permo-carboniferous carbonate platforms and reefs - Lock Haven ...

PERMO-CARBONIFEROUS CARBONATE 

PLATFORMS AND REEFS 

Volc 

W T-16 T-463 T-5246 T-5056 T-10 E 

500 m 

V.E. = 10 

5000 m 

Volc 

Unit 1 

Unit 2 

Unit 3 

Top Lvis 

Tul 

Bob 

Rad 

TSS 

HSS 

TSS HSS TSS HSS TSS 

Bash_SSB 

Serp_SSB 

Lvis13_csb 

Lvis2_MFS 

Lvis_SSB 

Lvis1_MFS 

Edited by: 

WAYNE M. AHR 

Department of Geology and Geophysics, Texas A&M University, College Station, Texas 77843-3115, U.S.A. 

PAUL M. (MITCH) HARRIS 

ChevronTexaco E&P Technology Company, 6001 Bollinger Canyon Road, 

San Ramon, California 94583-0746, U.S.A. 

WILLIAM A. MORGAN 

ConocoPhillips, Inc., P.O. Box 2197, Houston, Texas 77252-2197, U.S.A. 

AND 

IAN D. SOMERVILLE 

Department of Geology, University College - Dublin, Belfield, Dublin 4, Ireland 

Top Evis 

Evis_SSB 

Tour_MFS 

Fame_SSB 

Bash = Bashkirian 

Serp = Serpukhovian 

Tul = Tulsky 

Bob = Bobrikovsky 

Tour = Tournaisian 

Fame = Famennian (Devonian) Devonian 

Lvis = Late Visean Rad = Radaevsky SSB = Supersequence Boundary 

“Volc” = Thick Ash Interval Evis - Early Visean MFS = Maximum Flooding Surface 

TSS = Transgressive Sequence Set HSS = Highstand Sequence Set 

SEPM Special Publication 78 

AAPG Memoir 83 

Platform (Shallow & Deeper) 

Upper-Middle Slope & 

Deeper Platform 

Allochthonous Debris 

Lower Slope - 

Toe of Slope 

“Volc” Interval 

Basin 

Published jointly by 

SEPM (Society for Sedimentary Geology) and 

The American Association of Petroleum Geologists (AAPG) 

Tulsa, Oklahoma, U.S.A. 

Printed in the United States of America 

December, 2003 

Volcanic

SEPM, AAPG, and the authors are grateful to the following 

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Permo-Carboniferous Carbonate Platforms and Reefs 

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Published December, 2003

MISSISSIPPIAN RAMP, WEST VIRGINIA, U.S.A. 

SEQUENCE RESPONSE OF A DISTAL-TO-PROXIMAL FORELAND RAMP TO 

GLACIO-EUSTASY AND TECTONICS: MISSISSIPPIAN, APPALACHIAN BASIN, 

WEST VIRGINIA–VIRGINIA, U.S.A. 

AUS AL-TAWIL 

Saudi Aramco, Dhahran, Saudi Arabia 

AND 

THOMAS C. WYNN, AND J. FRED READ 

Department of Geological Sciences, Virginia Tech, Blacksburg Virginia 24061, U.S.A. 

ABSTRACT: This paper evaluates the limits of using well-cuttings data and wireline logs in conjunction with limited core and outcrop data 

to generate a regional, high-resolution sequence stratigraphy for upper Mississippian (Chesterian) Greenbrier carbonates, West Virginia, 

U.S.A. These data are then used to document the stratigraphic response of the distal to proximal foreland basin to tectonics and 

Carboniferous glacio-eustasy during the transition into ice-house times. The major mappable sequences are fourth-order sequences, a few 

meters to over a hundred meters thick. They consist of updip red beds and eolianites, lagoonal muddy carbonates, ooid grainstone and 

skeletal grainstone–packstone shoal complexes, open-ramp skeletal wackestone, and slope–basinal laminated argillaceous carbonates. The 

sequences are bounded downdip by lowstand sandstones and calcareous siltstones, and locally on the ramp by basal transgressive shales; 

only a few sequence boundaries are calichified, compared with updip sections in Kentucky, where caliche and breccias are common. 

Transgressive systems tracts range from thin units to others that constitute the lower half of the sequence. The highstand systems tracts 

contain significant grainstone units. Maximum flooding surfaces on the ramp slope occur at the base of slope or basinal facies that rest on 

lowstand to transgressive complexes, whereas on the ramp they occur beneath widespread grainstones that overlie nearshore shale or lime 

mudstone. In the Greenbrier succession, fourth-order sequences are arranged into weak third-order composite sequences bounded updip 

by red beds, and by lowstand sands and oolite along the ramp slope. The composite sequences contain three to four fourth-order sequences. 

Correlation of the foreland basin units with third-order global sea-level curves, and with high-frequency sequences within the 

intracratonic Illinois Basin, shows that in spite of differential subsidence rates ranging from 2 to 30 cm/ky across the foreland, global thirdand 

fourth-order sea-level changes whose amplitude increased with time were a strong influence on sequence development. Thrust-loadinduced 

differential subsidence of fault blocks of the foreland basement controlled the rapid basinward thickening of the depositional 

wedge, and modified the eustatic effects on the accumulating succession. Moderate-amplitude eustatic sea-level change and semiarid 

climate were the dominant causes of the widespread reservoirs of ooid grainstone and lowstand sands. The overall stratigraphy suggests 

upward increase in amplitudes of sea-level change, and cooling, which likely records the initiation of Gondwana ice-sheet growth. 

INTRODUCTION 

This paper mainly uses well data (especially well cuttings and 

wireline logs), along with limited outcrop data and core, to 

determine if they can be used to generate a high-resolution threedimensional, 

sequence stratigraphic framework for the Mississippian 

(Chesterian) carbonates, from 0 up to 200 m thick, in the 

eastern, Appalachian foreland basin (Colton, 1970; de Witt and 

McGrew, 1979) (Figs. 1–4). The study differs from Al-Tawil and 

Read (this volume), which was largely an outcrop-based study 

along two outcrop belts in the distal foreland, Kentucky. Its 

foreland-basin setting and presence of the ramp margin in the 

study area distinguish it from the studies of Smith and Read 

(1999, 2000, 2001) in the Illinois Basin, an intracratonic basin to the 

west in which the ramp margin is not present. The units in the 

study area show extremely complex facies heterogeneity, with 

rapid vertical and lateral facies changes that have hampered 

development of regional exploration models, and recognition of 

shoreline and ramp-margin trends. 

The paper provides the first detailed, high-resolution, regional 

sequence stratigraphic study of the mixed carbonate and 

siliciclastic facies extending from the slowly subsiding distal 

foreland to the more rapidly subsiding proximal foreland of the 

Appalachian Basin. Previously, the succession had been subdivided 

into several lithologically heterogeneous formations (Fig. 

4), with much of the correlation being based on a few semiregional 

siliciclastic marker beds (Reger, 1926; Wells, 1950; 

McFarlan and Walker, 1956; Rice et al., 1979). We subdivided the 

Chesterian succession into regionally traceable fourth-order high- 

Permo-Carboniferous Carbonate Platforms and Reefs 

SEPM Special Publication No. 78 and AAPG Memoir 83, Copyright © 2003 

SEPM (Society for Sedimentary Geology), American Association of Petroleum Geologists (AAPG), ISBN 1-56576-087-5, p. 11–34. 

11 

frequency depositional sequences, by tracing major landward 

and basinward shifts in facies, and mapping outcropping sequence-bounding 

disconformities, correlative conformities, and 

sequence-bounding siliciclastic units. This largely subsurface 

study illustrates how high-resolution sequence stratigraphic 

analysis can increase our understanding of facies stacking on the 

ramp (Ahr, 1973) and its response to foreland basin tectonics, 

glacio-eustatic sea-level changes related to onset of Gondwana 

glaciation (Crowell, 1978), and long-term climate change 

(Ettensohn et al., 1988). 

The study documents how, even in this tectonically active 

foreland-basin setting, the major role of tectonics has been one of 

generating differential accommodation across and along the 

foreland basin, and in localizing the position of structural highs, 

the ramp’s hinge line, and ramp-margin facies. In spite of the 

active tectonic regime, fourth-order and, to a lesser extent, thirdorder 

eustasy has been the overriding control on the generation 

of the regionally correlative sequences. 

GEOLOGIC SETTING 

Regional Structural Setting 

The Appalachian Basin is an elongate foreland basin bordered 

on the east by overthrusted Precambrian and early Paleozoic rocks 

and on the west by the Cincinnati Arch (Colton, 1970; de Witt and 

McGrew, 1979) (Figs. 1, 2). In the west, the Paleozoic rocks are 

undeformed to gently folded, whereas the eastern part of the 

basin is in the folded and thrusted Valley-and-Ridge Province.

12 

The Upper Mississippian carbonate strata in the Appalachian 

Basin (Arkle et al., 1979) are exposed in narrow, northeast– 

southwest trending outcrop belts, one along the distal foreland in 

Kentucky adjacent to the Cincinnati Arch (described by Al-Tawil 

and Read, this volume) and the others (described here) in the 

more proximal foreland, exposed along the eastern Appalachian 

Plateau and into the fold–thrust belt (Figs. 1, 2). Between these 

outcrop belts, the Mississippian rocks of the distal to proximal 

foreland basin lie in the subsurface beneath Pennsylvanian sediments 

and are penetrated by numerous wells, twenty of which 

were used along with a core and four outcrop sections for this 

study. 

The Appalachian foreland basin was tectonically active during 

the deposition of these Upper Mississippian carbonates, and 

the Mississippian rocks thicken into the foredeep in southwestern 

Virginia and Tennessee (Fig. 1). The depocenter of the Late 

Mississippian carbonate ramp is located in the overthrust belt in 

the Greendale Syncline, southwestern Virginia, where Mississippian 

units thicken to over 2000 meters of basinal, slope, and 

ramp-margin facies (Cooper, 1948; Bartlett and Webb, 1971). 

Several synsedimentary positive structures occur in West Virginia 

and eastern Kentucky (Yeilding and Dennison, 1986; Dever 

et al., 1990; Sable and Dever, 1990; Dever, 1995) (Fig. 1). The 

Cincinnati Arch, which borders the western edge of the basin, 

subsided very little during Chesterian time, while adjacent 

areas to the west (Illinois Basin) and the east (Appalachian 

AUS AL-TAWIL, THOMAS C. WYNN, AND J. FRED READ 

FIG. 1.—Geologic location map of study area, showing distribution of exposed Mississippian Greenbrier carbonate rocks (gray 

shading) in the Appalachian Basin, eastern U.S.A. Isopach contours (in meters) show thickness of Greenbrier Group, which 

thickens into the Appalachian foredeep to the southeast (modified from Flowers, 1956; Pryor and Sable, 1974; MacQuown and 

Pear, 1980; Yeilding and Dennison, 1986; Dever, 1986, 1995; Sable and Dever, 1990, Dever et al., 1990). Major structures within the 

study area are the ancestral Rome Trough border faults, and the Jessamine Dome and Cincinnati and Waverly arches, which 

bound the area to the northwest, and the Burning–Mann anticline and Virginia Promontory, which cut across the trend of the 

basin. Orogenic highlands bounded the basin to the southeast. 

Basin) subsided more rapidly (Pryor and Sable, 1974; deWitt 

and McGrew, 1979; Whitehead, 1984; Sable and Dever, 1990; 

Dever, 1995). The West Virginia Dome was a positive structure 

along the 38th parallel lineament in northeastern West Virginia, 

which was active during Early Chesterian time (Yeilding and 

Dennison, 1986). 

The basement in the study area is segmented by major faults 

into a series of blocks, some of which moved independently 

during Mississippian deposition (Shumaker and Wilson, 1996). 

These faults include northeast–southwest trending bounding 

faults of the Rome Trough, as well as several zones parallel to and 

to the east of these. The northeast–southwest trending basinal 

hinge occurs just to the west of the overthrust belt, and a smaller 

east–west trending hinge occurs in the northeast along the 400foot 

(120 m) Greenbrier isopach (Fig. 1). A northwest–southeast 

trending arch lies at right angles to the regional structure, in the 

vicinity of the “Virginia Promontory”, a long-lived Paleozoic 

high (Thomas, 1977). 

Stratigraphy and Biostratigraphy 

The Mississippian succession makes up a supersequence set 

(roughly 35 My duration) composed of at least three 

supersequences. The lower supersequence set boundary is the 

Devonian–Mississippian erosional unconformity (Bjerstadt and 

Kammer, 1988) (Figs. 3, 4). The lower supersequence consists of

WAVERLY 

ARCH 

ROME ROME TROUGH TROUGH 

APPALACHIAN 

BASIN 

OHIO 

the uppermost Devonian–lowermost Mississippian Cloyd Conglomerate–Berea 

Sandstone and the overlying Subury Shale– 

Price/Borden/Maccrady deltaic complex (Kinderhookian to 

mid-Osagean), which contains several clinoformed depositional 

sequences (Branson, 1912; Butts, 1940, 1941; Bjerstadt and 

Kammer, 1988; Bartlett, 1974). This is overlain southwest of the 

study area by the carbonate-prone Fort Payne–Salem 

supersequence (upper Osagean–lower Meramecian) (Fig. 4; 

Bjerstadt and Kammer, 1988; Sable and Dever, 1990; Khetani 

and Read, 2002). 

The upper Meramecian–Chesterian supersequence (Fig. 4) 

described in this paper overlies the lower two supersequences. 

The upper Meramecian Little Valley–Hillsdale sequence may 

have formed during lowered sea levels relative to the overlying 

units, given that it is the least extensive unit on the platform, and 

it contains relatively shallow oncolitic limestone facies even in the 

most basinward sections. The Chesterian sequences C1 to 11 

described in this paper make up the carbonate-prone, largely 

transgressive part of the supersequence. The shale-prone sequence 

C11 may be the initial sequence of the supersequence 

highstand, the bulk of the supersequence highstand being made 

up of the progradational Mauch Chunk red siliciclastics. The 

early Pennsylvanian unconformity marks the top boundary of 

the supersequence set. 

Regional Mississippian stratigraphy and biostratigraphy of 

the Mississippian in Virginia and West Virginia (Figs. 3, 4) are 

given in Butts (1922, 1940, 1941), Reger (1926), Wells (1950), 


40° N 83° W 40° N 80° W 

11 

BURNING-MANN 

ANTICLINE 

FOLD AND THRUST BELT 

WEST 

VIRGINIA 

12 

14 

13 

15 

17 

16 

18 20 

5 

19 

21 

22 

24 

9 

23 

25 26 

10 

VIRGINIA 

PROMONTORY 

WEST VIRGINIA 

DOME 

FIG. 2.—Location map showing outcrop sections and well sections used to construct the cross sections in the study. Location 

information on sections is in the Appendix. Outcrops of Mississippian Greenbrier Limestone is shaded. The northeast–southwest 

cross section, based on sections 1 to 11 (Fig. 6), includes data from the eastern outcrop belt, supplemented by nearby shallow wells 

and a core; the sections thicken southwestward down the basin axis. The northwest–southeast cross section extends across the 

foreland normal to the trend of the basin and is based on wells 12 to 26 (Fig. 7); it uses only subsurface well cuttings and wireline 

logs. The succession thickens markedly across the northeast–southwest trending hinge (just southeast of the 100 m isopach of 

Figure 1) that separates the distal foreland from the proximal foreland. Note that to the northeast this hinge curves to the east. 

Two minor east–west trending hinges occur to the southwest of the main hinge. Outcrops representative of the distal foreland 

of Kentucky (along strike to the southwest) are described in Al-Tawil and Read (this volume). 

4 

7 

8 

DISTAL FORELAND 

6 

2 

3 

1 

37° N 80° W 

VIRGINIA 

PROXIMAL FORELAND 

0 

100 km 

100 mi 

MISSISSIPPIAN 

OROGENIC 

HIGHLANDS 

13 

Flowers (1956), de Witt and McGrew (1979), Rice et al. (1979), and 

Maples and Waters (1987). The Greenbrier Group (Hillsdale to 

Alderson interval and the Lillydale Shale and Glenray Limestone 

of the Lower Bluefield Formation, Mauch Chunk Group) make 

up the study interval (0 to 1000 m thick). It unconformably 

overlies the Price–Maccrady deltaic–marine shelf deposits (Figs. 

3, 4). The basal unit studied is the upper Meramecian Little 

Valley–Hillsdale Limestone, which is a very cherty, pelletal– 

peloidal and skeletal packstone–wackestone with the distinctive 

corals Syringopora and Acrocyathus proliferus (Lithostrotionella) 

(Wells, 1950; DiRenzo, 1986). This is overlain by the Chesterian 

Denmar Formation, which in outcrop contains numerous exposure 

surfaces, consists of ooid grainstone, lime wackestone– 

mudstone, and minor quartz–peloidal grainstone, and has the 

crinoid Platycrinites penicillus and the earliest occurrence of the 

zonal conodonts G. bilineatus and C. charactus (Butts, 1922, 1940, 

1941; Wells, 1950; McFarlan and Walker, 1956; Dever et al., 1990). 

The Denmar Formation is overlain by the Taggard red beds, a 

regional marker consisting of red shale and siltstone and minor 

carbonates. 

The overlying ooid- and skeletal limestone–rich units and 

subordinate carbonate mudstone constitute the Gasper Formation, 

which contains the first appearance of the crinoid index 

fossil Talarocrinus (Butts, 1940, 1941; Wells, 1950). The interval is 

subdivided into the Pickaway, Union, Greenville Shale, and 

Alderson Formations (Reger, 1926). This is overlain by Lillydale 

Shale and the Glenray Limestone (Reger, 1926), consisting of

14 

PERIOD 

MISSISSIPPIAN 

DEV. 

EPOCH 

KIND. OSAGEAN MERAM. CHESTERIAN 

AGE 

327 

354 


CONODONT 

ZONES 

G. bilineatus - 

K. mehli 


C. altus 


C. charactus 

A. scalenus - 

Cavusgnathus 

Cavusgnathus - 

T. varians 

OTHER 

INDEX 

FOSSILS 

Talarocrinus 

spp. 

Platycrinites 

penicillus 

Acrocyathus spp. 

'Lithostrotion spp.' 

MAUCH CHUNK GP. 

G R E E N B R I A R G R O U P ("B I G L I M E") 

WEST VIRGINIA 

DROOP SS. 

REYNOLDS 

LIMESTONE 

GLENRAY 

LIMESTONE 

LILLYDALE 

SHALE 

ALDERSON 

LIMESTONE 

UNION 

LIMESTONE 

PICKAWAY 

LIMESTONE 

UPPER / LOWER 

TAGGARD 

DENMAR 

LIMESTONE 

VIRGINIA 

COVE 

CREEK 

LIMESTONE 

FIDO SS. 

HILLSDALE LST. 

LITTLE VALLEY LST. 

MACCRADY 

FORMATION 

PRICE 

FORMATION 

PENNINGTON 

FM. 

GASPER 

LIMESTONE 

STE. 

GENEVIEVE 

LIMESTONE 

LATE DEVONIAN AND 

EARLY MISSISSIPPIAN 

BLACK SHALE 

SEQ. 

THIS 

PAPER 

FIG. 3.—Chronostratigraphic chart, Mississippian interval, showing the major units in the study area of West Virginia and Virginia; 

studied interval is from the Little Valley–Hillsdale to the Glenray Formation, and includes the drillers’ “Big Lime” (Greenbrier 

Group) and the “Little Lime”. Chesterian sequences described in this paper are numbered C1 to C11; underlying Meramecian 

sequences are not numbered. The study interval is equivalent to the St. Louis to Glen Dean interval of Kentucky (Appalachian 

Basin) and the Illinois Basin. Modified from de Witt and McGrew (1979), Roberts et al (1995), Collinson et al (1971), Rexroad and 

Clarke (1960), Huggins (1983), Stamm (1997), Barlett and Webb (1971), Reger (1926), and Wells (1950). 

ooid and skeletal limestone and containing the zonal conodonts 

G. bilineatus and Kladognathus (Rexroad and Clarke, 1960). The 

succession is capped by Mauch Chunk Group red sandstone, 

siltstone, and mudrock, whose top is the Mississippian–Pennsylvanian 

unconformity (Brezinski, 1989). Regionally, the units 

thicken to over 600 m to the southeast into the foredeep, where 

they are dominated by thick units of dark skeletal wackestone– 

mudstone and basin–slope laminated argillaceous carbonate 

mudstones, and thin quartz sandstone and skeletal–ooid 

grainstone. In the subsurface, the Hillsdale to Alderson carbonates 

are the drillers’ “Big Lime” and the Glenray Limestone is 

the “Little Lime.” 

BLUEFIELD FM. 

C11 

C10 

C9 

C8 

C7 

C6 

C5 

C4 

C3 

C2 

C1 

FACIES AND REGIONAL DISTRIBUTION 

The facies developed in the eastern Appalachian Basin 

(Leonard, 1968; Carney and Smosna, 1989) are shown schematically 

on an idealized ramp model (Fig. 5) and resemble facies 

described elsewhere by Ettensohn et al. (1984) and Smith and 

Read (1999, 2001). The regional facies distributions are shown on 

the two regional cross sections (Figs. 6A, 7A). The first cross 

section extends southwestward down the axis of the foreland 

basin, and is based on five detailed outcrop sections, a core, and 

data from four wells. The second cross section is a dip section 

extending across the distal foreland onto the proximal foreland,

PERIOD 

MISSISSIPPIAN 

DEV. 

EPOCH 

KIND. OSAGEAN MERAM. CHESTERIAN 

AGE 

327 

354 

DISTAL PROXIMAL 

FORELAND 

MAUCH CHUNK 

SILICICLASTICS 

REYNOLDS LST. 

GLENRAY LST. 

LILLYDALE 

SHALE 

U. TAGGARD 

L. TAGGARD 

ALDERSON LST. 

UNION LST. 

PICKAWAY LST. 

R R R 

R R 

DENMAR LST. 

PRICE FM. 

SILICICLASTICS 


COVE CREEK LST. 

HILLSDALE LST. 

L. VALLEY LST. 

MACCRADY 

FM. 

LATE DEVONIAN AND EARLY 

MISSISSIPPIAN BLACK SHALE 

FIDO SS. 

FIG. 4.—Schematic lithostratigraphy showing the major Mississippian 

stratigraphic units in the area. The study interval 

(Little Valley–Hillsdale to the Glenray/Cove Creek Limestone), 

overlies Early Mississippian siliciclastics of the Price– 

Borden formations, and is overlain by and interfingers with 

the upper Mississippian Mauch Chunk siliciclastics. 

PALEOSOLS 

QUARTZ-PELOID 

GRAINSTONE / PACKSTONE 

RED BEDS 

SEQUENCE BOUNDARY 

PELOIDAL 


LIME WACKESTONE / MUDSTONE 

~ 200 km 

OOID 


GRAY SHALE/SILTSTONE 

15 

and consists solely of data from 15 wells (cuttings with 10 to 15 

foot sample intervals and wire-line logs). Figures 6B and 7B show 

the individual sequences from each cross section rehung using 

the top of each sequence as a datum, and using inferred water 

depths of the topmost facies to hang the ramp-margin sections. 

Facies are summarized in Tables 1 and 2. 

Paleosols and Erosion Surfaces 

Carbonate paleosols occur on disconformities on the distal 

foreland, southwest of the study area in Kentucky (Dever, 1973; 

Ettensohn, 1975; Harrison and Steinen, 1978; Al-Tawil, 1998) and 

West Virginia, but they are less common in the thicker West 

Virginia–Virginia units (Figs. 6, 7). The paleosols include laminar 

caliches and caliche breccia (Table 1). They formed under seasonally 

wet–dry semiarid conditions (Ettensohn et al., 1988). However, 

many contacts between sequences and parasequences in 

outcrop on the proximal foreland in West Virginia are sharp 

planar erosional to scalloped surfaces. They resemble microkarst 

or tidal solution basins. 

Red Beds (Marginal Marine) 

These are red, commonly structureless, unfossiliferous, fine 

siliciclastic mudrocks and laminated shale (Figs. 6, 7; Table 1). 

Updip in West Virginia, ripple lamination, wave lamination, and 

mud cracks are common, but burrows are rare. Some red beds 

have rooted horizons, and some layers are burrowed. The red 

beds are common in updip areas, where they are associated with 

eolianites, and restricted lagoonal and shoal-water grainstones. 

These updip facies are marginal marine to intertidal–supratidal 

red beds, which formed as highstand, basinward-shifting facies. 

Gray Shale, Quartz Sandstone, and Calcareous Siltstone 

(Tidally Influenced Open Siliciclastic Shelf) 

These facies include dark gray, poorly fissile, limy shales that 

are barren to variably fossiliferous, with small productid brachiopods, 

pelecypods, and bryozoa (Table 1). Some overlie pyritized 

bedding planes with abundant lacy bryozoans, especially at the 

tops of sequences C10 and C11 (Alderson and Glenray Formations). 

The shales form units a few centimeters thick that pinch 

out laterally, or are extensive sheets up to 30 m thick (e.g., in 

SKELETAL 


CALCAREOUS SILTSTONE 

SANDSTONE 

ARGILLACEOUS SKELETAL 

WACKESTONE / PACKSTONE 

LAMINATED SHALY 

LIME MUDSTONE /CALCAREOUS SILTSTONE 

FIG. 5.—Schematic facies profile for the Mississippian of the Appalachian Basin. Actual facies distributions are more complex, in that 

skeletal grainstone–packstone and ooid grainstone facies are developed not only on the ramp margin but also in local areas far 

back in the ramp interior. 

0 m 

20 m 

60 m

16 

150 MILES 

1 2 3 4 5 6 7 8 9 10 11 

CANAAN RANDOLPH 

GREENBRIER SUMMERS SUN OIL GREENBRIER 

MERCER 

VALLEY 

28 MINGO 

21 

7 

CORE 

15 UNION 

25 INGLESIDE HOLSTON 

FORMATION 

SEQ. 

GLENRAY LST. 

GREENVILLE 

SHALE 

BLUEFIELD 

FM. 

COVE CREEK 

LIMESTONE 

C11 

LILLYDALE SHALE 

DATUM 

FIDO SS. 

ALDERSON 

FM. 

C10 

C9 


UNION 

FM. 

C8 

RED-BEDS 

PICKAWAY 

FM. 

GRAY SHALE / SILTSTONE 

C7 


C6 

QUARTZ SANDSTONE 

TAGGARD 

FM. 

C5 

PRICE / MACCRADY FM. 

QUARTZ PELOID GRAINSTONE 

FINE-GRAINED PELLETAL PACKSTONE, 

LIME WACKESTONE / MUDSTONE 

OOID GRAINSTONE / PACKSTONE 

C4 

PELOID GRAINSTONE / PACKSTONE 

SKELETAL GRAINSTONE / PACKSTONE, MINOR WACKESTONE 

DENMAR 

FM. 

C3 

ARGILLACEOUS SKELETAL PACKSTONE / WACKESTONE 

200' 

LAMINATED SHALY LIME MUDSTONE/ 



C2 

PALEOSOL 

C1 

HILLSDALE 

FM. 

FIG. 6.—A) High-resolution sequence stratigraphic cross section along the eastern outcrop belt (parallel to the axis of the foreland basin), incorporating outcrop sections 

and nearby shallow wells, which were analyzed using cuttings and wireline logs. The cross section extends from West Virginia into Virginia, and was constructed 

by hanging sections from the regional Lillydale Shale marker; it shows the interpreted regional distribution of lithofacies, sequence boundaries, and their relation to 

the previous formational stratigraphy (modified from Al-Tawil, 1998). Legend is used for Figure 6B and 7B). Detailed information on well and outcrop locations is 

given in the Appendix.


1 2 3 4 5 6 7 8 9 11 

1 2 3 4 5 6 7 8 9 10 11 

100' 

FEET 

0 MILES 10 

100' 

0 MILES 10 

HILLSDALE 

FIG. 6.—B) Facies distributions within individual sequences in Figure 6A (axial cross section), constructed by rehanging each 

sequence from its top sequence boundary, which was assumed to be horizontal on the ramp, and allowing the ramp margin to 

slope into the basin so that water depths of facies are compatible with estimated water depths in the text. Dotted line shows 

estimated positions of maximum flooding surfaces. Legend is same as Figure 6A. 

FEET 

C5 

C4 

C3 

C2 

C1 

C11 

C10 

C9 

C8 

C7 

C6 

17

18 

23 24 25 26 

FORMATION 

SEQ. 

22 

21 

20 

? 

C12 

19 

18 

BLUEFIELD 

FM. 

17 

? 

GREENVILLE 

SHALE 

16 

15 

DROOP SS. 

13 14 

12 

C11 

? 

LILLYDALE SHALE 

DATUM 

ALDERSON 

FM. 

C10 


UNION 

FM. 

C9 

C8 

100 

PICKAWAY 

FM. 

FEET 

C7 

10 

20 = WYOMING 253 

21 = WYOMING 87 

22 = MCDOWELL 63 

23 = MCDOWELL 792 

24 = MERCER 16 

25 = MERCER 1 

26 = MERCER 27 

MILES 

0 

12 = WAYNE 277 

13 = WAYNE 1497 

14 = LINCOLN 155 

15 = LINCOLN 1375 

16 = MINGO 796 

17 = LOGAN 102 

18 = LOGAN 922 

19 = LOGAN 388 

12 13 14 15 16 17 18 19 20 21 22 23 24 25 

C6 

TAGGARD 

EQUIVALENT 

DATUM 

C5 

? 

? 

C4 

OOID GRAINSTONE / PACKSTONE 

RED BEDS 

DENMAR 

FM. 

C3 

PRICE / MACCRADY FM. 

PELOID GRAINSTONE / PACKSTONE 

GRAY SHALE / SILTSTONE 

SKELETAL GRAINSTONE / PACKSTONE, 

MINOR WACKESTONE 


QUARTZ SANDSTONE 

C2 

ARGILLACEOUS SKELETAL 

PACKSTONE / WACKESTONE 

LAMINATED SHALY LIME MUDSTONE/ 


QUARTZ-PELOID GRAINSTONE 

C1 

FINE-GRAINED PELLETAL 

PACKSTONE,LIME 

WACKESTONE / MUDSTONE 


DOLOMITE 

HILLSDALE 

FM. 

PALEOSOL 

FIG. 7.—A) High-resolution sequence stratigraphic cross section from updip to downdip in West Virginia based on subsurface wells analyzed using cuttings and wireline 

logs (based on Thomas Wynn’s Ph.D. study, in prep.). Upper units were hung from the Lillydale Shale marker, as in Figure 6, but the lower sequences were hung from 

the base of the upper Taggard equivalent (quartz–peloid eolianites–calcareous siltstones at base of sequence C6). Cross section shows the interpreted regional 

distribution of lithofacies, sequence boundaries, and their relation to the previous formational stratigraphy. Note the similarity of the sequence stratigraphy and facies 

distributions of this well-based cross section to the dominantly outcrop- and well-based cross section of Figure 6. Detailed well location information is in the Appendix.


12 13 14 15 16 17 18 19 20 21 22 23 24 25 26 

100 FT. 30 M 

16 KM. 

0 10 MILES 

FIG. 7.—B) Facies distributions within individual sequences in Figure 7A, constructed by rehanging each sequence from its top 

sequence boundary, which was assumed to be horizontal. Legend is same as Figure 7A. Dotted line shows estimated positions 

of maximum flooding surfaces. 

sequence C11) that grade up into argillaceous, skeletal limestones 

(Figs. 6, 7). Sandstone and/or calcareous siltstone forms units up 

to 10 meters thick along the ramp margin and slope. Calcareous 

siltstones and sandstones are festoon cross-bedded to waveripple 

cross laminated or massive in outcrop and generally are 

unfossiliferous (Table 1). 

The shales are quiet-water deposits. Poorly fossiliferous thin 

shales formed in lagoons shoreward of sandstone–siltstone units 

at bases of sequences (cf. Ettensohn et al., 1984). However, thick 

fossiliferous shales such as those in sequence C11 may be deeperramp, 

open marine units (cf. Ettensohn et al., 1984). High 

siliciclastic influx may have been due to more humid climate, 

19 

perhaps coupled with short-term sea-level lowering. This increased 

turbidity suppressed carbonate productivity on the ramp, 

generating widespread shale deposits. The calcareous siltstone 

contains sedimentary structures that suggest a high-energy waveand 

current-influenced coastal setting during shallowing phases 

on the outer ramp. 

Quartz–Peloid Grainstone 

(Coastal Eolianites, Minor Marine Sand Sheets) 

These facies generally are unfossiliferous, very fine to fine and 

less commonly medium-grained quartz–peloid grainstone com- 

? 

? 

? 

C11 

C10 

C9 

C8 

C7 

C6 

C5 

C4 

C3 

C2 

C1 

MERAMECIAN 

SEQUENCES

20 

Lithology 

& Depositional 

Environment 

Caliche, Breccia and 

Erosional Contacts 

(Subaerial) 

Color Paleosols yellow to 

brown 

Depositional 

Texture and 

Grain Types 

Sedimentary 

Structures 

Paleosols have 

cryptocrystalline and 

fibrous calcite crusts and 

fracture fills; rooted 

fabrics common, and 

patches of caliche-coated 

peloids and pisolites. 

Variably silicified. May 

be brecciated. 

Paleosols have 

undulating lamination, 

rootlet casts, and 

incipient brecciation. 

Erosional contacts are 

sharp, planar to 

scalloped surfaces on the 

underlying unit. 

Fossils Rootlet casts in 

paleosols. 

posed of well rounded peloids, rare ooids, skeletal fragments, 

and finer-grained subangular quartz concentrated at bases of 

laminae (Table 2; Figs. 6, 7). They are cross-stratified, with foresets 

of extremely planar to curved, paper-thin, millimeter-scale laminae, 

that commonly are inversely graded (from very fine to fine 

sand) and form a characteristic pinstripe pattern on the outcrop 

or in core. Where inverse grading is absent, laminae form couplets 

of alternating very fine and fine sand. These facies commonly 

are interbedded with red beds, lagoonal facies, and ooidshoal 

facies. 

These facies are interpreted as coastal eolianites on the basis 

of abundant quartz–peloid grainstone lithology, the well rounded 

carbonate grain shapes, planar to curved foresets with inverse 

grading, and very fine to fine sand couplets, pinstripe lamination, 

and lack of in situ fossils or burrows (Merkely, 1991; Hunter, 1993; 

Dodd et al., 1993; Smith, 1996). Some of the coarser-grained, lowangle 

cross-laminated units may be beach facies. 

Fine-Grained Lime Wackestone–Mudstone 

(Tidal Flat and Lagoon) 

This facies occurs in massive to laminated units 1 to 8 meters 

thick. They include thin to medium-bedded, light gray, finegrained 

lime wackestone–mudstone and pellet packstone, with a 

generally restricted assemblage of gastropods, ostracodes, and 

rare small fragments of echinoderms and brachiopods (Table 2; 

Figs. 6, 7). Bedding planes locally contain abundant small bur- 


TABLE 1.—Mississippian Paleosols and Siliciclastic Facies. 

Siliciclastic paleosols 

and red-beds 

(Subaerial to Marginal 

Marine) 

Red, maroon and green 

mottled 

Terrigenous mudrocks 

and siltstones. 

Redbeds massive, to 

laminated, rare currentand 

wave ripples on 

siltstone interbeds, rare 

mudcracks. Paleosols 

massive mudrocks with 

slickensides 

Root/ burrow traces in 

paleosols 

Gray shale 

(Low Energy Shallow 

Marine) 

Quartz sandstone/ 

calcareous siltstones 

(Coastal to Shallow 

Marine) 

Dark gray to olive green White to light gray 

Terrigenous clay and silt Well-sorted fine to 

medium -grained shaly 

sandstones and 

siltstones of quartz, 

lesser carbonate grains 

Poorly fissile to massive Cross-bedded to 

structureless; flaser, 

lenticular and wavy 

bedded locally 

Poorly fossiliferous to 

very fossiliferous. 

Mollusks, ostracods, 

some echinoderms 

brachiopods and 

bryozoa; biota sparse 

and restricted in updip 

shale 

Rare 

rows and trails. In the east, some of these facies are microbially 

laminated dolomitic mudstones with centimeter layering, rare 

mudcracks, and few fossils. These laminated units are common 

downdip on the ramp in the eastern outcrop belt, but more 

massive units are widespread on the ramp (Table 2; Figs. 6, 7). 

The laminated facies are tidal-flat deposits developed downdip 

on the ramp during periods of emergence. The massive, finegrained 

carbonates were formed in low-energy lagoonal settings 

on the ramp interior and between grainstone shoals on the ramp. 

Lack of open-marine fossils, and association with shallow-marine 

facies and exposure surfaces in updip areas, indicate a 

relatively restricted, possibly periodically hypersaline lagoonal 

setting. 

Peloid and Ooid Grainstone (Shoals and Tidal Bars) 

These are white, cross-bedded, well-sorted grainstones composed 

of medium-sand-size ooids, peloids, intraclasts, and 

abraded skeletal fragments of echinoderms, bryozoa, brachiopods, 

and mollusks. They form units from 1 to 12 meters thick 

(Table 2; Figs. 6, 7). In the subsurface, the oolites form diporiented 

northwest-to-southeast trending oolitic sand ridges over 

30 km long and 1.4 km wide along the ramp margin (Kelleher and 

Smosna, 1993; Smosna and Koehler, 1993). Both ooid and skeletal 

grainstones locally contain rare, small bryozoan mounds. 

These peloid–oolitic facies (Table 2) formed on high-energy 

tide-influenced and lesser wave-influenced ooid shoals in water

Facies 

& Depositional 

Environment. 

Quartz Peloid 

Grainstone 

(Coastal Eolianite, 

Minor Marine Sand 

Sheets) 

depths from the intertidal zone down to 10 meters (Kelleher and 

Smosna, 1993; Keith and Zuppman, 1990). 

Skeletal Grainstone–Packstone 

(Ramp Margin and Moderate-Energy Lagoon) 

Because they are difficult to differentiate in well cuttings, all 

the coarse-grained, skeletal grainstones, skeletal packstones, and 

grain-rich wackestones were lumped together. In outcrop, these 

are medium- to thick-bedded, and cross-bedded to structureless. 

They are medium- to coarse-grained skeletal units ranging from 

clean grainstones to muddy packstone and lesser wackestone. 

They consist of whole and fragmented crinoid, bryozoan, brachiopod, 

and lesser mollusk remains (Table 2). Some skeletal 

beds contain abundant whole columnals and calices. The skeletal 

grainstone units form extensive sheets seaward of oolitic 

grainstone shoals along the ramp margin in units up to 15 meters 

thick, which interfinger with wackestone and deep-water laminated 

marls (Figs. 6, 7). The more poorly washed facies form 

relatively extensive units on the ramp, interfingering laterally 

with and surrounded by poorly fossiliferous lime wackestone– 

mudstone (Figs. 6, 7). 

These facies include high-energy skeletal banks and sheets 

deposited in wave-agitated and tidal-current settings between 

and seaward of ooid shoals along the ramp margin, in water 

depths from a few meters to perhaps 20 meters. They also formed 

extensive units within the lagoon, in areas colonized by abundant 

crinoids and other skeletal organisms, and were subjected to low 

to moderate wave reworking. 

Argillaceous Skeletal Wackestone 

These are medium- to thick-bedded in outcrop and are dark 

gray shaly skeletal wackestones with some packstone interbeds. 

They are composed of medium- to very coarse-grained echinoderm, 

brachiopod, and bryozoan debris, with interstitial finely 

fragmented skeletal debris, lime mud, and argillaceous seams 

(Table 2). Some wackestone units have decimeter beds of skeletal 

grainstone–packstone. This facies locally contains very small 


TABLE 2. —Mississippian Carbonate Facies. 

Fine-Grained Lime 

Wackestone/ Mudstone, 

Minor Lime Packstone 

(Tidal Flat and Low- 

Energy Lagoon) 

Color Dark-gray weathering Light gray. Tan to yellow 

weathering 

Grain Types And 

Depositional 

Texture 

Rounded and abraded 

peloids and some ooids, 

abraded skeletal 

fragments, and 

subangular very fine to 

fine quartz up to 50%. 

Grainstone fabrics. 

Unfossiliferous to 

moderately fossiliferous 

fine wackestone and 

mudstone and pellet 

packstone. Skeletal debris 

fine grained; abundant lime 

mud, dolomite mud, and 

pellets; may contain quartz 

silt and clay. Locally 

cherty. 

Sedimentary 

Structures 

Wedge sets of planner 

and tangential crossbedding, 

with pinstripe 

lamination, and inverse 

grading, dips up to 20 

degrees 

Biota Abraded, rounded 

skeletal fragments. No 

in situ biota or marine 

trace fossils. 

Dolomitic mudstones 

massive to microbially 

laminated (centimeter-thick 

laminae, micrograded, 

rarely mudcracked). Finegrained, 

skeletal-rich 

wackestone-mudstone is 

typically massive. 

None to sparse; may have 

mollusks, small crinoid 

columnals, ostracods. May 

have trace-fossil horizons, 

small oncolites, rare corals, 

and brachiopods. 

Peloid and Ooid 

Grainstone 

(High-Energy 

Shoal) 

Skeletal Grainstone/ 

Packstone 

(Ramp Margin and 

Lagoonal Skeletal 

Sheets and Shoals) 

Argillaceous Skeletal 

Wackestone 

(Deep Subtidal) 

21 

bryozoan reefs, especially in sequences C10 and C11. Argillaceous 

packstone–wackestone is most abundant along the ramp 

margin (Figs. 6, 7). They formed below fair-weather wave base 

and above storm wave base on the open ramp in water depths of 

20 to 60 m by comparison with the Persian Gulf (Purser and 

Seibold, 1973). However, these depths to maximum storm wave 

base could have been shallower in the Appalachian Basin, given 

that dominant wind directions in the Persian Gulf are axial to the 

basin and thus have maximum fetch, whereas dominant winds in 

the Appalachian Basin blew across the basin (Golonka et al., 

1994). Siliciclastic muds may have washed out onto the ramp 

following floods or storms, or they may mark brief cessations in 

carbonate deposition during short periods of more humid climate 

and higher siliciclastic influx. 

Laminated Shaly Lime Mudstone (Ramp Slop to Basin) 

These are unfossiliferous, dark gray, evenly thin-laminated to 

massive, commonly silty or argillaceous lime-mudstone units up 

to 100 meters thick along the ramp slope (Fig. 6, section 11; Table 2). 

They consist of fine, silt-size carbonate and quartz, finely comminuted 

skeletal debris, and a matrix of argillaceous lime mudstone. 

The shaly lime mudstones are deep-water ramp-slope to 

basin-floor facies, which lack shallow-water sedimentary structures 

and biota. Fine carbonate sediment was brought onto the 

slope perhaps by storms and tidal currents and mixed with landderived 

terrigenous silts and clays. Lack of burrowing suggests 

deposition in anoxic waters below the photic zone in depths of 60 

to over 100 m; the photic zone in the basin perhaps extended 

down to only to 30 m, given the turbid basin waters, whereas 

wave base perhaps reached down to more than 30 m but less than 

60 m, by comparison with the present Persian Gulf, a similarsized 

and similarly situated foreland basin (Purser and Seibold, 

1973). 

SEQUENCE STRATIGRAPHY 

Laminated Shaly 

Lime Mudstone 

(Ramp slope- 

Basin) 

Light gray to white Light to medium gray Medium gray to dark gray Typically dark gray 

Well sorted, 

rounded, medium to 

coarse grainstone.of 

sand-size ooids, 

peloids, lesser 

skeletal fragments, 

and minor 

intraclasts. 

Cross-bedded to 

massive 

Crinoid , brachiopod, 

bryozoan, and 

mollusk fragments, 

forams 

Variably fragmented 

echinoderms, 

brachiopod, bryozoa, 

mollusks, and rare ooids. 

Mud-free grainstones to 

grain-rich packstones 

and minor wackestones, 

some with argillaceous 

seams. 

May be cross-bedded, 

thin- to medium-bedded 

or structureless 

Abundant echinoderms, 

brachiopods, and 

bryozoas; and lesser 

mollusks. 

Echinoderms, 

brachiopods, bryozoa, and 

lesser mollusks. Include 

wackestone and lesser 

packstone; abundant lime 

mud; terrigenous clay 

disseminated or in seams 

and stringers 

Massive, thick to thin 

bedded, structureless; 

may have graded storm 

beds 

Abundant echinoderm, 

common brachiopods and 

bryozoas, rare mollusks 

Carbonate and 

silisiclastic clay and 

silt 

Millimeter to 

centimeter even 

lamination to 

structureless 

None 

The sequence stratigraphic terms used are after Sarg (1988), 

Mitchum and Van Wagoner (1991), and Weber et al. (1995). Facies

22 

and depositional sequences are shown in Figure 6A (extending 

northeast to southwest down the axis of the foredeep and into the 

proximal foreland basin, with sections arranged from thinnest to 

thickest) and in Figure 7A (oriented east–west at right angles to 

the foreland basin, and extending across the distal foreland, and 

just beyond the hinge into the margin of the proximal foreland). 

Depositional sequences and their component facies were traceable 

from the outcrop sections into the subsurface using well 

cuttings and wireline logs, especially gamma-ray logs. 

Sequences were recognized on the basis of major landward 

and basinward shifts in diagnostic facies belts such as eolianites, 

red beds, siliciclastic sands and/or calcareous siltstones, and lime 

grainstones. Sequence boundaries were picked at erosional disconformities 

(commonly evident in outcrop sections) on top of 

shallowing-upward carbonate units, and below lowstand or 

transgressive siliciclastic red beds, sandstones, and shale (evident 

in both subsurface and outcrop). Along the ramp slope, 

sequence boundaries were placed beneath the shallowest-water 

facies within the dominantly deep-water successions. In wireline 

logs, many sequence boundaries (and some parasequence boundaries) 

are associated with significant gamma-ray kicks due to 

presence of terrigenous mudrocks, shale, or argillaceous dolomites. 

This was important in sections containing parasequences, 

because it allowed parasequences to be grouped into sequences 

on the basis of the characteristic upward-decreasing (as the 

limestones clean upward) to upward-increasing gamma-ray signature 

(as the carbonates become more siliciclastic-rich). Because 

it was not clear using the subsurface data whether siliciclastic 

units were late highstand, lowstand, or early transgressive, they 

have been arbitrarily referred to as lowstand if on the ramp slope 

and margin, and lowstand to early transgressive if updip. Where 

possible, maximum flooding surfaces were picked below the 

deepest marine facies within the sequence, but the subsurface 

data commonly prevented them being traced regionally; thus it 

was possible to separate the transgressive and highstand systems 

tracts only locally on the ramp. Some of the maximum flooding 

surfaces are associated with high gamma-ray values, especially 

where the carbonates are shaly, but many in pure carbonate 

sections are not. 

Fourth-order sequences (roughly between 0.1 and 0.5 My 

duration) are the dominant units mapped in the study interval in 

the Appalachian Basin and the Illinois Basin (Smith and Read, 

1999, 2001; Al-Tawil and Read, this volume). Third-order or 

composite sequences (Weber et al., 1995) are far less well developed, 

being masked by the profound facies shifts within the 

dominant fourth-order sequences. Individual fourth-order sequences 

and component facies are shown in Figures 6B and 7B, 

constructed by rehanging the sections from the top of each 

sequence. Major biostratigraphic markers include Acrocyathus in 

the Hillsdale Formation, the last appearance of Platycrinites penicillus 

at the top of the Denmar Formation, the first appearance of 

Talarocrinus in the Pickaway Formation, as well as the conodont 

zones (Fig. 3). Individual fourth-order sequences were below the 

resolution of the biostratigraphic data, which typically span 

several sequences, but they do form valuable biostratigraphic 

markers within the succession. 

The sparse radiometric data and their error bars, and the 

difficulties in intercontinental correlation between Mississippian 

brachiopod, conodont, and foraminiferal zones, make absolute 

age determination of the study interval difficult. The St. Louis– 

Hillsdale interval is probably about 5 My in duration (Ross and 

Ross, 1987); sequences C1 to C11 (the Denmar to Glenray interval) 

are about 3 to 8 My (Sando, 1984; Baxter and Brenckler, 1982; 

Roberts et al., 1995; Smith, 1996). Assigning duration of 3 to 8 My 

for the three third-order sequences in the Denmar to Glenray 


interval gives a duration of 1 to 2.5 My for each of the composite 

sequences and 280 ky to an extreme upper limit of 720 ky for each 

of the 11 fourth-order sequences, with the most likely durations 

being 300 to 400 ky. (Smith, 1996). This encompasses the estimated 

average duration of 375 ky for the time-equivalent fourthorder 

sequences in the Illinois Basin (Smith, 1996) and the 400 ky 

durations estimated for the overlying Mauch Chunk siliciclastic 

fourth-order sequences (Miller and Eriksson, 1999). 

Fourth-Order Depositional Sequences 

Meramecian Sequences.— 

These Meramecian sequences (Figs. 6, 7) form the basal part of 

the study interval and include the Little Valley and Hillsdale 

formations. They unconformably (?) overlie the Price–Maccrady 

red beds, resting on a widespread transgressive sandstone in 

many places. The Meramecian sequences thicken from 0 to over 

270 m across the hinges into the basin (Fig. 6, section 11; Bartlett 

and Webb, 1971). Two to four Meramecian sequences may be 

present on the ramp. The sequence boundaries are placed beneath 

widespread lowstand to early transgressive sandstone– 

calcareous siltstone on the ramp to the west (Fig. 7), which appear 

to pinch out to the north, where two sequences may make up the 

interval (sequences not plotted individually in Fig. 6B). The 

transgressive to highstand parts of the sequences consist mainly 

of commonly cherty, fine skeletal wackestone–mudstone, grading 

basinward into shaly lime wackestone or lime mudstone. In 

the thick, proximal foreland basin sections, several sequences 

may be present (not shown in Fig. 6). Lower sequences here 

(mapped as Little Valley Formation) consist of quartz sandstone 

overlain by dark gray, nearshore argillaceous skeletal wackestone– 

mudstone, whereas the overlying Hillsdale sequences consist of 

oncolitic skeletal wackestone–mudstone capped by skeletal 

grainstone (Huggins, 1983). 

Chesterian Sequences.— 

Sequence C1.—Sequence C1 is the basal sequence of the Denmar 

Limestone (Figs. 6, 7). It generally overlies the Upper Meramecian 

Hillsdale Limestone and is 0 to 55 m thick (Fig. 6). The basal 

sequence boundary is placed beneath widespread basal, thin 

siliciclastic-prone units. Sequence C1 has a basal lowstand unit 

of local red beds that grades seaward into quartz sandstone– 

siltstone a few meters thick along the ramp margin and ramp 

slope (Fig. 6, sections 7 to 9 and 11; Fig. 7, sections 25, 26). These 

grade updip onto the ramp interior into lowstand–transgressive 

sandstone–calcareous siltstone, shale, and quartz–peloid 

grainstones (Fig. 6; sections 3 to 6; Fig. 7, sections 19 to 26) that 

may be overlain by possible transgressive carbonates in the 

middle of the sequence in the west (Fig. 7). The transgressive 

systems tract along the ramp slope consists of up to 20 meters of 

skeletal wackestone–mudstone grading up into skeletal grainstone 

(Fig. 6, section 11). The highstand systems tract downdip consists 

of deeper-water, laminated mudstones and skeletal wackestone 

that passes upsection into skeletal grainstone. Updip the highstand 

systems tract consists of thin, ramp-margin skeletal–ooid 

grainstones that grade updip into lime wackestone–mudstone 

and discrete ooid–peloid grainstone shoal deposits (Fig. 6, sections 

1 to 11; Fig. 7, sections 19 to 26). 

Sequence C2.—Sequence C2, the next Denmar sequence, is 0 to 

75 m thick (Figs. 6, 7); it thickens markedly between sections 6 and 

7 and sections 10 and 11 (Fig. 6). The sequence boundary is placed 

beneath a thin quartz sandstone on the ramp margin and slope,

and updip beneath widespread siltstone to the west (Fig. 7) and 

beneath eolianites updip to the north (Fig. 6). It has a lowstand 

quartz sandstone a few meters thick along the ramp margin and 

ramp slope (Fig. 6, section 11; Fig. 7, sections 24 to 26). The 

transgressive systems tract along the ramp slope consists of 

skeletal wackestone (Fig. 6, section 11; Fig. 7, section 26), and 

within a broad sag on the ramp margin (Fig. 6, sections 6 to 10), 

it probably includes the lower third of the sequence, consisting of 

quartz peloid and ooid grainstone and lime wackestone–mudstone 

beneath the lowest skeletal grainstones. The lowstand to 

transgressive systems tract updip includes thin quartz–peloid 

eolianites, and the lower portion of ooid grainstones in the 

northeast (Fig. 6, sections 1 to 5); to the west, the lowstand to 

transgressive tract consists of the regional calcareous siltstones 

(Fig. 7). The highstand systems tract along the ramp slope consists 

of deep-water shaly lime mudstone overlain by skeletal 

wackestone (Fig. 6, section 11; Fig. 7, section 26). Farther updip it 

consists of ramp-margin skeletal and minor ooid grainstones 

(Fig. 6, sections 9, 10; Fig. 7, section 25) that interfinger with rampinterior 

lime wackestone–mudstone containing multilateral ooid 

grainstone units (Figs. 6, 7). The vertical repetition of facies 

suggests that two to three parasequences may make up the 

sequence between sections 6 to 10 (Fig. 6). 

Sequence C3.—This next Denmar sequence is 0 to 40 m thick 

(Figs. 6, 7). Its basal sequence boundary is placed beneath thin 

siliciclastic units or eolianite (Figs. 6, 7). The lowstand systems 

tract is a thin quartz sandstone on the ramp slope (Fig. 6, section 

11). The transgressive systems tract is a wedge of skeletal wackestone 

overlain by skeletal grainstone along the ramp slope (Fig. 6, 

section 11). Updip the transgressive systems tract may include 

the lower half of the sequence, and is composed of thin, basal 

siliciclastic and eolianite units and the overlying lower parts of 

the discontinuous ooid grainstone and lime wackestone–mudstone 

units (Figs. 6, 7). The highstand systems tract constitutes the 

upper part of the sequence and consists of a patchwork of skeletal, 

oolitic, and lime wackestone–mudstone facies (Figs. 6, 7). 

Sequence C4.—Sequence C4, the top Denmar sequence, ranges 

from 5 to 80 m thick with marked thickening between sections 10 

and 11 (Fig. 6) and sections 24 to 26 (Fig. 7). On the ramp margin, 

the sequence boundary is a double disconformity in outcrop (Fig. 

6, section 10), and on the ramp slope the sequence boundary is 

beneath lowstand quartz sandstone several meters thick (Fig. 6, 

section 11) or calcareous siltstone (Fig. 7, sections 24 to 26). The 

sequence boundary on the ramp interior is placed beneath discontinuous 

lowstand eolianites (Fig. 6, sections 3, 7) or thin basal 

siliciclastic or quartz–peloid grainstone units (Fig. 7). The transgressive 

systems tract is absent on the slope, but on the ramp it 

may include thin local units of ooid grainstone and quartz–peloid 

grainstone in the north and east (Fig. 6, sections 3, 7, 10); in the 

west, the lowstand to transgressive systems tract consists of an 

extensive basal siliciclastic sheet (Fig. 7, sections 19 to 23). The 

highstand systems tract makes up the bulk of the sequence, and 

consists of ramp-margin skeletal grainstone–packstone and ooid– 

peloid grainstone, and ramp-interior, discontinuous units of lime 

wackestone–mudstone that separate laterally, local ooid–peloid 

shoal deposits (Figs. 6, 7). On the ramp slope, the highstand tract 

consists of deeper-water shaly mudstones (Fig. 6, section 11). 

Sequence C5.—Sequence C5, 0 to 25 m thick, includes the 

lower Taggard red beds (Figs. 6, 7). It shows a considerable 

basinward shift in red siliciclastic facies compared with the 

underlying sequences. The sequence boundary is beneath widespread 

siliciclastics and local quartz–peloid eolianites. The 


23 

lowstand to early transgressive systems tract consists of red 

beds updip in the northeast (Fig. 6) and on the ramp margin in 

the west (Fig. 7, section 26); thin, local eolianites and sandstone– 

siltstone on ramp in the west (Fig. 7, sections 19 to 22); and up 

to 15 m of sandstone–calcareous siltstone and marine shale on 

the ramp margin and slope (Fig. 6, sections 8, 9; Fig. 7, sections 

20 to 25). No transgressive systems tract is recognized in the 

north and east (Fig. 6). In the west, however, the transgressive 

tract may include the lower half of the carbonate interval, 

composed of basal units of ooid and peloid grainstone and the 

overlying lime wackestone–mudstones, if these are shallow 

lagoonal facies (Fig. 7, sections 17 to 22). The highstand tract on 

the shallow ramp is dominated by lime wackestone–mudstone, 

with ramp-margin grainstone units, and localized updip ooid 

and skeletal grainstone bodies (Figs. 6, 7); the highstand systems 

tract on the slope consists of basinal shaly lime mudstone 

grading up into skeletal wackestone (Fig. 6, section 11). 

Sequence C6.—Sequence C6 is 0 to 40 m thick (Figs. 6, 7) and 

thickens rapidly between sections 5 and 6 (Fig. 6). It includes the 

upper Taggard red beds and the lower Pickaway Limestone. The 

sequence boundary is placed below extensive siliciclastics (Fig. 6) 

or beneath widespread quartz–peloid eolianites (Fig. 7). The 

lowstand to early transgressive deposits consist of regional red 

beds up to 10 m thick in the north and east (Fig. 6) or regional 

eolianites to the west (Fig. 7). Downdip these grade into up to 30 

m of lowstand marine sandstone locally split by a thin shale– 

calcareous siltstone (Fig. 6, sections 9 to 11; Fig. 7, sections 20 to 

26). The transgressive systems tract on the ramp slope consists of 

a 15-m-thick succession of skeletal grainstone capped by a thin 

ooid grainstone (Fig. 6, section 11), which appears to grade updip 

into peloid grainstone (Fig. 7, section 26). On the ramp, it is 

difficult to separate the transgressive systems tract from the 

highstand systems tract. These combined systems tracts are 0 to 

35 m thick. Along the ramp margin in the west, probable highstand 

deposits include a downdip ooid grainstone unit grading updip 

into skeletal wackestone and then into ramp-margin skeletal 

grainstone–packstone (Fig. 7, sections 23 to 26); these likely grade 

downdip into skeletal grainstone, skeletal wackestones, and then 

basinal lime mudstone (Fig. 6, section 11). On the ramp interior in 

the northeast, the transgressive to highstand systems tract is 

dominated by widespread units of lime wackestone–mudstone 

(Figs. 6, 7) with eolianites in the far west. Skeletal grainstone– 

packstone and minor ooid grainstones also are localized within 

the upper third of the sequence in a broad sag on the middle ramp 

in the northeast (Fig. 6, sections 6, 7). 

Sequence C7.—Sequence C7 includes the upper Pickaway 

Limestone. It is 0 to 50 m thick, thickening between sections 3 and 

4 (Fig. 6) and between sections 22 to 26 (Fig. 7). The C7 sequence 

boundary on the ramp and ramp margin is placed below lowstand 

to early transgressive sandstone, shale, calcareous siltstone, minor 

red beds, and quartz–peloid eolianites in the east and northeast 

(Fig. 6, sections 5 to 9; Fig. 7, sections 18 to 26), and beneath 

skeletal wackestone on the ramp slope (Fig. 6, section 11). The 

transgressive systems tract over much of the ramp in the northeast 

contains skeletal grainstone interfingering laterally with 

lime wackestone–mudstone, and is capped downdip by a thin 

shale and sandstone and updip by sandstone and ooid grainstone 

(Fig. 6, sections 3 to 9); the transgressive succession is not clearly 

defined in the west but may include a wedge of skeletal wackestone 

overlain by calcareous siltstone–shale along the ramp margin 

(Fig. 7, sections 24 to 26). On the ramp slope, the transgressive 

systems tract is a skeletal wackestone (Fig. 6, section 11). The 

highstand tract along the ramp slope is a deep-water shaly lime

24 

mudstone (Fig. 6, section 11). On the ramp, the highstand systems 

tract in the east contains two belts of skeletal–ooid grainstone, one 

along the ramp margin (Fig. 6, section 9) and the other updip to 

the north flanking a significant downwarp (Fig. 6, sections 4 to 6); 

these updip grainstones are bordered both updip and downdip 

by lime wackestone–mudstone. In the west, the highstand is a 

simple facies succession of skeletal grainstone and skeletal 

wackestone that passes updip into ooid limestone and then into 

lime wackestone–mudstone (Fig. 7). Three or more parasequences 

may be present in the sequence. 

Sequence C8.—Sequence C8 consists of the lower Union Limestone 

and is 0 to 55 m thick (Figs. 6, 7). It thickens between sections 

3 and 7 (Fig. 6). The C8 sequence boundary is placed beneath 

lowstand to early transgressive deeper-water skeletal wackestones 

along the lower ramp slope (Fig. 6, section 11), and beneath 

lowstand to transgressive sandstone, calcareous siltstone, and 

rare red beds on the upper ramp slope and ramp margin (Fig. 6, 

section 9; Fig. 7, sections 22 to 26). The lowstand to early transgressive 

systems tract on the ramp interior may include basal, 

thin, patchily developed shales, peloid grainstone, and quartz– 

peloid eolianites (Fig. 6, 7). The transgressive systems tract on the 

ramp margin in the east makes up the lower one third of the 

sequence, and includes skeletal wackestone (Fig. 6, section 11) 

grading upsection and updip into skeletal grainstone and lesser 

ooid grainstone (Fig. 6, sections 7 to 9) and thin siliciclastic units 

and quartz–peloid eolianite (Fig. 6, sections 3 to 5; Fig. 7, sections 

18 to 25). The highstand systems tract is thick and dominated by 

an upward-shallowing succession of deeper-water mudstone to 

skeletal wackestone along the ramp slope (Fig. 6, section 11). On 

the ramp margin and ramp interior, the highstand systems tract 

includes at least two shallowing-upwards successions of skeletal 

wackestone up into skeletal grainstone that is bordered updip by 

locally thick ooid grainstone (Fig. 6, sections 8, 9; Fig. 7, sections 

23 to 26); these grade updip into lime wackestone–mudstone with 

local units of ooid grainstone (Fig. 6), and in the west, some thick 

eolianites (Fig. 7, sections 12, 19) separated laterally by lime 

wackestone–mudstone and containing a local shale in the middle 

of the succession (Fig. 6; Fig. 7, sections 18, 19). 

Sequence C9.—Sequence C9 is 5 to 25 m thick, and includes the 

upper Union Limestone (Figs. 6, 7). Updip in outcrop, the basal 

sequence boundary is locally at the base of thin channeled sandstones 

and red beds (Fig. 6, sections 1, 2). Farther downdip, the 

sequence boundary is placed beneath a thin local shale and a 

widespread unit of calcareous siltstone and minor sandstone 

(Figs. 6, 7) and beneath a thin oolite on the ramp slope (Fig. 6, 

section 11). This thin ooid grainstone is the lowstand systems 

tract along the ramp slope. Updip, the lowstand to transgressive 

systems tract consists of calcareous siltstone and shale (Fig. 6, 

sections 5 to 9; Fig. 7, sections 15 to 26) and in the far north, the thin 

channeled sandstones capped by red beds (mid-Union red beds) 

(Fig. 6, sections 1, 2). The transgressive systems tract on the ramp 

slope is a thin skeletal wackestone above the lowstand oolite (Fig. 

6, section 11), but a distinct transgressive systems tract is not 

easily recognized updip. The highstand systems tract is an upward-shallowing 

succession of deeper-water shaly lime mudstone 

to skeletal wackestone on the lower ramp slope (Fig. 6, 

section 11; Fig. 7, section 26). These pass upslope to the north into 

extensive ramp-margin skeletal grainstone (Fig. 6, sections 7 to 9; 

Fig. 7, sections 24, 25). These ramp-margin grainstones pass 

landward into skeletal wackestone or lime wackestone–mudstone, 

and then into updip skeletal grainstones and ooid grainstone 

(Fig. 6, sections 2 to 5; Fig. 7, sections 20 to 22) and on the 

innermost ramp, lime wackestone–mudstone. 


Sequence C10.—Sequence C10, 10 to 60 m thick, consists of the 

Greenville Shale and Alderson Limestone (Figs. 6, 7). The C10 

basal sequence boundary is placed beneath a widespread thin 

shale or beneath calcareous siltstone and sandstone (Figs. 6, 7). 

The lowstand systems tract along the ramp slope is a thin sandstone–siltstone 

(Fig. 6, section 11; Fig. 7, sections 24 to 26). The 

lowstand to transgressive systems tract farther updip probably 

includes the basal shale and the calcareous siltstone unit (Figs. 6, 

7). The highstand systems tract consists of skeletal wackestone 

downdip, passing updip into ramp-margin skeletal grainstone– 

packstone (Fig. 6, sections 7 to 9; Fig. 7, sections 25, 26). These 

grade updip to the northeast into thin ooid grainstone units and 

widespread lime wackestone–mudstone (Fig. 6, sections 2 to 6) or 

in the west, into thick ooid–peloid grainstone and then into lime 

wackestone–mudstone and local units of skeletal grainstone– 

packstone (Fig. 7, sections 12 to 22). 

Sequence C11.—Sequence C11 includes the Lillydale Shale and 

the Glenray Limestone and is 0 to over 100 m thick thickening 

rapidly across the hinge (Fig. 6, sections 9 to 11; Fig. 7, sections 21 

to 22). The C11 sequence boundary is placed beneath thin, locally 

developed quartz sandstone–siltstones (Fig. 6, sections 1, 6, 8; Fig. 

7, sections 17 to 20), or beneath thin shales, red beds, or eolianites 

(Fig. 7, sections 12 to 16), or at the erosional contact on C10 

carbonates at section 2 (Fig. 6). In the west, the sequence boundary 

on the ramp may be locally incised, indicated by rapid 

thickening of the siltstone at the expense of underlying units (Fig. 

6, section 11; Fig. 7, section 18). The lowstand systems tract on the 

ramp slope is a cross-bedded lithoclastic sandstone (Fido Sandstone), 

which appears to be incised into the underlying unit (Fig. 

6, section 11). The transgressive systems tract is not defined 

downdip, but updip the lowstand to transgressive systems tract 

consists of thin, local sandstone–siltstones (which thicken locally 

into the possible incised valley at section 18, Fig. 7), and an 

overlying thin regional oolitic and skeletal limestone (Figs. 6, 7). 

The highstand systems tract includes 10 to 40 m of fossiliferous 

marine shale (Lillydale Shale) that grades basinward into a rampmargin 

complex of skeletal wackestone, skeletal grainstone, and 

interbedded sandstone–siltstone (Fig. 7, section 26) and then into 

basinal lime mudstone up to 80 m thick (Fig. 6, section 11). These 

are overlain by late highstand, thin, discontinuous skeletal limestone 

capped locally by ooid grainstone (Fig. 6, sections 1, 3, 6, 8, 

9, 11; Fig. 7, sections 22 to 23). The top sequence boundary of 

sequence C11 in outcrop is a paleosol updip and/or erosional 

surface updip (Fig. 6, sections 1, 3), but on the ramp slope (Fig. 6, 

section 11) the sequence boundary underlies a thin black shale 

capped by ooid grainstone that is conformable on C11 skeletal 

wackestone. Sequence C11 is overlain by highly progradational 

red and gray siliciclastics (Mauch Chunk Formation). 

Parasequences 

Fourth-order sequences typically have only one or two 

parasequences updip, but downdip, sequences commonly are 

composed of several parasequences. Detailed facies successions 

of parasequences and presence or absence of erosional disconformities 

capping units are best observed in the outcrop sections or 

core. No attempt has been made to show parasequences in 

Figures 6 and 7. The parasequences are difficult to recognize in 

the cuttings data in sections in which the parasequences are thin 

and of the same magnitude as the sample spacing of the cuttings. 

Siliciclastic-prone parasequences in updip areas commonly 

consist of marine sandstone, locally with a slightly channeled 

base, fining upward into siltstone and then into red mudrock. 

Others include thin ooid or peloid grainstone overlain by red or

gray siltstone and then red mudrock. Siliciclastic units commonly 

make up the lower two-thirds of parasequences at the bases of 

sequences. Disconformity-bounded parasequences (2 to 6 m thick) 

are recognizable even in thick successions of restricted lime 

wackestone–mudstone (as in C7) and are composed of very thin 

(or absent) basal grainstones overlain by wackestone–mudstone, 

in which layering increases in abundance and becomes thinner 

upward into the laminated caps, where present. 

Oolite-dominated sequences include asymmetrical, upwardshallowing 

units (3 to 10 m thick) bounded at the base or top by 

quartz–peloid eolianites, and are composed of skeletal wackestone 

(commonly absent), up into ooid grainstone (with rare channeled 

bases and local mounds), which may be capped by argillaceous 

dolomitic wackestone–mudstone or dolomitic mudstone. Erosional 

disconformities may occur on the top of the lime 

wackestone–mudstone units, or on the ooid grainstone units. 

Symmetrical upward-deepening to upward-shallowing oolitic 

parasequences contain basal lime wackestone–mudstone conformably 

overlain by the upward-shallowing succession. 

Parasequences in skeletal limestone–rich parts of sequences 

(sequences C8 to C10) commonly are 2 to 15 m thick, and 

generally are bounded by erosional or sharp planar contacts. 

Such parasequences include asymmetrical, upward-shallowing 

parasequences composed of thick skeletal wackestone– 

packstone grading up into ooid grainstone (may be absent), 

capped by dolomitic wackestone–mudstone. Other skeletal limestone–dominated 

parasequences are symmetrical and contain 

basal, thin transgressive ooid or skeletal grainstone, or basal 

upward-deepening units of lime wackestone–mudstone, 

interbedded with thin grainstones. Parasequences that straddle 

the maximum flooding surfaces of sequences commonly are 

thick and symmetrical. 

Third-Order or Composite Sequences 

Third-order or composite sequences (Weber et al., 1995) are 

relatively weakly developed in the study interval compared to 

the regionally mappable fourth-order sequences. The third-order 

sequences (Figs. 6, 7) are bounded by red beds and their downdip 

marine sandstones updip. They thicken downdip into the 

foredeep, where the third-order sequences are bounded along the 

ramp margin and ramp slope by extensive lowstand siliciclastic 

units. The third-order sequences commonly contain two or more 

high-frequency sequences. 

Composite Sequence 1.—Composite sequence 1 is composed of 

the Meramecian Little Valley–Hillsdale interval. It extends from 

the top of the Price–Maccrady red beds, up to the red beds of 

sequence C1 and its downdip equivalents (Figs. 4, 6, 7). It onlaps 

much of the ramp, pinching out to a feather edge updip. The 

Hillsdale composite sequence thickens into the foredeep to 270 m 

(Bartlett and Webb, 1971; Al-Tawil, 1998), where it may consist of 

sequences dominated by shallow-water, fine-grained carbonates 

and interbedded sandstone (not shown on cross sections in Figs. 

6, 7). 

Composite Sequence 2.—This composite sequence includes sequences 

C1 to 4, which make up the Denmar Limestone. It 

extends from the base of the basal red beds (and their downdip 

equivalents) of sequence C1, up to the base of the lower Taggard 

red beds (and downdip equivalents) of sequence C5 (Figs. 4, 6, 7). 

The lowstand systems tract is made up of the lowstand siliciclastics 

of sequence C1. The transgressive systems tract probably includes 

the rest of sequences C1 to C3. The maximum flooding 

surface is likely at the base of sequence C4 updip, and beneath the 


C4 deeper-water mudstones downdip. The highstand systems 

tract is made up of sequence C4. 

25 

Composite Sequence 3.—Sequences C5 to 10 may be the third 

composite sequence, extending from the basal C5 red beds up to 

the top of C10 (Figs. 4, 6, 7). It is composed of the Pickaway, Union, 

and Alderson Limestones. The lowstand systems tract of the 

composite sequence consists of sequences C5 and C6, which are 

red-bed prone, and show a major basinward shift in both distribution 

and facies. The transgressive systems tract sequence of the 

composite consists of sequence C7 and the lower part of sequence 

C8. The maximum flooding surface is marked by the most landward 

extent of open-marine skeletal wackestone in sequence C8. 

There is a significant development of siliciclastics at the base of 

C9, including an updip tongue of red beds (Fig. 6, sections 1, 2). 

The highstand systems tract is made up of the upper half of 

sequences C8, C9, and C10. 

Composite Sequence 4.—This composite sequence likely extends 

from the base of the incised Fido Sandstone–Lillydale Shale and 

includes sequence C11 (Lillydale–Glenray interval); the top of the 

composite sequence is above the study interval (Figs. 4, 6, 7). The 

lowstand systems tract may include the incised siliciclastic units at 

the base of C11, and the maximum flooding surface is above the 

thin limestones at the base of C11 and below the fossiliferous 

shales. The highstand systems tract probably includes the Lillydale 

Shale, the Glenray Limestone, the Bickford Shale, and the Reynolds 

Limestone, and possibly units above the study interval. 

CONTROLS ON SEQUENCE DEVELOPMENT 

ON THE FORELAND 

Tectonics and Subsidence Rates 

Accommodation rates calculated from the present-day stratigraphic 

thicknesses (not decompacted) were highest (at least 10 

to 30 cm/ky) in the basin compared to the ramp margin (a 

minimum of 5 to 15 cm/ky) and were lowest (at least 1 to 3 cm/ 

ky) in updip areas of the distal foreland. This promoted regional 

slopes into the foreland basin. Regional slopes per sequence can 

be calculated using the difference in thickness of the study 

interval from updip to downdip, and dividing this by the distance 

times the number of sequences. This gives slopes of at least 

5 cm/km/sequence on the distal foreland (Al-Tawil and Read, 

this volume) and average slopes of at least 20 cm/km/sequence 

down the plunging basin axis. The low distal foreland slopes 

resulted in the predominance of shallow-water facies over the 

whole of the distal foreland. Calculated slopes per sequence 

measured over the palinspastically restored dip direction (northwest 

to southeast) down the ramp margin and into the basin were 

higher, at least 30 cm/km. These calculated slopes are compatible 

with minimum estimated highstand water depths of at least 50 m 

on the ramp slope, which periodically shallowed to sea level 

during lowstands. 

It has been suggested that the region was tectonically quiescent 

during deposition of the study interval, and in a state of 

elevational equilibrium between the filled foreland basin and 

the beveled Devonian orogenic highlands (Ettensohn, 1994). 

However, this appears to be incompatible with the considerable 

differential subsidence shown by the thickening of the carbonates 

from a feather edge updip on the foreland to over 900 m in 

the basin (Figs. 6, 7). The thickening takes place not only across 

the regional hinge but also at various localities updip and 

downdip from the hinge at various times. This complex pattern 

may have resulted from the movement of basement fault blocks

26 

beneath the foreland, which parallel the orogenic front 

(Shumaker and Wilson, 1996). However, given the long wavelengths 

of deformation, it also could just relate to flexural 

downwarping. Such downwarping in the foreland basin over 

the Price delta system occurred too long (some 10 to 15 My) after 

the initial early Mississippian loading event (Ettensohn, 1994) to 

be related to relaxation. This suggests that accommodation 

space might have been initiated during the Meramecian and 

early Chesterian by renewed thrust loading followed by relaxation 

(cf. Quinlan and Beaumont, 1984; Al-Tawil and Read, this 

volume). This tectonic loading and differential subsidence during 

deposition of the Meramecian sequences and Chesterian 

sequences C1 to C6 caused rapid thickness changes (Ettensohn, 

1975, 1980; MacQuown and Pear, 1983; Dever, 1986, Dever et al., 

1990; Shumaker and Wilson, 1996). In addition, there was a 

northwestward migration (approximately 50 km) of the major 

hinge onto the distal foreland as well as a northeastward migration 

of smaller hinge zones during the Meramecian–Chesterian 

(Figs. 6, 7). This would seem to support continued advance of a 

thrust (or sediment) load onto the foreland, suggesting that 

Greenbrier deposition occurred during active tectonics rather 

than during a time of tectonic quiescence. 

The regional subsidence that initiated Chesterian deposition, 

coupled with long-term sea-level rise (Ettensohn, 1994), moved 

the siliciclastic shorelines far updip during fourth-order 

highstands, resulting in widespread carbonate deposition on the 

ramp and siliciclastic deposition only during fourth-order 

lowstands/early transgressions. The marked differential subsidence 

exhibited by the Meramecian sequence and Chesterian 

sequences C1 to C6 (Figs. 6, 7) contrasts with the more gradual 

thickness changes of the upper sequences C7 to C10, which 

subsided more as a result of regional flexuring, possibly as 

tectonic loading slowed and/or regional uplift by erosional unloading 

of the orogenic highlands was initiated. Perhaps the 

earlier faults on the foreland had locked by this time and the 

foreland underwent regional flexure. Rapid flexure over the 

hinge also characterizes sequence C11 (Figs. 6, 7) and sequence 

C12 (above the study interval) and may mark renewed loading of 

the foreland. Following deposition of the Chesterian carbonates, 

erosional unloading and uplift of the orogenic highlands caused 

the subsequent progradation of Upper Mississippian Mauch 

Chunk siliciclastic units (Ettensohn, 1994). 

Differential subsidence thus controlled thickness and facies 

variations across the ramp, not only across the major ramp 

margin hinge but also at various places on the distal foreland, and 

thus was an important control on the accommodation created 

during deposition of each sequence. Broad sags periodically 

developed on the foreland in the northeast (sequences C2, C6, C7; 

Fig. 6) and locally on the distal foreland in the west (as in 

Meramecian sequence 2, and sequences C2, C3, and C9; Fig. 7). 

These sags and adjacent highs may have influenced the position 

of updip shoal-water grainstones separated by lagoonal lime 

wackestone–mudstone from the ramp-margin grainstones along 

the hinge. Although the differential subsidence modified the 

magnitudes of the relative sea-level changes, it did not override 

the eustatic changes that formed the sequence-bounding, 

siliciclastic-prone units and the carbonate-prone transgressive 

and highstand units. 

Climate 

Climate Control on Facies.—During the early Mississippian, 

the Appalachian Basin had a humid climate, which favored 

siliciclastic deposition on the Price–Borden delta complex 

(Branson, 1912; Butts, 1940, 1941; Sable and Dever, 1990). This 


climate subsequently became more arid during deposition of 

the Maccrady red beds and evaporites. Evaporites in the 

Meramecian interval in Kentucky (southwest of the study area) 

indicate that arid and semiarid conditions persisted throughout 

the middle Mississippian (Sable and Dever, 1990; Khetani and 

Read, 2002). This semiarid climate persisted into the early 

Chesterian during deposition of the lower part of the study 

interval (indicated by red beds, widespread oolitic ramps, coastal 

eolianites, and caliche in West Virginia, Kentucky, and the 

Illinois Basin (Dever, 1973, 1995; Ettensohn et al., 1988; Smith 

and Read, 1999). However, the climate may have become more 

humid into the later Chesterian, when disconformities in the 

Appalachian and Illinois basins became characterized by 

paleokarst (Ettensohn et al., 1988), coarse siliciclastic influx 

increased (Swann, 1964; Smith, 1996), skeletal limestones became 

more widespread, and thick oolitic deposits (which on 

ramps require semi arid conditions; Read, 1985) became much 

less common. Also, regional meteoric aquifers developed during 

this late Mississippian–Early Pennsylvanian humid phase 

(Niemann and Read, 1988). These climatic changes are borne 

out by paleogeographic reconstructions (de Witt and McGrew, 

1979; Scotese and McKerrow, 1990) that show the North American 

plate gradually rotating counterclockwise and moving northward 

from the Pangean arid sub-equatorial belt during the 

Meramecian to Chesterian (latitudes of 25 degrees or less), 

reaching close to the Equator by Pennsylvanian time. 

By latest Mississippian time, semiarid climate returned at 

least intermittently with the deposition of Mauch Chunk red beds 

and semiarid paleosols (Caudill et al., 1996). 

Climate Control on Long-Term Global Sea Levels.—Oxygen 

stable-isotope data from the mid-continent United States show 

several major climate events in the Meramecian–Chesterian, 

with increased δ 18 O values being interpreted as cooling and/or 

ice buildup events and decreased δ 18 O values representing 

warming events (Mii et al., 1999). There is a mid-Meramecian 

cooling event that likely marks the regressive event at the base 

of the Upper Meramecian sequences. This was followed by 

warming in the early Chesterian, which promoted transgression 

and deposition of the Meramecian sequences and Chesterian 

sequences C1 to C4. This was followed by a cooling event that 

likely corresponds to deposition of the red beds of sequences C5 

and C6 (Taggard red beds). Subsequent warming probably 

caused the long-term transgression and deposition of sequences 

C7 through C10. The subsequent cooling trend (Mii et al., 1999; 

Veevers and Powell, 1987) and falling sea levels were accompanied 

by progradation of Mauch Chunk siliciclastic-prone sequences 

(C11 and younger units). 

Eustasy 

Evidence for Fourth-Order Eustasy.—A set of relative-sea-level 

curves for the basin (Fig. 8) were constructed by arbitrarily 

assigning equal time to each of the sequences and plotting the 

relative magnitude of updip–downdip migration of marker facies 

such as oolitic or skeletal grainstones and sandstones. The 

resultant curves were compared to the global sea-level curves of 

Ross and Ross (1987) and a simplified Illinois Basin curve modified 

from Smith (1996) and Smith and Read (2000) from which the 

higher-frequency fluctuations were removed. In addition to the 

biostratigraphic control based on conodonts and crinoids 

(Collinson et al., 1971; Huggins, 1983; deWitt and McGrew, 1979; 

Sable and Dever, 1990), correlation between the Appalachian and 

the Illinois basins was aided by the presence of relatively regional 

unconformities and siliciclastic markers in Kentucky (McFarlan

VISEAN NAMURIAN 

CHESTERIAN 

MERAMECIAN 

ST. LOUIS 

and Walker, 1956) and in West Virginia; such markers include the 

Taggard red beds, and the Greenville and Lillydale shales (Smith 

and Read, 1999, 2000, 2001). 

The Upper Meramecian sequences in the proximal foreland 

correlate with the upper and lower St. Louis Limestone of the 

Illinois Basin (Huggins, 1983) but only one of the Upper 

Meramecian Hillsdale sequences appears to reach the distal 

foreland in Kentucky (Al-Tawil and Read, this volume). 

Most of the Chesterian fourth-order sequences in eastern 

West Virginia can be traced into much thinner disconformitybounded 

sequences exposed on the distal foreland, eastern Kentucky 

(Al-Tawil, 1998; Al-Tawil and Read, this volume), and into 

the Illinois Basin (Ettensohn et al., 1984; Smith, 1996; Smith and 

Read, 1999; 2001) (Fig. 8). 

The four West Virginia–Virginia sequences (C1 to C4) in the 

Playtycrinites penicillus–bearing Denmar interval (G. bilineatus–C. 

charactus conodont zone; Huggins, 1983; Stamm, 1997) can be 

recognized in the time-equivalent Ste. Genevieve Limestone, 

exposed on the distal foreland in Kentucky. However, they are 

difficult to match with the three Ste. Genevieve sequences in the 

Illinois Basin, although four sequences may occur there if the two 

disconformity-bounded units in Smith and Read’s (1999) sequence 

C3 are raised to sequence status. 

West Virginia sequence C5 (containing the lower Taggard 

red beds), sequences C6 (bearing the upper Taggard red beds– 

lower Pickaway) and C7 of West Virginia (upper Pickaway 

interval), which contain the first occurrences of the crinoid 


ST. LOUIS 

STE. GENEVIEVE 

APPALACHIAN BASIN 

GLOBAL SEA LEVEL 

ROSS AND ROSS (1987) 

ILLINOIS BASIN 

SMITH (1996) 

KENTUCKY 

DISTAL FORELAND 

(AL-TAWIL AND READ, 

THIS VOLUME) 

WEST VIRGINIA 

PROXIMAL FORELAND 

(THIS STUDY) 

HIGH LOW HIGH LOW HIGH LOW HIGH LOW 

GLEN DEAN 

GLEN DEAN 10 

GLEN DEAN 

HARDINSBURG SS. 

C11 

C11 GLEN RAY 

HANEY 

HANEY 9 

HANEY C10 

C10 ALDERSON 

BEECH CK 

BEECH CK 8 

CYPRESS SS 

BEECH CK C9 

CAVE BRANCH SH. 

C9 UPPER UNION 

REELSVILLE 7 

SAMPLE SS. 

BEAVER BEND 6 

BETHEL SS. 

REELSV. C8B 

C8A 

BEAVER 

BEND 

C8B 

C8A 

LOWER 

UNION 

5 

C7 

C7 

PAOLI 

PAOLI 

C6 

C6 

4 

BRYANTSVILLE BRECCIA 

3 LEVIAS AUX 

VASE 

C5 

C4 

C5 

C4 

U. TAGGARD 

L. TAGGARD 


2 

SPAR MT. 

SS 

C3 


C3 

DENMAR 

1 

C2 

C2 

PAOLI / 

WARIX RUN 

C1 

ST. LOUIS 

HILLSDALE 

FIG. 8.—Relative third-order and fourth-order sea-level curves for the study interval in the West Virginia–Virginia Appalachians 

compared with curves from the distal Appalachian foreland in Kentucky (Al-Tawil and Read, this volume), the Illinois Basin 

(Smith, 1999; 2000; 2001), and the third-order curves of Ross and Ross (1987). Most of the fourth-order cycles can be correlated 

between the basins. The Ross and Ross third-order curves can be tied grossly to the Appalachian third-order cycles. The Illinois 

Basin curves show some additional large sea-level falls defined on the degree of incision, which are not recognized in the 

Appalachian Basin. 

27 

Talarocrinus, are correlated with the three sequences mapped as 

“Warix Run”/Paoli in Kentucky by Ettensohn et al. (1984), and 

the Talarocrinus-bearing Paoli sequences in the Illinois Basin 

(Butts, 1922, 1940, 1941; Wells, 1950; Ettensohn et al., 1984; Smith 

and Read, 1999). 

West Virginia sequence C8 (Lower Union) corresponds to the 

Beaver Bend–Reelsville cycle in Kentucky and in the Illinois Basin 

(de Witt and McGrew, 1979; Ettensohn et al., 1984; Smith and 

Read, 2001; Al-Tawil and Read, this volume). Sequence C9 of 

West Virginia (mid-Union red beds and Upper Union Limestone), 

containing G. bilineatus–C. altus conodonts, was correlated 

with the Reelsville limestone of the Illinois Basin (de Witt and 

McGrew, 1979), but it more likely correlates with the Cave Branch 

Shale–Beech Creek sequence in Kentucky and the Cypress Sandstone–Beech 

Creek–Big Clifty sequence (Illinois Basin), containing 

basal G. bilineatus–C. altus zonal conodonts (Collinson et al., 

1971; Ettensohn et al., 1984). 

West Virginia sequence C10 (Greenville Shale–Alderson Limestone), 

containing the uppermost occurrence of Talarocrinus (Butts, 

1940; Wells, 1950), correlates with the Haney sequence of Kentucky 

and the Illinois Basin (de Witt and McGrew, 1979; Ettensohn 

et al., 1984; Smith and Read, 2001; Al-Tawil and Read, this 

volume). West Virginia sequence C11 (Lillydale Shale/Fido Sandstone–Glenray 

Limestone), which contains the G. bilineatus–K. 

mehli zonal conodonts (Rexroad and Clark, 1960) correlates with 

the Hardinsburg–Glen Dean cycle of Kentucky and the Illinois 

Basin (de Witt and McGrew, 1979; Ettensohn et al., 1984). 

C1 

PICKAWAY

28 

The interbasinal extent of these sequences suggests that global 

eustasy played a major role in their development, rather than 

tectonics alone, which would operate locally as the basins subsided 

differentially. Although better age constraints are needed, 

the high-frequency sequences appear to relate to fourth-order 

sea-level cycles. If with better dating these do prove to be 400 ky, 

then they may relate to Milankovitch-driven, long-term eccentricity, 

which would be expected to predominate during times of 

global ice sheets (Imbrie and Imbrie, 1986; Ruddiman et al., 1986). 

This is compatible with the onset of late Paleozoic Gondwana 

glaciation in the Southern Hemisphere, in the Early Chesterian or 

even earlier (Horbury, 1989; Frakes et al., 1992; Smith and Dorobek, 

1993; Wright and Vanstone, 2001). 

The data from the Illinois and Appalachian basins suggest 

that there was an increase in the magnitude of the sea-level 

changes through the Chesterian, peaking in the Pennsylvanian 

with sea-level changes of up to 100 m (Heckel, 1986). However, 

the presence of significant incision on some of the basal Chesterian 

units in the Illinois Basin may suggest that there were some large 

sea-level falls even this early (Leetaru, 2000; Smith and Read, 

2000, 2001) and possibly earlier in the Meramecian (Smith and 

Dorobek, 1993; Wright and Vanstone, 2001; Horbury, 1989) (Fig. 

8). This upward increase in magnitude of Late Mississippian 

fourth-order sea-level cycles was superimposed on the long-term 

Chesterian transgression, resulting in progressive landward shift 

of fourth-order highstands. However, the increasing sea-level 

falls kept returning the lowstand sand bodies to the general 

vicinity of the ramp margin and slope, even though long-term 

transgression was occurring. 

Evidence for Third-Order Eustasy.—Some of the weak, composite 

(third-order) sequences appear to correspond to third-order 

sea-level cycles documented as global events by Ross and Ross 

(1987) (Fig. 8). The third-order cycle that formed the Little Valley– 

Hillsdale composite sequence correlates with the Late Meramecian 

(St. Louis) sea-level cycle (Huggins, 1983). The third-order cycle 

that formed the Denmar composite sequence correlates with the 

Ste. Genevieve (earliest Chesterian) sea-level cycle, with a major 

lowstand at the position of the Appalachian Basin sequences C5 

and C6 red beds (lower and upper Taggard red beds). However, 

the later composite sequences are difficult to relate to the global 

cycles, in part because of insufficient detail in the Ross and Ross 

(1987) cycle chart. However, they do show a third-order sea-level 

cycle extending from the top of the Ste. Genevieve cycle up to the 

top of the Glen Dean (top of our C11), whereas we placed the top 

of the cycle at the base of C11 (base of Glenray). 

There also are substantial differences between the major 

incisional events in the Appalachian and Illinois basins. In the 

Illinois Basin, Leetaru (2000) recognized major incision in the 

middle and near the top of the Ste. Genevieve carbonates, whereas 

we see major lowstands only above the equivalent (Denmar) 

carbonates in the Appalachian Basin (at bases of C5 and C6). 

Incisional events in the Illinois Basin (Smith and Read, 2000, 2001) 

for which there is little evidence in the Appalachian Basin also 

include the Bethel Sandstone incision (base of our C8 sequence), 

and the incision at the base of the Cypress Sandstone, Illinois 

Basin (base of our C9), although there is a red-bed tongue at this 

level updip in the Appalachian Basin. Incision at the base of the 

Hardinsburg Sandstone, in the Illinois Basin, corresponds to the 

incision observed at the base of our sequence C11 (Fido Sandstone). 

The differences between the amount of incision between 

the Appalachian and the Illinois basins (Smith and Read, 1999, 

2000, 2001) may reflect the effects of uplift or subsidence acting on 

weak third-order sea-level changes, coupled with siliciclastic 

influx and incision associated with significant fluvial systems in 


the Illinois Basin, which may have accentuated the effects of some 

of the eustatic falls in the Illinois Basin. The greater subsidence 

rates at least in the proximal foreland of the Appalachian Basin 

may have tended to inhibit incision, as did the relative paucity of 

siliciclastic systems there. 

Development of Fourth-Order Sequences 

Development of Sequence Boundaries.—The disconformities and 

associated caliches and breccias that mark sequence boundaries 

updip in Kentucky as well as in the Illinois Basin (Ettensohn et al., 

1984; Ettensohn et al., 1988; Smith and Read, 1999, 2001; Al-Tawil 

and Read, this volume), appear to be less developed in the West 

Virginia study area, although this may be more apparent than 

real and may reflect the dominance of well data over outcrop data 

in this paper. However, it could also be due to the rapid mantling 

of exposed carbonates by siliciclastic veneers (eolian or sheetwash 

deposition) that inhibited paleosol formation. Local, 

calichified and brecciated units downdip along the ramp margin 

may reflect the low gradient of the ramp top, which allowed it to 

be exposed during sea-level falls. Sharp, planar to hummocky 

erosion surfaces that are developed on some of the grainstones at 

parasequence and sequence boundaries in outcrops downdip 

could reflect widespread marine cementation followed by erosion 

in rock platform settings during transgression. Significant 

meteoric cementation beneath disconformities did not occur 

until later in the Mississippian, when the climate became more 

humid (Niemann and Read, 1988). 

Development of Lowstand to Early Transgressive Tracts.—Although 

regional 3-D data are needed, it appears that siliciclastics 

for the lowstand/early transgressive sand bodies were sourced 

from highlands to the northeast and cratonic sources to the north, 

and they traveled southward toward the ramp margin (Fig. 1). It 

is not yet clear whether any were derived from the highlands 

bordering the eastern side of the foredeep directly east of the 

study area, inasmuch as this either would have required very 

large sea-level falls to allow transport across the foredeep. The 

distribution of regional red beds in sequences C5 and C6 (Taggard 

red beds) suggests that siliciclastics from the northeast were 

funneled down the eastern side of the foreland, which may have 

been a subtle topographic low that deepened into the foredeep to 

the southwest. Sequences C5 and C6 red beds were not deposited 

on the topographically higher (?) distal foreland to the west, 

which was the site of eolian and early transgressive shallow 

marine siliciclastic deposition. Siliciclastic veneers at bases of the 

sequences on the distal foreland could have been transported 

onto the ramp during lowstand, by winds or flash floods, but 

most of the siliciclastic units probably were subsequently reworked 

into early transgressive marine sheetlike units. Quartz– 

peloid eolianites may have been preserved (as in the outcropping 

distal foreland deposits of Kentucky; Al-Tawil and Read, this 

volume), perhaps because of early incipient cementation, which 

inhibited reworking, but others could have been reworked into 

marine quartz–peloid grainstone sheets, which would be difficult 

to differentiate from eolian deposits in cuttings. Some of the 

lowstand to early transgressive sandstone–calcareous siltstones 

were deposited as barriers behind which thin, nearshore shales 

were deposited, as in sequences C7 to C10. 

Because of the poor constraints on the water depths of the 

facies, estimates of sea-level fall forming the lowstand deposits are 

crude. Where relative sea-level falls were small to moderate (say 10 

to 15 m or so), lowstand marine sandstones that are only slightly 

shallower than the underlying highstand facies were deposited, 

and there was only a small basinward shift in facies. These include

lowstands at the bases of sequences C2, C7, and C8. Where relative 

sea-level falls were larger (say 20 to 40 m or so), then forced 

regression caused a significant basinward shift of facies, and 

lowstand shallow-water ooid grainstones or marine sandstones 

were deposited on deeper-water skeletal wackestone or laminated 

shaly lime mudstone, with intervening facies not being deposited, 

as at the bases of sequences C1, C6, C9, C10, and C11. 

The localization of lowstand wedges on the ramp margin was 

controlled by high differential subsidence over the hinge and onto 

the proximal foreland during deposition of the lower sequences. 

During deposition of sequences C1 to C6 (Denmar and lower 

Pickaway sequences), where the underlying highstand skeletal 

bank facies were located close to the ramp margin, lowstand sands 

migrated onto the ramp slope as sea level fell. However, during 

deposition of sequences C7 and C8, the ramp margin was farther 

updip, as a probable response to the long-term sea-level rise and 

activation of hinges farther updip. Thus, sea level had to fall farther 

to expose the ramp margin and slope, so that lowstand siliciclastics 

reached only the ramp margin. However, the increasingly large 

sea-level falls during deposition of sequences C9 to C11, coupled 

with the broad flexuring of the ramp, rather than differential 

subsidence, could have returned deposition of lowstand quartz 

sandstone and oolite units to the ramp slope. The most well 

developed lowstand siliciclastic units appear to coincide with 

third-order sea-level falls, as in the base of sequence C1 (base of the 

Denmar), the bases of sequences C5 and C6 (Taggard red beds), 

and the base of sequence C11 (Fido Sandstone). 

Transgressive Systems Tract Development and Maximum Flooding.—Although 

the maximum flooding surfaces of each of the 

sequences were difficult to pick at all locations, given the mix of 

subsurface well data and higher-quality outcrop data, some 

generalizations can be made. Along the ramp slope, the coincidence 

of maximum flooding surfaces with the transgressive 

surface on the lowstand systems tract (as in sequences C 4, C10 

and C11) appears to reflect rapid deepening and drowning of the 

lowstand facies as relative sea level rose and the ramp was 

transgressed. In contrast, in most of the other sequences, transgressive 

wackestone units, or transgressive–regressive skeletal 

wackestone to skeletal grainstone (C1, C3, C6) were developed on 

the lowstand deposits, reflecting more gradual relative-sea-level 

rise during fourth-order transgression. If the skeletal units on the 

ramp slope developed solely under the influence of gradually 

rising sea level, they would be expected to be upward deepening 

into the overlying deeper-water facies. However, the skeletal 

banks appear to be shallowing-upward successions and probably 

reflect superimposed higher-frequency sea-level fluctuations 

or perhaps a slowing of subsidence. 

On the shallow ramp, some fourth-order sea-level rises reworked 

basal siliciclastic units but did not locally deposit a 

transgressive succession of carbonates, although such transgressive 

deposits formed within some of these same sequences elsewhere. 

Suppression of transgressive carbonate deposition could 

have been due to rapid rise rates that did not allow time for a 

significant transgressive carbonate to be deposited, or perhaps 

the shallow marine waters were too turbid with resuspended 

terrigenous silt. Deposition of a transgressive carbonate succession 

on the locally developed, eolian and marine reworked 

siliciclastic units could have been favored by slower rise of rates 

of relative sea level and waters becoming clearer as transgression 

proceeded. The presence of upward-shallowing units that make 

up the transgressive carbonates in some sequences that have 

significant transgressive systems tracts (sequence C2, C7, and C8; 

Fig. 6), suggests that transgressive carbonate deposition may 

have been favored by a short-term rise and fall superimposed on 


the fourth-order rise, which effectively delayed maximum flooding 

of the ramp. 

29 

Highstand Deposition.—During highstand deposition of the 

sequences, a complex facies mosaic developed, rather than a 

simple upward-shallowing, progradational succession that would 

result from seaward migration of shore-parallel facies belts. 

Skeletal grainstone–packstone and ooid grainstone were deposited 

not only along the ramp margin but also as extensive but 

discontinuous bodies far back from the margin, from which they 

were separated by extensive lagoonal (and locally tidal flat) lime 

wackestone–mudstone. Some of these updip skeletal and ooid 

bodies could have formed bordering highs adjacent to locally 

downwarped areas of the foreland (Carney, 1993), but others 

could have been localized by subtle depositional topography. 

Individual ramp-margin ooid bodies were shaped by tidal currents 

in a northeast–southwest direction normal to depositional 

strike (Kelleher and Smosna, 1993). 

Lagoonal lime wackestone–mudstones were widely deposited 

in protected lagoons behind ramp-margin ooid and skeletal 

grainstone shoals and far updip between and behind shoals. The 

absence of tidal-flat laminites from much of the distal foreland 

probably was due to the relatively large sea-level falls, which 

caused rapid seaward shift of the shoreline, faster than tidal flats 

could prograde. However, locally thick tidal-flat facies were 

deposited just behind the ramp margin as the lagoonal muds 

shallowed upward, possibly because of high sedimentation rates 

here, coupled with the low rates of relative sea-level fall, which 

would characterize late highstands, and the relatively high subsidence 

rates, which would have decreased the effects of eustatic 

fall. 

Development of Late Meramecian–Chesterian 

Composite Sequences 

In the study area, the Kinderhook–middle Osagean Borden– 

Price delta complex was not overlain by upper Osagean–lower 

Meramecian carbonates of the Fort Payne–Salem supersequence, 

which were deposited farther to the south. Thus, the upper 

Meramecian–Chesterian supersequence in this paper was deposited 

directly onto the earlier Price–Borden siliciclastic units, following 

rapid differential subsidence of the delta in the east, 

coupled with global sea-level rise (Dennison, 1985; Ettensohn, 

1994), which led to increased accommodation and long-term 

onlap of the upper Meramecian and Chesterian sequences onto 

the delta complex. 

Differential subsidence over the eastern hinge, coupled with 

relatively small third-order sea-level fluctuations which started 

from the low sea-level position at the end of the upper Osagean– 

lower Meramecian supersequence, limited deposition of composite 

sequence C1 (Little Valley–Hillsdale) to the margin of the 

distal foreland to the west. Deposition extended farther updip in 

the east, where subsidence was more rapid. The small superimposed 

fourth-order sea-level changes periodically flooded the 

ramp to shallow depths, resulting in deposition of dominantly 

muddy, restricted carbonates of composite sequence C1. 

The subsequent sea-level lowstand initiated localized red bed 

and/or eolianite deposition on the southern and eastern edge of 

the distal foreland, which continued to undergo differential 

subsidence over the hinge during the four transgressive–regressive 

events that formed composite sequence C2. Long-term relative 

sea-level rise caused progressive backstepping of the updip 

limits of sequences C1 to 4 onto the distal foreland, while considerable 

differential subsidence across the hinge promoted deepwater 

deposition in the southeast. The small, fourth-order sea-

30 

level changes only shallowly flooded the distal foreland to the 

west, with deposition of shallow lagoonal muds and intervening 

ooid shoals. More open marine skeletal facies were deposited in 

the more rapidly subsiding areas southeast of the hinge, and to 

the northeast updip along the basin axis. 

With long-term sea-level fall, widespread red beds and 

eolianites were deposited, initiating deposition of composite 

sequence C3 (Taggard red beds). The earlier hinge became less 

active during C5 deposition but was reactivated during C6 deposition, 

along with differential subsidence seaward of a local hinge 

updip to the northeast. The sea-level rises from the initial lowstand 

positions (base C5 and base C6) flooded the ramp only to shallow 

depths, so that restricted facies were deposited over much of the 

ramp. However, long-term sea-level rise, coupled with increasing-amplitude 

fourth-order eustasy, caused sequences C5 to C10 

to continue to onlap updip, with maximum flooding apparently 

occurring in the middle of sequence C8, followed by slight offlap 

during deposition of subsequent sequences (C9 and C10). These 

increased-amplitude fourth-order cycles, coupled with slight 

westward and northward migration of the hinges, promoted 

backstepping of ramp-margin facies and deposition of relatively 

open marine facies over the seaward edge of the distal foreland 

during C7 to C8 and perhaps C9 deposition. The decrease in 

deposition of ooid grainstone from sequence C9 to C11 may be 

due to the increasingly humid climate. As amplitudes of sea-level 

change increased, increased sea-level falls allowed lowstand 

siliciclastics and ooid grainstones to be deposited farther downdip 

during deposition of sequences C9 to base of C11. 

Composite sequence C4 was initiated with incision and deposition 

of sand bodies at the base of sequence C11, probably 

because of very low sea level. Increasingly humid conditions, 

coupled perhaps with uplift in the source areas to the northeast, 

generated the shale-prone sequences C11 and C12 (above the 

study interval). Thickening of at least sequence C11 into the 

foredeep may reflect a return to higher differential subsidence 

associated with the major hinge, culminating in thick, deep-water 

carbonates along the basin margin. Carbonate for these must 

have been shed from thick ramp-margin bank complexes along 

the ramp margin in the southeast (exposed along Pine Mountain, 

Kentucky; Al-Tawil and Read, this volume). Renewed, relatively 

deep submergence during deposition of sequence C11 (Figs. 6, 7) 

probably reflects this increased subsidence coupled with the 

relatively high-amplitude fourth-order sea-level changes. Shaleprone 

sequence C11, as well as the overlying, carbonate-poor 

Mauch Chunk siliciclastic sequences above the study interval, 

formed as long-term sea level started to fall, global climate cooled 

(Mii et al., 1999), climate became more humid, and erosional 

unloading of the orogenic belt occurred (Ettensohn, 1994), all of 

which promoted terrigenous influx onto the ramp and siliciclastic 

progradation. 

Moderate-Amplitude Glacio-Eustasy 

and Late Mississippian Ramps 

The study interval spans the onset of the Southern Hemisphere 

Gondwana glaciation (Fig. 9), and the estimated 

periodicities of the sequences and the facies developed suggest 

that moderate-amplitude glacio-eustasy forced by long-term eccentricity 

(Horbury, 1989; Frakes et al., 1992; Smith and Dorobek, 

1993; Miller and Eriksson, 1997; Smith and Read, 2000, 2001; 

Wright and Vanstone, 2001) may have controlled the formation of 

the sequences. This moderate-amplitude glacio-eustasy, which 

increased in magnitude with time, was transitional into the highamplitude 

glacio-eustasy of the Pennsylvanian (Fischer, 1982; 

Goldhammer et al., 1991; Heckel, 1986). The moderate-amplitude 


sea-level falls of the Late Mississippian caused large-scale basinward 

migration of siliciclastic shorelines, and caused lowstand to 

early transgressive siliciclastics to be deposited on the ramp, 

ramp margin, and ramp slope. The increasingly large, moderateamplitude 

sea-level rises superimposed on a long-term eustatic 

rise, and the load-induced foreland subsidence, caused extensive, 

periodic floodings of the ramp. However, the sea-level rises 

were not sufficiently large to drown the ramps, so that widespread 

grainstones were deposited not only on the ramp margin 

but also over large areas far updip on the ramp, where they 

interfinger laterally with muddy lagoonal carbonates. Ramps 

that formed under such moderate-amplitude eustasy contrast 

with greenhouse ramps (Fischer, 1982), on which significant 

grainstone units tend to be limited to the outer part of the shallow 

ramp, and parasequences are meter-scale cycles dominated by 

restricted, muddy shallowing-upward facies (Read, 1995; 

Koerschner and Read, 1989; Goldhammer et al., 1990). At the 

other end of the spectrum, the bulk of the sequences of the 

Mississippian ramp in West Virginia lack the widespread, onramp, 

deeper-water maximum-flooding black shale–lime mudstone 

facies of the Pennsylvanian cyclothems, typical of higheramplitude 

glacio-eustasy during icehouse times (Heckel, 1986; 

Smith and Read, 2000, 2001). 

CONCLUSIONS 

This study illustrates how well cuttings and wireline logs, 

integrated with limited core and outcrop data, can be used to 

generate high-resolution sequence stratigraphies of Mississippian 

ramps in the subsurface. The Meramecian to Chesterian 

Greenbrier succession in the West Virginia portion of the Appalachian 

Basin formed on a tropical, mixed carbonate–siliciclastic 

ramp that developed on a subsiding foreland that was undergoing 

differential subsidence, during the initiation of significant 

Southern Hemisphere, Gondwana, glaciation. The sequence 

stratigraphy was analyzed using two regional cross sections, one 

a dip section and the other an axial section, which were based on 

subsurface well cuttings and wireline logs from numerous shallow 

wells, several detailed outcrop sections, and a core. The 

Greenbrier succession forms the transgressive part of the upper 

Meramecian–Chesterian supersequence, and is made up of fourthorder 

depositional sequences stacked into variably developed 

third-order, composite sequences. The fourth-order depositional 

sequences reach thicknesses of tens of meters to a hundred 

meters. They are bounded by lowstand to early transgressive 

siliciclastic units and local updip eolianites. Transgressive banks 

developed along the ramp margin locally, but on the shallow 

ramp, transgressive shallow-water carbonates were recognized 

in some areas but not in others. Highstand facies are extensive 

lagoonal lime wackestone–mudstone, along with ooid grainstone 

and skeletal grainstone–packstone that developed along the ramp 

margin and as extensive isolated units on the ramp interior, and 

constitute the bulk of many sequences. Skeletal wackestone grading 

seaward into laminated, argillaceous lime mudstone developed 

on the deeper ramp and ramp slope. 

Deposition of the succession was influenced by differential 

subsidence of the foredeep, possibly marking a phase of late 

Paleozoic thrust loading. Paleoslopes on the distal foreland were 

a few centimeters per kilometer, increasing to 30 cm/km over the 

hinge. Accommodation rates were a few centimeters per kiloyear 

on the distal foreland increasing to 10 to 30 cm/ky on the 

proximal foreland. Tectonic subsidence generated a basinwardthickening 

wedge of sediments 0 to 900 m thick. Rapid thickening 

of many of the sequences occurs across the basin hinge, as well as 

across broad downwarps on the ramp, which were intermittently

active. This differential subsidence generated a complex mosaic 

of facies on the ramp; thus, sequences rarely show simple upward-deepening 

to upward-shallowing successions, or uniformly 

basinward-thickening trends. 

The fourth-order sea-level fluctuations that formed the sequences 

likely relate to glacially driven Milankovitch long-term 

eccentricity forcing. They were responsible for development of 

the characteristic sequence stacking with basal siliciclastic-prone 

units and upper carbonate-prone units, because they caused 

large shifts in the shoreline across the ramp. Sea-level changes 

became larger through time, which, along with the increasingly 

humid climate, caused an upsection change in the supersequence 

from dominantly restricted, ooid grainstone–prone sequences to 

more open marine, skeletal limestone–prone sequences with 

significant siliciclastic units. Third-order sea-level falls appear to 

have accentuated fourth-order falls, causing widespread deposition 

of red beds and marine siliciclastic units, and some incision 

of underlying units. Although subsidence rates on the proximal 

foreland were an order of magnitude greater than those acting on 

the distal foreland, fourth-order sea-level fluctuations (indicated 

by interbasinal correlation) still were the prime cause of the 

sequences, whereas tectonics strongly controlled regional and 

local thickness and facies variations. 

ACKNOWLEDGMENTS 

We would like to thank the West Virginia Economic and 

Geological Survey and especially Lee Avery, Dave Matchen, Pat 

Johns, and staff, for cuttings, cores, maps, and digital data files. We 

give special thanks to Taury Smith, whose outstanding Illinois 

Basin work helped clarify many of the problems of Mississippian 

facies and stratigraphy, as did the work of Garland Dever of the 

Kentucky Geological Survey. Bill Henika (Virginia Division of 

Mineral Resources) provided valuable insight into problems of 

Appalachian geology. Jim Molloy, Art Hall, Sarah Crews, Steve 

Turpin, Jane Roa, John Leone, Chad Haiar, Susan Barbor, Pat 

Barnhart, Don Thorsteson, Ashley Goodrich, Julia Caldwell, and 

Jen Hatfield helped collect and process data. Mike Pope, Anna 

Balog, Amy Khetani, Dan Miller, and Brian Coffey provided 

valuable discussion and help, and Ellen Mathena did much of the 

editing and typing. The study was supported by Petroleum Research 

Fund grant 34137-AC8, and National Science Foundation 

grant EAR-9725308 to J.F. Read, and funding from Mobil Oil with 

the strong support of Ron Kreisa and Jim Markello. We thank 

American Association of Petroleum Geologists, Geological Society 

of America, and the Society of Professional Well Log Analysts for 

additional support. Dave Hunt, Steve Dorobek, and Wayne Ahr 

greatly improved the paper with their insightful reviews. 

REFERENCES 

AHR, W.M., 1973, The carbonate ramp—an alternative to the shelf model: 

Gulf Coast Association of Geological Societies, Transactions, v. 23, p. 

221–225. 

AL-TAWIL, A., 1998, High resolution sequence stratigraphy of Late Mississippian 

carbonates in the Appalachian Basin: Unpublished Ph.D. 

dissertation, Virginia Tech, Blacksburg, 109 p. 

ARKLE, T., JR., BEISSELL, D.R., LARESE, R.E., NUHFER, E.B., PATCHEN, D.G., 

SMOSNA, R.A., GILLESPIE, W.H., LUND, R., NORTON, W., AND PFEFFERKORN, 

H.W., 1979, The Mississippian and Pennsylvanian (Carboniferous) 

Systems in the United States—West Virginia and Maryland: U.S. 

Geological Society, Professional Paper 1110-D, 35 p. 

BARTLETT, C.S., 1974, Anatomy of the Lower Mississippian Delta in southwestern 

Virginia: Unpublished Ph.D. dissertation, University of Tennessee, 

Knoxville, 373 p. 


31 

BARTLETT, C.S., JR., AND WEBB, H.W., 1971, Geology of the Bristol and 

Wallace quadrangles, Virginia: Virginia Division of Mineral Resources, 

Report of Investigations 25, 93 p. 

BAXTER, J.W., AND BRENCKLER, P.L., 1982, Preliminary statement on Mississippian 

calcareous foraminiferal successions of the midcontinent 

(U.S.A.) and their correlation to western Europe: Newsletters on 

Stratigraphy, v. 11, p. 136–153. 

BJERSTEDT, T.W., AND KAMMER, T.W., 1988, Genetic stratigraphy and depositional 

systems of the upper Devonian–lower Mississippian Price– 

Rockwell deltaic complex in the central Appalachians, U.S.A.: Sedimentary 

Geology, v. 54, p. 265–301. 

BRANSON, E.B., 1912, A Mississippian delta: Geological Society of America, 

Bulletin, v. 23, p. 447–456. 

BREZINSKI, D.K., 1989, Upper Mississippian depositional patterns in the 

north-central Appalachian Basin, and their implications to 

Chesterian hierarchical stratigraphy: Southeastern Geology, v. 30, 

p. 1–23. 

BUTTS, C., 1922, The Mississippian Series of eastern Kentucky: Kentucky 

Geological Survey, ser. 6, v. 7, 188 p. 

BUTTS, C., 1940, Geology of the Appalachian Valley in Virginia (Part I— 

Geologic Text and Illustrations): Virginia Geological Survey, Bulletin 

52, p. 1–568. 

BUTTS, C., 1941, Geology of the Appalachian Valley in Virginia (Part II— 

Fossil Plates and Explanations): Virginia Geological Survey, Bulletin 

52, p. 1–568. 

CARNEY, C., 1993, The drowning of ooid shoals: Mississippian Greenbrier 

Limestone near the West Virginia Dome, in Keith, B.D., and Zuppman, 

C.W., eds., Mississippian Oolites and Modern Analogs: American 

Association of Petroleum Geologists, Studies in Geology no. 35, p. 

141–148. 

CARNEY, C., AND SMOSNA, R., 1989, Carbonate deposition in a shallow 

marine gulf, the Mississippian Greenbrier Limestone of the Central 

Appalachian Basin: Southeastern Geology, v. 30, p. 25–48. 

CAUDILL, M.R., DRIESE, S.G., AND MORA, C.I., 1996, Preservation of a paleo- 

Vertisol and an estimate of late Mississippian paleoprecipitation: 

Journal of Sedimentary Research, v. 66, p. 58–70 

COLLINSON, C.W., REXROAD, C.B., AND THOMPSON, T.L., 1971, Conodont 

zonation of the North American Mississippian, in Sweet, W.C., and 

Bergstrom, S.M., eds., Symposium on Conodont Biostratigraphy: 

Geological Society of America, Memoir 127, p. 353–394. 

COLTON, G.W., 1970, The Valley and Ridge and Appalachian Plateau; 

stratigraphy and sedimentation; the Appalachian Basin; its depositional 

sequences and their geologic relationships, in Fisher, G.W., 

Pettijohn, F.J., Reed, J.C., Jr., and Weaver, K.N., eds., Studies of 

Appalachian Geology, Central and Southern: Wiley-Interscience, p. 

5–47. 

COOPER, B.N., 1948, Status of Mississippian stratigraphy in the central and 

northern Appalachian region, in Weller, J.M., ed., Symposium on 

Problems of Mississippian Stratigraphy and Correlation: Journal of 

Geology, v. 56, p. 255–263. 

CROWELL, J.C., 1978, Gondwanan glaciation, cyclothems, continental positioning, 

and climate change: American Journal of Science, v. 278, p. 

1345–1378. 

DE WITT, W., AND MCGREW, L.W., 1979, The Appalachian Basin, in Craig, 

L.C., and Connor, C.W., coordinators, Paleotectonic Investigations of 

the Mississippian System in the United States: U.S. Geological Survey, 

Professional Paper 1010, part I, p. 59–106. 

DENNISON, J.M., 1985, Catskill delta shallow marine strata, in Woodrow, 

D.L., and Sevon, W.D., eds., The Catskill Delta: Geological Society of 

America, Special Paper 201, p. 91–106. 

DEVER, G.R., JR., 1973, Stratigraphic relationships in the lower and 

middle Newman Limestone (Mississippian), east-central and northeastern 

Kentucky: Unpublished M.S. Thesis, University of Kentucky, 

Lexington, 121 p.

32 

DEVER, G.R., JR., 1986, Mississippian reactivation along the Irvine–Paint 

Creek Fault System in the Room Trough, east central Kentucky: 

Southeastern Geology, v. 27, no. 2, p. 8–18. 

DEVER, G.R., 1995, Tectonic implications of erosional and depositional 

features in upper Meramecian and lower Chesterian (Mississippian) 

carbonate rocks of south-central and east-central Kentucky: 

Unpublished Ph.D. Dissertation, University of Kentucky, Lexington, 

152 p. 

DEVER, G.R., GREB, S.F., MOODY, J.R., CHESTNUT, D.R., KEPFERLE, R.C., AND 

SERGEANT, R.E., 1990, Tectonic implications of depositional and 

erosional features in Carboniferous rocks of south-central Kentucky: 

Field Guide for Annual Field Conference of the Geological 

Society of Kentucky, 53 p. 

DIRENZO, V.N., JR., 1986, Petrology and depositional environments of 

the Little Valley Limestone (Upper Mississippian) Washington 

County, Virginia: Unpublished M.S. Thesis, East Carolina University, 

Greenville, 154 p. 

DODD, J.R., ZUPPMAN, C.W., HARRIS, C.D., LEONARD, K.W., AND BROWN, 

T.W., 1993, Petrologic method for distinguishing eolian and marine 

grainstones, Ste. Genevieve Limestone (Mississippian) of Indiana, 

in Keith, B.D., and Zuppman, C.W., eds., Mississippian Oolites and 

Modern Analogs: American Association of Petroleum Geologists, 

Studies in Geology no. 35, p. 49–59. 

ETTENSOHN, F.R., 1975, Stratigraphic and paleoenvironmental aspects of 

Upper Mississippian rocks (Upper Newman Group), east-central 

Kentucky: Unpublished Ph.D. Thesis, University of Illinois, Urbana, 

320 p. 

ETTENSOHN, F.R., 1980, An alternative to the barrier-shoreline model for 

deposition of Mississippian and Pennsylvanian rocks in northeastern 

Kentucky: Geological Society of America, Bulletin, v. 91, pt. 1, p. 

130–135; pt. 2, p. 934–1056. 

ETTENSOHN, F.R., 1994, Tectonic control on formation and cyclicity of 

major Appalachian unconformities and associated stratigraphic 

sequences, in Dennison, J.M., and Ettensohn, F.R., eds., Tectonic and 

Eustatic Controls on Sedimentary Cycles: SEPM, Concepts in Sedimentology 

and Paleontology, no. 4, p. 217–242. 

ETTENSOHN, F.R., DEVER, G.R., AND GROW, J.S., 1988, A paleosol interpretation 

for profiles exhibiting subaerial exposure crusts from the 

Mississippian of the Appalachian basin, in Reinhart, J., and Sigleo, 

W.R., eds., Paleosols and Weathering Through Geologic Time; Principles 

and Applications: Geological Society of America, Special 

Paper 216, p. 49–79. 

ETTENSOHN, F.R., RICE, C.L., DEVER, G.R., AND CHESTNUT, D.R., 1984, Slade 

and Paragon Formations—new stratigraphic nomenclature for Mississippian 

rocks along the Cumberland Escarpment in Kentucky: 

U.S. Geological Survey, Bulletin 1605-B, 37 p. 

FISCHER, A.G., 1982, Long-term climatic oscillations recorded in stratigraphy, 

in climate in Earth history: Washington, D.C., National 

Academy Press, p. 97–104. 

FLOWERS, R.R., 1956, A regional subsurface study of the Greenbrier 

Limestone in West Virginia: Unpublished M.S. Thesis, University of 

West Virginia, Morgantown, 220 p. 

FRAKES, L.A., FRANCIS, J.E., AND SYKTUS, J.L., 1992, Climate Modes of 

the Phanerozoic: Cambridge, U.K., Cambridge University Press, 

274 p. 

GOLDHAMMER, R., DUNN, K., AND HARDIE, L.A., 1990, Depositional cycles, 

composite sea level changes, cycle stacking patterns, and the hierarchy 

of stratigraphic forcing: examples from Alpine Triassic platform 

carbonates: Geological Society of America, Bulletin, v. 102, p. 535– 

562. 

GOLDHAMMER, R.K., OSWALD, E.J., AND DUNN, P.A., 1991, Hierarchy of 

stratigraphic forcing: Example from Middle Pennsylvanian shelf 

carbonates of the Paradox basin, in Franseen, E.K., Watney, W.L., 

Kendall, C.G.St.C., and Ross, W., eds., Sedimentary Modeling; Com- 


puter Simulations and Methods for Improved Parameter Definition: 

Kansas Geological Survey, Bulletin 233, p. 361–414. 

GOLONKA, J., ROSS, M.I., AND SCOTESE, C.R., 1994, Phanerozoic paleogeographic 

and paleoclimatic modeling maps, in Embry, A.F., 

Beauchamp, B., and Glass, D.J., eds., Pangea; Gobal Environments 

and Resources: Canadian Society of Petroleum Geologists, Memoir 

17, p. 1–47. 

HARRISON, R.S., AND STEINEN, R.P., 1978, Subaerial crusts, caliche profiles, 

and breccia horizons: comparison of some Holocene and Mississippian 

exposure surfaces, Barbados and Kentucky: Geological Society 

of America, Bulletin, v. 89, p. 385–396. 

HECKEL, P.H., 1986, Sea-level curve for Pennsylvanian eustatic marine 

transgressive–regressive depositional cycles along the midcontinent 

outcrop belt, North America: Geology, v. 14, p. 330–334. 

HORBURY, A.D., 1989, The relative roles of tectonism and eustasy in the 

deposition of the Urswick Limestone in south Cumbria and north 

Lancashire, in Arthurton, R.S., Gutteridge, P., and Nolan, S.C., eds., 

The Role of Tectonics in Devonian and Carboniferous Sedimentation 

in the British Isles: Yorkshire Geological Society, Occasional 

Publication no. 6, p. 153–169. 

HUGGINS, M.J., 1983, Meramecian conodonts and biostratigraphy of the 

(Upper Mississippian) Greenbrier Limestone (Hurricane Ridge and 

Greendale synclines), south-western Virginia and southern West 

Virginia: Unpublished Ph.D. Dissertation, Virginia Tech, Blacksburg, 

303 p. 

HUNTER, R.E., 1993, An eolian facies in the Ste. Genevieve Limestone of 

southern Indiana, in Keith, B.D., and Zuppman, C.W., eds., Mississippian 

Oolites and Modern Analogs: American Association of 

Petroleum Geologists, Studies in Geology no. 35, p. 31–48. 

IMBRIE, J., AND IMBRIE, K.P., 1986, Ice Ages; Solving the Mystery: Cambridge, 

Massachusetts, Harvard University Press, 224 p. 

KEITH, B.D., AND ZUPPMAN, C.W., 1990, eds., Mississippian Oolites and 

Modern Analogs: American Association of Petroleum Geologists, 

Studies in Geology no. 35, 265 p. 

KELLEHER, G.T., AND SMOSNA, R., 1993, Oolitic tidal-bar reservoirs in the 

Mississippian Greenbrier Group of West Virginia, in Keith, B.D., 

and Zuppman, C.W., eds., Mississippian Oolites and Modern Analogs: 

American Association of Petroleum Geologists, Studies in 

Geology no. 35, p. 163–173. 

KHETANI, A.B., AND READ, J.F., 2002, Sequence development of a mixed 

carbonate–siliciclastic high-relief ramp, Mississippian, Kentucky, 

U.S.A.: Journal of Sedimentary Research, v. 72, p. 657–672. 

KOERSCHNER, W.F., III, AND READ, J.F., 1989, Field and modelling studies 

of Cambrian carbonate cycles, Virginia Appalachians: Journal of 

Sedimentary Petrology, v. 59, p. 654–687. 

LEONARD, A.D., 1968, The petrology and stratigraphy of Upper Mississippian 

Greenbrier Limestones of eastern West Virginia: Unpublished 

Ph.D. Thesis, University of West Virginia, Morgantown, 

220 p. 

LEETARU, H.E., 2000, Sequence stratigraphy of the Aux Vases Sandstone: 

A major oil producer in the Illinois Basin: American Association of 

Petroleum Geologists, Bulletin, v. 84, p. 399–422. 

MACQUOWN, W.C., AND PEAR, J.L., 1983, Regional and local geologic 

factors control Big Lime stratigraphy and exploration for petroleum 

in eastern Kentucky, in Luther, M.K., ed., Proceedings of the technical 

sessions, Kentucky Oil and Gas Association, 44th Annual Meeting, 

June 11–13, 1980: Kentucky Geological Survey, ser. 11, Special 

Publication 9, p. 1–20. 

MAPLES, C.G., AND WATERS, J.A., 1987, Redefinition of the Meramecian/ 

Chesterian Boundary (Mississippian): Geology, v. 15, p. 647–651. 

MCFARLAN, A.C., SWANN, D.H., WALKER, F.H., AND NOSOW, E., 1955, Some 

old Chester problems—correlations of lower and middle Chester 

formations of western Kentucky: Kentucky Geological Survey Series 

IX, Bulletin 16, 37 p.

MCFARLAN, A.C., AND WALKER, F.H., 1956, Some old Chester problems— 

correlations along the eastern belt of outcrop: Kentucky Geological 

Survey, series IX, Bulletin 20, 36 p. 

MERKELY, P.A., 1991, Origin and distribution of carbonate eolianites in 

the Ste. Genevieve Limestone (Mississippian) of southern Indiana 

and northwestern Kentucky: Unpublished M.S. Thesis, Indiana 

University, Bloomington, 147 p. 

MII, H., GROSSMAN, E.L., AND YANCEY, T., 1999, Carboniferous isotope 

stratigraphies of North America: Implications for Carboniferous 

paleoceanography and Mississippian glaciation: Geological Society 

of America, Bulletin, v. 111, p. 960–973. 

MILLER, D.J., AND ERIKSSON, K.A., 1997, Late Mississippian prodeltaic 

rhythmites in the Appalachian Basin: A hierarchical record of tidal 

and climatic periodicities: Journal of Sedimentary Research, v. 67, p. 

653–660. 

MILLER, D.J., AND ERIKSSON, K.A., 1999, Linked sequence development 

and global climate change; the Upper Mississippian record in the 

Appalachian Basin: Geology, v. 27, p. 35–38. 

MITCHUM, R.M., JR., AND VAN WAGONER, J.C., 1991, High frequency 

sequences and their stacking patterns: sequence-stratigraphic evidence 

of high-frequency eustatic cycles: Sedimentary Geology, v. 

70, p. 131–160. 

NIEMANN, J.C., AND READ, J.F., 1988, Regional cementation from unconformity-recharged 

aquifer and burial fluids, Mississippian Newman 

Limestone, Kentucky: Journal of Sedimentary Petrology, v. 58, p. 

688–705. 

PRYOR, W.A., AND SABLE, E.G., 1974, Carboniferous of the Eastern Interior 

Basin, in Briggs, G., Carboniferous of the Southeastern United 

States: Geological Society of America, Special Paper 148, p. 281–313. 

PURSER B.H., AND SEIBOLD, E., 1973, The principal environmental factors 

influencing Holocene sedimentation and diagenesis in the Persian 

Gulf, in Purser, B.H., ed., The Persian Gulf: Holocene Carbonate 

Sedimentation and Diagenesis in a Shallow Epicontinental Sea: 

New York, Springer-Verlag, p. 1–10. 

QUINLAN, G.M., AND BEAUMONT, C., 1984, Appalachian thrusting, lithospheric 

flexure and Paleozoic stratigraphy of the eastern interior of 

North America: Canadian Journal of Earth Sciences, v. 21, p. 973–976. 

READ, J.F., 1985, Carbonate platform facies models: American Association 

of Petroleum Geologists, Bulletin, v. 69, p. 1–21. 

READ, J.F., 1995, Overview of carbonate platform sequences, cycle stratigraphy 

and reservoirs in greenhouse and icehouse worlds, in Read, 

J.F., Kerans, C., Weber, L.J., Sarg, J.F., and Wright F.W., Milankovitch 

Sea Level Changes, Cycles and Reservoirs on Carbonate Platforms 

in Greenhouse and Icehouse Worlds: SEPM, Short Course Notes no. 

35, p. 1–102. 

REGER, D.B., 1926, Mercer, Monroe, and Summer Counties: West Virginia 

Geological Survey, 963 p. 

RICE, C.L., SABLE, E.G., DEVER, G.R., JR., AND KEHN, T.M., 1979, The 

Mississippian and Pennsylvanian (Carboniferous) Systems in the 

United States—Kentucky: U.S. Geological Survey, Professional Paper 

1110-F, 32 p. 

REXROAD, C., AND CLARKE, C.E., 1960, Conodonts form the Glen Dean 

Formation of Kentucky and equivalent formations of Virginia and 

West Virginia: Journal of Paleontology, v. 34, p. 1202–1206. 

ROBERTS, J., CLAOUE-LONG, J., JONES, P.J., AND FOSTER, C.B., 1995, SHRIMP 

zircon age control of Gondwana sequences in Late Carboniferous 

and early Permian Australia, in Non-Biostratigraphical Methods of 

Dating and Correlation: Geological Society of London, Special Publication 

89, p. 145–174. 

ROSS, C.A., AND ROSS, J.R., 1987, Late Paleozoic sea-levels and depositional 

sequences, in Ross, C.A., and Haman, D., eds., Timing and 

Depositional History of Eustatic Sequences: Constraints on Seismic 

Stratigraphy: Cushman Foundation for Foraminiferal Research, 

Special Publication 24, p. 137–149. 


33 

RUDDIMAN, W.F., RAYMO, M., AND MACINTYRE, A., 1986, Matuyuma 41,000year 

cycles: North Atlantic Ocean and Northern Hemisphere icesheets: 

Earth and Planetary Science Letters, v. 80, p. 117–129. 

SABLE, E.G., AND DEVER, G.R., 1990, Mississippian Rocks in Kentucky: U.S. 

Geological Survey, Professional Paper 1503, 125 p. 

SANDO, W.J., 1984, Revised Mississippian time scale, Western Interior 

region, conterminous United States: U.S. Geological Survey, Stratigraphic 

Notes, p. A15-A26. 

SARG, J.F., 1988, Carbonate sequence stratigraphy, in Wilgus, C.K., Hastings, 

B.S., Kendall, C.G.St.C, Posamentier, H.W., Ross, C.A., and Van 

Wagoner, J.C., Sea-Level Changes: An Integrated Approach: SEPM, 

Special Publication 42, p. 155–181. 

SCOTESE, C.R., AND MCKERROW, W.S., 1990, Revised world maps and 

introduction, in McKerrow, W.S., and Scotese, C.R., eds., Palaeozoic 

Palaeogeography and Biogeography: Geological Society of London, 

Memoir 12, p. 1–24. 

SHUMAKER, R.C., AND WILSON, T.H., 1996, Basement structure of the Appalachian 

foreland in West Virginia: Its style and effect on sedimentation, 

in van der Pluijm, B.A., and Catacosinos, P.A., eds., Basement 

and Basins of Eastern North America: Geological Society of America, 

Special Paper 308, p. 139–155. 

SMITH, L.B., 1996, High resolution sequence stratigraphy of late Mississippian 

(Chesterian) mixed carbonates and siliciclastics, Illinois Basin: 

Unpublished Ph.D. Thesis, Virginia Tech, Blacksburg, 146 p. 

SMITH, L.B., AND READ, J.F., 1999, Application of high resolution sequence 

stratigraphy to tidally influenced Late Mississippian carbonates, 

Illinois Basin, in Harris, P.M., Saller, A.H., and Simo, J.A., eds., 

Advances in Carbonate Sequence Stratigraphy—Application to Reservoirs, 

Outcrops, and Models: SEPM, Special Publication 63, p. 107– 

126. 

SMITH, L.B., JR., AND READ, J.F., 2000, Rapid onset of Late Paleozoic glaciation: 

Evidence from Upper Mississippian strata of the midcontinent, 

United States: Geology, v. 28, p. 279–282. 

SMITH, L.B., AND READ, J.F., 2001, Discrimination of local and global effects 

on Upper Mississippian stratigraphy, Illinois Basin, U.S.A.: Journal of 

Sedimentary Research, v. 71, p. 985–1002. 

SMITH, T.M., AND DOROBEK, S.L., 1993, Stable isotopic composition of 

meteoric calcites: evidence of Early Mississippian climate change in 

the Mission Canyon Formation, Montana: Tectonophysics, v. 222, p. 

317–331. 

SMOSNA, R., AND KOEHLER, B., 1993, Tidal origin of a Mississippian oolite on 

the West Virginia Dome, in Keith, B.D., and Zuppman, C.W., eds., 

Mississippian Oolites and Modern Analogs: American Association of 

Petroleum Geologists, Studies in Geology no. 35, p. 149–162. 

STAMM, R., 1997, Late Mississippian conodont biostratigraphy of the 

Appalachian Basin: Preliminary correlations to the Eastern Interior 

Basin and eustatic curves (abstract): The Geological Society of America, 

31 st Annual South-Central and 50 th Rocky Mountain Sections, Abstracts 

with Programs, p. 48. 

SWANN, D.H., 1964, Late Mississippian rhythmic sediments of Mississippi 

Valley: American Association of Petroleum Geologists, Bulletin, v. 48, 

p. 637–658. 

THOMAS, W.A., 1977, Evolution of Appalachian salients and recesses from 

reentrants and promontories in the continental margin: American 

Journal of Science, v. 277, p. 1233–1278. 

VEEVERS, J.J., AND POWELL, C.M., 1987, Late Paleozoic glacial episodes in 

Gondwanaland reflected in transgressive–regressive sequences in 

Euramerica: Geological Society of America, Bulletin, v. 98, p. 475– 

487. 

WEBER, L.J., SARG, J.F., AND WRIGHT, F.M., 1995, Sequence stratigraphy 

and reservoir delineation of the Middle Pennsylvanian 

(Desmoinesian), Paradox basin and Aneth Field, southwestern USA, 

in Read, J.F., Kerans, C., Weber, L.J., Sarg, J.F., and Wright, F.M., 

eds., Milankovitch Sea-Level Changes, Cycles and Reservoirs on

34 

Carbonate Platforms in Greenhouse and Ice-house Worlds: SEPM, 

Short Course 35, Part 3, p. 1–81. 

WELLS, D., 1950, Lower Middle Mississippian of southeastern West Virginia: 

American Association of Petroleum Geologists, Bulletin, v. 34, 

p. 882–992. 

WHITEHEAD, N.H., III, 1984, Paleogeography and depositional environments 

of the Lower Mississippian of the East-Central United States, in 

Nassichuk, W.W., ed., Paleogeography and Paleotectonics: Neuvième 

Congrès International de Stratigraphie et de Géologie du Carbonifère 

1979, v. 3, Part 2, p. 280–290. 


APPENDIX. —Sections used in the study. 

WRIGHT, V.P., AND VANSTONE, S.D., 2001, Onset of Late Paleozoic glacioeustasy 

and the evolving climates of low latitude areas: A synthesis of 

current understanding: Geological Society of London, Journal, v. 158, 

p. 579–582. 

YEILDING, C.A., AND DENNISON, J.M., 1986, Sedimentary response to Mississippian 

tectonic activity at the east end of the 38th Parallel fracture 

zone: Geology, v. 14, p. 621–624. 

Section # Name County, State Latitude Longitude 

1 Canaan Valley Tucker County, West Virginia 39.002800 -79.467500 

2 Randolph 28 Randolph County, West Virginia 38.716940 -80.132780 

3 Mingo Randolph County, West Virginia 38.488200 -80.041700 

4 Greenbrier 21 Greenbrier County, West Virginia 37.965000 -80.627220 

5 Summers 7 Summers County, West Virginia 37.794170 -80.744170 

6 Sunoil Test Core #2 Pocahontas County, West Virginia 38.108333 -80.266667 

7 Greenbrier 15 Greenbrier County, West Virginia 37.751390 -80.586670 

8 Union Monroe County, West Virginia 37.610800 -80.587500 

9 Mercer 25 Mercer County, West Virginia 37.440000 -81.018610 

10 Ingleside Mercer County, West Virginia 37.313900 -81.141700 

11 Bristol Washington County, Virginia 36.75 -82.05 

12 Wayne 277 Wayne County, West Virginia 38.271111 -82.538611 

13 Wayne 1497 Wayne County, West Virginia 38.1975 -82.334167 

14 Lincoln 155 Lincoln County, West Virginia 38.306111 -82.202222 

15 Lincoln 1375 Lincoln County, West Virginia 38.163060 -82.130280 

16 Mingo 796 Mingo County, West Virginia 37.924170 -82.206670 

17 Logan 102 Logan County, West Virginia 37.96861 -81.973056 



20 Wyoming 253 Wyoming County, West Virginia 37.77167 -81.51806 

21 Wyoming 87 Wyoming County, West Virginia 37.525556 -81.637500 

22 McDowell 63 McDowell County, West Virginia 37.357500 -81.653611 

23 McDowell 792 McDowell County, West Virginia 37.44306 -81.400833 



26 Mercer 27 Mercer County, West Virginia 37.39722 -81.0525

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