Geology of the Shenandoah National Park Region - Csmres Jmu ...

39 th Annual Virginia Geological Field Conference 

October 2 nd - 3 rd , 2009 

Geology of the 

Shenandoah National Park Region 

Scott Southworth 

U. S. Geological Survey 

L. Scott Eaton 

James Madison University 

Meghan H. Lamoreaux 

College of William & Mary 

William C. Burton 


1 

Christopher M. Bailey 


Gregory Hancock 


Ronald J. Litwin 


Jennifer Whitten 

College of William & Mary

2009 

VGFC 

81 

38˚ 15’ N 

• 

Trayfoot 

Mtn. 

3 • 

78˚ 45’ W 

Blue 

Ridge 

Massanutten Mountain 

33 

South Fork Shenandoah River 

• 4 

Skyline Drive 

Pasture Fence Mtn. 

Roanoke 

Valley & Ridge 

340 

Rocky 

Mtn. • 

5 • 

• 

Big Flat 

Mtn. 

Washington 

VA 

Piedmont D.C. 

• Loft Mtn. 

• 

Fox Mtn. 

New 

Market Gap 

• 

Elkton 

340 

Shenandoah 

• 

Hanse 

Mtn. 

• Flattop 

Mtn. 

Swift Run • 

Gap 

2 • 

• 6 

Field Trip 

Stop 

• Hightop 

• Brokenback 

Mtn. 

• 7 

• 

Grindstone 

Mtn. 

Stanley 

Dean 

Mtn. • 

78˚ 30’ W 

N 

211 

Long Ridge 

• Lewis 

Mtn. 

Skyline Drive 

• Saddleback 

Mtn. 

Luray 

Hershberger 

Hill 

Big 

Meadows 

•7 

• Hazeltop 

Hawksbill • 

• 1 

Kirtley 

Mtn. 

Stony Man • 

33 

230 

Stanardsville 

Ruckersville 

Doubletop Mtn. 

Marys Rock • 

29 

0 5 10 

Thornton 

Gap 

• 

Old Rag 

Mtn. 

38˚ 30’ N 

Syria 

0 

kilometers 

5 10 

Figure 1. Shaded relief map of the Shenandoah National Park region and stop locations at the 2009 Virginia 

Geological Field Conference. 

2 

miles

Dedication 

The 39 th annual Virginia Geological Field Conference is dedicated to Tom Gathright whose work in western Virginia 

and Shenandoah National Park has educated and inspired many of us. 

Tom Gathright joined the Division of Mineral Resources in 1964. Originally a member of the Ground water 

Section, Tom conducted groundwater investigations in the Blue Ridge and Valley & Ridge provinces. During 

this time, he helped to locate and develop several major water supplies along the west flank of the Blue Ridge. 

Tom’s first project with the Geologic Mapping Section was to create a geologic map of Shenandoah National 

Park. This work was published in 1976 as Bulletin 86, “Geology of the Shenandoah National Park, Virginia.” 

It remains one of the most popular reports related to Virginia geology. Tom subsequently completed geologic 

maps of ten additional 7.5-minute quadrangles and helped to compile three geologic maps of three 30- x 

60-minute quadrangles. As the head of the geologic mapping section, he oversaw the completion of geologic 

mapping of the southwest Virginia Coalfields. Tom retired in 1991, but continues to work on an occasional 

basis as a geologic consultant and serves as a member of Virginia’s geologic mapping advisory committee. 

Tom framing a photograph on a crisp morning at Rockfish Gap. 

photos courtesy of Elizabeth Campbell (top) and Gerry Wilkes (bottom). 

3

Introduction 

Shenandoah National Park and its magnificent Skyline Drive lie astride the Blue Ridge Mountains in northcentral 

Virginia. Occupying more than 800 km 2 , Shenandoah is a long and narrow park with a highly irregular 

boundary. Established during the Great Depression, the Park was intended to serve as a leafy refuge for 

urbanites in eastern North America. Created from a patchwork of privately owned farms, orchards, and home 

sites, the Park has been reforested in the past 80 years with much of the region returned to a wilderness 

state. In modern times, the Park welcomes over two million visitors per year. 

The Blue Ridge Mountains in Shenandoah National Park form a distinctive highland that rises to elevations 

above 1,200 meters (~4,000’) with local topographic relief exceeding 900 meters (~3,000’). The crest of the 

range forms a drainage divide separating the Shenandoah (Potomac) River drainage to the west, in the Great 

Valley subprovince of the Valley & Ridge, from southeast-flowing streams of the James and Rappahannock 

river systems coursing into the foothills region of the Piedmont physiographic province (Fig. 1). 

Although a cloak of forest and mantle of soil commonly obscures the underlying bedrock, Shenandoah’s 

topography is dictated by its geology. Rocks exposed in the Blue Ridge Mountains are among the oldest 

in Virginia and bear witness to more than a billion years of magmatism, sedimentation, sea level oscillation, 

climate change, tectonic activity, and erosion. Bedrock includes a suite of Grenvillian basement rocks, 

metamorphosed Neoproterozoic sedimentary and volcanic rocks, and early Cambrian siliciclastic rocks. 

The Park is situated along the western margin of the Blue Ridge anticlinorium, a regional-scale Paleozoic 

structure developed at the hinterland edge of the Appalachian fold and thrust belt. The Blue Ridge highlands 

are the product of differential erosion in the Cenozoic, but post-Paleozoic tectonic activity has influenced 

the character of the Blue Ridge landscape in discernible ways. Quaternary surficial deposits are common 

throughout the Park, providing a record of past climate regimes as well as active processes shaping the 

modern landscape. 

Aspects of the Park’s geology have been studied since the 1930’s (Furcron, 1934; Jonas and Stose, 1939; 

King, 1950) and many of the counties that encompass the Park were mapped after World War II (Rockingham- 

Brent, 1960; Albemarle- Nelson, 1962; Greene and Madison- Allen, 1963; Page- Allen, 1967). The first 

comprehensive treatment of Shenandoah’s geology was undertaken by the Virginia Division of Mineral 

Resources in the late 1960s and culminated in the 1976 publication and geologic map (1:62,500 scale) by 

Tom Gathright. Gathright’s seminal work summarized many of the earlier studies and provided a unifying 

framework of the Park’s geology that is still informative in the 21 st century. The hydrogeologic setting of the 

Park is discussed by DeKay (1972). 

Robert Badger’s Roadside Geology of the Skyline Drive (1999) is a well-illustrated guide aimed at a broad 

audience. Tollo and others (2004), Eaton and others (2004), and Bailey and others (2006) provide technical 

field guides with stops in and near the Park. Web resources illustrating the geology of Shenandoah National 

Park include: William & Mary’s Geology of Virginia/Google Earth site (http://web.wm.edu/geology/virginia/ 

ge.php), James Madison’s Geology of Virginia site, (http://csmres.jmu.edu/geollab/vageol/vahist/) and Callan 

Bentley’s (Northern Virginia Community College) Geology of Shenandoah National Park site (http://www. 

nvcc.edu/home/cbentley/gol_135/shenandoah/index.htm). 

In the past fifteen years, a multitude of studies focusing on the geology of Shenandoah National Park have 

contributed new insights about 1) the complexity and chronology of the basement, 2) the structural geometry 

and deformation history of the region, and 3) Quaternary surficial processes and their efficacy. Nineteen 7.5’ 

quadrangles in and around Shenandoah National Park have been mapped (or remapped) at 1:24,000 scale 

in the last decade (Fig. 2). A new 1:100,000 scale geologic map of the Shenandoah National Park region and 

an accompanying report, compiling much of this research, was just published by the U. S. Geological Survey 

(Southworth and others, 2009; online at http://pubs.er.usgs.gov/usgspubs/ofr/ofr20091153). 

4

The 2009 Virginia Geological Field Conference will highlight new research, visiting exposures along and 

near the Skyline Drive. The trip will be broad, encompassing bedrock and structural geology as well as 

surficial geology and rates of erosion at local and regional scales. Shenandoah National Park is too vast for 

a single day traverse; thus the Conference will focus on the geology between Big Meadows (milepost 51) and 

Blackrock Summit (milepost 85) in the central and southern part of the Park. In addition to emphasizing new 

findings in Shenandoah we hope this trip will foster discussion about unanswered aspects about Blue Ridge 

geology and encourage future research in the Park. 

Weyers Cave 

38˚00’ N 

79˚00’ W 

Harrisonburg 

Gathright 

and others, 

1986 

Fort 

Defiance 

Gathright 


1978 

Waynesboro 

West 

Gathright 


1977 

38˚30’ N 

Grottoes 

Gathright 


1978 

Crimora 

Gathright 


1978 

Waynesboro 

East 

Gathright 


1977 

Tenth 

Legion 

Whitmeyer 

and Heller 

Elkton West 

Heller 

McGaheysville 

Browns 

Cove 

Lamoreaux 


2009 

Crozet 

Lederer 


2009 

Edinburg 

Hamburg 

Stanley 

Whitmeyer 

and others 

Elkton East Fletcher 

Swift Run 

Gap 

Bailey 


2009 

Free 

Union 

Southworth 

and Bailey 

Charlottesville 

West 

Big Meadows 

Whitmeyer 

and others 

Tollo and 

others 

Tollo and 

others, 2004 

Stanardsville 

Burton 

and Bailey, 

2009 

Earlysville 

Figure 2. Index map of 7.5 minute quadrangles surrounding Shenandoah National Park. Authors noted inside 

quadrangles mapped at 1:24,000 scale. Grey shading indicates recent work. 

5 

Strasburg 

Rader and 

Biggs, 1976 

Rileyville Bentonville 

Southworth 

and 

Tollo, 2006 

Luray Thornton 

Gap 

Tollo and 


Old Rag 

Mountain 

Tollo and 


Madison 

Bailey and 


Rochelle 

Bailey 

Front Royal 

Rader and 

Biggs, 1975 

Washington 

Tollo and 


Washington 

Tollo and 


Woodville 

Tollo and 

Bailey 

Brightwood 

Explanation 

Linden 

Lukert and 

Nichols, 

1976 

Flint Hill 

Lukert and 

Nichols, 

1976 

Massies 

Corner 

Lukert and 

Halladay, 

1980 

Shenandoah National Park 

Boundary 

0 1.5 3 6 9 12 Miles 

78˚00’ W 

39˚00’ N 

Castleton 

Tollo and Lowe, 1994 

Bailey and others, 2007

Bedrock Geology 

Stratigraphy 

Shenandoah National Park is underlain by three major geologic units: 1) Mesoproterozoic gneisses and 

granitoids that comprise the basement, 2) Neoproterozoic metasedimentary and metavolcanic rocks of the 

Swift Run and Catoctin formations, and 3) siliciclastic rocks of the Early Cambrian Chilhowee Group (Figs. 

3 and 4). Basement rocks crop out in the eastern Blue Ridge Mountains and the adjoining foothills, but also 

underlie peaks such as Old Rag Mountain, Mary’s Rock, and Roundtop. The Catoctin Formation forms much 

of the Blue Ridge’s high crest including Hawksbill, Stony Man, Mount Marshall, and Hightop. The Chilhowee 

Group, exposed in the western Blue Ridge, underlies steep mountains and ridges mantled with thin soil and 

abundant talus such as Grindstone, Rocky, Trayfoot, and Turk mountains. 

Basement rocks include granitoid gneisses and granitoids formed during the Grenville orogeny between 1.2 

and 1.0 Ga (Fig. 3). Prior to the 1980s rock units in the Blue Ridge basement were commonly mapped as 

formations and Gathright (1976) mapped two basement units, the Pedlar Formation and the Old Rag Granite, 

in the Park. Most technical studies published in the past two decades (Rader and Evans, 1993; Southworth 

and others, 2000; Bailey and others, 2003; Tollo and others, 2004b; Southworth and others, 2009) avoid 

using the formation terminology; rather Mesoproterozoic basement units are distinguished based on rock 

type, cross cutting relations, geochemistry, and geochronology. Many general interest publications and 

websites continue to use the formational lexicon, this is unfortunate because the rich geologic history of the 

basement is diluted (or worse still, misunderstood) using archaic formation names. For instance, recent 

mapping and geochronology in Shenandoah National Park demonstrates that the Pedlar Formation contains 

over 12 different units that that vary in age by ~150 million years (Southworth and others, 2009). 

Modern U-Pb zircon geochronology reveals that the basement complex is comprised of three temporally 

distinct groups of igneous rocks emplaced over a 150 million year interval in the Mesoproterozoic (Fig. 

3) (Aleinikoff and others, 2000; Tollo and others, 2004b; Southworth and others, 2009). The oldest group 

crystallized between 1,190 and 1,150 Ma and are granitoid gneisses with a compositional layering that 

developed under high-grade conditions. A volumetrically minor group of orthopyroxene-bearing granites 

crystallized between 1,120 and 1,110 Ma. The youngest group of granitoids was emplaced between 1,090 

and 1,020 Ma. Both the older and younger groups are chemically diverse and include pyroxene-bearing 

charnockitic rocks and alkali feldspar leucogranitoids. Tollo and others (2004) note that the basement suite 

was derived from melting of lower crustal sources in an intraplate setting. 

The Blue Ridge basement is almost entirely of igneous origin (orthogneisses), however recent mapping has 

identified small (

Recent mapping in the northeastern part of the Park reveals thin, discontinuous layers of felsic volcanic rock 

unconformably overlie the Mesoproterozoic basement and beneath the Catoctin Formation that yield U-Pb 

zircon ages of 720 to 710 Ma (Southworth and others, 2009). 

A late Neoproterozoic cover sequence of metasedimentary and metavolcanic rocks unconformably overlie 

the basement complex in the Shenandoah National Park region (Fig. 3). The Swift Run Formation is a 

heterogeneous clastic unit of highly variable thickness (absent to ~300 m) that crops out below metabasalts 

of the Catoctin Formation and to the east of the Park in outliers surrounded by basement (Gattuso and 

others, 2009). Common rock types include arkosic phyllite, meta-arkose, phyllite, laminated metasiltstone, 

and pebble to cobble metaconglomerate. Locally, compositionally mature, cross-bedded, quartz-rich 

metasandstone crops out. Gathright (1976) and Schwab (1986) interpret the Swift Run Formation to be a 

non-marine unit deposited in alluvial fan, floodplain, and lacustrine environments. A number of early workers 

report tuffaceous rocks from the Swift Run Formation, but recent research indicates these are primarily 

fine-grained metasedimentary rocks. Contemporaneous normal faulting likely influenced the deposition of 

Swift Run sediments in the outliers (Forte and others, 2005; Gattuso and others, 2009). At a number of 

locations clastic rocks are interlayered with metabasaltic greenstone, a geometry consistent with a coeval 

relationship between the Swift Run Formation and the lower Catoctin Formation (King, 1950; Gattuso and 

others, 2009). 

The Catoctin Formation forms an extensive unit in Shenandoah National Park and is characterized by 

metabasaltic greenstone with thin layers of meta-arkose, phyllite, and epiclastic breccia. Catoctin basalts 

were extruded over a large region (>4000 km 2 ) and generated from mantle-derived tholeiitic magmas (Badger 

and Sinha, 2004). In the Park, basalts extruded primarily as subaerial flows, evidenced by abundant columnar 

joints and flow-top breccias (Reed, 1955, 1969). In the central and northern part of the Park, nine to sixteen 

individual flows occur (Reed, 1955; Gathright, 1976; Badger, 1992). The Catoctin Formation is upwards 

of 700 meters thick in the Park and thins towards the west and southwest. Metadiabase dikes of similar 

composition to Catoctin metabasalts intrude the basement complex (as well as older Neoproterozoic rocks) 

and are likely feeder dikes for the overlying Catoctin lava flows. Dikes range from 0.5 to 5 m in thickness and 

are most common in the northern and central part of the Park. 

Badger and Sinha (1988) report a Rb-Sr isochron age of 570±36 Ma from exposures just south of Shenandoah 

National Park. Zircons from metarhyolite tuffs and dikes in the Catoctin Formation, exposed to the north 

of the Park, yield U-Pb ages between 570 and 550 Ma (Aleinikoff and others, 1995). Collectively, the 

geochronologic data suggest that Catoctin volcanism lasted between 15 and 20 million years and occurred 

during the Ediacaran period in the late Neoproterozoic. Paleomagnetic data from the Catoctin Formation 

are complex, but broadly compatible with a high southerly latitude (60˚ S) for the Virginia Blue Ridge during 

extrusion (Meert and others, 1994). 

The siliciclastic Chilhowee Group overlies the Catoctin Formation. In Shenandoah National Park, the 

Chilhowee Group includes the Weverton, Harpers, and Antietam formations and ranges from 500 to 800 

meters in aggregate thickness (Fig. 3). Gathright’s (1976) nomenclature for the Chilhowee Group includes 

the Weverton, Hampton (Harpers), and Erwin (Antietam) formations. Weverton, Harpers, and Antietam are 

units whose type locations occur along the Potomac River approximately 125 km to the northeast, whereas 

the Hampton and Erwin type locations are located northeastern Tennessee. The contact 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 and others, 

2007a). 

The Weverton Formation includes quartz metasandstone, granule to pebbly metaconglomerate, laminated 

metasiltstone, and quartzose phyllite. The Harpers Formation, in Shenandoah National Park, is dominated 

by green to gray phyllite and thinly bedded metasandstone. Less common rock types include well-cemented 

quartz arenite and ferrigunous metasandstone. Trace fossils (Skolithos and burrowed beds) are rare, but 

7

do occur in the upper part of the Harpers Formation. The Antietam Formation consists of distinctive wellcemented 

quartz arenite, with abundant Skolithos, and laminated metasiltstone. Collectively, the Chilhowee 

Group records a fluvial to shallow-marine transgressive sequence (2 nd -supersequence) (Simpson and 

Eriksson, 1990; Read and Eriksson, in press). Skolithos is the only well documented fossil in the Chilhowee 

Group from the Shenandoah National Park region, but along strike to the southwest, Rusophycus occurs 

near the base and Olenellus trilobites near the top, bracketing the age of the Chilhowee Group between the 

earliest Cambrian and early Middle Cambrian (

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. 

Mylonitic rocks occur in anastomosing zones, up to a kilometer thick, that cut basement rocks along the eastern 

margin of the Park (Figs. 4 and 5). Mylonite zones in the Park were first recognized by Gathright (1976) and 

aspects of these rocks have been discussed by Mitra (1977), Bailey and Simpson (1993), and Bailey and 

others (2002). Asymmetric structures consistently record a top-to-the-northwest sense of shear (i.e. hanging 

wall up movement). The reverse displacement across Blue Ridge high-strain zones accommodated crustal 

contraction enabling the relatively stiff basement complex to shorten while cover rocks were folded. 

A distinctive coarse foliation or compositional banding is developed in the older Mesoproterozoic basement 

units (Groups 1 and 2, >1,150 Ma) and defined by aligned feldspars and quartz aggregates that formed at 

upper amphibolite- to granulite-facies conditions. Younger Mesoproterozoic basement units (Group 3, 80°). Transverse faults cut previously folded contacts and 

the apparent strike-slip offset in map view is accomplished by dip-slip movement with maximum displacements 

of ~100 m. A number of transverse faults, including the recently recognized Simmons Gap and Sandy Bottom 

faults in Shenandoah National Park, correspond to well-developed topographic lineaments (Fig. 4; Bailey 

and others, 2009). 

North-northwest striking transverse faults and extension fractures are subparallel to a regional suite of Jurassic 

diabase dikes. Bailey and others (2006) interpreted many transverse faults in Shenandoah National Park 

as Jurassic structures noting that these faults record minor east-northeast directed extension and possibly 

developed in a transtensional stress regime. Wieczorek and others (2004) used LIDAR imagery to identify a 

9

topographic lineament, which they interpret as a north-northwest-striking fault (Harriston fault) near Grottoes 

and suggested this structure could have experienced Quaternary slip. Further research is required to evaluate 

the possibility of neotectonic activity along transverse faults. Many transverse faults are zones of structural 

weakness and strongly influenced the topographic character of the Blue Ridge Mountains; Simmons, Powell, 

Smith Roach, and Swift Run gaps are all located on transverse faults. 

Tectonic History 

Rocks and structures exposed in Shenandoah National Park formed in response to tectonic processes 

associated with the opening of ocean basins and the collision of continents throughout the past 1.2 billion 

years. Major tectonic events include the Mesoproterozoic Grenvillian orogen, Neoproterozoic Iapetan rifting, 

multiple episodes of Paleozoic collision, and Mesozoic crustal extension (Fig. 6). 

Mesoproterozoic granitoids and gneisses formed in the middle and lower crust along the margin of Laurentia 

during the long-lived Grenvillian orogeny that culminated in the amalgamation of the Rodinian supercontinent 

by 1,000 Ma. Older basement granitoids were emplaced during a magmatic interval 30 to 40 My in duration 

(1,180 – 1,140 Ma). A high-grade deformation event transformed these rocs into granitoid gneisses prior 

to the intrusion of the youngest plutonic suite by 1,080 Ma. The older deformation event is characterized by 

many northwest to east-west trending structures, but the kinematics of this event remain poorly understood. 

Lead isotopes in central and southern Appalachian basement massifs are distinctly different from other 

Laurentian Grenvillian provinces implying that Blue Ridge basement is not “native crust”, but rather was 

accreted to Laurentia during this long-lived orogeny. 

The small exposures of post-1,000 Ma paragneiss require uplift and erosion of the Blue Ridge basement 

followed by burial, metamorphism, and exhumation before Rodinia broke apart. From 730 to 700 Ma, 

plutonic and metavolcanic rocks of the Robertson River Igneous Suite were emplaced during an episode of 

crustal extensional. Although this tectonic event did not result in the formation of an ocean basin, it created 

significant accommodation space into which the thick sediments of the Fauquier, Lynchburg, and Mechum 

River sequence were deposited. The evidence of this first episode of Neoproterozoic crustal extension and 

rifting is preserved primarily to the east of Shenandoah National Park. 

Deposition of Swift Run sediments, emplacement of mafic dikes, and extrusion of Catoctin volcanics formed 

during a rift event that began ~570 Ma. Immediately southeast of the Park (in Greene and Albemarle 

counties), syn-sedimentary normal faults are associated with outliers of Swift Run and Catoctin formations. 

The regionally extensive Catoctin Formation may have been generated from plume-related processes that 

drove rifting along this segment of the Laurentian margin. The non-marine (Weverton) to nearshore-marine 

(Harpers and Antietam) rocks of the Chilhowee Group record the rift-to-drift transition as a passive continental 

margin developed along the Laurentian margin of the Iapetus Ocean in the Early Cambrian. Rocks in the 

Blue Ridge were progressively buried under a sequence of dominantly marine Cambrian and Ordovician 

sedimentary rocks (>8 km thick) deposited on the Laurentia continental margin. 

By the middle to late Ordovician, the Taconic orogeny had began in the Appalachians. Although numerous 

workers have argued for significant Taconic deformation in the Virginia Blue Ridge, recent geochronological 

studies provide little evidence of Ordovician activity. A late Ordovician to Silurian clastic wedge was deposited 

to the northwest of the modern Blue Ridge and significant subsidence continued into the middle Paleozoic 

(as preserved in Valley & Ridge sedimentary sequences). New Ar-Ar geochronology indicates that pervasive 

deformation in the basement and cover sequence occurred between ~360 and 310 Ma (late Devonian to 

Pennsylvanian). Deformation and metamorphism of this age falls between the classic Appalachian Devonian 

Acadian and Pennsylvanian-Permian Alleghanian orogenies. Zircon fission track data reveals the Blue 

Ridge had cooled through ~235˚ C by 300 Ma. The frontal Blue Ridge fault system slipped during the late 

Alleghanian (300 to 280 Ma) when the Blue Ridge massif was relatively cool. Evidence from the southern and 

central Appalachians suggests that middle to late Paleozoic tectonism was protracted and involved dextral 

10

PALEOZOIC MESOZOIC 

NEOPROTEROZOIC 

MESOPROT. 

AGE (Ma) 

180 

200 

220 

240 

260 

280 

300 

320 

340 

360 

380 

400 

420 

440 

460 

480 

500 

520 

540 

560 

580 

600 

change 

in scale 

650 

700 

750 

800 

change 

in scale 

900 

1000 

1100 

1200 

Stenian Tonian Cryogenian 

Ediacaran Cambrian Ordovician Silurian Devonian Carboniferous Permian Triassic Jurassic 

?? 

?? 

continued erosion 

Mafic magmatism 

Deposition of Culpeper Grp. 

exhumation and erosion 

brittle thrusting and emplacement of 

Blue Ridge sheet on Great Valley sheet 

contractional deformation, 

greenschist facies metamorphism 

high-strain zone movement, and 

foliation development in basement 

& cover sequence 

continued burial of Blue Ridge 

Continued Deposition 

in Valley and Ridge 

uplift, exhumation and erosion ?? 

continued burial of Blue Ridge 

Deposition of carbonates 

Deposition of Chilhowee Grp. 

Extrusion of Catoctin basalts 

Deposition of Swift Run Fm. 


Deposition of Mechum River Fm., 

Lynchburg Grp./Fauquier Fm. 


Intrusion of the Robertson 

River granitoids 


granulite facies metamorphism 

Intrusion of younger plutonic suite 

deformation and 

granulite facies metamorphism 

Intrusion of older plutonic suite 

11 

brittle 

normal 

faulting 

brittle 

normal 

faulting 

brittle/ 

ductile 

normal 

faulting 

?? 

Atlantic 

Rifting 

Alleghanian 

Orogeny 

Acadian 

Orogeny 

Taconian 

Orogeny 

Iapetan 

Rifting 

Crustal 

Extension 

Grenvillian 

Orogenesis 

Figure 6. Generalized geologic and tectonic history of the Blue Ridge (from Bailey and others, 2006).

contraction that transitioned to orthogonal convergence. The Alleghanian orogeny culminated with collision 

between North America and Africa, leading to the Pangean supercontinent by the close of the Paleozoic. 

In the early Mesozoic eastern North America and northwestern Africa began to rift apart, initiating a process 

that opened the Atlantic Ocean. In north-central Virginia, this event created the Culpeper basin in the eastern 

Blue Ridge and western Piedmont. Rifting generated basaltic magmas in the early Jurassic, and in the Blue 

Ridge to west-northwest-striking diabase dikes were intruded during a phase of modest east-west extension. 

Apatite fission-track ages reveal that rocks in the Shenandoah National Park region cooled below ~90˚ C 

during the Jurassic and Cretaceous indicating that Blue Ridge rocks have been near the surface for over 100 

million years. The topographic expression of the Blue Ridge Mountains developed during the late Cenozoic 

as differential erosion by Atlantic flowing rivers (Potomac, Rappahannock, and James) preferentially removed 

less resistant rocks in the Valley & Ridge and Piedmont. 

Surficial Geology 

Unconsolidated surficial materials and their resulting landforms cover large areas of the Shenandoah National 

Park region. Traditionally, geologists have underappreciated these deposits and their significance. An array 

of new studies focused on these deposits and surficial processes have yielded significant insights about 

the Blue Ridge landscape. Recent surficial investigations in the Shenandoah Park region include studies 

by Sherwood and others (1987), Whittecar (1992), Morgan and others (1999a,b), Litwin and others (2001, 

2003), Eaton and others (2003a,b), Smoot (2004), Morgan and others (2004), and Wieczorek and others 

(2004). Surficial materials contribute to soil character and play a key roll in land use, water resources, and 

hazards (sinkhole subsidence, debris flows, and flooding). Landforms and surficial deposits in and near the 

Park reflect processes active over a broad time spectrum and were formed under a wide range of climate 

conditions. 

Surficial deposits include those formed by (1) flowing water (alluvium, terrace deposits, alluvial-fan deposits, 

and alluvial-plain deposits), (2) gravity and high rainfall events on slopes (debris-flow and debris-fan 

deposits), (3) gravity and freeze-thaw processes on slopes (stratified slope deposits, colluvium, and blockfield 

deposits), and (4) chemical weathering (sinkholes and residuum) (Fig. 7). In general, alluvial-fan and 

alluvial-plain deposits are located on the lower slopes and valleys on the west and east sides of the Blue 

Ridge, respectively, and terrace deposits are located along major rivers in the valleys. Debris-fan deposits 

are concentrated in coves and hollows on the upper to lower slopes. Colluvium is concentrated in the 

Great Valley 

carbonate 

bedrock 

western 

Blue Ridge 

sinkhole 

S. Fork 

Shenandoah River 

block fields 

& talus 

deposits river terraces 

& alluvium 

alluvial fan 

siliciclastics 

upland 

alluvium 

periglacial 

deposits 

12 

Blue Ridge 

crest 

colluvium 

greenstone 

granitic 

bedrock 

Figure 7. Simplified model of surficial deposits in the Shenandoah National Park region. 

eastern 

Blue Ridge 

debris flow & 

debris flow fan 

alluvial plain

highlands. Debris-fan deposits are abundant in areas underlain by gneisses and metabasalts, whereas 

block-field deposits are most abundant in the areas underlain by siliciclastic rocks. Extensive alluvial-fan 

deposits cover carbonate bedrock immediately west of the Blue Ridge. The South Fork of the Shenandoah 

River has incised the alluvial fans to form terraces. In addition, sinkholes occur in the carbonate bedrock, 

including bedrock buried by alluvial fan deposits. In contrast, the lowlands east of the Blue Ridge are 

characterized by broad alluvial aprons along streams and by some terraces mantled with gravel. 

The age of surficial materials in the Shenandoah National Park region ranges from the late Pliocene to the 

Recent. Weathering characteristics of deposits in high terraces and fans are similar to late Pliocene deposits 

in the Fall Zone and Coastal Plain (Howard, 1993; Eaton and others, 2001). Debris fans have formed since 

at least the late Pleistocene, Eaton and others (2003a) obtained 14 C ages for organic matter in fans ranging 

from about 51,000 to 2,000 yr B.P. These old debris fan deposits were eroded and exposed by debris flows 

generated from a high rainfall event in 1995 that deposited much sediment on lower slopes and floodplains. 

Measurements of summit erosion rates in Shenandoah National Park using 10 Be 

Gregory Hancock and Jennifer Whitten, College of William & Mary 

Understanding the change in landscape morphology through time requires knowing erosion rates at various 

points within that landscape, particularly valley and summit erosion rates, and rates averaged over differing 

timescales. For instance, similar valley and summit erosion rates could suggest a state of dynamic equilibrium, 

wherein “all elements of the topography are mutually adjusted so that they are downwasting at the same 

rate” (Hack, 1960, p. 85). Hack (1960) proposed this concept from observations in the central and southern 

Appalachian landscape, and recent work in the Great Smoky Mountains suggests ridge crest, bare-bedrock 

erosion rates and longer-term exhumation rates determined from thermochronology are similar to basin 

average erosion rates, supporting a state of dynamic equilibrium in that landscape (Matmon and others, 

2003). 

While the concept of dynamic equilibrium is attractive given a lack of substantial orogenesis over the last 

~300 My in the Appalachians, disequilibrium may be produced by climatic effects on rates and patterns of 

erosion in mountainous topography. Peizhen and others (2001) suggest that the late Cenozoic transition to 

rapidly changing climate conditions 3–4 My ago led to a widespread acceleration of fluvial and glacial erosion 

rates, leading to relief changes as valley erosion rates increased relative to summit lowering rates in many 

mountainous regions (e.g, Molnar and England, 1990). Increasing relief has been documented in many other 

ranges during the late Cenozoic (e.g., Small and Anderson, 1995), which is consistent with disequilibrium 

induced by late Cenozoic climate change. Along the eastern North American margin, sedimentation rates 

doubled between the Late Miocene and the Quaternary (Poag and Sevon, 1989) and river incision rates 

across the region appear to have increased since the late Miocene (Mills, 2000). Hancock and Kirwan (2007) 

documented summit erosion rates on the eastern margin of the Appalachian Plateau in West Virginia that 

are well below measured fluvial erosion rates over a comparable period (~200 ky BP). This is consistent 

with increasing relief and landscape disequilibrium in at least in some areas of the central and southern 

Appalachians. In contrast to this proposed increased relief, other work in the central Appalachians calls 

for rapid summit erosion under periglacial climates, leading to an overall decrease in relief. Braun (1989) 

hypothesizes that periglaciation is the dominant erosional process in the central Appalachian uplands, is 

most efficient during glacial periods (and, hence, would be made possible as a consequence of late Cenozoic 

cooling), and occurs at rates greater than fluvial erosion rates. If true, the result of climate cooling in the 

central Appalachians has been the acceleration of upland erosion rates and a disequilibrium resulting in a 

decrease in relief during the Late Cenozoic. 

These three contrasting hypotheses pose the following question: are the central Appalachians in a state of 

dynamic equilibrium or not? We address this question by measuring rates of summit erosion at numerous 

locations within the park, and comparing these rates to previously determined basin-averaged and fluvial 

13

erosion rates. To determine erosion rates, we utilize the abundance of the cosmogenic radionuclide 10 Be 

accumulated on bare-bedrock summits. We focus on bare-bedrock summits as these are likely to be the 

most slowly eroding parts of the landscape, and therefore limit the overall lowering of elevation and relief 

within the range. To date, we have measured a total of 23 bare-bedrock erosion rates from nine summit 

locations located within four different lithologies in Shenandoah National Park (Fig. 8). 

Figure 8. Sampling locations within Shenandoah National Park. 

Blackrock Summit (Stop 3) is one of the sampling locations from which bare-bedrock erosion rates have 

been attained. The pale blue to purplish rocks here are well-cemented quartz sandstones and/or quartzites 

of the Lower Cambrian Harpers Formation (Gathright, 1976; Olney and others, 2007; Southworth and others, 

2009). Weathering of this rock produces well-defined fractures oriented along bedding planes. Lowering 

of this summit likely occurs through removal of blocks defined by these weathered sub- horizontal bedding 

planes coupled with vertical joints, as well as toppling of isolated tors and direct removal by rain and wind of 

granular material produced by weathering. Block sliding and tor toppling contribute blocks to an extensive 

downslope talus field comprised of regularly sized blocks (Gathright, 1976). The absence of vegetation, 

unstable blocks, and preferred block alignment may suggest recent and/or ongoing downslope creep of 

blocks in this talus field (Eaton, pers. comm.). 

To sample this site and others, we collected rock samples from nearly flat bedrock outcrops, and measured 

the abundance of the cosmogenic radionuclide (CRN) 10 Be in quartz extracted from the rock (Fig. 9). The 

surface area of the sampled bedrock outcrops range from a few 10’s to 100’s of square meters, and ~1-5 cm 

thick samples are collected from the bedrock surface. The abundance of 10 Be in surface samples is interpreted 

as a steady-state erosion rate (Bierman, 1994) and calculations of surface production rate corrected for 

elevation, latitude and horizon blockage are obtained using the methods of Balco and others (2008). 

14

A B 

Figure 9. Examples of sample locations. A) SH-12 on Blackrock Summit. B) SH-28 on Old Rag summit. 

The mean erosion rate all sampled summits in Shenandoah is 8.67 m/My, and rates varies from 1.56 ± 0.16 

m/My to 41.19 ± 8.9 m/My (Fig. 10). The average rate from the three surfaces sampled at Blackrock is 17.9 

± 5.1 m/My. Because erosion rates here and in other locations where block removal is important are not, 

strictly speaking, eroding steadily, we must consider the validity of the steady state erosion assumption. 

Small and others (1997) model 10 Be accumulation on bare-bedrock surfaces to determine how much error 

is introduced by applying a steady-state assumption to a location where erosion is caused by episodic block 

removal. The primary factors influencing the error are the long-term erosion rate and the thickness of the 

block removed (Small and others, 1997). Following the strategy of Small and others (1997), estimates of 

the maximum error introduced by our use of the steady-state assumption range from ~±1-18% of the actual 

long term erosion rate, with errors generally decreasing as estimated erosion rate increases (Whitten, 2009). 

Although this may appear to be a significant error, given that the lowest rates are on the order of a few m/ 

My, a 20% error does not matter when comparing these results to other rates of landscape erosion, as will be 

discussed later. In addition, such errors are nonsystematic, multiple samples should yield average erosion 

rates with less overall error relative to the actual average erosion rate (Small and others, 1997). We consider 

the erosion rates obtained to reflect maximum erosion rates. Most errors introduced by recent geologic 

events that would change production rates on sampled surfaces (e.g., devegetation and soil stripping) would 

tend to result in 10 Be abundances that would yield higher than average rates. 

Our measurements suggest that summit erosion rates in Shenandoah National Park are generally slow, and 

are closely related to rock type (Fig. 10). The Harpers Formation yields the highest average rate of ~23 m/ 

My, ~4 to 10 times higher than the other three lithologies sampled. The remaining three lithologies show very 

low erosion rates, and are similar to bedrock erosion rates in both periglacial environments and elsewhere in 

the Appalachians. Small and others (1997) found an average bare-bedrock erosion rate of 7.9 ± 4.1 m My -1 

on summit tors in the central Rocky Mountains and Sierra Nevada, USA. Our bare-bedrock summit rates are 

similar to non-fluvial erosion rates measured in this region, including on bare granite inselbergs in Georgia 

(2–10 m/My, Bierman and other, 1995); on bare sandstone and conglomerate capped ridges in Kentucky (~2 

m/My, Granger and others, 2001); on soil mantled hilltops in West Virginia (4–7 m/My, Clifton and Granger, 

2005); on bare sandstone summits along the margin of the Appalachian Plateau in West Virginia (2-10 m/ 

My, Hancock and Kirwan, 2007); and the bedrock-to-saprolite conversion rate beneath regolith in the Virginia 

piedmont (4.5–8 m/My, Pavich, 1989). In addition, our rates are comparable to an average of 7.8 m/My 

obtained by Duxbury (2009) obtained from five bare-bedrock summit samples collected within Shenandoah 

National Park. 

15

Figure 10. Summit erosion rates measured in Shenandoah National Park. Geology from Southworth and 

others (2009). 

These results suggest that bare-bedrock weathering rates on Shenandoah summits developed on lithologies 

other than the Harpers Formation are low, and are not significantly different from weathering rates obtained 

from a variety of other non-fluvial settings in the region. 

Comparison of our summit erosion rates to basin-averaged rates obtained by Duxbury (2009) suggests 

variation in the direction of relief change that is dependent on lithology. Overall, the average summit erosion 

rate (~8.8 m/My) is lower than the average of the basin-averaged rates (~13 m/My) of Duxbury (2009). More 

importantly, comparison of rates on differing lithologies yields evidence for relief change in some locations 

within the region. There are some difficulties in direct comparison between this study and Duxbury (2009), as 

we refer to the formations sampled while she refers to specific rock type sampled, but we can make general 

comparisons here by assuming the formation from which her rock types must have been derived. The basinaveraged 

erosion rate for granite-dominated watersheds is 15±4 m/My (Duxbury, 2009), while our summit 

erosion rates in similar lithologies are 5.5 m/My (Pedlar granitoids) and 2.4 m/My (Old Rag Granite Suite). The 

average for siliciclastic-floored basins is 11±4 m/My (Duxbury, 2009), while summit erosion rates in similar 

lithologies are 23±12 m/My (Harpers Formation). The average for quartzite-floored basins is 8.0±2 m/My 

(Duxbury, 2009), while summit erosion rates in similar lithologies are 4.5 ± 2.1 m/My (Antietam “quartzites”). 

Taken at face value, these comparisons imply a small but positive change in relief in granitic- and quartzitedominated 

watersheds, and a decline in relief in the siliciclastic-dominated watersheds. 

16

Importantly, the basin-averaged samples collected by Duxbury (2009) integrate across entire watersheds, 

thus including both slowly and rapidly eroding portions of the watersheds. The fact that the basin-averaged 

erosion rates in two of the sampled lithologies are higher than summit erosion rates suggest that sideslopes 

and/or channels must be eroding more rapidly than summits. The suggestion that channel erosion rates 

may be greater than the basin-averaged rates observed by Duxbury (2009) is supported by observations 

of fluvial incision rates averaged over similar timescales in this region. Fluvial incision rates compiled over 

a broad region of the east-central and southeastern U.S. average ~30–1000 m/My over ~10 4 to 10 6 years 

on Piedmont, Blue Ridge/Valley & Ridge, and Appalachian Plateau rivers (Mills, 2000). Incision rates into 

bedrock obtained from adjacent basins include ~50–160 m/My obtained from cave magnetostratigraphy and 

10 Be dating of terraces on the James River (Erickson and Harbor, 1998; Hancock and Harbor, 2003), ~20 m/ 

My from CRN dating of fluvial terraces on the New River (Ward and other, 2005), and ~600–800 m/My during 

the late Pleistocene from CRN dating of strath surfaces on the Potomac and Susquehanna Rivers (Reusser 

and others, 2004). If these rates are indicative of channel erosion rates within the Shenandoah National Park 

region, the difference between our bare-bedrock, summit-lowering rates and the fluvial incision rates imply 

that at least some parts of the Blue Ridge may be currently undergoing a period of growing relief. 

17

Acknowledgements 

Putting together a field conference with eight leaders (a.k.a.- co-conspirers) is an exciting proposition. We 

thank the Park Service for their warm welcome (and the fee waiver!) to Shenandoah as well as the use of 

the Byrd Visitors Center for the Friday evening session. Julena Campbell, Jim Schaberl, and Tim Taglauer 

were not only gracious, but extraordinarily helpful with logistics- thanks. VGFC officers, Amy Gilmer and 

Matt Heller, secured transportation and provisions for the conference. The U. S. Geological Survey has 

funded much of the research, in particular the EDMAP program has been instrumental in supporting the new 

mapping. 

Our colleagues, John Aleinikoff, Lorrie Coiner, Jorge Dinis, Amy Gilmer, Matt Heller, Bill Henika, Craig Kochel, 

Mick Kunk, Ben Morgan, Cullen Sherwood, Joe Smoot, Dick Tollo, and Steve Whitmeyer have added much 

with their insight. A legion of William & Mary students completed senior theses in and around the Park: 

significant contributions (of blood, sweat, and intellectual energy) have been proffered by Adam Forte, Joe 

Fuscaldo, Brian Hasty, Adam Gattuso, Colleen Keyser, Graham Lederer, Crystal Lemon, Joe Olney, Owen 

Nicholls, and Katie Wooton. 

18

ROAD LOG and Stop Descriptions 

The field trip road log begins at the bus parking area ~100 meters to the north of the Byrd Visitors Center at 

Big Meadows, Shenandoah National Park (mp 51). All field trip stops will be in the Park. Stops 4 and 7 involve 

hikes along well-maintained trails. Participants should be mindful of field conditions and exercise prudence. 

A number of stops include exposures along the Skyline Drive, although there is an ample shoulder at these 

outcrops participants must be cautious. Stay off the active road way and always watch for traffic. As 

with the flora and fauna, rocks in Shenandoah National Park are protected. No hammering or collecting 

of samples is permitted in the Park. At a few stops unweathered and fresh rock samples will be provided 

for closer inspection, but remember do not hammer on outcrops. Stop locations are given in decimal latitude 

and longitude using the NAD 27 datum. 

Directions and comments 

From the bus parking area follow leaders (on foot) to vantage 

point of Big Meadows. 

STOP 1 

Big Meadows (38.5168˚ N, 78.4363˚ W) 

Late Quaternary to Recent sediments, Neoproterozoic Catoctin Formation 

Litwin and Eaton 

The Big Meadows site (BMS) in Shenandoah National Park comprises Late Quaternary to Recent sediments 

that unconformably overlie the late Neoproterozoic Catoctin Formation (Allen, 1963; Gathright, 1976). Previous 

coring has determined that at least 6 m of unconsolidated sediment were deposited in the topographically 

lowest part of the BMS (Litwin and others, 2004a), although Quaternary deposits in the meadow are nonuniform 

in thickness and distribution. BMS is located at an elevation of 1,070 m (3500 ft) and experiences 

a cool microclimate; that in combination with a shallow perched water table within the alluvial sediments 

enable it to host a refugial cold-tolerant modern flora. The meadow supports small populations of balsam 

fir (Abies balsamea), red spruce (Picea rubens), Red-osier dogwood (Cornus stolonifera), and Canadian 

burnet (Sanguisorba canadensis). Each of these plants more typically thrives in environments at least 4-5˚ 

latitude north of this site, in New England and northward. It also hosts a pine species restricted to the 

cooler microclimate along the ridge crests of the Blue Ridge and Appalachians, Table Mountain pine (Pinus 

pungens; Little, 1971, 1976; W.B. Cass and R. Engle, NPS (SHEN), pers. comm.). The Virginia Blue Ridge 

generally is vegetated by a southern extension of the temperate Appalachian oak forest, and this forest 

surrounds the BMS (Küchler, 1964, 1975; Litwin and others, 2004a). Figure 11 illustrates the general forest 

structure presently established in eastern North America, and indicates the location of Shenandoah National 

Park with respect to it. From south to north, these are: 1) southern mixed forest, 2) oak-hickory-pine forest, 

3) Appalachian oak forest, 4) northern hardwoods forest, 5) northern hardwoods-spruce forest, 6) spruce-fir 

forest (minor type), 7) boreal forest, and 8) forest-tundra to tundra (Küchler, 1964, 1975). The modern mean 

annual temperature (MAT) isotherms for this same geographic area also are shown (Owenby and others, 

1992; NCDC, 2002: Litwin and others, 2004a). 

Fossil pollen evidence has been recovered from sparsely distributed, small depositional units across the 

eastern flank of the Blue Ridge, in and around Shenandoah National Park. Recent debris-flow denudation 

in the Blue Ridge (Morgan and Wieczorek, 1996; Morgan and others, 1999; Eaton, 1999; Wieczorek and 

others, 2000; Eaton and others, 2003) presented an opportunity to identify and collect climate information 

from these isolated Late Quaternary deposits. Most of these samples were dated directly using AMS 14 C 

analyses. These samples indicated that the vegetation in the Blue Ridge changed repeatedly with changes 

in Late Quaternary climate (Litwin and others, 2001, 2004a). Figure 12 shows a map of the Blue Ridge sites 

from which fossil pollen samples were collected (Litwin and others, 2004a). 

19

2 

1 

3 

5 

4 

70 

0 

75 

65 

0 

0 

Figure 11. Generalized forest zonation of eastern North America, showing modern mean annual temperature isotherms 

and forest zones. From south to north, these are: 1) southern mixed forest, 2) oak-hickory-pine forest, 3) 

Appalachian oak forest, 4) northern hardwoods forest, 5) northern hardwoods-spruce forest, 6) spruce-fir forest (minor 

type), 7) boreal forest, and 8) forest-tundra to tundra (Küchler, 1964, 1975). 

0 

78 30' 

0 

38 37.5' + 

+ 

N 

+ 

+ 

+ 

7 

60 

55 

0 

20 

0 

8 

6 

50 

25 

45 

0 

0 

0 

35 

0 

30 

Atlantic Ocean 

0 

40 

0 

35 

(after Kuchler, 1964, 1975) 

+ + 

+ + 

+ + 

+ 

+ 

0 

78 45' 

+ 

+ 

0 

38 22.5' + 

0 

78 30' 

SNP 

Moormans 

River 

0 

38 10' 

study area 

Shenandoah National Park 

Skyline Drive 

+ 

Pasture 

Fence Mt. 

0 

38 07.5' 

+ 

Big 

Meadows 

+ 

0 

78 42.5' 

+ 

core 

+ 

Hoover 

Camp 

Bluff 

Mt. 

study area+ 

Rapidan River 

Staunton River 

Kinsey 

Run 

Kirtley Mt. 

Wilson Run 

+ + 

Graves 

Mill 

+ 

(modified from Litwin et al., 2004a) 

+ 

+ 

0 

38 22.5' 

0 

78 15' 

0-1000' elev. 1000-2000' elev. 2000-3000' elev. 3000'+ elev. 

Rapidan River 

0 

78 15' 

+ 

0 

38 37.5' 

+ 

km 

+ 

0 1 

+ 

mi 

0 1 

Figure 12. Site map of study area, Shenandoah National Park and environs, Virginia. 

+ 

0

Figure 13 illustrates a preliminary relative temperature model for the pollen samples that were recovered, and 

estimates for the forest type that each pollen sample most closely resembled (forest types abbreviated). One 

of the first observations noted was that the dominant forest type covering the Blue Ridge today, Appalachian 

oak forest, did not exist in this area for most of the 45 ky of geologic record covered by these samples. The 

Appalachian oak forest likely first established itself on the Blue Ridge no earlier than ~15 ka, based on the 

calculated paleo-insolation curve (Litwin and others, 2001; Berger, 1978). 

525 

FODs 

Liquidambar 

? Morus 

Magnolia 

Decline in 

Picea 

Platanus 

Juglans 

Nyssa 

Ilex 

Liriodendron 

Castanea 

Carya 

Tsuga 

HOLOCENE 

NH 

NHS 

SF 

B 

? 

? 

PLEISTOCENE 

SM 

? 

OHP 

AO 

SOLAR (INSOLATION) 

TRENDS- 60 0 N LAT 

? ? LGM ? 

no relative thermal 

signal available 

14 C calibration limit 

0.1 

0 10 20 30 40 

(modified from Litwin et al., 2004a) 

14 C kyr Cal BP 

Figure 13. Blue Ridge dataset summary with reconstruction of prehistoric forest zones, observed first occurrence 

datums (FODs), and sequence and timing of climate-driven vegetation change (0- 45 ka).. 

This pollen evidence also suggested that during the Last Glacial Maximum (LGM, ~21 to 25 ka) the valleys 

and foothills surrounding the Blue Ridge most probably were covered completely with boreal forest. With 

the additional temperature reduction imposed by adiabatic cooling, the ridge crests along the Blue Ridge 

probably supported only alpine tundra, especially at high elevation sites such as Big Meadows. The modern 

boreal forest’s southerly limit in North America is positioned south of the southerly limit of permafrost. The 

southerly boreal forest limit occurs at approximately 1.5˚ C MAT (Ecoregions Working Group, 1989; Litwin 

and others, 2004a). The southerly limit of permafrost occurs at approximately -1.0˚ C MAT (Brown and 

others, 1997). The vegetation signal from the Blue Ridge pollen data suggested a minimum mean annual 

temperature shift of ~11˚ C (20˚ F), from the LGM to the present (Litwin and others, 2004a). This thermal 

difference is bounded between the coldest limit of the present forest in the BM area (Appalachian oak forest, 

at approximately 12.5˚ C (55˚ F)) and the warmest limit of the coldest forest type suggested by the pollen 

evidence at Kinsey Run (boreal forest, at approximately 1.5˚ C (35˚ F)). 

Evidence of LGM-induced cryogenic processes and cryogenic depositional features occurs throughout in 

and near Shenandoah National Park. In and around the study area stratified slope deposits are exposed 

at Hoover Camp (~700 m), and at Kinsey Run (~350’ m), which indicate that solifluction and gelifluction 

21 

calibrated sample 

sample older 

18-O 

than 

calibration limit 

Quercus/(Picea+1) 

interpolated age- older 

Quercus/ than (Picea+Abies)+1 

calibration limit 

AVERAGE MODERN BASELINE 

OF STUDY AREA 

1000 

100 

10 

1 

RELATIVE TEMPERATURE INDEX 

Warmer Colder 

(Quercus+ Fraxinus)/(Picea+ Abies+1)

processes previously were dominant depositional processes in this area (Smoot, 2004a,b; J.P. Smoot, pers. 

comm.). Dropstone fabrics (i.e., ice-rafted pebbles and cobbles) also are present in the Kinsey Run section 

(J.P. Smoot, unpub. data). At Big Meadows, the shallow surface of the BMS has been interpreted as relict 

sod-bound tundra (J.P. Smoot, pers. comm.). In low-angle light this landform is visible as a series of shallow 

steps and risers covering the bowl-shaped depression at Big Meadows. Cryogenic processes also dominated 

the deposition along the Blue Ridge at Black Rock (Eaton and others, 2002). The present landscape in 

and around the Blue Ridge commonly displays a Holocene-age geomorphic and textural overprinting that 

incompletely modifies a relict periglacial landscape. In contrast, the fossil pollen evidence suggests that the 

past floras that have grown on these relict surfaces appear to have transformed regularly in near-equilibrium 

with past climate shifts over a 45 ky time period (Litwin and others., 2001, 2004b). 

The fossil pollen evidence suggests that the vegetation changes on the eastern flank of the Blue Ridge since 

the LGM have been frequent, abrupt, and strongly expressed. Forests appear to have shifted approximately 

every ~200 years on average over the past 45 ky (Litwin and others, 2004a). Some of these shifts probably 

have been quite abrupt. Current analyses of new, long (35 m) high-resolution cores taken along the Potomac 

River drainage are now be used to test the findings from these Blue Ridge paleovegetation and paleoclimate 

analyses. 

Cumulative Point-to-Point 

Trip Mileage Mileage Directions and comments 

0.0 Load the buses, find a seat and sit down. Turn right out of bus 

parking area onto Big Meadows entrance road. 

0.1 0.1 Turn right (south) on Skyline Drive (mp 51.2). 

14.7 14.6 Cross U.S. Rt. 33 at Swift Run Gap (mp 65.8). 

Swift Run Gap is developed along a steeply dipping transverse 

fault. 

15.9 1.2 Mass wasting originating in basement rocks exposed 

on the southeast side of the Skyline Drive blocked the 

road in 2003. Although the Park Service has attempted to 

remediate the problem, debris commonly wastes off the cut. 

16.7 1.0 Turn right into Sandy Bottom Overlook (mp 67.8). Unload 

buses. 

STOP 2 

Sandy Bottom Overlook (38.3360˚ N, 78.5657˚ W) 

Basement Complex- charnockite 

Southworth and Bailey 

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 the gently plunging syncline. Recent work by Shufeldt and 

others (2008) demonstrates the synclinorium is cut by a number of northwest dipping back thrusts. The 

mountain front between the Elkton area and north to Stanley is quite sinuous (Figs. 1 and 4). In this area 

King (1950) recognized the Elkton embayment and Shenandoah salient, but did not place a fault at the foot 

of the range. 

South of Elkton, early Cambrian rocks of the Antietam Formation are brecciated and structurally overlie 

22

younger carbonate rocks. The sinuous nature of the mountain front likely reflects the low-angle nature of the 

Blue Ridge fault and offset associated with younger transverse faults. 

Hightop forms the mountain immediately to the east of the overlook (Fig. 14). The “great unconformity” 

between the Mesoproterozoic basement and Neoproterozoic cover sequence occurs just above the Skyline 

Drive. The basement is overlain by arkosic phyllite of the Swift Run Formation and at least 250 meters of 

Catoctin metabasalts. On Hightop, individual Catoctin lava flows dip gently to the southeast, such that the 

steep northwest flank of the mountain forms the anti-dip slope. The basement and cover sequence are 

folded into a broad open anticline (Roundtop anticline). Sandy Bottom is underlain by a northwest-striking 

transverse fault that places basement against the Catoctin Formation (Fig. 14). 

Figure 14. Inset from the Geologic map of the Swift Run Gap 7.5’ quadrangle with Stops 2 and 6 (from Bailey and 

others, 2009). 

The large roadcut exposes a distinctive orthopyroxene-bearing monzogranite. This charnockite is part of a 

compound pluton that occurs from south of Swift Run Gap to north of Skyland. The rock is composed of 10- 

30% alkali-feldspar, 30-50% plagioclase, 15-25% quartz, 10-15% orthopyroxene with minor amphibole and 

clinopyroxene. Rock textures range from porphyritic with alkali-feldspar megacrysts to equigranular. The 

unit is commonly massive, but does carry a weak foliation at some locations. A U-Pb zircon age of 1,049±9 

Ma was obtained from rock at this location and is interpreted as a crystallization age (Southworth and others, 

2009). Charnockite exposed immediately east of Hightop Mountain (4 km to the east) yielded a Ar-Ar age 

from amphibole of 950 Ma indicating that the basement complex in the western Blue Ridge has not been 

heated above the amphibole closure temperature (450˚ to 500˚ C) since the waning stages of the Grenvillian 

orogeny. 

23



16.7 0.0 Bear right (south) on Skyline Drive. 

22.3 5.6 Cross Simmons Gap. 

Simmons Gap sits along a 10 km-long, steeply dipping transverse 

fault that strikes 350˚ and extends across the Blue Ridge into the 

Great Valley. 

31.7 9.4 Outcrop on the west side of road exposes interbedded granule 

conglomerate to coarse quartz sandstone and metasiltone/phyllite. 

of the Weverton Formation. Beds are upright and dip gently to the 

northwest. Penetrative foliation is well developed in phyllite. 

33.7 2.0 Turn right in Blackrock Summit parking area (mp 84.8). Unload 

buses and assemble for hike to Blackrock Summit. 

STOP 3 

Blackrock Summit (38.2200˚ N, 78.7404˚ W) 

Harpers Formation- well-cemented quartz sandstone, Quaternary block field 

Eaton, Hancock, and Lamoreaux 

Blackrock Summit forms an isolated summit tor with outstanding views of the Blue Ridge and Shenandoah 

Valley. The bedrock crags are composed of a grayish to dusky purple, medium- to coarse-grained, wellcemented 

quartz sandstone in the Harpers Formation. Well-cemented sandstone (quartzite) is a minor 

rock type in the silt dominated Harpers Formation (Fig. 15), forming laterally discontinuous lenses up to a 

kilometer in length and 5 to 15 meters in thickness. 

Figure 15. Inset from the Geologic map of the Browns Cove 7.5’ quadrangle with Stops 3 and 4 (from Lamoreaux 

and others, 2009). 

24

Cross and plane beds are present and the sandstone likely represents a shoaling sequence in the marine. 

Bedding dips 10˚ to 20˚ to the northwest and is cut by two sets of subvertical fracture sets. 

Exposures of bedrock and block slopes near the summits of the Blue Ridge Mountains in central Virginia 

reveal a variety of deposits suggestive of periglacial slope processes. The summit of Blackrock at an elevation 

of 933 meters (3,060’), is a shattered tor consisting of gently dipping (12°), broken quartzite blocks and minor, 

less disturbed bedrock of the Harpers Formation (Fig. 16). The tor, consisting of isolated quartzite columns 

(3-8 m high), has undergone dislocation by cambering, leaving openings up to several hundred cm between 

prominent orthogonal joint sets (Fig. 17). Below the ridge top where the slopes gradually increase, some 

remaining displaced bedrock columns show rotational movement which probably facilitated toppling collapse 

and production of large amounts of blocky debris. Prominent block slopes extend downward from both the 

east and west sides of the tor. On the west (dip slope) side the block slope extends over 500 m downslope 

on gradients that vary between 18-35°. The longitudinal profile of the block slope is an undulating surface of 

weakly developed benches and escarpments (Fig. 18). 

Detailed measurements of clast volume and orientation show four sequences where boulder deposits exhibit 

an increase in mean clast volume downslope, consistent with discrete events and landforms. The thickest 

accumulations of debris are associated with large individual blocks that exceed 4 m in length (Fig. 18). The 

fabric of the deposits is largely open framework, some sites exhibit distinct layers of finer materials underlying 

coarser materials. The majority of the sites show a preferential downslope clast orientation of the long axis 

(Fig. 16). The lower half of this block slope narrows into a slightly sinuous path for ~150 m, and abruptly 

terminates with a distinct snout just above a small first order tributary of Paine Run. Additionally, circular 

shallow depressions measuring 1-3 m in diameter on the block slopes suggest an ice-rock mixture prior to 

melting and subsidence of these deposits, producing the depressions. 

These landforms and deposits are similar to solifluction and relict rock glacier deposits documented in the 

literature. Our continuing investigations will focus on the mechanisms and chronology of the slope processes, 

and their association with climatic events. 

Lowering of Blackrock Summit likely occurs through removal of blocks bound by weathered sub- horizontal 

bedding planes and vertical joints, as well as toppling of isolated tors and direct removal by rain and wind 

of granular material produced from weathering. To estimate the erosion rate at this site, rock samples were 

collected the surface of bedrock outcrops, and the abundance of the cosmogenic radionuclide (CRN) 10 Be in 

quartz (Fig. 9). 10 Be abundance in surface samples is interpreted as a steady-state erosion rate (Bierman, 

1994) and calculations of surface production rate are corrected for elevation, latitude and horizon blockage. 

The average erosion rate from three surfaces sampled at Blackrock is 17.9 ± 5.1 m/My, which is nearly 10m/ 

My greater than the average for eight other summit sites across the Park. 

25

B. 

= 1 m 3 

0 50 m 

2600 

A. 

2800 

2700 

2800 

2800 

2700 

STUDY AREA 

ROCKINGHAM COUNTY 

ALBEMARLE COUNTY 

2700 

2900 

APPALACHIAN TRAIL 

3000 

2900 

BLACKROCK 

SUMMIT 

2900 

26 

3000 

3100 

2600 

2800 

2700 

3000 

N 

0 1000 ft 

"West Face" Block Slope, Blackrock 

Figure 16. A. Base map of Blackrock Summit site. B. Detailed measurements of clast volume and orientation 

show four sequences where boulder deposits exhibit an increase in mean clast volume downslope, 

suggesting discrete events and landforms.

A. 

C. 

D. 

s 

lope p o 

r file 

B 

B 

scale 

E 

R I DG OW 

F U 

R 

R 

B 

A 

A 

A 

E. 

UNDISTURBED 

BLOCK FIELD 

SURFACE 

27 

1 

B. 

2 

25 

Figure 17 A. Shattered tor (height ~10 m) 

Quartzite blocks break along A) bedding planes 

and B) joint planes. B. Quartzite columns 

displaced by cambering and toppling. The tilt of 

the lines demonstrates the toppling of the nearly 

horizontal beds after cambering has opened 

and separated the blocks along joint planes. C. 

Prominent block slopes west of Blackrock 

summit. D. "Curious" ridge and furrow feature at 

the toe of the block slope shown in C. Note the 

geologists for scale. Landform may be a 

protalus rampart. E. Excavation pit of boulder 

front individual numbered boulders are identified 

in Figure 18. The line marks the upper 

boundary of the excavation. 

3 

7 

4 

5 

PIT 

FACE 

6 

BASE 

OF 

PIT

A. 

1.6 m 

A' 

A' 

~2 m 

B. 

PLAN VIEW 

Escarpment (knicks face downslope) 

Finer grained, sorted talus 

CROSS SECTION 

BEDROCK 

28 

2.4 m 

A 

A 

~10 m 

Figure 18. A. Plan and cross section views of the topography depicting weakly-developed benches 

and escarpments. B. Longitudinal profile of boulder front. The fabric of this deposit is dominantly 

open framework.



33.7 0.0 Load buses and go north on Skyline Drive (mp 84.8). 

34.8 1.1 Turn right into Dundo Picnic area (mp 83.7). 

35.1 0.3 Park and unload buses. 

STOP 4 

Dundo Picnic area- Lunch (38.2349˚ N, 78.7180˚ W) 

Weverton Formation- well-cemented coarse-grained sandstone 

Southworth and Lamoreaux 

The Dundo Picnic area will serve as our lunch stop; low outcrops of well-cemented coarse-grained sandstone 

to granule conglomerate in the Weverton Formation occur immediately to the northwest of the picnic area 

and along the Appalachian Trail to the east of the picnic area (Fig. 15). Coarse-grained sandstones in the 

Weverton Formation underlie much of the Blue Ridge’s crest in the area, although blue-gray quartzose 

phyllite makes up a significant part of the Weverton Formation in the southern part of Shenandoah National 

Park. Bedding dips gently to the northwest and the penetrative foliation dips to the southeast. 

In the 1960s wells were drilled to obtain water for the Dundo site. A 190-meter deep test well at the picnic 

area penetrated 42 meters of quartzose phyllite, metasandstone and metaconglomerate interpreted as the 

Weverton Formation, 30 meters of chlorite schist, phyllite, and conglomerate underlain by 116 meters of 

metabasaltic greenstone collectively interpreted as the Catoctin Formation (DeKay, 1972). The well yielded 

25 gallons per minute and the static water level was 110 meters below the surface. 

A purple to maroon spotted phyllite (with the spots being sericite and chlorite blebs) is commonly, but not 

everywhere present, at the top of the Catoctin Formation. The chemistry and texture of the spotted phyllite 

indicates that this rock originated as a basaltic metatuff. The absence of a laterally persistent spotted 

phyllite beneath the siliciclastic rocks of the Weverton Formation is consistent with an unconformable contact 

between the Catoctin Formation and Chilhowee Group. Traditionally, the Swift Run/Catoctin sequence 

upwards through the Chilhowee Group has been interpreted to record the rift to drift transition associated 

with the opening of the Iapetus Ocean, however the Catoctin/Weverton unconformity requires modification 

of this tectonic model. 



35.1 0.0 Load buses and continue out of picnic area 

35.5 0.4 Turn right (north) on to Skyline Drive. 

38.3 2.8 Outcrop on the southeast side of road exposes purple, porphyritic 

metabasalt of the Catoctin Formation. 

41.1 2.8 Turn left into Rockytop Overlook (mp 78.1). Unload buses. 

STOP 5 

Rockytop Overlook (38.2789˚ N, 78.6651˚ W) 

Harpers Formation- phyllitic siltstone and thinly bedded metasandstone 

Bailey and Hancock 

The view of the Big Run watershed (to the south and west) is expansive. Big Run forms the largest drainage 

29

A. 

78˚ 40’ W 

B. 

1 

Shenandoah Valley 

0 1 

kilometers 

2 

Big 

Run 

78˚ 42’ W 

Patterson Ridge 

Figure 19. A. Shaded relief map of Big Run drainage basin. B. Hypsometric curves for the Big Run (dashed line) 

and Doyles River (solid line). 

entirely within the Park (Fig. 19A). Most of the drainage basin is underlain by Chilhowee siliciclastics (mostly 

Harpers Formation metasiltstone), although greenstone of the Catoctin Formation crops out at the base of 

the slopes in the southern part of the basin. Bedding generally dips to the northwest, but a series of northnortheast 

plunging, asymmetric folds occur in the Antietam Formation along Brown and Rocky mountains. 

As is typical for drainage basins developed in the Chilhowee Group, slopes are steep with ~60% of the Big 

Run basin having slopes in excess of 35% (20˚ slope). The elevation/area relationship (hypsometric curve) 

for the Big Run basin is quite different from the neighboring Dolyes River drainage (developed primarily on 

the Catoctin greenstone and basement complex) (Fig. 19B). 

The large roadcut on the east side of the Skyline Drive exposes typical Harpers Formation lithologies. Brown 

to grayish phyllitic siltstone is predominant, but thinly bedded sandstone is also present. A penetrative 

northeast-striking foliation (cleavage) is well developed and dips moderately to the southeast, bedding is 

A. B. 

N=115 

Rocky Mountain 

Big Flat 

Mountain 

38˚ 18’ N 

ROCKYTOP 

OVERLOOK 

38˚ 16’ N 

N N 

Equal Area 

C.I. = 1, 3, 6, 

12 %/1% area 

fractures fracture strikes 

30 

0 

0 0.1 0.2 0.3 0.4 0.5 0.6 0.7 0.8 0.9 1 

Circle = 30 % 

Figure 20. A. Contoured stereogram of poles to fractures in the Harpers Formation between 

miles 76-79 B. Rose diagram of fracture strikes 10˚ bins. 

Height above basin outlet/Total relief 

0.9 

0.8 

0.7 

0.6 

0.5 

0.4 

0.3 

0.2 

0.1 

Drainage area below height above basin outlet 

/Total drainage area

upright and dips gently to the northwest at the southern end of the roadcut. A set of steeply dipping joints cut 

the rock; the dominant set strikes to the west-northwest (Fig. 20) and individual fracture surfaces are commonly 

ornamented with plumose structures (best viewed when the fracture surface is obliquely illuminated). Westnorthwest-striking 

extension fractures are common in the Chilhowee Group. This fracture set formed after 

the Blue Ridge cover sequence was folded and developed under a regional stress field with the maximum 

compressive stress oriented ~290˚±10˚. 



41.1 0.0 Load buses and Turn left (north) on to Skyline Drive. 

44.7 3.6 Loft Mountain Overlook (mp 74.5) 

A diverse array of lithologies are exposed in strongly deformed 

Catoctin Formation at the overlook. Nice view to the southeast 

with Simmons Gap fault between overlook and prominent foothills 

that are underlain by the basement complex. 

49.3 4.6 Cross Powell Gap. 

Powell Gap sits along a steeply dipping transverse fault bounding 

the southwest margin of the Bacon Hollow graben. The gap 

is underlain by granitoid gneiss; phyllitic arkosic of the Swift Run 

Formation is exposed in roadcuts 0.2 miles to the south. 

49.9 0.6 Turn right into Bacon Hollow Overlook (mp 69.3). Unload buses. 

STOP 6 

Bacon Hollow Overlook (38.3215˚ N, 78.5822˚ W) 

Basement complex- granitoid gneiss and low-silica charnockite 

Bailey and Southworth 

Bacon Hollow Overlook is cut into the southern slopes of Roundtop Mountain and affords views to the 

southeast. In the far distance, approximately 40 km away, the Southwestern Mountains form a low linear ridge 

underlain by the Catoctin Formation on the southeastern limb of the Blue Ridge anticlinorium. Neoproterozoic 

metasedimentary rocks of the Lynchburg Group and the basement complex underlie the low terrain in the 

middle distance. In the foreground, the southern rampart of Hightop is to the left, Bacon Hollow in the center, 

and Flattop Mountain to the right (note the eclectic array of buildings dotting the upper slopes of Flattop). 

Bacon Hollow forms a steep-walled, 500-meter deep valley drained to the south by the Roach River. The 

valley floor and lower slopes are underlain by granitoid gneiss with an east-striking high-grade foliation. The 

basement is overlain, on both Hightop and Flattop, by the Swift Run and Catoctin formations. A set of newly 

recognized, north-northwest striking faults bound the U-shaped valley and pass beneath Powell and Smith 

Roach gaps (Fig. 14) (Bailey and others, 2009). Slip on these steeply dipping faults appears to have “down 

dropped” rocks in Bacon Hollow (Fig. 14). Bacon Hollow is, in essence, a graben. Total displacement on the 

bounding faults is

Granitic gneiss is exposed at the overlook and along the large cuts on the Skyline Drive. The rock is composed 

of 30-40% perthite, 30-40% quartz, 5-20% plagioclase and up to 15% biotite. This rock forms an extensive 

unit along the eastern slopes of the Blue Ridge in southern Shenandoah National Park. The chemical 

composition of this unit ranges from alkali-feldspar granite to monzogranite and at some locations the rock 

is leucocratic. A U-Pb zircon age of 1,150±23 Ma was obtained from this exposure and the granitic gneiss is 

part of the older plutonic suite (Southworth and others, 2009). A charnockite dike intrudes the granitic gneiss 

in the cut ~120 meters southwest of the overlook and the summit of Roundtop is underlain by charnockite 

(observed at Stop 2, Fig. 14) 

A medium- to coarse-grained gneissic foliation is evident. This fabric is defined by mm- to cm-scale, elongate 

aggregates of feldspar and quartz (individual grains are subequant) and aligned biotite that formed under 

high-grade (but solid-state) conditions. At this location foliation strikes east-northeast to east and dips 25˚ to 

50˚ to the north. In the Swift Run Gap quadrangle high-temperature foliation in the granitic gneiss is folded 

(Fig. 21A); this fabric was formed and folded prior to the intrusion of the ~1,050 Ma charnockitic pluton. 

A younger foliation, defined by aligned greenschist minerals, is also developed at many locations in the 

basement complex. In contrast to the high-temperature fabric, the greenschist-facies foliation consistently 

strikes to the northeast and dips moderately southeast (Fig. 21B). 

N 

A. B. 

N=81 

high-temperature 

foliation 

π-axis 

027˚ 40˚ 

C. I. = 2σ 

Equal Area 

32 

N=160 

C. D. 

N=144 

N 

greenschist-facies 

foliation 

N N 

Equal Area 

C. I. = 2, 6, 

12%/ 1%area 

circle = 

30% of data 

fractures fracture strike 

mean foliation 023˚ 41˚ SE 

C. I. = 2σ 

Figure 21. A. Contoured stereogram of poles to high-temperature foliation. B. Contoured stereogram 

of poles to greenschist-facies foliation. C & D. Contoured stereogram of poles to fractures in the granitic 

gneiss

The granitic gneiss is cut by numerous fracture sets and in the exposures along the Skyline Drive northeaststriking 

fractures are common (Fig. 21C, D). These fractures display a wide range of dips. A number of 

northwest-striking and northeast-striking fractures are coated with Fe-oxide slickensides that record hanging 

wall down slip. The significance of these shear fractures is unclear. Regionally, the basement complex 

displays much greater variably in fracture orientation than the cover sequence. 



49.9 0.0 Load buses and turn right (north) on to Skyline Drive. 

53.4 3.5 Cross U.S. Rt. 33 at Swift Run Gap (mp 65.8). 

54.8 1.4 Hensley Hollow Overlook (mp 64.4). 

Outcrop on southeast side of road exposes well-foliated arkosic 

Phyllite in the Swift Run Formation. Immediately to the southwest 

of the overlook, foliated and altered basement charnockite is 

exposed. The contact is interpreted to be unconformable at this 

location. 

56.5 1.7 Turn right into South River picnic area (mp 62.7). 

56.9 0.4 Park and unload buses. 

REST STOP 

South River Picnic Area (38.3813˚ N, 78.5173˚ W) 

Catoctin Formation- foliated greenstone 

Burton 



56.9 0.0 Load buses and continue out of picnic area. 

57.1 0.2 Turn right (north) on to Skyline Drive. 

63.0 5.9 Pull off on right (east) side of Skyline Drive across from Meadow 

School parking area (mp 56.8). Unload buses and assemble for 

hike. 

STOP 7 

Bearfence Mountain Traverse 

Bailey 

Stop 7 involves a hike along the Appalachian Trail over Bearfence Mountain. The total distance of the 

hike, from the drop off point to the pick up point, is 2 km (1.4 mi.) with a vertical climb of ~100 m (300 ft). 

The Appalachian Trail is well maintained and the hike requires a modest level of effort. Those who do not 

wish to participate in the hike should remain with the buses and travel north to the pick up point. Figure 22 

is a 1:4,000 scale geologic map of the stop and should be consulted as a guide to locating the described 

outcrops. The decimal latitude and longitude of the outcrops is given using the NAD 27 datum. 

– – – – – – – – – – – – – – – – – – – – – – – – – – – – – – – – – – – – – – – – – – – 

Hike north along the Appalachian Trail. Prior to the first switchback there is abundant float of arkosic phyllite 

from the Swift Run Formation. Beyond the first switchback, small outcrops of basement granitoid occur. 

33

EXPLANATION 

Catoctin Formationmetasandstone 

Catoctin Formationmetabasalt 

Swift Run Formationarkosic 

metasandstone 

Charnockite and 

Granitic Gneiss 

drop off 

point 

0 

A T 

45 

feet 

100 

Skyline Drive 

Zsr 

meters 

fractional scale 1:4,000 

A T 

36 

15 49 

bedding foliation lineation Appalachian 

Trail 

0 200 400 600 

200 

49 

pickup 

point 

47 

48 

39 

34 

A T 

43 

Stop 7b 

38.4470˚ N, 78.4659˚ W 

15 

46 

45 

Figure 22. Geologic map of Bearfence Mountain area. 

Bearfence Mtn. 

parking area 

Y 

N 

49 

Zcb 

Zcb 

53 

8 

Stop 7d 

38.4522˚ N, 78.4654˚ W 

Zcs 

12 

44 

65 

49 

31 

21 

39 

50 

65 

52 

35 

39 

12 

41 

29 

45 

Stop 7a 

38.4463˚ N, 78.4655˚ W 

Bearfence Mtn. rock scramble 

58 

Zcb 

53 

71 

41 

38 

59 

71 

Stop 7c 

38.4474˚ N, 78.4657˚ W 

x Bearfence 

Mtn.summit 

Zsr 

52 

43 

Y

Pass the third switchback and continue ~160 m onwards to a large outcrop immediately south (right) and just 

below a rock wall on the trail (Fig. 22). 

STOP 7A 

(38.4463˚ N, 78.4655˚ W) 

Basement complex- altered granitoid gneiss 

The overhanging outcrop of granitoid gneiss is strongly epidotized (Fig. 23A). Epidote is a secondary mineral 

produced by the reaction of Fe-bearing minerals and feldspar in the presence of water. Prior to alteration 

the basement was a pyroxene-bearing charnockite, but few mafic minerals (other than epidote) are still 

present. The granitoid gneiss is cut by numerous hematite-stained fractures that are cut by white quartz 

veins. Foliation strikes to the north and dips ~40˚ east. Did the foliation develop prior to alteration? 

This exposure is less than 20 meters (map distance) from the overlying greenstone. One possibility is that 

epidote was generated by near-surface hydrothermal alteration when the overlying Catoctin lavas were 

erupted in the Neoproterozoic. Another possible scenery involves the formation of epidote during regional 

metamorphism and deformation in the Paleozoic. In Shenandoah National Park altered epidote-rich basement 

is common near the “great unconformity”, but is not confined to the contact zone. At a number of locations 

east of the Park epidote-rich coronas form around leucocratic basement inclusions that are surrounded by 

biotite-rich granitoids, and clearly developed during regional metamorphism (Wadman and others, 1997). 

Continue upwards along the Appalachian Trail. Pass low outcrops of altered basement along the trail, to the 

left (uphill) greenstone float of the Catoctin Formation is common. At ~140 m (0.1 mi) the trail goes through 

a switchback and outcrops of greenstone occur. At this location the Catoctin Formation is directly above the 

basement and siliciclastic rocks of the Swift Run Formation are not evident. At 60 m beyond the switchback 

the Appalachian Trail is joined on the right by a trail to the summit of Bearfence Mountain. Continue along 

A. 

qtz-filled 

fractures 

foliation 

the Appalachian Trail, the trail is nearly flat over this section. 120 m past the trail junction flagging denotes 

the way to outcrops 10 to 20 m southwest (left) of the Appalachian Trail (Fig. 22). 

STOP 7B 

(38.4470˚ N, 78.4659˚ W) 

Swift Run Formation- arkosic metasandstone 

Toppled blocks of Catoctin greenstone directly overlie outcrops of arkosic metasandstone in the Swift Run 

Formation. The medium- to coarse-grained Swift Run metasandstone contains abundant feldspar and quartz 

with rock fragments and heavy minerals forming a minor component. Cross and plane bedding is common 

B. 

cross 

beds 

35 

bedding 

Figure 23. A Altered granitoid gniess at stop 7a. B Arkosic meta-sandstone at stop 7b.

and graded bedding evident at a few locations (Fig. 23B). These rocks are clast supported, well to moderately 

sorted, and contain much less sericitic mica (phyllitic) than most Swift Run lithologies. The provenance is 

the granitic basement complex and sedimentary structures are consistent with a fluvial to fluvial braidplain 

depositional environment. 

Although the Swift Run Formation is missing to the southeast (near Stop 7A), there is at least 50 meters 

of metasandstone exposed to the west-northwest (down slope) above the basement complex. Dramatic 

thickness changes over a short distance are a characteristic of the Swift Run Formation and could reflect 

significant local topography at the time of deposition or later removal of an originally more laterally expansive 

unit. 

Bedding is upright and dips 10˚ to 20˚ to the northeast (Fig 23B). A penetrative foliation dips ~45˚ to the 

southeast and is refracted across grain-size changes in beds. Ar-Ar ages from syn-tectonic sericite in Swift 

Run lithologies exposed to the southeast of the Park range from 350 to 320 Ma (Bailey and others, 2007; 

Gattuso and others, 2009) and demonstrate that the penetrative deformation (foliation forming) event in the 

western Blue Ridge occurred prior to or at the earliest stage of the Alleghanian orogeny (325 to 280 Ma). 

– – – – – – – – – – – – – – – – – – – – – – – – – – – – – – – – – – – – – – – – – – – 

Return to the Appalachian Trail and continue north (left), passing the base of greenstone cliffs on the right. 

The trail descends gently and 140 m after stop 7B a side trail intersects on the right (Fig. 22). 

To journey on to Stop 7C (time and spirit permitting) take the side trail upwards to the right. 50 m on the trail 

joins the Bearfence Mountain summit trail, continue to the south for another 80 m to the rocky promontory 

(although great views also occur along the rock spine to the north). 

STOP 7C 

(38.4474˚ N, 78.4657˚ W) 

Catoctin Formation- foliated greenstone 

This location provides quality vistas of the western Blue Ridge and Massanutten Mountain beyond. The rock 

underfoot is Catoctin greenstone. Foliation strikes towards 030˚ and dips 45˚ southeast. Chlorite blebs, 

evident on exposed foliation surfaces, are elongate and define a downdip lineation that likely corresponds 

to the direction of maximum stretching. Among the greenstone cliffs on the Bearfence crest, foliation dips 

vary between 70˚ and 30˚, and in places, well-developed shear bands are present. Badger (1999) reports 

columnar joints from a number of spots along the Bearfence Mountain ridge crest. 

– – – – – – – – – – – – – – – – – – – – – – – – – – – – – – – – – – – – – – – – – – – 

Retrace steps and return to the Appalachian Trail. Bear right and continue north, the trail descends and 

greenstone cliffs are above and to the right of the trail. After ~440 m the Appalachian Trail intersects the 

Bearfence Mountain trail again. Turn left (northwest) onto this trail. After 20 m step off the trail to the right 

(Fig. 22). 

36

STOP 7D 

(38.4522˚ N, 78.4654˚ W) 

Catoctin Formation- arkosic metasandstone 

Laminated to thinly bedded quartzose metasandstone and pebble meta-conglomerate are exposed in the 

low outcrop. The rock is quite similar to the Swift Run metasandstone at Stop 7B. Bedding is upright and 

dips ~15˚ southeast, whereas foliation dips 45˚ to 50˚ southeast. Previous workers have interpreted these 

metasandstones as part of the Swift Run Formation exposed in the nose of a northeast-plunging overturned 

anticline. Field relations indicate that greenstone occurs both above and below the metasandstone layer 

(Fig. 22). An alternative interpretation, consistent with the field data, is that the metasandstone forms a layer 

within the Catoctin Formation. 

Metasedimentary rocks in the Swift Run and Catoctin formations are comparable and sourced from the 

same protolith; lithologic similarity and geometric relations indicate the Swift Run and lower Catoctin are 

contemporaneous units. Field relations at Bearfence Mountain serve to highlight the difficulty of mapping 

Blue Ridge formations based solely on lithology. 

– – – – – – – – – – – – – – – – – – – – – – – – – – – – – – – – – – – – – – – – – – – 

Continue west and gradually downhill for 150 m, emerge from the woods to the Skyline Drive and the 

Bearfence Mountain parking area. 



63.0 0.0 Buses will continue north on Skyline Drive. 

63.4 0.4 Bearfence Mountain parking area (mp 56.4) Load buses and 

proceed north on Skyline Drive. 

68.5 5.1 Turn left off Skyline Drive into Big Meadows complex. 

68.6 0.1 Turn right towards parking area for Byrd Visitors Center. 

68.8 0.2 Park in bus parking area. Unload buses and say goodbye. 

End of Road Log 

37

References 

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monazite from Mesoproterozoic granitic gneisses of the northern Blue Ridge, Virginia and Maryland: Precambrian 

Research, v. 99, no. 1, p. 113–146. 

Allen, R.M., 1963, Geology and mineral resources of Greene and Madison Counties: Virginia Division of Mineral 

Resources Bulletin 78, 102 p. 

Allen, R.M., 1967, Geology and mineral resources of Page County: Virginia Division of Mineral Resources Bulletin 81, 

78 p. 

Badger, R.L., 1999, Geology along Skyline Drive, Shenandoah National Park, Virginia: Helena, Mont., Falcon Publishing, 

100 p. 

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. 

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. 

Bailey, C.M., Peters, S.M., Morton, J., and Shotwell, N.L., 2007a, 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. 

Bailey, C.M., Kunk, M.J., Southworth, C.S., and Wooton, K.M., 2007b, Late Paleozoic orogenesis in the Virginia Blue 

Ridge and Piedmont: a view from the south: Geological Society of America Abstracts with Programs, v. 39, n. 1, p. 

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Bailey, C.M., Southworth, S., and Tollo, R.P., 2006a, Tectonic history of the Blue Ridge in north-central Virginia, in 

Pazzaglia, F.A., ed., Excursions in geology and history—Field trips in the Middle Atlantic States: Geological Society 

of America Field Guide 8, p. 113–134. 

Bailey, C.M., Hasty, B.A., and Kayser, C., 2006b, Transverse faults in the Virginia Blue Ridge: a legacy of Mesozoic 

tectonics: Geological Society of America Abstracts with Programs, Vol. 38, No. 3, p. 67. 

Bailey, C.M., Berquist, P.J., Mager, S.M., Knight, B.D., Shotwell, N.D., and Gilmer, A.K., 2003, Bedrock geology of the 

Madison 7.5’ quadrangle, Virginia: Virginia Division of Mineral Resources Publication 157, 22 p. 

Bailey, C. M. and Simpson, C. 1993. Extensional and contractional deformation in the Blue Ridge province, Virginia. 

Geological Society of America Bulletin, v. 105, p. 411-422. 

Balco, G., et al., 2008, A complete and easily accessible mean of calculating surface exposure ages or erosion rates 

from 10Be and 26Al measurements, Quaternary Geochronology, 3, 174-195. 

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42

v 

Jd 

v 

v 

Y1 

Zd 

limestone 

and dolomite 

metasiltstone 

and phyllite 

metasandstone 

and “quartzite” 

Ca 

metasandstone 

and conglomerate 

? ? 

Zp 

Y1 

v 

v 

43 

Tomstown Fm. 

Antietam Fm. 

Harpers Fm. 

Weverton Fm. 

Catoctin Fm. 

Swift Run Fm. 

Basement 

Complex 

PRINCIPAL ROCK TYPES CONTACTS 

v 

v 

v 

v 

v 

Ch 

Zsr 

Zc 

Y3 

phyllitic 

metatuff 

arkosic 

phyllite 

metabasalt/ 

greenstone 

metadiabase 

dike 

diabase 

dike 

granite, 

leucogranite, 

charnockite 

charnockitic 

granitoid 

granitoid gneiss 

charnockitic 

gneiss 

Figure 3. Generalized stratigraphic section of rock formations in the Shenandoah National Park region, Virginia. 

Based in part on Figure 6 from Gathright (1976). Zp- Neoproterozoic garnet-graphite paragneiss. 

Zd 

Ct 

Cw 

Y2 

Chilhowee Group 

MESOPROTEROZOIC NEOPROTEROZOIC CAMBRIAN 

stratigraphic 

or igneous 

significant 

unconformity 

fault

2009 

VGFC 

COc 

38˚ 15’ N 

Cc 

transverse 

fault 

reverse fault 

high-strain 

zone 

COc 

SD 

3 • 

Zsc 

78˚ 45’ W 

Blue 

Ridge 

Q 

• 4 

Roanoke 

Valley & Ridge 

Om 

N 

5 • 

Zsc 

Y 

Washington 

VA 

Piedmont D.C. 

Q 

SD 

Shenandoah 

Elkton 

Cc 

• 

6 

Zsc 

2 • 

Zsc 

Field Trip 

Stop 

• 7 

Y 

Y 

Y 

78˚ 30’ W 

Q 

Stanley 

Zsc 

Y 

Cc 

Y 

Cc 

Cc 

Luray 

COc 

Zsc 

Y 

Y 

Explanation 

Q 

Y 

Big 

Meadows 

SD 

Om 

COc 

Cc 

Zsc 

Zsm 

Zsc 

Syria 

Quaternary sediments 

Silurian-Devonian 

sedimentary rocks 

Ordovician Martinsburg 

Formation 

Cambrian-Ordovician 

carbonates 

Cc 

Early Cambrian 

Chilhowee Group 

Zsc 

Neoproterozoic Catoctin/ 

Swift Run formations 

Zsm 

Neoproterozoic Mechum 

River Formation 

Y 

Mesoproterozoic 

basement complex 

Figure 4. Simplified geologic map of the Shenandoah National Park region and stop locations at the 2009 Virginia 

Geological Field Conference. 

44 

• 7 

• 1 

0 5 10 

38˚ 30’ N 

0 

kilometers 

5 10 

miles 

Y

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