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PERMO-CARBONIFEROUS CARBONATE<br />
PLATFORMS AND REEFS<br />
Volc<br />
W T-16 T-463 T-5246 T-5056 T-10 E<br />
500 m<br />
V.E. = 10<br />
5000 m<br />
Volc<br />
Unit 1<br />
Unit 2<br />
Unit 3<br />
Top Lvis<br />
Tul<br />
Bob<br />
Rad<br />
TSS<br />
HSS<br />
TSS HSS TSS HSS TSS<br />
Bash_SSB<br />
Serp_SSB<br />
Lvis13_csb<br />
Lvis2_MFS<br />
Lvis_SSB<br />
Lvis1_MFS<br />
Edited by:<br />
WAYNE M. AHR<br />
Department of Geology <strong>and</strong> Geophysics, Texas A&M University, College Station, Texas 77843-3115, U.S.A.<br />
PAUL M. (MITCH) HARRIS<br />
ChevronTexaco E&P Technology Company, 6001 Bollinger Canyon Road,<br />
San Ramon, California 94583-0746, U.S.A.<br />
WILLIAM A. MORGAN<br />
ConocoPhillips, Inc., P.O. Box 2197, Houston, Texas 77252-2197, U.S.A.<br />
AND<br />
IAN D. SOMERVILLE<br />
Department of Geology, University College - Dublin, Belfield, Dublin 4, Irel<strong>and</strong><br />
Top Evis<br />
Evis_SSB<br />
Tour_MFS<br />
Fame_SSB<br />
Bash = Bashkirian<br />
Serp = Serpukhovian<br />
Tul = Tulsky<br />
Bob = Bobrikovsky<br />
Tour = Tournaisian<br />
Fame = Famennian (Devonian) Devonian<br />
Lvis = Late Visean Rad = Radaevsky SSB = Supersequence Boundary<br />
“Volc” = Thick Ash Interval Evis - Early Visean MFS = Maximum Flooding Surface<br />
TSS = Transgressive Sequence Set HSS = Highst<strong>and</strong> Sequence Set<br />
SEPM Special Publication 78<br />
AAPG Memoir 83<br />
Platform (Shallow & Deeper)<br />
Upper-Middle Slope &<br />
Deeper Platform<br />
Allochthonous Debris<br />
Lower Slope -<br />
Toe of Slope<br />
“Volc” Interval<br />
Basin<br />
Published jointly by<br />
SEPM (Society for Sedimentary Geology) <strong>and</strong><br />
The American Association of Petroleum Geologists (AAPG)<br />
Tulsa, Oklahoma, U.S.A.<br />
Printed in the United States of America<br />
December, 2003<br />
Volcanic
SEPM, AAPG, <strong>and</strong> the authors are grateful to the following<br />
for their generous contribution to the cost of publishing<br />
Permo-Carboniferous Carbonate Platforms <strong>and</strong> Reefs<br />
ExxonMobil Development Company<br />
Kansas Geological Survey<br />
Saudi Aramco<br />
University of Kansas, Department of Geology, Haas Fund<br />
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ISBN 1-56576-087-5<br />
Copyright © 2003 by<br />
SEPM (Society for Sedimentary Geology) <strong>and</strong> The American Association of Petroleum Geologists (AAPG)<br />
All Rights Reserved<br />
Printed in the United States of America<br />
Published December, 2003
MISSISSIPPIAN RAMP, WEST VIRGINIA, U.S.A.<br />
SEQUENCE RESPONSE OF A DISTAL-TO-PROXIMAL FORELAND RAMP TO<br />
GLACIO-EUSTASY AND TECTONICS: MISSISSIPPIAN, APPALACHIAN BASIN,<br />
WEST VIRGINIA–VIRGINIA, U.S.A.<br />
AUS AL-TAWIL<br />
Saudi Aramco, Dhahran, Saudi Arabia<br />
AND<br />
THOMAS C. WYNN, AND J. FRED READ<br />
Department of Geological Sciences, Virginia Tech, Blacksburg Virginia 24061, U.S.A.<br />
ABSTRACT: This paper evaluates the limits of using well-cuttings data <strong>and</strong> wireline logs in conjunction with limited core <strong>and</strong> outcrop data<br />
to generate a regional, high-resolution sequence stratigraphy for upper Mississippian (Chesterian) Greenbrier <strong>carbonate</strong>s, West Virginia,<br />
U.S.A. These data are then used to document the stratigraphic response of the distal to proximal forel<strong>and</strong> basin to tectonics <strong>and</strong><br />
Carboniferous glacio-eustasy during the transition into ice-house times. The major mappable sequences are fourth-order sequences, a few<br />
meters to over a hundred meters thick. They consist of updip red beds <strong>and</strong> eolianites, lagoonal muddy <strong>carbonate</strong>s, ooid grainstone <strong>and</strong><br />
skeletal grainstone–packstone shoal complexes, open-ramp skeletal wackestone, <strong>and</strong> slope–basinal laminated argillaceous <strong>carbonate</strong>s. The<br />
sequences are bounded downdip by lowst<strong>and</strong> s<strong>and</strong>stones <strong>and</strong> calcareous siltstones, <strong>and</strong> locally on the ramp by basal transgressive shales;<br />
only a few sequence boundaries are calichified, compared with updip sections in Kentucky, where caliche <strong>and</strong> breccias are common.<br />
Transgressive systems tracts range from thin units to others that constitute the lower half of the sequence. The highst<strong>and</strong> systems tracts<br />
contain significant grainstone units. Maximum flooding surfaces on the ramp slope occur at the base of slope or basinal facies that rest on<br />
lowst<strong>and</strong> to transgressive complexes, whereas on the ramp they occur beneath widespread grainstones that overlie nearshore shale or lime<br />
mudstone. In the Greenbrier succession, fourth-order sequences are arranged into weak third-order composite sequences bounded updip<br />
by red beds, <strong>and</strong> by lowst<strong>and</strong> s<strong>and</strong>s <strong>and</strong> oolite along the ramp slope. The composite sequences contain three to four fourth-order sequences.<br />
Correlation of the forel<strong>and</strong> basin units with third-order global sea-level curves, <strong>and</strong> with high-frequency sequences within the<br />
intracratonic Illinois Basin, shows that in spite of differential subsidence rates ranging from 2 to 30 cm/ky across the forel<strong>and</strong>, global third<strong>and</strong><br />
fourth-order sea-level changes whose amplitude increased with time were a strong influence on sequence development. Thrust-loadinduced<br />
differential subsidence of fault blocks of the forel<strong>and</strong> basement controlled the rapid basinward thickening of the depositional<br />
wedge, <strong>and</strong> modified the eustatic effects on the accumulating succession. Moderate-amplitude eustatic sea-level change <strong>and</strong> semiarid<br />
climate were the dominant causes of the widespread reservoirs of ooid grainstone <strong>and</strong> lowst<strong>and</strong> s<strong>and</strong>s. The overall stratigraphy suggests<br />
upward increase in amplitudes of sea-level change, <strong>and</strong> cooling, which likely records the initiation of Gondwana ice-sheet growth.<br />
INTRODUCTION<br />
This paper mainly uses well data (especially well cuttings <strong>and</strong><br />
wireline logs), along with limited outcrop data <strong>and</strong> core, to<br />
determine if they can be used to generate a high-resolution threedimensional,<br />
sequence stratigraphic framework for the Mississippian<br />
(Chesterian) <strong>carbonate</strong>s, from 0 up to 200 m thick, in the<br />
eastern, Appalachian forel<strong>and</strong> basin (Colton, 1970; de Witt <strong>and</strong><br />
McGrew, 1979) (Figs. 1–4). The study differs from Al-Tawil <strong>and</strong><br />
Read (this volume), which was largely an outcrop-based study<br />
along two outcrop belts in the distal forel<strong>and</strong>, Kentucky. Its<br />
forel<strong>and</strong>-basin setting <strong>and</strong> presence of the ramp margin in the<br />
study area distinguish it from the studies of Smith <strong>and</strong> Read<br />
(1999, 2000, 2001) in the Illinois Basin, an intracratonic basin to the<br />
west in which the ramp margin is not present. The units in the<br />
study area show extremely complex facies heterogeneity, with<br />
rapid vertical <strong>and</strong> lateral facies changes that have hampered<br />
development of regional exploration models, <strong>and</strong> recognition of<br />
shoreline <strong>and</strong> ramp-margin trends.<br />
The paper provides the first detailed, high-resolution, regional<br />
sequence stratigraphic study of the mixed <strong>carbonate</strong> <strong>and</strong><br />
siliciclastic facies extending from the slowly subsiding distal<br />
forel<strong>and</strong> to the more rapidly subsiding proximal forel<strong>and</strong> of the<br />
Appalachian Basin. Previously, the succession had been subdivided<br />
into several lithologically heterogeneous formations (Fig.<br />
4), with much of the correlation being based on a few semiregional<br />
siliciclastic marker beds (Reger, 1926; Wells, 1950;<br />
McFarlan <strong>and</strong> Walker, 1956; Rice et al., 1979). We subdivided the<br />
Chesterian succession into regionally traceable fourth-order high-<br />
Permo-Carboniferous Carbonate Platforms <strong>and</strong> Reefs<br />
SEPM Special Publication No. 78 <strong>and</strong> AAPG Memoir 83, Copyright © 2003<br />
SEPM (Society for Sedimentary Geology), American Association of Petroleum Geologists (AAPG), ISBN 1-56576-087-5, p. 11–34.<br />
11<br />
frequency depositional sequences, by tracing major l<strong>and</strong>ward<br />
<strong>and</strong> basinward shifts in facies, <strong>and</strong> mapping outcropping sequence-bounding<br />
disconformities, correlative conformities, <strong>and</strong><br />
sequence-bounding siliciclastic units. This largely subsurface<br />
study illustrates how high-resolution sequence stratigraphic<br />
analysis can increase our underst<strong>and</strong>ing of facies stacking on the<br />
ramp (Ahr, 1973) <strong>and</strong> its response to forel<strong>and</strong> basin tectonics,<br />
glacio-eustatic sea-level changes related to onset of Gondwana<br />
glaciation (Crowell, 1978), <strong>and</strong> long-term climate change<br />
(Ettensohn et al., 1988).<br />
The study documents how, even in this tectonically active<br />
forel<strong>and</strong>-basin setting, the major role of tectonics has been one of<br />
generating differential accommodation across <strong>and</strong> along the<br />
forel<strong>and</strong> basin, <strong>and</strong> in localizing the position of structural highs,<br />
the ramp’s hinge line, <strong>and</strong> ramp-margin facies. In spite of the<br />
active tectonic regime, fourth-order <strong>and</strong>, to a lesser extent, thirdorder<br />
eustasy has been the overriding control on the generation<br />
of the regionally correlative sequences.<br />
GEOLOGIC SETTING<br />
Regional Structural Setting<br />
The Appalachian Basin is an elongate forel<strong>and</strong> basin bordered<br />
on the east by overthrusted Precambrian <strong>and</strong> early Paleozoic rocks<br />
<strong>and</strong> on the west by the Cincinnati Arch (Colton, 1970; de Witt <strong>and</strong><br />
McGrew, 1979) (Figs. 1, 2). In the west, the Paleozoic rocks are<br />
undeformed to gently folded, whereas the eastern part of the<br />
basin is in the folded <strong>and</strong> thrusted Valley-<strong>and</strong>-Ridge Province.
12<br />
The Upper Mississippian <strong>carbonate</strong> strata in the Appalachian<br />
Basin (Arkle et al., 1979) are exposed in narrow, northeast–<br />
southwest trending outcrop belts, one along the distal forel<strong>and</strong> in<br />
Kentucky adjacent to the Cincinnati Arch (described by Al-Tawil<br />
<strong>and</strong> Read, this volume) <strong>and</strong> the others (described here) in the<br />
more proximal forel<strong>and</strong>, exposed along the eastern Appalachian<br />
Plateau <strong>and</strong> into the fold–thrust belt (Figs. 1, 2). Between these<br />
outcrop belts, the Mississippian rocks of the distal to proximal<br />
forel<strong>and</strong> basin lie in the subsurface beneath Pennsylvanian sediments<br />
<strong>and</strong> are penetrated by numerous wells, twenty of which<br />
were used along with a core <strong>and</strong> four outcrop sections for this<br />
study.<br />
The Appalachian forel<strong>and</strong> basin was tectonically active during<br />
the deposition of these Upper Mississippian <strong>carbonate</strong>s, <strong>and</strong><br />
the Mississippian rocks thicken into the foredeep in southwestern<br />
Virginia <strong>and</strong> Tennessee (Fig. 1). The depocenter of the Late<br />
Mississippian <strong>carbonate</strong> ramp is located in the overthrust belt in<br />
the Greendale Syncline, southwestern Virginia, where Mississippian<br />
units thicken to over 2000 meters of basinal, slope, <strong>and</strong><br />
ramp-margin facies (Cooper, 1948; Bartlett <strong>and</strong> Webb, 1971).<br />
Several synsedimentary positive structures occur in West Virginia<br />
<strong>and</strong> eastern Kentucky (Yeilding <strong>and</strong> Dennison, 1986; Dever<br />
et al., 1990; Sable <strong>and</strong> Dever, 1990; Dever, 1995) (Fig. 1). The<br />
Cincinnati Arch, which borders the western edge of the basin,<br />
subsided very little during Chesterian time, while adjacent<br />
areas to the west (Illinois Basin) <strong>and</strong> the east (Appalachian<br />
AUS AL-TAWIL, THOMAS C. WYNN, AND J. FRED READ<br />
FIG. 1.—Geologic location map of study area, showing distribution of exposed Mississippian Greenbrier <strong>carbonate</strong> rocks (gray<br />
shading) in the Appalachian Basin, eastern U.S.A. Isopach contours (in meters) show thickness of Greenbrier Group, which<br />
thickens into the Appalachian foredeep to the southeast (modified from Flowers, 1956; Pryor <strong>and</strong> Sable, 1974; MacQuown <strong>and</strong><br />
Pear, 1980; Yeilding <strong>and</strong> Dennison, 1986; Dever, 1986, 1995; Sable <strong>and</strong> Dever, 1990, Dever et al., 1990). Major structures within the<br />
study area are the ancestral Rome Trough border faults, <strong>and</strong> the Jessamine Dome <strong>and</strong> Cincinnati <strong>and</strong> Waverly arches, which<br />
bound the area to the northwest, <strong>and</strong> the Burning–Mann anticline <strong>and</strong> Virginia Promontory, which cut across the trend of the<br />
basin. Orogenic highl<strong>and</strong>s bounded the basin to the southeast.<br />
Basin) subsided more rapidly (Pryor <strong>and</strong> Sable, 1974; deWitt<br />
<strong>and</strong> McGrew, 1979; Whitehead, 1984; Sable <strong>and</strong> Dever, 1990;<br />
Dever, 1995). The West Virginia Dome was a positive structure<br />
along the 38th parallel lineament in northeastern West Virginia,<br />
which was active during Early Chesterian time (Yeilding <strong>and</strong><br />
Dennison, 1986).<br />
The basement in the study area is segmented by major faults<br />
into a series of blocks, some of which moved independently<br />
during Mississippian deposition (Shumaker <strong>and</strong> Wilson, 1996).<br />
These faults include northeast–southwest trending bounding<br />
faults of the Rome Trough, as well as several zones parallel to <strong>and</strong><br />
to the east of these. The northeast–southwest trending basinal<br />
hinge occurs just to the west of the overthrust belt, <strong>and</strong> a smaller<br />
east–west trending hinge occurs in the northeast along the 400foot<br />
(120 m) Greenbrier isopach (Fig. 1). A northwest–southeast<br />
trending arch lies at right angles to the regional structure, in the<br />
vicinity of the “Virginia Promontory”, a long-lived Paleozoic<br />
high (Thomas, 1977).<br />
Stratigraphy <strong>and</strong> Biostratigraphy<br />
The Mississippian succession makes up a supersequence set<br />
(roughly 35 My duration) composed of at least three<br />
supersequences. The lower supersequence set boundary is the<br />
Devonian–Mississippian erosional unconformity (Bjerstadt <strong>and</strong><br />
Kammer, 1988) (Figs. 3, 4). The lower supersequence consists of
WAVERLY<br />
ARCH<br />
ROME ROME TROUGH TROUGH<br />
APPALACHIAN<br />
BASIN<br />
OHIO<br />
the up<strong>permo</strong>st Devonian–lowermost Mississippian Cloyd Conglomerate–Berea<br />
S<strong>and</strong>stone <strong>and</strong> the overlying Subury Shale–<br />
Price/Borden/Maccrady deltaic complex (Kinderhookian to<br />
mid-Osagean), which contains several clinoformed depositional<br />
sequences (Branson, 1912; Butts, 1940, 1941; Bjerstadt <strong>and</strong><br />
Kammer, 1988; Bartlett, 1974). This is overlain southwest of the<br />
study area by the <strong>carbonate</strong>-prone Fort Payne–Salem<br />
supersequence (upper Osagean–lower Meramecian) (Fig. 4;<br />
Bjerstadt <strong>and</strong> Kammer, 1988; Sable <strong>and</strong> Dever, 1990; Khetani<br />
<strong>and</strong> Read, 2002).<br />
The upper Meramecian–Chesterian supersequence (Fig. 4)<br />
described in this paper overlies the lower two supersequences.<br />
The upper Meramecian Little Valley–Hillsdale sequence may<br />
have formed during lowered sea levels relative to the overlying<br />
units, given that it is the least extensive unit on the platform, <strong>and</strong><br />
it contains relatively shallow oncolitic limestone facies even in the<br />
most basinward sections. The Chesterian sequences C1 to 11<br />
described in this paper make up the <strong>carbonate</strong>-prone, largely<br />
transgressive part of the supersequence. The shale-prone sequence<br />
C11 may be the initial sequence of the supersequence<br />
highst<strong>and</strong>, the bulk of the supersequence highst<strong>and</strong> being made<br />
up of the progradational Mauch Chunk red siliciclastics. The<br />
early Pennsylvanian unconformity marks the top boundary of<br />
the supersequence set.<br />
Regional Mississippian stratigraphy <strong>and</strong> biostratigraphy of<br />
the Mississippian in Virginia <strong>and</strong> West Virginia (Figs. 3, 4) are<br />
given in Butts (1922, 1940, 1941), Reger (1926), Wells (1950),<br />
MISSISSIPPIAN RAMP, WEST VIRGINIA, U.S.A.<br />
40° N 83° W 40° N 80° W<br />
11<br />
BURNING-MANN<br />
ANTICLINE<br />
FOLD AND THRUST BELT<br />
WEST<br />
VIRGINIA<br />
12<br />
14<br />
13<br />
15<br />
17<br />
16<br />
18 20<br />
5<br />
19<br />
21<br />
22<br />
24<br />
9<br />
23<br />
25 26<br />
10<br />
VIRGINIA<br />
PROMONTORY<br />
WEST VIRGINIA<br />
DOME<br />
FIG. 2.—Location map showing outcrop sections <strong>and</strong> well sections used to construct the cross sections in the study. Location<br />
information on sections is in the Appendix. Outcrops of Mississippian Greenbrier Limestone is shaded. The northeast–southwest<br />
cross section, based on sections 1 to 11 (Fig. 6), includes data from the eastern outcrop belt, supplemented by nearby shallow wells<br />
<strong>and</strong> a core; the sections thicken southwestward down the basin axis. The northwest–southeast cross section extends across the<br />
forel<strong>and</strong> normal to the trend of the basin <strong>and</strong> is based on wells 12 to 26 (Fig. 7); it uses only subsurface well cuttings <strong>and</strong> wireline<br />
logs. The succession thickens markedly across the northeast–southwest trending hinge (just southeast of the 100 m isopach of<br />
Figure 1) that separates the distal forel<strong>and</strong> from the proximal forel<strong>and</strong>. Note that to the northeast this hinge curves to the east.<br />
Two minor east–west trending hinges occur to the southwest of the main hinge. Outcrops representative of the distal forel<strong>and</strong><br />
of Kentucky (along strike to the southwest) are described in Al-Tawil <strong>and</strong> Read (this volume).<br />
4<br />
7<br />
8<br />
DISTAL FORELAND<br />
6<br />
2<br />
3<br />
1<br />
37° N 80° W<br />
VIRGINIA<br />
PROXIMAL FORELAND<br />
0<br />
100 km<br />
100 mi<br />
MISSISSIPPIAN<br />
OROGENIC<br />
HIGHLANDS<br />
13<br />
Flowers (1956), de Witt <strong>and</strong> McGrew (1979), Rice et al. (1979), <strong>and</strong><br />
Maples <strong>and</strong> Waters (1987). The Greenbrier Group (Hillsdale to<br />
Alderson interval <strong>and</strong> the Lillydale Shale <strong>and</strong> Glenray Limestone<br />
of the Lower Bluefield Formation, Mauch Chunk Group) make<br />
up the study interval (0 to 1000 m thick). It unconformably<br />
overlies the Price–Maccrady deltaic–marine shelf deposits (Figs.<br />
3, 4). The basal unit studied is the upper Meramecian Little<br />
Valley–Hillsdale Limestone, which is a very cherty, pelletal–<br />
peloidal <strong>and</strong> skeletal packstone–wackestone with the distinctive<br />
corals Syringopora <strong>and</strong> Acrocyathus proliferus (Lithostrotionella)<br />
(Wells, 1950; DiRenzo, 1986). This is overlain by the Chesterian<br />
Denmar Formation, which in outcrop contains numerous exposure<br />
surfaces, consists of ooid grainstone, lime wackestone–<br />
mudstone, <strong>and</strong> minor quartz–peloidal grainstone, <strong>and</strong> has the<br />
crinoid Platycrinites penicillus <strong>and</strong> the earliest occurrence of the<br />
zonal conodonts G. bilineatus <strong>and</strong> C. charactus (Butts, 1922, 1940,<br />
1941; Wells, 1950; McFarlan <strong>and</strong> Walker, 1956; Dever et al., 1990).<br />
The Denmar Formation is overlain by the Taggard red beds, a<br />
regional marker consisting of red shale <strong>and</strong> siltstone <strong>and</strong> minor<br />
<strong>carbonate</strong>s.<br />
The overlying ooid- <strong>and</strong> skeletal limestone–rich units <strong>and</strong><br />
subordinate <strong>carbonate</strong> mudstone constitute the Gasper Formation,<br />
which contains the first appearance of the crinoid index<br />
fossil Talarocrinus (Butts, 1940, 1941; Wells, 1950). The interval is<br />
subdivided into the Pickaway, Union, Greenville Shale, <strong>and</strong><br />
Alderson Formations (Reger, 1926). This is overlain by Lillydale<br />
Shale <strong>and</strong> the Glenray Limestone (Reger, 1926), consisting of
14<br />
PERIOD<br />
MISSISSIPPIAN<br />
DEV.<br />
EPOCH<br />
KIND. OSAGEAN MERAM. CHESTERIAN<br />
AGE<br />
327<br />
354<br />
AUS AL-TAWIL, THOMAS C. WYNN, AND J. FRED READ<br />
CONODONT<br />
ZONES<br />
G. bilineatus -<br />
K. mehli<br />
G. bilineatus -<br />
C. altus<br />
G. bilineatus -<br />
C. charactus<br />
A. scalenus -<br />
Cavusgnathus<br />
Cavusgnathus -<br />
T. varians<br />
OTHER<br />
INDEX<br />
FOSSILS<br />
Talarocrinus<br />
spp.<br />
Platycrinites<br />
penicillus<br />
Acrocyathus spp.<br />
'Lithostrotion spp.'<br />
MAUCH CHUNK GP.<br />
G R E E N B R I A R G R O U P ("B I G L I M E")<br />
WEST VIRGINIA<br />
DROOP SS.<br />
REYNOLDS<br />
LIMESTONE<br />
GLENRAY<br />
LIMESTONE<br />
LILLYDALE<br />
SHALE<br />
ALDERSON<br />
LIMESTONE<br />
UNION<br />
LIMESTONE<br />
PICKAWAY<br />
LIMESTONE<br />
UPPER / LOWER<br />
TAGGARD<br />
DENMAR<br />
LIMESTONE<br />
VIRGINIA<br />
COVE<br />
CREEK<br />
LIMESTONE<br />
FIDO SS.<br />
HILLSDALE LST.<br />
LITTLE VALLEY LST.<br />
MACCRADY<br />
FORMATION<br />
PRICE<br />
FORMATION<br />
PENNINGTON<br />
FM.<br />
GASPER<br />
LIMESTONE<br />
STE.<br />
GENEVIEVE<br />
LIMESTONE<br />
LATE DEVONIAN AND<br />
EARLY MISSISSIPPIAN<br />
BLACK SHALE<br />
SEQ.<br />
THIS<br />
PAPER<br />
FIG. 3.—Chronostratigraphic chart, Mississippian interval, showing the major units in the study area of West Virginia <strong>and</strong> Virginia;<br />
studied interval is from the Little Valley–Hillsdale to the Glenray Formation, <strong>and</strong> includes the drillers’ “Big Lime” (Greenbrier<br />
Group) <strong>and</strong> the “Little Lime”. Chesterian sequences described in this paper are numbered C1 to C11; underlying Meramecian<br />
sequences are not numbered. The study interval is equivalent to the St. Louis to Glen Dean interval of Kentucky (Appalachian<br />
Basin) <strong>and</strong> the Illinois Basin. Modified from de Witt <strong>and</strong> McGrew (1979), Roberts et al (1995), Collinson et al (1971), Rexroad <strong>and</strong><br />
Clarke (1960), Huggins (1983), Stamm (1997), Barlett <strong>and</strong> Webb (1971), Reger (1926), <strong>and</strong> Wells (1950).<br />
ooid <strong>and</strong> skeletal limestone <strong>and</strong> containing the zonal conodonts<br />
G. bilineatus <strong>and</strong> Kladognathus (Rexroad <strong>and</strong> Clarke, 1960). The<br />
succession is capped by Mauch Chunk Group red s<strong>and</strong>stone,<br />
siltstone, <strong>and</strong> mudrock, whose top is the Mississippian–Pennsylvanian<br />
unconformity (Brezinski, 1989). Regionally, the units<br />
thicken to over 600 m to the southeast into the foredeep, where<br />
they are dominated by thick units of dark skeletal wackestone–<br />
mudstone <strong>and</strong> basin–slope laminated argillaceous <strong>carbonate</strong><br />
mudstones, <strong>and</strong> thin quartz s<strong>and</strong>stone <strong>and</strong> skeletal–ooid<br />
grainstone. In the subsurface, the Hillsdale to Alderson <strong>carbonate</strong>s<br />
are the drillers’ “Big Lime” <strong>and</strong> the Glenray Limestone is<br />
the “Little Lime.”<br />
BLUEFIELD FM.<br />
C11<br />
C10<br />
C9<br />
C8<br />
C7<br />
C6<br />
C5<br />
C4<br />
C3<br />
C2<br />
C1<br />
FACIES AND REGIONAL DISTRIBUTION<br />
The facies developed in the eastern Appalachian Basin<br />
(Leonard, 1968; Carney <strong>and</strong> Smosna, 1989) are shown schematically<br />
on an idealized ramp model (Fig. 5) <strong>and</strong> resemble facies<br />
described elsewhere by Ettensohn et al. (1984) <strong>and</strong> Smith <strong>and</strong><br />
Read (1999, 2001). The regional facies distributions are shown on<br />
the two regional cross sections (Figs. 6A, 7A). The first cross<br />
section extends southwestward down the axis of the forel<strong>and</strong><br />
basin, <strong>and</strong> is based on five detailed outcrop sections, a core, <strong>and</strong><br />
data from four wells. The second cross section is a dip section<br />
extending across the distal forel<strong>and</strong> onto the proximal forel<strong>and</strong>,
PERIOD<br />
MISSISSIPPIAN<br />
DEV.<br />
EPOCH<br />
KIND. OSAGEAN MERAM. CHESTERIAN<br />
AGE<br />
327<br />
354<br />
DISTAL PROXIMAL<br />
FORELAND<br />
MAUCH CHUNK<br />
SILICICLASTICS<br />
REYNOLDS LST.<br />
GLENRAY LST.<br />
LILLYDALE<br />
SHALE<br />
U. TAGGARD<br />
L. TAGGARD<br />
ALDERSON LST.<br />
UNION LST.<br />
PICKAWAY LST.<br />
R R R<br />
R R<br />
DENMAR LST.<br />
PRICE FM.<br />
SILICICLASTICS<br />
MISSISSIPPIAN RAMP, WEST VIRGINIA, U.S.A.<br />
COVE CREEK LST.<br />
HILLSDALE LST.<br />
L. VALLEY LST.<br />
MACCRADY<br />
FM.<br />
LATE DEVONIAN AND EARLY<br />
MISSISSIPPIAN BLACK SHALE<br />
FIDO SS.<br />
FIG. 4.—Schematic lithostratigraphy showing the major Mississippian<br />
stratigraphic units in the area. The study interval<br />
(Little Valley–Hillsdale to the Glenray/Cove Creek Limestone),<br />
overlies Early Mississippian siliciclastics of the Price–<br />
Borden formations, <strong>and</strong> is overlain by <strong>and</strong> interfingers with<br />
the upper Mississippian Mauch Chunk siliciclastics.<br />
PALEOSOLS<br />
QUARTZ-PELOID<br />
GRAINSTONE / PACKSTONE<br />
RED BEDS<br />
SEQUENCE BOUNDARY<br />
PELOIDAL<br />
GRAINSTONE / PACKSTONE<br />
LIME WACKESTONE / MUDSTONE<br />
~ 200 km<br />
OOID<br />
GRAINSTONE / PACKSTONE<br />
GRAY SHALE/SILTSTONE<br />
15<br />
<strong>and</strong> consists solely of data from 15 wells (cuttings with 10 to 15<br />
foot sample intervals <strong>and</strong> wire-line logs). Figures 6B <strong>and</strong> 7B show<br />
the individual sequences from each cross section rehung using<br />
the top of each sequence as a datum, <strong>and</strong> using inferred water<br />
depths of the topmost facies to hang the ramp-margin sections.<br />
Facies are summarized in Tables 1 <strong>and</strong> 2.<br />
Paleosols <strong>and</strong> Erosion Surfaces<br />
Carbonate paleosols occur on disconformities on the distal<br />
forel<strong>and</strong>, southwest of the study area in Kentucky (Dever, 1973;<br />
Ettensohn, 1975; Harrison <strong>and</strong> Steinen, 1978; Al-Tawil, 1998) <strong>and</strong><br />
West Virginia, but they are less common in the thicker West<br />
Virginia–Virginia units (Figs. 6, 7). The paleosols include laminar<br />
caliches <strong>and</strong> caliche breccia (Table 1). They formed under seasonally<br />
wet–dry semiarid conditions (Ettensohn et al., 1988). However,<br />
many contacts between sequences <strong>and</strong> parasequences in<br />
outcrop on the proximal forel<strong>and</strong> in West Virginia are sharp<br />
planar erosional to scalloped surfaces. They resemble microkarst<br />
or tidal solution basins.<br />
Red Beds (Marginal Marine)<br />
These are red, commonly structureless, unfossiliferous, fine<br />
siliciclastic mudrocks <strong>and</strong> laminated shale (Figs. 6, 7; Table 1).<br />
Updip in West Virginia, ripple lamination, wave lamination, <strong>and</strong><br />
mud cracks are common, but burrows are rare. Some red beds<br />
have rooted horizons, <strong>and</strong> some layers are burrowed. The red<br />
beds are common in updip areas, where they are associated with<br />
eolianites, <strong>and</strong> restricted lagoonal <strong>and</strong> shoal-water grainstones.<br />
These updip facies are marginal marine to intertidal–supratidal<br />
red beds, which formed as highst<strong>and</strong>, basinward-shifting facies.<br />
Gray Shale, Quartz S<strong>and</strong>stone, <strong>and</strong> Calcareous Siltstone<br />
(Tidally Influenced Open Siliciclastic Shelf)<br />
These facies include dark gray, poorly fissile, limy shales that<br />
are barren to variably fossiliferous, with small productid brachiopods,<br />
pelecypods, <strong>and</strong> bryozoa (Table 1). Some overlie pyritized<br />
bedding planes with abundant lacy bryozoans, especially at the<br />
tops of sequences C10 <strong>and</strong> C11 (Alderson <strong>and</strong> Glenray Formations).<br />
The shales form units a few centimeters thick that pinch<br />
out laterally, or are extensive sheets up to 30 m thick (e.g., in<br />
SKELETAL<br />
GRAINSTONE / PACKSTONE<br />
CALCAREOUS SILTSTONE<br />
SANDSTONE<br />
ARGILLACEOUS SKELETAL<br />
WACKESTONE / PACKSTONE<br />
LAMINATED SHALY<br />
LIME MUDSTONE /CALCAREOUS SILTSTONE<br />
FIG. 5.—Schematic facies profile for the Mississippian of the Appalachian Basin. Actual facies distributions are more complex, in that<br />
skeletal grainstone–packstone <strong>and</strong> ooid grainstone facies are developed not only on the ramp margin but also in local areas far<br />
back in the ramp interior.<br />
0 m<br />
20 m<br />
60 m
16<br />
150 MILES<br />
1 2 3 4 5 6 7 8 9 10 11<br />
CANAAN RANDOLPH<br />
GREENBRIER SUMMERS SUN OIL GREENBRIER<br />
MERCER<br />
VALLEY<br />
28 MINGO<br />
21<br />
7<br />
CORE<br />
15 UNION<br />
25 INGLESIDE HOLSTON<br />
FORMATION<br />
SEQ.<br />
GLENRAY LST.<br />
GREENVILLE<br />
SHALE<br />
BLUEFIELD<br />
FM.<br />
COVE CREEK<br />
LIMESTONE<br />
C11<br />
LILLYDALE SHALE<br />
DATUM<br />
FIDO SS.<br />
ALDERSON<br />
FM.<br />
C10<br />
C9<br />
AUS AL-TAWIL, THOMAS C. WYNN, AND J. FRED READ<br />
UNION<br />
FM.<br />
C8<br />
RED-BEDS<br />
PICKAWAY<br />
FM.<br />
GRAY SHALE / SILTSTONE<br />
C7<br />
CALCAREOUS SILTSTONE<br />
C6<br />
QUARTZ SANDSTONE<br />
TAGGARD<br />
FM.<br />
C5<br />
PRICE / MACCRADY FM.<br />
QUARTZ PELOID GRAINSTONE<br />
FINE-GRAINED PELLETAL PACKSTONE,<br />
LIME WACKESTONE / MUDSTONE<br />
OOID GRAINSTONE / PACKSTONE<br />
C4<br />
PELOID GRAINSTONE / PACKSTONE<br />
SKELETAL GRAINSTONE / PACKSTONE, MINOR WACKESTONE<br />
DENMAR<br />
FM.<br />
C3<br />
ARGILLACEOUS SKELETAL PACKSTONE / WACKESTONE<br />
200'<br />
LAMINATED SHALY LIME MUDSTONE/<br />
CALCAREOUS SILTSTONE<br />
SEQUENCE BOUNDARY<br />
C2<br />
PALEOSOL<br />
C1<br />
HILLSDALE<br />
FM.<br />
FIG. 6.—A) High-resolution sequence stratigraphic cross section along the eastern outcrop belt (parallel to the axis of the forel<strong>and</strong> basin), incorporating outcrop sections<br />
<strong>and</strong> nearby shallow wells, which were analyzed using cuttings <strong>and</strong> wireline logs. The cross section extends from West Virginia into Virginia, <strong>and</strong> was constructed<br />
by hanging sections from the regional Lillydale Shale marker; it shows the interpreted regional distribution of lithofacies, sequence boundaries, <strong>and</strong> their relation to<br />
the previous formational stratigraphy (modified from Al-Tawil, 1998). Legend is used for Figure 6B <strong>and</strong> 7B). Detailed information on well <strong>and</strong> outcrop locations is<br />
given in the Appendix.
MISSISSIPPIAN RAMP, WEST VIRGINIA, U.S.A.<br />
1 2 3 4 5 6 7 8 9 11<br />
1 2 3 4 5 6 7 8 9 10 11<br />
100'<br />
FEET<br />
0 MILES 10<br />
100'<br />
0 MILES 10<br />
HILLSDALE<br />
FIG. 6.—B) Facies distributions within individual sequences in Figure 6A (axial cross section), constructed by rehanging each<br />
sequence from its top sequence boundary, which was assumed to be horizontal on the ramp, <strong>and</strong> allowing the ramp margin to<br />
slope into the basin so that water depths of facies are compatible with estimated water depths in the text. Dotted line shows<br />
estimated positions of maximum flooding surfaces. Legend is same as Figure 6A.<br />
FEET<br />
C5<br />
C4<br />
C3<br />
C2<br />
C1<br />
C11<br />
C10<br />
C9<br />
C8<br />
C7<br />
C6<br />
17
18<br />
23 24 25 26<br />
FORMATION<br />
SEQ.<br />
22<br />
21<br />
20<br />
?<br />
C12<br />
19<br />
18<br />
BLUEFIELD<br />
FM.<br />
17<br />
?<br />
GREENVILLE<br />
SHALE<br />
16<br />
15<br />
DROOP SS.<br />
13 14<br />
12<br />
C11<br />
?<br />
LILLYDALE SHALE<br />
DATUM<br />
ALDERSON<br />
FM.<br />
C10<br />
AUS AL-TAWIL, THOMAS C. WYNN, AND J. FRED READ<br />
UNION<br />
FM.<br />
C9<br />
C8<br />
100<br />
PICKAWAY<br />
FM.<br />
FEET<br />
C7<br />
10<br />
20 = WYOMING 253<br />
21 = WYOMING 87<br />
22 = MCDOWELL 63<br />
23 = MCDOWELL 792<br />
24 = MERCER 16<br />
25 = MERCER 1<br />
26 = MERCER 27<br />
MILES<br />
0<br />
12 = WAYNE 277<br />
13 = WAYNE 1497<br />
14 = LINCOLN 155<br />
15 = LINCOLN 1375<br />
16 = MINGO 796<br />
17 = LOGAN 102<br />
18 = LOGAN 922<br />
19 = LOGAN 388<br />
12 13 14 15 16 17 18 19 20 21 22 23 24 25<br />
C6<br />
TAGGARD<br />
EQUIVALENT<br />
DATUM<br />
C5<br />
?<br />
?<br />
C4<br />
OOID GRAINSTONE / PACKSTONE<br />
RED BEDS<br />
DENMAR<br />
FM.<br />
C3<br />
PRICE / MACCRADY FM.<br />
PELOID GRAINSTONE / PACKSTONE<br />
GRAY SHALE / SILTSTONE<br />
SKELETAL GRAINSTONE / PACKSTONE,<br />
MINOR WACKESTONE<br />
CALCAREOUS SILTSTONE<br />
QUARTZ SANDSTONE<br />
C2<br />
ARGILLACEOUS SKELETAL<br />
PACKSTONE / WACKESTONE<br />
LAMINATED SHALY LIME MUDSTONE/<br />
CALCAREOUS SILTSTONE<br />
QUARTZ-PELOID GRAINSTONE<br />
C1<br />
FINE-GRAINED PELLETAL<br />
PACKSTONE,LIME<br />
WACKESTONE / MUDSTONE<br />
SEQUENCE BOUNDARY<br />
DOLOMITE<br />
HILLSDALE<br />
FM.<br />
PALEOSOL<br />
FIG. 7.—A) High-resolution sequence stratigraphic cross section from updip to downdip in West Virginia based on subsurface wells analyzed using cuttings <strong>and</strong> wireline<br />
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<br />
the base of the upper Taggard equivalent (quartz–peloid eolianites–calcareous siltstones at base of sequence C6). Cross section shows the interpreted regional<br />
distribution of lithofacies, sequence boundaries, <strong>and</strong> their relation to the previous formational stratigraphy. Note the similarity of the sequence stratigraphy <strong>and</strong> facies<br />
distributions of this well-based cross section to the dominantly outcrop- <strong>and</strong> well-based cross section of Figure 6. Detailed well location information is in the Appendix.
MISSISSIPPIAN RAMP, WEST VIRGINIA, U.S.A.<br />
12 13 14 15 16 17 18 19 20 21 22 23 24 25 26<br />
100 FT. 30 M<br />
16 KM.<br />
0 10 MILES<br />
FIG. 7.—B) Facies distributions within individual sequences in Figure 7A, constructed by rehanging each sequence from its top<br />
sequence boundary, which was assumed to be horizontal. Legend is same as Figure 7A. Dotted line shows estimated positions<br />
of maximum flooding surfaces.<br />
sequence C11) that grade up into argillaceous, skeletal limestones<br />
(Figs. 6, 7). S<strong>and</strong>stone <strong>and</strong>/or calcareous siltstone forms units up<br />
to 10 meters thick along the ramp margin <strong>and</strong> slope. Calcareous<br />
siltstones <strong>and</strong> s<strong>and</strong>stones are festoon cross-bedded to waveripple<br />
cross laminated or massive in outcrop <strong>and</strong> generally are<br />
unfossiliferous (Table 1).<br />
The shales are quiet-water deposits. Poorly fossiliferous thin<br />
shales formed in lagoons shoreward of s<strong>and</strong>stone–siltstone units<br />
at bases of sequences (cf. Ettensohn et al., 1984). However, thick<br />
fossiliferous shales such as those in sequence C11 may be deeperramp,<br />
open marine units (cf. Ettensohn et al., 1984). High<br />
siliciclastic influx may have been due to more humid climate,<br />
19<br />
perhaps coupled with short-term sea-level lowering. This increased<br />
turbidity suppressed <strong>carbonate</strong> productivity on the ramp,<br />
generating widespread shale deposits. The calcareous siltstone<br />
contains sedimentary structures that suggest a high-energy wave<strong>and</strong><br />
current-influenced coastal setting during shallowing phases<br />
on the outer ramp.<br />
Quartz–Peloid Grainstone<br />
(Coastal Eolianites, Minor Marine S<strong>and</strong> Sheets)<br />
These facies generally are unfossiliferous, very fine to fine <strong>and</strong><br />
less commonly medium-grained quartz–peloid grainstone com-<br />
?<br />
?<br />
?<br />
C11<br />
C10<br />
C9<br />
C8<br />
C7<br />
C6<br />
C5<br />
C4<br />
C3<br />
C2<br />
C1<br />
MERAMECIAN<br />
SEQUENCES
20<br />
Lithology<br />
& Depositional<br />
Environment<br />
Caliche, Breccia <strong>and</strong><br />
Erosional Contacts<br />
(Subaerial)<br />
Color Paleosols yellow to<br />
brown<br />
Depositional<br />
Texture <strong>and</strong><br />
Grain Types<br />
Sedimentary<br />
Structures<br />
Paleosols have<br />
cryptocrystalline <strong>and</strong><br />
fibrous calcite crusts <strong>and</strong><br />
fracture fills; rooted<br />
fabrics common, <strong>and</strong><br />
patches of caliche-coated<br />
peloids <strong>and</strong> pisolites.<br />
Variably silicified. May<br />
be brecciated.<br />
Paleosols have<br />
undulating lamination,<br />
rootlet casts, <strong>and</strong><br />
incipient brecciation.<br />
Erosional contacts are<br />
sharp, planar to<br />
scalloped surfaces on the<br />
underlying unit.<br />
Fossils Rootlet casts in<br />
paleosols.<br />
posed of well rounded peloids, rare ooids, skeletal fragments,<br />
<strong>and</strong> finer-grained subangular quartz concentrated at bases of<br />
laminae (Table 2; Figs. 6, 7). They are cross-stratified, with foresets<br />
of extremely planar to curved, paper-thin, millimeter-scale laminae,<br />
that commonly are inversely graded (from very fine to fine<br />
s<strong>and</strong>) <strong>and</strong> form a characteristic pinstripe pattern on the outcrop<br />
or in core. Where inverse grading is absent, laminae form couplets<br />
of alternating very fine <strong>and</strong> fine s<strong>and</strong>. These facies commonly<br />
are interbedded with red beds, lagoonal facies, <strong>and</strong> ooidshoal<br />
facies.<br />
These facies are interpreted as coastal eolianites on the basis<br />
of abundant quartz–peloid grainstone lithology, the well rounded<br />
<strong>carbonate</strong> grain shapes, planar to curved foresets with inverse<br />
grading, <strong>and</strong> very fine to fine s<strong>and</strong> couplets, pinstripe lamination,<br />
<strong>and</strong> lack of in situ fossils or burrows (Merkely, 1991; Hunter, 1993;<br />
Dodd et al., 1993; Smith, 1996). Some of the coarser-grained, lowangle<br />
cross-laminated units may be beach facies.<br />
Fine-Grained Lime Wackestone–Mudstone<br />
(Tidal Flat <strong>and</strong> Lagoon)<br />
This facies occurs in massive to laminated units 1 to 8 meters<br />
thick. They include thin to medium-bedded, light gray, finegrained<br />
lime wackestone–mudstone <strong>and</strong> pellet packstone, with a<br />
generally restricted assemblage of gastropods, ostracodes, <strong>and</strong><br />
rare small fragments of echinoderms <strong>and</strong> brachiopods (Table 2;<br />
Figs. 6, 7). Bedding planes locally contain abundant small bur-<br />
AUS AL-TAWIL, THOMAS C. WYNN, AND J. FRED READ<br />
TABLE 1.—Mississippian Paleosols <strong>and</strong> Siliciclastic Facies.<br />
Siliciclastic paleosols<br />
<strong>and</strong> red-beds<br />
(Subaerial to Marginal<br />
Marine)<br />
Red, maroon <strong>and</strong> green<br />
mottled<br />
Terrigenous mudrocks<br />
<strong>and</strong> siltstones.<br />
Redbeds massive, to<br />
laminated, rare current<strong>and</strong><br />
wave ripples on<br />
siltstone interbeds, rare<br />
mudcracks. Paleosols<br />
massive mudrocks with<br />
slickensides<br />
Root/ burrow traces in<br />
paleosols<br />
Gray shale<br />
(Low Energy Shallow<br />
Marine)<br />
Quartz s<strong>and</strong>stone/<br />
calcareous siltstones<br />
(Coastal to Shallow<br />
Marine)<br />
Dark gray to olive green White to light gray<br />
Terrigenous clay <strong>and</strong> silt Well-sorted fine to<br />
medium -grained shaly<br />
s<strong>and</strong>stones <strong>and</strong><br />
siltstones of quartz,<br />
lesser <strong>carbonate</strong> grains<br />
Poorly fissile to massive Cross-bedded to<br />
structureless; flaser,<br />
lenticular <strong>and</strong> wavy<br />
bedded locally<br />
Poorly fossiliferous to<br />
very fossiliferous.<br />
Mollusks, ostracods,<br />
some echinoderms<br />
brachiopods <strong>and</strong><br />
bryozoa; biota sparse<br />
<strong>and</strong> restricted in updip<br />
shale<br />
Rare<br />
rows <strong>and</strong> trails. In the east, some of these facies are microbially<br />
laminated dolomitic mudstones with centimeter layering, rare<br />
mudcracks, <strong>and</strong> few fossils. These laminated units are common<br />
downdip on the ramp in the eastern outcrop belt, but more<br />
massive units are widespread on the ramp (Table 2; Figs. 6, 7).<br />
The laminated facies are tidal-flat deposits developed downdip<br />
on the ramp during periods of emergence. The massive, finegrained<br />
<strong>carbonate</strong>s were formed in low-energy lagoonal settings<br />
on the ramp interior <strong>and</strong> between grainstone shoals on the ramp.<br />
Lack of open-marine fossils, <strong>and</strong> association with shallow-marine<br />
facies <strong>and</strong> exposure surfaces in updip areas, indicate a<br />
relatively restricted, possibly periodically hypersaline lagoonal<br />
setting.<br />
Peloid <strong>and</strong> Ooid Grainstone (Shoals <strong>and</strong> Tidal Bars)<br />
These are white, cross-bedded, well-sorted grainstones composed<br />
of medium-s<strong>and</strong>-size ooids, peloids, intraclasts, <strong>and</strong><br />
abraded skeletal fragments of echinoderms, bryozoa, brachiopods,<br />
<strong>and</strong> mollusks. They form units from 1 to 12 meters thick<br />
(Table 2; Figs. 6, 7). In the subsurface, the oolites form diporiented<br />
northwest-to-southeast trending oolitic s<strong>and</strong> ridges over<br />
30 km long <strong>and</strong> 1.4 km wide along the ramp margin (Kelleher <strong>and</strong><br />
Smosna, 1993; Smosna <strong>and</strong> Koehler, 1993). Both ooid <strong>and</strong> skeletal<br />
grainstones locally contain rare, small bryozoan mounds.<br />
These peloid–oolitic facies (Table 2) formed on high-energy<br />
tide-influenced <strong>and</strong> lesser wave-influenced ooid shoals in water
Facies<br />
& Depositional<br />
Environment.<br />
Quartz Peloid<br />
Grainstone<br />
(Coastal Eolianite,<br />
Minor Marine S<strong>and</strong><br />
Sheets)<br />
depths from the intertidal zone down to 10 meters (Kelleher <strong>and</strong><br />
Smosna, 1993; Keith <strong>and</strong> Zuppman, 1990).<br />
Skeletal Grainstone–Packstone<br />
(Ramp Margin <strong>and</strong> Moderate-Energy Lagoon)<br />
Because they are difficult to differentiate in well cuttings, all<br />
the coarse-grained, skeletal grainstones, skeletal packstones, <strong>and</strong><br />
grain-rich wackestones were lumped together. In outcrop, these<br />
are medium- to thick-bedded, <strong>and</strong> cross-bedded to structureless.<br />
They are medium- to coarse-grained skeletal units ranging from<br />
clean grainstones to muddy packstone <strong>and</strong> lesser wackestone.<br />
They consist of whole <strong>and</strong> fragmented crinoid, bryozoan, brachiopod,<br />
<strong>and</strong> lesser mollusk remains (Table 2). Some skeletal<br />
beds contain abundant whole columnals <strong>and</strong> calices. The skeletal<br />
grainstone units form extensive sheets seaward of oolitic<br />
grainstone shoals along the ramp margin in units up to 15 meters<br />
thick, which interfinger with wackestone <strong>and</strong> deep-water laminated<br />
marls (Figs. 6, 7). The more poorly washed facies form<br />
relatively extensive units on the ramp, interfingering laterally<br />
with <strong>and</strong> surrounded by poorly fossiliferous lime wackestone–<br />
mudstone (Figs. 6, 7).<br />
These facies include high-energy skeletal banks <strong>and</strong> sheets<br />
deposited in wave-agitated <strong>and</strong> tidal-current settings between<br />
<strong>and</strong> seaward of ooid shoals along the ramp margin, in water<br />
depths from a few meters to perhaps 20 meters. They also formed<br />
extensive units within the lagoon, in areas colonized by abundant<br />
crinoids <strong>and</strong> other skeletal organisms, <strong>and</strong> were subjected to low<br />
to moderate wave reworking.<br />
Argillaceous Skeletal Wackestone<br />
These are medium- to thick-bedded in outcrop <strong>and</strong> are dark<br />
gray shaly skeletal wackestones with some packstone interbeds.<br />
They are composed of medium- to very coarse-grained echinoderm,<br />
brachiopod, <strong>and</strong> bryozoan debris, with interstitial finely<br />
fragmented skeletal debris, lime mud, <strong>and</strong> argillaceous seams<br />
(Table 2). Some wackestone units have decimeter beds of skeletal<br />
grainstone–packstone. This facies locally contains very small<br />
MISSISSIPPIAN RAMP, WEST VIRGINIA, U.S.A.<br />
TABLE 2. —Mississippian Carbonate Facies.<br />
Fine-Grained Lime<br />
Wackestone/ Mudstone,<br />
Minor Lime Packstone<br />
(Tidal Flat <strong>and</strong> Low-<br />
Energy Lagoon)<br />
Color Dark-gray weathering Light gray. Tan to yellow<br />
weathering<br />
Grain Types And<br />
Depositional<br />
Texture<br />
Rounded <strong>and</strong> abraded<br />
peloids <strong>and</strong> some ooids,<br />
abraded skeletal<br />
fragments, <strong>and</strong><br />
subangular very fine to<br />
fine quartz up to 50%.<br />
Grainstone fabrics.<br />
Unfossiliferous to<br />
moderately fossiliferous<br />
fine wackestone <strong>and</strong><br />
mudstone <strong>and</strong> pellet<br />
packstone. Skeletal debris<br />
fine grained; abundant lime<br />
mud, dolomite mud, <strong>and</strong><br />
pellets; may contain quartz<br />
silt <strong>and</strong> clay. Locally<br />
cherty.<br />
Sedimentary<br />
Structures<br />
Wedge sets of planner<br />
<strong>and</strong> tangential crossbedding,<br />
with pinstripe<br />
lamination, <strong>and</strong> inverse<br />
grading, dips up to 20<br />
degrees<br />
Biota Abraded, rounded<br />
skeletal fragments. No<br />
in situ biota or marine<br />
trace fossils.<br />
Dolomitic mudstones<br />
massive to microbially<br />
laminated (centimeter-thick<br />
laminae, micrograded,<br />
rarely mudcracked). Finegrained,<br />
skeletal-rich<br />
wackestone-mudstone is<br />
typically massive.<br />
None to sparse; may have<br />
mollusks, small crinoid<br />
columnals, ostracods. May<br />
have trace-fossil horizons,<br />
small oncolites, rare corals,<br />
<strong>and</strong> brachiopods.<br />
Peloid <strong>and</strong> Ooid<br />
Grainstone<br />
(High-Energy<br />
Shoal)<br />
Skeletal Grainstone/<br />
Packstone<br />
(Ramp Margin <strong>and</strong><br />
Lagoonal Skeletal<br />
Sheets <strong>and</strong> Shoals)<br />
Argillaceous Skeletal<br />
Wackestone<br />
(Deep Subtidal)<br />
21<br />
bryozoan <strong>reefs</strong>, especially in sequences C10 <strong>and</strong> C11. Argillaceous<br />
packstone–wackestone is most abundant along the ramp<br />
margin (Figs. 6, 7). They formed below fair-weather wave base<br />
<strong>and</strong> above storm wave base on the open ramp in water depths of<br />
20 to 60 m by comparison with the Persian Gulf (Purser <strong>and</strong><br />
Seibold, 1973). However, these depths to maximum storm wave<br />
base could have been shallower in the Appalachian Basin, given<br />
that dominant wind directions in the Persian Gulf are axial to the<br />
basin <strong>and</strong> thus have maximum fetch, whereas dominant winds in<br />
the Appalachian Basin blew across the basin (Golonka et al.,<br />
1994). Siliciclastic muds may have washed out onto the ramp<br />
following floods or storms, or they may mark brief cessations in<br />
<strong>carbonate</strong> deposition during short periods of more humid climate<br />
<strong>and</strong> higher siliciclastic influx.<br />
Laminated Shaly Lime Mudstone (Ramp Slop to Basin)<br />
These are unfossiliferous, dark gray, evenly thin-laminated to<br />
massive, commonly silty or argillaceous lime-mudstone units up<br />
to 100 meters thick along the ramp slope (Fig. 6, section 11; Table 2).<br />
They consist of fine, silt-size <strong>carbonate</strong> <strong>and</strong> quartz, finely comminuted<br />
skeletal debris, <strong>and</strong> a matrix of argillaceous lime mudstone.<br />
The shaly lime mudstones are deep-water ramp-slope to<br />
basin-floor facies, which lack shallow-water sedimentary structures<br />
<strong>and</strong> biota. Fine <strong>carbonate</strong> sediment was brought onto the<br />
slope perhaps by storms <strong>and</strong> tidal currents <strong>and</strong> mixed with l<strong>and</strong>derived<br />
terrigenous silts <strong>and</strong> clays. Lack of burrowing suggests<br />
deposition in anoxic waters below the photic zone in depths of 60<br />
to over 100 m; the photic zone in the basin perhaps extended<br />
down to only to 30 m, given the turbid basin waters, whereas<br />
wave base perhaps reached down to more than 30 m but less than<br />
60 m, by comparison with the present Persian Gulf, a similarsized<br />
<strong>and</strong> similarly situated forel<strong>and</strong> basin (Purser <strong>and</strong> Seibold,<br />
1973).<br />
SEQUENCE STRATIGRAPHY<br />
Laminated Shaly<br />
Lime Mudstone<br />
(Ramp slope-<br />
Basin)<br />
Light gray to white Light to medium gray Medium gray to dark gray Typically dark gray<br />
Well sorted,<br />
rounded, medium to<br />
coarse grainstone.of<br />
s<strong>and</strong>-size ooids,<br />
peloids, lesser<br />
skeletal fragments,<br />
<strong>and</strong> minor<br />
intraclasts.<br />
Cross-bedded to<br />
massive<br />
Crinoid , brachiopod,<br />
bryozoan, <strong>and</strong><br />
mollusk fragments,<br />
forams<br />
Variably fragmented<br />
echinoderms,<br />
brachiopod, bryozoa,<br />
mollusks, <strong>and</strong> rare ooids.<br />
Mud-free grainstones to<br />
grain-rich packstones<br />
<strong>and</strong> minor wackestones,<br />
some with argillaceous<br />
seams.<br />
May be cross-bedded,<br />
thin- to medium-bedded<br />
or structureless<br />
Abundant echinoderms,<br />
brachiopods, <strong>and</strong><br />
bryozoas; <strong>and</strong> lesser<br />
mollusks.<br />
Echinoderms,<br />
brachiopods, bryozoa, <strong>and</strong><br />
lesser mollusks. Include<br />
wackestone <strong>and</strong> lesser<br />
packstone; abundant lime<br />
mud; terrigenous clay<br />
disseminated or in seams<br />
<strong>and</strong> stringers<br />
Massive, thick to thin<br />
bedded, structureless;<br />
may have graded storm<br />
beds<br />
Abundant echinoderm,<br />
common brachiopods <strong>and</strong><br />
bryozoas, rare mollusks<br />
Carbonate <strong>and</strong><br />
silisiclastic clay <strong>and</strong><br />
silt<br />
Millimeter to<br />
centimeter even<br />
lamination to<br />
structureless<br />
None<br />
The sequence stratigraphic terms used are after Sarg (1988),<br />
Mitchum <strong>and</strong> Van Wagoner (1991), <strong>and</strong> Weber et al. (1995). Facies
22<br />
<strong>and</strong> depositional sequences are shown in Figure 6A (extending<br />
northeast to southwest down the axis of the foredeep <strong>and</strong> into the<br />
proximal forel<strong>and</strong> basin, with sections arranged from thinnest to<br />
thickest) <strong>and</strong> in Figure 7A (oriented east–west at right angles to<br />
the forel<strong>and</strong> basin, <strong>and</strong> extending across the distal forel<strong>and</strong>, <strong>and</strong><br />
just beyond the hinge into the margin of the proximal forel<strong>and</strong>).<br />
Depositional sequences <strong>and</strong> their component facies were traceable<br />
from the outcrop sections into the subsurface using well<br />
cuttings <strong>and</strong> wireline logs, especially gamma-ray logs.<br />
Sequences were recognized on the basis of major l<strong>and</strong>ward<br />
<strong>and</strong> basinward shifts in diagnostic facies belts such as eolianites,<br />
red beds, siliciclastic s<strong>and</strong>s <strong>and</strong>/or calcareous siltstones, <strong>and</strong> lime<br />
grainstones. Sequence boundaries were picked at erosional disconformities<br />
(commonly evident in outcrop sections) on top of<br />
shallowing-upward <strong>carbonate</strong> units, <strong>and</strong> below lowst<strong>and</strong> or<br />
transgressive siliciclastic red beds, s<strong>and</strong>stones, <strong>and</strong> shale (evident<br />
in both subsurface <strong>and</strong> outcrop). Along the ramp slope,<br />
sequence boundaries were placed beneath the shallowest-water<br />
facies within the dominantly deep-water successions. In wireline<br />
logs, many sequence boundaries (<strong>and</strong> some parasequence boundaries)<br />
are associated with significant gamma-ray kicks due to<br />
presence of terrigenous mudrocks, shale, or argillaceous dolomites.<br />
This was important in sections containing parasequences,<br />
because it allowed parasequences to be grouped into sequences<br />
on the basis of the characteristic upward-decreasing (as the<br />
limestones clean upward) to upward-increasing gamma-ray signature<br />
(as the <strong>carbonate</strong>s become more siliciclastic-rich). Because<br />
it was not clear using the subsurface data whether siliciclastic<br />
units were late highst<strong>and</strong>, lowst<strong>and</strong>, or early transgressive, they<br />
have been arbitrarily referred to as lowst<strong>and</strong> if on the ramp slope<br />
<strong>and</strong> margin, <strong>and</strong> lowst<strong>and</strong> to early transgressive if updip. Where<br />
possible, maximum flooding surfaces were picked below the<br />
deepest marine facies within the sequence, but the subsurface<br />
data commonly prevented them being traced regionally; thus it<br />
was possible to separate the transgressive <strong>and</strong> highst<strong>and</strong> systems<br />
tracts only locally on the ramp. Some of the maximum flooding<br />
surfaces are associated with high gamma-ray values, especially<br />
where the <strong>carbonate</strong>s are shaly, but many in pure <strong>carbonate</strong><br />
sections are not.<br />
Fourth-order sequences (roughly between 0.1 <strong>and</strong> 0.5 My<br />
duration) are the dominant units mapped in the study interval in<br />
the Appalachian Basin <strong>and</strong> the Illinois Basin (Smith <strong>and</strong> Read,<br />
1999, 2001; Al-Tawil <strong>and</strong> Read, this volume). Third-order or<br />
composite sequences (Weber et al., 1995) are far less well developed,<br />
being masked by the profound facies shifts within the<br />
dominant fourth-order sequences. Individual fourth-order sequences<br />
<strong>and</strong> component facies are shown in Figures 6B <strong>and</strong> 7B,<br />
constructed by rehanging the sections from the top of each<br />
sequence. Major biostratigraphic markers include Acrocyathus in<br />
the Hillsdale Formation, the last appearance of Platycrinites penicillus<br />
at the top of the Denmar Formation, the first appearance of<br />
Talarocrinus in the Pickaway Formation, as well as the conodont<br />
zones (Fig. 3). Individual fourth-order sequences were below the<br />
resolution of the biostratigraphic data, which typically span<br />
several sequences, but they do form valuable biostratigraphic<br />
markers within the succession.<br />
The sparse radiometric data <strong>and</strong> their error bars, <strong>and</strong> the<br />
difficulties in intercontinental correlation between Mississippian<br />
brachiopod, conodont, <strong>and</strong> foraminiferal zones, make absolute<br />
age determination of the study interval difficult. The St. Louis–<br />
Hillsdale interval is probably about 5 My in duration (Ross <strong>and</strong><br />
Ross, 1987); sequences C1 to C11 (the Denmar to Glenray interval)<br />
are about 3 to 8 My (S<strong>and</strong>o, 1984; Baxter <strong>and</strong> Brenckler, 1982;<br />
Roberts et al., 1995; Smith, 1996). Assigning duration of 3 to 8 My<br />
for the three third-order sequences in the Denmar to Glenray<br />
AUS AL-TAWIL, THOMAS C. WYNN, AND J. FRED READ<br />
interval gives a duration of 1 to 2.5 My for each of the composite<br />
sequences <strong>and</strong> 280 ky to an extreme upper limit of 720 ky for each<br />
of the 11 fourth-order sequences, with the most likely durations<br />
being 300 to 400 ky. (Smith, 1996). This encompasses the estimated<br />
average duration of 375 ky for the time-equivalent fourthorder<br />
sequences in the Illinois Basin (Smith, 1996) <strong>and</strong> the 400 ky<br />
durations estimated for the overlying Mauch Chunk siliciclastic<br />
fourth-order sequences (Miller <strong>and</strong> Eriksson, 1999).<br />
Fourth-Order Depositional Sequences<br />
Meramecian Sequences.—<br />
These Meramecian sequences (Figs. 6, 7) form the basal part of<br />
the study interval <strong>and</strong> include the Little Valley <strong>and</strong> Hillsdale<br />
formations. They unconformably (?) overlie the Price–Maccrady<br />
red beds, resting on a widespread transgressive s<strong>and</strong>stone in<br />
many places. The Meramecian sequences thicken from 0 to over<br />
270 m across the hinges into the basin (Fig. 6, section 11; Bartlett<br />
<strong>and</strong> Webb, 1971). Two to four Meramecian sequences may be<br />
present on the ramp. The sequence boundaries are placed beneath<br />
widespread lowst<strong>and</strong> to early transgressive s<strong>and</strong>stone–<br />
calcareous siltstone on the ramp to the west (Fig. 7), which appear<br />
to pinch out to the north, where two sequences may make up the<br />
interval (sequences not plotted individually in Fig. 6B). The<br />
transgressive to highst<strong>and</strong> parts of the sequences consist mainly<br />
of commonly cherty, fine skeletal wackestone–mudstone, grading<br />
basinward into shaly lime wackestone or lime mudstone. In<br />
the thick, proximal forel<strong>and</strong> basin sections, several sequences<br />
may be present (not shown in Fig. 6). Lower sequences here<br />
(mapped as Little Valley Formation) consist of quartz s<strong>and</strong>stone<br />
overlain by dark gray, nearshore argillaceous skeletal wackestone–<br />
mudstone, whereas the overlying Hillsdale sequences consist of<br />
oncolitic skeletal wackestone–mudstone capped by skeletal<br />
grainstone (Huggins, 1983).<br />
Chesterian Sequences.—<br />
Sequence C1.—Sequence C1 is the basal sequence of the Denmar<br />
Limestone (Figs. 6, 7). It generally overlies the Upper Meramecian<br />
Hillsdale Limestone <strong>and</strong> is 0 to 55 m thick (Fig. 6). The basal<br />
sequence boundary is placed beneath widespread basal, thin<br />
siliciclastic-prone units. Sequence C1 has a basal lowst<strong>and</strong> unit<br />
of local red beds that grades seaward into quartz s<strong>and</strong>stone–<br />
siltstone a few meters thick along the ramp margin <strong>and</strong> ramp<br />
slope (Fig. 6, sections 7 to 9 <strong>and</strong> 11; Fig. 7, sections 25, 26). These<br />
grade updip onto the ramp interior into lowst<strong>and</strong>–transgressive<br />
s<strong>and</strong>stone–calcareous siltstone, shale, <strong>and</strong> quartz–peloid<br />
grainstones (Fig. 6; sections 3 to 6; Fig. 7, sections 19 to 26) that<br />
may be overlain by possible transgressive <strong>carbonate</strong>s in the<br />
middle of the sequence in the west (Fig. 7). The transgressive<br />
systems tract along the ramp slope consists of up to 20 meters of<br />
skeletal wackestone–mudstone grading up into skeletal grainstone<br />
(Fig. 6, section 11). The highst<strong>and</strong> systems tract downdip consists<br />
of deeper-water, laminated mudstones <strong>and</strong> skeletal wackestone<br />
that passes upsection into skeletal grainstone. Updip the highst<strong>and</strong><br />
systems tract consists of thin, ramp-margin skeletal–ooid<br />
grainstones that grade updip into lime wackestone–mudstone<br />
<strong>and</strong> discrete ooid–peloid grainstone shoal deposits (Fig. 6, sections<br />
1 to 11; Fig. 7, sections 19 to 26).<br />
Sequence C2.—Sequence C2, the next Denmar sequence, is 0 to<br />
75 m thick (Figs. 6, 7); it thickens markedly between sections 6 <strong>and</strong><br />
7 <strong>and</strong> sections 10 <strong>and</strong> 11 (Fig. 6). The sequence boundary is placed<br />
beneath a thin quartz s<strong>and</strong>stone on the ramp margin <strong>and</strong> slope,
<strong>and</strong> updip beneath widespread siltstone to the west (Fig. 7) <strong>and</strong><br />
beneath eolianites updip to the north (Fig. 6). It has a lowst<strong>and</strong><br />
quartz s<strong>and</strong>stone a few meters thick along the ramp margin <strong>and</strong><br />
ramp slope (Fig. 6, section 11; Fig. 7, sections 24 to 26). The<br />
transgressive systems tract along the ramp slope consists of<br />
skeletal wackestone (Fig. 6, section 11; Fig. 7, section 26), <strong>and</strong><br />
within a broad sag on the ramp margin (Fig. 6, sections 6 to 10),<br />
it probably includes the lower third of the sequence, consisting of<br />
quartz peloid <strong>and</strong> ooid grainstone <strong>and</strong> lime wackestone–mudstone<br />
beneath the lowest skeletal grainstones. The lowst<strong>and</strong> to<br />
transgressive systems tract updip includes thin quartz–peloid<br />
eolianites, <strong>and</strong> the lower portion of ooid grainstones in the<br />
northeast (Fig. 6, sections 1 to 5); to the west, the lowst<strong>and</strong> to<br />
transgressive tract consists of the regional calcareous siltstones<br />
(Fig. 7). The highst<strong>and</strong> systems tract along the ramp slope consists<br />
of deep-water shaly lime mudstone overlain by skeletal<br />
wackestone (Fig. 6, section 11; Fig. 7, section 26). Farther updip it<br />
consists of ramp-margin skeletal <strong>and</strong> minor ooid grainstones<br />
(Fig. 6, sections 9, 10; Fig. 7, section 25) that interfinger with rampinterior<br />
lime wackestone–mudstone containing multilateral ooid<br />
grainstone units (Figs. 6, 7). The vertical repetition of facies<br />
suggests that two to three parasequences may make up the<br />
sequence between sections 6 to 10 (Fig. 6).<br />
Sequence C3.—This next Denmar sequence is 0 to 40 m thick<br />
(Figs. 6, 7). Its basal sequence boundary is placed beneath thin<br />
siliciclastic units or eolianite (Figs. 6, 7). The lowst<strong>and</strong> systems<br />
tract is a thin quartz s<strong>and</strong>stone on the ramp slope (Fig. 6, section<br />
11). The transgressive systems tract is a wedge of skeletal wackestone<br />
overlain by skeletal grainstone along the ramp slope (Fig. 6,<br />
section 11). Updip the transgressive systems tract may include<br />
the lower half of the sequence, <strong>and</strong> is composed of thin, basal<br />
siliciclastic <strong>and</strong> eolianite units <strong>and</strong> the overlying lower parts of<br />
the discontinuous ooid grainstone <strong>and</strong> lime wackestone–mudstone<br />
units (Figs. 6, 7). The highst<strong>and</strong> systems tract constitutes the<br />
upper part of the sequence <strong>and</strong> consists of a patchwork of skeletal,<br />
oolitic, <strong>and</strong> lime wackestone–mudstone facies (Figs. 6, 7).<br />
Sequence C4.—Sequence C4, the top Denmar sequence, ranges<br />
from 5 to 80 m thick with marked thickening between sections 10<br />
<strong>and</strong> 11 (Fig. 6) <strong>and</strong> sections 24 to 26 (Fig. 7). On the ramp margin,<br />
the sequence boundary is a double disconformity in outcrop (Fig.<br />
6, section 10), <strong>and</strong> on the ramp slope the sequence boundary is<br />
beneath lowst<strong>and</strong> quartz s<strong>and</strong>stone several meters thick (Fig. 6,<br />
section 11) or calcareous siltstone (Fig. 7, sections 24 to 26). The<br />
sequence boundary on the ramp interior is placed beneath discontinuous<br />
lowst<strong>and</strong> eolianites (Fig. 6, sections 3, 7) or thin basal<br />
siliciclastic or quartz–peloid grainstone units (Fig. 7). The transgressive<br />
systems tract is absent on the slope, but on the ramp it<br />
may include thin local units of ooid grainstone <strong>and</strong> quartz–peloid<br />
grainstone in the north <strong>and</strong> east (Fig. 6, sections 3, 7, 10); in the<br />
west, the lowst<strong>and</strong> to transgressive systems tract consists of an<br />
extensive basal siliciclastic sheet (Fig. 7, sections 19 to 23). The<br />
highst<strong>and</strong> systems tract makes up the bulk of the sequence, <strong>and</strong><br />
consists of ramp-margin skeletal grainstone–packstone <strong>and</strong> ooid–<br />
peloid grainstone, <strong>and</strong> ramp-interior, discontinuous units of lime<br />
wackestone–mudstone that separate laterally, local ooid–peloid<br />
shoal deposits (Figs. 6, 7). On the ramp slope, the highst<strong>and</strong> tract<br />
consists of deeper-water shaly mudstones (Fig. 6, section 11).<br />
Sequence C5.—Sequence C5, 0 to 25 m thick, includes the<br />
lower Taggard red beds (Figs. 6, 7). It shows a considerable<br />
basinward shift in red siliciclastic facies compared with the<br />
underlying sequences. The sequence boundary is beneath widespread<br />
siliciclastics <strong>and</strong> local quartz–peloid eolianites. The<br />
MISSISSIPPIAN RAMP, WEST VIRGINIA, U.S.A.<br />
23<br />
lowst<strong>and</strong> to early transgressive systems tract consists of red<br />
beds updip in the northeast (Fig. 6) <strong>and</strong> on the ramp margin in<br />
the west (Fig. 7, section 26); thin, local eolianites <strong>and</strong> s<strong>and</strong>stone–<br />
siltstone on ramp in the west (Fig. 7, sections 19 to 22); <strong>and</strong> up<br />
to 15 m of s<strong>and</strong>stone–calcareous siltstone <strong>and</strong> marine shale on<br />
the ramp margin <strong>and</strong> slope (Fig. 6, sections 8, 9; Fig. 7, sections<br />
20 to 25). No transgressive systems tract is recognized in the<br />
north <strong>and</strong> east (Fig. 6). In the west, however, the transgressive<br />
tract may include the lower half of the <strong>carbonate</strong> interval,<br />
composed of basal units of ooid <strong>and</strong> peloid grainstone <strong>and</strong> the<br />
overlying lime wackestone–mudstones, if these are shallow<br />
lagoonal facies (Fig. 7, sections 17 to 22). The highst<strong>and</strong> tract on<br />
the shallow ramp is dominated by lime wackestone–mudstone,<br />
with ramp-margin grainstone units, <strong>and</strong> localized updip ooid<br />
<strong>and</strong> skeletal grainstone bodies (Figs. 6, 7); the highst<strong>and</strong> systems<br />
tract on the slope consists of basinal shaly lime mudstone<br />
grading up into skeletal wackestone (Fig. 6, section 11).<br />
Sequence C6.—Sequence C6 is 0 to 40 m thick (Figs. 6, 7) <strong>and</strong><br />
thickens rapidly between sections 5 <strong>and</strong> 6 (Fig. 6). It includes the<br />
upper Taggard red beds <strong>and</strong> the lower Pickaway Limestone. The<br />
sequence boundary is placed below extensive siliciclastics (Fig. 6)<br />
or beneath widespread quartz–peloid eolianites (Fig. 7). The<br />
lowst<strong>and</strong> to early transgressive deposits consist of regional red<br />
beds up to 10 m thick in the north <strong>and</strong> east (Fig. 6) or regional<br />
eolianites to the west (Fig. 7). Downdip these grade into up to 30<br />
m of lowst<strong>and</strong> marine s<strong>and</strong>stone locally split by a thin shale–<br />
calcareous siltstone (Fig. 6, sections 9 to 11; Fig. 7, sections 20 to<br />
26). The transgressive systems tract on the ramp slope consists of<br />
a 15-m-thick succession of skeletal grainstone capped by a thin<br />
ooid grainstone (Fig. 6, section 11), which appears to grade updip<br />
into peloid grainstone (Fig. 7, section 26). On the ramp, it is<br />
difficult to separate the transgressive systems tract from the<br />
highst<strong>and</strong> systems tract. These combined systems tracts are 0 to<br />
35 m thick. Along the ramp margin in the west, probable highst<strong>and</strong><br />
deposits include a downdip ooid grainstone unit grading updip<br />
into skeletal wackestone <strong>and</strong> then into ramp-margin skeletal<br />
grainstone–packstone (Fig. 7, sections 23 to 26); these likely grade<br />
downdip into skeletal grainstone, skeletal wackestones, <strong>and</strong> then<br />
basinal lime mudstone (Fig. 6, section 11). On the ramp interior in<br />
the northeast, the transgressive to highst<strong>and</strong> systems tract is<br />
dominated by widespread units of lime wackestone–mudstone<br />
(Figs. 6, 7) with eolianites in the far west. Skeletal grainstone–<br />
packstone <strong>and</strong> minor ooid grainstones also are localized within<br />
the upper third of the sequence in a broad sag on the middle ramp<br />
in the northeast (Fig. 6, sections 6, 7).<br />
Sequence C7.—Sequence C7 includes the upper Pickaway<br />
Limestone. It is 0 to 50 m thick, thickening between sections 3 <strong>and</strong><br />
4 (Fig. 6) <strong>and</strong> between sections 22 to 26 (Fig. 7). The C7 sequence<br />
boundary on the ramp <strong>and</strong> ramp margin is placed below lowst<strong>and</strong><br />
to early transgressive s<strong>and</strong>stone, shale, calcareous siltstone, minor<br />
red beds, <strong>and</strong> quartz–peloid eolianites in the east <strong>and</strong> northeast<br />
(Fig. 6, sections 5 to 9; Fig. 7, sections 18 to 26), <strong>and</strong> beneath<br />
skeletal wackestone on the ramp slope (Fig. 6, section 11). The<br />
transgressive systems tract over much of the ramp in the northeast<br />
contains skeletal grainstone interfingering laterally with<br />
lime wackestone–mudstone, <strong>and</strong> is capped downdip by a thin<br />
shale <strong>and</strong> s<strong>and</strong>stone <strong>and</strong> updip by s<strong>and</strong>stone <strong>and</strong> ooid grainstone<br />
(Fig. 6, sections 3 to 9); the transgressive succession is not clearly<br />
defined in the west but may include a wedge of skeletal wackestone<br />
overlain by calcareous siltstone–shale along the ramp margin<br />
(Fig. 7, sections 24 to 26). On the ramp slope, the transgressive<br />
systems tract is a skeletal wackestone (Fig. 6, section 11). The<br />
highst<strong>and</strong> tract along the ramp slope is a deep-water shaly lime
24<br />
mudstone (Fig. 6, section 11). On the ramp, the highst<strong>and</strong> systems<br />
tract in the east contains two belts of skeletal–ooid grainstone, one<br />
along the ramp margin (Fig. 6, section 9) <strong>and</strong> the other updip to<br />
the north flanking a significant downwarp (Fig. 6, sections 4 to 6);<br />
these updip grainstones are bordered both updip <strong>and</strong> downdip<br />
by lime wackestone–mudstone. In the west, the highst<strong>and</strong> is a<br />
simple facies succession of skeletal grainstone <strong>and</strong> skeletal<br />
wackestone that passes updip into ooid limestone <strong>and</strong> then into<br />
lime wackestone–mudstone (Fig. 7). Three or more parasequences<br />
may be present in the sequence.<br />
Sequence C8.—Sequence C8 consists of the lower Union Limestone<br />
<strong>and</strong> is 0 to 55 m thick (Figs. 6, 7). It thickens between sections<br />
3 <strong>and</strong> 7 (Fig. 6). The C8 sequence boundary is placed beneath<br />
lowst<strong>and</strong> to early transgressive deeper-water skeletal wackestones<br />
along the lower ramp slope (Fig. 6, section 11), <strong>and</strong> beneath<br />
lowst<strong>and</strong> to transgressive s<strong>and</strong>stone, calcareous siltstone, <strong>and</strong><br />
rare red beds on the upper ramp slope <strong>and</strong> ramp margin (Fig. 6,<br />
section 9; Fig. 7, sections 22 to 26). The lowst<strong>and</strong> to early transgressive<br />
systems tract on the ramp interior may include basal,<br />
thin, patchily developed shales, peloid grainstone, <strong>and</strong> quartz–<br />
peloid eolianites (Fig. 6, 7). The transgressive systems tract on the<br />
ramp margin in the east makes up the lower one third of the<br />
sequence, <strong>and</strong> includes skeletal wackestone (Fig. 6, section 11)<br />
grading upsection <strong>and</strong> updip into skeletal grainstone <strong>and</strong> lesser<br />
ooid grainstone (Fig. 6, sections 7 to 9) <strong>and</strong> thin siliciclastic units<br />
<strong>and</strong> quartz–peloid eolianite (Fig. 6, sections 3 to 5; Fig. 7, sections<br />
18 to 25). The highst<strong>and</strong> systems tract is thick <strong>and</strong> dominated by<br />
an upward-shallowing succession of deeper-water mudstone to<br />
skeletal wackestone along the ramp slope (Fig. 6, section 11). On<br />
the ramp margin <strong>and</strong> ramp interior, the highst<strong>and</strong> systems tract<br />
includes at least two shallowing-upwards successions of skeletal<br />
wackestone up into skeletal grainstone that is bordered updip by<br />
locally thick ooid grainstone (Fig. 6, sections 8, 9; Fig. 7, sections<br />
23 to 26); these grade updip into lime wackestone–mudstone with<br />
local units of ooid grainstone (Fig. 6), <strong>and</strong> in the west, some thick<br />
eolianites (Fig. 7, sections 12, 19) separated laterally by lime<br />
wackestone–mudstone <strong>and</strong> containing a local shale in the middle<br />
of the succession (Fig. 6; Fig. 7, sections 18, 19).<br />
Sequence C9.—Sequence C9 is 5 to 25 m thick, <strong>and</strong> includes the<br />
upper Union Limestone (Figs. 6, 7). Updip in outcrop, the basal<br />
sequence boundary is locally at the base of thin channeled s<strong>and</strong>stones<br />
<strong>and</strong> red beds (Fig. 6, sections 1, 2). Farther downdip, the<br />
sequence boundary is placed beneath a thin local shale <strong>and</strong> a<br />
widespread unit of calcareous siltstone <strong>and</strong> minor s<strong>and</strong>stone<br />
(Figs. 6, 7) <strong>and</strong> beneath a thin oolite on the ramp slope (Fig. 6,<br />
section 11). This thin ooid grainstone is the lowst<strong>and</strong> systems<br />
tract along the ramp slope. Updip, the lowst<strong>and</strong> to transgressive<br />
systems tract consists of calcareous siltstone <strong>and</strong> shale (Fig. 6,<br />
sections 5 to 9; Fig. 7, sections 15 to 26) <strong>and</strong> in the far north, the thin<br />
channeled s<strong>and</strong>stones capped by red beds (mid-Union red beds)<br />
(Fig. 6, sections 1, 2). The transgressive systems tract on the ramp<br />
slope is a thin skeletal wackestone above the lowst<strong>and</strong> oolite (Fig.<br />
6, section 11), but a distinct transgressive systems tract is not<br />
easily recognized updip. The highst<strong>and</strong> systems tract is an upward-shallowing<br />
succession of deeper-water shaly lime mudstone<br />
to skeletal wackestone on the lower ramp slope (Fig. 6,<br />
section 11; Fig. 7, section 26). These pass upslope to the north into<br />
extensive ramp-margin skeletal grainstone (Fig. 6, sections 7 to 9;<br />
Fig. 7, sections 24, 25). These ramp-margin grainstones pass<br />
l<strong>and</strong>ward into skeletal wackestone or lime wackestone–mudstone,<br />
<strong>and</strong> then into updip skeletal grainstones <strong>and</strong> ooid grainstone<br />
(Fig. 6, sections 2 to 5; Fig. 7, sections 20 to 22) <strong>and</strong> on the<br />
innermost ramp, lime wackestone–mudstone.<br />
AUS AL-TAWIL, THOMAS C. WYNN, AND J. FRED READ<br />
Sequence C10.—Sequence C10, 10 to 60 m thick, consists of the<br />
Greenville Shale <strong>and</strong> Alderson Limestone (Figs. 6, 7). The C10<br />
basal sequence boundary is placed beneath a widespread thin<br />
shale or beneath calcareous siltstone <strong>and</strong> s<strong>and</strong>stone (Figs. 6, 7).<br />
The lowst<strong>and</strong> systems tract along the ramp slope is a thin s<strong>and</strong>stone–siltstone<br />
(Fig. 6, section 11; Fig. 7, sections 24 to 26). The<br />
lowst<strong>and</strong> to transgressive systems tract farther updip probably<br />
includes the basal shale <strong>and</strong> the calcareous siltstone unit (Figs. 6,<br />
7). The highst<strong>and</strong> systems tract consists of skeletal wackestone<br />
downdip, passing updip into ramp-margin skeletal grainstone–<br />
packstone (Fig. 6, sections 7 to 9; Fig. 7, sections 25, 26). These<br />
grade updip to the northeast into thin ooid grainstone units <strong>and</strong><br />
widespread lime wackestone–mudstone (Fig. 6, sections 2 to 6) or<br />
in the west, into thick ooid–peloid grainstone <strong>and</strong> then into lime<br />
wackestone–mudstone <strong>and</strong> local units of skeletal grainstone–<br />
packstone (Fig. 7, sections 12 to 22).<br />
Sequence C11.—Sequence C11 includes the Lillydale Shale <strong>and</strong><br />
the Glenray Limestone <strong>and</strong> is 0 to over 100 m thick thickening<br />
rapidly across the hinge (Fig. 6, sections 9 to 11; Fig. 7, sections 21<br />
to 22). The C11 sequence boundary is placed beneath thin, locally<br />
developed quartz s<strong>and</strong>stone–siltstones (Fig. 6, sections 1, 6, 8; Fig.<br />
7, sections 17 to 20), or beneath thin shales, red beds, or eolianites<br />
(Fig. 7, sections 12 to 16), or at the erosional contact on C10<br />
<strong>carbonate</strong>s at section 2 (Fig. 6). In the west, the sequence boundary<br />
on the ramp may be locally incised, indicated by rapid<br />
thickening of the siltstone at the expense of underlying units (Fig.<br />
6, section 11; Fig. 7, section 18). The lowst<strong>and</strong> systems tract on the<br />
ramp slope is a cross-bedded lithoclastic s<strong>and</strong>stone (Fido S<strong>and</strong>stone),<br />
which appears to be incised into the underlying unit (Fig.<br />
6, section 11). The transgressive systems tract is not defined<br />
downdip, but updip the lowst<strong>and</strong> to transgressive systems tract<br />
consists of thin, local s<strong>and</strong>stone–siltstones (which thicken locally<br />
into the possible incised valley at section 18, Fig. 7), <strong>and</strong> an<br />
overlying thin regional oolitic <strong>and</strong> skeletal limestone (Figs. 6, 7).<br />
The highst<strong>and</strong> systems tract includes 10 to 40 m of fossiliferous<br />
marine shale (Lillydale Shale) that grades basinward into a rampmargin<br />
complex of skeletal wackestone, skeletal grainstone, <strong>and</strong><br />
interbedded s<strong>and</strong>stone–siltstone (Fig. 7, section 26) <strong>and</strong> then into<br />
basinal lime mudstone up to 80 m thick (Fig. 6, section 11). These<br />
are overlain by late highst<strong>and</strong>, thin, discontinuous skeletal limestone<br />
capped locally by ooid grainstone (Fig. 6, sections 1, 3, 6, 8,<br />
9, 11; Fig. 7, sections 22 to 23). The top sequence boundary of<br />
sequence C11 in outcrop is a paleosol updip <strong>and</strong>/or erosional<br />
surface updip (Fig. 6, sections 1, 3), but on the ramp slope (Fig. 6,<br />
section 11) the sequence boundary underlies a thin black shale<br />
capped by ooid grainstone that is conformable on C11 skeletal<br />
wackestone. Sequence C11 is overlain by highly progradational<br />
red <strong>and</strong> gray siliciclastics (Mauch Chunk Formation).<br />
Parasequences<br />
Fourth-order sequences typically have only one or two<br />
parasequences updip, but downdip, sequences commonly are<br />
composed of several parasequences. Detailed facies successions<br />
of parasequences <strong>and</strong> presence or absence of erosional disconformities<br />
capping units are best observed in the outcrop sections or<br />
core. No attempt has been made to show parasequences in<br />
Figures 6 <strong>and</strong> 7. The parasequences are difficult to recognize in<br />
the cuttings data in sections in which the parasequences are thin<br />
<strong>and</strong> of the same magnitude as the sample spacing of the cuttings.<br />
Siliciclastic-prone parasequences in updip areas commonly<br />
consist of marine s<strong>and</strong>stone, locally with a slightly channeled<br />
base, fining upward into siltstone <strong>and</strong> then into red mudrock.<br />
Others include thin ooid or peloid grainstone overlain by red or
gray siltstone <strong>and</strong> then red mudrock. Siliciclastic units commonly<br />
make up the lower two-thirds of parasequences at the bases of<br />
sequences. Disconformity-bounded parasequences (2 to 6 m thick)<br />
are recognizable even in thick successions of restricted lime<br />
wackestone–mudstone (as in C7) <strong>and</strong> are composed of very thin<br />
(or absent) basal grainstones overlain by wackestone–mudstone,<br />
in which layering increases in abundance <strong>and</strong> becomes thinner<br />
upward into the laminated caps, where present.<br />
Oolite-dominated sequences include asymmetrical, upwardshallowing<br />
units (3 to 10 m thick) bounded at the base or top by<br />
quartz–peloid eolianites, <strong>and</strong> are composed of skeletal wackestone<br />
(commonly absent), up into ooid grainstone (with rare channeled<br />
bases <strong>and</strong> local mounds), which may be capped by argillaceous<br />
dolomitic wackestone–mudstone or dolomitic mudstone. Erosional<br />
disconformities may occur on the top of the lime<br />
wackestone–mudstone units, or on the ooid grainstone units.<br />
Symmetrical upward-deepening to upward-shallowing oolitic<br />
parasequences contain basal lime wackestone–mudstone conformably<br />
overlain by the upward-shallowing succession.<br />
Parasequences in skeletal limestone–rich parts of sequences<br />
(sequences C8 to C10) commonly are 2 to 15 m thick, <strong>and</strong><br />
generally are bounded by erosional or sharp planar contacts.<br />
Such parasequences include asymmetrical, upward-shallowing<br />
parasequences composed of thick skeletal wackestone–<br />
packstone grading up into ooid grainstone (may be absent),<br />
capped by dolomitic wackestone–mudstone. Other skeletal limestone–dominated<br />
parasequences are symmetrical <strong>and</strong> contain<br />
basal, thin transgressive ooid or skeletal grainstone, or basal<br />
upward-deepening units of lime wackestone–mudstone,<br />
interbedded with thin grainstones. Parasequences that straddle<br />
the maximum flooding surfaces of sequences commonly are<br />
thick <strong>and</strong> symmetrical.<br />
Third-Order or Composite Sequences<br />
Third-order or composite sequences (Weber et al., 1995) are<br />
relatively weakly developed in the study interval compared to<br />
the regionally mappable fourth-order sequences. The third-order<br />
sequences (Figs. 6, 7) are bounded by red beds <strong>and</strong> their downdip<br />
marine s<strong>and</strong>stones updip. They thicken downdip into the<br />
foredeep, where the third-order sequences are bounded along the<br />
ramp margin <strong>and</strong> ramp slope by extensive lowst<strong>and</strong> siliciclastic<br />
units. The third-order sequences commonly contain two or more<br />
high-frequency sequences.<br />
Composite Sequence 1.—Composite sequence 1 is composed of<br />
the Meramecian Little Valley–Hillsdale interval. It extends from<br />
the top of the Price–Maccrady red beds, up to the red beds of<br />
sequence C1 <strong>and</strong> its downdip equivalents (Figs. 4, 6, 7). It onlaps<br />
much of the ramp, pinching out to a feather edge updip. The<br />
Hillsdale composite sequence thickens into the foredeep to 270 m<br />
(Bartlett <strong>and</strong> Webb, 1971; Al-Tawil, 1998), where it may consist of<br />
sequences dominated by shallow-water, fine-grained <strong>carbonate</strong>s<br />
<strong>and</strong> interbedded s<strong>and</strong>stone (not shown on cross sections in Figs.<br />
6, 7).<br />
Composite Sequence 2.—This composite sequence includes sequences<br />
C1 to 4, which make up the Denmar Limestone. It<br />
extends from the base of the basal red beds (<strong>and</strong> their downdip<br />
equivalents) of sequence C1, up to the base of the lower Taggard<br />
red beds (<strong>and</strong> downdip equivalents) of sequence C5 (Figs. 4, 6, 7).<br />
The lowst<strong>and</strong> systems tract is made up of the lowst<strong>and</strong> siliciclastics<br />
of sequence C1. The transgressive systems tract probably includes<br />
the rest of sequences C1 to C3. The maximum flooding<br />
surface is likely at the base of sequence C4 updip, <strong>and</strong> beneath the<br />
MISSISSIPPIAN RAMP, WEST VIRGINIA, U.S.A.<br />
C4 deeper-water mudstones downdip. The highst<strong>and</strong> systems<br />
tract is made up of sequence C4.<br />
25<br />
Composite Sequence 3.—Sequences C5 to 10 may be the third<br />
composite sequence, extending from the basal C5 red beds up to<br />
the top of C10 (Figs. 4, 6, 7). It is composed of the Pickaway, Union,<br />
<strong>and</strong> Alderson Limestones. The lowst<strong>and</strong> systems tract of the<br />
composite sequence consists of sequences C5 <strong>and</strong> C6, which are<br />
red-bed prone, <strong>and</strong> show a major basinward shift in both distribution<br />
<strong>and</strong> facies. The transgressive systems tract sequence of the<br />
composite consists of sequence C7 <strong>and</strong> the lower part of sequence<br />
C8. The maximum flooding surface is marked by the most l<strong>and</strong>ward<br />
extent of open-marine skeletal wackestone in sequence C8.<br />
There is a significant development of siliciclastics at the base of<br />
C9, including an updip tongue of red beds (Fig. 6, sections 1, 2).<br />
The highst<strong>and</strong> systems tract is made up of the upper half of<br />
sequences C8, C9, <strong>and</strong> C10.<br />
Composite Sequence 4.—This composite sequence likely extends<br />
from the base of the incised Fido S<strong>and</strong>stone–Lillydale Shale <strong>and</strong><br />
includes sequence C11 (Lillydale–Glenray interval); the top of the<br />
composite sequence is above the study interval (Figs. 4, 6, 7). The<br />
lowst<strong>and</strong> systems tract may include the incised siliciclastic units at<br />
the base of C11, <strong>and</strong> the maximum flooding surface is above the<br />
thin limestones at the base of C11 <strong>and</strong> below the fossiliferous<br />
shales. The highst<strong>and</strong> systems tract probably includes the Lillydale<br />
Shale, the Glenray Limestone, the Bickford Shale, <strong>and</strong> the Reynolds<br />
Limestone, <strong>and</strong> possibly units above the study interval.<br />
CONTROLS ON SEQUENCE DEVELOPMENT<br />
ON THE FORELAND<br />
Tectonics <strong>and</strong> Subsidence Rates<br />
Accommodation rates calculated from the present-day stratigraphic<br />
thicknesses (not decompacted) were highest (at least 10<br />
to 30 cm/ky) in the basin compared to the ramp margin (a<br />
minimum of 5 to 15 cm/ky) <strong>and</strong> were lowest (at least 1 to 3 cm/<br />
ky) in updip areas of the distal forel<strong>and</strong>. This promoted regional<br />
slopes into the forel<strong>and</strong> basin. Regional slopes per sequence can<br />
be calculated using the difference in thickness of the study<br />
interval from updip to downdip, <strong>and</strong> dividing this by the distance<br />
times the number of sequences. This gives slopes of at least<br />
5 cm/km/sequence on the distal forel<strong>and</strong> (Al-Tawil <strong>and</strong> Read,<br />
this volume) <strong>and</strong> average slopes of at least 20 cm/km/sequence<br />
down the plunging basin axis. The low distal forel<strong>and</strong> slopes<br />
resulted in the predominance of shallow-water facies over the<br />
whole of the distal forel<strong>and</strong>. Calculated slopes per sequence<br />
measured over the palinspastically restored dip direction (northwest<br />
to southeast) down the ramp margin <strong>and</strong> into the basin were<br />
higher, at least 30 cm/km. These calculated slopes are compatible<br />
with minimum estimated highst<strong>and</strong> water depths of at least 50 m<br />
on the ramp slope, which periodically shallowed to sea level<br />
during lowst<strong>and</strong>s.<br />
It has been suggested that the region was tectonically quiescent<br />
during deposition of the study interval, <strong>and</strong> in a state of<br />
elevational equilibrium between the filled forel<strong>and</strong> basin <strong>and</strong><br />
the beveled Devonian orogenic highl<strong>and</strong>s (Ettensohn, 1994).<br />
However, this appears to be incompatible with the considerable<br />
differential subsidence shown by the thickening of the <strong>carbonate</strong>s<br />
from a feather edge updip on the forel<strong>and</strong> to over 900 m in<br />
the basin (Figs. 6, 7). The thickening takes place not only across<br />
the regional hinge but also at various localities updip <strong>and</strong><br />
downdip from the hinge at various times. This complex pattern<br />
may have resulted from the movement of basement fault blocks
26<br />
beneath the forel<strong>and</strong>, which parallel the orogenic front<br />
(Shumaker <strong>and</strong> Wilson, 1996). However, given the long wavelengths<br />
of deformation, it also could just relate to flexural<br />
downwarping. Such downwarping in the forel<strong>and</strong> basin over<br />
the Price delta system occurred too long (some 10 to 15 My) after<br />
the initial early Mississippian loading event (Ettensohn, 1994) to<br />
be related to relaxation. This suggests that accommodation<br />
space might have been initiated during the Meramecian <strong>and</strong><br />
early Chesterian by renewed thrust loading followed by relaxation<br />
(cf. Quinlan <strong>and</strong> Beaumont, 1984; Al-Tawil <strong>and</strong> Read, this<br />
volume). This tectonic loading <strong>and</strong> differential subsidence during<br />
deposition of the Meramecian sequences <strong>and</strong> Chesterian<br />
sequences C1 to C6 caused rapid thickness changes (Ettensohn,<br />
1975, 1980; MacQuown <strong>and</strong> Pear, 1983; Dever, 1986, Dever et al.,<br />
1990; Shumaker <strong>and</strong> Wilson, 1996). In addition, there was a<br />
northwestward migration (approximately 50 km) of the major<br />
hinge onto the distal forel<strong>and</strong> as well as a northeastward migration<br />
of smaller hinge zones during the Meramecian–Chesterian<br />
(Figs. 6, 7). This would seem to support continued advance of a<br />
thrust (or sediment) load onto the forel<strong>and</strong>, suggesting that<br />
Greenbrier deposition occurred during active tectonics rather<br />
than during a time of tectonic quiescence.<br />
The regional subsidence that initiated Chesterian deposition,<br />
coupled with long-term sea-level rise (Ettensohn, 1994), moved<br />
the siliciclastic shorelines far updip during fourth-order<br />
highst<strong>and</strong>s, resulting in widespread <strong>carbonate</strong> deposition on the<br />
ramp <strong>and</strong> siliciclastic deposition only during fourth-order<br />
lowst<strong>and</strong>s/early transgressions. The marked differential subsidence<br />
exhibited by the Meramecian sequence <strong>and</strong> Chesterian<br />
sequences C1 to C6 (Figs. 6, 7) contrasts with the more gradual<br />
thickness changes of the upper sequences C7 to C10, which<br />
subsided more as a result of regional flexuring, possibly as<br />
tectonic loading slowed <strong>and</strong>/or regional uplift by erosional unloading<br />
of the orogenic highl<strong>and</strong>s was initiated. Perhaps the<br />
earlier faults on the forel<strong>and</strong> had locked by this time <strong>and</strong> the<br />
forel<strong>and</strong> underwent regional flexure. Rapid flexure over the<br />
hinge also characterizes sequence C11 (Figs. 6, 7) <strong>and</strong> sequence<br />
C12 (above the study interval) <strong>and</strong> may mark renewed loading of<br />
the forel<strong>and</strong>. Following deposition of the Chesterian <strong>carbonate</strong>s,<br />
erosional unloading <strong>and</strong> uplift of the orogenic highl<strong>and</strong>s caused<br />
the subsequent progradation of Upper Mississippian Mauch<br />
Chunk siliciclastic units (Ettensohn, 1994).<br />
Differential subsidence thus controlled thickness <strong>and</strong> facies<br />
variations across the ramp, not only across the major ramp<br />
margin hinge but also at various places on the distal forel<strong>and</strong>, <strong>and</strong><br />
thus was an important control on the accommodation created<br />
during deposition of each sequence. Broad sags periodically<br />
developed on the forel<strong>and</strong> in the northeast (sequences C2, C6, C7;<br />
Fig. 6) <strong>and</strong> locally on the distal forel<strong>and</strong> in the west (as in<br />
Meramecian sequence 2, <strong>and</strong> sequences C2, C3, <strong>and</strong> C9; Fig. 7).<br />
These sags <strong>and</strong> adjacent highs may have influenced the position<br />
of updip shoal-water grainstones separated by lagoonal lime<br />
wackestone–mudstone from the ramp-margin grainstones along<br />
the hinge. Although the differential subsidence modified the<br />
magnitudes of the relative sea-level changes, it did not override<br />
the eustatic changes that formed the sequence-bounding,<br />
siliciclastic-prone units <strong>and</strong> the <strong>carbonate</strong>-prone transgressive<br />
<strong>and</strong> highst<strong>and</strong> units.<br />
Climate<br />
Climate Control on Facies.—During the early Mississippian,<br />
the Appalachian Basin had a humid climate, which favored<br />
siliciclastic deposition on the Price–Borden delta complex<br />
(Branson, 1912; Butts, 1940, 1941; Sable <strong>and</strong> Dever, 1990). This<br />
AUS AL-TAWIL, THOMAS C. WYNN, AND J. FRED READ<br />
climate subsequently became more arid during deposition of<br />
the Maccrady red beds <strong>and</strong> evaporites. Evaporites in the<br />
Meramecian interval in Kentucky (southwest of the study area)<br />
indicate that arid <strong>and</strong> semiarid conditions persisted throughout<br />
the middle Mississippian (Sable <strong>and</strong> Dever, 1990; Khetani <strong>and</strong><br />
Read, 2002). This semiarid climate persisted into the early<br />
Chesterian during deposition of the lower part of the study<br />
interval (indicated by red beds, widespread oolitic ramps, coastal<br />
eolianites, <strong>and</strong> caliche in West Virginia, Kentucky, <strong>and</strong> the<br />
Illinois Basin (Dever, 1973, 1995; Ettensohn et al., 1988; Smith<br />
<strong>and</strong> Read, 1999). However, the climate may have become more<br />
humid into the later Chesterian, when disconformities in the<br />
Appalachian <strong>and</strong> Illinois basins became characterized by<br />
paleokarst (Ettensohn et al., 1988), coarse siliciclastic influx<br />
increased (Swann, 1964; Smith, 1996), skeletal limestones became<br />
more widespread, <strong>and</strong> thick oolitic deposits (which on<br />
ramps require semi arid conditions; Read, 1985) became much<br />
less common. Also, regional meteoric aquifers developed during<br />
this late Mississippian–Early Pennsylvanian humid phase<br />
(Niemann <strong>and</strong> Read, 1988). These climatic changes are borne<br />
out by paleogeographic reconstructions (de Witt <strong>and</strong> McGrew,<br />
1979; Scotese <strong>and</strong> McKerrow, 1990) that show the North American<br />
plate gradually rotating counterclockwise <strong>and</strong> moving northward<br />
from the Pangean arid sub-equatorial belt during the<br />
Meramecian to Chesterian (latitudes of 25 degrees or less),<br />
reaching close to the Equator by Pennsylvanian time.<br />
By latest Mississippian time, semiarid climate returned at<br />
least intermittently with the deposition of Mauch Chunk red beds<br />
<strong>and</strong> semiarid paleosols (Caudill et al., 1996).<br />
Climate Control on Long-Term Global Sea Levels.—Oxygen<br />
stable-isotope data from the mid-continent United States show<br />
several major climate events in the Meramecian–Chesterian,<br />
with increased δ 18 O values being interpreted as cooling <strong>and</strong>/or<br />
ice buildup events <strong>and</strong> decreased δ 18 O values representing<br />
warming events (Mii et al., 1999). There is a mid-Meramecian<br />
cooling event that likely marks the regressive event at the base<br />
of the Upper Meramecian sequences. This was followed by<br />
warming in the early Chesterian, which promoted transgression<br />
<strong>and</strong> deposition of the Meramecian sequences <strong>and</strong> Chesterian<br />
sequences C1 to C4. This was followed by a cooling event that<br />
likely corresponds to deposition of the red beds of sequences C5<br />
<strong>and</strong> C6 (Taggard red beds). Subsequent warming probably<br />
caused the long-term transgression <strong>and</strong> deposition of sequences<br />
C7 through C10. The subsequent cooling trend (Mii et al., 1999;<br />
Veevers <strong>and</strong> Powell, 1987) <strong>and</strong> falling sea levels were accompanied<br />
by progradation of Mauch Chunk siliciclastic-prone sequences<br />
(C11 <strong>and</strong> younger units).<br />
Eustasy<br />
Evidence for Fourth-Order Eustasy.—A set of relative-sea-level<br />
curves for the basin (Fig. 8) were constructed by arbitrarily<br />
assigning equal time to each of the sequences <strong>and</strong> plotting the<br />
relative magnitude of updip–downdip migration of marker facies<br />
such as oolitic or skeletal grainstones <strong>and</strong> s<strong>and</strong>stones. The<br />
resultant curves were compared to the global sea-level curves of<br />
Ross <strong>and</strong> Ross (1987) <strong>and</strong> a simplified Illinois Basin curve modified<br />
from Smith (1996) <strong>and</strong> Smith <strong>and</strong> Read (2000) from which the<br />
higher-frequency fluctuations were removed. In addition to the<br />
biostratigraphic control based on conodonts <strong>and</strong> crinoids<br />
(Collinson et al., 1971; Huggins, 1983; deWitt <strong>and</strong> McGrew, 1979;<br />
Sable <strong>and</strong> Dever, 1990), correlation between the Appalachian <strong>and</strong><br />
the Illinois basins was aided by the presence of relatively regional<br />
unconformities <strong>and</strong> siliciclastic markers in Kentucky (McFarlan
VISEAN NAMURIAN<br />
CHESTERIAN<br />
MERAMECIAN<br />
ST. LOUIS<br />
<strong>and</strong> Walker, 1956) <strong>and</strong> in West Virginia; such markers include the<br />
Taggard red beds, <strong>and</strong> the Greenville <strong>and</strong> Lillydale shales (Smith<br />
<strong>and</strong> Read, 1999, 2000, 2001).<br />
The Upper Meramecian sequences in the proximal forel<strong>and</strong><br />
correlate with the upper <strong>and</strong> lower St. Louis Limestone of the<br />
Illinois Basin (Huggins, 1983) but only one of the Upper<br />
Meramecian Hillsdale sequences appears to reach the distal<br />
forel<strong>and</strong> in Kentucky (Al-Tawil <strong>and</strong> Read, this volume).<br />
Most of the Chesterian fourth-order sequences in eastern<br />
West Virginia can be traced into much thinner disconformitybounded<br />
sequences exposed on the distal forel<strong>and</strong>, eastern Kentucky<br />
(Al-Tawil, 1998; Al-Tawil <strong>and</strong> Read, this volume), <strong>and</strong> into<br />
the Illinois Basin (Ettensohn et al., 1984; Smith, 1996; Smith <strong>and</strong><br />
Read, 1999; 2001) (Fig. 8).<br />
The four West Virginia–Virginia sequences (C1 to C4) in the<br />
Playtycrinites penicillus–bearing Denmar interval (G. bilineatus–C.<br />
charactus conodont zone; Huggins, 1983; Stamm, 1997) can be<br />
recognized in the time-equivalent Ste. Genevieve Limestone,<br />
exposed on the distal forel<strong>and</strong> in Kentucky. However, they are<br />
difficult to match with the three Ste. Genevieve sequences in the<br />
Illinois Basin, although four sequences may occur there if the two<br />
disconformity-bounded units in Smith <strong>and</strong> Read’s (1999) sequence<br />
C3 are raised to sequence status.<br />
West Virginia sequence C5 (containing the lower Taggard<br />
red beds), sequences C6 (bearing the upper Taggard red beds–<br />
lower Pickaway) <strong>and</strong> C7 of West Virginia (upper Pickaway<br />
interval), which contain the first occurrences of the crinoid<br />
MISSISSIPPIAN RAMP, WEST VIRGINIA, U.S.A.<br />
ST. LOUIS<br />
STE. GENEVIEVE<br />
APPALACHIAN BASIN<br />
GLOBAL SEA LEVEL<br />
ROSS AND ROSS (1987)<br />
ILLINOIS BASIN<br />
SMITH (1996)<br />
KENTUCKY<br />
DISTAL FORELAND<br />
(AL-TAWIL AND READ,<br />
THIS VOLUME)<br />
WEST VIRGINIA<br />
PROXIMAL FORELAND<br />
(THIS STUDY)<br />
HIGH LOW HIGH LOW HIGH LOW HIGH LOW<br />
GLEN DEAN<br />
GLEN DEAN 10<br />
GLEN DEAN<br />
HARDINSBURG SS.<br />
C11<br />
C11 GLEN RAY<br />
HANEY<br />
HANEY 9<br />
HANEY C10<br />
C10 ALDERSON<br />
BEECH CK<br />
BEECH CK 8<br />
CYPRESS SS<br />
BEECH CK C9<br />
CAVE BRANCH SH.<br />
C9 UPPER UNION<br />
REELSVILLE 7<br />
SAMPLE SS.<br />
BEAVER BEND 6<br />
BETHEL SS.<br />
REELSV. C8B<br />
C8A<br />
BEAVER<br />
BEND<br />
C8B<br />
C8A<br />
LOWER<br />
UNION<br />
5<br />
C7<br />
C7<br />
PAOLI<br />
PAOLI<br />
C6<br />
C6<br />
4<br />
BRYANTSVILLE BRECCIA<br />
3 LEVIAS AUX<br />
VASE<br />
C5<br />
C4<br />
C5<br />
C4<br />
U. TAGGARD<br />
L. TAGGARD<br />
STE. GENEVIEVE<br />
2<br />
SPAR MT.<br />
SS<br />
C3<br />
STE. GENEVIEVE<br />
C3<br />
DENMAR<br />
1<br />
C2<br />
C2<br />
PAOLI /<br />
WARIX RUN<br />
C1<br />
ST. LOUIS<br />
HILLSDALE<br />
FIG. 8.—Relative third-order <strong>and</strong> fourth-order sea-level curves for the study interval in the West Virginia–Virginia Appalachians<br />
compared with curves from the distal Appalachian forel<strong>and</strong> in Kentucky (Al-Tawil <strong>and</strong> Read, this volume), the Illinois Basin<br />
(Smith, 1999; 2000; 2001), <strong>and</strong> the third-order curves of Ross <strong>and</strong> Ross (1987). Most of the fourth-order cycles can be correlated<br />
between the basins. The Ross <strong>and</strong> Ross third-order curves can be tied grossly to the Appalachian third-order cycles. The Illinois<br />
Basin curves show some additional large sea-level falls defined on the degree of incision, which are not recognized in the<br />
Appalachian Basin.<br />
27<br />
Talarocrinus, are correlated with the three sequences mapped as<br />
“Warix Run”/Paoli in Kentucky by Ettensohn et al. (1984), <strong>and</strong><br />
the Talarocrinus-bearing Paoli sequences in the Illinois Basin<br />
(Butts, 1922, 1940, 1941; Wells, 1950; Ettensohn et al., 1984; Smith<br />
<strong>and</strong> Read, 1999).<br />
West Virginia sequence C8 (Lower Union) corresponds to the<br />
Beaver Bend–Reelsville cycle in Kentucky <strong>and</strong> in the Illinois Basin<br />
(de Witt <strong>and</strong> McGrew, 1979; Ettensohn et al., 1984; Smith <strong>and</strong><br />
Read, 2001; Al-Tawil <strong>and</strong> Read, this volume). Sequence C9 of<br />
West Virginia (mid-Union red beds <strong>and</strong> Upper Union Limestone),<br />
containing G. bilineatus–C. altus conodonts, was correlated<br />
with the Reelsville limestone of the Illinois Basin (de Witt <strong>and</strong><br />
McGrew, 1979), but it more likely correlates with the Cave Branch<br />
Shale–Beech Creek sequence in Kentucky <strong>and</strong> the Cypress S<strong>and</strong>stone–Beech<br />
Creek–Big Clifty sequence (Illinois Basin), containing<br />
basal G. bilineatus–C. altus zonal conodonts (Collinson et al.,<br />
1971; Ettensohn et al., 1984).<br />
West Virginia sequence C10 (Greenville Shale–Alderson Limestone),<br />
containing the up<strong>permo</strong>st occurrence of Talarocrinus (Butts,<br />
1940; Wells, 1950), correlates with the Haney sequence of Kentucky<br />
<strong>and</strong> the Illinois Basin (de Witt <strong>and</strong> McGrew, 1979; Ettensohn<br />
et al., 1984; Smith <strong>and</strong> Read, 2001; Al-Tawil <strong>and</strong> Read, this<br />
volume). West Virginia sequence C11 (Lillydale Shale/Fido S<strong>and</strong>stone–Glenray<br />
Limestone), which contains the G. bilineatus–K.<br />
mehli zonal conodonts (Rexroad <strong>and</strong> Clark, 1960) correlates with<br />
the Hardinsburg–Glen Dean cycle of Kentucky <strong>and</strong> the Illinois<br />
Basin (de Witt <strong>and</strong> McGrew, 1979; Ettensohn et al., 1984).<br />
C1<br />
PICKAWAY
28<br />
The interbasinal extent of these sequences suggests that global<br />
eustasy played a major role in their development, rather than<br />
tectonics alone, which would operate locally as the basins subsided<br />
differentially. Although better age constraints are needed,<br />
the high-frequency sequences appear to relate to fourth-order<br />
sea-level cycles. If with better dating these do prove to be 400 ky,<br />
then they may relate to Milankovitch-driven, long-term eccentricity,<br />
which would be expected to predominate during times of<br />
global ice sheets (Imbrie <strong>and</strong> Imbrie, 1986; Ruddiman et al., 1986).<br />
This is compatible with the onset of late Paleozoic Gondwana<br />
glaciation in the Southern Hemisphere, in the Early Chesterian or<br />
even earlier (Horbury, 1989; Frakes et al., 1992; Smith <strong>and</strong> Dorobek,<br />
1993; Wright <strong>and</strong> Vanstone, 2001).<br />
The data from the Illinois <strong>and</strong> Appalachian basins suggest<br />
that there was an increase in the magnitude of the sea-level<br />
changes through the Chesterian, peaking in the Pennsylvanian<br />
with sea-level changes of up to 100 m (Heckel, 1986). However,<br />
the presence of significant incision on some of the basal Chesterian<br />
units in the Illinois Basin may suggest that there were some large<br />
sea-level falls even this early (Leetaru, 2000; Smith <strong>and</strong> Read,<br />
2000, 2001) <strong>and</strong> possibly earlier in the Meramecian (Smith <strong>and</strong><br />
Dorobek, 1993; Wright <strong>and</strong> Vanstone, 2001; Horbury, 1989) (Fig.<br />
8). This upward increase in magnitude of Late Mississippian<br />
fourth-order sea-level cycles was superimposed on the long-term<br />
Chesterian transgression, resulting in progressive l<strong>and</strong>ward shift<br />
of fourth-order highst<strong>and</strong>s. However, the increasing sea-level<br />
falls kept returning the lowst<strong>and</strong> s<strong>and</strong> bodies to the general<br />
vicinity of the ramp margin <strong>and</strong> slope, even though long-term<br />
transgression was occurring.<br />
Evidence for Third-Order Eustasy.—Some of the weak, composite<br />
(third-order) sequences appear to correspond to third-order<br />
sea-level cycles documented as global events by Ross <strong>and</strong> Ross<br />
(1987) (Fig. 8). The third-order cycle that formed the Little Valley–<br />
Hillsdale composite sequence correlates with the Late Meramecian<br />
(St. Louis) sea-level cycle (Huggins, 1983). The third-order cycle<br />
that formed the Denmar composite sequence correlates with the<br />
Ste. Genevieve (earliest Chesterian) sea-level cycle, with a major<br />
lowst<strong>and</strong> at the position of the Appalachian Basin sequences C5<br />
<strong>and</strong> C6 red beds (lower <strong>and</strong> upper Taggard red beds). However,<br />
the later composite sequences are difficult to relate to the global<br />
cycles, in part because of insufficient detail in the Ross <strong>and</strong> Ross<br />
(1987) cycle chart. However, they do show a third-order sea-level<br />
cycle extending from the top of the Ste. Genevieve cycle up to the<br />
top of the Glen Dean (top of our C11), whereas we placed the top<br />
of the cycle at the base of C11 (base of Glenray).<br />
There also are substantial differences between the major<br />
incisional events in the Appalachian <strong>and</strong> Illinois basins. In the<br />
Illinois Basin, Leetaru (2000) recognized major incision in the<br />
middle <strong>and</strong> near the top of the Ste. Genevieve <strong>carbonate</strong>s, whereas<br />
we see major lowst<strong>and</strong>s only above the equivalent (Denmar)<br />
<strong>carbonate</strong>s in the Appalachian Basin (at bases of C5 <strong>and</strong> C6).<br />
Incisional events in the Illinois Basin (Smith <strong>and</strong> Read, 2000, 2001)<br />
for which there is little evidence in the Appalachian Basin also<br />
include the Bethel S<strong>and</strong>stone incision (base of our C8 sequence),<br />
<strong>and</strong> the incision at the base of the Cypress S<strong>and</strong>stone, Illinois<br />
Basin (base of our C9), although there is a red-bed tongue at this<br />
level updip in the Appalachian Basin. Incision at the base of the<br />
Hardinsburg S<strong>and</strong>stone, in the Illinois Basin, corresponds to the<br />
incision observed at the base of our sequence C11 (Fido S<strong>and</strong>stone).<br />
The differences between the amount of incision between<br />
the Appalachian <strong>and</strong> the Illinois basins (Smith <strong>and</strong> Read, 1999,<br />
2000, 2001) may reflect the effects of uplift or subsidence acting on<br />
weak third-order sea-level changes, coupled with siliciclastic<br />
influx <strong>and</strong> incision associated with significant fluvial systems in<br />
AUS AL-TAWIL, THOMAS C. WYNN, AND J. FRED READ<br />
the Illinois Basin, which may have accentuated the effects of some<br />
of the eustatic falls in the Illinois Basin. The greater subsidence<br />
rates at least in the proximal forel<strong>and</strong> of the Appalachian Basin<br />
may have tended to inhibit incision, as did the relative paucity of<br />
siliciclastic systems there.<br />
Development of Fourth-Order Sequences<br />
Development of Sequence Boundaries.—The disconformities <strong>and</strong><br />
associated caliches <strong>and</strong> breccias that mark sequence boundaries<br />
updip in Kentucky as well as in the Illinois Basin (Ettensohn et al.,<br />
1984; Ettensohn et al., 1988; Smith <strong>and</strong> Read, 1999, 2001; Al-Tawil<br />
<strong>and</strong> Read, this volume), appear to be less developed in the West<br />
Virginia study area, although this may be more apparent than<br />
real <strong>and</strong> may reflect the dominance of well data over outcrop data<br />
in this paper. However, it could also be due to the rapid mantling<br />
of exposed <strong>carbonate</strong>s by siliciclastic veneers (eolian or sheetwash<br />
deposition) that inhibited paleosol formation. Local,<br />
calichified <strong>and</strong> brecciated units downdip along the ramp margin<br />
may reflect the low gradient of the ramp top, which allowed it to<br />
be exposed during sea-level falls. Sharp, planar to hummocky<br />
erosion surfaces that are developed on some of the grainstones at<br />
parasequence <strong>and</strong> sequence boundaries in outcrops downdip<br />
could reflect widespread marine cementation followed by erosion<br />
in rock platform settings during transgression. Significant<br />
meteoric cementation beneath disconformities did not occur<br />
until later in the Mississippian, when the climate became more<br />
humid (Niemann <strong>and</strong> Read, 1988).<br />
Development of Lowst<strong>and</strong> to Early Transgressive Tracts.—Although<br />
regional 3-D data are needed, it appears that siliciclastics<br />
for the lowst<strong>and</strong>/early transgressive s<strong>and</strong> bodies were sourced<br />
from highl<strong>and</strong>s to the northeast <strong>and</strong> cratonic sources to the north,<br />
<strong>and</strong> they traveled southward toward the ramp margin (Fig. 1). It<br />
is not yet clear whether any were derived from the highl<strong>and</strong>s<br />
bordering the eastern side of the foredeep directly east of the<br />
study area, inasmuch as this either would have required very<br />
large sea-level falls to allow transport across the foredeep. The<br />
distribution of regional red beds in sequences C5 <strong>and</strong> C6 (Taggard<br />
red beds) suggests that siliciclastics from the northeast were<br />
funneled down the eastern side of the forel<strong>and</strong>, which may have<br />
been a subtle topographic low that deepened into the foredeep to<br />
the southwest. Sequences C5 <strong>and</strong> C6 red beds were not deposited<br />
on the topographically higher (?) distal forel<strong>and</strong> to the west,<br />
which was the site of eolian <strong>and</strong> early transgressive shallow<br />
marine siliciclastic deposition. Siliciclastic veneers at bases of the<br />
sequences on the distal forel<strong>and</strong> could have been transported<br />
onto the ramp during lowst<strong>and</strong>, by winds or flash floods, but<br />
most of the siliciclastic units probably were subsequently reworked<br />
into early transgressive marine sheetlike units. Quartz–<br />
peloid eolianites may have been preserved (as in the outcropping<br />
distal forel<strong>and</strong> deposits of Kentucky; Al-Tawil <strong>and</strong> Read, this<br />
volume), perhaps because of early incipient cementation, which<br />
inhibited reworking, but others could have been reworked into<br />
marine quartz–peloid grainstone sheets, which would be difficult<br />
to differentiate from eolian deposits in cuttings. Some of the<br />
lowst<strong>and</strong> to early transgressive s<strong>and</strong>stone–calcareous siltstones<br />
were deposited as barriers behind which thin, nearshore shales<br />
were deposited, as in sequences C7 to C10.<br />
Because of the poor constraints on the water depths of the<br />
facies, estimates of sea-level fall forming the lowst<strong>and</strong> deposits are<br />
crude. Where relative sea-level falls were small to moderate (say 10<br />
to 15 m or so), lowst<strong>and</strong> marine s<strong>and</strong>stones that are only slightly<br />
shallower than the underlying highst<strong>and</strong> facies were deposited,<br />
<strong>and</strong> there was only a small basinward shift in facies. These include
lowst<strong>and</strong>s at the bases of sequences C2, C7, <strong>and</strong> C8. Where relative<br />
sea-level falls were larger (say 20 to 40 m or so), then forced<br />
regression caused a significant basinward shift of facies, <strong>and</strong><br />
lowst<strong>and</strong> shallow-water ooid grainstones or marine s<strong>and</strong>stones<br />
were deposited on deeper-water skeletal wackestone or laminated<br />
shaly lime mudstone, with intervening facies not being deposited,<br />
as at the bases of sequences C1, C6, C9, C10, <strong>and</strong> C11.<br />
The localization of lowst<strong>and</strong> wedges on the ramp margin was<br />
controlled by high differential subsidence over the hinge <strong>and</strong> onto<br />
the proximal forel<strong>and</strong> during deposition of the lower sequences.<br />
During deposition of sequences C1 to C6 (Denmar <strong>and</strong> lower<br />
Pickaway sequences), where the underlying highst<strong>and</strong> skeletal<br />
bank facies were located close to the ramp margin, lowst<strong>and</strong> s<strong>and</strong>s<br />
migrated onto the ramp slope as sea level fell. However, during<br />
deposition of sequences C7 <strong>and</strong> C8, the ramp margin was farther<br />
updip, as a probable response to the long-term sea-level rise <strong>and</strong><br />
activation of hinges farther updip. Thus, sea level had to fall farther<br />
to expose the ramp margin <strong>and</strong> slope, so that lowst<strong>and</strong> siliciclastics<br />
reached only the ramp margin. However, the increasingly large<br />
sea-level falls during deposition of sequences C9 to C11, coupled<br />
with the broad flexuring of the ramp, rather than differential<br />
subsidence, could have returned deposition of lowst<strong>and</strong> quartz<br />
s<strong>and</strong>stone <strong>and</strong> oolite units to the ramp slope. The most well<br />
developed lowst<strong>and</strong> siliciclastic units appear to coincide with<br />
third-order sea-level falls, as in the base of sequence C1 (base of the<br />
Denmar), the bases of sequences C5 <strong>and</strong> C6 (Taggard red beds),<br />
<strong>and</strong> the base of sequence C11 (Fido S<strong>and</strong>stone).<br />
Transgressive Systems Tract Development <strong>and</strong> Maximum Flooding.—Although<br />
the maximum flooding surfaces of each of the<br />
sequences were difficult to pick at all locations, given the mix of<br />
subsurface well data <strong>and</strong> higher-quality outcrop data, some<br />
generalizations can be made. Along the ramp slope, the coincidence<br />
of maximum flooding surfaces with the transgressive<br />
surface on the lowst<strong>and</strong> systems tract (as in sequences C 4, C10<br />
<strong>and</strong> C11) appears to reflect rapid deepening <strong>and</strong> drowning of the<br />
lowst<strong>and</strong> facies as relative sea level rose <strong>and</strong> the ramp was<br />
transgressed. In contrast, in most of the other sequences, transgressive<br />
wackestone units, or transgressive–regressive skeletal<br />
wackestone to skeletal grainstone (C1, C3, C6) were developed on<br />
the lowst<strong>and</strong> deposits, reflecting more gradual relative-sea-level<br />
rise during fourth-order transgression. If the skeletal units on the<br />
ramp slope developed solely under the influence of gradually<br />
rising sea level, they would be expected to be upward deepening<br />
into the overlying deeper-water facies. However, the skeletal<br />
banks appear to be shallowing-upward successions <strong>and</strong> probably<br />
reflect superimposed higher-frequency sea-level fluctuations<br />
or perhaps a slowing of subsidence.<br />
On the shallow ramp, some fourth-order sea-level rises reworked<br />
basal siliciclastic units but did not locally deposit a<br />
transgressive succession of <strong>carbonate</strong>s, although such transgressive<br />
deposits formed within some of these same sequences elsewhere.<br />
Suppression of transgressive <strong>carbonate</strong> deposition could<br />
have been due to rapid rise rates that did not allow time for a<br />
significant transgressive <strong>carbonate</strong> to be deposited, or perhaps<br />
the shallow marine waters were too turbid with resuspended<br />
terrigenous silt. Deposition of a transgressive <strong>carbonate</strong> succession<br />
on the locally developed, eolian <strong>and</strong> marine reworked<br />
siliciclastic units could have been favored by slower rise of rates<br />
of relative sea level <strong>and</strong> waters becoming clearer as transgression<br />
proceeded. The presence of upward-shallowing units that make<br />
up the transgressive <strong>carbonate</strong>s in some sequences that have<br />
significant transgressive systems tracts (sequence C2, C7, <strong>and</strong> C8;<br />
Fig. 6), suggests that transgressive <strong>carbonate</strong> deposition may<br />
have been favored by a short-term rise <strong>and</strong> fall superimposed on<br />
MISSISSIPPIAN RAMP, WEST VIRGINIA, U.S.A.<br />
the fourth-order rise, which effectively delayed maximum flooding<br />
of the ramp.<br />
29<br />
Highst<strong>and</strong> Deposition.—During highst<strong>and</strong> deposition of the<br />
sequences, a complex facies mosaic developed, rather than a<br />
simple upward-shallowing, progradational succession that would<br />
result from seaward migration of shore-parallel facies belts.<br />
Skeletal grainstone–packstone <strong>and</strong> ooid grainstone were deposited<br />
not only along the ramp margin but also as extensive but<br />
discontinuous bodies far back from the margin, from which they<br />
were separated by extensive lagoonal (<strong>and</strong> locally tidal flat) lime<br />
wackestone–mudstone. Some of these updip skeletal <strong>and</strong> ooid<br />
bodies could have formed bordering highs adjacent to locally<br />
downwarped areas of the forel<strong>and</strong> (Carney, 1993), but others<br />
could have been localized by subtle depositional topography.<br />
Individual ramp-margin ooid bodies were shaped by tidal currents<br />
in a northeast–southwest direction normal to depositional<br />
strike (Kelleher <strong>and</strong> Smosna, 1993).<br />
Lagoonal lime wackestone–mudstones were widely deposited<br />
in protected lagoons behind ramp-margin ooid <strong>and</strong> skeletal<br />
grainstone shoals <strong>and</strong> far updip between <strong>and</strong> behind shoals. The<br />
absence of tidal-flat laminites from much of the distal forel<strong>and</strong><br />
probably was due to the relatively large sea-level falls, which<br />
caused rapid seaward shift of the shoreline, faster than tidal flats<br />
could prograde. However, locally thick tidal-flat facies were<br />
deposited just behind the ramp margin as the lagoonal muds<br />
shallowed upward, possibly because of high sedimentation rates<br />
here, coupled with the low rates of relative sea-level fall, which<br />
would characterize late highst<strong>and</strong>s, <strong>and</strong> the relatively high subsidence<br />
rates, which would have decreased the effects of eustatic<br />
fall.<br />
Development of Late Meramecian–Chesterian<br />
Composite Sequences<br />
In the study area, the Kinderhook–middle Osagean Borden–<br />
Price delta complex was not overlain by upper Osagean–lower<br />
Meramecian <strong>carbonate</strong>s of the Fort Payne–Salem supersequence,<br />
which were deposited farther to the south. Thus, the upper<br />
Meramecian–Chesterian supersequence in this paper was deposited<br />
directly onto the earlier Price–Borden siliciclastic units, following<br />
rapid differential subsidence of the delta in the east,<br />
coupled with global sea-level rise (Dennison, 1985; Ettensohn,<br />
1994), which led to increased accommodation <strong>and</strong> long-term<br />
onlap of the upper Meramecian <strong>and</strong> Chesterian sequences onto<br />
the delta complex.<br />
Differential subsidence over the eastern hinge, coupled with<br />
relatively small third-order sea-level fluctuations which started<br />
from the low sea-level position at the end of the upper Osagean–<br />
lower Meramecian supersequence, limited deposition of composite<br />
sequence C1 (Little Valley–Hillsdale) to the margin of the<br />
distal forel<strong>and</strong> to the west. Deposition extended farther updip in<br />
the east, where subsidence was more rapid. The small superimposed<br />
fourth-order sea-level changes periodically flooded the<br />
ramp to shallow depths, resulting in deposition of dominantly<br />
muddy, restricted <strong>carbonate</strong>s of composite sequence C1.<br />
The subsequent sea-level lowst<strong>and</strong> initiated localized red bed<br />
<strong>and</strong>/or eolianite deposition on the southern <strong>and</strong> eastern edge of<br />
the distal forel<strong>and</strong>, which continued to undergo differential<br />
subsidence over the hinge during the four transgressive–regressive<br />
events that formed composite sequence C2. Long-term relative<br />
sea-level rise caused progressive backstepping of the updip<br />
limits of sequences C1 to 4 onto the distal forel<strong>and</strong>, while considerable<br />
differential subsidence across the hinge promoted deepwater<br />
deposition in the southeast. The small, fourth-order sea-
30<br />
level changes only shallowly flooded the distal forel<strong>and</strong> to the<br />
west, with deposition of shallow lagoonal muds <strong>and</strong> intervening<br />
ooid shoals. More open marine skeletal facies were deposited in<br />
the more rapidly subsiding areas southeast of the hinge, <strong>and</strong> to<br />
the northeast updip along the basin axis.<br />
With long-term sea-level fall, widespread red beds <strong>and</strong><br />
eolianites were deposited, initiating deposition of composite<br />
sequence C3 (Taggard red beds). The earlier hinge became less<br />
active during C5 deposition but was reactivated during C6 deposition,<br />
along with differential subsidence seaward of a local hinge<br />
updip to the northeast. The sea-level rises from the initial lowst<strong>and</strong><br />
positions (base C5 <strong>and</strong> base C6) flooded the ramp only to shallow<br />
depths, so that restricted facies were deposited over much of the<br />
ramp. However, long-term sea-level rise, coupled with increasing-amplitude<br />
fourth-order eustasy, caused sequences C5 to C10<br />
to continue to onlap updip, with maximum flooding apparently<br />
occurring in the middle of sequence C8, followed by slight offlap<br />
during deposition of subsequent sequences (C9 <strong>and</strong> C10). These<br />
increased-amplitude fourth-order cycles, coupled with slight<br />
westward <strong>and</strong> northward migration of the hinges, promoted<br />
backstepping of ramp-margin facies <strong>and</strong> deposition of relatively<br />
open marine facies over the seaward edge of the distal forel<strong>and</strong><br />
during C7 to C8 <strong>and</strong> perhaps C9 deposition. The decrease in<br />
deposition of ooid grainstone from sequence C9 to C11 may be<br />
due to the increasingly humid climate. As amplitudes of sea-level<br />
change increased, increased sea-level falls allowed lowst<strong>and</strong><br />
siliciclastics <strong>and</strong> ooid grainstones to be deposited farther downdip<br />
during deposition of sequences C9 to base of C11.<br />
Composite sequence C4 was initiated with incision <strong>and</strong> deposition<br />
of s<strong>and</strong> bodies at the base of sequence C11, probably<br />
because of very low sea level. Increasingly humid conditions,<br />
coupled perhaps with uplift in the source areas to the northeast,<br />
generated the shale-prone sequences C11 <strong>and</strong> C12 (above the<br />
study interval). Thickening of at least sequence C11 into the<br />
foredeep may reflect a return to higher differential subsidence<br />
associated with the major hinge, culminating in thick, deep-water<br />
<strong>carbonate</strong>s along the basin margin. Carbonate for these must<br />
have been shed from thick ramp-margin bank complexes along<br />
the ramp margin in the southeast (exposed along Pine Mountain,<br />
Kentucky; Al-Tawil <strong>and</strong> Read, this volume). Renewed, relatively<br />
deep submergence during deposition of sequence C11 (Figs. 6, 7)<br />
probably reflects this increased subsidence coupled with the<br />
relatively high-amplitude fourth-order sea-level changes. Shaleprone<br />
sequence C11, as well as the overlying, <strong>carbonate</strong>-poor<br />
Mauch Chunk siliciclastic sequences above the study interval,<br />
formed as long-term sea level started to fall, global climate cooled<br />
(Mii et al., 1999), climate became more humid, <strong>and</strong> erosional<br />
unloading of the orogenic belt occurred (Ettensohn, 1994), all of<br />
which promoted terrigenous influx onto the ramp <strong>and</strong> siliciclastic<br />
progradation.<br />
Moderate-Amplitude Glacio-Eustasy<br />
<strong>and</strong> Late Mississippian Ramps<br />
The study interval spans the onset of the Southern Hemisphere<br />
Gondwana glaciation (Fig. 9), <strong>and</strong> the estimated<br />
periodicities of the sequences <strong>and</strong> the facies developed suggest<br />
that moderate-amplitude glacio-eustasy forced by long-term eccentricity<br />
(Horbury, 1989; Frakes et al., 1992; Smith <strong>and</strong> Dorobek,<br />
1993; Miller <strong>and</strong> Eriksson, 1997; Smith <strong>and</strong> Read, 2000, 2001;<br />
Wright <strong>and</strong> Vanstone, 2001) may have controlled the formation of<br />
the sequences. This moderate-amplitude glacio-eustasy, which<br />
increased in magnitude with time, was transitional into the highamplitude<br />
glacio-eustasy of the Pennsylvanian (Fischer, 1982;<br />
Goldhammer et al., 1991; Heckel, 1986). The moderate-amplitude<br />
AUS AL-TAWIL, THOMAS C. WYNN, AND J. FRED READ<br />
sea-level falls of the Late Mississippian caused large-scale basinward<br />
migration of siliciclastic shorelines, <strong>and</strong> caused lowst<strong>and</strong> to<br />
early transgressive siliciclastics to be deposited on the ramp,<br />
ramp margin, <strong>and</strong> ramp slope. The increasingly large, moderateamplitude<br />
sea-level rises superimposed on a long-term eustatic<br />
rise, <strong>and</strong> the load-induced forel<strong>and</strong> subsidence, caused extensive,<br />
periodic floodings of the ramp. However, the sea-level rises<br />
were not sufficiently large to drown the ramps, so that widespread<br />
grainstones were deposited not only on the ramp margin<br />
but also over large areas far updip on the ramp, where they<br />
interfinger laterally with muddy lagoonal <strong>carbonate</strong>s. Ramps<br />
that formed under such moderate-amplitude eustasy contrast<br />
with greenhouse ramps (Fischer, 1982), on which significant<br />
grainstone units tend to be limited to the outer part of the shallow<br />
ramp, <strong>and</strong> parasequences are meter-scale cycles dominated by<br />
restricted, muddy shallowing-upward facies (Read, 1995;<br />
Koerschner <strong>and</strong> Read, 1989; Goldhammer et al., 1990). At the<br />
other end of the spectrum, the bulk of the sequences of the<br />
Mississippian ramp in West Virginia lack the widespread, onramp,<br />
deeper-water maximum-flooding black shale–lime mudstone<br />
facies of the Pennsylvanian cyclothems, typical of higheramplitude<br />
glacio-eustasy during icehouse times (Heckel, 1986;<br />
Smith <strong>and</strong> Read, 2000, 2001).<br />
CONCLUSIONS<br />
This study illustrates how well cuttings <strong>and</strong> wireline logs,<br />
integrated with limited core <strong>and</strong> outcrop data, can be used to<br />
generate high-resolution sequence stratigraphies of Mississippian<br />
ramps in the subsurface. The Meramecian to Chesterian<br />
Greenbrier succession in the West Virginia portion of the Appalachian<br />
Basin formed on a tropical, mixed <strong>carbonate</strong>–siliciclastic<br />
ramp that developed on a subsiding forel<strong>and</strong> that was undergoing<br />
differential subsidence, during the initiation of significant<br />
Southern Hemisphere, Gondwana, glaciation. The sequence<br />
stratigraphy was analyzed using two regional cross sections, one<br />
a dip section <strong>and</strong> the other an axial section, which were based on<br />
subsurface well cuttings <strong>and</strong> wireline logs from numerous shallow<br />
wells, several detailed outcrop sections, <strong>and</strong> a core. The<br />
Greenbrier succession forms the transgressive part of the upper<br />
Meramecian–Chesterian supersequence, <strong>and</strong> is made up of fourthorder<br />
depositional sequences stacked into variably developed<br />
third-order, composite sequences. The fourth-order depositional<br />
sequences reach thicknesses of tens of meters to a hundred<br />
meters. They are bounded by lowst<strong>and</strong> to early transgressive<br />
siliciclastic units <strong>and</strong> local updip eolianites. Transgressive banks<br />
developed along the ramp margin locally, but on the shallow<br />
ramp, transgressive shallow-water <strong>carbonate</strong>s were recognized<br />
in some areas but not in others. Highst<strong>and</strong> facies are extensive<br />
lagoonal lime wackestone–mudstone, along with ooid grainstone<br />
<strong>and</strong> skeletal grainstone–packstone that developed along the ramp<br />
margin <strong>and</strong> as extensive isolated units on the ramp interior, <strong>and</strong><br />
constitute the bulk of many sequences. Skeletal wackestone grading<br />
seaward into laminated, argillaceous lime mudstone developed<br />
on the deeper ramp <strong>and</strong> ramp slope.<br />
Deposition of the succession was influenced by differential<br />
subsidence of the foredeep, possibly marking a phase of late<br />
Paleozoic thrust loading. Paleoslopes on the distal forel<strong>and</strong> were<br />
a few centimeters per kilometer, increasing to 30 cm/km over the<br />
hinge. Accommodation rates were a few centimeters per kiloyear<br />
on the distal forel<strong>and</strong> increasing to 10 to 30 cm/ky on the<br />
proximal forel<strong>and</strong>. Tectonic subsidence generated a basinwardthickening<br />
wedge of sediments 0 to 900 m thick. Rapid thickening<br />
of many of the sequences occurs across the basin hinge, as well as<br />
across broad downwarps on the ramp, which were intermittently
active. This differential subsidence generated a complex mosaic<br />
of facies on the ramp; thus, sequences rarely show simple upward-deepening<br />
to upward-shallowing successions, or uniformly<br />
basinward-thickening trends.<br />
The fourth-order sea-level fluctuations that formed the sequences<br />
likely relate to glacially driven Milankovitch long-term<br />
eccentricity forcing. They were responsible for development of<br />
the characteristic sequence stacking with basal siliciclastic-prone<br />
units <strong>and</strong> upper <strong>carbonate</strong>-prone units, because they caused<br />
large shifts in the shoreline across the ramp. Sea-level changes<br />
became larger through time, which, along with the increasingly<br />
humid climate, caused an upsection change in the supersequence<br />
from dominantly restricted, ooid grainstone–prone sequences to<br />
more open marine, skeletal limestone–prone sequences with<br />
significant siliciclastic units. Third-order sea-level falls appear to<br />
have accentuated fourth-order falls, causing widespread deposition<br />
of red beds <strong>and</strong> marine siliciclastic units, <strong>and</strong> some incision<br />
of underlying units. Although subsidence rates on the proximal<br />
forel<strong>and</strong> were an order of magnitude greater than those acting on<br />
the distal forel<strong>and</strong>, fourth-order sea-level fluctuations (indicated<br />
by interbasinal correlation) still were the prime cause of the<br />
sequences, whereas tectonics strongly controlled regional <strong>and</strong><br />
local thickness <strong>and</strong> facies variations.<br />
ACKNOWLEDGMENTS<br />
We would like to thank the West Virginia Economic <strong>and</strong><br />
Geological Survey <strong>and</strong> especially Lee Avery, Dave Matchen, Pat<br />
Johns, <strong>and</strong> staff, for cuttings, cores, maps, <strong>and</strong> digital data files. We<br />
give special thanks to Taury Smith, whose outst<strong>and</strong>ing Illinois<br />
Basin work helped clarify many of the problems of Mississippian<br />
facies <strong>and</strong> stratigraphy, as did the work of Garl<strong>and</strong> Dever of the<br />
Kentucky Geological Survey. Bill Henika (Virginia Division of<br />
Mineral Resources) provided valuable insight into problems of<br />
Appalachian geology. Jim Molloy, Art Hall, Sarah Crews, Steve<br />
Turpin, Jane Roa, John Leone, Chad Haiar, Susan Barbor, Pat<br />
Barnhart, Don Thorsteson, Ashley Goodrich, Julia Caldwell, <strong>and</strong><br />
Jen Hatfield helped collect <strong>and</strong> process data. Mike Pope, Anna<br />
Balog, Amy Khetani, Dan Miller, <strong>and</strong> Brian Coffey provided<br />
valuable discussion <strong>and</strong> help, <strong>and</strong> Ellen Mathena did much of the<br />
editing <strong>and</strong> typing. The study was supported by Petroleum Research<br />
Fund grant 34137-AC8, <strong>and</strong> National Science Foundation<br />
grant EAR-9725308 to J.F. Read, <strong>and</strong> funding from Mobil Oil with<br />
the strong support of Ron Kreisa <strong>and</strong> Jim Markello. We thank<br />
American Association of Petroleum Geologists, Geological Society<br />
of America, <strong>and</strong> the Society of Professional Well Log Analysts for<br />
additional support. Dave Hunt, Steve Dorobek, <strong>and</strong> Wayne Ahr<br />
greatly improved the paper with their insightful reviews.<br />
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Section # Name County, State Latitude Longitude<br />
1 Canaan Valley Tucker County, West Virginia 39.002800 -79.467500<br />
2 R<strong>and</strong>olph 28 R<strong>and</strong>olph County, West Virginia 38.716940 -80.132780<br />
3 Mingo R<strong>and</strong>olph County, West Virginia 38.488200 -80.041700<br />
4 Greenbrier 21 Greenbrier County, West Virginia 37.965000 -80.627220<br />
5 Summers 7 Summers County, West Virginia 37.794170 -80.744170<br />
6 Sunoil Test Core #2 Pocahontas County, West Virginia 38.108333 -80.266667<br />
7 Greenbrier 15 Greenbrier County, West Virginia 37.751390 -80.586670<br />
8 Union Monroe County, West Virginia 37.610800 -80.587500<br />
9 Mercer 25 Mercer County, West Virginia 37.440000 -81.018610<br />
10 Ingleside Mercer County, West Virginia 37.313900 -81.141700<br />
11 Bristol Washington County, Virginia 36.75 -82.05<br />
12 Wayne 277 Wayne County, West Virginia 38.271111 -82.538611<br />
13 Wayne 1497 Wayne County, West Virginia 38.1975 -82.334167<br />
14 Lincoln 155 Lincoln County, West Virginia 38.306111 -82.202222<br />
15 Lincoln 1375 Lincoln County, West Virginia 38.163060 -82.130280<br />
16 Mingo 796 Mingo County, West Virginia 37.924170 -82.206670<br />
17 Logan 102 Logan County, West Virginia 37.96861 -81.973056<br />
18 Logan 922 Logan County, West Virginia 37.835280 -81.960000<br />
19 Logan 388 Logan County, West Virginia 37.78972 -81.784444<br />
20 Wyoming 253 Wyoming County, West Virginia 37.77167 -81.51806<br />
21 Wyoming 87 Wyoming County, West Virginia 37.525556 -81.637500<br />
22 McDowell 63 McDowell County, West Virginia 37.357500 -81.653611<br />
23 McDowell 792 McDowell County, West Virginia 37.44306 -81.400833<br />
24 Mercer 16 Mercer County, West Virginia 37.50833 -81.17639<br />
25 Mercer 1 Mercer County, West Virginia 37.328056 -81.224444<br />
26 Mercer 27 Mercer County, West Virginia 37.39722 -81.0525
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