Road to descend into the Loyalsock Creek drainage. Turn right onto Wallis Run Road and continue down Wallis Run to Loyalsock Creek. Turn left onto Rt. 973 and immediately cross Loyalsock Creek and turn right onto Rt. 87, continuing downstream. We will now be crossing the broad late Wisconsinan outwash terraces of Loyalsock Creek. Pass under I-180 and immediately turn right onto ramp to I-180 east. We will now go around the plunging anticlinal nose of Bald Eagle Mountain to our left (west). At I-180 exit 5 bear right onto ramp to Rt.54 and turn left going under I-180. Park on right just beyond the exit intersection next to trucking warehouse parking lot fence (Lat. 41° 07′ 14.85″N, Long. 76° 47′ 57.67″W).

runs along 10–30 km northeast of the site (Braun, 1994a, 2004a) (Fig. 1). To the southeast in the Anthracite coal region, a border between the Illinoian or older margin and the farthest edge of glaciation, the early Pleistocene MIS 22+ margin, can be drawn on the basis of normal versus reversed magnetism of the deposits (Braun, 1994a, 2004a; Sasowsky, 1994). In this area there is yet no paleomagnetic information, but there is a chain of patches of much thicker till and kame deposits, the MIS 16? line on Figure 1. Also along that line there is a subtle change in the continuity of the pre-Illinoian glacial deposits with the patches being larger and more continuous to the northeast of the line and smaller and less continuous southwest of the line. Future mapping work should better define this MIS 16? line and possibly delineate yet another glacial margin farther to the southwest within the MIS 22+ line. Turn around and then turn left onto ramp to I-180 south. After crossing I-80, I-180 becomes Rt. 147. Take Rt. 147 south to Northumberland, turn right onto U.S. 11 and cross the West Branch Susquehanna River. Turn left onto combined U.S. 11–15 and retrace the previous day’s route; this time heading downstream along the Susquehanna River. Bear right onto ramp to

Stop 12. Mid-Pleistocene Kame Remnant Sitting on a Hill Top along I-180 near Turbotville, Pennsylvania At this site we are standing on a broad hill top in a gently rolling-hill landscape mostly underlain by shale within the breached nose of the White Deer portion of the Nittany anticlinorium (Inners, 1997). The hill is capped by 40 ft of sand and gravel that extends for a mile, north to south along the top the hill and I-180 (Fig. 29). The material is deeply leached and oxidized, high chroma 5YR color, stratified silty sands and sandy gravels (Inners, 1997). The material is intensely oxidized to a depth of at least 30 ft, and the upper 5 ft or so is a roundstone diamict produced by the cryoturbation and bioturbation of the sand and gravel. Much of the fine matrix in the diamict has been produced by destruction of the more weatherable local sandstone, siltstone, and shale clasts by the intense weathering and mixing of the material. Remnant clasts are primarily of resistant-to-weathering quartzose sandstone, quartz pebble conglomerate, and chert, with some of those clasts rubified throughout. The near surface diamict had often been misinterpreted as till by early geologists (Braun and Kaktins, 1986). Proceeding downward for 5–10 ft beneath the diamict, remnant bedding becomes more prominent downward and weathered clast “ghosts” become less common downward. Below 15–20 ft, the material is just oxidized and the clasts are dominated by readily weatherable locally derived sandstone, siltstone, and shale clasts. Weather and snow cover permitting, the material can be observed in cuts along side a trucking company parking lot immediately southeast of the I-180 interchange. This site is interpreted to be a frontal kame or head of outwash at the terminus of a pre-Illinoian, mid-Pleistocene glacier, most likely of MIS 16 age (Fig. 1). A chain of similar kame deposits marking the Illinoian or older (MIS 6 or 12) glacial limit

Figure 29. Surficial geology map (Inners, 1997) of the area around the pre-Illinoian kame at Stop 12. Abbreviations: af—artificial fill, Qac—alluvium-colluvium, Qc—colluvium, Qpt—pre-Illinoian till, Qpsd—pre-Illinoian stratified drift. Numbers—thickness of surficial deposits in meters in wells; double circles—erratics.

Susquehanna River Basin U.S. 22–322 south to Harrisburg. Cross the Susquehanna River again, go through the water gaps, and bear right onto I-81 west to cross the river one more time. Immediately after crossing the river, bear right onto ramp to U.S. 11–15 south to Enola and Camp Hill. In a couple of miles one will be driving along the west bank of the Susquehanna River. Where U.S. 11–15 turns right, continue straight ahead along the river on N. Front St. Where the Market St. bridge enters from the left, continue ahead on Market St. and curve right to ascend out of the Susquehanna Valley. Turn left onto South 3rd St., turn right onto Hummel Ave., turn left onto South 10th St., and continue straight ahead onto ramp for I-83 south. Take I-83 south to I-695 (Baltimore beltway), bearing right onto combined I-83 and I-695. Bear right onto ramp for I-83 south (Jones Falls Expressway) and take I-83 to where its ends, bearing right onto E. Fayette St. Continue ahead onto W. Fayette St. and the Sheraton Baltimore City Center Hotel at 101 W. Fayette St. (Lat. 39° 17′ 24.73″ N, Long. 76° 33′ 01.55″ W). REFERENCES CITED Andreus, E., 1993, Surficial geology and stratigraphy of the Antes Fort area (senior research project): Bloomsburg, Bloomsburg University Geography and Earth Science Department, 28 p. Ashley, G.H., 1935, Studies in Appalachian mountain structure: Geological Society of America Bulletin, v. 46, p. 1395–1436. Bascom, F., 1921, Cycles of erosion in the Piedmont province of Pennsylvania: The Journal of Geology, v. 29, p. 540–559, doi: 10.1086/622809. Beaumont, C., Kooi, H., and Willett, S., 2000, Coupled tectonic-surface process models with applications to rifted margins and collisional orogens, in Summerfield, M.A., ed., Geomorphology and Global Tectonics: Chirchester, John Wiley and Sons, p. 28–55. Braun, D.D., 1988, Glacial geology of the anthracite and North Branch Susquehanna lowland regions, in Inners, J., ed., Bedrock and Glacial Geology of the North Branch Susquehanna Lowland and the Eastern Middle Anthracite Field, Northeastern Pennsylvania: Harrisburg, 53rd Annual Field Conference of Pennsylvania Geologists Guidebook, p. 3–25. Braun, D.D., 1989a, Glacial and periglacial erosion of the Appalachians, in Gardner, T.W. and Sevon, W.D., eds., Appalachian geomorphology: Geomorphology, v. 2, no. 1–3, p. 233–258. Braun, D.D., 1989b, A revised Pleistocene glaciation sequence in eastern Pennsylvania: Support for limited early Wisconsinan ice and a single late Illinoian advance beyond the late Wisconsin border: 28th International Geological Congress, v. 1, p. 196–197. Braun, D.D., 1990, Negative evidence of late Wisconsinan catastrophic flooding down the Susquehanna River: Geological Society of America Abstracts with Programs, v. 22, p. 6. Braun, D.D., 1994a, Late Wisconsinan to pre-Illinoian (G?) glacial events in eastern Pennsylvania, in Braun, D.D., ed., Late Wisconsinan to PreIllinoian (G?) glacial and periglacial events in eastern Pennsylvania: 57th Field Conference of the Friends of the Pleistocene (Northeastern Section) Guidebook, U.S. Geological Survey Open-File Report 94-434, p. 1–21. Braun, D.D., 1994b, The ubiquitous boulder colluvium of eastern Pennsylvania, a relict periglacial mobilization of the entire land surface, in Braun, D.D., ed., Late Wisconsinan to Pre-Illinoian (G?) glacial and periglacial events in eastern Pennsylvania: 57th Field Conference of the Friends of the Pleistocene (Northeastern Section) Guidebook, U.S. Geological Survey Open-File Report 94-434. p. 34–37. Braun, D.D., 1995, Geomorphologic analysis and soil descriptions, in Bohlin, J., Geomorphological Assessment of the Proposed Water Pipeline to Antes Fort: Appalachian Archaeological Research, p. 2-1–2-13. Braun, D.D., 1996, Pseudo-morainic topography of the Allentown area of eastern Pennsylvania: Harrisburg, 15th Annual Field Trip of the Harrisburg Area Geological Society Guidebook, 28 p. Braun, D.D., 1999, Pleistocene geology of the Allentown area, in Sevon, W.D., and Fleeger, G.M., eds., Economic and Environmental Geology and

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Topography in the Allentown-Bethlehem area: Harrisburg, 64th Annual Field Conference of Pennsylvania Geologists Guidebook, p. 31–40. Braun, D.D., 2004a, The glaciation of Pennsylvania, USA, in Ehlers, J., and Gibbard, P.L., eds., Quaternary Glaciations Extent and Chronology, Part II: North America: Developments in Quaternary Science, Elsevier, p. 237–242. Braun, D.D., 2004b, Surficial geology of the Picture Rocks 7.5-Minute Quadrangle, Lycoming and Sullivan Counties, Pennsylvania: Pennsylvania Geological Survey, 4th series, Open-File Report OFSM 04-01.1., 21 p. (PDF). Braun, D.D., 2004c, Surficial geology of the Sonestown 7.5-minute quadrangle, Lycoming and Sullivan Counties, Pennsylvania: Pennsylvania Geological Survey, 4th series, Open-File Report OFSM 04-02.1, 21 p., Portable Document Format (PDF). Braun, D.D., 2006a, Deglaciation of the Appalachian Plateau, northeastern Pennsylvania—Till shadows, till knobs forming “beaded valleys”: Revisiting systematic stagnation-zone retreat, in Fleisher, P.J., Knuepfer, L.K., and Butler, D.R., eds., Ice Sheet Geomorphology—Past and Present Processes and Landforms: Geomorphology, v. 75, p. 248–265. Braun, D.D., 2006b, Surficial geology of the Bodines 7.5-minute quadrangle, Lycoming County, Pennsylvania: Pennsylvania Geological Survey, 4th series, Open-File Report OFSM 06-09.1, 23 p., Portable Document Format (PDF). Braun, D.D., 2007, Surficial geology of the Trout Run 7.5-minute quadrangle, Lycoming County, Pennsylvania: Pennsylvania Geological Survey, 4th series, Open-File Report OFSM 07-11.0, 26 p., Portable Document Format (PDF). Braun, D.D., 2008, The Pleistocene record in the middle and lower Susquehanna River Basin and the longer term evolution of the Susquehanna Basin landscape: University Park, 20th Biennial Meeting of the American Quaternary Association, Fieldtrip Guidebook, 58 p. Braun, D.D., and Kaktins, T., 1986, Diamicts from the weathering of preWisconsinan glaciofluvial and fluvial deposits in central Pennsylvania: Geological Society of America Abstracts with Programs, v. 17, no. 1, p. 6. Braun, D.D., Miller, D.S., Miller, J.C., and Inners, J.D., 1984, Abandoned valley of the North Branch Susquehanna River at Mifflinville, Pennsylvania: Evidence for a pre-late Wisconsinan ice margin: Geological Society of America Abstracts with Programs, v. 15, p. 5. Braun, D.D., Gardner, T.W., and Pazzaglia, F.J., 1993, Field trip overview and summary, in Gardner, T.W., Braun, D.D., Pazzaglia, F.J., and Sevon, W.D., Late Cenozoic landscape evolution of the Susquehanna River basin: Third International Geomorphology Conference, Post-Conference Field Trip Guidebook, 288 p. Braun, D.D., Pazzaglia, F.J., and Potter, N., Jr., 2003, Margin of Laurentide ice to the Atlantic Coastal Plain: Miocene–Pleistocene landscape evolution in the central Appalachians, in Easterbrook, D.J., ed., Quaternary geology of the United States: Reno, Nevada, International Union for Quaternary Research 2003 Field Guide volume, Desert Research Institute, p. 219–244. Bucek, M.F., 1975, Pleistocene geology and history of the West Branch of Susquehanna River Valley near Williamsport, Pennsylvania (Ph.D. thesis): College Park, Pennsylvania State University, 197 p. Campbell, M.R., 1903, Geographic development of northern Pennsylvania and southern New York: Geological Society of America Bulletin, v. 14, p. 277–296. Campbell, M.R., 1929, Late geologic deformation of the Appalachian Piedmont as determined by river gravels: Proceedings of the National Academy of Sciences of the United States of America 15, p. 156–161. Campbell, M.R., 1933, Chambersburg (Harrisburg) peneplain in the Piedmont of Maryland and Pennsylvania: Geological Society of America Bulletin, v.44, p. 553–573. Ciolkosz, E.J., 1999, Pennsylvania State University Soil Characterization Laboratory Database: University Park, Crop and Soil Sciences Department, Pennsylvania State University. Ciolkosz, E.J., Cronce, R.C., Sevon, W.D., and Waltman, W.J., 1995, Genesis of Pennsylvania’s limestone soils: Pennsylvania State University Agronomy Series 135, 28 p. Clark, G.M., and Ciolkosz, E.J., 1988, Periglacial geomorphology of the Appalachian highlands and interior highlands south of the glacial border—A review: Geomorphology, v. 1, p. 191–220, doi: 10.1016/0169 -555X(88)90014-1. Clark, G.M., Behling, R.E., Braun, D.D., Ciolkosz, E.J., Kite, J.S., and Marsh, B., 1992, Central Appalachian periglacial geomorphology: Field

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excursion guidebook for 27th International Geographical Congress, Pennsylvania State University Agronomy Series 120, 248 p. Crowl, G.H. and Sevon, W.D., 1980, Glacial border deposits of Late Wisconsinan age in northeastern Pennsylvania: Pennsylvania Geological Survey, 4th series, General Geology Report 71, 68 p., 1 plate. Davis, W.M., 1889, The rivers and valleys of Pennsylvania: National Geographic, v. 1, p. 183–253. Davis, W.M., 1899, The geographical cycle: The Geographical Journal, v. 14, p. 481–504, doi: 10.2307/1774538. Davis, W.M., and Wood, J.W., Jr., 1890, The geographic development of northern New Jersey: Boston Society of Natural History Proceedings, v. 24, p. 365–423. Denny, C.S., and Lyford, W.H., 1963, Surficial geology and soils of the ElmiraWilliamsport region, New York and Pennsylvania: U.S. Geological Survey Professional Paper 379, 60 p. Engel, S.A., Gardner, T.W., and Ciolkosz, E.J., 1996, Quaternary soil chronosequences on the lower terraces of the Susquehanna River, Pennsylvania: Geomorphology, v. 17, p. 273–294, doi: 10.1016/0169-555X(96)00005-0. Faill, R.T., Wells, R.B., and Sevon, W.D., 1977, Geology and mineral resources of the Salladasburg and Cogan Station quadrangles, Lycoming County, Pennsylvania: Pennsylvania Geologic Survey, 4th series, Atlas 133cd, 44 p., 2 plates. Fenneman, N.M., 1938, Physiography of Eastern United States: New York, McGraw-Hill Book Company, 689 p. Gardner, T.W., Braun, D.D., Pazzaglia, F.J., and Sevon, W.D., 1993, Late Cenozoic landscape evolution of the Susquehanna River basin: Third International Geomorphology Conference, Post-Conference Field Trip Guidebook, 288 p. Gardner, T.W., Sasowsky, I.D., and Schmidt, V.A., 1994, Reversed polarity glacial sediments and revised glacial chronology, West Branch Susquehanna River Valley, central Pennsylvania: Quaternary Research, v. 42, p. 131– 135, doi: 10.1006/qres.1994.1062. Garner, H.F., 1974, The Origin of Landscapes: New York, Oxford University Press, 734 p. Hack, J.T., 1960, Interpretation of erosional topography in humid temperate regions: American Journal of Science, v. 258-A, p. 80–97. Hack, J.T., 1965, Geomorphology of the Shenandoah Valley, Virginia and West Virginia, and origin of the residual ore deposits: U.S. Geological Survey Professional Paper 484, 84 p. Hack, J.T., 1973, Drainage adjustment in the Appalachians, in Morisawa, M., ed., Fluvial Geomorphology: Binghamton, 4th Annual Geomorphology Symposia Proceedings, State University of New York, Publications in Geomorphology, p. 51–74. Hack, J.T., 1975, Dynamic equilibrium and landscape evolution, in Melhorn, W.N., and Flemal, R.C., eds., Theories of Landform Development: Binghamton, Sixth Annual Geomorphology Symposia Proceedings, State University of New York, Publications in Geomorphology, p. 87–102. Hoskins, D.M., 1987, The Susquehanna River water gaps near Harrisburg, Pennsylvania, in Roy, D.C., ed., Northeastern Section: Geological Society of America, Centennial Field Guide, v. 5, p. 47–50. Inners, J.D., 1997, Geology and mineral resources of the Allenwood and Milton quadrangles, Union and Northumberland Counties, Pennsylvania: Pennsylvania Geological Survey, 4th series, Atlas 144cd, 135 p., 3 plates. Johnson, D.W., 1931, Stream sculpture on the Atlantic slope: A study in the evolution of Appalachian rivers: New York, Columbia University Press, 142 p. Jonas, A.I. and Stose, G.W., 1930, Geology and mineral resources of the Lancaster quadrangle, Pennsylvania: Pennsylvania Geologic Survey, 4th series, Atlas 168, 106 p. Judson, S., 1975, Evolution of Appalachian topography, in Melhorn, W.N., and Flemal, R.C., eds., Theories of Landform Development: Binghamton, Sixth Annual Geomorphology Symposia Proceedings, State University of New York, Publications in Geomorphology, p. 1–28. Leverett, F., 1934, Glacial deposits outside the Wisconsin terminal moraine in Pennsylvania: Pennsylvania Geological Survey, 4th series, General Geology Report 7, 123 p. Levine, J.R., 1986, Deep burial of coal-bearing strata, Anthracite region, Pennsylvania: Sedimentation or tectonics?: Geology, v. 14, p. 577–580, doi: 10.1130/0091-7613(1986)14<577:DBOCSA>2.0.CO;2. Lewis, H.C., 1884, Report on the terminal moraine in Pennsylvania and western New York: Pennsylvania Geological Survey, 2nd series, Report Z, 299 p. Lobeck, A.K., 1932, Atlas of American geology: New York, Columbia University, The Geographical Press, sheets 31 and 34.

MacLachlan, D.B., 1985, Pennsylvania anthracite as foreland effect of Alleghanian thrusting: Geological Society of America Abstracts with Programs, v. 17, no. 1, p. 53. Marchand, D.E., 1978, Quaternary deposits and Quaternary history, in Marchand, D.E., Ciolkosz, E.J., Bucek, M.F., and Crowl, G.H., eds., Quaternary Deposits and Soils of the Central Susquehanna Valley of Pennsylvania: University Park, 41st Friends of the Pleistocene Field Conference, Pennsylvania State University Agronomy Series 52, p. 1–19. Marchand, D.E., and Crowl, G.H., 1991, Surficial geologic map of parts of Union and Snyder Counties, Pennsylvania: U.S. Geological Survey Miscellaneous Investigations Map I-2051. Marsh, B., 1999, Paleoperiglacial landscapes of central Pennsylvania: Lewisburg, 62nd Northeast Friends of the Pleistocene Field Trip Guidebook, 69 p. Mathews, E.B., 1917, Submerged “deeps” in the Susquehanna River: Geological Society of America Bulletin, v. 28, p. 335–346. Meyerhoff, H.A., 1972, Postorogenic development of the Appalachians: Geological Society of America Bulletin, v. 83, p. 1709–1728, doi: 10.1130/0016-7606(1972)83[1709:PDOTA]2.0.CO;2. Naeser, N.D., Naeser, C.W., Morgan, B.A., III, Schultz, A.P., and Southworth, C.S., 1999, Cooling history of the Blue Ridge Province, Virginia, North Carolina, and Tennessee, from apatite and zircon fission-track analysis: Geological Society of America Abstracts with Programs, v. 31, p. 32. Naeser, C.W., Naeser, N.D., Kunk, M.J., Morgan, B.A., Schultz, A.P., Southworth, C.S., and Weems, R.E., 2001, Paleozoic through Cenozoic uplift, erosion, stream capture, and deposition history in the Valley and Ridge, Blue Ridge, Piedmont, and Coastal Plain provinces of Tennessee, North Carolina, Virginia, Maryland, and District of Columbia: Geological Society of America Abstracts with Programs, v. 33, p. 312. Oberlander, T.M., 1985, Origin of drainage traverse to structures in orogens, in Morisawa, M., and Hack, J.T., eds., Tectonic Geomorphology: New York, Allen and Unwin, p. 156–182. Page, L.V., and Shaw, L.C., 1973, Floods of June 1972 in the Harrisburg area, Pennsylvania: U.S. Geological Survey Hydrologic Investigations Atlas HA-530. Pazzaglia, F.J., and Brandon, M.T., 1996, Macrogeomorphic evolution of the post-Triassic Appalachian Mountains determined by deconvolution of the offshore basin sedimentary record: Basin Research, v. 8, p. 255–278, doi: 10.1046/j.1365-2117.1996.00274.x. Pazzaglia, F.J., and Gardner, T.W., 1993, Fluvial terraces of the lower Susquehanna River: Geomorphology, v. 8, p. 83–113, doi: 10.1016/ 0169-555X(93)90031-V. Pazzaglia, F.J., and Gardner, T.W., 1994, Late Cenozoic flexural deformation of the middle U.S. Atlantic passive margin: Journal of Geophysical Research, v. 99, no. B6, p. 12,143–12,157, doi: 10.1029/93JB03130. Pazzaglia, F.J., and Gardner, T.W., 2000, Late Cenozoic large-scale landscape evolution of the U.S. Atlantic passive margin, in Summerfield, M., ed., Geomorphology and Global Tectonics: New York, John Wiley, p. 283–302. Pazzaglia, F.J., Gardner, T.W., and Merritts, D., 1998, Bedrock fluvial incision and longitudinal profile development over geologic time scales determined by fluvial terraces, in Wohl, E., and Tinkler, K., eds., Bedrock Channels: American Geophysical Union, Geophysical Monograph Series, v. 107, p. 207–235. Pazzaglia, F.J., Braun, D.D., Pavich, M., Bierman, P., Potter, N., Jr., Merritts, D., Walter, R., and Germanoski, D., 2006, Rivers, glaciers, landscape evolution, and active tectonics of the central Appalachians, Pennsylvania and Maryland, in Pazzaglia, F.J., ed., Excursions in Geology and History: Field Trips in the Middle Atlantic States: Geological Society of America Field Guide 8, p. 169–197. Peltier, L.C., 1949, Pleistocene terraces of the Susquehanna River, Pennsylvania: Pennsylvania Geological Survey, 4th series, General Geology Report 23, 151 p. Poag, C.W., and Sevon, W.D., 1989, A record of Appalachian denudation in postrift Mesozoic and Cenozoic sedimentary deposits of the U.S. Middle Atlantic continental margin, in Gardner, T.W., and Sevon, W.D., eds., Appalachian Geomorphology: Geomorphology, v. 2, p. 119–157. Ramage, J.M., Gardner, T.W., and Sasowsky, I.D., 1998, Early Pleistocene glacial Lake Lesley, West Branch Susquehanna River Valley, central Pennsylvania: Geomorphology, v. 22, p. 19–37, doi: 10.1016/S0169 -555X(97)00053-6. Reccelli, H., 1993a, Age delineation of glacial deposits that diverted Lycoming Creek to the present course at Powys, Lycoming County, Pennsylvania: Pennsylvania Academy Science, v. 67, p. 194.

Susquehanna River Basin Reccelli, H., 1993b, Age delineation of glacial deposits that diverted Lycoming Creek to the present course at Powys, Lycoming County, Pennsylvania (Senior Research Project): Bloomsburg, Pennsylvania, Bloomsburg University Geography and Geosciences, 36 p. Reusser, L.J., Bierman, P., Pavich, M., Larsen, J., and Finkel, R., 2006, An episode of rapid bedrock channel incision during the last glacial cycle, measured with 10Be: American Journal of Science, v. 306, p. 69–102, doi: 10.2475/ajs.306.2.69. Ridge, J.C., Braun, D.D., and Evenson, E.B., 1990, Does the Altonian drift exist in Pennsylvania and New Jersey?: Quaternary Research, v. 33, p. 253– 258, doi: 10.1016/0033-5894(90)90023-E. Sasowsky, I.D., 1994, Paleomagnetism of Glacial Sediments from three locations in eastern Pennsylvania, in Braun, D.D., ed., Late Wisconsinan to pre-Illinoian (G?) glacial and periglacial events in eastern Pennsylvania: 57th Field Conference of the Friends of the Pleistocene, Northeastern Section Guidebook, U.S. Geological Survey Open-File Report 94-434, p. 21–23. Sevon, W.D., 1985, Pennsylvania’s polygenic landscape: Harrisburg, 4th Annual Field Trip, Harrisburg Area Geological Society Guidebook, 55 p. Sevon, W.D., 1986, Susquehanna River water gaps: Many years of speculation: Pacific Geology, v. 17, no. 3, p. 4–7. Sevon, W.D., 1989a, Erosion in the Juniata River drainage basin, Pennsylvania, in Gardner, T.W., and Sevon, W.D., eds., Appalachian Geomorphology: Geomorphology, v. 2, p. 303–318. Sevon, W.D., 1989b, The rivers and valleys of Pennsylvania, then and now, in Sevon, W.D., ed., The rivers and valleys of Pennsylvania, then and now: Harrisburg, Harrisburg Area Geological Society and the 20th Annual Geomorphology Symposium Guidebook, p. 1–16. Sevon, W.D., 1999, Cenozoic History, in Shultz, C.H., ed., The Geology of Pennsylvania: Pennsylvania Geologic Survey, 4th series, Special Publication 1, p. 450–455. Shaw, J., 1989, Drumlins, subglacial meltwater floods, and ocean responses: Geology, v. 17, p. 853–856, doi: 10.1130/0091-7613(1989)017<0853: DSMFAO>2.3.CO;2. Slingerland, R., and Furlong, K.P., 1989, Geodynamic and geomorphic evolution of the Permo-Triassic Appalachian Mountains, in Gardner, T.W., and Sevon, W.D., eds., Appalachian Geomorphology: Geomorphology, v. 2, p. 23–37. Smoot, J.P., 1999, Early Mesozoic sedimentary rocks, in Shultz, C.H., ed., The Geology of Pennsylvania: Pennsylvania Geological Survey, 4th series, Special Publication 1, p. 180–201.

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Stanford, S.D., 1997, Pliocene-Quaternary geology of northern New Jersey: An overview, in Stanford, S.D., and Witte, R.W., eds., Pliocene-Quaternary Geology of Northern New Jersey: New Brunswick, 60th Annual Reunion of the Northeast Friends of the Pleistocene Guidebook, p. 1.1–1.26. Stanford, S.D., Ashley, G.M., and Brenner, G.J., 2001, Late Cenozoic fluvial stratigraphy of the New Jersey Piedmont: A record of glacioeustasy, planation, and incision on a low-relief passive margin: Geology, v. 109, p. 265–276, doi: 10.1086/319242. Stose, G.W., 1928, High gravels of the Susquehanna River above Columbia, Pennsylvania: Geological Society of America Bulletin, v.39, p. 1073–1086. Stose, G.W. and Jonas, A.I., 1939Geology and mineral resources of York County: Pennsylvania Geologic Survey, 4th series, County Report C-67, 199 p. Thiesen, J.P., 1983, Is there a fault in our gap?: Pacific Geology, v. 14, no. 3, p. 5–11. Thompson, G.H., 1985, The preglacial great falls of the Susquehanna River: Geological Society of America Abstracts with Programs, v. 19, no. 7, p. 734. Thompson, G.H., 1990, Geomorphology of the lower Susquehanna River, in Scharnberger, C.K., ed., Carbonates, Schists, and Geomorphology in the Vicinity of the Lower Reaches of the Susquehanna River: Harrisburg, 55th Field Conference of Pennsylvania Geologists Guidebook, p. 86–106. Thompson, H.D., 1949, Drainage evolution in the Appalachians of Pennsylvania: Annals of the New York Academy of Sciences, v. 52, p. 31–63. White, W.B., 2000, Dissolution of limestone from field observation, in Kimchouk, A.B., Ford, D.C., Palmer, A.N., and Dreybrodt, W., eds., Speleogenesis, Evolution of Karst Aquifers: National Speleological Society, January 2000 edition, p. 149–155. Williams, E.H., Jr., 1895, Notes on the southern ice limit in eastern Pennsylvania: American Journal of Science, v. 49, p. 174–185. Williams, E.H., Jr., 1917, Pennsylvania glaciation first phase: Woodstock, Vermont, privately published, 101 p. Williams, E.H., Jr., 1920, The deep Kansan poundings in Pennsylvania and the deposits therein: American Philosophical Society Proceedings, v. 59, p. 49–84. Zhang, E., and Davis, A., 1993, Coalification patterns of the Pennsylvania coal measures in the Appalachian foreland basin, western and south-central Pennsylvania: Geological Society of America Bulletin, v. 105, p. 162– 174, doi: 10.1130/0016-7606(1993)105<0162:CPOTPC>2.3.CO;2.

MANUSCRIPT ACCEPTED BY THE SOCIETY 2 DECEMBER 2009

Printed in the USA

The Geological Society of America Field Guide 16 2010

Stratigraphy, structure, and tectonics: An east-to-west transect of the Blue Ridge and Valley and Ridge provinces of northern Virginia and West Virginia Lynn S. Fichter Steven J. Whitmeyer Department of Geology and Environmental Science, James Madison University, 800 S. Main Street, Harrisonburg, Virginia 22807, USA Christopher M. Bailey Department of Geology, College of William & Mary, Williamsburg, Virginia, USA William Burton U.S. Geological Survey, Reston, Virginia 22092, USA

ABSTRACT This field guide covers a two-day east-to-west transect of the Blue Ridge and Valley and Ridge provinces of northwestern Virginia and eastern West Virginia, in the context of an integrated approach to teaching stratigraphy, structural analysis, and regional tectonics. Holistic, systems-based approaches to these topics incorporate both deductive (stratigraphic, structural, and tectonic theoretical models) and inductive (field observations and data collection) perspectives. Discussions of these pedagogic approaches are integral to this field trip. Day 1 of the field trip focuses on Mesoproterozoic granitoid basement (associated with the Grenville orogeny) and overlying Neoproterozoic to Early Cambrian cover rocks (Iapetan rifting) of the greater Blue Ridge province. These units collectively form a basement-cored anticlinorium that was thrust over Paleozoic strata of the Valley and Ridge province during Alleghanian contractional tectonics. Day 2 traverses a foreland thrust belt that consists of Cambrian to Ordovician carbonates (Iapetan divergent continental margin), Middle to Upper Ordovician immature clastics (associated with the Taconic orogeny), Silurian to Lower Devonian quartz arenites and carbonates (inter-orogenic tectonic calm), and Upper Devonian to Lower Mississippian clastic rocks (associated with the Acadian orogeny). Alleghanian structural features include the Little North Mountain thrust, Cacapon Mountain anticlinorium, Broad Top synclinorium, and Wills Mountain anticlinorium. Within the road log of this field guide we include both planned and optional stops, so that readers can explore the pedagogic concepts discussed herein in more detail, if desired.

Fichter, L.S., Whitmeyer, S.J., Bailey, C.M., and Burton, W., 2010, Stratigraphy, structure, and tectonics: An east-to-west transect of the Blue Ridge and Valley and Ridge provinces of northern Virginia and West Virginia, in Fleeger, G.M., and Whitmeyer, S.J., eds., The Mid-Atlantic Shore to the Appalachian Highlands: Field Trip Guidebook for the 2010 Joint Meeting of the Northeastern and Southeastern GSA Sections: Geological Society of America Field Guide 16, p. 103–125, doi: 10.1130/2010.0016(05). For permission to copy, contact [email protected]. ©2010 The Geological Society of America. All rights reserved.

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INTRODUCTION This field guide focuses on an east-to-west, cross-orogen transect of the Blue Ridge and Valley and Ridge provinces of northwestern Virginia and eastern West Virginia, generally following Route 33 from Ruckersville, Virginia, to the Allegheny Front near Judy Gap, West Virginia. Over the past few decades several field trip guides have documented the stratigraphy and depositional environments of this region (e.g., Diecchio, 1986; Fichter, 1986; Fichter and Diecchio, 1993), as well as the structural geometry and tectonic history of this part of the Appalachian orogen (e.g., Perry and DeWitt, 1977; Spencer et al., 1989; Bailey et al., 2006). It is not the purpose of this field guide to reiterate all of this prior work, except to briefly summarize pertinent models and interpretations in the context of field locations that are highlighted. Instead, this field guide focuses on the pedagogy of an integrated approach to teaching stratigraphy and structural analysis within the overarching framework of tectonics. In the text that follows we present a holistic, systems approach to teaching these topics from both inductive and deductive perspectives that are rooted in field observations and data collection. In 2007, the James Madison University Geology and Environmental Science Department initiated a new course for earth science majors, entitled “Stratigraphy, Structure, Tectonics,” or “SST.” The course builds on a central theme of the earth science degree, which is to simultaneously think about, evaluate, and integrate the relative importance of many variables that influenced the geology and natural history of a region. In general, we want students to learn to evaluate how different systems influence each other, rather than think about each system individually. Thus, components of traditional stratigraphic and structural analyses techniques are incorporated within the broad umbrella of tectonics, in a course co-taught by specialists in these respective fields. A major component of the SST course is a semester-long project based on several field trips that collectively transect the Blue Ridge and Valley and Ridge geologic provinces of northern Virginia and West Virginia. During the field trips, structural and stratigraphic data are collected at each outcrop, with students and instructors evaluating the data to assemble a regional tectonic history. Stratigraphic data include detailed facies analysis and environmental interpretations; structural data include orientation measurements and geometric relationships. Students incorporate field observations, data, and interpretations into a final report that consists of two principal parts: (1) a cross section from the Blue Ridge to the Allegheny Front that incorporates multiple scales of stratigraphic and structural field evidence, and (2) a comprehensive tectonic and geologic history of the region that uses stratigraphic and structural field observations and is informed by theoretical models. For this report it is not enough to simply recognize and identify specific outcrop features. Observations and conclusions must be both inductively and deductively valid, and to the degree possible, internally consistent and rooted in the overarching mechanisms and driving forces of tectonics.

The Geological Society of America (GSA) field trip described in this guidebook is based on the SST semester-long project described above. We hope that this field guide provides a provocative way of thinking about student field trips and field education in general, as well as serving as a useful resource for educators leading their own field trips through this region. GEOLOGIC SETTING Blue Ridge The first day of this field trip begins south of Madison, Virginia, and proceeds west to Route 33, crosses the Blue Ridge at Swift Run Gap, and ends in Harrisonburg, Virginia. In northcentral Virginia, the Blue Ridge province forms a basementcored anticlinorium that was thrust over Paleozoic strata of the Valley and Ridge province during late Paleozoic (Alleghanian) contractional tectonics. Mesoproterozoic granitoid basement and a cover sequence of Neoproterozoic to Early Cambrian rocks underlie the Blue Ridge. The basement complex formed at middle to lower crustal depths during and after the assembly of the Rodinian supercontinent, while the cover sequence records rifting, Rodinian breakup, and the subsequent transgression of the Iapetus Ocean at the onset of the Phanerozoic. The basement suite formed between 1.2 and 1.0 Ga and includes three temporally distinct groups of granitoids (Fig. 1). The oldest group crystallized between 1190 and 1150 Ma, the middle group between 1120 and 1110 Ma, and the youngest group between 1090 and 1020 Ma (Aleinikoff et al., 2000; Tollo et al., 2004; Southworth et al., 2009). The older group is gneissic and characterized by a high-grade compositional layering. Both the older and younger groups are chemically diverse and range from low-silica charnockites to alkali feldspar leucogranitoids. The volumetrically less significant middle group is restricted to orthopyroxene-bearing granites. The felsic 730–700 Ma Robertson River plutonic and volcanic complex cuts Mesoproterozoic basement and a thick sequence of Neoproterozoic metasedimentary rocks (Lynchburg, Fauquier, and Mechum River units) overlies older rocks (Tollo and Aleinikoff, 1996; Bailey et al., 2007). This sequence of plutonic, volcanic, and sedimentary rocks formed during an episode of rifting that created significant accommodation space but did not generate ocean crust. A later Neoproterozoic cover sequence of metasedimentary and metavolcanic rocks unconformably overlies the basement complex in the western Blue Ridge and includes the clastic Swift Run Formation and predominantly basaltic greenstones of the Catoctin Formation (Fig. 1). The Swift Run Formation is heterogeneous with a highly variable thickness. The Catoctin Formation forms a thick unit extruded over a large area (>4000 km2) and was generated from mantle-derived tholeiitic magmas (Badger and Sinha, 2004). Metadiabase dikes of similar composition to Catoctin metabasalts intrude the basement complex (and older Neoproterozoic rocks) and are likely feeder dikes for the

Blue Ridge and Valley and Ridge Provinces of northern Virginia and West Virginia overlying Catoctin lava flows. Badger and Sinha (1988) report a Rb-Sr isochron age of 570 ± 36 Ma, and zircons from metarhyolite tuffs and dikes in the Catoctin Formation yield U-Pb ages between 570 and 550 Ma (Aleinikoff et al., 1995). The Early Cambrian Chilhowee Group overlies the Catoctin Formation and in north-central Virginia includes the Weverton, Harpers, and Antietam formations (Figs. 1 and 2). The contact

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between the underlying Catoctin metabasalts and the overlying Chilhowee Group has traditionally been interpreted as an unconformity, but its significance remains uncertain (King, 1950; Gathright, 1976; Southworth et al., 2009). The Weverton Formation includes quartz metasandstone, metaconglomerate, laminated metasiltstone, and quartzose phyllite. The Harpers Formation is dominated by phyllite and thinly bedded metasandstone with

Figure 1. Generalized stratigraphic section of rock formations in the western Blue Ridge anticlinorium, Virginia. Based in part on figure 6 from Gathright (1976). Zp—Neoproterozoic garnet-graphite paragneiss.

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Figure 2. Generalized stratigraphic column for the central and northern Shenandoah Valley of Virginia and eastern West Virginia, encompassing the lithologic units visited on Day 2 of the field trip. After Fichter and Diecchio (1993).

Blue Ridge and Valley and Ridge Provinces of northern Virginia and West Virginia minor well-cemented quartz arenite and ferruginous metasandstone. The Antietam Formation consists of well-cemented quartz arenite with abundant Skolithos. Collectively, the Chilhowee Group records shallow-marine sequences along the Laurentian margin during the Cambrian (Simpson and Eriksson, 1989). A thick package of Cambrian to Ordovician carbonate and shale stratigraphically overlies the Chilhowee Group and crops out in the Shenandoah Valley west of the Blue Ridge (Figs. 1 and 2). Unmetamorphosed diabase dikes cut Mesoproterozoic basement and Neoproterozoic to early Cambrian cover rocks. The olivine normative, low-titanium diabase is similar to mafic dikes and sills exposed in the Mesozoic Culpeper basin east of Shenandoah National Park. A diabase dike along the Potomac River, ~100 km north of the Park, yielded an Ar-Ar age of 200 Ma (Kunk et al., 1992). The Blue Ridge anticlinorium is an imbricated stack of basement thrust sheets emplaced over Valley and Ridge rocks along a family of related, low-angle thrusts. The frontal Blue Ridge fault system truncates early fold structures and penetrative fabrics in both the hanging-wall and footwall rocks; the deformation style is typically brittle, and breccias are well developed at a number of localities. In the core of the anticlinorium mylonitic rocks occur in km-thick anastomosing zones. Asymmetric structures in these high-strain zones consistently record a top-tothe-northwest sense of shear (i.e., hanging wall up movement). Displacement across Blue Ridge mylonite zones accommodated crustal contraction enabling the relatively stiff basement complex to shorten while cover rocks were folded. Folds in western Blue Ridge cover rocks are typically asymmetric northwest-verging structures. Axial planar foliation (cleavage) is well developed in fine-grained rocks and dips gently to moderately southeast. A distinctive coarse foliation or compositional banding is developed in the older Mesoproterozoic basement units (>1150 Ma) and formed at upper amphibolite- to granulite-facies conditions. Younger Mesoproterozoic basement units (Group 3, <1090 Ma) typically lack a high-temperature fabric. A younger foliation is variably developed in both the basement and cover sequence, this foliation characteristically strikes northeast, dips to the southeast, and is defined by aligned greenschist-facies minerals. New Ar-Ar geochronology indicates that pervasive defor-

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mation in the basement and cover sequence occurred between ~380 and 310 Ma (late Devonian to Pennsylvanian). A suite of north-northwest (330°–350°) and west-northwest (280°–300°) transverse faults cut the regional structural grain and may be related to transtensional stresses during Mesozoic rifting. Valley and Ridge The second day of this field trip begins in Harrisonburg, Virginia, and follows Route. 33 west to Judy Gap, West Virginia, before terminating at the Allegheny Front at Briery Gap. Lithologic units west of the frontal Blue Ridge fault system include Cambrian and Ordovician carbonate rocks of the Iapetus (protoAtlantic) Ocean, deposited after the rifting of Rodinia (Fig. 2). West of Elkton, Virginia, these change into clastic sequences of the mid- to late Ordovician Taconic orogeny (Cooper and Cooper, 1946; Woodward, 1951; Diecchio, 1993), and in Fort Valley of the Massanutten synclinorium, Devonian black shales of the Acadian orogeny (Rader and Biggs, 1975, 1976). West of Harrisonburg, Cambrian to Ordovician carbonates are again apparent in the hanging wall of the west-directed Little North Mountain thrust fault. West of the Little North Mountain fault, lithologies are characterized by Upper Ordovician clastic rocks of the Taconic sequence (strikingly different from those exposed east of the fault), Silurian to Lower Devonian quartz arenites and carbonates deposited during the orogenic calm between the Taconic and Acadian orogenies (Dennison and Head, 1975; Rader and Perry, 1976; Smosna et al., 1977), and a thick sequence of Upper Devonian to Lower Mississippian clastic rocks (the Catskill clastic wedge) that represent the dominant expression of the Acadian orogeny in the mid-Atlantic region (Woodward, 1943; Dennison, 1970). This region is considered a type example of a foreland thrust belt, representing the foreland of the Alleghanian collision of the African part of Gondwana with Laurentia (ancestral North America). Major structural features include the Little North Mountain thrust in the east, progressing west through the Cacapon Mountain anticlinorium, Broad Top synclinorium, and Wills Mountain anticlinorium to the Allegheny Front (Fig. 3). The major fold structures are underlain by duplexes of Cambrian to Ordovician carbonate and clastic rocks (Kulander and Dean, 1986, Mitra, 1986; Evans,

Figure 3. NW-SE cross section across the central Appalachian foreland from Dunne (1996) that shows the subsurface imbricate duplex structure, interpreted to underlie the surface fold and fault pattern traversed during Day 2. WMA—Wills Mountain anticlinorium, BTS—Broad Top synclinorium, CMA—Cacapon Mountain anticlinorium, NMT—North Mountain thrust.

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1989), which implies that the surface rocks are the roof sequence of a blind thrust that likely propagates into the subsurface of the Allegheny Plateau (Rogers, 1970; Wilson and Shumaker, 1992; Dunne, 1996). Surface structures are apparent throughout the region at many scales, from outcrop-scale folds and thrusts with centimeters to meters of displacement, to folds with a wavelength of several kilometers, such as the breached Wills Mountain anticline where the floor of Germany Valley is exposed. The scale and geometry of structural features are controlled, in part, by mechanical aspects of alternating strong and weak lithologic sequences, such that significant westward transport likely occurs along weak, shale-dominated units (Martinsburg, Rose Hill, and Wills Creek formations), while outcrop structures are more apparent in the stronger sandstone and dolomitic units (Oswego, Tuscarora, Williamsport, and Oriskany formations). INTRODUCTION TO PEDAGOGIC APPROACHES Stratigraphy and structural geology are often presented as distinctly different subjects within the geoscience curriculum. Their vocabularies are distinct, and each requires examination of different features at the outcrop. For a stratigrapher, structural deformation is commonly viewed as “noise” that masks or gets in the way of the pertinent information, while structural geologists view stratigraphic units as mechanical layers that respond differentially to external stress, with little regard to specific depositional features. Yet what is common to both stratigraphic and structural analyses is that everything we observe in the rocks has resulted from energy transfer driven by tectonics. For many geologists, energy transfer is a partially abstract concept. The final products of the transfer of energy are apparent, but there are typically many pathways that energy could have traveled, as addressed by Chamberlin’s (1890) method of multiple working hypotheses. Conceptually, stratigraphic and structural analyses provide two distinct, but interrelated, approaches to evaluating tectonic evidence. In structural analyses the evidence of energy transfer is often direct: folds and faults are visually apparent or extrapolated from outcrop measurements, and energy transfer can be inferred from the degree of deformation or strain seen in the rocks. Conversely, in stratigraphic analyses based on flow regimes (the energy required to transport and deposit sediments), interpretations are not obvious from outcrop evidence. Most interpretations must be based on observations of modern depositional processes that may or may not be correlative with ancient environments. Stratigraphic interpretations that are further removed from direct outcrop evidence include the size, shape, and history of the basins in which sediments were originally deposited. General characteristics of basin subsidence may be apparent, but specific details on how far, how fast, and how deep the basin subsided are not directly obvious. Tectonic processes that cause depositional basins to subside also result in structural deformation of the rocks, but generally these large-scale processes cannot be fully analyzed from evidence in a single outcrop. Nevertheless,

the tectonic history of a region underpins fundamental questions that structural geologists and stratigraphers consider when evaluating data from individual outcrops. The approach taken in this field guide is to present outcrop-based stratigraphic and structural analyses techniques jointly in the larger context of tectonics, reflecting the approaches developed in the SST course. DEDUCTIVE (TOP-DOWN) AND INDUCTIVE (BOTTOM-UP) STRATEGIES This field guide includes a brief description of the theoretical framework and pedagogic strategies used in the SST course, where stratigraphy and structural analysis are synthesized in the service of tectonics and basin interpretation. It is not meant to be a theoretical and/or historical summary of the geology for professionals, but rather a pedagogic framework for field-oriented investigations and education. With this in mind, we consider two strategies for interpreting geology: deductive (top-down) and inductive (bottom-up). In our experience we have found that students often have trouble distinguishing between these two approaches, especially on the outcrop. A deductive approach is predominantly theoretical and revolves around conceptual models, such as plate tectonic regimes, geophysical models based on remote sensing data, and basin development models derived from ancient and modern examples. In the field, theoretical models are used in two ways: first, to guide observations on the outcrop by taking fragmentary information apparent in single outcrops and placing them in a regional context; and second, to test theoretical models by making predictions of what should be observed based on what has already been observed. For example, after observing an outcrop containing hummocky units, one could predict that down section the features would devolve into deep shelf or submarine fan deposits, while up section they would evolve into shoreline and terrestrial deposits. Correspondingly, structural analysis integrates theories of stress and strain with rheologic models of materials to interpret and predict deformed structures at all scales of observation. For example, small-scale folds theoretically mimic regional-scale folds that are too large to be seen at a single outcrop, such that Z and S folds in an outcrop are often used to infer the location of the axis of a larger fold. An inductive approach is empirical, where observations on the outcrop are the primary data for developing an understanding of geologic processes and larger-scale features. In addition, bottom-up observations are used to test the inferences and predictions of theoretical models. In practice real-world observations may diverge far enough from theoretical models for us to suspect that the theoretical models are either incomplete or wrong. This discrepancy often opens avenues to refining or expanding our understanding. The pedagogic strategy advocated here is to continuously pit top-down deductions and bottom-up inductions against each other until a logical contradiction emerges, or evidence and/or theory have been exhausted. This approach forces us to rethink theoretical models and/or gather more observational data from the

Blue Ridge and Valley and Ridge Provinces of northern Virginia and West Virginia outcrop. For the professors, the ideal outcome of these investigations is for each student, informed by their knowledge of theoretical models and using their stratigraphic and structural field observations, to build a complete and comprehensive tectonic and geologic history of the region. In the sections below we develop stratigraphic and structural theoretical models pertinent to this field trip and then discuss their pedagogic application in the field. THE TECTONICS OF STRATIGRAPHY: BASIN ANALYSIS Sedimentation requires the creation of accommodation space, and the creation and filling of accommodation space requires the transfer of energy. This energy is of two kinds: (1) tectonic energy, including heat from the Earth’s interior that drives plate motion, and associated gravitational energy; and (2) solar energy that drives atmospheric circulation, powers the hydrologic cycle, and results in climatic zonation. Once accommodation space is created, deposition at any one locality is controlled by these two sources of energy: (1) dip-fed energy of tectonic origin: water flowing down a slope, whether subaerial or subaqueous; and (2) strike-fed energy of solar origin: water flowing more or less horizontally, as along a beach (acknowledging tidal currents are also important). The production of accommodation space is controlled by five processes, although these typically work together. (1) Isostatic control—vertical response to loading and unloading of the crust. (2) Tectonic control—plate tectonic energy expressed in cycles of collision and extension, producing uplift and subsidence (basin formation). (3) Eustatic control—worldwide sea level changes, driven by both tectonics and climate. (4) Sediment supply—affecting how fast the accommodation space is filled. (5) Climate—most obviously during glacial cycles, but in lesser ways under other conditions. Basin analysis is the investigation and classification of basins based on how they form, how they evolve, what controls their evolution, and their geologic records. The depositional records occur in a nested hierarchy of patterns (called signatures) beginning with individual laminations, and working up through beds, bedsets, parasequences, systems tracts, sequences, and basinfilling sequences. The filling sequence of any basin results from several energy sources. The initial driving energy is tectonic, influencing the rate and degree of subsidence, followed by dipfed and strike-fed processes that fill the accommodation space, influenced by sediment supply and sea level changes. In this field guide we focus on two regional examples of basin development: (1) the rift-to-drift sequence, as represented by the Chilhowee group that overlies Blue Ridge basement rocks in north-central Virginia, and (2) the evolution of a foreland basin, as represented by the Silurian–Devonian clastic wedge in the Valley and Ridge province of western Virginia and eastern West Virginia.

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Theoretical Rift-to-Drift Stratigraphic Basin Evolution The theoretical model for a typical rift-to-drift sequence begins with uplift associated with the initiation of a hot spot. This produces tensional stresses, which results in horizontal extension and the formation of a horst (highland) and graben (rift) system (Fig. 4). A typical rift system consists of an axial graben and parallel lateral half-graben that lie on either side of the axial graben. The central axial graben may initiate as a subaerial feature, but it ultimately collapses below sea level (Fig. 4A). Shortly thereafter the drift phase commences, where oceanic crust starts to upwell along one of the central rifts, forming a new ocean basin (Fig. 4B). As this occurs the source of heat that lifted the hot spot migrates out to the center of the opening ocean, away from the new continental margins, which cool and subside. Initially, the subsidence is driven by thermal decay (cooling leading to increasing density and isostatic sinking), but as time goes by sediment loading contributes to the subsidence. Subsidence decays exponentially, such that subsidence in the rift phase is very rapid and slows progressively into the drift

Figure 4. Theoretical model of a rift sequence (cross sections A, B, and C), broadly applicable to the Catoctin, Swift Run, and Chilhowee Group (Weverton, Harpers, Antietam) stratigraphic record.

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phase (Fig. 4C). Correspondingly, in the rift phase accommodation increases very rapidly, more rapidly than sediment can keep up with, leading to a deep water environment. Then, as subsidence slows, sediment filling catches up with the accommodation space, leading to a coarsening and shallowing upward sequence. The transition of the rift phase to the drift phase produces a divergent continental margin (DCM) and a transgression of the shoreline across the former location of the rift. At this stage deposition is able to keep up with subsidence, and mostly coastal and shallow to deep shelf deposition results. Predictive Stratigraphic Model The predictive stratigraphic model presented here (based on Fichter and Diecchio, 1986) assumes that several dynamic processes are going on simultaneously, or at least overlap in time. First assumption: stratigraphic development in the rift-to-drift system is basically a coarsening and shallowing upward sequence. Second assumption: the Chilhowee Group stratigraphic record we are dealing with spans both the uplift and subsidence stages of the rifting. Third assumption: graben-filling occurred on the Blue Ridge western flank, based on visible Bouma sequences. Fourth assumption: unlike a typical fan delta system that finishes with a subaerial alluvial fan (which could still be there based on large planar cross-bedding in the Weverton Formation), the submarine fan grades up section into a shallow storm shelf (based on hummocky sequences in the Harpers Formation), and then beach facies of the Antietam Formation. The sequence below accounts for all of these conditions and processes. (1) Initial horst uplift and graben subsidence stemming from hot spot activity results in high relief. Depositional systems develop on both sides of the rift continental terrace (horst that borders the axial rift on each side) and flow in opposite directions, one toward the axial rift, the other toward the continent. Toward the rift, fan deltas develop in the graben, while on the opposite side rivers flow toward the continent. To our knowledge, no remnants of such river systems have been described for the midAtlantic region. The graben fill results in a coarsening upward sequence, from basin shales, through submarine fan, and (ideally) into an alluvial fan. High relief results in rapid erosion and deposition. (2) With time, erosion lowers the horsts and deposition fills the accommodation space, lessening the relief. Submarine fan turbidite deposits (Bouma sequences) shallow upward. (3) At about the same time, the continental margin is subsiding, eventually sinking below sea level resulting in a transgression. Shallow marine shelf deposits appear in the section. These may deposit on top of submarine fan Bouma sequences, or may transgress across fluvial systems (top of the fan delta). Both are possible in different locations. Tops of horsts may appear as islands. (4) As the continental margin subsides and the horsts erode (and/or are buried in debris) continental drainage

reverses. Where the gradient and flow used to be toward the continent, the gradient and flow are now toward the DCM. This taps the deposits of lag quartz sands found on continental cratons and brings them down to the new coastal systems. (5) Finally, as the relief diminishes, the continental margin subsides and the shore transgresses. Sea level fluctuations begin to more strongly influence the stratigraphic record. Local Rift Stratigraphic Record On the west flank of the Blue Ridge anticlinorium, the formations representing rift events are the Catoctin metabasalts (initiation of proto-Atlantic oceanic crust formation), Swift Run siliciclastics (interlayered with metabasalts), Weverton submarine fan (Bouma sequences), Harpers shelf (hummocky sequences), Antietam (beach system), and Shady (tidal dolomite). Outcrops of the Swift Run, Weverton, Harpers, and Antietam formations are sparse and not contiguous. On this field trip, we will see all of these formations except the Shady dolomite. A Harpers outcrop on Route 211 near Luray (north of our field trip route) contains hummocky sequences. However, we will visit a Harpers locality at the base of the Blue Ridge on Route 33 that contains Bouma sequences (Stop 1-8). Theoretical Foreland Basin Stratigraphic Basin Evolution In our theoretical model, foreland basin development is preceded by a tectonic calm, either a divergent continental margin (DCM) or an interorogenic calm (Fig. 5, Stage I). Tectonic calms are recognized by quartz arenite and/or carbonates, both of which are stratigraphically thin (tens to hundreds of feet), as compared with the stratigraphic thickness of orogenic events (thousands of feet). The early stages of a collisional orogeny produce rapid foreland subsidence (Fig. 5; Stages II–III) that exponentially slows with time. Similar to rift basins, foreland basins initiate with the formation of deep water environments followed by long-term filling of the accommodation space by progradation and filling (Fig. 5; Stages IV–V). This pattern is best developed adjacent to the maximum relief produced by mountain building and diminishes as one moves away from the orogenic core. Predictive Stratigraphic Model The ideal model of a progradation and filling sequence follows Walther’s Law: the horizontal sequence of environments will stack vertically in the same order during a transgression or regression (Middleton, 1973; Fig. 6). The order of environments is controlled by the systematic changes in the way energy is transferred. Following initial tectonic subsidence, the main energy driving the system switches to sedimentation (dip-fed and strike-fed) and sea level changes. Figure 6 shows the preferential stacking order for the filling sequence of an ideal foreland basin. This sequence becomes the top-down basis for making interpretations and predictions on the outcrop.

Blue Ridge and Valley and Ridge Provinces of northern Virginia and West Virginia

Figure 5. History of an ideal foreland basin stratigraphic filling sequence; after Fichter and Poché (2001).

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Local Foreland Basin Stratigraphic Record Two distinct foreland basin records are encountered within the region covered by this field guide, one for the Ordovician Taconic orogeny and the other for the Devonian Acadian orogeny (Fig. 2). The Acadian sequence (Catskill clastic wedge) in this part of Virginia and West Virginia is almost a perfect match for the theoretical model (Fichter, 1986). The Needmore Formation shows anoxic features, fossil rich layers, and condensed sections that collectively represent the early stages of a rapidly subsiding basin. The Millboro Formation is composed of anoxic, thinly laminated shales that indicate the deepest environment of the section. The laminated shales grade up section into Bouma sequences of the Brallier Formation, which transition into hummocky sequences of the Greenland Gap Formation and then point bar sequences of the Hampshire Formation. The red beds of the Hampshire Formation indicate a muddy shoreline environment, which is a slight deviation from the theoretical model. Simplistically we would expect to see a sandy shoreline, either because that is what most people see when they visit the beach or because many coastlines today consist of sandy beach and barrier island systems. The absence of a sandy beach and presence of a muddy shoreline is useful pedagogically as it prompts on-the-outcrop discussions of coastal energy regimes and climatic zones. In contrast to the almost ideal Acadian foreland basin record, the Taconic stratigraphic record deviates substantially from the theoretical model. Three distinct Taconic sequences occur along a west-to-east transect of the Valley and Ridge province: Germany Valley (westernmost section), Brocks Gap (middle section along Little North Mountain), and the Massanutten Synclinorium section (Fig. 7). Due to time constraints, we only see the Germany Valley section on this field trip. Oblique collisions of the Taconic terrane(s) left the present-day northern Virginia region in a protected reentrant that experienced only slight flexure fold-

ing (Thomas, 1977; Drake et al., 1989). This divided the Taconic foreland basin into a rapidly subsiding deep water eastern basin and a slowly subsiding shallower western basin (Diecchio, 1986, 1993). These basins remained isolated from each other until deposition of the Oswego Formation in the late Ordovician. Thus, the western basin (the Germany Valley section of this field guide) never developed a deep water system. Instead, carbonate hummocky sequences (Black River and Trenton groups) grade into clastic hummocky sequences (Reedsville Formation), and shoreline sandbars (Oswego Formation). This is overlain by a clastic tidal system (Juniata Formation red beds) that does not transition into a fluvial system, but rather transitions into a beach and barrier island system (Tuscarora Formation). The details of the Taconic stratigraphic sequence are compared and contrasted with the Acadian sequence, which prompts discussions of the relative merits of deductive versus inductive analysis techniques and their relative applicability to field investigations. Equating Inductive and Deductive Approaches to Stratigraphic Analysis One of the consequences of evaluating the geologic record as an energy-transferring evolutionary system is that stratigraphic sequences can be viewed as patterns that are replicated at many scales: an example of natural fractal organization. For large-scale tectonic synthesis, geologists and students need to know where they are in the hierarchy of stratigraphic sequences, know the energy processes that produced each sequence, and know how and why the sequences are nested in a particular order. On the outcrop this can be evaluated inductively (bottom-up). For example, fining upward Bouma sequences that are nested within fining upward parasequences (representing sea level changes), that are nested within composite parasequence sets (systems tracts, either

Figure 6. Ideal progradational and corresponding stacking sequence of depositional environments.

Blue Ridge and Valley and Ridge Provinces of northern Virginia and West Virginia coarsening or fining upward), that are nested within an overall coarsening upward sequence, collectively represent a prograding submarine fan environment. Similar patterns exist in shelf sequences, tidal sequences, and fluvial sequences. However, by the time larger scale components of this hierarchy are considered, the patterns can no longer be viewed at a single outcrop. At this point, patterns are deductively evaluated within the framework of a large-scale theoretical model to see how effectively they correspond to the outcrop evidence. Stratigraphic sequences in the Blue Ridge and Valley and Ridge provinces of northern Virginia and West Virginia exhibit a rift-to-drift sequence followed by several orogenic events (Taconic, Acadian, and Alleghanian), followed by a second riftto-drift sequence (opening of the Atlantic Ocean). Each of the orogenic events represents a pulse of tectonic activity that decays exponentially over time and includes systematic filling of the accommodation space during the decay periods. With each pulse

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of tectonic activity the stratigraphic record responds in generally predictable ways. Thus, theoretical models that are based on how energies are transferred predict that particular sequences (e.g., Bouma sequences; Fig. 8) will not appear haphazardly in the stratigraphic column. By evaluating outcrop evidence bottom-up, with ideal basin-filling sequences as testable models, we and our students can build plausible regional-scale tectonic histories from the visible stratigraphic record. THE TECTONICS OF STRUCTURAL ANALYSIS: DEFORMATION AT MULTIPLE SCALES In the late nineteenth century, Raphael Pumpelly articulated his “rule” that the axes of minor folds are typically congruent with major fold structures of the same phase of deformation (Pumpelly et al., 1894). This is often generalized by stating that small-scale structures mimic larger scale structures, a theoretical

Figure 7. Interpretive stratigraphic and tectonic cross sections for the Taconic sequences of Germany Valley, Little North Mountain (Brocks Gap), and the Massanutten Synclinorium. Top: cross section after Diecchio (1993).

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principal that fits within fractal analyses. Structural geologists use Pumpelly’s Rule to infer regional structures that are not observable at a single outcrop from detailed outcrop measurements, an example of integrating theoretical models with inductive observations. However, a critical assumption is that structures at all scales resulted from a single tectonic event. This principle of patterns of deformation at multiple scales is abundantly in evidence along the Valley and Ridge transect described by this field guide, where folds and faults are apparent at the outcrop scale and inferred at larger, regional scales. The concept of energy transfer is demonstrable along the east-to-west progression of field stops, such that the WNW-directed collision of Gondwana with this region of eastern Laurentia produced deep-rooted, penetrative deformation in the basement rocks of the Blue Ridge, which shallows to thin-skinned deformation across the Valley and Ridge, before finally dissipating in the subsurface of the Allegheny Plateau. Ductile deformation is primarily observed in basement rocks east of the Blue Ridge thrust system along the western margin of Shenandoah National Park. Mylonitic shear zones have a similar NE-SW orientation to, and can be temporally equated with, brittle faults of the foreland thrust belt of the Valley and Ridge to the west (Bailey et al., 2006). The geometry of folds and faults is controlled, in part, by the mechanical strength of rock units, with most of the westward transport accommodated along shale-dominated units like the Martinsburg Formation (Dunne, 1996). The scale of deformation structures range from those visible at a single outcrop (e.g., Stop 2-19), to folds that span several kilometers (e.g., Germany Valley anticline, Stop 2-24). Seismic data suggests that multiple-scale deformation patterns extend into the subsurface, where fault-bounded horses underlie much of the visible surface patterns (Evans, 1989; Wilson and Shumaker, 1992; Fig. 3). This field guide documents stops that exhibit ductile and brittle folds and faults, of at least three distinguishable orders of magni-

tude. We consider first-order structures to be at the several-kilometer scale, and thus not directly observable at a single outcrop; the Germany Valley anticline (Stop 2-24) is a good example. Evidence for second-order structures is apparent at several outcrops (Stops 2-12 and 2-19), and the scale of a specific structure may be tens of meters to a kilometer. Third-order structures are a few meters in scale and readily apparent at a single outcrop, such as Stops 2-10 and 2-19. Mega-scale structures, like the subsurface duplexes in Figure 3, and microscopic-scale structures are inferred from other data sets, such as the mylonitic structures in Blue Ridge basement rocks from Day 1, but will not be directly observed on this field trip. CONCLUSIONS The deductive and inductive approaches to stratigraphic and structural analyses outlined above provide a multifaceted toolbox for outcrop-based field investigations. Our integrated approach forces novice geoscientists to utilize all of these techniques for the overarching goal of synthesizing the tectonic history of the northwestern Virginia and eastern West Virginia region. As students work toward this goal, it becomes apparent to them that outcrop evidence exists not only at several spatial scales but also at several temporal scales. Stratigraphic data provide information about the Taconic and Acadian orogenic events, while structural data predominantly reflect deformation from the Alleghanian orogeny. The task of interpreting and synthesizing observations at multiple spatial and temporal scales is daunting for both novices and experts alike, but it is ultimately what geoscientists are required to do in order to build compelling tectonic histories. It is our good fortune that this region is rich in outcrops that facilitate multidisciplinary investigations of several tectonic episodes that collectively cover most of the Appalachian orogenic cycle. For students completing the SST course, the expectation is that the theoretical models and concepts presented here, once

Figure 8. Theoretical model showing how specific sequences, such as Bouma sequences, are predicted from the transfer of tectonic and stratigraphic energy.

Blue Ridge and Valley and Ridge Provinces of northern Virginia and West Virginia absorbed, allow one to evaluate the preserved stratigraphic and structural record, and then predict what might be seen laterally in nearby outcrops. The deductive experiences on the outcrop temper this expectation with the reality that theoretical models do not always match field observations. However, these discrepancies between models and outcrop evidence provide an opportunity to refine the models. From a pedagogic perspective, what is important about this approach is that students begin to construct a framework of interpretation and understanding, and with more and more experience, begin to think about the geologic and tectonic record in terms of linked systems and energy.

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Stop 1-1. Pavement Outcrops in Stanardsville Yard (Older Pyroxene-Bearing Gneiss) Mile 10.4 (N38° 18′ 02″, W78° 26′ 06″)

The road log for the Blue Ridge part of the field trip (Day 1) starts at the junction of U.S. Route 29 and Virginia Route 230, in Ruckersville, south of Madison, Virginia (N38° 20′ 50″, W78° 17′ 06″; mile 0.0). Day 1 stops are shown on a relief map of the area (Fig. 9) and on a simplified geologic map (Fig. 10).

Home of Jackie Pamenter. Pavement outcrops in center of driveway loop and behind outbuilding of coarse-grained, orthopyroxene-bearing Mesoproterozoic granitoid gneiss that is one of the older basement lithologies in this part of the Blue Ridge, with a crystallization age of 1177 + 11 Ma. The gneiss displays a well-developed, high-temperature tectonic foliation, defined by mm- to cm-scale feldspar and quartz aggregates with elongate clots of orthopyroxene and hornblende—lithologies with ages <1100 Ma lack such a strong foliation. The gneissic foliation strikes northeast and dips to the northwest; elsewhere in the Blue Ridge, northwest strikes are not uncommon for rocks with similar gneissic fabrics of “Grenville” age. Note orange-weathering boulders to left of shed—this distinctive weathering rind is characteristic of pyroxene-bearing granitoid rocks. The sculpture in the center of the loop features a piece of Catoctin metabasalt.

Cumulative Point-to-point mileage mileage Directions and comments

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ROAD LOG, DAY 1: BLUE RIDGE

Proceed westbound on VA Rt. 230. Cross Conway River; enter Greene County. Turn right into driveway; proceed to driveway loop and park.

Turn left (north) on VA Rt. 230 (Octonia Road). Turn left (northwest) onto Rt. 637. Turn left (south) onto Tall Pines Drive.

Figure 9. Shaded relief map with major towns and roads, showing the field trip stops for Day 1.

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12.5

0.3

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Bear right onto Greene Acres Road North. Park on left.

Stop 1-2. Roadcut on Greene Acres Road North (Basement-Derived Mylonite) Mile 13.6 (N38° 19′ 42″, W78° 27′ 54″)

Stop 1-3. Roadside Outcrop on Turkey Ridge Road (Younger Mesoproterozoic Charnockite) Mile 15.0 (N38° 20′22″, W78° 27′ 58″)

Protomylonite and mylonite are well exposed in this cut. The protolith for these strongly deformed rocks is younger Mesoproterozoic charnockite (see next stop), and this high-strain zone can be traced for 10 km to the north. Foliation dips to the east and a strongly developed penetrative lineation plunges obliquely downdip. Kinematic indicators such as asymmetric porphyroclasts and shear bands indicate a top-to-the west shear sense (reverse fault); small, enigmatic outliers nearby of Catoctin metabasalt on basement suggest that the shear zone had an earlier, top-to-the-east shear sense (normal fault).

This exposure shows massive-weathering, coarsegrained, orthopyroxene-bearing granitoid, also known as charnockite. This lithology is one of the youngest Mesoproterozoic lithologies in the Blue Ridge, with a crystallization age of ~1050 Ma that indicates emplacement near the end of the Grenville orogeny. Accordingly, the rock lacks a tectonic foliation; faint foliation observed in some charnockite outcrops is probably magmatic in origin. The contacts between this young charnockite and the older gneisses are primarily intrusive in nature; in a roadcut on U.S. Route 33, a narrow dike of this type of charnockite intrudes wellfoliated granitoid gneiss. Similar to the rocks at Stop 1-1, this pyroxene-bearing rock weathers to produce a distinctive orange rind, as seen in nearby boulders.

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Continue west on Greene Acres Road North. Turn right (west) onto Greene Acres Road South. Turn right (north) onto Rt. 638 (Turkey Ridge Road). Pull off on right at slight bend.

Turn left (west) onto Rt. 634 (Bull Yearling Road). Turn right (west) onto U.S. Rt. 33. Pull off into small area on the right (north) side of U.S. Rt. 33. Walk northward for 0.3 mi. to the exposure.

Figure 10. Simplified geologic map of a region of north-central Virginia. Field trip stops for Day 1 are indicated.

Blue Ridge and Valley and Ridge Provinces of northern Virginia and West Virginia Stop 1-4. Cliffs along Tributary of East Swift Run (Swift Run Formation) Mile 20.2 (N38° 20′ 56″, W78° 30′ 53″) The low cliffs form one of the best exposures of the Swift Run Formation in Shenandoah National Park. The original type location (Jonas and Stose, 1939) was 1.6 km to the west, but that outcrop was destroyed during the widening of U.S. Rt. 33, and this outcrop makes an admirable neostratotype for the Swift Run Formation. Approximately 10–12 m of arkosic metasandstone, metaconglomerate, and mudstone are exposed. Bedding dips gently southeast and is cut by a more steeply dipping foliation. Cross and plane bedding is common, and graded bedding and mudstone rip-up clasts are evident at a few locations. Coarse sandstone layers commonly truncate layers of maroon mudstone. The provenance for these sediments is the granitic basement complex, and sedimentary structures are consistent with a fluvial to fluvial braidplain depositional environment. Cores were drilled for paleomagnetic analysis in an effort to determine the paleolatitude during Swift Run deposition, but the remnant magnetic signature in these rocks is complex. Cumulative Point-to-point mileage mileage Directions and comments 20.2

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Bear right (northwest) onto U.S. Rt. 33. Swift Run Gap cross under the Skyline Drive. Turn right into Shenandoah National Park entrance and proceed through entrance station. Turn right (south) on the Skyline Drive. Turn left into Smith Roach Gap parking area. Hike eastward on old road for 0.3 mi to outcrop.

Stop 1-5. South Flank of Hightop Mountain (Catoctin Formation) Mile 26.1 (N38° 19′ 35″, W78° 34′ 11″) The walk from the parking area takes us across foliated greenstone interlayered with thin layers of arkosic phyllite in the Catoctin Formation. At ~400 m a large outcrop of greenstone crops out immediately to the north of the trail. Columnar joints are well developed and transected by a southeast-dipping foliation. The long axes of the columns plunge steeply to the southeast, consistent with a gently dipping flow geometry. On the south flank of Hightop the Catoctin Formation dips gently to the southeast. Cumulative Point-to-point mileage mileage Directions and comments 26.1

0.0

Turn right (north) onto Skyline Drive.

26.9

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Turn left into Sandy Bottom Overlook.

Stop 1-6. Sandy Bottom Overlook (Charnockite in the Basement Complex) Mile 26.9 (N38° 20′ 10″, W78° 33′ 57″) Sandy Bottom Overlook provides a commanding view of the Blue Ridge (northeast), the Shenandoah Valley (west), and Massanutten Mountain (west/northwest). Well-cemented quartz arenites of Silurian age underlie the long linear ridges of the Massanutten Mountain complex. The canoe-shaped prow at the southwestern end of Massanutten reflects the geometry of this gently plunging syncline. The mountain front between the Elkton area and north to Stanley is sinuous and likely reflects the lowangle nature of the Blue Ridge fault and offset associated with younger transverse faults. The “great unconformity” between Mesoproterozoic basement and the Neoproterozoic cover sequence occurs above the Skyline Drive. Basement is overlain by arkosic phyllite of the Swift Run Formation and at least 250 m of Catoctin metabasalts capping Hightop Mountain. Sandy Bottom (immediately to the northwest) is underlain by a northwest-striking transverse fault that places basement against the Catoctin Formation. The large roadcut across from the overlook exposes a distinctive orthopyroxene-bearing monzogranite. A U-Pb zircon age of 1049 ± 9 Ma was obtained from rock at this location and is interpreted as a crystallization age (Southworth et al., 2009). This charnockitic pluton is part of the younger group of Grenvillian granitoids. Cumulative Point-to-point mileage mileage Directions and comments 26.9 29.1

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Turn left (north) onto Skyline Drive. Turn left onto Shenandoah National Park entrance road. Turn right (west) onto U.S. Rt. 33. Turn left (southwest) onto Rt. 628. Park on the right across from driveway.

Stop 1-7. Roadcut on Route 628 (Weverton Formation) Mile 34.9 (N38° 20′ 14″, W78° 37′ 16″) The long roadcut exposes a coarse arkosic metasandstone and quartzose phyllite of the Weverton Formation. Bedding is upright and dips 15° to 25° to the northwest. A southeast-dipping foliation is well developed in the phyllite. In the Elkton area finegrained rocks are common in both the Weverton and overlying Harpers Formation making the distinction between units difficult. Fine-grained white mica in the quartzose phyllite yields an Ar/Ar plateau age of ~340 Ma and is interpreted to be either a cooling or growth age. In either case, penetrative deformation in this rock occurred prior to the Alleghanian orogeny.

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Two sets of fractures cut the Weverton Formation at these exposures. The older set strikes to the northeast and dips normal to bedding, consistent with fracturing prior to folding. The younger subvertical set strikes to the northwest, and fractures are adorned with plumose structure; this set may have developed under a west-northwest oriented principal stress associated with late Paleozoic contraction. Cumulative Point-to-point mileage mileage Directions and comments 34.9

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Retrace route to the northeast on Rt. 628. Turn left (west) on U.S. Rt. 33. Park on right shoulder, cross to south side of the divided highway.

Stop 1-8. Roadcut on U.S. Route 33 (Harpers Formation) Mile 39.0 (N38° 22′ 31″, W78° 35′ 18″) This roadcut exposes laminated quartzose metasandstone of the Harpers Formation. Bedding dips gently to the northwest, and a southeast-dipping penetrative foliation (cleavage) is well developed in the fine-grained rocks. Depositional units are proximal Bouma sequences, mostly TAE and TABE, from one to a few dm thick. TA units with scoured bottoms, ranging from coarse quartzose sands to granule and pebble gravels with graded bedding; all consistent with a steep gradient and deeper water marine depositional environment. Cumulative Point-to-point mileage mileage Directions and comments 39.0 41.8 42.0

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Continue northwest of U.S. Rt. 33. Exit to the right off U.S. Rt. 33. Turn left onto U.S. Rt. 340 at the end of ramp. Park on the right (northwest), and carefully cross road.

Stop 1-9. Roadcut on U.S. Route 340, South of Elkton (Antietam Formation) Mile 45.7 (N38° 23′ 03″, W78° 37′ 42″) The exposure to the southeast side of U.S. Rt. 340 is cut into extensively fractured and brecciated quartz arenite of the Antietam Formation. These rocks are in the hanging wall of the Blue Ridge fault that is buried beneath alluvium of the South Fork of the Shenandoah River immediately to the northwest. Wells drilled on the Merck chemical plant penetrated middle Cambrian rocks of the Elbrook Formation (King, 1950) within 0.5 km of this outcrop, and the intervening Waynesboro and Tomstown formations are absent. Cumulative Point-to-point mileage mileage Directions and comments 45.7

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U-turn and go north on U.S. Rt. 340.

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Cross under U.S. Rt. 33 and enter Elkton. Turn right into Shenandoah town park.

Stop 1-10. Outcrops in Shenandoah Town Park (Beekmantown Group and Conococheague Formation) Mile 52.9 (N38° 28′ 56″, W78° 37′ 02″) Exposures of strongly deformed limestone, with abundant sheared stylolites, crop out on the north side of Gem pond, in Shenandoah town park on the east side of Rt. 340. Outcrops at the west end of the pond are interpreted as the Stonehenge member of the Beekmantown Group, while outcrops at the eastern end of the pond are Conococheague Formation. Cliffs along the eastern side of the pond exhibit several layer-parallel zones of fault gouge, interpreted to be splays of the northwest-directed Stanley fault. The main fault surface is probably the erosional notch at the northeastern corner of the pond; this is also potentially the contact between the Ordovician Beekmantown Group and the Cambrian Conococheague Formation. End of Day 1 field stops. Drive back south on Rt. 340 (5.4 miles), and turn right onto ramp to join Rt. 33 west. Drive ~17 miles to Harrisonburg, Virginia. At about mile 10 along this drive you will pass the prow of Massanutten Mountain, just north of Rt. 33, which is defined by Silurian Massanutten Formation sandstones overlying the Ordovician Cub sandstone (hummocky sequences) and Martinsburg Formation (Bouma sequences) and represents the western syncline of the Massanutten Synclinorium. The eastern syncline is apparent as the mountain of Lairds Knob, which is east and a bit north of Massanutten Mountain. At about mile 13 you cross over the west-directed Staunton thrust fault (not exposed), which juxtaposes Cambrian Elbrook Formation limestones (hanging wall) against Ordovician Beekmantown Group carbonates (footwall). The city of Harrisonburg predominantly sits on subhorizontal Ordovician carbonates and shales of the Edinburg and Martinsburg formations. ROAD LOG, DAY 2: VALLEY AND RIDGE The road log for the Valley and Ridge part of the field trip (Day 2) includes many more stops than we will have time to visit in a single day. We have included optional stops for the possible future use of field trip participants and field guide readers. Day 2 stops we hope to visit are shown on a relief map of the area (Fig. 11) and on a simplified geologic map (Fig. 12). Note that we visit all of the stops listed below in the SST course, but it takes us two to three days to accomplish this. Readers will also note that the stop descriptions are extremely terse. This format mimics the information given to the SST students, where most of the details are deciphered during examinations and discussions on location at

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Figure 11. Google Earth relief map showing the field trip stops for Day 2.

Blue Ridge and Valley and Ridge Provinces of northern Virginia and West Virginia

Figure 12. Simplified geologic map of a region of northwestern Virginia (after VDMR, 2003) and eastern West Virginia (after WVGES, 1998), draped over Google Earth relief map. Field trip stops for Day 2 are indicated.

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Blue Ridge and Valley and Ridge Provinces of northern Virginia and West Virginia each outcrop. Our intention is to duplicate this approach during this field trip, and therefore we have limited these stop descriptions to only enough basic information to get discussions initiated. Day 2 of the field trip starts in Harrisonburg, Virginia, at the junction of High St. and Rt. 33 (N38° 27′ 00″, W78° 52′ 22″; mile 0.0). Drive west on Rt. 33 (Rawley Springs Rd.), passing the Frazier quarry on the right (mile 0.7) in the Ordovician Edinburg Formation. Black, micritic limestones are quarried here, known locally as “bluestone.” Continue driving west down section through carbonate rocks dipping shallowly east. Pass Mole Hill in the left (mile 3.9), an Eocene basaltic volcanic plug that exhibits pyroxene ± olivine phenocrysts and occasional quartz sandstone(?) xenoliths. Stop 2-1 (Optional). Small Outcrop along the North Side of Route 33 (Conococheague Formation) Mile 7.0 (N38° 28′ 37″, W78° 59′ 24″) Small limestone outcrop of Cambrian–Ordovician Conococheague Formation. This is the hanging wall of the North Mountain thrust; footwall rocks can be seen in the ridges just west of here. Pass very small limestone outcrop of Conococheague Formation along the south side of Rt. 33 at mile 7.7.

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Stop 2-4. Extensive Roadcut on the North Side of Route 33 (Hampshire Formation) Mile 22.6 (N38° 35′ 19″, W79° 10′ 05″) Shallowly east-dipping siltstones, sandstones, and gravel conglomerates, featuring point bars, levees, and crevasse splays of a meandering river system. Chlorite slickenlines can be seen on some bedding planes. The next few stops progress down stratigraphically through the Devonian clastic wedge. Stop 2-5. Large Roadcut on the North Side of Route 33 (Contact between Greenland Gap and Hampshire Formations) Mile 24.2 (N38° 35′ 53″, W79° 10′ 35″) Variable shales, siltstones, sandstones, and a thick quartz pebble conglomerate of the Greenland Gap Formation transition into red sands and gravels of the Hampshire Formation. This location represents the transition from a marine shelf to a muddy shoreline depositional environment. Small contraction faults can be seen at the western end of the roadcut. Stop 2-6 (Optional). Overgrown Roadcut on the North Side of Route 33 (Greenland Gap Formation) Mile 26.8 (N38° 36′ 41″, W79° 13′ 01″)

Stop 2-2. Small Abandoned Quarry on the North Side of Route 33 (Greenland Gap Formation) Mile 9.3 (N38° 29′ 45″, W79° 01′ 26″)

Open anticline of shale, siltstone, and sandstone beds of the Greenland Gap Formation, with a possible east-directed, highangle backthrust at the eastern end of the roadcut.

You have just crossed over the west-directed North Mountain thrust and are now at an old quarry in the Devonian Greenland Gap Formation (hummocky sequences). This is the proximal footwall to the thrust, and the footwall syncline drag fold can be seen at the eastern end of the quarry. Beds are steeply overturned along the eastern wall of the quarry and transition to subhorizontal at the western end.

Stop 2-7. Quarried Outcrop behind the West Virginia Department of Transportation (WV DOT) Shed on the North Side of Route 33 (Brallier Formation) Mile 28.0 (N38° 37′ 19″, W79° 14′ 08″)

Stop 2-3 (Optional). Cliffs on the North Side of Route 33 (Pocono Formation) Mile 11.4 (N38° 30′ 27″, W79° 02′ 51″) Cross-bedded sandstone of the Mississippian Pocono Formation, dipping shallowly to the east. The Pocono Formation is the youngest unit along this traverse. Proceed west on Rt. 33 up the long incline that leads to the top of Shenandoah Mountain, passing many outcrops of Pocono Formation and dark-red Hampshire Formation (Devonian). Note that the incline of the road is roughly equivalent to the shallow east dip of the rocks, thus you don’t cross much section through here. The crest of Shenandoah Mountain (mile 21.6; N38° 34′ 44″, W79° 10′ 04″) is Hampshire Formation and also the Virginia–West Virginia state boundary.

Abundant Bouma (turbidite) sequences are apparent in the large quarried cliff face behind the equipment shed and gravel piles. Dip is shallow to the east. Turn right at small town of Brandywine (mile 28.6), staying on Rt. 33. Stop 2-8. Large Roadcuts on Both Sides of Route 33 (Millboro Formation) Mile 29.9 (N38° 37′ 59″, W79° 13′ 55″) Park on east side of Rt. 33, just south of the roadcuts, where the road curves. Dark-gray to black shale and siltstone beds dip shallowly to the east. Abundant fauna can be seen on some bedding planes, including dwarf brachiopods, clams, and nautiloids. Occasional concretions are apparent, and convolute bedding (soft-sediment deformation) can be seen in the brown, silty beds near the base of the outcrop on the east side of Rt. 33.

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Stop 2-9. Large Roadcut on the South Side of Route 33 (Needmore Formation) Mile 33.1 (N38° 39′ 32″, W79° 14′ 01″)

Stop 2-13 (Optional). Roadcuts on Both Sides of 90° Bend in Route 33 (Rose Hill Formation) Mile 36.0 (N38° 40′ 24″, W79° 16′ 34″)

Subhorizontal dark-gray shales and silts of the Needmore Formation with condensed sections (black) and fossiliferous horizons (gastropods, cephalopods, brachiopods, and trilobites). Laminated sands occur near the top of the outcrop.

Dangerous location along Rt. 33, but access problems can be mitigated by parking on Ravenswood Lane just west of the roadcut, and hiking above the roadcut on the south side of Rt. 33 to view the roadcut on the north side of Rt. 33. A somewhat eroded example of a complex anticlinal structure of smallscale (third-order) folds and faults. Pass small-scale folds in the Helderberg Group carbonate rocks at mile 36.9.

Stop 2-10. Long Outcrop of Cliffs on South Side of Uphill Section of Route 33 (Oriskany Formation, Helderberg Group, Tonoloway Formation) Mile 33.5 (N38° 39′ 34″, W79° 14′ 18″) The base of the uphill road incline consists of cliffs of Oriskany Formation quartz arenite dipping steeply to the east; brachiopod molds are apparent on bedding planes. Progressing uphill to the northwest along the south side of the road, the beach and near-shore facies of the Oriskany Formation transition into subtidal units of the Helderberg Group. These include (moving uphill, but down section): the Licking Creek and Mandata formations (poorly exposed cherts and carbonates), New Scotland Formation (thick-bedded, fossiliferous limestone with chert bands), Coeymans Formation (bryozoan-rich calcareous siltstones), and Keyser Formation (two zones of coarse-grained calcarenite, separated by the poorly exposed Big Mountain shale). The top of the hill consists of algal-laminated micritic limestone of the Tonoloway Formation. Tonoloway beds at the top of the hill are horizontal, indicating the core of an anticline. Parasitic, west-vergent folds are also apparent just before the top of the incline. Note that the Helderberg Group is the base of the Devonian, and the Tonoloway Formation is the uppermost Silurian unit. This is the eastern flank of a first-order fold, the western flank of which is apparent at Stop 2-14. Stop 2-11 (Optional). Outcrop on North Side of Route 33 (Tuscarora Formation) Mile 35.0 (N38° 40′ 01″, W79° 15′ 41″) Clean white sandstones of the Tuscarora Formation, dipping steeply to the east. Stop 2-12. Low Cliffs Just behind the Trees on the North Side of Route 33 (Tuscarora Formation) Mile 35.5 (N38° 40′ 08″, W79° 16′ 09″) Clean white sandstone of the Silurian Tuscarora Formation dipping steeply to the west. You have just driven through an anticline cored by the red shale and sandstone of the Juniata Formation, which were visible in roadcuts on the north side of Rt. 33. Note the subhorizontal Tuscarora cliffs at the top of the distant hillside above you to the east.

Stop 2-14 (Optional). Cliffs on the South Side of Route 33 (Contact between Oriskany and Needmore Formations) Mile 37.7 (N38° 40′ 25″, W79° 17′ 55″) Sandstone beds (cliff formers) of the Oriskany Formation and shales of the Needmore Formation all dip steeply to the west. This is the western flank of a first-order fold, for which the eastern flank is apparent at Stop 2-10. Stop 2-15 (Optional). Low Roadcuts on Both Sides of Route 33 (Needmore Formation) Mile 38.6 (N38° 39′ 57″, W79° 18′ 46″) Moderately southeast-dipping shales of the Needmore Formation; last Devonian outcrop until the final Briery Gap section. Lunch in Franklin. Stop 2-16 (Optional). Roadcuts on Both Sides of Route 33 at Western Outskirts of Franklin, West Virginia. Subhorizontal Beds of Oriskany Formation Sandstones Mile 40.6 (N38° 39′ 11″, W79° 19′ 51″)

Stop 2-17 (Optional). Roadcuts on North Side of Route 33 (Contact between Helderberg Group and Oriskany Formation) Mile 41.9 (N38° 39′ 07″, W79° 21′ 04″) Carbonate and sandstone beds dip steeply to the west. This section of the traverse has several small outcrops of Oriskany Formation and/or Helderberg Group, with dips that often alternate between east and west, suggesting undulating small-scale folds. Stop 2-18 (Optional). Quarry on North Side of Route 33 (Tonoloway Formation) Mile 43.6 (N38° 39′ 19″, W79° 22′ 51″) Moderately east-dipping beds of algal-laminated micrites.

Blue Ridge and Valley and Ridge Provinces of northern Virginia and West Virginia Stop 2-19. Roadcut on North Side of Route 33 (Oriskany Formation) Mile 44.4 (N38° 39′ 30″, W79° 23′ 38″) Park at small pullout on south side of Rt. 33. The roadcut shows an outcrop-scale faulted fold in sandstone beds of the Devonian Oriskany Formation. The faulted fold is a break thrust, with a well-defined hanging-wall anticline and footwall syncline. This is a good example of the smallest scale (thirdorder) deformation structure that we will see during the Day 2 traverse. Pass an outcrop of east-dipping Tuscarora Formation sandstone at mile 44.8. Stop 2-20. Roadcut on North Side of Route 33 (Juniata Formation) Mile 45.1 (N38° 39′ 50″, W79° 24′ 17″) Parking is problematic; best to park around the bend in the road (west of the outcrop) and then walk back to the roadcut. Juniata Formation red shale and sandstone beds show a welldefined anticline, interpreted as a ramp anticline over a cryptic west-directed thrust. The fault trace is probably located at the northward bend in Rt. 33, just west of the roadcut. This is an example of intermediate scale (second-order) deformation structures along the Day 2 traverse. Stop 2-21. Roadcuts on East Slope of North Fork Mountain (Contact between Juniata and Tuscarora Formations) Mile 47.8 (N38° 41′ 54″, W79° 24′ 18″) Juniata red beds (south end of outcrop) transition into clean quartz arenite of the Tuscarora Formation. This contact is considered the Ordovician–Silurian boundary. Vertical Skolithos burrows occur near the base of the Tuscarora Formation, followed by massive cross-beds, followed by more Skolithos burrows. An interval of thin sandstone and dark-gray shale beds contain Arthophycus (bedding plane trails). Dip is moderately to the east. Follow Rt. 33 west over the crest of the North Fork Mountain ridge (Tuscarora Formation). Stop 2-22 (Optional). Roadcuts on East Side of Route 33 (Juniata Formation) Mile 49.3 (N38° 42′ 32″, W79° 24′ 29″) Juniata Formation red beds of alternating shale and sandstone dip shallowly to the east. The beds are cyclical, with cross-bedded, fine-grained sublitharenite at the base, followed by sublithic wacke with vertical burrows, and culminating with a bioturbated mudstone. These are interpreted as fining-upward tidal parasequences.

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Stop 2-23 (Optional). Roadcut on South Side of Route 33 (Contact between Reedsville Formation and Oswego Formation) Mile 49.5 (N38° 42′ 29″, W79° 24′ 39″) This outcrop shows the gradational transition between the upper Reedsville Formation, a bioturbated, fossiliferous mudstone, and the Oswego sandstone. The Oswego is an immature sandstone (sublitharenite) that represents the culmination of a marine coarsening-upward sequence that started at the base of the Reedsville. Stop 2-24. Germany Valley Overlook on North Side of Route 33 Mile 49.7 (N38° 42′ 29″, W79° 24′ 47″) The view to the north shows a several kilometer-scale breached anticline (Wills Mountain anticlinorium), with Germany Valley visible below. Ridges to the east and west are Silurian Tuscarora Formation, and the valley is composed of Ordovician carbonates. Closure of the anticline is visible along the horizon to the north where the ridges connect. This is representative of the largest scale (first-order) deformation structures that we see during the Day 2 traverse. At this locality we can see the whole structure from one overlook. Roadcuts on South Side of Route 33 (Middle Reedsville Formation) Shale and calcareous siltstone that represent mid-shelf deposition; includes thin beds of bioclastic calcarenite. Stop 2-25 (Optional). Roadcut on South Side of Route 33 (Trenton Formation) Mile 50.4 (N38° 42′ 22″, W79° 25′ 21″) Subhorizontal, thin hummocky sequences (distal shelf) with zones of fossil hash. Stop 2-26. Judy Gap (Tuscarora and Juniata Formations) Mile 53.6 (N38° 42′ 22″, W79° 27′ 50″) Parallel bands of vertical cliffs on both sides of Rt. 33 consist of clean Tuscarora sandstone and red beds of the Juniata Formation, apparently duplicated. Contraction fault visible in streambed on the north side of Rt. 33. Vertical beds of the Tonoloway Formation are found to the west in a quarry on the south side of Rt. 33. Turn right at intersection to follow Rt. 33 to the north. Stop 2-27. Junction of Route 33 and Briery Gap Road (Oriskany Formation) Mile 54.7 (N38° 43′ 03″, W79° 27′ 35″)

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Vertical to slightly overturned (east-dipping) beds of Oriskany Formation sandstone; brachiopod molds apparent on bedding planes. Turn left (west) onto Briery Gap road (leave Rt. 33). Stop 2-28. Partway up Hill on Briery Gap Road (Greenland Gap and Hampshire Formations) Mile 55.7 (N38° 43′ 48″, W79° 28′ 04″) Park at pullout on north side of road. Just east of the pullout (downhill) is the contact between the Greenland Gap Formation and Hampshire Formation (mid-shelf parasequences). Note that the beds gradually dip more shallowly to the west as you drive uphill. Stop 2-29 (Optional). Just below T Intersection at the Top of Briery Gap Road (Pocono Formation) Mile 56.3 (N38° 44′ 01″, W79° 28′ 37″) Shallowly west-dipping, cross-bedded quartz conglomerate sandstone. This is essentially the Allegheny Front: the western boundary of the Valley and Ridge province and the western limit of visible Alleghanian deformation. End of field trip! Turn around, and head back to Baltimore. ACKNOWLEDGMENTS The authors thank Rick Diecchio and John Haynes for helpful reviews of this field guide, and all of the SST students that have “field tested” this approach to integrating stratigraphic, structural, and tectonic analyses in the field. REFERENCES CITED Aleinikoff, J.N., Zartman, R.E., Walter, M., Rankin, D.W., Lyttle, P.T., and Burton, W.C., 1995, U-Pb ages of metarhyolites of the Catoctin and Mount Rogers Formations, central and southern Appalachians: Evidence for two pulses of Iapetan rifting: American Journal of Science, v. 295, p. 428–454. Aleinikoff, J.N., Burton, W.C., Lyttle, P.T., Nelson, A.E., and Southworth, C.S., 2000, U-Pb geochronology of zircon and monazite from Mesoproterozoic granitic gneisses of the northern Blue Ridge, Virginia and Maryland: Precambrian Research, v. 99, no. 1, p. 113–146, doi: 10.1016/S0301-9268 (99)00056-X. Badger, R.L., and Sinha, A.K., 1988, Age and Sr isotopic signature of the Catoctin volcanic province—Implication for subcrustal mantle evolution: Geology, v. 16, p. 692–695, doi: 10.1130/0091-7613(1988)016<0692: AASISO>2.3.CO;2. Badger, R.L., and Sinha, A.K., 2004, Geochemical stratigraphy and petrogenesis of the Catoctin volcanic province, central Appalachians, in Tollo, R.P., Corriveau, L., McLelland, J., and Bartholomew, M.J., eds., Proterozoic Tectonic Evolution off the Grenville Orogen in Eastern North America: Geological Society of America Memoir 197, p. 435–458. Bailey, C.M., Southworth, S., and Tollo, R.P., 2006, Tectonic history of the Blue Ridge, north-central Virginia, in Pazzaglia, F.J., ed., Excursions in Geology and History: Field Trips in the Middle Atlantic States: Geological Society of America Field Guide 8, p. 113–134, doi: 10.1130/2006 .fld008(07). Bailey, C.M., Peters, S.M., Morton, J., and Shotwell, N.L., 2007, Structural geometry and tectonic significance of the Neoproterozoic Mechum River Formation, Virginia Blue Ridge: American Journal of Science, v. 307, no. 1, p. 1–22, doi: 10.2475/01.2007.01.

Chamberlin, T.C., 1890, The method of multiple working hypotheses: Science, v. 15, no. 366, p. 92–96, doi: 10.1126/science.ns-15.366.92. Cooper, B.N., and Cooper, G.A., 1946, Lower middle Ordovician stratigraphy of the Shenandoah Valley, Virginia: Geological Society of America Bulletin, v. 57, p. 35–114, doi: 10.1130/0016-7606(1946)57[35:LMOSOT] 2.0.CO;2. Dennison, J.M., 1970, Stratigraphic divisions of Upper Devonian Greenland Gap Group (“Chemung Formation”) along Allegheny Front in West Virginia, Maryland and Highland County, Virginia: Southeastern Geology, v. 12, p. 53–82. Dennison, J.M., and Head, J.W., 1975, Sea level variations interpreted from the Appalachian Basin Silurian and Devonian: American Journal of Science, v. 275, p. 1089–1120. Diecchio, R.J., 1986, Taconian clastic sequence and general geology in the vicinity of the Allegheny Front in Pendleton County, West Virginia: Geological Society of America Centennial Field Guide–Southeastern Section, p. 85–90. Diecchio, R.J., 1993, Stratigraphic interpretation of the Ordovician of the Appalachian basin and implications for Taconian flexural modeling: Tectonics, v. 12, p. 1410–1419, doi: 10.1029/93TC01791. Drake, A.A., Jr., Sinha, A.K., Laird, J., and Guy, R.E., 1989, The Taconic orogen, in Hatcher, R.D., Jr., Thomas, W.A., and Viele, G.W., eds., The Appalachian-Ouachita orogen in the United States: Geological Society of America, The Geology of North America, v. F-2, p. 101–177. Dunne, W.M., 1996, The role of macroscale thrusts in the deformation of the Alleghanian roof sequence in the central Appalachians: A re-evaluation: American Journal of Science, v. 296, p. 549–575. Evans, M.A., 1989, The structural geometry and evolution of foreland thrust systems, northern Virginia: Geological Society of America Bulletin, v. 101, p. 339–354, doi: 10.1130/0016-7606(1989)101<0339:TSGAEO> 2.3.CO;2. Fichter, L.S., 1986, The Catskill clastic wedge (Acadian orogeny) in eastern West Virginia: Geological Society of America Centennial Field Guide– Southeastern Section, p. 91–96. Fichter, L.S., and Diecchio, R.J., 1986, Stratigraphic model for timing the opening of the proto-Atlantic Ocean in Northern Virginia: Geology, v. 14, p. 307–309, doi: 10.1130/0091-7613(1986)14<307:SMFTTO>2.0.CO;2. Fichter, L.S., and Diecchio, R.J., 1993, Evidence for the progressive closure of the proto-Atlantic ocean in the Valley and Ridge province of northern Virginia and eastern West Virginia: National Association of Geology Teachers Field Trip Guidebook, p. 27–49. Fichter, L.S., and Poché, D.J., 2001, Ancient environments and the interpretation of geologic history, Third edition: Upper Saddle River, Prentice-Hall, 310 p. Gathright, T.M., II, 1976, Geology of the Shenandoah National Park, Virginia: Virginia Division of Mineral Resources Bulletin 86, 93 p. Jonas, A.I., and Stose, G.W., 1939, Age relation of the Precambrian rocks in the Catoctin Mountain–Blue Ridge and Mount Rogers anticlinoria in Virginia: American Journal of Science, v. 237, no. 8, p. 575–593. King, P.B., 1950, Geology of the Elkton area, Virginia: U.S. Geological Survey Professional Paper 230, 82 p. Kulander, B.R., and Dean, S.L., 1986, Structure and tectonics of central and southern Appalachian Valley and Ridge and Plateau provinces, West Virginia and Virginia: AAPG Bulletin, v. 70, p. 1674–1684. Kunk, M.J., Froelich, A.J., and Gottfried, D., 1992, Timing of emplacement of diabase dikes and sheets in the Culpeper Basin and vicinity, Virginia and Maryland: Ar/Ar age spectrum results from hornblende and K-feldspar in granophyres: Geological Society of America Abstracts with Programs, v. 25, no. 2, p. 31. Middleton, G.V., 1973, Johannes Walther’s Law of the Correlation of Facies: Geological Society of America Bulletin, v. 84, p. 979–988, doi: 10.1130/0016-7606(1973)84<979:JWLOTC>2.0.CO;2. Mitra, S., 1986, Duplex structures and imbricate thrust systems: Geometry, structural position, and hydrocarbon potential: AAPG Bulletin, v. 70, p. 1087–1112. Perry, W.J., Jr., and DeWitt, W., Jr., 1977, A field guide to thin-skinned tectonics in the central Appalachians: Washington, D.C., American Association of Petroleum Geologists Annual Convention, Field Trip 4, 54 p. Pumpelly, R., Wolff, J.E., and Dale, T.N., 1894, Geology of the Green Mountains in Massachusetts: Washington, D.C., U.S. Geological Survey Monograph 23, 206 p. Rader, E.K., and Biggs, T.H., 1975, Geology of the Front Royal quadrangle, Virginia: Virginia Division of Mineral Resources Report of Investigation 40, 91 p.

Blue Ridge and Valley and Ridge Provinces of northern Virginia and West Virginia Rader, E.K., and Biggs, T.H., 1976, Geology of the Strasburg and Toms Brook quadrangles, Virginia: Virginia Division of Mineral Resources Report of Investigation 45, 104 p. Rader, E.K., and Perry, W.J., Jr., 1976, Stratigraphy as key to arch related origin of Little North Mountain structural front, Virginia and West Virginia: AAPG Bulletin, v. 60, p. 1623. Rogers, J., 1970, The tectonics of the Appalachians: New York, John Wiley and Sons, 271 p. Simpson, E.L., and Eriksson, K.A., 1989, Sedimentology of the Unicoi Formation in southern and central Virginia: Evidence for Late Proterozoic to Early Cambrian rift-to-passive margin transition: Geological Society of America Bulletin, v. 101, p. 42–54, doi: 10.1130/0016-7606(1989)101<0042: SOTUFI>2.3.CO;2. Smosna, R.A., Patchen, D.G., Washauer, S.M., and Perry, W.J., Jr., 1977, Relationships between depositional environments, Tonoloway Limestone and distribution of evaporites in the Salina Formation, West Virginia, in Fisher, J.H., ed., Reefs and Evaporites—Concepts and Depositional Models: American Association of Petroleum Geologists Studies in Geology 5, p. 125–143. Southworth, S., Aleinikoff, J.N., Bailey, C.M., Burton, W.C., Crider, E.A., Hackley, P.C., Smoot, J.P., and Tollo, R.P., 2009, Geologic map of the Shenandoah National Park region Virginia: U.S. Geological Survey OpenFile Report 2009-1153, 96 p., 1 plate, scale 1:100,000. Spencer, E.W., Bell, J.D., and Kozak, S.J., 1989, Valley and Ridge and Blue Ridge traverse, central Virginia: Washington, D. C., 28th International Geologic Conference, Field Guide for Field Trip 157, 69 p. Thomas, W.A., 1977, Evolution of Appalachian-Ouachita salients and recesses from reentrants and promontories in the continental margin: American Journal of Science, v. 277, p. 1233–1278.

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Tollo, R.P., and Aleinikoff, J.N., 1996, Petrology and U-Pb geochronology of the Robertson River Igneous Suite, Blue Ridge province, Virginia: Evidence for multistage magmatism associated with an early episode of Laurentian rifting: American Journal of Science, v. 296, p. 1045–1090. Tollo, R., Aleinikoff, J., Borduas, E., and Hackley, P., 2004, Petrologic and geochronologic evolution of the Grenville orogen, northern Blue Ridge province, Virginia, in Tollo, R.P., Corriveau, L., McLelland, J., and Bartholomew, M.J., eds., Proterozoic Tectonic Evolution off the Grenville Orogen in Eastern North America: Geological Society of America Memoir 197, p. 647–695. Virginia Division of Mineral Resources (VDMR), 2003, Publication 174: Digital Representation of the 1993 Geologic Map of Virginia, CD ROM (ISO-9660). West Virginia Geological and Economic Survey (WVGES), 1998, Digital Geologic Map of West Virginia: ftp://ftp.wvgis.wvu.edu/pub/Clearinghouse/geoscience/ geologicalMapOfWV_WVGES_1986_utm83_coverage_shp.zip. Wilson, T.H., and Shumaker, R.C., 1992, Broad Top thrust sheet: An extensive blind thrust in the central Appalachians: AAPG Bulletin, v. 76, p. 1310–1324. Woodward, H.P., 1943, Devonian system of West Virginia: West Virginia Geological Survey, v. 15, 655 p. Woodward, H.P., 1951, Ordovician system of West Virginia: West Virginia Geological Survey, v. 21, 627 p.

MANUSCRIPT ACCEPTED BY THE SOCIETY 2 DECEMBER 2009

Printed in the USA

Contents Preface 1. The Peach Bottom area in the Pennsylvania-Maryland Piedmont R. T. Fail! and R. C. Smith II

2. Soils, geomorphology, landscape evolution, and land use in the Virginia Piedmont and Blue Ridge WC. Sherwood, A.S. Hartsho rn, and L.S. Eaton

3. Magmatic layering and intrusive plumbing in the Jurassic Morgantown Sheet, Central Atlantic Magmatic Province L. Srogi, T. Lutz, L.D. Dickson, M. Pollock, K. Gimson, and N. Lynde

4. The early through late Pleistocene record in the Susquehanna River Basin D.D. Braun

5. Stratigraphy, structure, and tectonics: An east-to-west transect of the Blue Ridge and Valley and Ridge provinces of northern Virginia and West Virginia L.S. Fichter, S.J. Whitmeyer, C.M. Bailey, and W. Burton

. . THE GEOLOGICAL SOCIETY . O F AMERICA 3300 Penrose Place • P.O. Box 9140 Boulder, Colorado 80301-9140, USA

ISBN 978 - 0 - 8137-0016-8