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60
PENNSYLVANIAN-PERMIAN VEGETATIONAL
CHANGES IN TROPICAL EURAMERICA
William A. DiMichele, C. Blaine Cecil, Dan S. Chaney, Scott D. Elrick,
Spencer G. Lucas, Richard Lupia, W. John Nelson, and Neil J. Tabor
INTRODUCTION
Vegetational changes across the Pennsylvanian-Permian boundary are recorded in several
largely terrestrial basins across the Euramerican portions of equatorial Pangea. For the
purposes of this paper, these include the Bursum-Abo Formation transition and its equivalents
in several small basins in New Mexico, the Halgaito Formation of southeastern Utah, Markley
Formation of the eastern shelf of the Midland Basin in north-central Texas, Council Grove
Group of northern Oklahoma and southern Kansas, and Dunkard Group of the central
Appalachian Basin. This transition also is recorded in numerous basins in Europe, reviewed by
Roscher and Schneider (2006), based on paleoclimate indicators preserved in those regions.
Collectively, these deposits form a west-to-east transect across the Pangean paleotropics and
thus provide a paleogeographic setting for examination of both temporal and spatial changes in
vegetation across the Pennsylvanian-Permian boundary (Figures 1 and 2).
The Pennsylvanian-Permian transition records the change from wetland vegetation as the
predominant assemblages found in the plant fossil record, to seasonally dry vegetation. This
has often been called the “Paleophytic-Mesophytic” transition, a concept that is flawed
DiMichele, W. A., Cecil, C. B., Chaney, D. S., Elrich, S. D., Lucas, S. G., Lupia, R., Nelson, W. J., and Tabor, N. J., 2011,
Pennsylvanian-Permian vegetational changes in tropical Euramerica, in Harper, J. A., ed., Geology of the Pennsylvanian-
Permian in the Dunkard basin: Guidebook, 76th Annual Field Conference of Pennsylvania Geologists, Washington, PA, p. 60-
102.
Figure 1. Continental configuration at the Pennsylvanian-Permian boundary. Yellow ovals indicate the
principal areas discussed herein: Left – New Mexico and Utah, Center – Texas and Oklahoma, Right –
Central Appalachians/Dunkard. Map courtesy of Ron Blakey, Northern Arizona University.
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conceptually but does reflect a pattern of vegetational change at a certain coarse level of
resolution (DiMichele et al., 2008). Interpreted simplistically, this change is asynchronous
across the Pangean tropics. In detail, however, it is clear that “dry” floras appear as early as the
Middle Pennsylvanian in central and western Pangea (Scott et al., 2010; Falcon-Lang et al.,
2009; Plotnick et al., 2009; Galtier et al., 1992; Dolby et al., 2011), and that coals are
effectively gone from the stratigraphic record of far western Pangea after the end of the Atokan
(Bolsovian, mid-Moscovian) (Lucas et al., 2009a). In contrast, organic facies persist, even if
much reduced in thickness, quality and abundance, together with associated wetland plants, into
the Early Permian in central and east-central Pangea, at the same time as deposits dominated by
plants of seasonally dry environments are increasing in frequency (e.g., Martin, 1998; Kerp and
Fichter, 1985).
When the complexity of this transition is embraced, it is found that its nature takes on
different qualities depending on the spatial and temporal scale at which it is resolved and
Figure 2. Stratigraphic sections for the areas discussed herein, positioned West-to-East along the Pangean
paleoequator. Stars indicate the position of the first appearance of seasonally dry flora in the local
stratigraphic section. Solid lines are accepted temporal boundaries or boundaries of formations or groups.
Dotted lines are inferred temporal boundaries.
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examined. On the temporal scale of glacial-interglacial cycles, floras consisting of species
tolerant of seasonal moisture deficits are found intercalated with floras of humid environments,
resulting in the alternation of taxonomically distinctive assemblages, at least on the northern
side of the Appalachian-Variscan central Pangean mountain range (Mapes and Gastaldo, 1986;
Broutin et al., 1990; DiMichele and Aronson, 1992 – the history and dynamics of this mountain
belt were complex, such that the mountain range was not contemporaneously present across the
Pangean continent, as is frequently shown in paleogeographic reconstructions – see discussion
in Roscher and Schneider, 2006). In some instances, this floristic alternation can occur between
different beds of the same outcrop (DiMichele et al., 2005a; Falcon-Lang et al., 2009). The
changing floras reflect climate contrasts that have been interpreted to change in concert with
glacial-interglacial cyclicity (Falcon-Lang, 2004; DiMichele et al., 2010c; Falcon-Lang and
DiMichele, 2010). Such vegetational cyclicity is interpreted in this paper to reflect shifts
between cool-wet (glacial) and warm-dry (interglacial) (Cecil and Dulong, 2003) tropical
climate states. Furthermore, the cyclicity observed for vegetation, and inferred for climate,
occurs within depositional packages interpreted to result from sea-level changes, recorded in the
classic Pennsylvanian cyclothems. Also seen are climate trends on longer time scales of 2-6
million years (Darrah, 1969; Cecil, 1990; Falcon-Lang et al., 2011a; Allen et al., 2011) linked
to extended periods of drier climate, higher mean sea-level, and reduced polar ice volume
(Fielding et al., 2008; Rygel et al., 2008; Eros et al., in press). At the longest temporal scale is
the long-recognized trend toward drier tropics manifested in the Pennsylvanian-Permian
transition in the western and central parts of equatorial Pangea (Remy, 1975 for discussion from
a paleobotanical perspective). This longest-term trend is not monotonically unidirectional, but
is an average of an oscillatory trend ultimately leading to a reduction in the number and length
of intervals of significantly wet tropical climate. Within this longer-term trend it might be said
that climate fluctuations continue, but that, on average, the wet periods become less wet and the
dry periods become more dry. Cecil et al. (2011) suggest that this trend is due either to the final
assembly of Pangea or to the rise of the central Pangean mountains. The trend happens,
inexplicably, at a time when evidence has been adduced for the return of large amounts of ice at
the beginning of the Permian (e.g., Montañez et al., 2007; Fielding et al., 2008).
BASES OF INFERENCE/RATIONALE
The environmental preferences of fossil plants can be resolved in many environmental
dimensions. Very commonly, however such preferences are broadly construed as favoring
habitats with year-round high soil-moisture content or environments where there are seasonal
deficits of moisture. These preferences are determined from plant architecture, growth habit,
and environmental correlates of occurrence, including the sedimentary environments in which
the fossils are found. Such inferences come from the literature and from the personal
experiences of the authors, which are both published and unpublished. The extent to which
such patterns are reflective of regional climates is contextual, dependent on broader spatial and
temporal patterns of floristic occurrence, the sedimentological context of the floras, and the
extent of the depositional systems of which they are a part. Consequently, the focus herein is
more on the plants and less on the accessory information that accompanies the paleobotanical
samples. That said, there are clear preservational biases in the fossil record of plants. Paleosols
tell us that vegetation was present during the times of drier climate, but we have a limited
understanding of that vegetation, at least at present, despite a considerable number of
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publications that discuss it (for recently published papers see, for example, Feldman et al.,
2005; Falcon-Lang et al., 2009; Plotnick et al., 2009; Scott et al., 2010; DiMichele et al.,
2010c; Falcon-Lang et al., 2011a; Dolby et al., 2011). The dry soil plants do become well
represented in the Permian fossil record of the equatorial region, reflecting, perhaps, changed
sedimentary systems and changed paleobotanical search images (Kerp, 1996).
To the extent possible, we consider information from paleosols relevant to rainfall and
water table, sedimentary patterns, geochemical proxies for various aspects of climate, climate
modeling, near-field studies of the ice record, and other sources of information (e.g., Cecil et
al., 2003; Feldman et al., 2005; Montañez et al., 2007; Tabor and Montañez, 2004; Poulsen et
al., 2007; Tabor and Poulsen, 2008; Tabor et al., 2002, 2008; DiMichele et al., 2010c; Horton et
al., 2010).
In a glacial world, there also is very likely to be a close, mechanistic relationship among
tropical climate, sediment transport patterns, global sea-level dynamics, polar ice-volume, and
the composition of the atmosphere. This is framed broadly, though not entirely, by considering
rock stratigraphic sequences in terms of a combined climate-sea level-tectonics framework
(Cecil and Dulong, 2003; Poulsen et al., 2007; Elrick and Nelson, 2010; Allen et al., 2011; Eros
et al., in press). Here we are operating under the presumption that sea-level lowstand,
particularly from its midpoint into the earliest phases of sea-level rise, is associated with the
wettest equatorial climates. If peats/coals are present in a cycle, this is the most likely time for
them to have formed. The driest equatorial climates occur at sea-level highstand and continue,
diminishing, into the phases of sea-level fall.
It is during these stages of the cycle that limestones, evaporites, and paleosols indicative of
seasonal climates are most likely to form. This model (Cecil et al., 2003; Poulsen et al., 2007;
Eros et al., in press) has been slow to emerge for several reasons. Perhaps foremost is
incumbency. Earlier models, such as that of Bohacs and Suter (1997), which call upon sea-
level rise to initiate and sustain cratonic-basin blanket-peat formation, treat climate as a
constant (e.g., p. 1618, where they note that “only accommodation varies significantly and most
other variables remain constant (flora, climate, environment, etc.)”). This, in effect, removes
climate as a variable in their sequence stratigraphic model, without specifying the state of that
invariant climate. Such models, therefore, attribute all changes in lithological sequences to sea-
level changes. When considered within a “total evidence” framework, however, these models
fall short. For example, were sea-level rise the driver of peat formation, it is then necessary to
explain why, during the late Paleozoic, peat beds did not form in the western parts of the
Pangean equatorial region while, at the same time, extensive peats were forming in the
Midcontinent and eastern coal basins (e.g., Cecil et al., 2003; Bishop et al., 2010).
Furthermore, modern sea-level rise has not created vast peat-forming swamplands along
coastlines worldwide.
The prevailing understanding that the modern tropics are wettest during interglacials and
dry during glacial phases of glacial-interglacial cycles is not as much an iron-clad truth as
assumed, and is changing as greater study of historical Amazonian climate patterns emerges.
The Amazonian lake record is consistent with an equatorial band of the Amazon being wettest
at or near glacial maximum, possibly caused by constraint of the intertropical convergence zone
(ITCZ) (e.g., Bush et al., 2002). In addition, the great majority of modern Amazonian
interglacial rainfall reflects recycling of water by the tropical rainforest (e.g., Bush and Flenley,
2007), something that may be entirely dependent on the presence of angiosperms and their
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unique water transport and evapotranspirative capacities (Boyce and Lee, 2010; Boyce et al.,
2010). The latter point means that pre-angiospermous systems, such as those of the late
Paleozoic, may not have been able to create “rainforest” levels of moisture over large areas
during inter-glacial episodes (Remy, 1975). Without the powerhouse evapo-transpiration
created by the angiosperms, the implication is that tropical wetness must have been generated
by a largely physical mechanism (e.g., the restriction of the ITCZ by polar ice and attendant
atmospheric conditions at the ice front – Cecil et al., 2003).
SPATIAL PATTERNS WITHIN THE PALEOEQUATORIAL REGION
The main geomorphological trend that is most easily observed and tracked in space and
time is the presence or absence, or relative abundance, of coal and the subjacent “underclay”
paleosols (Cecil et al, 2003; Cecil et al., 2011). Well developed “underclay” paleosols can
often be traced in somewhat elevated areas well beyond the occurrence of the superjacent coal
beds where soils with aquic to udic soil moisture regimes inhibited or even excluded the
development of peraquic conditions that are conducive to the development of histosols (peat
precursors to coal) on the land surface (Cecil et al., 2011). The physical, chemical, and
mineralogical properties of the underclay paleosols beneath coal beds clearly indicate low stand
sea level conditions and pedogenesis under a humid to perhumid paleoclimate. Thus, coal
represents the “wettest” (glacial) climate phase of a glacial-interglacial cycle when rainfall was
sufficient to perpetually maintain the water table at the surface.
Also relatively readily observed as indicators of climate are seasonally-dry floral elements.
It is possible to assess: (a) their “on-average” abundance or the commonness with which they
are encountered in a particular geological unit, such as a formation, (b) their degree of mixing
with the more hygromorphic (wetland) floral elements, and (c) the time, or times, at which they
first appear and/or become the most commonly encountered vegetational elements. With regard
to this latter point, “dominance” is a term we will use with caution and constrain to relative
abundance within a single fossil assemblage, which is considered herein a proxy for a natural
plant community. The term “dominance” when expanded to larger temporal or spatial scales
implies the existence of a single vegetation type or species pool in equilibrium with, or
responding to, a monotonic climate. Rather, we see two or more distinctive, climatically
characteristic species pools reflective of different mean climate states (sensu Cecil et al., 2011).
The presence of these different species pools in sedimentary basins is a response to changes in
soil moisture. Soil moisture changes, allowing for the ever-present effects of local edaphics,
may reflect changes in regional climate brought about by glacial-interglacial cyclicity. As
mentioned previously, we consider climatic cyclicity within a single glacial-interglacial cycle to
be wet/cool during sea level lowstands (maximum ice volume) and seasonally dry during sea
level highstands (minimum ice volume). These glacial-interglacial patterns, on the 105-year
time scale, will be superimposed on the longer term changes in climate, on the 106, 2 to 6
million year scale (Allen et al., 2011), when mean climate state was overall warmer-drier or
cooler-wetter, reflective of changes in mean global ice volume and sea-level. And patterns on
both these time scales will be encapsulated within the long term Pennsylvanian-Permian climate
change, on the 107-year time scale.
At the 107-year, Pennsylvanian-Permian time scale, there clearly is a trend toward overall
drier climate beginning in the late Middle Pennsylvanian and continuing through the Early
Permian (Remy, 1975; Roscher and Schneider, 2006; Tabor and Poulsen, 2008), perhaps
65
attributable to the final assembly of the mega-continent Pangea, the northward drift of the
continental landmass, and changes in atmospheric and oceanic circulation (Cecil, 1990; Parrish,
1993; Roscher and Schneider, 2006; Tabor et al., 2008). During glacial-interglacial
oscillations, this means that the wetter phases became less wet and the drier phases became
even drier. This pattern, ultimately, permits the survival in basinal lowlands of plants more
tolerant of moisture seasonality, at all phases of the cycle. As basinal landscapes “dry out” over
time, the wetland specialist plants become less common and often are spatially concentrated,
occurring together as habitat specialists in reduced diversity assemblages in landscape “wet
spots” (DiMichele et al., 2006).
Westernmost Pangea: New Mexico and Southeastern Utah
During the late Paleozoic, present day New Mexico was located in the western part of
equatorial Pangea as an archipelago of basement cored uplifts surrounded by shallow marine
and, in some cases, locally terrestrial basin floors of the Ancestral Rocky Mountain basin and
range (Figure 1). This region of the supercontinent appears to have been persistently more
seasonally dry than more central parts of the megacontinent throughout the later Pennsylvanian
and into the Early Permian, with a distinct monsoonal pattern of precipitation, evidently sourced
from Panthalassa rather than Tethys (Parrish, 1993; Tabor and Montañez, 2004; Tabor and
Poulsen, 2008). Glacial-interglacial scale lithological cyclicity is more obscure in this region
than in the central Pangean coal basins due in large part to the near absence of coal/underclay as
a low-stand marker bed. Cyclicity can be detected in some of the marine (Elrick and Scott,
2010) and terrestrial Pennsylvanian sections, and in mixed marine and terrestrial Lower
Permian rocks (Mack et al., 2010). It also can be detected as the higher-level, stage-scale
fluctuations in climate state (sensu Cecil et al., 2011), e.g., a much drier Missourian
(Kasimovian) than either the earlier Desmoinesian (Moscovian) or the later Virgilian (Falcon-
Lang et al., 2011a), and in the same overall drying trend that characterizes the Pennsylvanian-
Permian transition in general (Remy, 1975; Cecil et al., 1985; Cecil, 1990).
Our studies in New Mexico reveal the shift from wetter-to-drier climates earlier than any
other area we have examined in detail. In the central part of the state, Socorro County, a nearly
complete lithological transition, with associated flora and both marine and terrestrial fauna is
preserved from the Middle Pennsylvanian to the Early Permian (Lucas et al., 2009a; Krainer
and Lucas, 2009). The Pennsylvanian-Permian transition is well exposed in both this central
region, and to the north in areas near Albuquerque (Lucas and Krainer, 2004; Krainer and
Lucas, 2004; Lucas and Zeigler, 2004), and still farther north in areas to the west of Sierra
Nacimiento, east of Cuba (Lucas et al., 2010a, b; Krainer and Lucas, 2010; DiMichele et al.,
2010a). Seasonally-dry floral elements begin to appear in the Middle to Late Pennsylvanian in
each of these areas (depending on the age of the oldest exposed terrestrial rocks) and are a part
of nearly every plant assemblage collected thereafter, becoming outright the overwhelmingly
most common vegetation type either in the later part of the Late Pennsylvanian, in the Bursum
Formation or in the wholly Lower Permian Abo Formation (e.g., Mamay and Mapes, 1992;
Utting et al., 2004; Tidwell and Ash, 2004; DiMichele et al., 2004; DiMichele et al., 2007;
Lucas et al., in press a). Geological ages throughout New Mexico are among the best
constrained of any of the Pennsylvanian-Permian transition regions considered in this paper, all
accurately calibrated with conodonts and/or fusulinids from adjacent and intercalated marine
strata (e.g., Lucas et al., 2009a).
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Floras from these areas require some explanation because of their mode of preservation.
In the Atokan (Bolsovian/middle Moscovian) Sandia Formation, coal beds are represented only
by a few isolated lenses, geographically very localized (e.g., Thompson, 1942; Armstrong et al.,
1979; Kues and Giles, 2004; Krainer et al., 2011) and not thick enough or laterally extensive
enough to be of economic significance. They are associated with wetland floras (Lucas et al.,
2009b), but are not even remotely comparable in extent to coal-forming paleoenvironments of
similar age in the eastern USA and in Western Europe. Paleosol deposits, so called “fire clays”,
associated with the coal beds have been mined locally for brick and pottery (Lucas et al.,
2009b), and indicate periods of humid to perhumid climates and intense weathering.
The Desmoinesian (Asturian/late Moscovian) in New Mexico reflects a time of elevated
sea levels relative to the earlier Pennsylvanian, as across much of the Ancestral Rocky
Mountain zone, with local tectonics creating an essentially non-marine basin floor setting in an
area east of Socorro, central New Mexico, that persisted through the late Missourian (Lucas et
al., 2009a). In this area, glacial-interglacial-scale oscillation of wetland and seasonally-dry
floras can be detected in parts of the section. The wetland floras, although rare, are
authochthonous to parautochthonous and completely dominated by typically Pennsylvanian
wetland plants, such as pteridosperms and Psaronius tree ferns. These assemblages appear to
be preserved in lowstand coastal-plain deposits, based on their sequence position (lowstand),
and sedimentary characteristics (thin, widespread, autochthonous, rooted siltstone deposits).
They are intercalated with shales that appear to represent marine high-stand to falling-stage
shallow shelf deposits, containing brackish-to-marine fauna and allochthonous plant remains
(Figure 3). These latter floras contrast sharply with those of the lowstand, wetter phases, and
contain predominantly the xeromorphic/seasonally dry plants, such as Sphenopteris germanica,
Charliea, and walchian conifers, with admixtures of wetland elements, mainly pteridosperms
(DiMichele et al., 2010b). The Missourian of this area is particularly dry (Lerner et al. 2009;
Falcon-Lang et al, 2011a) (Figure 4), as it appears to be across the North American part of the
Pangean tropics (Cecil et al., 1985; Cecil, 1990; Martino and Blake, 2001; Martino, 2004;
DiMichele et al., 2010b; Cecil et al., 2011). Floras preserved in Socorro County vary from
completely dominated by seasonally-dry substrate species, including conifers, callipterids,
Charliea, Taeniopteris and Sphenopteris germanica (Falcon-Lang et al., 2011a), to mixed, with
wetland elements forming a background to primarily seasonally-dry substrate taxa (Lerner et
al., 2009). A coeval flora of Missourian age is known from 100 km (62 mi) to the north, near
Albuquerque, from the Kinney Brick Company quarry in the Manzanita Mountains (Mamay
and Mapes, 1992). This deposit was formerly thought to be of middle Virgilian age, is
dominated by wetland plants, but has a well represented and diverse array of seasonally dry-
substrate taxa, including Sphenopteris germanica (Sphenopteridium manzanitanum, Mamay,
1992), Charliea manzanitana (Mamay, 1990), Dicranophyllum readii (Mamay, 1981), several
forms of conifer and pteridosperms typically found in drier substrate settings (Lucas et al., in
press a).
At the Pennsylvanian-Permian transition, including the Bursum (and its equivalents) and
lower Abo formations (Lucas and Krainer, 2004; Krainer et al., 2009; Krainer and Lucas, 2010;
Mack et al., 2010), floras are regionally variable and show different mixtures of wetland and
seasonally dry elements. In Socorro County, Bursum floras are preserved in coastal plain
environments (Figure 5A) that are not substantially different from those of the Desmoinesian
and Missourian in the same area. These floras are, however, dominated by dryland species,
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including walchian conifers and cordaitaleans. The similar sedimentary environments permit an
isotaphonomic comparison (sensu Behrensmeyer and Hook, 1992) and suggest that the wettest
intervals of climate cycles were significantly drier than during earlier time periods in the
Desmoinesian through Virgilian. In the Lucero Uplift, 80 km (50 mi) to the NNW of the
Figure 3. Desmoinesian (Middle Pennsylvanian) age exposure and plants from the Bartolo and Tinajas
members of the Atrasado Formation, Socorro County, New Mexico. A—Bartolo Member – two lower shales
separated by calcareous sandstone, each contain allochthonous, seasonally dry flora. Crest of ridge marks
the position of the Amado Member, immediately above the Desmoinesian-Missourian boundary. B—
Walchian conifer twig with leaves (USNM 543964; USNM loc 43527). C—Charliea pinnule (USNM 543967;
USNM loc 43469). D—Sphenopteris germanica pinnae (USNM 508828; USNM loc 43469). All specimens
approximately 3 X magnification.
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Socorro area, latest Pennsylvanian floras from the Red Tanks Member of the Bursum
Formation are also preserved in lacustrine sediments within a coastal plain setting. Conditions
were rarely humid enough to engender peat/coal formation (Krainer and Lucas, 2004). The
plants are predominantly wetland taxa but with a large admixture of conifers, callipterids and
taeniopterids (Carrizo Arroyo flora, see Tidwell and Ash, 2004), all elements of seasonally dry
vegetation. Such a mixed flora suggests that the wetland species were confined to wet places
on an otherwise seasonally dry landscape (as suggested for the regional climate by Krainer and
Lucas, 2004). The northern-most Late Pennsylvanian plant-bearing strata (El Cobre Canyon
Formation, a Bursum equivalent, and Lower Permian Arroyo del Agua Formation) (Lucas et al.,
2010b), were deposited in braided river systems (Krainer and Lucas, 2010). The floras are
numerically dominated by wetland plants, like those from the approximately equivalent aged
Red Tanks Member in the Lucero Uplift, but, as there, these wetland elements appear to have
been drawn from stream banks or from locally colonized areas within the shifting braidplains.
Most of the assemblages contain uncommon, but consistently present, plants from substrates
that were likely seasonally dry, such as walchian conifers, Taeniopteris and cordaitaleans
(DiMichele et al., 2010a).
In all regions of New Mexico, the Early Permian rocks record the long-term transition to a
seasonally drier climate mean at all stages of climate cycles (Mack et al., 1991; Krainer and
Lucas, 2010). Depending on facies, the flora of these Early Permian rocks varies considerably.
A major facies change takes place as a consequence of a pulse of the Ancestral Rocky
Mountain uplift beginning in the early Wolfcampian (latest Pennsylvanian). This produces the
synorogenic Abo Formation, which formed on a vast coastal plain, represented by channel-form
fluvial deposits and overbank siltstones and mudstones (Figures 5B, 6). Abo Formation
exposures are intercalated with marine deposits of the Hueco Group in southern New Mexico,
Figure 4. Missourian (Late Pennsylvanian) gypsum-carbonate sabkha deposit, Tinajas Member of the
Atrasado Formation, Socorro County, New Mexico. The bases of coniferopsid (conifers or cordaitaleans)
trees are preserved within this deposit, rooted in micritic limestone and buried by gypsum and carbonate
(see Falcon-Lang et al, 2011a). A—Large tree base with attached roots, extending into the basal micrite,
exposed by erosion of the gypsum-carbonate bed. B—Tree trunk extending upward into the gypsum-
carbonate bed (scale = 50 cm).
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Figure 5. Wolfcampian (latest Pennsylvanian) age Bursum Formation and Early Permian age Abo
Formation, Socorro County, New Mexico. A—Bursum Formation plant-bearing shales, perhaps deposited in
a floodplain lake. Most of the Bursum consists of paleosols and limestones, and is without plant fossils.
B—Succession in the Socorro County area: Atrasado Formation (white, limestone in foreground), Bursum
Formation (gray, low slope and hills in middle foreground), Abo Formation (red hills), Yeso Formation
(yellow-gray hills on skyline). Pennsylvanian-Permian boundary lies at the Bursum-Abo contact just above
the center of the photograph.
Figure 6. Early Permian age Abo Formation, Socorro County, New Mexico. A—Siltstone channel fill
characteristic of deposits in which plants are commonly found. B—Lower and Middle portions of the Abo
formation shown in stream cutbank. Siltstones with strong pedogenic overprint characterize the lower
portions of the Abo Formation. Thin, sheet siltstones and fine grained sandstones are generally more
common in the middle and upper part of the Abo Formation.
(Lucas et al., 1995), but otherwise the Abo is entirely terrestrial, with thick calcic Vertisols and
stream deposits, with loessites further to the north (Mack et al., 1991; Kessler et al., 2001; Mack
et al., 2003; Mack, 2007). The flora from these beds comes primarily from siltstones of the
channel facies and is of remarkably low diversity (DiMichele et al., 2007), consisting of patches
of various walchian conifers and the peltasperm Supaia, with a small, and generally rare,
70
admixture of callipterids and other plants of seasonally dry habitats (Figure 7). This is
particularly true for the central and southern areas of the state. In the northern areas, such as
Arroyo del Agua, the flora is more diverse than in the typical Abo redbeds, but of lower
diversity than the Late Pennsylvanian floras of the same area. The most common elements
include conifers, callipterid peltasperms, and Taeniopteris, all seasonally dry plants. However,
there also are the common wetland elements, calamitaleans, pteridosperms, and tree ferns, that
probably grew within the floodplains of braided stream systems (DiMichele and Chaney, 2005).
Approximately 500 km (311 mi) to the NW of Socorro County, and perhaps at the margin
of the paleotropical, equatorial belt (Dubiel et al., 1996), are deposits at the Pennsylvanian-
Permian boundary in southeastern Utah (Figure 8). The sedimentary systems in this region
reflect greater aridity at all times than those in New Mexico (Soreghan et al., 2002; Tabor et al.,
2008), with regional climate varying from arid (aeolian dune sands, loessites and Calsisols) to
subhumid (rivers and flood plains with Calcisols, colonized by plants). Keeping in mind that
this area is paleolatitudinally north of the equatorial region, there appears to be an overall trend
Figure 7. Plant fossils typical of seasonally dry flora, from the Bursum and Abo formations, Socorro
County, New Mexico. A—Walchian conifer branch (field photo; specimen not collected). B—Rhachiphyllum
sp., callipterid peltasperm frond segment (USNM 543963; USNM loc. 43460). C—Autunia conferta, callipterid
peltasperm frond segment (tip) (USNM 543955; USNM loc. 42251). D—Supaia anomala, supaioid peltasperm
leaf fragment (USNM 543956; USNM loc 42255).
71
from semiarid during the early Late Pennsylvanian to seasonally wetter conditions in the Early
Permian (Soreghan et al., 2002). Like the Abo Formation, in New Mexico, the Early Permian
deposits in southeastern Utah are significantly more “inland” than older rocks of the Late
Pennsylvanian.
In a simplified stratigraphy from this area, the Middle and Late Pennsylvanian strata are
assigned mostly to the Honaker Trail Formation, the uppermost part of which is represented by
a cyclic series of paleosols, channel or braidplain sandstones, non-marine limestones and shales
possibly reflective of a strong global-scale, allocyclic eustatic signal (Hite and Buckner, 1981;
Soreghan et al., 2002). We did not find any plant macrofossils, other than a few unidentifiable
axes, in these latest Pennsylvanian rocks, though a flora composed entirely of Pennsylvanian
wetland species was reported from clastic rocks in the Honaker Trail Formation, 150 km (93
mi) to the north, near Moab, Utah (Tidwell, 1988). Palynological analyses of clastic and
evaporitic rocks from the Paradox Formation of SE Utah, which is of Desmoinesian age
(Middle Pennsylvanian), were carried out by Rueger (1996). Substantial climate cycling at this
time is detectable from these analyses, as it was in similar age rocks from New Mexico.
Overall, the SE Utah palynological assemblages are dominated, on a percent-abundance basis,
by striate bisaccate monosaccate grains, generally considered suggestive of seasonally dry flora.
However, these grains cycled in abundance through halites and their clastic interbeds. Typical
wetland-flora palynomorphs were found to be rare but present (including Lycospora), and were
negatively correlated with striate bisaccate grains.
The Honaker Trail Formation is overlain by the Halgaito Formation, which in the Valley
of the Gods, in SE Utah (~ 16 km [~10 mi] north of Mexican Hat, Utah), consists of eolian
sandstones and loessites, non-marine limestones, paleosols indicative of seasonally dry
climates, and water-lain sandstones of various geometries, from thin bedded and flat-bottomed
to bar forms of moderately large channels (Soreghan et al., 2002; Tabor et al., 2008). The
Halgaito is thought to cross the Pennsylvanian-Permian boundary (e.g., Condon, 1997), and
contains well preserved plant fossils. It is overlain by eolian sands and evaporites of the Cedar
Mesa Formation (Loope, 1984, 1985, 1988).
The Halgaito Formation in the Valley of the Gods and Monument Valley contains both
plants and non-marine aquatic and terrestrial vertebrate fossils. The vertebrates are found
mainly in carbonate facies and conglomeratic deposits in small channels (Vaughn, 1962; Scott
Figure 8. Setting Hen Butte, Valley of the Gods, southeastern Utah. Stratigraphic succession. Camera
angle causes the apparent overlap in the units. Late Pennsylvanian Honaker Trail Formation.
Pennsylvanian-Permian Halgaito Formation. Early Permian Cedar Mesa Formation.
72
and Sumida, 2004), and are represented by taxa typically found in the Early Permian from
elsewhere in the southwestern and eastern United States. Plants have been known from this
region for many years (Vaughn, 1962, 1973; Mamay and Breed, 1970), though no thorough
paleobotanical study of them has been undertaken. Specifically, in the Valley of the Gods, these
plants are confined, for the most part, to various facies of a widely traceable interval consisting
of channel cut-and-fill features. The flora was predominantly calamitalean sphenopsids and
tree ferns, mainly preserved as trunks of relatively large size (10-15 cm [4-6 in] in diameter)
that had been only locally transported. As with deposits of similar age in New Mexico, the
flora also contained a background of plants typical of seasonally dry substrate environments,
including walchian conifers. This, again, leads us to conclude that the preservation of
parautochthonous, well preserved, typically wetland plants within channel facies strongly
suggests a riparian flora growing along and perhaps within these channels where soil moisture
was high for much of the year. However, the background of drier-site plants suggests that the
landscape outside of the riparian zone was populated by plants tolerant of seasonal drought; this
is consistent with the record of the paleosols, as well, which indicate only subhumid climates at
the wettest (Soreghan et al., 2002). The most unusual thing about these channel and paleosol
deposits is their presence within what otherwise appears to be a succession of arid to hyper-arid
deposits.
Vegetational temporal patterns across the entire western Pangean region parallel those in
areas farther to the east. For example, xeromorphic plants typical of seasonally dry
environments have been reported from the Middle Pennsylvanian in the Illinois coal basin
(Falcon-Lang et al., 2009; Plotnick et al., 2009), and recently a callipterid has been reported
from the Middle Pennsylvanian in this basin, as well (Pšenička et al., 2011). Conifer pollen has
been found to dominate assemblages from the late Middle Pennsylvanian in Atlantic Canada
(Dolby et al., 2011). These patterns indicate that seasonally dry, subhumid climates alternated
on glacial-interglacial scales (Milankovitch) with intervals of humid climate that characterized
times of peat formation. At a still larger scale, the early Late Pennsylvanian (Missourian/
Kasimovian) has been characterized as a time of relatively high sea-level and significantly
greater climatic dryness across the western and west central Pangean tropics, including Nevada
(Bishop et al., 2010), New Mexico (Falcon-Lang et al., 2011a), Utah (Soreghan et al., 2002),
and the Applachian Basin (Cecil et al., 1985; Cecil, 1990; Joeckel, 1995; Martino and Blake,
2001; Martino, 2004; Greb et al., 2008). Wetter conditions appear to have returned in the later
Late Pennsylvanian (Virgilian/Gzhelian), indicated by the mixed floras and even a rare coal in
New Mexico (Krainer and Lucas, 2004), and major coal beds in the Appalachians (e.g., Cecil et
al., 1985; Cecil, 1990; Ruppert et al., 1999) with associated typically wetland Pennsylvanian
floras (Blake et al., 2002). By the Permian, those floras that are known from the western
Pangean regions are uniformly composed of species typical of seasonally dry habitats, with
local patches of wetland plants surviving in those parts of landscapes that had semi-permanent,
probably ground-water-based, soil moisture (see papers cited above, plus Lucas et al., in press
b).
Moving East: North-central Texas
The Eastern Shelf of the Midland Basin of north-central Texas preserves one of the
“classic” Permian red beds successions in North America (Figure 9). This area has been the
source of numerous studies, mainly in Early Permian vertebrate paleontology, which brought
73
the terrestrial geology of the area into focus (e.g., Romer, 1935; Olson, 1952, Sander, 1989;
Nelson et al., 2001; Montañez et al., 2007). Various aspects of paleobotany also have been
investigated, mainly systematics (e.g., White, 1912; Mamay, 1967, 1968, 1976, 1986), but also
stratigraphy (Read and Mamay, 1964). The Pennsylvanian portion of the section has received
less paleontological acclaim, but has been studied in terms of its geology (Moore and Plummer,
1922; Brown, 1967; Feray, 1967; Galloway and Brown, 1973; Hentz, 1988; Tabor and
Montañez, 2004) and paleobotany (e.g., Gupta, 1977; DiMichele et al., 2005a). This area was
in the southwestern portion of equatorial Pangea during the Pennsylvanian and Permian. As
such, the area was subject to atmospheric circulation patterns similar to those affecting areas
farther to the paleo-west, in New Mexico and Utah. During the Virgilian (Gzhelian) the wetter
parts of climate cycles, presumably driven by glacial-interglacial cyclicity, were usually wet
enough for coals to form, though these were generally of low quality, high in ash, thin, and
limited in areal extent (e.g., Gennett and Ravn, 1993). These coaly facies become thinner, less
widely distributed and ultimately disappear from the stratigraphic succession near the
Pennsylvanian-Permian boundary (Hentz, 1988).
The Pennsylvanian-Permian boundary on the Eastern Shelf as defined by conodonts,
occurs at about the level of the Saddle Creek Limestone Member of the Harpersville Formation,
just below its contact with the Pueblo Formation (Wardlaw, 2005). Farther to the north, where
we have studied the paleobotanical succession, the geological section is predominantly
terrestrial, with different formational names, but not enough marine limestones to place the
Pennsylvanian-Permian boundary accurately with conodont (or any other marine invertebrate)
biostratigraphy. However, it can be projected to be in the upper part of the Markley Formation
(Hentz, 1988).
Fossil plants are preserved across the Pennsylvanian-Permian boundary, occurring in a
variety of facies that preserve both wetland assemblages and those from seasonally dry or better
drained substrates. There is very little mixing of these assemblages, based on study of the
facies distribution of macrofossils, which appear to have occurred at different stages of glacial-
interglacial cycles (DiMichele et al., 2005a). The fossil assemblages are preserved in
abandoned channel deposits on landscapes that are represented primarily by paleosols (clastic
deposits with a strong pedogenic overprint – see Tabor and Montañez, 2004), or channel-form
Figure 9. North-central Texas, Early Permian (Kungurian) red beds deposit, formed in ox-bow lake.
Deposits such as these have yielded both flora and vertebrate fauna.
74
Figure 10. North-central Texas Pennsylvanian and Early Permian age outcrops of the Markley
Formation. A—Typical succession of beds from the latest Pennsylvanian. From base:
paleosol, quartz-kaolinite siltstone (containing seasonally dry flora), organic shale/coal bed
(containing wetland swamp flora), gray siltstones (containing wet floodplain flora), sandstone
(represented by float blocks, rarely containing seasonally dry flora). B—Same location as in
(A), viewed from a distance. Note the prominent white, quartz-kaolinite siltstone and overlying
organic shale bed in the center of the outcrop. C—Earliest Permian plant bearing outcrop,
consisting of floodplain siltstone and channel-form sandstone deposits.
75
sandstones. Plant-bearing deposits occur mainly in association with various coal bed horizons,
across which coals/organic shales are discontinuous and variable in thickness and organic
content.
A typical outcrop in the lower part of the Markley Formation (Figure 10 A, B) has a
distinctive lithological sequence, which is strongly associated with plant biofacies patterns. It
must be stressed that this pattern has been observed at multiple outcrops and throughout the
Pennsylvanian portion of the Markley Formation, through several cycles, marked by distinct
coal-bed horizons. Closer to the Pennsylvanian-Permian boundary, coaly facies are no longer
present in outcrops, but the rest of the facies patterns and plant compositional aspects remain
the same. These patterns are described by DiMichele et al. (2005a), and from base to top of a
typical outcrop they are as follows:
1. A basal paleosol, evidencing wet but well drained conditions, sometimes
containing roots that cannot be attributed to specific plant groups. Paleosols may
be several meters thick and, throughout the section, evidence a drying trend in the
later Pennsylvanian and into the earliest Permian (Tabor and Montañez, 2004).
2. A kaolinite-quartz siltstone bed that may vary in thickness from a few centimeters
to over a meter. This bed generally rests unconformably on the upper surface of the
paleosol. It contains a flora composed of Sphenopteris germanica and walchian
conifers, with rare elements including Charliea-like pinnules and various
pteridosperms. Typically wetland plants have not been identified in this facies.
The flora is characteristic of Pennsylvanian seasonally dry climates.
3. An organic-clastic bed that may vary from an organic-rich, finely laminated clay
shale to a normally bright-banded coal bed. Where plants are found in this facies,
the most common and abundant are typically Macroneuropteris scheuchzeri,
Pecopteris of various species, and Sigillaria brardii. This low diversity
assemblage represents the typical Pennsylvanian clastic-to-peat swamp flora of
flooded substrates with high water tables throughout most of the year. Such beds
may vary from less than a meter in thickness to several meters.
4. A sequence of gray to buff siltstone deposited in shallow scours with erosional
bases. The siltstones commonly contain a weak pedogenic overprint. Where
pedogenesis has not proceeded too far, plant fossils often are preserved in the
channel scours. The flora preserved in the fill of any one small scour can vary
greatly in species richness, but the overall flora of any given outcrop, consisting of
several such subdeposits, is generally the most diverse of any of the plant-bearing
facies, consisting of greater than 30 species. These species are principally tree
ferns and pteridosperms, typical of Pennsylvanian wetland floras.
5. Sandstone, often many meters in thickness (difficult to measure because of talus
formation and vegetational cover) with an erosional basal contact with the
floodplain shales may, on occasion, contain plant fossils. Such sands probably are
at the base of the subsequent cycle but, because of the nature of weathering in the
area, they tend to support hillslopes and occur at the tops of exposed sections.
Only rarely have plants been identified in these rocks. Where they have occurred,
the plants are cordaitaleans and conifers, both xeromorphic, coniferalean taxa that
have been identified as elements of seasonally dry floras. Cordaitaleans are an
extremely diverse group, and appear to occur from coastal environments
76
(Raymond, 1988; Falcon-Lang, 2005; Raymond et al., 2010) into remote interior
areas of tropical latitudes (Falcon-Lang, 2000). However, the appearance of
conifers and cordaitaleans in this channel facies is consistent with the appearance
of such elements, typical of seasonally dry floras, throughout the Middle and Late
Pennsylvanian in coal-bearing rock sequences of the Western Interior basin
(Feldmann et al., 2005), Eastern Interior (Illinois) basin (Falcon-Lang et al., 2009;
Plotnick et al., 2009), Central Appalachian basin (Martino and Blake, 2001) and
the Maritimes area of Canada (Dolby et al., 2011).
The over-riding paleobotanical pattern observed in the Texas deposits is the co-
occurrence, at the outcrop scale, of wetland flora and seasonally dry flora (Figure 11),
throughout the Virgilian and into the earliest parts of the Wolfcampian. Both vegetation types
occupied the coastal plain environment, but they are principally unmixed, occurring in
lithologically distinct beds, in stereotypical patterns of succession. Thus, the degree to which
these floras overlapped in space and time continues to be an open question. They were
certainly likely to have been present contemporaneously in the equatorial region, but did they
coexist on lowland landscapes in close spatial proximity, during the latest Pennsylvanian and
earliest Permian, as we have surmised from assemblages in New Mexico and Utah? In the case
of north-central Texas, the evidence is most parsimoniously interpreted, in our opinion, as
indicating they did not share the lowlands contemporaneously. In this area, the wetland plants
do not appear to have been confined to channels or channel belts, surrounded by seasonally dry
flora. Rather, the confinement of the Texas floras to distinctly different lithofacies, which
themselves are of types closely tied to broader indicators of regional climate and sea-level or
local base-level, argues for differentiation in space and time.
In light of the climate-rock framework and models described in the introductory portion of
this paper, we interpret the floristic and lithofacies patterns in north-central Texas as follows.
Coal and organic shale beds reflect the wettest portions of the climate cycles and likely formed
during lowstand to late lowstand on a flat lying coastal plain, developing under moist subhumid
to humid paleoclimate, differentially in lower areas where standing water was most likely to
accumulate for long periods, including channel belts, formed during the earlier phases of sea-
level fall. Overlying these organic beds are floodplain deposits, which frequently are in
gradational contact with the organic facies. We interpret these as the early phases of sea-level
rise, during which climate was still moist subhumid, favorable to development of a wetland
flora. These sections contain few or no limestones. So, these low areas of the landscape may
have been embayed or served as areas of active drainage during highstand, surrounded by
interfluves. From highstand, through falling stages and into early lowstand, paleosols formed
on these floodplain sediments, resulting in deep pedogenesis. In addition, shallow channels
developed across the landscapes, in which sandstone channel bars were preserved. The climate
during these phases, based on the plants, appears to have been seasonal, but still likely dry
subhumid. The kaolinite beds have flow features associated with them, and the contained flora
is, in most instances fragmentary and evidences at least local transport. We interpret these
deposits as representing sediments washed into the channels during early lowstand as rainfall is
increasing on the landscape, under the transition from dry to moist subhumid conditions.
The Pennsylvanian-Permian transitional interval in north-central Texas is marked by a
strong floristic discontinuity (DiMichele et al., 2010c). Coal beds become progressively thinner
and localized in development prior to disappearing entirely in the latest Pennsylvanian,
77
suggesting a diminishment of moisture in the wettest phases of glacial-interglacial cycles. In
the Permian, there is a change to much simplified lithological patterns on outcrop, and organic
beds are lacking (Figure 10 C). Channel sediments, where plants have been identified, are
dominated entirely by a flora that resembles that typical of the kaolinite-quartz beds of the
Pennsylvanian outcrops, that is: Sphenopteris germanica, walchian conifers, and various
pteridosperms, but with an admixture of callipterids (such as Autunia conferta) and other
seasonally dry elements such as cordaitaleans and Taeniopteris. This kind of flora remains the
most commonly encountered throughout the Archer City Formation, which lies immediately
above the Markley and, together with the upper 10 % or so of Markley thickness, is primarily
Asselian and Sakmarian in age (Montañez et al., 2007), a time interpreted to be one of major
glaciation in the southern hemisphere (Fielding et al., 2008). Montañez et al., (2007)
Figure 11. North-central Texas, seasonally dry flora from the Pennsylvanian-Permian Markley and Early
Permian Archer City formations. A. Rhachiphyllum cf. schenkii, callipterid peltasperm frond segment, Early
Permian, floodplain siltstone (USNM 536518; USNM loc 40037). B. Sphenopteris germanica, quartz-kaolinite
siltstone bed (USNM 543965; USGS loc 9998). C. Walchian conifer, Early Permian ox-bow lake deposit
(USNM 543966; USNM loc 40027). D. Walchian conifer, quartz-kaolinite siltstone bed (USNM 543968; USNM
loc 39998).
78
demonstrated strong vegetational tracking of climate changes and dynamics associated with this
and subsequent intervals of inferred glaciation and the intervening non-glacial periods.
However, we hasten to point out that significant unconformities indicative of prolonged sea
level fall associated with long-term, abundant continental ice have not been identified in this
region or on the North American craton in general.
Palynology of the north-central Texas section through the Pennsylvanian-Permian
boundary (Gupta, 1977) suggests a more gradual turnover than that found in the macrofossils
(leading Gupta to place the palynological boundary well above that for either fusilinids or
conodonts, at the time his paper was written). These palynological data suggest that the wetter
and drier elements of the flora were in reasonably close proximity on the Eastern Shelf, even
though this is not supported by the macrofossils, which do not show the patterns of wet-flora/
dry-flora co-occurrence seen in the deposits from the more westerly regions of the equatorial
belt that suggest spatial adjacency.
In summary, the vegetational transition in north-central Texas, based on macrofossils,
appears to be detectable on the temporal scales of both glacial-interglacial oscillations and on
the longer scale of the general Pennsylvanian-Permian drying trend. However, the changeover
to a more consistently sampled seasonally dry flora is superficially later in Texas, and to the
north in Oklahoma (see below), than in New Mexico and Utah, appearing not to have occurred
until very near to the Pennsylvanian-Permian boundary. This may, in no small part, reflect a
taphonomic sampling bias: wetland deposits, and their associated flora, which may have
represented a much smaller temporal fraction of lowland basin occupancy than that of
seasonally dry flora (Falcon-Lang et al., 2009; Falcon-Lang and DiMichele, 2010), are more
abundant and more easily located by collectors than the seasonally dry flora (often represented
either by paleosols or by scattered deposits with poorly preserved fossils), therefore giving the
appearance of greater concentration of wetlands on the landscape in time and space than is
warranted by fact. Floral samples in the overlying Archer City Formation of Texas, which
lacks indicators of climates as wet as or wetter than moist subhumid, are rare and widely
distributed in time and space. These floras are typically of the seasonally dry type, which may,
in fact, be those of the wettest climate phases. The Markley Formation actually may record
only the fact that the wetter parts of glacial-interglacial cycles, such as they were in the latest
Pennsylvanian, were simply wetter and more capable of generating widespread wetland
landscapes than were similarly much drier wettest phases in equivalent age rocks of New
Mexico and Utah. If this is the case, the longer term drying trend in these two areas is less
conspicuously offset than would appear, though it is still present.
Northern Oklahoma and Southern Kansas
Understanding of Late Pennsylvanian and Early Permian paleobotany of north-central
Oklahoma and adjacent south-central Kansas derives primarily from palynology. This
transition occurs in rocks of the latest Pennsylvanian through earliest Permian Wabunsee,
Admire, Council Grove, and Chase groups. Macrofossils, although recorded, have not been
collected or studied systematically or in their stratigraphic context. In Oklahoma, foliage has
been found in cores from the Wood Siding Formation (Wabunsee Group, latest Virgilian) and
the Winfield Limestone (Chase Group, Early Permian) (Lupia, 2010), and also has been
reported in outcrop in the Doyle Shale (Chase Group) (Chaplin 2010).
The Council Grove and Chase groups, of latest Pennsylvanian through Early Permian age,
79
contain alternating shales and limestones associated with epicontinental sea-level fluctuations,
set against a pattern of overall drying through the Early Permian (West et al., 2010). The
Pennsylvanian-Permian boundary is placed in the Red Eagle Limestone of the Council Grove
Group on the basis of conodonts (Sawin et al., 2006). Abundant evidence of cyclicity
consistent with glacial-interglacial drivers has been documented across the Penn-Permian
boundary by Olszewski and Patzkowsky (2003), who conclude that eustatic lows were intervals
of dry climates whereas eustatic highs were intervals of humid climates, but this may reflect
terminological differences from the interpretations presented here. They also note that thin, but
persistent coals occur above and below the boundary. Mazullo et al. (2007) also concluded that
depositional sequences in these rocks reflect glacial-interglacial drivers, based on examination
of isotopic evidence from brachiopods, which they inferred to track changes in ice-volume
rather than local ocean temperature.
Previously, two comprehensive palynological sampling programs have studied the
Pennsylvanian-Permian boundary interval in this area. Clendening (1970, 1975) investigated
Kansas palynology for comparison with his Dunkard studies. Wilson and Rashid (1971)
likewise sampled Virgilian through Lower Permian sediments to establish the boundary. In
both cases, they concluded that the vegetational change occurred far above the conodont-based
boundary assignment in the Red Eagle Limestone, placing it within the lower Chase Group or at
its top, respectively. Although lacking in the resolution necessary to detect oscillations, both
suggest that Early Permian vegetation was substantially similar to the underlying Late
Pennsylvanian. More recently this interval has been reinvestigated (Lupia, 2010), specifically
in the context of the conodont-based boundary, and has affirmed these prior large-scale findings
as well as observed smaller-scale oscillations in palynological content across the boundary,
without a sharp vegetational discontinuity. Rather, wetter-drier phase oscillations are
superimposed on a longer-term, coarser scale “drying” trend. The abundance of striate
bisaccate pollen, indicative of taxa from seasonally dry habitats (e.g., conifers and peltasperms),
is negatively correlated with the abundance of spores representative of wetland taxa (e.g., ferns,
horsetails). In the course of these analyses, considerable lateral variation in lithologies was
found, suggesting differences in local conditions, not unlike the regional variations found in the
Dunkard Group (see below; Cecil et al., 2011; Fedorko and Skema, 2011). One core showed a
Virgilian section that was overall quite dark grey and organic-rich, whereas another, from less
than 32 km (20 mi) away, showed dominantly red bed shales in equivalent Virgilian strata
assuming correct correlation.
The Central Appalachian Basin – Dunkard Group
The Dunkard Group, of the central Appalachian Basin, was located in the western central
portion of Pangea during the Pennsylvanian-Permian transition. In this position, it was well
inland and backed up against the early phases of tectonic uplift of the Appalachians, which was
the source of most Dunkard sediment (Martin, 1998). The occurrence of linguloids in the
Washington coal zone (Cross and Schemel, 1956; Berryhill, 1963) suggests that the basin may
have had a drainage to the open ocean, but the lack of unequivocally marine fossils in any
Dunkard strata strongly suggests that marine waters never actually entered the basin. In fact,
the last certain occurrence of marine conditions in the Applachians is the early Virgilian
(Gzhelian) Ames Limestone, which has a date securely established by conodonts (Barrick et al.,
2008), and lies approximately 100 meters below the base of the Dunkard Group. Dunkard
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strata are dominantly fine-grained, clay to silt-sized clastic sediments with high mica and low
feldspar content, perhaps suggesting a distant source or one poor in such minerals. Common to
abundant non-marine limestones, vertic to calcic-vertic, deeply developed paleosols, and
evidence of fluctuating water tables indicate a prevalence of subhumid climates, moist and dry,
with some periods of greater aridity (Cecil et al., 2011). Coals, indicators of humid climate, are
thickest in the lower parts of the Dunkard, where the Waynesburg and Washington coals have
been mined commercially. However, above the Washington coal zone, coals tend to be thin
and often discontinuous (Fedorko and Skema, 2011). These coals are high in ash and sulfur
content (Eble et al., 2011), consistent with their formation under climatic conditions that
favored topogenous peat swamps, barely within the climatic and sedimentological window
favorable for peat formation (Cecil et al., 1985; Cecil et al., 2011).
Dunkard sedimentation has been described as “cyclic” (Beerbower, 1961, see also
Fedorko and Skema, 2011). These patterns are most detectable in the northern portions of the
Dunkard Group where facies diversity is highest. On the southern and western margins of the
basin, the section consists mainly of interbedded channel sandstones/siltstones and paleosols
(Martin, 1998; Fedorko and Skema, 2011). In the deeper portions of the basin, the lithofacies
included in this cyclicity indicate the oscillation of regional climate extremes: at the wettest end
are the coal beds, indicative of widespread, high-soil-moisture conditions, probably under moist
subhumid to humid conditions. At the other extreme are non-marine limestones, indicative of
high evapotranspiration and a range of conditions varying from dry subhumid to semi-arid to
occasionally arid. Also alternating are conditions that favored widespread vertic paleosol
development, often with calcareous deposits, reflective of a strongly seasonal distribution of
moisture under subhumid climatic conditions, and deposits of sandstone, siltstone, and
mudstone, the grain sizes and geometries of which indicate deposition in standing water bodies
of considerable areal extent (Martin, 1998; Cecil et al., 2011; Fedorko and Skema, 2011).
It can be concluded from these observations that the Dunkard Basin was subject to cyclic
variations in water table and atmospheric moisture delivery on several different temporal scales.
Even at its “wettest” (Waynesburg and Washington coals) there were still intervals of moisture
fluctuation and drought of variable duration, sometimes enough to interrupt peat formation and
return the region temporarily to high levels of clastic transport (see Cecil and Dulong, 2003, for
a discussion of the relationship between climate and sediment transport). Even at its “driest”,
the region appears to have been wet enough to support fluvial siliciclastic sediment transport,
reflected in features such as frequent clastic partings within limestone, high clastic content of
some limestones (making them almost calcareous shales) and the occasional preservation of
plants in the limestones themselves, indicative of wet-substrate conditions (specimens of the
tree fern foliage Pecopteris have been found within the Windy Gap Limestone of the upper
Greene Formation).
The Dunkard flora is well established and is characterized overwhelmingly by
assemblages of typically Late Pennsylvanian (late Virgilian: Gzhelian) character (see Gillespie
and Pfefferkorn, 1979; Blake et al., 2002), typical of wetland habitats. This was clear even
from the illustrated macroflora of Fontaine and White (1880), the original description of the
flora, from Gillespie et al. (1975), who reinterpreted the flora to some degree, and from the
commentaries of White (1936) and Darrah (1969, 1975). Palynological analyses (Clendening,
1972, 1974, 1975; Clendening and Gillespie, 1972) present a similar picture, though they are
focused primarily on those rocks representative of the wettest portions of climate cycles (Eble
81
et al., 2011), the coals and coal roof shales.
In our own collecting, we have found
wetland species, particularly pecopterid
ferns, in nearly every facies (except clastic
paleosols), including limestones,
sandstones, the latter representing both high
-water deltaic deposits (flat-bottomed sands)
and erosionally based channel-form deposits
representing low stand or falling stages of
base-level, floodplain mudstones, and organic-rich swamp deposits.
The exception to the wetland Dunkard flora is the occurrence of callipterid
peltasperms and conifers in some floodplain mudstones and clastic partings in limestones,
occurring sporadically (at current levels of stratigraphic resolution) from the Washington coal
zone, in the Washington Formation, to the level of the Nineveh coal in the middle Greene
Formation (DiMichele et al., 2011) (Figures 12 & 13). These are exceptional occurrences, of
which perhaps as many as 10 have been reported (see Darrah, 1975). There appear to be few
other taxa in the callipterid assemblages. Most significantly, Darrah (1969, 1975) reports
conifers, including what he identified as Lebachia, from at least one site. No illustration was
provided, and the location of Darrah’s collections is unknown to us. In the David White
collections made in 1902, and those of Aureal Cross, made in the late 1940s, other taxa,
including the pteridosperm Odontopteris, calamitalean stems and Annularia foliage, are rare.
However, most of the collections are small and it is not certain that the non-callipterid material
was collected from the same beds as the callipterids (based on examination of the matrix and on
what can be gleaned from surviving field notes and notes in collections).
The spatio-temporal distinctiveness of the wetland and callipterid floras most likely
reflects environmental control. In keeping with the more widely known and documented
occurrences of callipterids (e.g., Kerp and Fichter, 1985; Read and Mamay, 1964) and of
peltaspermous seed plants in general (DiMichele et al., 2005b), these plants are likely
representative of seasonality of moisture distribution with periods of soil moisture deficit,
probably in subhumid climatic regimes. Numerous calcic paleosols, throughout the Dunkard,
document climatic intervals with a probable ustic soil moisture regime indicative of an intense
dry season for most of the year (Cecil, et al, 2011). And for most such soils, the surface
vegetation is not known – it might have been a callipterid-conifer assemblage.
Poor exposure of Dunkard strata (Figure 12) limits our understanding of the environmental
context of both the common wetland and rare seasonally dry floras, but much more so the latter.
Wetland floras are represented by many more collections and collecting sites than are
seasonally dry floras, which provides a “statistically” richer picture of the context of the
wetland vegetation, despite the generally limited outcrop exposure. However, the cyclicity of
Dunkard lithotypes, and environments they represent (Beerbower, 1961), suggest that the
callipterids may (1) characterize one part of any given climate-deposystem cycle, and (2) that
there was an increase in climate contrasts in the Dunkard beginning near the time of deposition
Figure 12. Northern West Virginia, Brown’s Bridge
locality from which the original Fontaine and
White (1899) callipterids were collected. Site of
railroad grade. Modern exposure is heavily
vegetated.
82
Figure 13. Callipterids from Brown’s Bridge locality, West Virginia. Collected by David White, 1902. These
species, typical of seasonally dry habitats, were one of the major reasons that Fontaine and White inferred
a Permian age for the Dunkard Group. A—Autunia conferta (USNM 543957; USGS loc 2926), 4 X
magnification. B—Lodevia oxydata (USNM 543958; USGS loc 2926), 3X magnification.
83
of the Waynesburg A coal bed, compared to the oldest portions of the Dunkard and the
underlying Monongahela Formation. This climate contrast continued at least through
deposition of the lower to middle Greene Formation. Evidence of this increasing contrast
between the wetter and drier portions of climate-deposystem cycles is seen as early as the time
of formation of the paleosol seat-earth that lies beneath the Waynesburg A coal bed, which is
locally a thick, well developed, calcic Vertisol, evidencing strong moisture seasonality with
periods of high evapotranspiration. This paleosol is succeeded by a return to humid conditions
with the formation of the peat that formed the Waynesburg A coal bed. Paleosols of this type
have been found, intermittently, well up into the Greene Formation (such as the exposures on
Great House Hill Road, near Wylieville, West Virginia, which is stratigraphically near the
position of the Lower Rockport Limestone: Fedorko and Skema, 2011).
There is no definitive marine or radioisotopic evidence by which the Pennsylvanian-
Permian boundary can be located within the Dunkard. There are, however, strongly suggestive
paleontological and lithological data suggesting that it may lie close to the level of the
Washington coal. Such evidence include the occurrence of brackish fauna at that horizon
(Berryhill, 1963), consistent with a latest Pennsylvanian marine high stand (Davydov et al.,
2010) at that time, non-marine ostracode data (Tibert et al., in press; Tibert, 2011) and tetrapod
vertebrate data (Lucas, 2011), both of which suggest a Permian age, possibly beginning around
the time of deposition of the Washington coal complex. Thus, the interval within which
callipterids appear is, indeed, not an unreasonable candidate for placement of a systemic
boundary. Such a placement would be consistent with the early thinking of David White
(1904), based on plant fossils (though he later joined I.C. White in considering the entire
Dunkard Group to be Permian – White, 1936). That upward diminishment of coaly, or organic-
rich, facies throughout the Greene Formation is a similar pattern to that of the other American
basins we have discussed. In the Dunkard, the organic facies persist longer than in the basins
from the more western regions of the Pangean equatorial belt. This pattern suggests that the
wetter end of climate cycles remained moist subhumid to humid longer in a progressive west-to
-east direction. As a consequence, a wetland flora remains the most commonly encountered
throughout the Dunkard section, even into the Early Permian.
DISCUSSION
There are clear trends in the spatial and temporal patterns of vegetational change across
the Pennsylvanian-Permian boundary of the Pangean equatorial region in North America.
Wetland species, characteristic of humid and moist subhumid climates, are the most commonly
encountered in the terrestrial fossil record of most of the Pennsylvanian, a pattern that has long
been recognized. Taxa typical of seasonally dry environments, those not typically associated
with organic-rich deposits or in association with physical indicators of humid climates, are
characteristically the most commonly encountered plants of Permian age rocks. This pattern
appears to be time transgressive. It begins first in the western equatorial regions of Pangea and
occurs progressively later in time along the paleoequator to the east, at least on the northern and
western side of the Central Pangean mountain belt. As in the Dunkard Group, strata in Western
Europe, such as the Lower Rotliegends, which transgress the Pennsylvanian-Permian boundary,
have numerous organic-rich deposits, often well into what is interpreted as the Lower Permian,
depending on the particular basin and its tectonic setting (see Roscher and Schneider, 2006).
These may preserve intercalated wetland and seasonally dry floras, with elements of seasonally
84
dry floras appearing in the Pennsylvanian and, conversely, wetland elements persisting into the
Permian (e.g., Remy, 1975; Broutin, 1977; Wagner and Martinez-García, 1982; Wagner, 1983;
Kerp and Fichter, 1985; Jerzykiewicz, 1987; Kerp et al., 1990; Popa, 1999; Steyer et al., 2000;
Cassinis and Ronchi, 2002; Wagner and Mayoral, 2007; Boyarina, 2010).
Underlying the average vegetational trend is the concept of a progressive “drying” from
the Pennsylvanian into the Permian, which is correct at a coarse resolution (see summary in
Remy, 1975). However, as resolution is increased, it can be seen that this general drying trend
contains a great deal of finer scale climate variation, including that likely to be orbitally forced
on the temporal scales encompassed by Milankovitch cyclicity (e.g., Heckel, 2008; Connolly
and Stanton, 1992; Rasbury et al., 1998; “stage scale” of Cecil et al., 2011), consistent with
glacial-interglacial cycles of the 105-year scale. And, on longer time frames, such as that
evaluated by Allen et al. (2011), these Milankovich-scale fluctuations in climate can be seen to
be superimposed on broader fluctuations in moisture that range in duration from 2 to 6 million
years, corresponding broadly to fluctuations in southern hemisphere ice volume (Fielding et al.,
2008).
At the glacial-interglacial scale of 105 years, increased contrasts can be identified between
the intervals of peat formation (humid climate) and those of paleosol formation below the peat
beds (subhumid climate) (Cecil et al., 2003; DiMichele et al., 2010c). This began at least by the
Middle Pennsylvanian across much of the Pangean tropics west of the Central Pangean
mountain belt, which may have simply been highlands at that time, in what is now North
America (Roscher and Schneider, 2006). It is reflected in the nature of the fossil-plant
assemblages found in association with coals vs. those deposits formed during the drier inter-
coal time intervals (Galtier et al., 1992; Falcon-Lang et al., 2009; Plotnick et al., 2009; Scott et
al., 2010; Falcon-Lang and DiMichele, 2010). Climate contrasts on the glacial-interglacial
scale become even more accentuated during the drier of these intervals, such as the Missourian
(early Late Pennsylvanian), where excellently preserved “Permian”-type floras may be found
and lithologies record strong contrasts between wetter and drier parts of climate-sealevel-
deposystem cycles (Cridland and Morris, 1963; Darrah, 1975; Martino and Blake, 2001;
Martino, 2004; Martino and Greb, 2009; Falcon-Lang et al., 2011a; Allen et al., 2011; Lucas et
al., in press a).
The physical and biological aspects of these inferred glacial-interglacial oscillations in
climate are manifested differently among Pangean depositional basins, depending on regional
climatic means, and this affects our understanding of the floras in those respective areas. For
example, in western Pangea (New Mexico and Utah), even at those times of wettest climate,
peat formation was at best sporadic and generally did not occur. The regional climate system
was shifted to the drier end of the “wet-dry” spectrum. Cyclicities can be recognized in the
western basins at multiple temporal scales, and they follow the same basic patterns as in more
easterly parts of the Pangean tropics: indicators of sea-level change can be identified, there are
fluctuations in terrestrial floristic composition associated with these changes, and longer-term
changes can be identified associated with changing polar ice volume. The “apparency” of
vegetational change, that is our ability to discern such change from the plant fossil record, is,
however, strongly affected by regional climate differentiation. The lack of coal in western
Pangea removes a focal point for both recognition of and preservation of wetland vegetation in
autochthonous and parautochthonous accumulations. As a consequence, the field search images
for locating plant-fossil-bearing deposits are very different in a central Pangean coal basin than
85
in areas of western Pangea, where coaly rocks are rare to absent. In the west, we often find
plants, after considerable searching and excavation, in rock units that preserve mainly
allochthonous material and that formed during portions of the sea-level and climate cycle, such
as deltaic deposits formed at highstand or falling stages of sea level, that are considerably
different from those represented by coals and the sedimentary rocks immediately associated
with them (such as floor and roof shales with which they intergrade). Strikingly, a search for
like kinds of lithotypes in coal basins often reveals the same kinds of non-wetland plants as are
found in western Pangea, and often in allochthonous assemblages (White, 1912; Arnold, 1941;
Cridland and Morris, 1963; Feldmann et al., 2005; Plotnick et al., 2009; Falcon-Lang et al.,
2009; Pšenička et al., 2011).
Major changes in the volume of polar ice occur on temporal scales that are considerably
longer than those driven by orbital cyclicity (Allen et al., 2011). Glacial-interglacial cycles are
superimposed on variations in mean ice volume, and thus on the change in climate state they
cause (Cecil et al., 2011). The larger scale changes do not appear to occur rhythmically, as if
driven by various combinations of orbital forcing factors, but this is presently uncertain. Such
changes appear to have occurred near the Atokan-Desmoinesian (Bolsovian-Asturian, mid-
Moscovian) boundary and the Desmoinesian-Missourian (approximately the Moscovian-
Kasimomvian) boundary, and to have resulted in threshold-like changes in climate state at those
boundaries.
At the Atokan-Desmoinesian boundary, there appears to have been an overall shift toward
greater dryness at all phases of glacial-interglacial cycles. This resulted in a shift from an
oscillation between perhumid climates during peat formation and moist subhumid climates
during the intervening periods, to humid peat-forming climates and dry subhumid climates
between periods of peat formation (Cecil, 1990). The result was a change from raised,
ombrotrophic peat swamps to planar, minerotrophic swamps (e.g., Cecil et al., 1985; Greb et
al., 2002; Eble et al., 2001, 2003), an increase in sulfate and carbonate minerals in the system
(Cecil et al., 1985), changed sedimentary patterns (Cecil, 1990; Bertier et al., 2008), and
changes in the dominant vegetation of wetland intervals (Peppers, 1996; Cleal, 2007).
Major vegetational change also took place at the Desmoinesian-Missourian (Westphalian-
Stephanian in traditional useage, approximately Moscovian-Kasimovian) boundary (Phillips et
al., 1974; DiMichele and Phillips, 1996). This vegetational change is closely correlated with,
and was seemingly caused by, climatic warming and drying that began earlier in the Middle
Pennsylvanian and culminated near the Desmoinesian-Missourian boundary (Cecil, 1990;
Phillips and Peppers, 1984), resulting in great diminishment of polar ice (Fielding et al., 2008;
Allen et al., 2011) and a high-stand of sea level (Rygel et al., 2008; Heckel, 2008). One of the
driving forces of this warming may have been CO2 (Cleal et al., 1999), which, though not
shown definitively for the Desmoinesian-Missourian boundary, has been demonstrated for
similar changes in Permian ice volume by Birgenheier et al. (2010), which may serve as a
model for this older time interval. Similarly, model studies of late Paleozoic atmospheric
composition, ice volume and vegetational feedbacks implicate fluctuations in CO2 as a major
controlling variable (Horton and Poulsen, 2009; Horton et al., 2010). Another exacerbating
factor may have been the rise of the Central Pangean Variscan mountain range across Central
Europe, which both exposed earlier-deposited coals to erosion and reduced dramatically the
area of basinal depocenters (Cleal and Thomas, 2005; Cleal et al., 2009). The Missourian/
Kasimovian warm-dry period transitioned to a seemingly wetter Virgilian/Gzhelian, based on
86
patterns of coal thickness and areal distribution in the central Pangean coal basins (Cecil, 1990),
and from patterns seen in parts of the western US, but this phase in particular, as has been
discussed in this paper, was spatially quite heterogeneous. The changes at the Desmoinesian-
Missourian boundary are instructive when viewed in comparison with Pennsylvanian-Permian
boundary. In both cases there is a trend toward drying and perhaps warming through time.
However, in the Missourian, this warming and drying appears to be correlated with a significant
diminishment in polar ice volume (Fielding et al., 2008) and a distinct global rise in mean sea-
level (Rygel et al., 2008). In stark contrast, there is evidence of a major reappearance of ice
during the Early Permian, the first of several Permian ice intervals (Fielding et al., 2008;
Montañez et al., 2007), though evidence of widespread, long-term sea-level fall has not been
demonstrated. Thus, the changes at these two time periods do not appear to be replicates of the
same basic kind of global event or set of drivers, and the vegetational consequences certainly
were different (DiMichele et al., 2009).
The longest-term climatic trend considered here is that characterizing the general drying
beginning in the later Middle Pennsylvanian and continuing throughout the Permian.
Difficulties in picking a lithological boundary between the Carboniferous and Permian reflect
the gradual, directionally oscillatory, lithological changes that occurred in much of the world
between the two Periods – there is no “natural” break, in most instances (see discussion in
Wardlaw et al., 2004). This lack of a natural lithological break (lacuna) argues against a global
lowstand of sea level, which is consistent with the paleobotanical changes, as described herein.
Whereas there certainly is an “on average” difference between a Pennsylvanian tropical flora
and one from the Permian, there is always the chance of finding a “Permian”-type assemblage
in the Pennsylvanian and vice versa. This has been a point of confusion for many years for
paleobotanically-based stratigraphic frameworks. Consider, for example, the remarks of W.C.
Darrah in the “Discussion” following Remy’s paper (Remy, 1975), regarding the Garnett,
Kansas Missourian flora, which Darrah insisted had to be Permian, or the remarks of Remy
(1975) himself, who could not accept explanations of high-frequency climate changes to
explain the intercalation of wetland and seasonally dry floras within Pennsylvanian and
Permian sections, despite arguing forcefully for strong climatic controls on plant distribution.
Such disagreements have continued to the present (e.g., Wagner and Lyons, 1997; Falcon-Lang
et al., 2011b). Certainly, the recognition of “xeromorphic” or seasonally dry floral elements as
facies fossils, representative of drier climatic settings, has a long history (Elias, 1936, 1970;
Arnold 1941).
Long-term drying is the spatio-temporal scale that is most frequently discussed when
referring to Pennsylvanian-Permian climatic changes (see, for example, Remy, 1975, who
recognizes changes at multiple scales, but focuses strongly on the long-term patterns). It is
often discussed, despite clear awareness of the patterns at more resolved temporal scales, as if
there is a monotonic trend toward drier and warmer climates, accompanied by a simple
vegetational transition. In so doing, an unintended consequence is lack of acknowledgement of
the tremendous complexity of this interval, even in an area confined to the low latitudes,
encompassing climate and sea level changes, changes in continental positions, the rise and
subsequent erosion of mountain ranges, atmospheric circulation patterns and composition, and
the interaction of all of these factors. The effects of this multiplicity of variables, many
changing in concert at one or more scales of resolution, on the distribution of terrestrial
vegetation should be expected to be enormous, and there is no reason to believe that the tropics
87
were not populated by several distinct, barely overlapping species pools, including both plants
and animals, that migrated in and out of lowland, basinal settings where the likelihood of
preservation was greatest.
This raises the issue of “apparency” in vegetational sampling, and what may be described
as a taphonomic megabias (sensu Behrensmeyer and Hook, 1992; Behrensmeyer et al., 2000),
that is a large scale, persistent bias in the basic structure of the fossil record. In this case, we
use “apparency” to refer to what is most likely to be discovered in the course of normal field
activities when prospecting for and collecting plant fossils. For example, there is unambiguous
evidence of seasonally dry vegetation living in basinal lowlands beginning at least by the early
Middle Pennsylvanian (Atokan/Bolsovian/early Moscovian), from every major terrestrial basin
across North America and some in Europe (discussed, with citations, in the preceding
paragraphs), that is, not simply in allochthonous assemblages transported into the basins from
local “uplands”. Yet, that vegetation is poorly represented and any random sampling of likely
host rocks will yield almost invariably a wetland flora. This reflects the strong differences in
the preservational conditions during the wetter times of climate-sea level-deposystem cycles in
comparison to those when seasonally dry climates prevailed – paleosols may be our best
evidence of the plants that once occupied the relatively dry landscapes. Yet, given estimated
times of accumulation of peat vs. the development of thick, calcic vertic paleosols, it is not
unreasonable to believe that seasonally dry floras were the dominant vegetation of basinal
lowlands for extended periods of time during the late Middle and Late Pennsylvanian
(DiMichele et al., 2010c), perhaps even occupying these areas for longer than the wetland
vegetation that is such an iconic representation of the “coal age” (Falcon-Lang and DiMichele,
2010; Dolby et al., 2011). Conversely, there is ample evidence of the survival of wetland floras
well into the Permian, even into the classic Kungurian-age red beds of north-central Texas
(DiMichele et al., 2006; Chaney and DiMichele, 2007), demonstrating continued presence on
the landscapes, even in the drier parts of equatorial western Pangea.
Asynchonous patterns of vegetational change appear, unsurprisingly, to characterize the
Permian itself, as much as the transition from the Pennsylvanian to the Permian. Recent papers
on early occurrences of such plants as gigantopterids (Ricardi-Branco, 2008; Booi et al, 2009a)
and comiods (Booi et al., 2009b), and certain groups of conifers (Looy, 2007) extend the
impression that these plants evolved in tropical envionments and spread out through time, into
western Pangea (e.g., Chaney et al., 2009) and still later into such extra-tropical areas as Angara
(Mamay et al., 2009). Similarly, early occurrences of typically Mesozoic plants (e.g.,
DiMichele et al., 2001; Kerp et al., 2006) indicate that there was a tremendous amount of
evolutionary and ecological dynamics in the terrestrial landscape that either escapes detection
entirely, appears only in brief glimpses as conditions in preservational basins create exceptional
windows of opportunity, or appears millions of years after the actual evolutionary origin of the
plants and their ecological associations. Thus, the rare occurrence of exotic plant fossils may
provide more information about the dynamics of climate than do the abundant wetland floras.
Caution is called for when interpreting the fossil record without understanding, or even just an
awareness, of the local and regional climatic and sedimentological context of a collection.
CONCLUDING REMARKS
The characterization of changes in terrestrial vegetation across the Pennsylvanian-Permian
88
boundary is literally a problem of “seeing the forest for the trees.” One’s understanding of it
depends on the scale or scales of space and time at which one resolves the data empirically. It
also depends on the geographic region with which one is most familiar. Smaller scale, glacial-
interglacial oscillations can be detected in most basins along the paleo-equator. However, the
patterns and timing of the longer-scale trends – the prominence and areal extent of aridity, say,
are spatially patterned and largely time transgressive from west-to-east along the equatorial
region. Additionally, given the approximately 10 My of the Pennsylvanian-Permian transition,
plant evolution also must be factored into paleoecological analyses. The ultimate conclusion
that we offer is not unique, but is similar to that of many who have looked at the Carboniferous-
Permian transition, lithologically, faunistically, or floristically: the end points are distinct, but
between these different terminal conditions lies a long transition interval, probably as long as
either of the better characterized periods. The rationale for a “Dyassic” Period, long ago
discussed as a time interval of gradational change between the last coal beds of the
Carboniferous and gypsum deposits of the Permian, is clear.
ACKNOWLEDGMENTS
Support for field investigations, provided by the National Museum of Natural History
small grants program and the Smithsonian Institution Endowment Funds, is gratefully
acknowledged.
REFERENCES
Allen, J. P., Fielding, C. R., Gibling, M. R., and Rygel, M. C., 2011, Fluvial response to paleo-
equatorial climate fluctuations during the late Paleozoic ice age: Geological Society of
America Bulletin,v. 123, p. 1524-1538.
Armstrong, A. K., Kottlowski, F. E., Stewart, W. J., Mamet, B. L., Baltzer, E. H., Siemers, W.
T. and Thompson, S., 1979, The Mississippian and Pennsylvanien (Carboniferous)
System in the United States - New Mexico: U.S. Geological Survey Professional Paper
1110-W, p. W1-W27.
Arnold, C. A., 1941, Some Paleozoic plants from central Colorado and their stratigraphic
significance: Contributions, Museum of Paleontology, University of Michigan, v. 6, p.
59-70.
Barrick, J. E., Heckel, P. H., and Boardman, D. R., 2008, Revision of the conodont
Idiognathodus simulator (Ellison, 1941), the marker species for the base of the Late
Pennsylvanian global Gzhelian Stage: Micropaleontology, v. 54, p. 125-137.
Beerbower, J. R., 1961, Origin of cyclothems of the Dunkard Group (Upper Pennsylvanian-
Lower Permian) in Pennsylvania, West Virginia, and Ohio: Geological Society of
America Bulletin, v. 72, p. 1029-1050.
Behrensmeyer, A. K. and Hook, R. W., rapporteurs, 1992, Paleoenvironmental contexts and
taphonomic modes, in Behrensmeyer, A. K., Damuth, J. D., DiMichele, W. A., Potts,
R., Sues, H.-D., and Wing, S. L., eds., Terrestrial Ecosystems Through Time: University
of Chicago Press, Chicago, p. 15-136.
Behrensmeyer, A. K., Kidwell, S. M., and Gastaldo, R. A., 2000, Taphonomy and
paleobiology: Paleobiology, v. 26, p. 103-147.
Berryhill, H. L., 1963, Geology and coal resources of Belmont County, Ohio: U.S. Geological
89
Survey Professional Paper 380, 113 p.
Bertier, P., Swennen, R., Lagrou, D., Laenen, B., and Kemps, R., 2008, Palaeoclimate
controlled diagenesis of the Westphalian C & D fluvial sandstones in the Campine
Basin (north-east Belgium): Sedimentology, v. 55, p. 1375–1417.
Birgenheier, L. P., Frank, T. D., Fielding, C. R., and Rygel, M. C., 2010, Coupled carbon
isotopic and sedimentological records from the Permian system of eastern Australia
reveal the response of atmospheric carbon dioxide to glacial growth and decay during
the late Palaeozoic Ice Age: Palaeogeography, Palaeoclimatology, Palaeoecology, v.
286, p. 178-193.
Bishop, J. W., Montañez, I. P., and Osleger, D. A., 2010, Dynamic Carboniferous climate
change, Arrow Canyon, Nevada: Geosphere, v. 6, p. 1-34.
Blake, Jr., B. M., Cross, A. T., Eble, C. F., Gillespie, W. H., and Pfefferkorn, H. W., 2002,
Selected plant megafossils from the Carboniferous of the Appalachian region, United
States, in Hills, L., Henderson, C. M., and Bamber, E.W., eds., Carboniferous and
Permian of the World: Canadian Society of Petroleum Geologists Memoir 19, p. 259-
335.
Bohacs, K. M. and Suter, J., 1997, Sequence stratigraphic distribution of coaly rocks;
fundamental controls and paralic examples: American Association of Petroleum
Geologists Bulletin, v. 81, p. 1612-1639.
Booi, M., van Waveren, I. M., and van Konijnenburg-van Cittert, J. H. A., 2009a, The Jambi
gigantopterids and their place in gigantopterid classification: Botanical Journal of the
Linnean Society, v. 161, p. 302–328.
Booi, M., van Waveren, I. M., and van Konijnenburg-van Cittert, J. H. A., 2009b, Comia and
Rhachiphyllum from the early Permian of Sumatra, Indonesia: Review of Palaeobotany
and Palynology, v. 156, p. 418-435.
Boyarina, N., 2010, Late Gzhelian pteridosperms with callipterid foliage of the Donets Basin,
Ukraine: Acta Palaeontologica Polonica, v. 55, p. 343-359.
Boyce, C. K. and Lee, J.-E., 2010, An exceptional role for flowering plant physiology in the
expansion of tropical rainforests and biodiversity: Proceedings of the Royal Society B,
v. 277, p. 3437-3443.
Boyce, C. K., Lee, J.-E., Feild, T. S., Brodribb, T. J. and Zwieniecki, M. A., 2010, Angiosperms
helped put the rain in the rainforests: The impact of plant physiological evolution on
tropical biodiversity: Annals of the Missouri Botanical Garden, v. 97, p. 527-540.
Broutin, J., 1977, Nouvelles données sur la flore des basins autuno-stéphaniens des environs de
Guadalcanal (Province de Séville, Espagne): Cuadernos Geología Ibérica, v. 4, p. 91-
98.
Broutin, J., Doubinger, J., Farjanel, G., Freytet, F., Kerp, H., Langiaux, J., Lebreton, M.-L.,
Sebban, S., and Satta, S., 1990, Le renouvellement des flores au passage Carbonifère
Permien: Approaches stratigraphique, biologique, sédimentologique: Comptes Rendus
de l’Académie des Sciences Paris, v. 311, p. 1563-1569.
Brown, L. F., 1967, Virgil-Lower Wolfcamp repetitive depositional environments in north-
central Texas, in Elam, J. C. and Chuber, S., eds., Cyclic Sedimentation in the Permian
Basin: West Texas Geological Society Publication 56, p. 115-134.
Bush, M. B. and Flenley, J. R., 2007, Tropical Rainforest Responses to Climate Change:
Springer/Praxis, Chichester, UK, 440 p.
Bush, M. B., Miller, M. C., De Oliveira, P. E., and Colinvauz, P. A., 2002, Orbital forcing
90
signal in sediments of two Amazonian lakes: Journal of Paleolimnology, v. 27, p. 341-
352.
Cassinis, G. and Ronchi, A., 2002, The (late-) Post-Variscan continental success of Sardinia:
Rendiconti della Società Paleontologica Italiana 1, p. 77-92.
Cecil, C. B., 1990, Paleoclimate controls on stratigraphic repetition of chemical and siliciclastic
rocks: Geology, v. 18, p. 533–536.
Cecil, C. B. and Dulong, F. T., 2003, Precipitation models for sediment supply in warm
climates: SEPM Special Paper 77, p. 21-27.
Cecil, C. B., Stanton, R. W., Neuzil, S. G., Dulong, F. T., Ruppert, L. F., and Pierce, B. S.,
1985, Paleoclimate controls on late paleozoic sedimentation and peat formation in the
central appalachian basin (U.S.A.): International Journal of Coal Geology, v. 5, p. 195-
230.
Cecil, C. B., Dulong, F. T., West, R. C., Stamm, R., Wardlaw, B. A., and Edgar, N. T., 2003,
Climate controls on stratigraphy of a Middle Pennsylvanian cyclothem in North
America: SEPM Special Paper 77, p. 151-180.
Cecil, C. B., DiMichele, W., Fedorko, N., and Skema, V., 2011, Autocyclic and allocyclic
controls on the origin of the Dunkard Group, in Harper, J. A., ed., Geology of the
Pennsylvanian-Permian in the Dunkard basin. Guidebook, 76th Annual Field
Conference of Pennsylvania Geologists, Washington, PA, p. 26-45.
Chaney, D. S. and DiMichele, W. A., 2007, Paleobotany of the classic redbeds (Clear Fork
Group – Early Permian) of north central Texas, in Wong, T. E., ed., Proceedings of the
XVth International Congress on Carboniferous and Permian Stratigraphy, p. 357-366.
Chaney, D. S., Mamay, S. H., DiMichele, W. A., and Kerp, H., 2009, Auritifolia gen. nov.,
probable seed plant foliage with comioid affinities from the Early Permian of Texas,
U.S.A: International Journal of Plant Sciences, v. 170, p. 247-266.
Chaplin, J. R., 2010, Stratigraphic analysis of the Permian Chase Group in northern Oklahoma–
outcrop analogs of reservoir rocks in the Hugoton Embayment of northwestern
Oklahoma and southwestern Kansas: Oklahoma Geological Survey, Guidebook 36, 111
p.
Cleal, C. J., 2007, The Westphalian–Stephanian macrofloral record from the South Wales
Coalfield, UK: Geological Magazine, v. 144, p. 465-486.
Cleal, C. J. and Thomas, B. A., 2005, Palaeozoic tropical rainforests and their effect on global
climates: Is the past the key to the present? Geobiology, v. 3, p. 13-31.
Cleal, C. J., James, R. M., and Zodrow, E. L., , Variation in stomatal density in the Late
Carbonifereous gymnosperm frond, Neuropteris ovata: Palaios, v. 14, p. 180-185.
Cleal, C. J., Opluštil, S., Thomas, B. A., and Tenchov, Y., eds., 2009, Late Moscovian
terrestrial biotas and palaeoenvironments of Variscan Euramerica: Netherlands Journal
of Geosciences, v. 88, p. 181-278.
Clendening, J. A., 1970, Sporological evidence on the geological age of the Dunkard strata in
the Appalachian Basin: Unpublished Ph.D. dissertation, West Virginia University, 453
p.
Clendening, J. A., 1972, Stratigraphic placement of the Dunkard according to palynological
assemblages: Castanea, v. 37, p. 258-287.
Clendening, J. A., 1974, Palynological evidence for a Pennsylvanian age assignment of the
Dunkard Group in the Appalachian basin: Part 2: West Virginia Geological and
Economic Survey, Coal Geology Bulletin 3, 105 pp.
91
Clendening, J. A., 1975, Palynological evidence for a Pennsylvanian age assignment of the
Dunkard Group in the Appalachian basin: Part 1, in Barlow, J. A. and Burkhammer, S.,
eds., 1975, Proceedings of the first I. C. White Symposium “The age of the Dunkard:”
West Virginia Geological and Economic Survey, Morgantown, WV, p. 195-216.
Clendening, J. A. and Gillespie, W. H., 1972, Stratigraphic placement of the Dunkard – A
review of the paleobotanical and other evidence: Castanea, v. 37, p. 26-48.
Condon, S. M., 1997, Geology of the Pennsylvanian and Permian Cutler Group and Permian
Kaibab Limestone in the Paradox basin, southeastern Utah and southwestern Colorado:
U.S. Geological Survey Bulletin 2000-P, 46 p.
Connolly, W.M. and Stanton, Jr., R. J., 1992, Interbasinal cyclostratigraphic correlation of
Milankovitch band transgressive-regressive cycles: Correlation of Desmoinesian-
Missourian strata between southeastern Arizona and the midcontinent of North
America: Geology, v. 20, p. 999-1002.
Cridland, A. A. and Morris, J. E., 1963, Taeniopteris, Walchia, and Dichophyllum, in the
Pennsylvanian System of Kansas: University of Kansas Science Bulletin 44, p. 71-85.
Cross, A. T. and Schemel, M. P., 1956, Geology of the Ohio River Valley in West Virginia, in
Geology and Economic Resources of the Ohio River Valley in West Virginia. West
Virginia Geological and Economic Survey XXII: 1-149.
Darrah, W. C., 1969, A critical review of the Upper Pennsylvanian floras of eastern United
States with notes on the Mazon Creek flora of Illinois. Privately printed, Gettysburg,
PA, 220 p.
Darrah, W. C., 1975, Historical aspects of the Permian flora of Fontaine and White, in Barlow,
J. A. and Burkhammer, S, eds., 1975, Proceedings of the First I. C. White Symposium,
“The Age of the Dunkard”: West Virginia Geological and Economic Survey,
Morgantown, WV, p. 81-99.
Davydov, V.I., Crowley, J.L., Schmitz, M.D. and Poletaev, V.I. 2010. High precision U-Pb
zircon age calibration of the global Carboniferous time scale and Milankovitch-band
cyclicity in the Donets Basin, eastern Ukraine. Geochemistry, Geophysics, Geosystems
11: 1–22.
DiMichele, W. A. and Aronson, R. B., 1992, The Pennsylvanian–Permian vegetational
transition: A terrestrial analogue to the onshore-offshore hypothesis: Evolution, v. 46,
p. 807–824.
DiMichele, W. A. and Chaney, D. S., 2005, Pennsylvanian-Permian fossil floras from the
Cutler Group, Cañon del Cobre and Arroyo del Agua areas in northern New Mexico:
New Mexico Museum of Natural History and Science Bulletin 31, p. 26-33.
DiMichele, W. A. and Phillips, T. L., 1996, Climate change, plant extinctions and vegetational
recovery during the Middle–Late Pennsylvanian Transition: The case of tropical peat-
forming environments in North America, in Hart, M. L., ed., Biotic Recovery from
Mass Extinctions: Geological Society of London, Special Publication 102, p. 201–221.
DiMichele, W. A., Mamay, S. H., Chaney, D. S., Hook, R. W., and Nelson, W. J., 2001, An
Early Permian flora with Late Permian and Mesozoic affinities from north-central
Texas: Journal of Paleontology, v. 75, p. 449-460.
DiMichele, W. A., Chaney, D. S., Kerp, J. H., and Lucas, S. G., 2010a. Late Pennsylvanian
floras in Western Equatorial Pangea, Cañon del Cobre, New Mexico: New Mexico
Museum of Natural History and Science Bulletin 49, p. 75-113.
DiMichele, W. A., Tabor, N. J., and Chaney, D. S., 2005a, Outcrop-scale environmental
92
heterogeneity and vegetational complexity in the Permo-Carboniferous Markley
Formation of North Central Texas, in Lucas, S. G. and Zeigler, K. E., eds., The
Nonmarine Permian: New Mexico Museum of Natural History and Science Bulletin 30,
p. 60-66.
DiMichele, W. A., Kerp, H., Krings, M., and Chaney, D. S., 2005b, The Permian peltasperm
radiation: Evidence from the southwestern United States, in Lucas, S. G. and Zeigler,
K. E., eds., The Nonmarine Permian: New Mexico Museum of Natural History and
Science Bulletin 30, p. 67-79.
DiMichele, W.A., Tabor, N.J., Chaney, D.S. and Nelson, W.J. 2006. From wetlands to wet
spots: Environmental tracking and the fate of Carboniferous elements in Early Permian
tropical floras. Geological Society of America Special Papers 399: 223-248.
DiMichele, W.A., Chaney, D.S., Nelson, W.J., Lucas, S.G., Looy, C.V., Quick, K. and Wang J.
2007. A low diversity, seasonal tropical landscape dominated by conifers and
peltasperms: Early Permian Abo Formation, New Mexico. Review of Palaeobotany and
Palynology 145: 249-273.
DiMichele, W. A., Kerp, H., Tabor, N., and Looy, C. V., 2008, The so-called ‘‘Paleophytic-
Mesophytic’’ transition in equatorial Pangea: Multiple biomes and vegetational
tracking of climate change through geological time: Palaeogeography,
Palaeoclimatology, Palaeoecology, v. 268, p. 152–163.
DiMichele, W. A., Montañez, I. P., Poulsen, C. J., and Tabor, N. J., 2009, Climate and
vegetational regime shifts in the late Paleozoic ice age earth: Geobiology, v. 7, p. 200-
226.
DiMichele, W. A., Chaney, D. S., Kerp, H., and Lucas, S. G., 2010a, Late Pennsylvanian floras
in western equatorial Pangea, Cañon del Cobre, New Mexico: New Mexico Museum of
Natural History and Science Bulletin 49, p. 75-113.
DiMichele, W. A., Falcon-Lang, H. J., Chaney, D. S., Nelson, W. J., and Lucas, S. G., 2010b,
Impact of Late Pennsylvanian (Early Kasimovian) global warming was greater on
Wetland than Dryland tropical floras: Geological Society of America, Abstracts with
Programs, v. 42, no. 5, p. 193.
DiMichele, W. A., Cecil, C. B., Montañez, I. P., and Falcon-Lang, H. J., 2010c, Cyclic changes
in Pennsylvanian paleoclimate and effects on floristic dynamics in tropical Pangaea:
International Journal of Coal Geology, v. 83, p. 329-344.
DiMichele, W. A., Blake, Jr., B. M., Cecil, B., Fedorko, N., Kerp, H., and Skema, V., 2011,
Callipterids of the Dunkard Group of the Appalachian basin: Their identity and
paleoenvironmental significance, in Harper, J. A., ed., Geology of the Pennsylvanian-
Permian in the Dunkard basin. Guidebook, 76th Annual Field Conference of
Pennsylvania Geologists, Washington, PA, p. 144-166..
Dolby, G., Falcon-Lang, H. J., and Gibling, M. R., 2011, A conifer-dominated palynological
assemblage from Pennsylvanian (late Moscovian) alluvial drylands in Atlantic Canada:
Implications for the vegetation of tropical lowlands during glacial phases: Journal of the
Geological Society, v. 168, p. 571-584.
Dubiel, R. F., Huntoon, J. E., Condon, S. M., and Stanesco, J. D., 1996, Permian deposystems,
paleogeography, and paleoclimate of the Paradox Basin and vicinity, in Longman, M.
W. and Sonnenfeld, M. D., eds., Paleozoic Systems of the Rocky Mountain Region:
SEPM (Society for Sedimentary Geology), Rocky Mountain Section, p. 427–444.
Eble, C. F., Greb, S. F., and Williams, D. A. 2001, The geology and palynology of Lower and
93
Middle Pennsylvanian strata in the Western Kentucky Coal Field: International Journal
of Coal Geology, v. 47, p. 189-206.
Eble, C. F., Pierce, B. S., and Grady W. C., 2003, Palynology, petrography and geochemistry of
the Sewickley coal bed (Monongahela Group, Late Pennsylvanian), northern
Appalachian basin, USA: International Journal of Coal Geology, v. 55, p. 187-204.
Eble, C. F., Grady, W. C., and Blake, B. M., 2011, Compositional characteristics of Dunkard
Group coal beds: Palynology, coal petrography and geochemistry, in Harper, J. A., ed.,
Geology of the Pennsylvanian-Permian in the Dunkard basin: Guidebook, 76th Annual
Field Conference of Pennsylvania Geologists, Washington, PA, p. 46-59.
Elias, M. K., 1936, Late Paleozoic plants of the midcontinent region as indicators of time and of
environment: Report of the XVI International Geological Congress (Washington, D.C.,
1933), v. 1, p. 691-700.
Elias, M. K., 1970, Progress in correlation of Carboniferous rocks: Compte Rendu, 6th
International Carboniferous Congress, v. 2, p. 695-715.
Elrick, S. D. and Nelson, W. J., 2010, Facies relationships of the Middle Pennsylvanian
Springfield Coal and Dykersburg Shale: Constraints on sedimentation, development of
coal splits and climate change during transgression: Geological Society of America
Abstracts with Programs, v. 42, no. 2, p. 51.
Elrick, M. and Scott, L. A., 2010, Carbon and oxygen isotope evidence for high-frequency (104
-105 yr) and My-scale glacio-eustasy in Middle Pennsylvanian cyclic carbonates (Gray
Mesa Formation), central New Mexico: Palaeogeography, Palaeoclimatology,
Palaeoecology, v. 285, p. 307-320.
Eros, J. M., Montañez, I. P. Osleger, D., Davydov, V. Nemyrovska, T., Poletaev, V., and
Zhykalyak, M., in press, Upper Carboniferous sequence stratigraphy and relative sea
level history, Donets Basin, Ukraine: Palaeogeography, Palaeoclimatology,
Palaeoecology.
Falcon-Lang, H. J., 2000, Upland ecology of some Late Carboniferous cordaitalean trees from
Nova Scotia and England: Palaeogeography, Palaeoclimatology, Palaeoecology, v. 156,
p. 225-242.
Falcon-Lang, H. J., 2004, Pennsylvanian tropical rainforests responded to glacial-interglacial
rhythms: Geology, v. 32, p. 689–692.
Falcon-Lang, H. J., 2005, Small cordaitalean trees in a marine-influenced coastal habitat in the
Pennsylvanian Joggins Formation, Nova Scotia: Journal of the Geological Society, v.
162, p. 485-500.
Falcon-Lang, H. J. and DiMichele, W. A., 2010, What happened to the coal forests during
Pennsylvanian glacial phases? Palaios, v. 25, p. 611–617.
Falcon-Lang, H. J., Nelson, J., Elrick, S., Looy, C., Ames, P., and DiMichele, W. A., 2009,
Incised channel-fills containing conifers imply that seasonally-dry vegetation dominated
Pennsylvanian tropical lowlands: Geology 37: 923–926.
Falcon-Lang, H. J., Jud, N. A., Nelson, W. J., DiMichele, W. A., Chaney, D. S., and Lucas, S.
G., 2011a, Pennsylvanian coniferopsid forests in sabkha facies reveal the nature of
seasonal tropical biome: Geology, v. 39, p. 371-375.
Falcon-Lang, H. J., Heckel, P. H., DiMichele, W. A., Blake, Jr., B. M., Easterday, C. R., Eble,
C. F., Elrick, S., Gastaldo, R. A., Greb, S. F., Martino, R. L., Nelson, W. J., Pfefferkorn,
H. W., Phillips, T. W., and Rosscoe, S. J., 2011b, No major stratigraphic gap exists near
the Middle-Upper Pennsylvanian (Desmoinesian–Missourian) boundary in North
94
America: Palaios, v. 26, p. 125–139.
Fedorko, N. and Skema, V., 2011, Stratigraphy of the Dunkard Group in West Virginia and
Pennsylvania, in Harper, J. A., ed., Geology of the Pennsylvanian-Permian in the
Dunkard basin. Guidebook, 76th Annual Field Conference of Pennsylvania Geologists,
Washington, PA, p. 1-25.
Feldman, H. R., Franseen, E. K., and Joeckel, R. M., 2005, Impact of longer-term modest
climate shifts on architecture of high-frequency sequences (cyclothems) in the
Pennsylvanian of midcontinent USA: Journal of Sedimentary Research, v. 75, p. 350–
368.
Feray, D. E., 1967, Tectonic versus eustatic control of Pennsylvanian cyclical sedimentation in
north-central Texas, in Elam, J. C. and Chuber, S., eds., Cyclic Sedimentation in the
Permian basin: West Texas Geological Society Publication 56, p. 81-82.
Fielding, C. R., Frank, T. D., and Isbell, J. L., 2008, The late Paleozoic ice age – A review of
current understanding and synthesis of global climate patterns, in Fielding, C. R., Frank,
T. D., and Isbell, J. L., eds., 2008, Resolving the Late Paleozoic Ice Age in Time and
Space: Geological Society of America Special Paper 441, p. 343-354.
Fontaine, W. M. and White, I. C., 1880, The Permian or Upper Carboniferous flora of West
Virginia and S. W. Pennsylvania: Pennsylvania Geological Survey, 2nd ser., Report PP,
143 p.
Galloway, W. F. and Brown, L. F., 1973, Depositional systems and shelf-slope selations on
cratonic basin margin, uppermost Pennsylvanian of North-Central Texas: American
Association of Petroleum Geologists Bulletin, v. 57, p. 1185-1218.
Galtier, J., Scott, A.C., Powell, J. H., Glover, B. W., and Waters, C. N., 1992, Anatomically
preserved conifer-like stems from the Upper Carboniferous of England: Proceedings of
the Royal Society of London B, v. 247, p. 211-214.
Gennett, J. A. and Ravn, R. L., 1993, Palynology of the Upper Pennsylvanian Dalton Coal, Palo
Pinto County, Texas, U.S.A: Palynology, v. 17, p. 115-122.
Gillespie, W. H. and Pfefferkorn, H. W., 1979, Distribution of commonly occurring plant
megafossils in the proposed Pennsylvanian System stratotype, in Englund, K. J., Arndt,
H. H., and Henry, T. W., eds., Proposed Pennsylvanian System Stratotype: Virginia and
West Virginia: Guidebook for Ninth International Congress of Carboniferous
Stratigraphy and Geology, Field Trip No. 1, American Geological Institute, Selected
Guidebook Series No. 1, p. 87-96.
Gillespie, W. H., Hennen, G. J., and Balasco, C., 1975, Plant megafossils from Dunkard strata
in northwestern West Virginia and southwestern Pennsylvania, in Barlow, J. A. and
Burkhammer, S., eds., 1975, Proceedings of the first I. C. White Symposium “The age
of the Dunkard:” West Virginia Geological and Economic Survey, Morgantown, WV, p.
223-248.
Greb, S. F., Eble, C. F., Hower, J. C., and Andrews, W. M., 2002, Multiple-bench architecture
and interpretations of original mire phases – Examples from the Middle Pennsylvanian
of the central Appalachian basin, USA: International Journal of Coal Geology, v. 49, p.
147-175.
Greb, S. F., Pashin, J. C., Martino, R. L., and Eble, C. F., 2008, Appalachian sedimentary
cycles during the Pennsylvanian: Changing influences of sea level, climate and
tectonics, in Fielding, C. R., Frank, T. D., and Isbell, J. L., eds., 2008, Resolving the
Late Paleozoic Ice Age in Time and Space: Geological Society of America Special
95
Paper 441, p. 235-248.
Gupta, S., 1977, Miofloral succession and interpretation of the base of the Permian System in
the eastern shelf of north central Texas, U.S.A: Review of Palaeobotany and Palnology,
v. 24, p. 49-66.
Heckel, P. H., 2008, Pennsylvanian cycothem in midcontinent North America as far-field
effects of waxing and waning of Gondwana ice sheets, in Fielding, C. R., Frank, T. D.,
and Isbell, J. L., eds., 2008, Resolving the Late Paleozoic Ice Age in Time and Space:
Geological Society of America Special Paper 441, p. 275-289.
Hentz, T. F., 1988, Lithostratigraphy and paleoenvironments of Upper Paleozoic continental red
beds, north-central Texas: Bowie (new) and Wichita (revised) Groups: Bureau of
Economic Geology, Report of Investigations 170, p. 1-49.
Hite, R. J. and Buckner, D. H., 1981, Stratigraphic correlations, facies concepts, and cyclicity in
Pennsylvanian rocks of the Paradox basin, in Weigand, D. L., ed., Geology of the
Paradox basin: Rocky Mountain Association of Geologists, Denver, CO, p. 147–159.
Horton, D. E. and Poulsen, C. J., 2009, Paradox of late Paleozoic glacioeustacy: Geology, v.
37, p. 715-718.
Horton, D. E., Poulsen, C. J., and Pollard D., 2010, Influence of high-latitude vegetation
feedbacks on late Palaeozoic glacial cycles: Nature Geoscience, v. 3, p. 572-577.
Jerzykiewicz, J., 1987, Latest Carboniferous (Stephanian) and Early Permian (Autunia)
palynological assemblages from the Intrasudetic basin, southwestern Poland:
Palynology, v. 11, p. 117-131.
Joeckel, R. M., 1995, Paleosols below the Ames marine unit (Upper Pennsylvanian,
Conemaugh Group) in the Appalachian Basin, U.S.A.; Variability on an ancient
depositional landscape: Journal of Sedimentary Research, v. 65, p. 393-407.
Kerp, H., 1996, Post-Variscan late Palaeozoic northern hemisphere gymnosperms: The onset to
the Mesozoic: Review of Palaeontology and Palynology, v. 90, p. 263-285.
Kerp, H. and Fichter, J., 1985, Die Makrofloren des saarpfälzischen Rotliegenden (? Ober-
Karbon – Unter-Perm; SW-Deutschland): Mainzer. Geowiss. Mitt., v. 14, p. 159-286.
Kerp, J. H. F., Poort, R. J., Swinkels, H. A. J. M., and Verwer, R., 1990, Aspects of Permian
palaeobotany and palynology: IX. Conifer-dominated Rotliegend floras from the Saar-
Nahe basin (? Late Carboniferous-Early Permian; SW-Germany) with special reference
to the reproductive biology of early conifers: Review of Palaeobotany and Palynology,
v. 62, p. 205-248.
Kerp, H., Abu Hamad, A., Vörding, B., and Bandel, K., 2006, Typical Triassic Gondwanan
floral elements in the Upper Permian of the paleotropics: Geology, v. 34, p. 265-268.
Kessler, J. L. P., Soreghan, G. S., and Wacker, H. J., 2001, Equatorial aridity in western
Pangea: Lower Permian loessite and dolomitic paleosols in northeastern New Mexico,
U.S.A.: Journal of Sedimentary Research, v. 71, p. 817-832.
Krainer, K. and Lucas, S. G., 2004, The Upper Pennsylvanian Red Tanks Member of the
Bursum Formation at Carrizo Arroyo, Central New Mexico: Transition from shallow
marine to nonmarine facies: New Mexico Museum of Natural History and Science
Bulletin 25, p, 53-69.
Krainer, K. and Lucas, S. G., 2009, Cyclic sedimentation of the Upper Pennsylvanian (Lower
Wolfcampian) Bursum Formation, Central New Mexico: Tectonics versus
glacioeustacy: New Mexico Geological Society, 60th Field Conference, Geology of the
Chupadera Mesa Region, p. 167-182.
96
Krainer, K. and Lucas, S, G., 2010, Sedimentology of the Pennsylvanian-Permian Cutler Group
and Lower Permian Abo Formation, northern New Mexico: New Mexico Museum of
Natural History and Science Bulletin 49, p. 25-36.
Krainer, K., Lucas, S. G., and Vachard, D., 2009, The Upper Carboniferous Bursum Formation
at Abo Pass, Socorro County, New Mexico: New Mexico Geological Society, 60th Field
Conference, Geology of the Chupadera Mesa Region, p. 98-102.
Krainer, K., Lucas, S.G., and Vachard, D., 2011, Lectostratotype section of the Pennsylvanian
Sandia Formation, Bernalillo County, New Mexico: New Mexico Museum of Natural
History and Science Bulletin 53, p. 71-85.
Kues, B. S. and Giles, K. A., 2004, The late Paleozoic ancestral Rocky Mountains system in
New Mexico, in Mack, G. H. and Giles, K.A., eds., The geology of New Mexico. A
Geologic History: New Mexico Geological Society, Socorro, NM, p. 95-136.
Lerner, A. J., Lucas, S. G., Spielmann, J. A., Krainer, K., DiMichele, W. A., Chaney, D. S.,
Ivanov, A., and Schneider, J., 2009, The biota and paleoecology of the Upper
Pennsylvanian (Missourian) Tinajas Arroyo locality, Socorro County, New Mexico:
New Mexico Geological Society, 60th Field Conference, Geology of the Chupadera
Mesa Region, p. 267-280.
Loope, D. B., 1984, The origin of upper Paleozoic sandstones, southeastern Utah: Journal of
Sedimentary Research, v. 54, p. 563-580.
Loope, D. B., 1985, Episodic deposition and preservation of eolian sands: A late Paleozoic
example from southeastern Utah: Geology, v. 13, p. 73-76.
Loope, D. B., 1988, Rhizoliths in acient eolianites: Sedimentary Geology, v. 56, p. 301-314.
Looy, C., 2007, Extending the range of derived Late Paleozoic conifers: Lebowskia gen. nov.
(Majonicaceae): International Journal of Plant Sciences, v. 168, p. 957-972.
Lucas, S. G., 2011, Tetrapod fossils and the age of the Upper Paleozoic Dunkard Group,
Pennsylvania-West Virginia-Ohio, in Harper, J. A., ed., Geology of the Pennsylvanian-
Permian in the Dunkard basin. Guidebook, 76th Annual Field Conference of
Pennsylvania Geologists, Washington, PA, p. 172-178.
Lucas, S. G. and Krainer, K., 2004, The Red Tanks Member of the Bursum Formation in the
Lucero Uplift and regional stratigraphy of the Bursum Formation in New Mexico: New
Mexico Museum of Natural History and Science Bulletin 25, p. 43-52.
Lucas, S. G. and Zeigler, K. E., 2004, Permian stratigraphy in the Lucero Uplift, Central New
Mexico: New Mexico Museum of Natural History and Science Bulletin 25, p. 71-82.
Lucas, S. G., Anderson, O. J., Heckert, A. B., and Hunt, A. P., 1995, Geology of Early Permian
tracksites, Robledo Mountains, south-central New Mexico, in Lucas, S. G. and Heckert,
A. B., eds., Early Permian Footprints and Facies: New Mexico Museum of Natural
History and Science Bulletin 6, p. 13-32.
Lucas, S. G., Krainer, K., and Barrick, J., 2009a, Pennsylvanian stratigraphy and conodont
biostratigraphy in the Cerros de Amado, Socorro County, New Mexico: New Mexico
Geological Society, 60th Field Conference, Geology of the Chupadera Mesa Region, p.
183-212.
Lucas, S. G., DiMichele, W. A., Krainer, K., Chaney, D. S., and Spielmann, J. A., 2009b, A
coal-measure forest near Socorro, New Mexico: New Mexico Geological Society
Guidebook, 60th Field Conference, Geology of the Chupadera Mesa Region, p. 235-
242.
Lucas, S. G., Krainer, K., and Spielmann, J. A., 2010a, Summary of geology of Cañon del
97
Cobre, Rio Arriba County, New Mexico: New Mexico Museum of Natural History and
Science Bulletin 49, p. 15-24.
Lucas, S. G., Harris, S. K., Spielmann, J. A., Rinehart, L. F., Berman, D. S., Henrici, A. C., and
Krainer, K., 2010b, Vertebrate paleontology, biostratigraphy and biochronology of the
Pennsylvanian-Permian Cutler Group, Cañon del Cobre, northern New Mexico: New
Mexico Museum of Natural History and Science Bulletin 49, p. 115-124.
Lucas, S. G., Allen B. D., Krainer, K., Barrick, J., Vachard, D., Schneider, J. W., DiMichele,
W. A., and Bashforth, A. R., in press a, Precise age and biostratigraphic significance of
the Kinney Brick Quarry Lagerstätte, Pennsylvanian of New Mexico, USA:
Stratigraphy.
Lucas, S. G., Harris, S. K., Speilmann, J. A., Berman, D. S., Henrici, A., Krainer, K., Rinehart,
L. F., DiMichele, W. A., Chaney, D. S., and Kerp, H., in press b, Lithostratigraphy,
paleontology, biostratigraphy and age of the Upper Paleozoic Abo Formation near
Jemez Springs, northern New Mexico, USA: Bulletin of the Carnegie Museum of
Natural History.
Lupia, R., 2010, Palynology of the Pennsylvanian-Permian transition, Wood Siding (Virgilian)
to Wellington (Leonardian) Formations, in North-Central Oklahoma: Geological
Society of America, Abstracts with Programs, v. 42, no. 5, p. 193.
Mack, G. H., 2007, Sequence stratigraphy of the Lower Permian Abo Member in the Robledo
and Doña Ana Mountains near Las Cruces, New Mexico: New Mexico Geology, v. 29,
p. 3-12.
Mack, G. H., Cole, D. R., Giordano, T. H., Schaal, W. C., and Barcelos, J. H., 1991,
Paleoclimatic controls on stable oxygen and carbon isotopes in caliche of the Abo
Formation (Permian), south-central New Mexico, U.S.A.: Journal of Sedimentary
Petrology, v. 61, p. 458–472.
Mack, G. H., Leeder, M., Perez-Arlucea, M., and Bailey, B. D. J., 2003, Early Permian silt-bed
fluvial sedimentation in the Orogrande basin of the Ancestral Rocky Mountains, New
Mexico, USA.: Sedimentary Geology, v. 160, p. 159-178.
Mack, G. H., Tabor, N. J., and Zollinger, H. J., 2010, Palaeosols and sequence stratigraphy of
the Lower Permian Abo Member, south-central New Mexico, USA: Sedimentology, v.
57, p. 1566-1583.
Mamay, S. H., 1967, Lower Permian plants from the Arroyo Formation in Baylor County, north
-central Texas: U.S. Geological Survey Professional Paper 575-C, p. 120- 126.
Mamay, S. H., 1968, Russelites, new genus, a problematical plant from the Lower Permian of
Texas: U.S. Geological Survey Professional Paper 593-I, p. I1-I15.
Mamay, S. H., 1976, Paleozoic origin of the cycads. U.S. Geological Survey Professional Paper
934, p. 1-48.
Mamay, S. H. ,1981, An unusual new species of Dicranophyllum Grand'Eury from the Virgilian
(Upper Pennsylvanian) of New Mexico, USA: The Palaeobotanist, v. 28-29, p. 86-92.
Mamay, S. H., 1986, New species of Gigantopteridaceae from the Lower Permian of Texas:
Phytologia, v. 61, p. 311- 315.
Mamay, S. H., 1990, Charliea manzanitana, n. gen., n. sp., and other enigmatic parallel-veined
foliar forms from the Upper Pennsylvanian of New Mexico and Texas: American
Journal of Botany, v. 77, p. 858-866.
Mamay, S. H., 1992, Sphenopteridium and Telangiopsis in a Diplopteridium-like association
from the Virgilian (Upper Pennsylvanian) of New Mexico: American Journal of
98
Botany, v. 79, p. 1092-1101.
Mamay, S. H. and Breed, W. J., 1970, Early Permian plants from the Cutler Formation in
Monument Valley, Utah: U.S. Geological Survey Professional Paper 700B, p. 109-117.
Mamay, S. H. and Mapes, G., 1992, Early Virgilian plant megafossils from the Kinney Brick
Company Quarry, Manzanita Mountains, New Mexico, in Zidek, J., ed., Geology and
Paleontology of the Kinney Brick Quarry, Late Pennsylvanian, Central New Mexico:
New Mexico Bureau Mines and Mineral Resources Bulletin 138, p. 61-86.
Mamay, S. H., Chaney, D. S., and DiMichele, W. A., 2009, Comia, a seed plant of
peltaspermous affinity: A brief review of the genus and description of two new species
from the Early Permian (Artinskian) of Texas, C. greggii sp. nov. and C. craddockii sp.
nov: International Journal of Plant Sciences, v. 170, p. 267-282.
Mapes, G. and Gastaldo, R. A., 1986, Upper Paleozoic non-peat accumulating floras, in
Broadhead, T. W., ed., Fossil Land Plants: Notes for a Short Course (sponsored by The
Paleontological Society): University of Tennessee, Department of Geological Sciences,
Studies in Geology 15, p. 115-127.
Martin, W. D., 1998, Geology of the Dunkard Group (Upper Pennsylvanian – Lower Permian)
in Ohio, West Virginia, and Pennsylvania: Ohio Geological Survey Bulletin 73, 49 p.
Martino, R. L., 2004, Sequence stratigraphy of the Glenshaw Formation (Middle–Late
Pennsylvanian) in the central Appalachian basin, in Pashin, J. and Gastaldo, R.A., eds.,
Coal-Bearing Strata: Sequence Stratigraphy, Paleoclimate, and Tectonics: American
Association of Petroleum Geologists Studies in Geology 51, p. 1–28.
Martino, R. L., and Blake, Jr., B. M., 2001, Walchian conifers from the middle Conemaugh
Group (Late Pennsylvanian), Wayne County, West Virginia: West Virginia Academy of
Science Abstracts with Proceedings, v. 73, no. 1, p. 30-31.
Martino, R.L. and Greb, S. F., 2009, Walking trails of the giant terrestrial arthropod
Arthropleura from the Upper Carboniferous of Kentucky: Journal of Paleontology, v.
83, p. 140-146.
Mazzulo, S. J., Boardman, D. R., Grossman, E. L., and Dominick-Wells, K., 2007, Oxygen-
carbon isotope stratigraphy of Upper Carboniferous to Lower Permian marine deposits
in midcontinent U.S.A. (Kansas and NE Oklahoma): Implications for seawater
chemistry and depositional cyclicity: Carbonates and Evaporites, v. 22, p. 55-72.
Montañez, I.P., Tabor, N. J., Niemeier, D., DiMichele, W. A., Frank, T. D., Fielding, C. R.,
Isbell, J. L., Birgenheier, L. P., and Rygel, M. C., 2007, CO2-forced climate and
vegetation instability during late Paleozoic deglaciation: Science, v. 315, p. 87-91.
Moore, R. C. and Plummer, F. B., 1922, Pennsylvanian stratigraphy of North Central Texas:
Journal of Geology, v. 30, p. 18-42.
Nelson, W. J., Hook, R. W., and Tabor, N., 2001, Clear Fork Group (Leonardian. Lower
Permian) of north-central Texas: Oklahoma Geological Survey, Circular 104, p. 167-
169.
Olson, E. C., 1952, The evolution of a Permian vertebrate chronofauna: Evolution, v. 6, p. 181-
196.
Olszewski, T. D. and Patzkowsky, M. E., 2003, From cyclothems to sequences: The record of
eustasy and climate on an icehouse epeiric platform (Pennsylvanian-Permian, North
American midcontinent): Journal of Sedimentary Research, v. 73, p. 15-30.
Parrish, J. T., 1993, Climate of the supercontinent Pangea: Journal of Geology, v. 101, p. 215–
233.
99
Peppers, R. A., 1996, Palynological Correlation of Major Pennsylvanian (Middle and Upper
Carboniferous) Chronostratigraphic Boundaries in the Illinois and Other Basins:
Geological Society of America Memoir 188, 111 p.
Phillips, T. L. and Peppers, R. A., 1984, Changing patterns of Pennsylvanian coal-swamp
vegetation and implications of climatic control on coal occurrence: International
Journal of Coal Geology, v. 3, p. 205-255.
Phillips, T. L., Peppers, R. A., Avcin, M. J., and Laughnan, P. F., 1974, Fossil plants and coal:
Patterns of change in Pennsylvanian coal swamps of the Illinois Basin: Science, v. 184,
p. 1367-1369.
Plotnick, R. E., Kenig, F., Scott, A. C., Glasspool, I. J., Eble, C. F., and Lang, W. J., 2009,
Pennsylvanian paleokarst and cave fills from northern Illinois, USA: A window into
late Carboniferous environments and landscapes: Palaios, v. 24, p. 627-637.
Popa, M. E., 1999, The Early Permian megaflora from the Reşiţa Basin, South Carpathians,
Romania: Acta Palaebotanica, Proceedings of the 5th EPPC, Supplement 2, p. 47-57.
Poulsen, C. J., Pollkard, D., Montanez, I. P., and Rowley, D., 2007, Late Paleozoic tropical
climate response to Gondwanan deglaciation: Geology, v. 35, p. 771–774.
Pšenička, J., Kerp, H., Opluštil, S., Elrick, S., DiMichele, W. A., Bek, J., Nelson, W. J., and
Ames, P. R., 2011, Preliminary report on the earliest callipterid assignable to the
morphogenus Rhachiphyllum Kerp from the late Moscovian (Asturian) Farmington
Shale (Illinois Basin, USA): XVII International Carboniferous and Permian Congress,
Programme and Abstracts, Geological Survey of Western Australia, Record 2011/20, p.
105.
Rasbury, E. T., Hanson, G. N., Meyers, W. J., Holt, W. E., Goldstein, R. H., and Saller, A. H.,
1998, U-Pb dates of paleosols: Constraints on late Paleozoic cycle durations and
boundary ages: Geology, v. 26, p. 403-406.
Raymond, A., 1988, The paleoecology of a coal-ball deposit from the middle Pennsylvanian of
Iowa dominated by cordaitalean gymnosperms: Review of Palaeobotany and
Palynology, v. 53, p. 233-250.
Raymond, A., Lambert, L., Costanza, S., Stone, E.J., and Cutlip, P. C., 2010, Cordaiteans in
paleotropical wetlands: An ecological re-evaluation: International Journal of Coal
Geology, v. 83, p. 248-265.
Read, C. B. and Mamay, S. H., 1964, Upper Paleozoic floral zones and floral provinces of the
United States: U.S. Geological Survey Professional Paper 454-K, 35 p.
Remy, W., 1975, The floral changes at the Carboniferous-Permian boundary in Europe and
North America, in Barlow, J. A. and Burkhammer, S., eds., 1975, Proceedings of the
first I. C. White Symposium “The age of the Dunkard:” West Virginia Geological and
Economic Survey, Morgantown, WV, p. 305-352.
Ricardi-Branco, F., 2008, Venezuelan paleoflora of the Pennsylvanian-Early Permian:
Paleobiogeographical relationships to central and western equatorial Pangea:
Gondwana Research, v. 14, p. 297-305.
Romer, A. S., 1935, Early history of Texas red-beds vertebrates: Geological Society of
America Bulletin, v. 46, p. 1597-1658.
Roscher, M. and Schneider, J. W., 2006, Permo-Carboniferous climate: Early Pennsylvanian to
Late Permian climate development of central Europe in a regional and global context, in
Lucas, S. G., Cassinis, G., and Schneider, J. W., eds., Non-Marine Permian
Biostratigraphy and Biochronology: Geological Society of London, Special
100
Publications 265, p. 95-136.
Rueger, B. F., 1996, Palynology and its relationship to climatically induced depositional cycles
in the middle Pennsylvanian (Desmoinesian) Paradox Formation of southeastern Utah:
U.S. Geological Survey Bulletin 2000-K, p. 1-22.
Ruppert, L. F., Twealt, S. J., Bragg, L. J., and Wallack, R. N., 1999, A digital resource model of
the Upper Pennsylvanian Pittsburgh coal bed, Monongahela Group, northern
Appalachian basin coal region, USA: International Journal of Coal Geology, v. 41, p. 3-
24.
Rygel, M. C., Fielding, C. R., Frank, T. D., and Birgenheier, L. P., 2008, The magnitude of late
Paleozoic glacioeustatic fluctuations: A synthesis: Journal of Sedimentary Research, v.
78, p. 500-511.
Sander, P. M., 1989, Early Permian depositional environments and pond bonebeds in central
Archer County, Texas: Palaeogeography, Palaeoclimatology, Palaeoecology, v. 69, p. 1
-21.
Sawin, R. S., West, R. R., Franseen, E. K., Watney, W. L., and McCauley, J. R., 2006,
Carboniferous-Permian Boundary in Kansas, midcontinent, USA: Kansas Geological
Survey, Bulletin 252, part 2, p. 1-13.
Scott, A. C., Kenig, F., Plotnick, R. E., Glasspool, I. J., Chaloner, W. G., and Eble, C. F., 2010,
Evidence of multiple late Bashkirian to early Moscovian (Pennsylvanian) fire events
preserved in contemporaneous cave fills: Palaeogeography, Palaeoclimatology,
Palaeoecology, v. 291, p. 72-84.
Scott, K. M. and Sumida, S. S., 2004, Permo-Carboniferous vertebrate fossils from the Halgaito
Shale, Cutler Group, southeastern Utah: Geological Society of America, Abstracts with
Programs, v. 36, no. 5, p. 230.
Soreghan, G. S., Elmore, R. D., and Lewchuk, M. T., 2002, Sedimentologic-magnetic record of
western Pangean climate in upper Paleozoic loessite (lower Cutler beds, Utah):
Geological Society of America Bulletin, v. 114, p. 1019-1035.
Steyer, J. S., Escuillie, F., Pouillon, J.-M., Broutin, J., Debriette, P., Freytet, P., Gand, G.,
Poplin, C., Rage, J. C., Schneider, J. W., Stamberg, S., Werneburg, R., and Cuny, G.,
2000, New data on the flora and fauna from the ?uppermost Carboniferous-Lower
Permian of Buxieres-les-Mines, Bourbon l’Archambault Basin (Allier, France); a
preliminary report: Bulletin de la Société Géologique de France, v. 171, p. 239-249.
Tabor, N. J. and Poulsen, C. J., 2008, Paleoclimate across the Late Pennsylvanian-Early
Permian tropical paleolatitudes: A review of climate indicators, their distribution, and
relation to paleophysiographic climate factors: Palaeobiology, Palaeoclimatology,
Palaeoecology, v. 268, p. 293-310.
Tabor, N. J. and Montañez, I. P., 2004, Morphology and distribution of fossil soils in the Permo
-Pennsylvanian Wichita and Bowie Groups, north-central Texas, USA: Implications for
western equatorial Pangean palaeoclimate during icehouse-greenhouse transition:
Sedimentology, v. 51, p. 851-884.
Tabor, N. J, Montañez, I. P., and Southard, R. J., 2002, Paleoenvironmental reconstruction from
chemical and isotopic compositions of Permo-Pennsylvanian pedogenic minerals:
Geochimica et Cosmochimica Acta, v. 66, p. 3093-3107.
Tabor, N. J., Montanez, I. P., Scotese, C. R., Poulsen, C. J., and Mack, G. H., 2008, Paleosol
archives of environmental and climate history in paleotropical western Euramerica
during the latest Pennsylvanian through Early Permian, in Fielding, C. R., Frank, T. D.,
101
and Isbell, J. L., eds., 2008, Resolving the Late Paleozoic Ice Age in Time and Space:
Geological Society of America Special Paper 441, p. 291-304.
Thompson, M. L., 1942, Pennsylvanian System in New Mexico: New Mexico Bureau of Mines
and Mineral Resources Bulletin 17, p. 1-92.
Tibert, N. E., 2011, Ostracode distribution across the Pennsylvanian-Permian boundary interval
in the Dunkard Group of the central Appalachian basin USA, in Harper, J. A., ed.,
Geology of the Pennsylvanian-Permian in the Dunkard basin. Guidebook, 76th Annual
Field Conference of Pennsylvania Geologists, Washington, PA, p. 179-187.
Tibert, N. E., Dewey, C. P., and Skema, V., in press. Taxonomic notes on the nonmarine
Carboniferous and Permian Ostracoda from the Appalachian Basin, USA: Microp
aleontology.
Tidwell, W. D., 1988, A new Upper Pennsylvanian or Lower Permian flora from southeastern
Utah: Brigham Young University Geology Studies, v. 35, p. 33-56.
Tidwell, W. D. and Ash, S. R., 2004, Synopsis of the flora in the Red Tanks Formation, Carrizo
Arroyo, New Mexico: New Mexico Museum of Natural History and Science Bulletin
25, p. 97-103.
Utting, J., Hartkopf-Fröder, C., Lucas, S. G., and Traverse, A., 2004, Palynological
investigation of the Upper Pennsylvanian Red Tanks Member, Bursum Formation,
Carrizo Arroyo, New Mexico, U.S.A.: New Mexico Museum of Natural History and
Science Bulletin 25, p. 89-96.
Vaughn, P. P., 1962, Vertebrates from the Halgaito tongue of the Cutler Formation, Permian of
San Juan County, Utah: Journal of Vertebrate Paleontology, v. 36, p. 529-539.
Vaughn, P. P., 1973, Vertebrates from the Cutler Group of Monument Valley and Vicinity, in
James, H. L., ed., Guidebook of Monument Valley and Vicinity Arizona and Utah:
New Mexico Geological Society, 24th Field Conference, p. 99-105.
Wagner, R. H., 1983, A lower Rotliegend flora from Ayrshire: Scottish Journal of Geology, v.
19, p. 135-155.
Wagner, R. H and Lyons, P. C., 1997, A critical analysis of the higher Pennsylvanian
megafloras of the Appalachian region: Review of Palaeobotany and Palynology, v. 95,
p. 255-283.
Wagner, R. H. and Martinez-García, E., 1982, Description of an Early Permian flora from
Asturias and comments on similar occurrences in the Iberian Peninsula: Trabajos
Geologicos, Universidad de Ovidedo, v. 12, p. 273-287.
Wagner, R. H and Mayoral, E. J., 2007, The Early Permian of Valdeviar in Sevilla province,
SW Spain: Basin history and climatic/palaeogeographic implications: Journal of Iberian
Geology, v. 33, p. 93-124.
Wardlaw, B. R., 2005, Age assignment of the Pennsylvanian-Early Permian succession of
North Central Texas: Permophiles, v. 46, p. 21-22.
Wardlaw, B. R., Davydov, V., and Gradstein, F. M., 2004, The Permian Period, in Gradstein,
F., Ogg, J., and Smith, A., eds., A Geologic Time Scale 2004: Cambridge University
Press, Cambridge, UK, p. 249-270.
West, R. R., Miller, K. B., and Watney, W. L., 2010, The Permian System of Kansas: Kansas
Geological Survey, Bulletin 257, p. 1-82.
White, D., 1904, Permian elements in the Dunkard flora: Geological Society of America
Bulletin, v. 14, p. 538-542.
White, D., 1912, The characters of the fossil plant Gigantopteris Schenk and its occurrence in
102
North America: Proceedings of the United States National Museum, v. 41, p. 493-516.
White, D., 1936, Some features of the Early Permian flora of America: 16th International
Geological Congress (1933) Proceedings, v. 1, p. 679-689.
Wilson, L. R. and Rashid, M. A., 1975, Palynological evidence for a Pennsylvanian age
assignment of the Gearyan Series in Kansas and Oklahoma, in Barlow, J. A. and
Burkhammer, S., eds., 1975, Proceedings of the first I. C. White Symposium “The age
of the Dunkard:” West Virginia Geological and Economic Survey, Morgantown, WV, p.
183.
