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15 Sequential Development of Platform to Off-platform Facies of the Great American Carbonate Bank in the Central Appalachians

David K. Brezinski Maryland Geological Survey, Baltimore, Maryland, U.S.A.

John F. Taylor Geoscience Department, Indiana University of Pennsylvania, Indiana, Pennsylvania, U.S.A.

John E. Repetski U.S. Geological Survey, Reston, Virginia, U.S.A.

ABSTRACT

In the central Appalachians, carbonate deposition of the great American carbonate bank began during the Early Cambrian with the creation of initial ramp facies of the Vintage For- mation and lower members of the Tomstown Formation. Vertical stacking of bioturbated subtidal ramp deposits (Bolivar Heights Member) and dolomitized microbial boundstone (Fort Duncan Member) preceded the initiation of platform sedimentation and creation of a sand shoal facies (Benevola Member) that was followed by the development of peritidal cyclicity (Dargan Member). Initiation of peritidal deposition coincided with the development of a rimmed platform that would persist throughout much of the Cambrian and Early Ordo- vician. At the end of deposition of the Waynesboro Formation, the platform became subaerially exposed because of the Hawke Bay regression, bringing the Sauk I supersequence to an end. In the Conestoga Valley of eastern Pennsylvania, Early Cambrian ramp deposition was suc- ceeded by deposition of platform-margin and periplatform facies of the Kinzers Formation.

The basal Sauk II transgression during the early Middle Cambrian submerged the plat- form and reinitiated the peritidal cyclicity that had characterized the pre-Hawke Bay depo- sition. This thick stack of meter-scale cycles is preserved as the Pleasant Hill and Warrior Formations of the Nittany arch, the Elbrook Formation of the Great Valley, and the Zooks Corner Formation of the Conestoga Valley. Deposition of peritidal cycles was interrupted during deposition of the Glossopleura and Bathyriscus-Elrathina Biozones by third-order deep- ening episodes that submerged the platform with subtidal facies. Regressive facies of the Sauk

383

Brezinski, David K., John F. Taylor, and John E. Repetski, 2012, Sequential

development of platform to off-platform facies of the great American

carbonate bank in the central Appalachians, in J. R. Derby, R. D. Fritz,

S. A. Longacre, W. A. Morgan, and C. A. Sternbach, eds., The great American

carbonate bank: The geology and economic resources of the Cambrian–

Ordovician Sauk megasequence of Laurentia: AAPG Memoir 98, p. 383 –420.

DOI:10.1306/13331500M983500

II supersequence produced platform-wide restrictions and the deposition of the lower sandy member of the Gatesburg Formation, the Big Spring Station Member of the Conococheague Formation, and the Snitz Creek Formation. Resubmergence of the platform was initiated during the late Steptoean (Elvinia Zone) with the expansion of extensive subtidal thrombolitic boundstone facies. Vertical stacking of no fewer than four of these thrombolite-dominated intervals records third-order deepening episodes separated by intervening shallowing episodes that produced peritidal ribbony and laminated mudcracked dolostone.

The maximum deepening of the Sauk III transgression produced the Stonehenge For- mation in two separate and distinct third-order submergences. Circulation restriction during the Sauk III regression produced a thick stack of meter-scale cycles of the Rockdale Run Formation (northern Virginia to southern Pennsylvania), the upper Nittany Dolomite, the Epler Formation, and the lower Bellefonte Dolomite of the Nittany arch (central Pennsyl- vania). This regressive phase was interrupted by a third-order deepening event that produced the oolitic member of the lower Rockdale Run and the Woodsboro Member of the Grove Formation in the Frederick Valley. Restricted circulation continued into the Whiterockian, with deposition of the upper Rockdale Run and the Pinesburg Station Dolomite in the Great Valley and the middle and upper parts of the Bellefonte Dolomite in the Nittany Arch region. This deposition was continuous from the Ibexian into the Whiterockian; the succession lacks significant unconformities and there are no missing biozones through this interval, the top of which marks the end of the Sauk megasequence.

During deposition of the Tippecanoe megasequence, the peritidal shelf cycles were reestablished during deposition of the St. Paul Group. The vertical stacking of lithologies in the Row Park and New Market Limestones represents transgressive and regressive facies of a third-order deepening event. This submergence reached its maximum deepening within the lower Row Park Limestone and extended into the Nittany arch region with deposition of the equivalent Loysburg Formation. Shallow tidal-flat deposits were bordered to the south and east by deep-water ramp deposits of the Lincolnshire Formation. The St. Paul Group is succeeded upsection by ramp facies of the Chambersburg and the Edinburg Formations in the Great Valley, whereas shallow-shelf sedimentation continued in the Nittany arch area with the deposition of the Hatter Limestone and the Snyder and Linden Hall Formations. Carbonate deposition on the great American carbonate bank was brought to an end when it was buried beneath clastic flysch deposits of the Martinsburg Formation. Foundering of the bank was diachronous, as the flysch sediments prograded from east to west.

INTRODUCTION

The outcrop belts that expose Cambrian and Ordovi- cian strata in the central Appalachians (Figure 1) pro- vide an excellent opportunity to examine the transition from platform to off-platform facies of the great Amer- ican carbonate bank (GACB) of Ginsburg (1982). Plat- form facies are exposed in the Nittany arch region of central Pennsylvania and the Great Valley, which stretches from northern Virginia (Shenandoah Valley) northeastward to eastern Pennsylvania (Lebanon Val- ley) and adjacent New Jersey (Paulinskill Valley). Rapid subsidence in the Pennsylvania depocenter (Read, 1989a, b) resulted in the accumulation of more than 4000 m (>13,120 ft) of carbonate-dominated strata in these out- crop belts. Comparably thick wedges of periplatform stratathataccumulatedincontinentalslopeandcontinen- tal rise environments are preserved in the western parts

of the Conestoga Valley of Pennsylvania and the Fred- erick Valley of Maryland (Figure 1). In the easternmost exposures of these valleys, much of the sequence is con- densed into comparatively thin packages of black shale and shaly limestone that were deposited in sediment- starved basinal environments.

The goal of this chapter is to summarize the results of research conducted during the last two decades on the history of sea level change and derivative sequence stratigraphy and cyclostratigraphy in the deposits of the GACB in the central Appalachian region (i.e., the Pennsylvania depocenter of Read, 1989a, b). We have continued to subdivide this immense stack of Cam- brian and Ordovician platform and periplatform carbon- ates on an increasingly finer scale through an integrated approach best described as a biostratigraphically con- strained event stratigraphy. Recently recovered trilobite and conodont faunas from all of the outcrop belts have

384 Brezinski et al.

expedited recognition and correlation of numerous third-order (grand cycle-scale) transgressive-regressive cycles across the entire platform. Some of these third- order events also affected the style of deposition in off-platform environments and/or in other shallow- marine depositional basins, facilitating correlation into coeval successions in other regions. In many poorly studied regions, however, the biostratigraphic control is sparse and the correlations are offered only as hypoth- eses to be tested as new information becomes available.

The numerical orders of cycles and sequences used in this chapter are intended only as categories of con- venience (see Schlager, 2004) based entirely on scale (thickness and implied duration), with no implication of a specific forcing mechanism or presumption of eu- static origin. Many of the third-order sequence bound- aries are mappable because of the contrast in the un- derlying and overlying lithofacies, typical of flooding surfaces and unconformities. Consequently, many of these horizons have been used as formation or member boundaries to produce more highly refined lithostra- tigraphy both for surface mapping and for correlation in the subsurface. Many of the figures provided here are updated cross sections that show the distribution of lithofacies and correlation of the sequences that they define. Most depict the distribution of sequences de-

lineated in the Sauk megasequence, from which most of our new information were collected. A brief and more tentative treatment of the cycle and sequence stratig- raphy of the Middle Ordovician (Tippecanoe megase- quence) carbonates is provided at the end of the chapter to carry the history of the GACB through to completion with its destruction during Taconic orogenesis.

THE SAUK MEGASEQUENCE

As in other areas of Laurentian North America, the carbonates deposited on and adjacent to the GACB in the central Appalachians constitute much of the Sauk and Tippecanoe sequences of Sloss (1963). In recent studies (Golonka and Kiessling, 2002; Miller et al., 2004), these large-scale first-order sequences or cycles are re- ferred to as megasequences. Palmer (1981) demonstrated that the Sauk megasequence is divisible into three sub- sequences (now supersequences) that he termed Sauk I, Sauk II, and Sauk III, in ascending order (Figure 2). We interpret each of these Sauk supersequences as second-order transgressive-regressive cycles. Read (1989) further subdivided Sauk I and Sauk III into two parts, thereby delineating five transgressive-regressive cycles or sequences within the Sauk megasequence

Figure 1. Outcrop belts (shaded) and specific areas of exposure (white lettering) of Cambrian and Ordovician strata of the great American carbonate bank in the central Appalachians. 50 km (31 mi).

Sequential Development of Platform to Off-platform Facies 385

in this region. The basal sequence (sequence 1) of Read (1989), termed the preplatform shelf, is composed of siliciclastic deposits that accumulated in rift-to-drift shelf and basinal environments on the southern margin of Laurentia before initiation of carbonate deposition during passive margin creation. Hence, the oldest car- bonate strata at the base of sequence 2 (in the middle of the Sauk I supersequence of Palmer, 1981) record the birth of the GACB in the central Appalachians (Cecil et al., 2004).

Sauk I Ramp to Shelf Facies Transition

Recent detailed studies of the Cambrian stratigraphy of the eastern Great Valley in Maryland (Brezinski, 1992) and the Conestoga Valley in southeastern Pennsylva- nia (Taylor and Durika, 1990; Taylor et al., 1997) sup-

port and clarify the major elements of the depositional history in the Early Cambrian reconstructed by Read (1989a) for this region. They bear out the initial devel- opment of a carbonate ramp, followed by evolution of the margin into a high-relief shelf with widespread shale deposition landward of a narrow rim of microbial reefs and ooid shoals. However, new biostratigraphic data, in conjunction with a refined member-level litho- stratigraphy, reveal inaccuracies in some previous inter- pretations and miscorrelation of some units within Pen- nsylvania and Maryland. Detailed mapping facilitated by subdivision of the Tomstown Formation into four members in the Great Valley (Brezinski, 1992) resulted in the discovery of complex structures that had been overlooked in previous studies (Figure 3). The lowest Tomstown strata are made up of finely laminated marble within the basal Bolivar Heights Member

Figure 2. The correlation of groups and formations in the Sauk and Tippecanoe megase- quences in the central Appala- chians. Dol = Dolomite; Fm = Formation; Gp = Group.

386 Brezinski et al.

(Figure 4A). This marble is traceable along the western margin of the Blue Ridge from Pennsylvania to Vir- ginia. This unit, named the Keedysville marble bed, is a mylonite that marks a detachment zone along which the entire Cambrian and Ordovician carbonate stack appears to have been decoupled from the underlying ChilhoweeGroupclasticsduring later Paleozoic orogen- esis (Brezinski, 1992; Brezinski et al., 1996; Campbell and Anderson, 1996). Jonas and Stose (1930, 1944) de- scribed what appears tobe an analogous tectonitewithin the correlative lower Vintage Formation of the Conestoga Valley.

Some formation contacts in or near this detachment zone that had been interpreted as depositional in na- ture (Reinhardt and Wall, 1975; Read, 1989a) are now known to have resulted from juxtaposition of units by faulting. This is significant because cyclic peritidal carbonates reported from the base of the Tomstown Formation by Reinhardt and Wall (1975) led Read (1989) to conclude that shallow peritidal deposition was initiated very early in the deposition of sequence 2 in Maryland, in contrast to more protracted deposi- tion of deep subtidal ramp facies at the base of this sequence in Virginia. The discovery that the cyclic fa- cies directly overlying the Chilhowee clastics in Mary-

land represents the highest member (Dargan Member) of the Tomstown Formation, emplaced along a fault nappe (Brezinski, 1992), eliminates the evidence for a contrast in depositional conditions between Maryland and Virginia, as well as the interpreted facies replace- ment of nearshore clastics (Antietam Formation) by peritidal cycles without an intervening phase of sub- tidal ramp deposition in Maryland. In fact, where non- deformed, the Bolivar Heights and Fort Duncan Mem- bers of the Tomstown Formation in Maryland consist mostly of burrow-mottled, noncyclic, subtidal ramp car- bonates similar to those in the lower part of the Sauk I supersequence (sequence 2 of Read, 1989) in Virginia (Patterson Creek and Austinville Members of the Shady Dolomite) and Pennsylvania (Vintage Formation). The mottled fabric of the Bolivar Heights Member is un- questionably the product of bioturbation (Figure 4B); the origin of fabrics in the dolomite of the overlying Fort Duncan Member is less certain. At least some parts of the Fort Duncan Member display relict fabrics that resemble stromatactoids and fenestrate thrombolitic boundstone (Figure 4C). This suggests that the Fort Duncan Member records the early stages of develop- ment of the microbial reefs that eventually created a narrow carbonate rim at the seaward margin of the

Figure 3. Lateral variations of lithologies and facies and members of the Tomstown Formation along depositional strike in the eastern Great Valley from Virginia to southern Pennsylvania, with derivative relative sea level curve. 10 m (33 ft); 1 km (0.6 mi).

Sequential Development of Platform to Off-platform Facies 387

shale-dominated shelf during deposition of the upper- most part of Sauk I. The light-colored dolomite of the overlying Benevola Member (Figure 4D) strongly re- sembles the Ledger Formation in southeastern Penn- sylvania and the Austinville Member of the Shady Dolomite of Virginia, which accumulated in ooid sand shoals at the very edge of the shelf (Figure 3). Both the Benevola Member and lower Ledger display some relict cross-stratification and are quarried extensively because of their exceptional purity. As previously noted, the overlying Dargan Member is characterized by well-

developed meter-scale peritidal cycles (Figure 4E) similar to those that are ubiquitous in the Middle and Upper Cambrian units in this region. It is likely that the cycles of the Dargan Member formed through pe- riodic shoreward progradation of broad outer-shelf banks similar to those that produced its slightly youn- ger counterparts (Figure 3).

A slightly different style of deposition prevailed on the outer shelf in the latest Early Cambrian when the pure carbonate belt shrank to form a narrow rim only 10 to 15 km (6.2–9.3 mi) wide during deposition of the

Figure 4. The Sauk I ramp and platform lithologies. (A) The laminated tectonite Keedysville marble bed at the base of the Tomstown Formation. (B) The typical burrow-mottled limestone of the of the Bolivar Heights Member ramp facies of the Toms- town Formation. (C) The clotted microbial fabric of the Fort Duncan Member of the Tomstown Forma- tion. (D) The massive, fractured, light-gray dolomitized sand shoal facies of the Benevola Member of the Tomstown Formation. (E) The meter-scale limestone-dolomite cycles (arrows) within the Dargan Member of the Tomstown For- mation. (F) The shaly limestone- dolomite cycles of the Red Run Member of the Waynesboro For- mation. (G) The massive dolo- mitized mud biostrome of the upper Cavetown Member of the Waynesboro Formation. (H) The transition from subtidal limestone to peritidal cycles at the contact between the Cavetown (Cwak) and Chewsville (Cwac) Members of the Waynesboro Formation.

388 Brezinski et al.

shale-rich Waynesboro Formation (Read, 1989a). The Waynesboro Formation is a succession of siliciclastic- rich carbonate strata that was subdivided into three members in the eastern Great Valley by Brezinski (1992): the Red Run, Cavetown, and Chewsville Members, in ascending order (Figure 5). The basal and upper mem- bers (Red Run and Chewsville) are dominated by peri- tidal cycles that consist of carbonates interbedded with red and green siliciclastics. A typical mixed clastic- carbonate meter-scale cycle comprises reddish to light- gray Skolithos-bearing sandstone that grades upward into tan laminated dolomite and red-brown mudcracked siltstone and shale. The intervening Cavetown Mem- ber lacks the clastic components of the other members. This middle member contains massive dolomitic lime mudstone, in packages as much as 20 m (66 ft) thick at the base and near the top of the unit. The remainder of the member consists of peritidal cycles in which gray, thick-bedded, burrow-mottled dolomitic limestones grade upward into tan, laminated, mudcracked dolo- mites (Figure 5).

Within the Great Valley, Lower Cambrianstrata of the Waynesboro Formation (Bonnia-Olenellus Biozone) are directly overlain by middle Middle Cambrian (Glosso- pleura Zone) strata of the Elbrook Formation (Brezinski, 1996a). The absence of the Poliella and Albertella Zones is attributable to the Hawke Bay event (Palmer and James, 1979), a major craton-wide regressive episode that re- sulted in a sizable unconformity between Sauk I and Sauk II in all but the most rapidly subsiding depo- centers (Palmer, 1981) (Figure 5).

Within the Nittany arch of central Pennsylvania, only thin slices of the Lower Cambrian section are ex- posed in the hanging walls of several large thrust faults. The red siltstones and shales resembling the Chews- ville Member of the Waynesboro Formation exposed in that area are the only Lower Cambrian strata exposed west of the Great Valley in the central Appalachian transect.

Sauk I Platform-margin and Off-platform Facies

Although some uncertainty remains regarding a shelf- break origin for the members of the Tomstown For- mation, the formations in the upper part of Sauk I and at the base of Sauk II in the Conestoga Valley undoubt- edly represent shelf-marginal and off-platform envi- ronments (Rodgers, 1968; Gohn, 1976; Reinhardt, 1977; Taylor and Durika, 1990; Taylor et al., 1996; De Wet et al., 2004). Figure 6 shows the lithostratigraphic and biostratigraphic units recognized in the Lower and Mid- dle Cambrian of the eastern Great Valley and Conestoga Valley, along with the biozones established for this in- terval. Figure 7 illustrates the lateral facies relationships

between the carbonate-rich units of the eastern Great Valley and the periplatform and off-platform facies of the Conestoga Valley.

The Vintage Formation, which directly overlies the Chilhowee clastics in the Conestoga Valley, consists of burrow-mottled carbonates similar to those within the Bolivar Heights and Fort Duncan Members of the TomstownFormation(Figure8A).Likethelowermembers

Figure 5. The stratigraphic column of the Waynesboro Formation with relative sea level curve based on stacking of lithologic components. The amplitude and frequency of fourth-order and smaller cycles were approximately portrayed. 50 m (164 ft).

Sequential Development of Platform to Off-platform Facies 389

of the Tomstown Formation, the Vintage Formation ap- parently formed through deposition on a subtidal ramp. The overlying Kinzers Formation, which has no counter- part in the Great Valley succession, has been interpreted as a wedge of periplatform sediment that accumu- lated in slope-to-rise environments (Gohn, 1976; Taylor and Durika, 1990), seaward of the shelf-margin ooid shoals on which the pure dolomite of the overlying Ledger Formation was deposited. Marked changes in lithofacies and profound thickening of the Kinzers Formation from the eastern to western Cones- toga Valley indicate that the platform had evolved into a high-relief constructional rimmed shelf by the time of Kinzers deposition. In the eastern Conestoga Valley, the entire Kinzers is approximately 70 m (~230 ft) thick and comprises a lower member dominated by dark- gray shale (Emigsville Member) (Figure 8B), an un- named middle member consisting of bluish gray mot- tled limestone with shaly interbeds (Jonas and Stose, 1930; Campbell, 1969) (Figure 8C), and a shale-rich upper member (Longs Park member). This relatively thin eastern Kinzers accumulated in a sediment-starved ba- sinal setting. In the western Conestoga Valley, the Emigsville Member triples in thickness to approxi- mately 60 m (~197 ft) and the overlying carbonate part of the Kinzers, referred to here as the upper member, attains a thickness of more than 500 m (>1640 ft) (Ganis and Hopkins, 1990).

A deep-water periplatform origin for the greatly thickened upper Kinzers carbonates is reflected in the

local occurrence of limestone cobble to boulder con- glomerates, which are interpreted as proximal debris- flow or bypass-channel deposits (Figure 8D). In addi- tion, the upper member contains several thin (10–15 m [33–49 ft]) intervals of dark impure limestone with cosmopolitan trilobite taxa found elsewhere only in the toe-of-slope limestone conglomerates in the northern Appalachians (Taylor and Durika, 1990; Taylor et al., 1997). In previous studies (Jonas and Stose, 1930; Ganis and Hopkins, 1990; Taylor and Durika, 1990), attempts to correlate various Kinzers sections throughout the Conestoga Valley via key bed stratigraphy suffered from the false assumption that only a single interval of dark impure carbonate (Upper Kinzers Sandstone of Jonas and Stose, 1930, i.e., Greenmount Member of Ganis and Hopkins, 1990, and Taylor and Durika, 1990) exists at the top of the formation (Figure 8F). Sub- sequent discovery (Taylor et al., 1997) that several such intervals are present, each with a different Lower or Middle Cambrian fauna, required substantial revision of the previous correlation schemes. The current model, showing multiple tongues of deeper water carbonate in the upper member of the Kinzers, is provided as Figure 7. So far, none of the faunas have been recovered from more than one locality, so more such intervals might await discovery. None of the tongues identified so far in the upper member of the western Conestoga Valley has yielded the Middle Cambrian fauna of the Longs Park member in the condensed eastern Kinzers to tie the sections together across the valley.

Figure 6. The revised lithostratigraphy of carbonate units of the Sauk I and Sauk II in the Conestoga Valley and Great Valley and relationship to trilobite biozonation.

390 Brezinski et al.

Only the Emigsville Member at the base of the Kinzers can be confidently correlated across the Conestoga Val- ley, where it plays a significant role as a major aqui- clude in the hydrogeology of the area (Meisler and Becher, 1971). As a dominantly siliciclastic unit, it also represents a significant departure from the primarily carbonate deposition that prevailed in this area from the end of Chilhowee deposition in the Early Cambrian well into the Late Ordovician. The unit remains enig- matic in many respects. Despite its exceptionally pre- served Lower Cambrian fauna in the eastern part of the valley (Dunbar, 1925; Jonas and Stose, 1930; Resser and Howell, 1938; Campbell and Kauffman, 1969; Conway Morris, 1985; Skinner, 2004), its age has not been constrained beyond some part of the very thick Bonnia-Olenellus Zone, which encompasses all the fossil- bearing Lower Cambrian formations in the central Ap- palachians. Lacking any more precise age informa- tion, it is not even clear what horizon or interval in the relatively nearby eastern Great Valley succession (left column of Figure 7) formed at the time that carbonate deposition was suspended in the Conestoga Valley

to create the Emigsville Member. Figure 7 suggests equivalence with the middle of the Tomstown Forma- tion (Benevola Member) and, hence, the initial devel- opment of the rimmed shelf. But this is not well con- strained and any horizon or unit within Sauk I in the Great Valley could be correlated with the Emigsville Member without conflicting with the available bio- stratigraphic data.

Nonetheless, what makes this correlation particu- larly inviting is a strikingly similar stratigraphic suc- cession documented in coeval Lower and Middle Cam- brian platform-margin deposits in southern Virginia (Barnaby and Read, 1990; Betzner and Read, 2009). In that succession, a dark quartzose shaly unit known as the Taylor marker separates the ramp carbonates of the underlying Patterson Member of the Shady Dolo- mite from the overlying periplatform deposits of the upper Shady Dolomite, which formed in front of a high-relief rimmed shelf in the latest Early and ear- liest Middle Cambrian. Although more rigorous evalua- tion of the posited equivalence of the Emigsville Mem- ber and the Taylor marker (and many other plausible

Figure 7. The stratigraphic cross section showing distribution of lithofacies in Sauk I and II in the eastern Great Valley (column A) and the western (column B) and eastern (column C) Conestoga Valley of Pennsylvania. 50 m (164 ft); 40 km (25 mi); m–C = Middle Cambrian; L–C = Lower Cambrian.

Sequential Development of Platform to Off-platform Facies 391

correlations within Sauk I) awaits subdivision of the Bonnia-Olenellus Zone, the virtually identical lithofacies succession within the uppermost part of that zone in the two areas is fairly compelling in itself.

Interpretations of the depositional environment for the Emigsville Member have also varied considerably. In most studies, a base-of-slope or basinal setting was envisioned, invoked by the significant reduction in car- bonate and, perhaps, by association with other occur- rences of exceptional Burgess Shale-type preservation in deep-marine facies. Recently, however, evidence for an alternate interpretation of the Emigsville Member as

a deep-shelf to upper-slope facies, within reach of storm-wave base, has been proposed (Skinner, 2004). The interpretation of the probably coeval Taylor marker as the product of a sea level fall, followed by a short- lived drowning event and backstepping of the margin with inception of the rimmed shelf (Barnaby and Read, 1990), suggests that the Emigsville Member might bear the imprint of both shallow- and deep-water conditions.

In contrast to the break at the Tomstown-Waynesboro contact in the platform carbonates of the Great Valley, no stratigraphic gap attributable to the Hawke Bay event has yet been documented in the western Conestoga

Figure 8. The lithologies within the Sauk I platform edge and periplatform deposits in the Con- estoga Valley. (A) The Keedysville marblelike tectonite at the base of the Vintage Formation. (B) The burrow-mottled fabric character- istic of the Vintage Formation ramp deposits. (C) The dolomitized void fillings in microbrialite(?) Vintage- type section. (D) The typical phyl- litic character of the Emigsville Member of the Kinzers Formation. (E) Thin intervals of off-shelf black limestone (bottom black facies [BB]), interfingering with purer periplatform limestone at Thomas- ville, York County, Pennsylvania. (F) The debris-flow cobble con- glomerate or breccia in periplat- form carbonates of the upper member of the Kinzers (from lo- cation of arrow in photograph E). (G) The burrow-mottled deep- ramp limestone of Kinzers For- mation (upper member) at type section. (H) The thinly bedded silty dolomites of the upper Kinzers Formation, Thomasville, York County, Pennsylvania.

392 Brezinski et al.

Valley. In fact, the discovery in that thick periplatform succession of the Poliella trilobite Zone, an interval generally not present because of Hawke Bay erosion, suggests that little or no interruption in sedimentation occurred during deposition of the uppermost Kinzers in the western Conestoga Valley (Taylor et al., 1997). The biostratigraphic data published for the eastern Con- estoga Valley succession (Campbell 1969, 1971) place the Sauk I-Sauk II boundary somewhere within the greatly condensed Longs Park member at the top of the Kinzers. Barnaby and Read (1990) reported a similar situation in the Shady Dolomite in Virginia.

The Kinzers can be interpreted as a western autoctho- nous facies of the mostly resedimented Conestoga For- mation. The latter unit consists of limestone-shale rhyth- mites and limestone conglomerates. These lithologies formed in off-platform environments seaward of the peritidal bank that rimmed the shelf in the Middle to Late Cambrian during deposition of Sauk II and Sauk III. The conglomerates that characterize the Conestoga Formation are interpreted as proximal to distal debris- flow deposits that formed in a toe-of-slope setting and became interstratified with deep-water hemipelagic rhythmites (Figure 8G).

In the Frederick Valley of Maryland, off-shelf facies assignable to Sauk I are found in the Araby and the lowest Frederick Formations. The Araby Formation (Reinhardt, 1974) crops out along the eastern border of the Frederick Valley and is a sequence of fine-grained bioturbated siltstone and fine-grained sandstone that Reinhardt (1974) correlated with the upper Chilhowee Group of the Great Valley. The Araby Formation is con- formably overlain by the basal member of the Frederick, the Monocacy Member, which consists of a succession of approximately 70 m (~230 ft) of black to dark-gray shale interbedded with rhythmic and brecciated lime- stones and shale (Brezinski, 2004). A Lower Cam- brian fauna that includes Olenellus and the enigmatic conical fossil Salterella has been recovered from the lower part of the Monocacy Member (Reinhardt, 1974; Brezinski, 2004). This fauna indicates equivalence to some parts of the Antietam, Tomstown, and/or Waynesboro Formations of the Great Valley. How- ever, a much younger Middle Cambrian (upper Sauk II) fauna occurs in the highest beds of the Monocacy Mem- ber. Consequently, Brezinski (2004) postulated that the Araby Formation and Monocacy Member of the Frederick Formation represent sediment-starved ba- sinal facies deposited in a deep-ocean floor setting. As in the eastern Conestoga Valley, the Sauk I-Sauk II boundary in the distal off-platform succession in the eastern part of the Frederick Formation lies somewhere within a greatly condensed interval of dark shale and basinal limestone.

Sauk II Platform Facies

The reexpansion of the carbonate platform after the Hawke Bay regression resulted in deposition of hun- dreds of meters of peritidal carbonates on the central Appalachian platform through the Middle Cambrian, most of them packaged in meter-scale cycles. These carbonates constitute sequence 3 of Read (1989), which is equivalent to the Sauk II supersequence of Palmer (1981). The entire succession has been mapped as the Elbrook Formation in the Great Valley and as the Zooks Corner Formation in the Conestoga Valley (Meisler and Becher, 1971; Brezinski, 1996a). In the Nittany arch, Sauk II is divided into two units, the Pleasant Hill and Warrior Formations (Figure 2). Brezinski (1996a) divided the Elbrook Formation into three informal members. The lower member consists of 200 m (656 ft) of tan shaly dolomite, interbedded with thin (5 m [16 ft]), gray, bioturbated lime mudstone intervals. This mem- ber exhibits peritidal cycles that, in most cases, are to- tally dolomitic (Figure 9A). The middle member con- sists of as much as 70 m (230 ft) of mostly noncyclic bioturbated lime mudstone with only a few thin dolo- mitic laminites (Figure 9B). The upper member is 500 m (1640 ft) thick and consists entirely of meter-scale peritidal cycles. A typical cycle in the upper member consists of a thin (<1 m [<3.3 ft]) microbial boundstone or thinly bedded lime mudstone that grades upsec- tion into ribbon lime mudstone, which in turn is capped by a tan laminated dolomite (Brezinski, 1996a) (Figure 9C).

The lower member of the Elbrook Formation spans the Glossopleura and Ehmaniella Zones (Rasetti, 1965; Brezinski, 1996a) (Figure 8). The subtidal facies of the overlying middle member of the Elbrook Formation and the correlative Pleasant Hill Formation in the Nittany arch represent a third-order deepening episode that submerged much of the central Applachians plat- form during deposition of the Bolaspidella Zone. The cyclic deposits of the upper member of the Elbrook Formation and the coeval Warrior Formation of the Nittany arch indicate that shallower peritidal condi- tions returned in the central Appalachians perhaps as early as late in Bolaspidella Zone time and continued throughout deposition of the Cedaria and Crepicephalus Zones (figure 12 of Brezinski, 1996a). This relative sea level history differs from that reconstructed for Sauk II by Read (1989a) (Figure 10), which records maximum deepening later, during deposition of the Crepicephalus Zone. The discrepancy suggests that local tectonics in one or both areas resulted in slightly different timing of the deepest conditions produced during deposition of Sauk II strata in the Pennsylvania and Tennessee depocenters.

Sequential Development of Platform to Off-platform Facies 393

In the Conestoga Valley, Sauk II peritidal facies equivalent to the Elbrook Formation of the Great Val- ley are preserved as the Zooks Corner Formation. This unit consists of meter-scale limestone and dolomite cy- cles identical with those characteristic of the lower member of the Elbrook Formation (Figure 9D).

The regression that created the Sauk II-Sauk III boundary produced an unconformity and a condensed interval of dolomite and sandstone at the base of the Conococheague Formation in the outer-shelf succes- sion in the central Appalachians. In the Great Valley,

the sandy and dolomitic package is mapped as the Big Spring Station Member (Wilson, 1952); in the Lebanon- Conestoga Valley areas, it is referred to as the Snitz Creek Member (Geyer et al., 1963; Palmer, 1971). Both units represent eastern tongues of the cyclically inter- bedded sandstone and dolomite inner-shelf facies that constitutes the bulk of the Gatesburg Formation in the Nittany arch of central Pennsylvania. Specifically, the Big Spring Station and Snitz Creek Members represent the thinned seaward extension of the lower sandy member of the Gatesburg Formation (Wilson, 1952).

Figure 9. The Sauk II lithologies of the central Appalachians. (A) The shaly dolomitic cycles of restricted platform facies of the lower member of the Elbrook Formation. (B) The subtidal plat- form facies consisting of bur- rowed limestone of the middle member of the Elbrook Forma- tion. (C) Thrombolite-based meter-scale cycles (arrows) char- acteristic of the upper Elbrook Formation. (D) High-frequency limestone/dolomite cycles (arrows) of the upper Zooks Corner For- mation of the Conestoga Valley. (E) The light-gray, thick-bedded, coarse-grained dolomites of the Ledger Formation with relict cross-bedding. (F) Shelf-break fenestral grainstone (fg) facies in the Ledger Formation showing meter-scale laminar void filled with internal sediment (is), fibrous marine cement (fmc), and saddle dolomite (sd), along with micro- bialitic boundstone rising as erect masses (em) and descending as pendant masses (pm) from base and top of cavity. (G) The high- angle cleavage (vertical fabric) intersecting limestone and shale rhythmites of the Conestoga Formation. (H) The massive debris- flow breccia of lower and more proximal facies of the Conestoga Formation.

394 Brezinski et al.

The nature and placement of the Sauk II-Sauk III super- sequence boundary are discussed below in the section dealing with Sauk III platform facies.

Sauk II Shelf-edge and Off-shelf Facies

As previously noted, the boundary between Sauk I and Sauk II lies somewhere in possibly, conformable succession near the top of the upper member of the Kinzers Formation in the proximal periplatform

deposits on the western side of the Conestoga Valley, and within the thin Longs Park member at the top of the Kinzers in the thin distal off-platform succession on the eastern side of the valley. Overlying the basal Middle Cambrian strata at the top of the Kinzers in both areas are thick shelf-break deposits of the Ledger Formation; these are in turn overlain by basinal facies assigned to the Conestoga Formation. The very few fossils that have been recovered from strata that are either Ledger or Conestoga assign these units to the Middle Cambrian, and hence, to Sauk II. The nature of the contact between the Kinzers and the Ledger in the few areas where this boundary has been exposed suggests a conformable relationship; the contact marks a transition from shaly upper-slope facies to perva- sively dolomitized shelf-break carbonate sands. At Longs Park in the eastern Conestoga Valley, the contact is within thin interbeds of Ledger-like dolomite occur- ring in the highest part of the Longs Park member. In cores and quarry exposures of the western Conestoga Valley, large flame structures and other features attri- butable to loading of the muddy sediment of uppermost Kinzers by shelf-edge sands of the Ledger character- ize the boundary. The Kinzers-Ledger contact, there- fore, apparently records the progradation of shelf mar- gin very early during the deposition of Sauk II.

In contrast, the contact between the Ledger and over- lying Conestoga appears everywhere sharp and is prob- ably unconformable, recording a backstep of the shelf edge later in Sauk II deposition. Recovery of a trilobite fauna tentatively assigned to the Middle Cambrian Bolaspidella Zone from low in the Conestoga Forma- tion (Taylor et al., 1997) suggests that this retreat might be attributable to the transgression that resulted in de- position of the middle member of the Elbrook Forma- tion, which records maximum deepening in the Great Valley during deposition of Sauk II. Gohn (1976), in the most detailed study of the Conestoga to date, reports that massive toe-of-slope limestone conglomerates characterize the base of the formation and give way upsection to more distal basin-margin facies dominated by limestone-shale rhythmites (Figure 7, right column; Figure 8H, G). The formation includes an even more distal, condensed, basinal, black shale facies in the south- easternmost exposures of the Conestoga Valley. De- positional models and stratigraphic cross sections in virtually all studies since the work of Rodgers (1968) depict the Conestoga Formation, or a unit of similar lithologic character, as accumulating seaward of the central Appalachian shelf margin throughout the Cam- brian and Early Ordovician. This facies pattern appears valid for Sauk II and for most of the Upper Cambrian, in as much as the Frederick Formation in the Frederick Valley comprises virtually identical basin-margin facies.

Figure 10. The composite stratigraphic column of the Elbrook Formation in the central Appalachians, with inferred relative sea level curve modified from Brezinski (1996a). The amplitude and frequency of fourth-order and smaller cycles were approximately depicted. 100 m (328 ft).

Sequential Development of Platform to Off-platform Facies 395

However, it is difficult to argue that such a carbonate- rich basin-margin facies was deposited seaward of the shelf margin during deposition of the uppermost Kinzers when the starved basin facies of the Longs Park member was forming.

Several recent studies (Ganis and Hopkins, 1990; Taylor and Durika, 1990; Taylor and Krawiec, 1993; De Wet et al., 1999, 2004, 2012) provided new information regarding the original lithofacies and internal stratig- raphy of the Ledger Formation based on drill cores and new quarry exposures in the western Conestoga Valley. Ganis and Hopkins (1990) proposed a tripartite member stratigraphy for the formation, separating the pure dolomite packages at the top and base of the for- mation, with a middle member to which they assigned the name Willis Run Member. Some of the quarry ex- posures that originally were assigned to this middle member have since (Taylor et al., 1997) been re-inter- preted as part of the Kinzers Formation. However, much of the limestone that formed the basis of Ganis and Hopkin’s (1990) Willis Run Member is without a doubt a non-dolomitized part of the Ledger Formation that locally produces the useful tripartite subdivision pro- posed by those authors. Given the local nature of the limestone package within the Ledger, and difficulty of precise correlation between exposures in the western Conestoga Valley, werefer herein to the nondolomitized portion of the Ledger (fenestral grainstone in Figure 7) as the Willis Run facies, rather than member.

The basal part of the Ledger Formation, directly underlying the Willis Run facies, is exceptionally pure refractory-grade dolomite that owes its low insoluble content (at least in part) to deposition in well-washed shelf-break oolite shoals. Relict cross-stratification and oolitic fabric are discernible in parts of the unit (Figure 9E). The dolomite package that overlies the Willis Run facies is similar, but of slightly lesser purity. The Willis Run facies is predominantly limestone, and the fabrics preserved are quite unusual. The dominant lithology is a dark well-winnowed grainstone to boundstone with abundant millimeter- to centimeter- size fenestrae. Also present are flattened decimeter- to meter-size voids filled with a complex array of internal sediment, fibrous marine cements, and the remains of a thriving community of cavity-dwelling microbes, specifically, dendritic Renalcis. At a much later stage, coarse baroque and/or saddle dolomite occluded the last remaining pore space in most of the large voids (Figure 9F) (see article by De Wet et al., 2012).

The stratigraphic-paleogeographic context of this unusual facies (Taylor and Krawiec, 1993) and its sedimentological geochemical attributes (De Wet et al., 1999; 2004; 2012) leave no doubt that it formed at the very edge of the rimmed shelf to upper slope. Its un-

usual features reflect extensive syndepositional ce- mentation of the subtidal carbonate sands and peri- odic buckling of the cemented surficial sediment. This created the large flattened cavities as the unstable shelf- break sediment shifted in response to stresses gener- ated by storms and/or seismic tremors. The Willis Run shelf-break facies contains very few fossils, but a few trilobites were recovered and identify this part of the Ledger as lower Middle Cambrian Glossopleura Zone (Taylor et al., 1997). Intriguingly, the only other known occurrence of this distinctive shelf-margin facies is also from the lower Middle Cambrian, spe- cifically the next older Albertella Zone of the Cathedral Formation, in western Canada (Pratt, 2002). Like the Willis Run facies, the interval of the Cathedral Forma- tion characterized by the buckled and shingled fenes- tral crusts, that is, teepees in the broad sense, lies atop and shoreward of an unequivocal shelf-break ooid shoal facies that fringes the Cathedral escarpment (figure 2 of Pratt, 2002).

A very thin interval of off-platform Sauk II strata is preserved within the upper part of the Monocacy Mem- ber of the Frederick Formation. As previously noted, this 70-m (230-ft)-thick shaly unit contains a Lower Cambrian Olenellus fauna at its base and an upper Mar- juman Crepicephalus Zone (highest Middle Cambrian) fauna at the top. Therefore, although more than 700 m (>2297 ft) of peritidal sediments accumulated to create the Elbrook Formation of the Great Valley and the Pleas- ant Hill and Warrior Formations of the Nittany arch, only about 70 m (~230 ft) of black shale and alloc- thonous carbonate of the Monocacy Member of the Frederick Formation were deposited in the sediment- starved condensed basinal setting to the east (Brezinski, 1996b, 2004).

Sauk III Platform Facies

Upper Cambrian Sauk III deposits are represented in the central Appalachians by the Gatesburg Formation in the Nittany arch, and the Conococheague Forma- tion or Group of the Great Valley and Conestoga Val- ley (Figure 2). Both formations are strongly cyclic, comprising vertically stacked facies associations that record waxing and waning of sea level and constitute third-, fourth-, and fifth-order depositional sequences (Figure 11). Deposition of peritidal cyclic facies was interrupted in the earliest Ordovician by an expan- sive submergence and increase in subsidence that produced the Stonehenge Formation, the basal unit of the Beekmantown Group (Taylor et al., 1992; Read, 1989a). This event, the Stonehenge transgression (Taylor et al., 1992), is arguably the transgressive peak of

396 Brezinski et al.

the Sauk III supersequence (Brezinski et al., 1999) (Figure 12).

Cyclic peritidal deposition resumed after accumu- lation of the Stonehenge and remained the dominant style of deposition throughout the regressive phase of the Sauk III supersequence, which is represented by the remainder of the Beekmantown Group. However, deepening sufficient to prevent the development of meter-scale cycles occurred at the peaks of at least three third-order cycles during deposition of the remainder of the Beekmantown Group, facilitating geologic mapping and correlation of Lower Ordovician units across the central Appalachians and into other sedimentary basins.

In studies of the uppermost Cambrian and basal Ordovician strata in New York (Landing et al., 2003; Landing and Westrop, 2006) and southern Virginia and Tennessee (Bova and Read, 1987; Read, 1989a; Pope and Read, 1998), sandstone-rich intervals and miss- ing biozones near the base of the Ordovician formed

the foundation for treatment of the Cambrian and Ordovician parts of Sauk III as separate superse- quences. Deposition of a more complete succession in the Pennsylvania depocenter renders precise placement of a supersequence boundary near the Cambrian- Ordovician boundary in the central Appalachians more problematic. Sandstone-rich intervals interpreted as regressive peaks occur at several different levels within the highest Cambrian, and no unequivocal gap in the biostratigraphy has been documented at any of them. Nonetheless, the lithostratigraphic units and deposi- tional sequences in this area do lend themselves to a separate discussion of the depositional history of the Upper Cambrian and Lower Ordovician Series; they are treated accordingly in the following sections. More detailed information on the ambiguity of the data and difficulty of identifying one horizon (or thin interval) as a supersequence boundary that divides Sauk III into two parts in this area is provided in the section on the Lower Ordovician platform facies.

Figure 11. The composite strati- graphic section of the Conoco- cheague Formation of the central Appalachians illustrating third- order sequences and ranges of species of the trilobite genus Plethopeltis. The amplitude and frequency of interpreted fifth- order sea level units were ap- proximately depicted. 50 m (164 ft).

Sequential Development of Platform to Off-platform Facies 397

Upper Cambrian Sauk III Platform Facies

Third-order transgressions and regressions altered the style of deposition on the platform several times dur- ing the Late Cambrian. This is reflected to some degree within the member stratigraphy of the Gatesburg For- mation in Pennsylvania, but the very poor exposures and scarcity of fossils in that unit severely limit the extent to which the history of Late Cambrian sea level can be reconstructed from that inner-shelf succession. In contrast, the Conococheague Formation provides more numerous fossiliferous horizons, and the unit is almost completely exposed in cliffs along the Chesapeake and Ohio (C&O) Canal along the north- ern side of the Potomac River in the Great Valley (Figure 11).

A significantly improved trilobite-based zonation within the thick (600–900 m [1970–2950 ft]) peritidal bank-margin deposits of the Conococheague has re- vealed the lateral persistence, across tens of kilometers,

of four third-order transgressive cycle peaks within the formation (Figure 12). Each transgression provided sufficient accommodation space to allow for deposition of a complex of thick microbial (thrombolitic) reefs with intervening and interbedded trilobite-bearing grain- stone and noncyclic ribbon carbonate (Figure 13). These transgressive (TSTs) to early highstand (HSTs) systems tracts alternate with thick intervals of cyclically inter- bedded planar microbial laminite, deep prism- cracked laminites, and minor ribbon carbonate with thin (<1 m [<3.3 ft]) discontinuous microbial reefs that are interpreted as late highstand to lowstand systems tracts (LSTs) (Figure 13). These regressive intervals correspond to the cyclic facies of Demicco (1985). Each couplet, comprising a TST to early HST reef-grain- stone package combined with the cyclic or laminite late HST to LST that gradationally overlies it, constitutes a third-order sequence. These sequences range from approximately 110 to 150 m (�360–490 ft) in thickness. As is typical for carbonate successions that

Figure 12. The lateral extent and distribution of thrombolite intervals in measured sections of the Conococheague Formation along depositional strike in the central Appalachians. Thrombolitic intervals are interpreted as third-order deepening episodes. Section A, Waynecastle Dairy; section B, Chesapeake and Ohio (C&O) Canal; milepost 108w, section C, Cooke Pasture; section D, Sheetz Farm. P = Plethopeltis. 10 km (6.2 mi); 50 km (31 mi).

398 Brezinski et al.

accumulated on broad flat-topped platforms (Montanez and Osleger, 1993), the boundaries between these se- quences are not surfaces, but thin (5–10 m) zones. In the Conococheague, these sequence boundary zones (SBZs) (Montanez and Osleger, 1993), equivalent to the minimium accommodation zones of Lehrmann and Goldhammer (1999), are characterized by numer- ous thin cycles dominated by the shallowest litho- facies, specifically dolomitic laminite and sandstone. For a thorough treatment of the Conococheague cy- cles, including their constituent lithofacies, variability, and relationship to sea level, see Koerschner and Read (1989).

For convenient reference, we have numbered the thick reef or grainstone packages (maximum accom- modation zones of Lehrmann and Goldhammer, 1999) within the Conococheague from bottom to top as throm- bolites I to IV (Figures 11–13). For a detailed descrip- tion of the fine-scale sequence stratigraphy and fauna of one of these packages (thrombolite III), see Taylor et al. (2009). Although the individual thrombolitic in- tervals maintain a fairly consistent thickness along de-

positional strike (Figure 12), they exhibit a dramatic increase in thickness from west to east (Figure 13; columns B and C). This pronounced thickening of in- dividual thrombolitic intervals and a concurrent re- duction in the prominence of the intervening prism- cracked dololaminite-rich cyclic facies clearly reflect significantly more rapid subsidence of the outermost part of the outer-shelf bank.

The shallowest and most restricted conditions re- corded within the Upper Cambrian in the Great Val- ley are those represented by the basal 70 to 90 m (230– 295 ft) of the Conococheague, a package of dolomite and sandstone set apart by Wilson (1952) as the Big Spring Station Member of that formation (Figure 14A). The same package was identified as the Snitz Creek Member at the base of the Conococheague Group in the Conestoga Valley (Geyer et al., 1963; Palmer, 1971). This regressive interval was created at some point dur- ing the transition between Sauk II and Sauk III and represents a thick SBZ. However, lithologic and geo- chemical evidence recently acquired from sections in Maryland and southern Pennsylvania confirms that this

Figure 13. The distribution of Upper Cambrian facies across depositional strike in the central Appalachians, from the Nittany arch to Frederick Valley. Section A, Nittany arch; B, Chesapeake and Ohio (C&O) Canal, milepost 108w; C, C&O Canal, milepost 83; D, Frederick Valley (modified from Brezinski, 2004). P = Plethopelis. 10 km (6.2 mi); 100 m (328 ft).

Sequential Development of Platform to Off-platform Facies 399

SBZ is marked at its base by an unconformity (see be- low). The lowstand during Sauk II–Sauk III regression occurred during deposition of the basal Upper Cam- brian (Steptoean Stage) Aphelaspis and Dunderbergia Zones in other areas of North America. These zones are either represented by the barren dolomite and sandstone of the Big Spring Station Member, or they are absent from the platform succession in the central Ap- palachians, their corresponding chrons falling within the lacuna of an unconformity at the base of the Conococheague.

Previous syntheses of the Cambrian in this region (Palmer, 1971; Read, 1989a, b) show the Sauk II-Sauk III

boundary near the middle of the Big Spring Station Member, based on reports (Wilson, 1952; Rasetti, 1961) of trilobites characteristic of the Crepicephalus Zone (highest zone of the Marjuman Stage at the top of the Middle Cambrian by modern usage; previously part of the Dresbachian Stage at the base of the Upper Cam- brian) from the lower half of the member south of Waynesboro, Pennsylvania. Recent work on the Elbrook- Conococheague boundary interval, where it is contin- uously exposed along the C&O Canal (Brezinski, 1996a; J. F. Taylor, unpublished data), has revealed that the upper member of the Elbrook Formation includes a few thin intervals with abundant quartz sand tens of

Figure 14. The Sauk III platform lithologies. (A) The restricted- circulation platform facies of the Big Spring Station Member of the Conococheague Formation. (B) The thrombolite bioherm within the middle part of the Con- cocheague Formation. (C) The ribbony with deep prism cracks dolomite within the regressive fa- cies of the Conococheague For- mation. (D) The ribbony texture of subtidal thinly bedded limestone of the basal Stoufferstown Mem- ber of the Stonehenge Limestone. (E) The massive thrombolite with flanking strata in the middle member of the Stonehenge Lime- stone. (F) The ribbony strata of the upper member of the Stone- henge. (G) High-frequency lime- stone or dolomite shallowing cycles (triangles) of the lower Rockdale Run Formation. (H) Do- lomitic shallowing cycles (trian- gles) within the Pinesburg Station Dolomite.

400 Brezinski et al.

meters below the sharp contact with the overlying Conococheague. Unaware of the thin sandy zones in the Elbrook Formation, Wilson (1952) mistakenly placed the formation contact in the poorly exposed section south of Waynesboro at the lowest occurrence of quartz sand. The true formation contact lies roughly 30 m (�98 ft) higher in that section at a sharp bedding plane bound- ary where thrombolite-based limestone-dolomite cy- cles of the upper member of the Elbrook Formation are overlain by tens of meters of sandstone and dolo- mite of the Big Springs Station Member. Consequently, all of the horizons from which the Crepicephalus Zone faunas were reported in previous studies lie within the Elbrook Formation; no trilobites or other diag- nostic fossils have yet been recovered from the Big Spring Station Member.

Moreover, the carbonate-carbon isotopic profile gen- erated in a detailed study of the isotope stratigraphy through the section south of Waynesboro (Glumac, 2002) reveals a sharp positive shift of 2% precisely at the true base of the Big Spring Station Member; this strongly supports the interpretation of the sharp con- tact as an unconformity. In addition, the relatively high values (>2%) obtained from the strata above the sur- face effectively confirm that they are Steptoean in age. The highest Marjuman strata of the Crepicephalus Zone range between 0 and +1%, well below the values ob- tained from the basal Conococheague strata. The case is now strong, therefore, that the sharp contact be- tween the Elbrook and Conococheague is the uncon- formable boundary between the Marjuman strata at the top of Sauk II and the Steptoean strata at the base of Sauk III. However, the isotopic values alone can- not resolve whether the Big Spring Station Member was deposited early in the Steptoean when Aphelaspis and/or Dunderbergia Zone strata were being depos- ited elsewhere, or whether it accumulated late in the Steptoean during deposition of some part of the Elvinia Zone. Both intervals are characterized by values about +2%, so the lacuna might include only a small part of the Steptoean, or it might include well more than half of that age.

These findings also have implications for the place- ment of the Sauk II-Sauk III boundary in the Nittany arch. With recognition of the base of the Conococheague as the unconformable supersequence boundary mark- ing the point of maximum regression, the equivalent horizon in the Nittany arch is most likely the base of the lower sandy member of the Gatesburg Formation, where sandstone and dolomite cycles replace the pure dolomite of the underlying Stacy Dolomite Member.

After deposition of the Big Spring Station and lower sandy members, the Sauk III transgression produced a backstepping of the mixed sandstone-dolomite cy-

clic facies westward. The result was a spreading of pure subtidal carbonates across the entire central Ap- palachian shelf during deposition of the middle and upper parts of the Elvinia Zone (Wilson, 1952). The re- sult in the Great Valley was deposition of thrombolite I, the lowest and thickest of the third-order transgressive cycle peaks (HST) in the Late Cambrian (Figure 14B). This thick reef and grainstone package, which con- tains thrombolitic boundstone intervals as much as 5 m (16 ft) thick, is the basal facies of the middle member of the Conococheague Formation in Maryland (Bell, 1993) and the lowest component of the Zullinger Formation of the Conococheague Group in adjacent Pennsylvania (Root, 1968). Concurrent deposition in the Nittany arch produced the reef-rich lower half of the Ore Hill Mem- ber of the Gatesburg Formation (Wilson, 1951, 1952; Loch and Taylor, 1995, 2004; Taylor et al., 1999).

After flourishing throughout deposition of the Elvinia Zone, microbial reefs disappeared entirely during depo- sition of the overlying Taenicephalus Zone (Figure 14C). High-resolution sampling of the Ore Hill Member (Loch and Taylor, 1995, 2004; Taylor et al., 1999) has estab- lished that the reef disappearance does coincide pre- cisely with this zonal boundary, which is also the boundary between the Pterocephaliid and Ptychaspid biomeres (Taylor, 2006). At this horizon, the top of the Irvingella major Subzone of the Elvinia Zone and base of the Parabolinoides Subzone of the overlying Taenicephalus Zone, platform trilobite diversity is ex- tremely low, with a nearly monogeneric fauna and, effectively, no biofacies differentiation (Westrop and Cuggy, 1999) across the great American carbonate bank. Apparently, the environmental factors that erad- icated the few taxa of the Pterocephaliid biomere that had survived the stage-boundary extinctions at the end of the Steptoean (base of the I. major Subzone), reducing platform trilobite diversity to its minimum in the Parabolinoides Subzone, also inhibited microbial reef development in some way. These effects were not lim- ited to the Appalachian shelf. The destruction of outer- shelf microbial reef complexes at the Pterocephaliid- Ptychaspid biomere in Wyoming (Saltzman et al., 1995) and a total absence of reefs from the Parabolinoides Sub- zone in Alberta (Westrop, 1986) indicate that the reef growth was inhibited across a much larger part of the Laurentian platform than just the central Appalachian region.

Reefs gradually recovered in the affected areas, re- appearing first as small (submeter-scale) isolated mounds and only later forming large complexes again near the middle of the Lower Sunwaptan Substage. In the Great Valley, the full recovery came about with the deposition of the thrombolite II maximum accommo- dation zone. The discovery of Ptychaspis granulosa very

Sequential Development of Platform to Off-platform Facies 401

near the base of thrombolite II dates that transgres- sion as a medial Early Sunwaptan event (Taylor et al., 2007), equivalent in age to middle Ptychaspis-Prosaukia Zone strata in the upper Mississippi Valley (Nelson, 1951) and the base of the Ellipsocephaloides Zone in Alberta, where reefs once again occur in profusion (Westrop, 1986). Apparently, the conditions that pro- moted microbial reef growth to produce thrombolite II on the southern Laurentian margin triggered a simi- lar development of thrombolitic reefs on the northern shelf of the paleocontinent.

Trilobites and conodonts from within and just be- low thrombolite III suggest that the reef complex(es) created by the third Late Cambrian transgression spans the boundary between the Lower and Upper Sunwaptan Substages (i.e., the base of the Illaenurus Zone in western North America) (Ludvigsen and Westrop, 1985; Westrop, 1986). The basal reefs in this interval yield Plethopeltis saratogensis and other species of the uppermost Lower Sunwaptan Saratogia Zone, where- as the upper part of the package contains Plethopeltis stitti, a common component of lowest Saukia Zone (basal Upper Sunwaptan) faunas in Oklahoma (Stitt, 1971). To the east in the Frederick Valley of Maryland, new bio- stratigraphic data in coeval shelf-break and off-platform deposits (Grove and Frederick Formations) reveal profound changes in formation thickness over short distances both across and along depositional strike. Strike-parallel thickening of some units provides strong evidence for an embayed platform margin in that area. Among the more significant changes in style of deposition recorded in the Frederick Valley succes- sion is that observed at the base of the Adamstown Member of the Frederick Formation, where coarse gravity-flow deposits are overlain by dark shale and lime mudstone. The timing and extent of this facies change, which most likely resulted from the thrombo- lite III transgression, are discussed later in the Sauk III shelf-break and off-platform facies section. For an ex- panded treatment of this event, designated the Adams- town submergence event, see Taylor et al. (2009).

A relatively rich trilobite collection from thrombo- lite IV containing species of Euptychaspis, Conococheaguea, and Prosaukia establishes that this highest transgres- sive cycle peak in the Conococheague is medial Late Sunwaptan in age, equivalent to strata elsewhere as- signed to the Saukiella junia Subzone of the Saukia Zone.

Lower Ordovician Sauk III Platform Facies

Overlying the Conococheague Formation is the Stone- henge Formation, an interval as much as 300 m (984 ft) thick, composed predominantly of limestone. In the

Great Valley, its type area (Stose, 1908; Sando, 1958), the Stonehenge is subdivided into three members. The basal Stoufferstown Member consists of 50 to 70 m (164–230 ft) of thinly bedded to ribbony, medium- to dark-gray, locally intraclastic limestone. Individual beds are separated by thin, wispy, dolomitic or shaly laminae (Figure 14D). Unlike the ribbony intervals within the Conococheague, the Stoufferstown Mem- ber contains no finely laminated dolomite prism or mudcracked laminites, or any of the other intertidal to supratidal lithologies typical of meter-scale peritidal cycles. Noting this absence, Nguyen (1986) interpreted the noncyclic facies of the Stoufferstown Member as having been deposited under deeper subtidal condi- tions. Using trilobites and conodonts from just below and above the base of the Stoufferstown Member to date and correlate the third-order transgression re- corded by the base of the Stonehenge, Taylor et al. (1992) identified formation and member boundaries of the same age in various areas of Laurentian North America, all of them reflecting submergence by what they named the Stonehenge transgression. They re- ported that faunas and lithofacies in widely scattered sedimentary basins from North Greenland to Texas and Utah record a rapid transgression during depo- sition of the middle of the Symphysurina trilobite Zone and within the relatively thin Cordylodus angulatus co- nodont Zone. Evidence from more recent studies has supported this claim, adding to the list of deeper or more offshore facies introduced by the Stonehenge transgression the basal strata of the Oneota Dolomite atop the Jordan Sandstone in the upper Mississippi Valley (Runkel et al., 2007), and even a thin interval of graptolitic shale within the condensed inner-shelf glaucarenite facies of the Bliss Formation in southern New Mexico (Taylor and Repetski, 1995; Taylor et al., 2004). Although the label of Stonehenge transgression was subsequently adopted in a few later studies (Miller et al., 2004; Nielsen, 2004) for a longer term deepening trend, we herein retain the original, more restrictive definition as a rapid third-order submer- gence in the earliest Ordovician (C. angulatus Zone) and do not consider any subsequent third-order trans- gressions to be a part of it. This transgressive peak was also identified by James et al. (1989) as event 6 in their synthesis of the early Paleozoic development of the northern Appalachian shelf in eastern Canada. In that area, the Stonehenge transgression is recorded in the Watts Bight Formation at the base of the St. George Group. In their discussion of the event, James et al. (1989) argued for a eustatic origin, noting the consis- tency in the sedimentologic response between the St. George Group in Newfoundland (Knight and James, 1987) and the Beekmantown Group in Virginia (Bova

402 Brezinski et al.

and Read, 1987). The Stonehenge transgression corre- sponds to the third-order event identified as O-2 in the sea level curve provided by Read (1989a) (Figure 10). Although his placement of this event within the bio- stratigraphic and lithostratigraphic framework is ac- curate, the magnitude of the transgression is not ade- quately portrayed.

Because of very rapid subsidence, evidence in the Pennsylvania depocenter for a significant sea level drop before the Stonehenge transgression is equivocal (Taylor et al., 1992). But unconformities and regres- sive units occur at precisely this level in many other basins; the unconformity and overlying arenaceous Sprakers Member at the base of the Tribes Hill For- mation in eastern New York (Landing et al., 1996; Landing et al., 2003), the Permin Land Formation in North Greenland (Bryant and Smith, 1990), the Jordan Sandstone in Minnesota and Wisconsin (Runkel et al., 2007), the Gunter Sandstone Member at the base of the Gasconade Dolomite in the Ozarks region (Kurtz, 1981; Thompson, 1991), and the unconformity at the base of the Garden City Formation in Utah and Idaho (Taylor and Repetski, 1985) all seem to indicate that a significant drawdown preceded the Stonehenge transgression.

As previously noted, the preserved signal of sea level change through the Cambrian-Ordovician boundary interval in the Pennsylvania depocenter is more com- plicated and ambiguous than in New York (Landing and Westrop, 2006) and southern Virginia (Bova and Read, 1987; Read, 1989a), where a specific prominent sandstone-rich interval with an associated biostrati- graphic break have been used to place the systemic boundary. Moreover, the ages of the sandy intervals in New York and Virginia differ. Conodonts from the Chepultepec Formation in Virginia (Bova and Read, 1987) indicate that the sandstone at the base of that unit, the horizon identified by Read (1989a) as the super- sequence boundary that separates the lower (Cam- brian) and upper (Ordovician) divisions of Sauk III, corresponds closely to the base of the Ibexian Series, which was defined at the base of the Cordylodus proavus conodont Zone (Ross et al., 1997) (i.e., conodont fauna A of Ethington and Clark, 1971). This zonal boundary served for some time as the base of the Ordovician System in North America (Ross et al., 1997) before rat- ification of the base of the younger Iapetognathus cono- dont Zone as the internationally recognized Cambrian- Ordovician boundary (Cooper et al., 2001). The Sprakers Member at the base of the Tribes Hill Formation in New York lies unconformably atop strata at the very top of the C. proavus Zone and is overlain by limestone that yields a Lower Ordovician Rossodus manitouensis conodont Zone fauna (i.e., fauna C of Ethington and

Clark, 1971). Thus, the drawdown that produced the Sprakers Member at the Cambrian-Ordovician bound- ary in New York significantly postdated the regres- sion responsible for deposition of the sandstone late in the Cambrian in Virginia and Tennessee.

A thin (5–10 m [16–33 ft]) SBZ of sandstone-capped cycles precisely correlative with the basal Ibexian sand- stone in Virginia does occur high in the Conococheague Formation in northern Virginia, but it is not unique. A comparable package occurs approximately 100 m (328 ft) lower in the formation low in the Saukia trilo- bite Zone and the upper Proconodontus tenuiserratus to lower Proconodontus posterocostatus conodont Zone (J. F. Taylor and J. E. Repetski, unpublished data). Orndorff (1988) also reported sandstone from very high in the Conococheague Formation in association with cono- donts typical of the Cordylodus intermedius Zone. This somewhat younger sandy interval might represent the same regressive episode as the Sprakers Member of the Tribes Hill Formation, but the biostratigraphic data available from the Conococheague lack the precision necessary to test that hypothesis. The central Appala- chian carbonate succession, therefore, includes multi- ple sandy intervals whose relative significance cannot be assessed with the available data to allow selection of one as representing the point of maximum draw- down. Athough some conodont subzones originally were thought to be missing from the SBZ that spans the base of the Ibexian in the Conococheague (Orndorff, 1988), additional sampling confirmed that these units actually are present (Taylor et al., 1992).

Although the point of maximum shallowing in the uppermost Cambrian and basal Ordovician in the cen- tral Appalachians cannot be determined with precision, the maximum flooding interval is clearly represented by the Stoufferstown Member of the Stonehenge For- mation. The facies and the faunas of that unit confirm that maximum deepening occurred during deposi- tion of the highest beds of the C. angulatus Zone and/ or basal strata of the R. manitouensis conodont Zone, an interval correlative with the lower Tribes Hill For- mation (just above the Sprakers Member) in New York (Landing et al., 1996) and the middle member of the Chepultepec Formation in Virginia (Bova and Read, 1987). Overlying the Stoufferstown Member in its type area, the Hagerstown Valley, is a middle member, as much as 150 m (492 ft) thick and characterized by mas- sive to thickly bedded thrombolitic reefs with steep- sided, coarse grainstone-filled channels (Figure 14E). Sando (1957) informally designated this massive unit the lower member and used its base as the contact with the underlying Conococheague, as Stose (1908) had done. At that time, he opted to arbitrarily assign the distinct package of noncyclic thinly bedded strata

Sequential Development of Platform to Off-platform Facies 403

that ultimately became the Stoufferstown Member to the top of the Conococheague. However, on recover- ing Ordovician trilobites from that package, he chose to reassign it to the base of the Stonehenge, naming it the Stoufferstown Member (Sando, 1958). Sando (1958) concurrently referred to all Stonehenge strata above the Stoufferstown Member as the upper member, the label he had used previously (Sando, 1957) for the interval of thinly bedded limestone at the top of the Stonehenge and above the massive thrombolitic lower member. To avoid confusion, we retain the original use of upper member and refer to the reefal package (the original lower member) as the middle member.

The boundary between the Stoufferstown Member and middle member is conformable and gradational, recording the shallowing out of the deepest conditions introduced by the Stonehenge transgression. Conse- quently, those two members together constitute a sin- gle third-order sequence. Individual bioherms as much as 15 m (49 ft) thick dominate the lower half of the mid- dle member. The reefs become progressively thinner, and the relative abundance of interstratified thinly bedded to ribbony wackestone increases upward through the member. A few thin (<0.5 m [<1.6 ft]) do- lomite beds occur very near the top of the middle bio- hermal member (Sando, 1957). The relatively sharp boundary of the middle member with the overlying upper member, which consists of thinly bedded, non- cyclic, subtidal limestone (Figure 14F) very similar to that of the Stoufferstown Member, records the second Early Ordovician third-order transgression. The upper member, approximately 100 m (�328 ft) thick, is dom- inated by ribbony and intraclastic limestone inter- stratified with intervals of oolitic grainstone. Although present, thrombolites are scarce and much thinner (typically <1 m [<3.3 ft]) in the upper member. The oolitic intervals characteristic of this member can at- tain a thickness of 5 m (16 ft). The gradational contact with the overlying Rockdale Run Formation repre- sents the return of meter-scale cyclicity, which is marked by the reappearance of tan laminated dolo- mite cycle caps.

The three members recognized within the Stone- henge Formation in the type area are also recogniz- able in the inner-shelf succession of the Nittany arch, although the thickness and character of each differ slightly because of greater proximity to the paleoshore- line, less rapid subsidence, and pervasive later stage dolomitization in some areas. In many locations, the interval above the Mines Formation Member of the Gatesburg Formation and below the Nittany Dolomite is no longer limestone of the Stonehenge Formation, but has been dolomitized to produce the Larke Dolomite. At other localities, only parts of the Stone-

henge have been converted to Larke Dolomite, so most geologic maps merge the two units in a single map unit. The somewhat random distribution of late-stage dolomite also resulted in the base of the Stonehenge Formation being placed some distance below the base of the flooding surface created by the Stonehenge transgression. This resulted in the addition of a basal unit, below the Stoufferstown Member, that is charac- terized by meter-scale cycles (the Spring Creek member of Donaldson, 1959). Similarly, in the part of the Great Valley in eastern Pennsylvania known as the Lebanon Valley (figure 18-1 of Gray and Root, 1999) and in adjacent New Jersey (Markewicz and Dalton, 1977), a dolomitization front descended through the Stonehenge and replaced much, or all, of the three members with a mappable succession of dolomite designated as the Rickenbach Dolomite (Hobson, 1963). Significant var- iation in the depth in the stratigraphic column to which this dolomitization front descended has caused con- siderable confusion in the mapping and correlation of formations in the Beekmantown Group between eastern Pennsylvania and New Jersey. The key to correlation across the entire central Appalachian transect appears to be the lateral persistence of the Stoufferstown Member and a major turnover in cono- dont faunas at the base of the R. manitouensis Zone a short distance above the base of that unit.

Medium- to thick-bedded intraclastic rudstones (flat- pebble conglomerates), most likely deposited as tem- pestites on the subtidal shelf after the Stonehenge transgression, are very common within the Stouffers- town Member in Pennsylvania. They occur as interbeds within the characteristically wavy thin- to irregularly bedded wackestone to very fine grainstone of the Stoufferstown Member throughout its area of expo- sure. The noncyclic subtidal succession manifested by these lithologies has been identified by different names in different areas. It is the informal Graysville mem- ber (Donaldson, 1959) in the Nittany arch, the middle member of the Stonehenge Formation in the Lebanon Valley, and the Big Springs Member of the New Jersey Epler Formation (Hobson,1963). Seeing no need for multiple designations, we refer to this package through- out the central Appalachians as the Stoufferstown Member.

The addition of the thin package of cyclic carbo- nates termed the Spring Creek member at the base of the Stonehenge Formation in the Nittany arch (Donald- son, 1959) is a reasonable, pragmatic, cartographic adap- tation that exploits the traceability of the contact be- tween chert-rich dolomite of the underlying Mines Dolomite Member of the Gatesburg Formation and the limestone-based cycles of the basal Spring Creek. But assignment of the package mapped by Hobson (1963) in

404 Brezinski et al.

the Lebanon Valley as the Stonehenge Formation is less easily justified because his lower member, which accounts for approximately half the thickness of the formation, is composed of meter-scale cycles litholog- ically indistinguishable from the underlying Conoco- cheague Formation. Only a very thin (10–15 m [33–49 ft]) Stoufferstown (i.e., middle member of Hobson, 1963) and the lowest 18 m (59 ft) of the thrombolitic reefs of the middle member (our use is equivalent to upper member of Hobson, 1963) survived below the dolomi- tization front that created the overlying Rickenbach. On the margin of the Pennsylvania depocenter in eastern- most Pennsylvania and adjacent New Jersey, the entire Stonehenge was lost to dolomitization. Nonetheless, the shaly character, flat-pebble conglomerates, and conodont fauna of the Stoufferstown Member again facilitate recognition of that key interval in New Jersey, where it was mapped as the Big Springs Member in the middle of what has incorrectly been identified (Mar- kewicz and Dalton, 1977) as the Epler Formation (see discussion below).

In the Nittany arch, the boundary between the Stoufferstown Member and middle member is a well- preserved, gradational, midcycle boundary that records shallowing, with the appearance of isolated thrombo- litic bioherms a few meters or decameters in width in the upper half to one third of the Stoufferstown Mem- ber. Because of the shallower inner-shelf environment, the middle member contains numerous prism-cracked finely laminated dolomites that alternate with blocky, thickly bedded microbial biostromes. One-meter (3.3-ft)- thick dololaminite, with abundant quartz sand near the top of the middle member, records what apparently were the shallowest conditions attained during the re- gression. This likely correlates in time with the highest few meters of the middle member in the Great Valley where the only two dolomitic laminites found in the member in that area occur (Nguyen, 1986). The sharp contact with the overlying upper member is a flood- ing surface marked by the replacement of interbedded thrombolitic reefs and dololaminites with subtidal, thinly bedded, fine grainstone and medium-bedded flat-pebble conglomerates (i.e., lithologies that strongly resemble those at the base of the Stoufferstown Mem- ber). As in the Great Valley, the upper member con- tains intervals of oolitic grainstone with a very well- preserved macrofauna dominated by trilobites of the Bellefontia Zone. Interbeds of tan finely crystalline do- lomite occur near the top of the member in some sec- tions, and meter-scale cycles are abundant low in the Forge Union Member of the overlying Nittany Dolo- mite. However, exposures of this contact are few, typi- cally poor, and in all but a few sections, the underlying unit is the Larke Dolomite, instead of the limestone of

the Stonehenge Formation (Donaldson, 1959; Spelman, 1966). Consequently, precise placement of this forma- tion contact, a gradational midcycle boundary like the base of the middle member of the Stonehenge but strongly overprinted by later dolomitization, is prob- lematic. The timing and nature of the transition into peritidal cyclic facies within this second third-order sequence in the Lower Ordovician can be more easily established in the Great Valley of Maryland where the Stonehenge grades upward into the Rockdale Run Formation (Sando, 1958; Brezinski et al., 1999).

The Rockdale Run, which is more than 850 m (>2789 ft) thick, was subdivided by Sando (1957) into three lith- ologically distinct intervals (Figure 15). The basal 60 m (197 ft) consist of limestone-dolomite cycles in which the thin thrombolitic reefs, which comprise the lime- stone bases of the cycles, have been partially replaced by chert. This lower cyclic package is the regressive phase of the upper Stonehenge third-order cycle. This interval is overlain by about 70 m (�230 ft) of medium-bedded oolitic packstone to grainstone that Sando (1957) informally named the oolitic member. This interval represents a third-order deepening event that resulted in deposition of a deeper subtidal facies.

Although the name has been used extensively in the limestone-floored valleys in northern New Jersey (Markewicz and Dalton, 1977), the interval identified as Epler in that area is now known to comprise a sig- nificantly older interval than the true Epler in Penn- sylvania. As previously mentioned, the lithologies and conodont faunas recovered from the middle Big Springs Member of the Epler in New Jersey prove beyond a doubt that it is actually the dolomitized Stoufferstown Member of the Stonehenge Formation. Consequently, the use of the term Epler to refer to these rocks should be discontinued, as should the term Rickenbach for dolomites even lower in the section that clearly are not equivalent to the type Rickenbach in its type area in the Lebanon Valley.

A fourth Early Ordovician deepening episode that moderated the restrictive conditions during deposi- tion of the upper half of the Rockdale Run Formation produced a thin interval of limestone approximately 250 m (�820 ft) above the top of the oolitic member. The same transgression is recorded in the Epler Formation of the Lebanon Valley (Hobson, 1963), which includes two third-order transgressive peaks. This fourth event produced the subtidal package at the top of that forma- tion. The subtidal facies at the base of the Epler formed as a result of the previous transgression, which created the oolitic member of the Rockdale Run Formation in Maryland. Peritidal cyclic deposits that make up the middle of the Epler formed during the regressive phase

Sequential Development of Platform to Off-platform Facies 405

Figure 15. The stratigraphic column of the Lower Ordovician units in Sauk III, show- ing the positions, relative to conodont and trilobite zones, of subtidal packages that represent third-order deepening episodes in the Stonehenge and Rockdale Run Formations of the Great Valley. The relative sea level curve (modified from Brezinski, et al., 1999). The amplitude and frequency of individual fourth- and fifth-order sea level units are approximately depicted. Ph. = Phragmodus; H. = Histiodella; N. = Neomultioistodus; T. = Tricladiodus; 100 m (328 ft).

406 Brezinski et al.

of the lower Epler transgressive-regressive sequence. Age-equivalent cyclic sandy carbonates produced during this mid-Epler regression are preserved in the Nittany arch as the Axemann Formation, which lies between the Nittany Dolomite and the Bellefonte Dolomite.

Roughly, the upper half of the Rockdale Run For- mation consists of stacked meter-scale cycles and is 450 m (1476 ft) thick. Individual cycles consist of a lower subcycle composed of thin-bedded limestone to thrombolitic boundstone and an upper subcycle con- sisting of ribbony dolomitic limestone and tan lami- nated dolomite (Brezinski et al., 1999; Cecil et al., 2004) (Figure 14G). The relative thickness of the limestone and dolomite parts of the cycles varies with respect to their position within the formation. Within the lower 150 m (492 ft), typically, the limestone part of the cycle is well developed and thicker than the dolomitic caps, which are commonly no more than 0.5 m (V1.6 ft) thick. Higher in the formation, dolomite dominates each cycle and the limestone parts become increasingly thin (0.5 m [1.6 ft]) to nonexistent. This vertical change in the lithology of the cycles led Sando (1957) to desig- nate an informal dolomite member for the uppermost part of the formation. This upsection increase in dolo- mite records increasingly restricted conditions toward the top of the formation, resulting from exposure of the platform during the regressive phase of the Sauk III supersequence (Brezinski et al., 1999) (Figure 13).

The restriction that created the dolomite member at the top of the Rockdale Run Formation became even more severe during deposition of the overlying Pinesburg Station Dolomite (Figure 15), a sequence of medium-gray, tan-weathering, medium- to thick-bedded cherty dolomite that is 130 to 170 m (427–558 ft) thick (Sando, 1957). Alternations of thick-bedded dolomite with dololaminites attest to the cyclic nature of the original sediment (Figure 14H). Commonly, this mas- sive to thickly bedded dolomite is highly fractured to brecciated. Although much of the pervasive fracturing that characterizes this unit can be attributed to tectonic deformation, at least some of the brecciated dolomite layers formed through subaerial exposure and disso- lution by fresh waters (Mussman and Read, 1986).

Shoaling of the Sauk III brought to a close the Sauk megasequence in the central Appalachians and led to exposure of much of the platform. This exposure resulted in the creation of the Knox-Beekmantown unconformity (Mussman and Read, 1986; Ryder et al., 1992) in most regions of the Appalachians. In the central part of the Great Valley, deposition of shallow to peritidal carbonate sediment continued and the magnitude of Knox-Beekmantown hiatus for this area is in question; physical evidence of subaerial expo-

sure has not been documented, and the record of conodont biozones is unbroken through this interval (Harris and Repetski, 1982a). The Sauk-Tippecanoe megasequence boundary in western Maryland thus coincides precisely with the formational boundary of the [limestone] base of the St. Paul Group on the [dolostone] top of the Pinesburg Station Dolomite (Figure 16).

The base of the dolomite member of the upper Rockdale Run Formation would mark the level at which Finney et al. (2007) chose to place the Sauk-Tippecanoe boundary. Finney et al. (2007) pointed out that the be- ginning of the significant sea level drawdown over Laurentia is marked by the initiation of floods of craton- derived sands onto thick successions of continental slope black shales in contintental margin settings, which they dated by both graptolites and conodonts in the western United States Cordillera and the Ouachita Mountains of Arkansas. The appearance of these sands marks a type I sequence boundary (Van Wagoner et al., 1988), which thus provides a logical and correlatable boundary between the two supersequences. This bound- ary falls at a level high in the Reutterodus andinus cono- dont Zone, a position very near the top of the Ibexian. Although the conodont faunas from the upper Rockdale Run are quite sparse and of low diversity (and this part of the section is notoriously devoid of other identi- fiable fossils), the first appearance of Middle Ordovi- cian conodonts is very close to the base of the dolomite member of the Rockdale Run (Harris and Repetski, 1982b). Although the precision of correlation is less than ideal, this level is consistent both depositionally and biostratigraphically with the sequence boundary level described by Finney et al. (2007). Conceding the merits of placing the Sauk-Tippecanoe boundary at the ear- liest stage of the Blackhillsian-Rangerian interregional regression, herein however, we choose to place that supersequence boundary at the subsequent (relatively rapid) flooding of the craton that occurred during deposition of the upper part of the Histiodella holo- dentata or basal part of the Phragmodus polonicus cono- dont zone (Figures 2, 16). This higher level is more recognizable across a larger area, and it is even more correlatable because, across most exposed areas of Laurentia, the time between the broad initiation of sea level drop and this subsequent rapid flooding is represented as nonrecord, an unconformity, or as a succession of thin shallow-water lithofacies units sep- arated by numerous unconformities.

Sauk III Platform-margin and Periplatform Facies

Although the shallow-water carbonates of Sauk III were being deposited to the west, shelf-break and periplatform

Sequential Development of Platform to Off-platform Facies 407

sediments accumulated within the Frederick Valley, as the Grove and Frederick Formations, in what is now Maryland’s western Piedmont (Reinhardt, 1974, 1977). The Frederick Formation is interpreted as a relatively complete succession of deep-water carbonates that formed through continuous deposition during the late Steptoean and entire Sunwaptan. As previously noted, the basal unit, the Araby Formation, formed as a very thin distal equivalent of the Chilhowee clastics to the west, and the basal Monocacy Member of the Frederick accumulated slowly through Sauk II as starved basin deposits (Brezinski, 2004) (Figure 17A). However, evidence of an oversteepened platform mar- gin to the west during deposition of Sauk III is pro- vided by the overlying Rocky Springs Station Member of the Frederick Formation, a thick wedge of carbon- ates characterized by intervals of polymictic breccia and arenaceous grainstone (Reinhardt, 1974, 1977; Demicco,

1985; Brezinski, 2004) (Figure 17B). The Rocky Springs Station breccias range in thickness from 1 to 10 m (3.3– 33 ft) on the western side of the valley, but rarely exceed more than 1 m (>3.3 ft) in thickness on the eastern side. Individual breccia beds commonly exhibit nor- mal or reverse-grading, fining upsection into calcar- eous sandstone, then flaggy bedded peloidal limestone, and finally into thinly bedded hemipelagic limestones (Figure 17C). The breccia facies associations commonly exhibit a cyclic stacking of one unit on top of another (Brezinski, 2004) (Figure 18). The vertically repeated breccia facies associations are consistent with subaque- ous debris flow and turbidity currents (Reinhardt, 1974, 1977; Brezinski, 2004). The lack of rounding of carbonate clasts and local inverse grading are consistent with subaqueous cohesionless sediment flows (Nemec and Steel, 1984). Also present within the flaggy and thin- bedded hemipelagic limestones are imbricated clasts and buckled and contorted strata that are the result of translational sliding. The presence of these transported layers within the Rocky Springs Station Member indi- cates deposition on a relatively steep slope (Cook, 1979) (Figure 17D). In its entirety, the member is an amal- gamated composite of several LSTs that formed during several major drawdowns in the Late Cambrian. Recov- ery of a trilobite fauna assignable to the lower Steptoean Dunderbergia Zone from low in the member (Rasetti, 1961) identifies those strata as part of the LST that ac- cumulated in a proximal slope setting during the re- gression that created the Big Spring Station Member and/or the unconformity at the base of the Conoco- cheague to the west.

The wedge-shaped apron of slide masses that charac- terize the Rocky Springs Station Member on the west- ern side of the Frederick Valley interfinger distally with thin-bedded to laminated shaly limestone and hemi- pelagic basinal shales and shaly limestone to the east (Brezinski, 2004). The wedgelike geometry of the Rocky Springs Station Member is produced by eastward thin- ning of the member from more than 2300 m (>7546 ft) to approximately 300 m (�984 ft). The accumulation of this wedge of sediment apparently reduced the steep- ness of the platform edge so that progressively thinner slide masses were deposited upsection (through time). Continued reduction in the steepness of the slope by aggradation and progradation might have played some part in changing the style of deposition from high- energy debris-flow deposition in the Rocky Springs Station Member to lower energy accumulation of very thinly bedded shaly lime mudstone during deposition of the overlying Adamstown Member. However, the age of this important transition, provided by trilobites recovered from the top of the Rocky Springs Station Member (Taylor et al., 2009), suggests a different extrinsic

Figure 16. The stratigraphic section of the Tippecanoe megasequence of the Great Valley of Maryland (modified from Brezinski et al., 1999). The amplitude and frequency of individual small-scale sea level units have not been accurately determined or portrayed. A. = Amorphognathus; C. = Cahabagnathus; Ph. = Phragmodus; H. = Histiodella; N. = Neomultioistodus; T. = Tricladiodus; 100 m (328 ft).

408 Brezinski et al.

cause. The sharp contact between these members marks a termination of the transport of carbonate sands and gravels from the shelf margin to lower slope environ- ments. A trilobite fauna collected from the highest few meters of the Rocky Springs Station Member contains Keithiella depressa and species of Onchonotus and Stenopi- lus. Correlations between the Frederick Formation and the Conococheague Formation are facilitated by comparison with the Shallow Bay Formation in Newfoundland and Hoyt Limestone in New York. Similarity of trilobite faunas suggests a widespread sea level rise that produced thrombolite III and ter- minated downslope transport of carbonate debris. At

Hickey Cove, Newfoundland, K. depressa and one of the stratigraphically restricted species of Plethopeltis from the Conococheague were recovered just above the highest occurrence of P. saratogensis, above the highest conglomerate and arenaceous grainstone of the Downes Point Member, and just below the contact with shaly, low-energy, distal-slope deposits of the Tuckers Cove Member. James et al. (1989) attribute this onlap of more fine-grained distal-slope facies onto the prox- imal debris apron as the result of a eustatic rise in sea level in the middle of the Sunwaptan Stage. This is precisely the age indicated for the sea level rise that producedthromboliteIIIintheConococheagueFormation

Figure 17. The Sauk III periplat- form lithologies. (A–E) Frederick Formation. (A) The black lami- nated shale of condensed facies of the upper Monocacy Member. (B) The coarsening-upward (white triangle) sequence of poly- mictic debris-flow breccia of the Rocky Springs Station Member. (C) Imbricate slabs of limestone indicating penecontemporaneous slide deposits, Rocky Springs Sta- tion Member. (D) Thinly bedded limestone or dolomite rhythmites of the Adamstown Member. (E) The burrow-mottled limestone of the upper Lime Kiln Member ramp facies. (F–H) The Grove Forma- tion. (F) The cross-bedded arena- ceous dolomitic grainstone of the Ceresville Member. (G) The dolo- mitized thrombolite texture char- acteristic of the Fountain Rock Member. (H) The thinly bedded, shaly, upper slope limestone of the Woodsboro Member.

Sequential Development of Platform to Off-platform Facies 409

at the top of the range of P. saratogensis. Consequently, we interpret the comparable onlap of distal-slope facies of the Adamstown Member over the proximal-slope facies of the Rocky Springs Station Member as the product of the thrombolite III transgression, the medial Sunwaptan eustatic event identified by James et al. (1989) as their event 5.

As periplatform deposition continued through ac- cumulation of the Adamstown Member, aggradation brought the sediment surface closer to sea level. Even- tually, the shallowing was manifested in the onset of bioturbation at the base of the overlying Lime Kiln Member (Reinhardt, 1974). The prevalence of biotur- bation increases upward through the Lime Kiln, and the appearance of some microbial boundstone near the top of the member (Figure 17E) suggests emergence into the euphotic zone. The highest strata of the Lime Kiln Member constitute a relatively shallow-water ramp facies consisting of interbedded medium-gray, biotur- bated, thin-bedded, lime mudstone and associated thrombolitic and stromatolitic microbial boundstone (Figure 17E).

By the Early Ordovician, deep-water deposits that had accumulated in the periplatform setting had ag- graded, or very nearly, to sea level (Figure 19). The tran- sition to shallow-shelf deposition is recorded in lithol- ogies at the base of the Grove Formation (Brezinski, 2004). The basal Ceresville Member of the Grove For- mation consists of 50 to 65 m (164–213 ft) of dolomi-

tized, cross-bedded, arenaceous grainstone and throm- bolitic boundstone (Figure 17F). The Ceresville Member is overlain by the much thicker Fountain Rock Mem- ber, which comprises interbedded, massive, dolomitic, thrombolitic boundstone and tan fractured and lami- nated dolomite (Figure 17G). Some thrombolitic layers are more than 5 m (>16 ft) thick, whereas the inter- bedded dolomite layers are rarely more than 1 m thick (V3.3 ft). Also present are localized herringbone cross- bedded quartzose sandstones. Thickness estimates for the Fountain Rock Member, based on incomplete mea- sured sections and outcrop band width, range from 750 to 900 m (2461–2953 ft) (Taylor et al., 1996; Brezinski, 2004). This great thickness of thrombolitic boundstone is interpreted as the product of repeated episodes of sea level rise that provided accommodation space for the accumulation of a thick succession of algal bioherms at the edge of the ramp (Figure 17G).

The highest Paleozoic strata preserved within the Frederick Valley are dark, shaly, thin-bedded, and only slightly bioturbated lime mudstone very similar in to the Lime Kiln Member of the Frederick Formation (Taylor et al., 1996). Named the Woodsboro Member of the Grove Formation by Brezinski (2004), this unit is be- tween 100 and 150 m (328 and 492 ft) thick and rep- resents a return to the upper-slope conditions like those that produced the upper Frederick Formation (Figure 17H). Conodonts recovered from the Woodsboro Member indicate that it formed during Macerodus dianae

Figure 18. The section of stacked poly- mictic breccias within the Rocky Springs Station Member on the eastern side of the Frederick Formation outcrop belt. The sequence illustrates the cyclic nature of the fining-upward debris-flow deposits (modified from Brezinski, 2004). 5 m (16.4 ft).

410 Brezinski et al.

Figure 19. The stratigraphic cross section from the Nittany arch of Pennsylvania to the Frederick Valley of Maryland, illustrating lithofacies changes across the platform during Sauk III regression and subtidal packages marking four third-order transgressions in the earliest Ordovician. 100 m (328 ft); 1 km (0.6 mi).

Seq u en tial

D evelo

p m en t o f P latfo

rm to

O ff-p

latfo rm

Facies 4 1 1

Zone time. It was during deposition of this zone that the dark subtidal facies bearing an asaphid trilobite fauna in the lower part of the Epler Formation and the oolitic member of the Rockdale Run Formation formed; both were products of the same third-order deepen- ing episode (Hardie, 1989; Brezinski et al., 1999). Thus, the Woodsboro Member appears to represent a medial Early Ordovician submergence event that not only drowned the platform edge thrombolitic boundstone facies in the Frederick Valley, but also interrupted the deposition of meter-scale cycles in the Rockdale Run Formation of the Great Valley to the west. This medial Ibexian submergence event is also recorded in the southwestern United States by the José Oolite, a pack- age of dark burrow-mottled oolitic packstone within the M. dianae Zone in the El Paso Group of southern New Mexico and west Texas (Taylor et al., 2004). Like the oolitic member of the Rockdale Run, the José Oolite (treated as a member of the Hitt Canyon Formation in recent articles) reflects deepening that inhibited micro- bial reef growth and interrupted the deposition of the peritidal cyclic facies across vast areas of the Lauren- tian platform. This submergence event also correlates well with the O-5 third-order deepening episode of Read (1989a) in the southern Appalachians. In north- eastern Tennessee, this interval comprises the upper- most part of the Chepultepec Dolomite (Upper Lime- stone of Read, 1989a) through much of the overlying Kingsport Formation, an interval that contains M. dianae Zone conodonts (Repetski, 1985).

THE TIPPECANOE MEGASEQUENCE

The peritidal to supratidal facies that characterize the upper dolomite member of the Rockdale Run Forma- tion and the overlying Pinesburg Station Dolomite are overlain sharply by a succession of limestone assigned to the Row Park and New Market Limestones and col- lectively known as the St. Paul Group (Neuman, 1951). This sharp contact is unconformable and separates the Beekmantown Group from the overlying Middle Ordo- vician limestones (Figure 20A). Farther to the south, this contact has been interpreted as the Knox-Beekmantown unconformity, whose lacuna expands to as much as 10 m.y. (Read, 1980). However, the region of our study was at or near the depocenter of the central Appalachian basin, and deposition appears to have been essentially continuous through the uppermost Beekmantown and into the St. Paul Group. No evidence of significant un- conformity is present in this region, nor is there bio- stratigraphic evidence of any significant hiatus (Harris and Repetski, 1982a). In the Frederick Valley, the Sauk-

Tippecanoe boundary interval has been removed by subsequent erosion.

Shelf Facies

The St. Paul Group is 70 to 100 m (230–328 ft) thick and consists of four lithofacies. These facies consist of a massive fenestral lime mudstone (Figure 20C); a skel- etal packstone-grainstone; a laminated dolomitic lime mudstone; and a stromatolitic boundstone (Demicco and Mitchell, 1982; Hardie, 1989) (Figure 20C, D). The facies succession, both vertically and laterally, is inter- preted as being produced by parts of two third-order deepening episodes (Hardie, 1989) (Figure 16). The first deepening is represented by the stacking of subtidal skeletal packstone-grainstone facies on top of the shal- low subtidal pond deposits of the fenestral limestone facies. Shoaling from this event produced the upsec- tion stacking of the tidal deposits of laminated dolo- mitic lime mudstone and then the stromatolitic bound- stone facies. Deepening of the second third-order event produced a reversed facies succession from below, the laminated lime mudstone being replaced upsection by nonfenestral massive lime mudstone. Hardie (1989, their figure 14) interpreted this sequence as being produced by conversion of a restricted-circulation lagoon to an open-shelf lagoon followed by regression-induced transition to tidal-flat and coastal-lake environments. The second deepening event continued through the deposition of the top of the New Market Limestone and into the overlying Chambersburg Formation (Figure 16).

The Row Park and New Market Limestones are cor- relative with the Loysburg Formation of the Nittany arch region. Much of the Loysburg Formation is char- acterized by cyclic limestone consisting of intraclastic and bioclastic grainstone interstratified with ribbony and laminated dolomitic limestone. The banding of the ribbony facies gives this unit its informal name, the Tiger Stripe (Wagner, 1966). The cyclic ribbony and laminated facies of the Loysburg is similar to that pre- sent in the New Market Limestone of the Great Val- ley and is interpreted as a tidal-flat deposit (Laughrey et al., 2004).

Ramp Facies

To the south, in northern Virginia, the light-colored up- per limestone strata of the New Market Limestone inter- finger with, and are replaced by, a dark-gray, thin- to medium-bedded, siliceous, argillaceous lime wacke- stone termed the Lincolnshire Formation (Read, 1980; Rader and Read, 1989). The Lincolnshire Formation

412 Brezinski et al.

represents subtidal shelf to ramp facies. This deeper water facies bordered the New Market tidal flats on the south (Read, 1980).

Overlying the Row Park and New Market Lime- stones of the Great Valley is a sequence of dark-gray, thin- to medium-bedded, locally nodular limestone known as the Chambersburg Formation in Pennsyl- vania and Maryland and the Edinburg Formation in Virginia. The Chambersburg Formation is 100 to 150 m (328–492 ft) thick, and its basal 15 m (49 ft) are grada- tional with the underlying New Market Limestone

and Lincolnshire Formation (Figure 20E). These thin- bedded, dark-gray, argillaceous basal strata grade up- section into 15 to 20 m (49–66 ft) of dark-gray, shaly, nodular limestone that are informally termed the Echinosphaerites beds (Neuman, 1951; Brezinski, 1996b) (Figure 20F). These nodular beds grade upsection into thin-bedded argillaceous limestone and then into a thickly bedded, bioturbated lime wackestone that is as much as 10 m (33 ft) thick in the middle of the forma- tion (Figure 20G). The bioturbated strata, in turn, are re- placed upsection by the thin-bedded facies and then

Figure 20. The lithologies of the Tippecanoe megasequence. (A) The unconformable contact be- tween the brecciated dolomite of the Pinesburg Station (Ops) and overlying fenestral lime mudstone of the Row Park Limestone of the St. Paul Group (Osp). (B) The massive fenestral lime mudstone of the basal Row Park Limestone of the St. Paul Group. (C, D) The thick-bedded high-frequency cy- clic facies of the New Market Limestone. (E) The medium- bedded, cherty, subtidal- to inner- ramp facies of the Lincolnshire Formation. (F) The nodular-bedded outer-ramp facies of the Echino- sphaerites beds of the lower Chambersburg Formation. (G) The ramp facies of the Chambersburg Formation. (H) The Stickley Run Member of the Martinsburg For- mation illustrating foundering of the carbonate ramp and bank by shales.

Sequential Development of Platform to Off-platform Facies 413

by the nodular-bedded facies at the top of the formation.

The sequence of facies in the Chambersburg For- mation records two separate third-order deepening episodes (Figure 16). The earlier deepening episode began in the upper New Market Limestone, with the upsection transition from cyclic tidal-flat facies to sub- tidal, medium-bedded, bioturbated limestone. This deepening continued into shallow-ramp facies of the basal Chambersburg Formation, reaching maximum deepening in the deep-ramp facies of the Echinosphae- rites interval. Shoaling back to shallow subtidal envi- ronments from this first deepening is recorded upsection by the overlying thin-bedded facies followed by the thickly bedded middle part of the Chambers- burg Formation (Figure 20G). This shallowing was fol- lowed by another deepening into deeper ramp facies at the top of the formation (Figure 20H).

In the Nittany arch region, the Chambersburg Formation is replaced by a succession of carbonate for- mations that illustrate the foundering of the Late Or- dovician shelf, through the deposition of the Black River and Trenton Groups. The Black River Group re- cords the transition from tidal-flat facies within the Hatter Limestone at its base, to agitated shallow- to deeper subtidal shelf facies in the Snyder, and then to deep-subtidal upper-ramp facies of the Linden Hall Formation (Laughrey et al., 2004) (Figure 21). Overall deepening of the succession continued into the over- lying Trenton Group, with the deposition of the wavy- to nodular-bedded ramp wackestones of the Nealmont Formation, and the dark-gray medium-bedded rhyth- mites of the Salona and Coburn Formations (Figure 21).

DEATH OF THE GREAT AMERICAN CARBONATE BANK IN THE CENTRAL APPALACHIANS

As many as ten third-order deepening and shallowing episodes have been recognized in the Black River and Trenton Groups of the Nittany arch. This compares with only three currently identified in the Great Valley of Maryland. The reasons for this difference lie both in the more nearshore position of the Nittany section, and the equivalence of the upper Trenton Group with the Martinsburg to the east. Conodont biostratigraphy illustrates that the Edinburg-Chambersburg lithologies were buried beneath the Martinsburg, whereas the car- bonates of the Salona and Coburn Formations were being deposited farther to the west. This diachronous overstepping of the carbonate shelf by the Taconic flysch sediments is illustrated in Figure 22. Thus, whereas deep-ramp facies were being buried be-

neath clastics of the Martinsburg Formation in northern Virginia, shallow-shelf deposition and up- per-ramp facies continued to be deposited in the Nittany arch region (Swartz, 1948).

ACKNOWLEDGMENTS

We thank Bill Spinrad, Marie Sauder, and Scott Bell of the U.S. National Park Service for expediting the issu- ance of the scientific collecting permit to sample ex- emplary stratigraphic sections along the C&O Canal National Historical Park. This chapter was greatly im- proved with suggestions provided by J. F. Read, J. T. Dutro, and C. A. Kertis. J. F. Taylor acknowledges the

Figure 21. The stratigraphy and relative sea level curve for formations of the Tippecanoe megasequence in the Nittany arch (modified from Laughrey et al., 2004). 50 m (164 ft).

414 Brezinski et al.

donors of the Petroleum Research Fund administered by the American Chemical Society for partial support of this research and numerous students at the Indiana University of Pennsylvania (IUP) who participated in the research, particularly, M. Arford, J. D. Bader, C. A. Burger, J. A. Camargo, M. L. Druso, N. J. Durika, J. W. Fulton, B. C. Klingensmith, T. D. Lavanga, J. D. Loch, M. P. Morgan, P. J. Perfetta, H. J. Renyck, C. A. Roebuck, C. A. Roth, B. K. Sell, N. M. Welsh, and M. L. Zelawski. Financial support was also received through grants from the IUP University Senate and the Faculty Pro- fessional Development Committee of the Pennsylvania State System of Higher Education. We also benefited from assistance in the field and shared insight on strat- igraphic relationships from G. R. Ganis, D. Hopkins, C. T. Nguyen, R. C. Orndorff, and the late J. Reinhardt and W. J. Sando.

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