CO2-forced climate and vegetation instability during Late Paleozoic deglaciation.

Abstract

The late Paleozoic deglaciation is the vegetated Earth's only recorded icehouse-to-greenhouse transition, yet the climate dynamics remain enigmatic. By using the stable isotopic compositions of soil-formed minerals, fossil-plant matter, and shallow-water brachiopods, we estimated atmospheric partial pressure of carbon dioxide (pCO2) and tropical marine surface temperatures during this climate transition. Comparison to southern Gondwanan glacial records documents covariance between inferred shifts in pCO2, temperature, and ice volume consistent with greenhouse gas forcing of climate. Major restructuring of paleotropical flora in western Euramerica occurred in step with climate and pCO2 shifts, illustrating the biotic impact associated with past CO2-forced turnover to a permanent ice-free world.

Full text

Page 1

Earth and Atmospheric Sciences, Department of
Papers in the Earth and Atmospheric Sciences
University of Nebraska - LincolnYear 
CO2-Forced Climate and Vegetation
Instability During Late Paleozoic
Deglaciation
Isabel P. Monta˜ nez∗
William A. DiMichele∗∗
John L. Isbell§
Neil J. Tabor†
Tracy D. Frank††
Lauren P. Birgenheier¶
Deb Niemeier‡
Christopher R. Fielding‡‡
Michael C. Rygel?
∗University of California, Davis, montanez@geology.ucdavis.edu
†Southern Methodist University, Dallas, TX
‡University of California, Davis
∗∗Smithsonian Museum of Natural History, Washington, DC
††University of Nebraska at Lincoln, tfrank2@unl.edu
‡‡University of Nebraska-Lincoln, cfielding2@unl.edu
§University of Wisconsin, Milwaukee
¶University of Nebraska-Lincoln
?University of Nebraska-Lincoln
This paper is posted at DigitalCommons@University of Nebraska - Lincoln.
http://digitalcommons.unl.edu/geosciencefacpub/104

Page 2

A
mospheric partial pressure of CO2 (pCO2)
and surface temperatures with changing
global ice volume (1, 2). Although the pre-
cise mechanistic link between atmospheric
greenhouse gases and climate is debated,
there remains little doubt that high concen-
trations of atmospheric CO2 have strongly
amplified Earth’s past climates. Anthropo-
genic CO2
spheric CO2 to concentrations higher than
at any time in at least the past 650,000 years
and could increase it to more than 2000
parts per million by volume (ppmv) as ac-
decade of studying Pleistocene
ice cores has unequivocally doc-
umented a strong coupling of at-
emissions have increased atmo-
cessible fossil fuel reservoirs are exhausted
(3). The last time such concentrations were
seen on Earth was at the onset of our mod-
ern icehouse [~40 to 34 million years ago
(Ma)], a transition from ice-free to glacial
conditions characterized by repeated C cy-
cle perturbation, large magnitude changes
in atmospheric pCO2, and major ephemeral
warmings (4, 5). As our climate system de-
parts from the well-studied Pleistocene gla-
cial-interglacial cycles, a deep-time perspec-
tive of pCO2-climate-glaciation linkages is
essential for a fuller understanding of what
may be the Earth’s most epic deglaciation.
We present here the results of a multi-
pronged investigation that provides ev-
idence for significantly changing atmo-
spheric CO2 concentrations and surface
temperatures during a 40-million-year pe-
riod of the late Paleozoic (~305 to 265 Ma),
which encompasses the deterioration of the
most widespread and long-lived icehouse
of the last half-billion years (6). This global
warming event accompanied a permanent
transition to an ice-free world, a condition
that arguably lasted until the current gla-
cial state. These results, when integrated
with a newly emerging glaciation history
for southern Gondwana (7–11), indicate
strong linkages between pCO2, climate, and
ice-mass dynamics during the final stages
of the Late Paleozoic Ice Age (end of LPIA).
Published in Science, vol. 315. no. 5808 (January 5, 2007), pp. 87 - 91; doi: 10.1126/science.1134207 Copyright © 2007 by the American
Association for the Advancement of Science. Used by permission. http://www.sciencemag.org/cgi/content/full/315/5808/87
Submitted August 22, 2006; accepted November 20, 2006.
CO2-Forced Climate and Vegetation Instability
During Late Paleozoic Deglaciation
Isabel P. Montañez,1 Neil J. Tabor,2 Deb Niemeier,3 William A. DiMichele,4 Tracy D. Frank,5
Christopher R. Fielding,5 John L. Isbell,6 Lauren P. Birgenheier,5 and Michael C. Rygel 5
1 Department of Geology, University of California, Davis, CA 95616, USA
2 Department of Geological Sciences, Southern Methodist University, Dallas, TX 75275, USA
3 Department of Civil and Environmental Engineering, University of California, Davis, CA 95616, USA
4 Department of Paleobiology, Smithsonian Museum of Natural History, Washington, DC 20560, USA
5 Department of Geosciences, 214 Bessey Hall, University of Nebraska–Lincoln, Lincoln, NE 68588, USA
6 Department of Geosciences, University of Wisconsin, Post Office Box 413, Milwaukee, WI 53201, USA
Present address for M. C. Rygel: Department of Geology, State University of New York, College at Potsdam,
Potsdam, NY 13676, USA
Corresponding author — I. P. Montañez, email montanez@geology.ucdavis.edu
Abstract
The late Paleozoic deglaciation is the vegetated Earth’s only recorded icehouse-to-greenhouse transition, yet the cli-
mate dynamics remain enigmatic. By using the stable isotopic compositions of soil-formed minerals, fossil-plant mat-
ter, and shallow-water brachiopods, we estimated atmospheric partial pressure of carbon dioxide (pCO2) and trop-
ical marine surface temperatures during this climate transition. Comparison to southern Gondwanan glacial records
documents covariance between inferred shifts in pCO2, temperature, and ice volume consistent with greenhouse
gas forcing of climate. Major restructuring of paleotropical flora in western Euramerica occurred in step with cli-
mate and pCO2 shifts, illustrating the biotic impact associated with past CO2-forced turnover to a permanent ice-
free world.
87

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88 Montañez et al. in Science 315 (2007)
Integration of these climate proxy records
with our newly developed tropical paleo-
botanical records shows repeated climate-
driven ecosystem restructuring in western
paleo-equatorial Euramerica.
The CO2 contents of ancient atmospheres
can be estimated from the carbon stable iso-
tope values (δ13C) of ancient soil-formed car-
bonates and goethites with an uncertainty of
≤ ±500 ppmv (12, 13). These minerals are the
proxy of choice when pCO2
ppmv), whereas the method’s sensitivity de-
creases at lower pCO2 (<800 ppmv) (14, 15).
The precision of pCO2
variable assumptions used for each pCO2
calculation (16), which can be further re-
fined if the δ13C of coexisting organic mat-
ter is available and if quantitative estimates
of paleosoil-respired CO2 content and pale-
otemperatures can be inferred from modern
analogs or independently derived geochem-
ical proxies (15).
To reconstruct atmospheric CO2 during
the end of the LPIA, we measured the δ13C
values of soil-formed calcites (δ13Ccarb) col-
lected from mature, well-drained profiles
from the Eastern Shelf of the Midland Ba-
sin; the Pedregosa, Anadarko, and Paradox
Basins; and the Grand Canyon Embayment
of western paleoequatorial Euramerica (Fig-
ure S1 and table S1) (17). We consider mea-
sured paleosol δ13Ccarb
proxy of soil-water CO2 during formation,
given the lack of evidence for mineral re-
crystallization and overgrowth and their
overall shallow and low-temperature burial
histories (18). Furthermore, we consider the
δ13C of well-preserved fossil plant matter
(δ13Corg) tobea faithful proxy of the C iso-
tope composition of soil-respired CO2 and,
in turn, of atmospheric CO2 (19, 20). Com-
pression and permineralized fossil plants,
cuticles, coal, and charcoal were collected
from mudstone deposits of abandoned flu-
vial channels and floodplains, which are
stratigraphically intercalated (on a sub-10-
m resolution) with carbonate-bearing pa-
leosols (table S2). The use of measured
δ13Corg rather than penecontemporane-
ous marine carbonates as a proxy of atmo-
spheric δ13C reflects a growing appreciation
of local-scale C cycling effects on the δ13C
values of epicontinental marine carbon-
ates (21). The terrestrial δ13Ccarb and δ13Corg
time series have an average sampling in-
terval of <1 million years (My) and de-
fine long-term trends that exhibit system-
atic variability (Figure 1, A and B). That the
long-term δ13Corg trend records first-order
variations in atmospheric δ13C is supported
by its similarity to time-equivalent δ13Corg
records of Permo-Carboniferous coals from
the North China Platform (22) and by a
narrow range, throughout the study area,
in the ratio of intracellular, pi, and atmo-
spheric, pa, partial pressures of CO2 in pa-
leoflora [0.46 to 0.57 ± 0.3 (2 SE)], which
were estimated by using measured δ13Corg
values of fossil plants and δ13Ccarb values
of contemporaneous marine brachiopods
(17). These factors indicate that changes in
is high (≥ 1000
estimates reflects the

values to be a robust
geomorphic or environmental conditions



















































Figure 1. Temporal distribution of carbonate (A) and fossil plant (B) δ13C values used to construct
best estimate of Permo-Carboniferous atmospheric pCO2 (C). Individual points in (A) and (B) are the
average of analyses from suites of contemporaneous paleosols (from 5 to 18) and associated plant lo-
calities (from 3 to 21); “cand p” encompasses all compression and permineralized plant matter, coals,
and charcoals. Vertical bars are ±2 SE around the mean. PDB, Pee Dee belemnite. (B) Solid curve is
three-point weighted running average through samples from the Eastern Shelf, Midland Basin. Gray
band is δ13Corg of Permo-Carboniferous coals from three correlated successions in North China Plat-
form (22). Overlapping δ13Corg trends but different δ13Corg values are interpreted to reflect overall
wetter conditions for the North China Platform relative to western paleoequatorial Euramerica in the
Permian. Data and pCO2 presented on an age model (51) developed for the terrestrial composite sec-
tion by linearly interpolating between known biostratigraphic boundaries. (C) Best estimate of paleo-
pCO2 (black curve) from Monte Carlo simulation of chronostratigraphically well-constrained sample
populations; uncertainty in pCO2 estimates (gray curves) reflects variability in δ13Ccarb and δ13Corg, in-
terpreted to record inter- and intrabasinal variations in soil conditions, vegetation, and climate. Vertical
bars are published goethite-based CO2 estimates from the same set of paleosols (25).

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Co2-ForCed CliMate and Vegetation instability during late PaleozoiC deglaCiation  89






















































in the study area were secondary to atmo-
spheric δ13C in influencing measured fossil-
plant δ13Corg values.
Ranges of paleosoil-respired CO2 con-
tent were inferred from the morphologies of
suites of contemporaneous paleosols (23) by
comparison with modern analogs, address-
ing a major source of uncertainty in previ-
ous applications of the CO2 paleobarome-
ter (table S3) (14, 15). Paleosol temperatures
were inferred from the oxygen and hydro-
gen isotopic compositions of pedogenic
phyllosilicates and Fe oxides obtained from
the same set of paleosols (18, 24). The best
estimate of paleoatmospheric pCO2 was
defined by using Monte Carlo simulation






















































involving 1000 randomly drawn samples
for each variable for each time-location
combination (17). Monte Carlo simulation
uses random sampling techniques to sto-
chastically solve physical process problems,
in this case quantitatively estimating paleo-
pCO2 and the associated uncertainty by in-
tegrating across all of the inferred and mea-
sured input variables.
Modeled CO2 concentrations (Figure 1C
and table S4) define a long-term rise from
an average of present atmospheric levels
(PAL = 280 ppmv) in the earliest Permian to
values of up to 3500 ppmv by the late Early
Permian. A substantial decline in pCO2 into
the early Middle Permian is corroborated by
independently derived goethite-based esti-
mates of Permian pCO2 (25). A short-lived
(~2 My) drop in pCO2 to near PAL, defined
by contemporaneous paleosols, punctuates
the Early Permian rise. Modeled pCO2 sug-
gests that PAL values were limited to the
earliest Permian after latest Carboniferous
levels of up to 1000 ppmv, in accord with
pCO2
(26) and with southern Gondwanan sedi-
mentologic and geochemical evidence for
latest Carboniferous warming (9, 27). Our
record refines the structure of well-estab-
lished pCO2 reconstructions, which indicate
sustained PAL values throughout much of
the Permo-Carboniferous (15, 28, 29). The
higher-frequency oscillations revealed by
this study would be below the temporal res-
olution (5 to 20 My time-averaging) of those
long-term CO2
In order to evaluate the nature of the
CO2-climate relationship, we developed a
time-equivalent record of paleotropical sea-
surface temperatures (SSTs) by using δ18O
values from a global compilation of well-
preserved latest Carboniferous through
Middle Permian tropical shallow-water bra-
chiopods (table S5) (30); brachiopods have
diagenetically resistant, low-Mg calcitic
shells that incorporate oxygen isotopes in
equilibrium with seawater (31). The resid-
ual brachiopod δ18O record (Figure 2A) dis-
plays clear isotopic fluctuations, with inter-
vals of maximum values corresponding to
Permian glacial maxima or marked coolings
in Antarctica and/or Australia (10, 11) and,
to the degree afforded by geochronologic
dates, with the younger periods of inferred
glacial maxima in the Karoo Basin (8, 32),
southern Argentina (9), and Tasmania (33).
Intervals of minimum δ18Ocarb values corre-
spond with independently inferred periods
of marked warming and sea-level rise (7–9,
34)(Figure 2C).
Inferring secular
from δ18Ocarb requires careful consideration
of the compound effects on values of con-
tinental ice volume, local hydrography, and
SST, as well as any vital effects and postde-
positional alteration (31, 35). The eustatic
component in the Permo-Carboniferous
brachiopod δ18O record due to ice volume
variability likely accounts for far less than 2
per mil (‰) of the observed δ18O variation
given reconstructed amplitudes (10 to <100
m) of Permo-Carboniferous glacioeustasy
(10) and an O isotope composition of seawa-
ter (δ18Osw)–sea level relationship of 0.1‰
per 10 m of sea level change (36). The resid-
ual secular δ18Ocarb signal is interpreted to
record changes in temperature, salinity, and
pH. Local hydrographic variations in trop-
ical epicontinental seas would have damp-
ened the magnitude of δ18Ocarb shifts, given
hypothe-sized heightened freshwater dis-
charge to continental shelves (decreased sa-
linity and lowered δ18Osw) during late Pa-
leozoic periods of maximum glaciation, and
increased evaporation (increased salinity
and δ18Osw) during drier, highly seasonal
inferred from marine carbonate δ13C
records.
paleotemperatures
Figure 2. Relationship among Permo-Carboniferous pCO2, climate, and cryosphere. Temporal dis-
tribution of glacial maxima and/or cool periods based on stratigraphic distribution of diamictites,
rhythmites, and dropstone and keel turbate structures in Antarctica and Australian glacigenic depos-
its (10, 11). (A) Three-point weighted running average (blue curve) and ±2 SE (dashed curves) of de-
trended δ18Obrachiopod values binned into 1- to 3-My windows (green triangles). Error bars indicate ±2
SE around the mean δ18Obrachiopod values. (B) Inferred paleotropical SSTs (red interval) (40) are re-
ported as temperature anomalies given the potential effects of local and regional environmental and
diagenetic influences on brachiopod δ18O. Paleo-SST anomalies (relative to 17.5° C) were calculated
from a three-point weighted running average (± 2 SE) through δ18O-based paleotemperature esti-
mates (table S5). Blue curves are best estimate (heavy) and uncertainty (light) of paleo-pCO2.(C)
Relative sea-level curve compiled from (8, 53); distribution of warm intervals, from (7–9) and (34).

Page 5

90 Montañez et al. in Science 315 (2007)



























































glacial minima (36). Moreover, paleo-SSTs
under elevated pCO2 may be under-esti-
mated by up to 2°C, given that lowered sea-
water pH would have shifted δ18Ocarb to less
negative values (38, 39).
The amplitudes of the reconstructed
SSTshifts (40) indicate substantial changes
in the mean state of tropical climate during
the end of the LPIA, with glacial tropical
oceans at least 4° to 7°C cooler than those
of intervening glacial minima (Figure 2B).



























































Inferred periods of elevated tropical SSTs
and pCO2 coincide with independently rec-
ognized intervals of warmer temperate con-
ditions in high-latitude southern Gondwana
(Figure 2C) indicated by the accumulation
of nonglacial sediments, including exten-
sive kaolin and bauxite deposits in Aus-
tralia during peak (Artinskian) warming
and pCO2 (7) and increased faunal diver-
sity in Australia and South America (7, 11,
41). The covariance among inferred shifts
in paleotropical SSTs, pCO2, and variations
in high-latitude Gondwanan glaciation and
climate implies a strong CO2-climate-gla-
ciation linkage during the Permian. Al-
though our coupled records suggest atmo-
spheric CO2 may have played a direct role
in forcing Early to Middle Permian climate
and ice mass stability, a determination of
phase relationships between these parame-
ters is precluded by the uncertainties in the
age models. The inferred variations in trop-
ical SSTs between periods of glacial max-
ima and minima, however, are consistent
with the range predicted by Permian cli-
mate simulations for a change in radiative
CO2 forcing from 1 to 8 PAL (42).
Permo-Carboniferous
blages from western paleoequatorial Eura-
merica archive a mechanistic vegetational
response to late Paleozoic pCO2 and cli-
mate change. Reconstructed plant commu-
nities from the same terrestrial successions
that host the pedogenic mineral-bearing pa-
leosols document major dominance-diver-
sity changes corresponding one-for-one to
inferred changes in paleotropical climate,
pCO2, and glacial extent (Figure 3 and table
S6). Four tropical biomes appear in succes-
sion, composed of increasingly xeromorphic
species, representing progressively more
seasonally moisture-stressed environments.
These biomes are floristically distinct, shar-
ing only a few opportunistic ferns and sphe-
nopsids (43). Typical latest Carboniferous
flora, rich in marattialean ferns, medullosan
pteridosperms, sphenopsids, and sigillarian
lycopsids, was replaced essentially instan-
taneously by one rich in conifers {Walchia
and Ernestiodendron; compare with Brachy-
phyllum (44), callipterids (Rhachiphyllum),
cycadophytes (Russellites), and other seed
plants [Cordaites, Sphenopteridium (45)]}. This
floristic shift is synchronous with an abrupt
continental climate transition from everwet
to semi-arid conditions (Figure 3A), charac-
terized by increased temperatures (18, 24)
and seasonal moisture availability inferred
from paleosol morphologies (23).
Conifers and callipterids diversified in
seasonally dry habitats during the initial
Early Permian (Sakmarian) rise in CO2 and
the warm period of glacial minima, spatially
replacing the tree fern–rich and the pterido-
sperm-rich wetland floras (Figure 3). Tree
fern–rich floras reappeared during wetter,
cooler conditions of the mid-Early Permian
(Artinskian) glaciation, stratigraphically in-
tercalated but not mixed, with conifer-cal-
lipterid floras. These two glacial floras show
limited species overlap and oscillated at the
103- to 105-year scale, reflecting short-lived
pluvials (46). Dramatic floristic changes also
occurred during the cold period at the close
of the Early Permian (Kungurian), with the
migration into lowland basins of unique
seed-plant assemblages not observed again
until the Late Permian (conifers) and the Me-
sozoic (cycads) (47). These temporally suc-
cessive floras tracked climatic conditions and
contained progressively more evolutionarily
plant assem-
Figure 3. Patterns of abundance change in major flora of study area (A and B) and comparison to
independently derived Permo-Carboniferous climate and pCO2 (C). Plants from 49 sampling locali-
ties on the Eastern Shelf, Midland Basin, are rank ordered: 1, rare (occurs in <10% of sampling quad-
rats at any given locality), 2, common (occurs in 10 to 50% of sampling quadrats), and 3, abundant
(occurs in >50% of sampling quadrats). (A) Tree ferns and pteridosperms are hygromorphic and oc-
cur in deposits with sedimentologic and pedogenic indicators of everwet to subhumid seasonal
conditions. Red climate curve for paleoequatorial western Euramerica defined by using soil mois-
ture regimes and degree of seasonality inferred from paleosol morphologies (23); zigzag pattern in-
dicates short-term (103 to 105 year) climate cycles inferred from intervals of polygenetic soils that
exhibit climatically out-of-phase superposition of calcic and argillic horizons. (B) Conifers and pelta-
sperms are xeromorphic and typically are found in association with sedimentologic and pedogenic
indicators of moisture limitation.

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Co2-ForCed CliMate and Vegetation instability during late PaleozoiC deglaCiation  91
advanced lineages. This suggests that evolu-
tionary innovation, the appearance of new
plant body plans, occurred in extrabasinal
areas and was revealed by climate-driven
floral migration into lowland basins.
The history of latest Carboniferous to
Middle Permian climate provides a unique
deep-time perspective on the precarious
balance between icehouse and greenhouse
states during major climate transitions,
which are coupled to changing atmospheric
CO2 content. Maximum expansion of Gond-
wanan continental ice sheets occurred dur-
ing earliest Permian time (10) under the
lowest paleoatmospheric CO2 levels and pa-
leotropical SSTs. Widespread Early Permian
(mid-Sakmarian) collapse of ice sheets (8,
10) coincided with the onset of rising atmo-
spheric CO2 levels, after which time tropi-
cal SSTs and pCO2 rose. Subsequent glacial
influence was restricted to eastern Austra-
lia (6), with resurgent ice masses occurring
during three more episodes (11) of lowered
atmospheric pCO2 before the permanent
transition to an ice-free world (260 Ma). Our
study indicates that ice buildup in Australia
during subsequent cold periods, however,
was progressively less widespread, with the
two youngest glacials generally confined to
local valleys or mountain ice caps along the
polar margin of Australian Gondwana. No-
tably, SSTs and pCO2 did not return to ear-
liest Permian levels during these post-Sak-
marian glacial periods.
Our reconstructed pCO2, paleotemper-
atures, and inferred glacial history depict
an Early Permian atmosphere that system-
atically increased from PAL to levels simi-
lar to those predicted to exist if fossil fuels
are exhausted. Although global-scale degla-
ciation was unrelenting under rising Early
Permian atmospheric CO2, transient peri-
ods of icehouse stability and glacial resur-
gence returned during short-lived intervals
of low pCO2, perhaps until a CO2 threshold
and greenhouse stability precluded the re-
establishment of glacial conditions [com-
pare with (48)]. This late Paleozoic climate
behavior mimics, in reverse, the magni-
tude and temporal scale of atmospheric
CO2 changes and ephemeral warmings that
foreshadowed the transition into our pres-
ent glacial state (4, 5), further documenting
the degree of climate variability, carbon cy-
cle perturbation, and tropical ecosystem re-
structuring that has been associated with
past CO2-forced climate transitions.

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28. R. A. Berner, Z. Kothavala, Am. J. Sci. 301, 182
(2001).
29. C. I. Mora, S. G. Driese, L. A. Colarusso, Science
271, 1105 (1996).
30. Published δ18O data are from biostratigraphi-
cally constrained and diagenetically screened
calcitic brachiopods (31, 35, 49); δ18O values
< –8‰ were excluded. The δ18O values were
detrended to remove the effects of the long-
term linear Phanerozoic trend by removing,
from each data point, the least squares linear
fit calculated using SPlus.
31. H. Mii, E. Grossman, T. E. Yancey, Geol. Soc. Am.
Bull. 111, 960 (1999).
32. B. Bangert, R. Armstrong, H. Stollhofen, V.
Lorenz, J. Afr. Earth Sci. 29, 33 (1999).
33. M. R. Banks, in Earth’s Pre-Pleistocene Glacial
Records, M. J. Hambrey, W. B. Harland, Eds.
(Cambridge Univ. Press, Cambridge, 1981),
pp. 495–501.
34. J. N. J. Visser, Sedimentology 44, 507 (1997).
35. C. Korte, T. Jasper, H. W. Kozur, J. Veizer, Pal-
aeogeogr. Palaeoclimatol. Palaeoecol. 224, 333
(2005).
36. D. P. Schrag et al., Quat. Sci. Rev. 21, 331 (2002).
37. C. B. Cecil et al., in Climate Controls on Stratigra-
phy, C. B. Cecil, T. N. Edgar, Eds. (SEPM Special
Publication 77, Society for Sedimentary Ge-
ology, Tulsa, OK, 2003), pp. 151–180.
38. D. L. Royer, R. A. Berner, I. P. Montañez, N. J. Ta-
bor, D. J. Beerling, Geol. Soc. Am. Today 14, 3
(2004).
39. W. C. Beck, E. L. Grossman, J. W. Morse, Geo-
chim. Cosmochim. Acta 69, 3493 (2005).
40. Detrended δ18O values were binned into 1-
to 3-My windows. Average values (±2 SE)
were translated to a range of paleotropical
SSTs by using a quadratic equation (50) and a
constant (0‰) or variable (–0.5 to + 1.5‰)
δ18Osw. The influence of local and environ-
mental factors and diagenesis on the long-
term δ18Ocarb trend is considered secondary,
given the overlap in individual published da-
tasets and Tethyan and Panthalassan brachio-
pod δ18O values; secondary influences are
likely recorded in the δ18Obrachiopod variability
within contemporaneous populations.
41. O. R. Lopez-Gamundi, in Late Glacial and Post-
glacial Environmental Changes: Quaternary, Car-
boniferous-Permian, and Proterozoic, I. P. Mar-
tini, Ed. (Oxford Univ. Press, Oxford, 1997),
pp. 147–168.
42. A. M. E. Winguth et al., Paleoceanography 17,
1057 (2002).
43. W. A. DiMichele, R. B. Aronson, Evolution 46,
807 (1992).
44. Compare with Brachyphyllum after S. H. Mamay,
U.S. Geol. Surv. Prof. Pap. 575-C (1967), p. 120.
45. Sphenopteridium, after S. H. Mamay, Am. J. Bot.
79, 1092 (1992).
46. W. A. DiMichele, N. J. Tabor, D. S. Chaney, W. J.
Nelson, Geol. Soc. Am. Spec. Pub. 299 (2006),
pp. 223–248.
47. W. A. DiMichele, D. S. Chaney, W. H. Dixon, W. J.
Nelson, R. W. Hook, Palaios 15, 524 (2000).
48. L. R. Kump, Nature 436, 333 (2005).
49. J. Veizer et al., Chem. Geol. 161, 59 (1999).
50. S. T. Kim, J. R. O’Neil, Geochim. Cosmochim. Acta
61, 3461 (1997).
51. A chronostratigraphic framework for the East-
ern Shelf of Texas (23) was used as the basis
for the composite section into which other
records were integrated with use of available
biostratigraphy and lithostratigraphic corre-
lation. Calibration of the Eastern Shelf suc-
cession to the Gradstein et al. 2004 geologic
time scale (52) is based, for the study area,
on the stratigraphic position of the stage
boundaries determined by conodont, fusuli-
nid, and ammonite biostratigraphy.
52. F. M. Gradstein, J. G. Ogg, A. G. Smith, Eds., A
Geologic Time Scale 2004 (Cambridge Univ.
Press, Cambridge, 2005).
53. C. A. Ross, J. R. P. Ross, in Sea-Level Changes:
An Integrated Approach, C. K. Wilgus et al.,
Eds. (SEPM Special Publication 42, Society for
Sedimentary Geology, Tulsa, OK, 1988), pp.
227–247.PP-0126086, and ANT-0440919.
Supporting Material follows: Materials and Methods, Figure S1, and Tables S1 to S6;
it can also be found online @ www.sciencemag.org/cgi/content/full/315/5808/87/DC1

Page 7

SUPPLEMENTAL ONLINE MATERIAL for Montañez et al.
Materials & Methods
Calcite nodules and rhizoliths, phyllosilicates, and Fe-oxide samples were collected
from at least 0.5 m beneath the surfaces of mature paleosols within latest Pennsylvanian
through Middle Permian terrestrial successions shown in Figure S1. We present here only
the methodology for analyses and CO2 estimates not previously published. The
methodology for C isotope analysis of goethites and their application to paleo-pCO2
reconstruction are presented in detail in S1. The methodology for D and O isotope
analysis of phyllosilicates and Fe-oxides and their application to continental paleo-
temperature reconstruction are presented in S2 – S4.
Carbonate
micromorphologies (S5) and no evidence of secondary dissolution and reprecipitation
based on petrographic criteria (S6) were microsampled using a Merchantek automated
microsampler. Approximately 50 µg of carbonate were roasted at 375°C in vacuum for
three hours to remove organics. Samples were analysed on a Fisons Optima IRMS using
an Isocarb common acid bath autocarbonate system at 90°C at the Stan Margolis Stable
Isotope Laboratory, University of California, Davis (UCD). External precision for the
δ13C measurements based on standards and replicates is better than ±0.06‰. Results are
shown in Table S1.
Stable Isotope Analysis: Micritic calcite exhibiting pedogenic
Organic Matter Stable Isotope Analysis: Organic matter samples were hand-picked with
cleaned metal probe and tweezers, cleaned in Nanopupre H2O, and placed into acid-
washed Pyrex test tubes. Samples were treated with 1 N HCl overnight at room
temperature and then rinsed four times with Nanopupre H2O to remove any carbonate and
hydrolysable C. Between 60 and 100 µg of cleaned and dried sample was loaded in tin
boats. C isotope analysis was carried out on an Isoprime isotope ratio mass spectrometer
interfaced to a Carlo Erba elemental analyzer at the Plant Sciences Stable Isotope
Laboratory and Dept. of Geology at UCD. External precision for the δ13C measurements
based on standards and replicates is better than ±0.3‰. Results are shown in Table S2.
Estimating Pi/Pa indices from Fossil Plant Matter: Pi/pa, the ratio of partial pressures
of CO2 in the substomatal cavity of plants and the atmosphere, is an established proxy of
water-use efficiency of plant growth that is influenced by several environmental variables
(S7). We estimated pi/pa ratios using the well-validated theoretical model of Farquhar et
al. (S8) according to:
Plant Δ13C = a + (b – a) X pi/pa,
where a, the fractionation that occurs during stomatal diffusion, is 4.4‰, and b, the
kinetic fractionation that occurs during the fixation of CO2 by Rubisco, is 27‰.
Plant Δ13C was calculated using:
(δ13Catm – δ 13Cplant matter)
Δ13C = ___________________________________
(1 + δ 13Cplant matter/1000)

Page 8

Estimated pi/pa ratios throughout the 40-million year interval define a narrow range
within each depositional basin (0.46 to 0.57 ±0.3 (2 std error of the mean) for the Eastern
Shelf; 0.55 to 0.58 ±0.2 for the Pedregosa Basin; 0.48 to 0.56 ±0.02 for the Anadarko
Basin), regardless of landscape or stratigraphic position. The low values record the
overall dry conditions in the region and post-depositional alteration; their narrow range,
however, is interpreted to record the low degree of geomorphic or environmental
influence on fossil plant δ13Corg values.
Monte Carlo Modeling: To estimate the best-fit and standard deviation curves for paleo-
pCO2 we used Monte Carlo simulation methods, which can be applied to evaluate
statistical estimators with computational algorithms to simulate a population. For each of
the individual time-location combinations (Table S3), data summary statistics were
calculated for each variable used in the Monte Carlo simulation (details of modeling and
error analysis presented in S9). In addition, variables were also assigned a statistical
distribution. For each time-location combination, 1000 samples were drawn for each of
the Monte Carlo variables to generate alternative scenarios. That is, each variable was
assumed to be independent and values were randomly drawn based on assumed normal or
uniform distributions. Together, these scenarios gave a range of possible solutions, some
of which were considered more probable and some less probable. That is, the result is
considered a distribution of results from which summary statistics (e.g., expected values)
can be estimated. Once the sample of 1000 values had been generated for each time-
location combination, individual pCO2 values were calculated along with the overall
mean and standard deviations (Table S4).


Figure S1. Location of sampled successions (triangles) on the Eastern Shelf of the
Midland Basin (north-central TX) and the Pedregosa (south-central NM), southern
Anadarko (south-central OK), Paradox (south-eastern UT), and Grand Canyon
Embayment (north-western AZ). The “gap” intersecting NM and CO corresponds to
inferred distance of shortening from Laramide and Sevier orogenies. The solid black lines

Page 9

delineate the approximate position of the Early Permian coastlines along western
paleoequatorial Pangaea (modified from S10).

Table S1. Pedogenic carbonate δ13C values used for paleo-pCO2 reconstruction.
Samples Age1
(Ma)
δ δ13Cped carb
measured

δ δ13Cped carb
Avg. (±2 s.e.)

Honaker Trail Fm, UT
HT 1A-1
HT 1B-1
HT3-1
HT 3A
HT2
HT1-1
HT1-2
HT1-3
Hermit Shale Fm, Kohl’s Ranch, AZ
N1-1
N1-2
Bursum Fm, Socorro, NM
BF1-1
BF1-2
BF4-1
BF4-2
BF7
upper Halgaito Fm, AZ
HAL 1A
HAL 2
HAL 4A
HAL 7
HAL 9A
HAL 9B
HAL 12
HAL 15
HAL 17
HAL 18
HAL 22
HAL 25B
HAL 25B-B2
HAL 26
HAL 48
HAL 48 R


300.3








300.0


299.8





298.6


















-3.4 (±0.6)








-4.2 (±1.1)


-6.1 (±0.3)





-4.0 (±0.2)


















-2.8
-3.9
-4.0
-3.2
-4.3
-2.4
-3.7
-3.1

-4.8
-3.7

-5.7
-6.7
-5.9
-5.8
-6.1

-4.0
-4.1
-4.2
-3.6
-3.4
-3.5
-4.4
-3.0
-4.5
-4.2
-4.1
-4.1
-4.2
-3.9
-4.7
-4.5

Page 10

Table S1. continued



Samples Age1
(Ma)
δ δ13Cped carb
measured

δ δ13Cped carb
Avg. (±2 s.e.)

Hermit Shale Fm, south of Flagstaff, AZ
S2-1
S2-2
S4-1
S4-2
S4-3
middle Archer City Fm, stratotype (SS5),
nc-TX
AC 27
AC 28
AC 29
AC 31
AC 32
upper Archer City Fm bonebed (SS8), nc-TX
ABBA 79 - 9C
ACBBH/R
ABBA 3C-1
ABBA 3C-2
basal Nocona Fm, coprolite bonebed, nc-TX
NBCX
NBC4-1
NBC4F
middle Nocona Fm, nc-TX
NLAD 99-1A
NLAD 99-1B
NLAD 2C
Abo Fm, Socorro to Las Cruces, NM
A7-1
A9-1
A10
A3-1
A3-2
A4-1
A4-2
A7-2
A8
A9-2
A9-3
upper Hermit Shale Fm, Nash Pt & Fossil
Springs, AZ
NP2-1
NP2-2
HS1
HS5
297.7





-5.9 (±0.6)





-6.6
-6.7
-5.6
-5.2
-5.3

296.0





295.1




292.4



290.0



285.2











-9.3 (±2.1)





-9.4 (±1.0)




-9.5 (±1.0)



-8.6 (±0.8)



-5.6 (±0.6)











-10.1
-10.9
-11.1
-5.3
-9.0

-10.7
-9.9
-8.5
-8.7

-9.6
-8.2
-10.6

-8.3
-8.3
-9.4

-6.7
-6.2
-4.1
-4.5
-4.0
-6.3
-6.3
-6.0
-5.0
-6.1
-5.9

285.1




-5.1 (±1.1)




-4.4
-4.4
-6.2
-5.3

Page 11

Table S1. continued



Samples Age1
(Ma)
δ δ13Cped carb
measured

-2.1
-4.4
-2.1
-4.0
-4.0
-3.0
-3.4
-2.6

δ δ13Cped carb
Avg. (±2 s.e.)

Cedar Mesa Fm, se-UT
CM 1A
CM 2A
CM 5
CM3A-1
CM3B
CM3A-2
CM3B-2
CM5
uppermost Nocona Fm, lower Parkeys Oil
Patch, nc- TX
NoPlat1
NoPlat 2C-H
NoPlat 2C-C
NoPlat 4J
NoPlat 4L
NoPlat 4P
NoPlat 4O-1
NoPlat 4O-2
NoPlat 4M
upper Wellington Fm, Waurika, OK
WAR 1B.1
WAR 2B.1
WAR 2B.2
WAR 4B.1
WAR 4B.2
WAR 4C.1
WAR 4C.2
WAR 4C.3
basal Petrolia Fm, upper Parkeys Oil Patch,
nc- TX
POP 1E
POP 1C
POP 1G
POP 2D
POP 4GA
POP 4GB
POP 4H
upper Wellington Fm, OK
OK 2.1.1
285.0








-3.2 (±0.6)








285.0









284.6








-5.6 (±0.9)









-8.6 (±0.6)
-5.9
-9.0
-5.9
-4.1
-5.2
-5.2
-5.1
-5.2
-5.1

-8.3
-8.2
-9.3
-7.4
-7.7
-9.6
-9.3
-9.1









284.2







283.8

-6.4 (±1.0)







-8.7 (±0.1)

-7.7
-6.6
-8.7
-5.1
-5.4
-5.5
-5.9

-8.7
OK 2.1.2
OK 2.3.1
OK 2.3.2





-8.8
-8.6
-8.6

Page 12

Table S1. continued


Samples Age1
(Ma)
δ δ13Cped carb
measured

-10.0
-9.8
-10.5
-10.4
-10.8
-10.6
-9.4
-8.9
-8.2
-6.9

-7.3
-7.4
-7.3
-7.2
-7.8
-7.1
-7.0

δ δ13Cped carb
Avg. (±2 s.e.)

middle Petrolia Fm, Amphitheater, nc- TX
PAM 2D
PAM 2DR
PAM1C
PAM1C2
PAM 2E
PAM 2F
PAM 3C-1
PAM 3C-2
PAM 1O
PAW1
upper Petrolia Fm, Hentz stratotype, nc- TX
PSTL 2E
PSTL 2E
283.4










282.2







-9.5 (±0.8)










-7.3 (±0.2)







PSTL 3G2
PSTL 3G3
PSTL 3G5
PSTL 3H
PSTL 3H
uppermost Petrolia Fm, Castle Hollow,
nc- TX
CPU 12
CPU 12
CPU 9B-1
CPU 9B-2
CPU 8
281.9





-5.1 (±0.5)


-5.4
-5.2
-4.7
-4.5
-5.9




Organ Rock, se-UT
OR2-1
OR2-2
middle Waggoner Ranch Fm, Franklin Bend,
nc- TX
WFBU 1D
WFBU10E-1
WFBU10E2
WFBU4D-1
WFBU4D-2
WFBU4E-1
WFBU4E-2
281.8


-2.5 (±0.1)
-2.5
-2.6



280.9







-5.8 (±1.0)







-4.8
-3.8
-4.9
-4.8
-6.2
-6.7
-6.9
WFBU4D-3





-7.8

Page 13

Table S1. continued


Samples Age1
(Ma)
δ δ13Cped carb
measured

δ δ13Cped carb
Avg. (±2 s.e.)

Upper Waggoner Ranch Fm, Mitchell Creek,
nc-TX
MCLRUS
MC0
MC 25
MC 28B
WCB1A
WCB 2-1
WCB5
WCB 7
WCB 30
WCB 33
WCB 36
WCB 38
WCB 40
WCB44
WCB 50
uppermost Garber Fm, OK
OK 3.1.1
OK 3.3.1
OK 3.3.2
Leuders Fm, nc-TX
TWA
TWB
LL1
LL2
LL3
LL-4
L-99-3
L-99-2
L-99-2R
L-99-4
MLEU1
ULEU 1
ULEU 2
basal Clear Fork Gp, Craddock Ranch, nc-
TX
CFC 2E-1
CFC 2F
280.2















280.1



279.8













-5.4 (±0.5)















-6.0 (±0.4)



-4.5 (±0.7)













-4.6
-4.6
-5.4
-5.5
-5.7
-5.7
-6.0
-5.9
-7.9
-5.2
-4.6
-4.8
-4.8
-5.3
-5.2

-6.4
-5.7
-5.8

-5.4
-5.3
-5.0
-5.2
-5.3
-5.2
-2.8
-3.8
-3.6
-1.4
-4.1
-5.8
-6.1

279.3


-4.3 (±1.1)


-3.2
-6.1
CFC 3-1

-3.8

CFC 3D-1



-4.0



CFC 1G-1 -4.1

Page 14

Table S1. continued


Samples Age1
(Ma)
δ δ13Cped carb
measured

δ δ13Cped carb
Avg. (±2 s.e.)

Arroyo Fm, Clear Fork Gp, Hog Creek,
nc-TX
HC2(A)
HC2(B)
HC8A(1)
HC8A(2)
HC9A
HC11A
basal Vail Fm, Clear Fork Gp, Hog Creek,
nc-TX
CF-1A-1
CF-1A-2
CF-1A-R
CF-1B
99-CF-2
Vail Fm, Clear Fork Gp, Wichita River,
nc-TX
WRVF8
WRVF12
WRVF20
WRVF22
WRVF26
Choza Fm, Clear Fork Gp, Montgomery
Ranch, nc-TX
MG3A
MG11A
MG5
MG 16A
278.9






-5.0 (±0.6)






-6.4
-5.3
-4.6
-4.7
-4.1
-5.0

278.1





-5.1 (±0.9)





-4.9
-5.1
-5.0
-4.0
-6.5
277.1






-4.5 (±0.8)





-3.4
-4.0
-4.8
-4.7
-5.6
274.4

-4.8 (±0.2)

-4.6
-4.9
-4.8
-5.1







1Ages of localities based on stratigraphic age model developed in this study and
calibrated to the Geologic Time-Scale 2004 (S11).

Page 15

Table S2. Fossil organic matter δ13C values used for paleo-pCO2 reconstruction.
Formation & Locality
Age1
(Ma)
Sample ID Material Analyzed
δ δ13Corg
measured
2
δ δ13Corg
Fm Avg.
3
δ δ13Corg
±2 X std. err. (count)
Markley Formation, nc-Texas:
-22.8 ±0.3 (33)
Brannon Mine
Newcastle & Coal Mountain
Newcastle
Bloodworth
Bloodworth, Cooper
Bloodworth, Cooper
Bloodworth, Cooper
Bloodworth, Cooper
Cooper
Cooper
Cooper
Cooper
Cooper upper beds
Cooper lower shale
Gillespies
Loving West
Bloodworth, Cooper
Lycopod
Lycopod B
Bloodworth Bed 12
Williamson Dr.
Williamson Dr
Williamson Dr
Williamson Dr
Williamson Dr
Williamson Dr
Voyles Coal
Malone Ranch &Antelope
Walker
Walker
Walker
Walker
301.3
300.8
300.8
300.3
300.3
300.3
300.3
300.3
300.3
300.3
300.3
300.3
300.3
300.3
300.3
300.3
300.3
300.3
300.3
300.3
300.3
300.3
300.3
300.3
300.3
300.3
298.6
298.5
298.5
298.5
298.2
298.3
A3
A4
A17
A1
WAD 1
WAD 1R
WAD 2
WAD 2R
WAD 5
WAD6R
WAD6
A5
A6
A7
A8
A9
A10
A11
A12
A13
A16
WAD 3
WAD 4
WAD 4R
A16
A15
A14
A2
MBW1A
MBW1B
WAD8
WAD 7
carbonized wood
compression plant
bulk organic matter
compression plant
compression plant
compression plant
compression plant
compression plant
compression plant
compression plant
compression plant
bulk organic matter
compression plant
bulk organic matter
compression plant
compression plant
coal
compression plant
coal
compression plant
charcoal
compression plant
compression plant
compression plant
charcoal
charcoal
charcoal
compression plant
compression plant
compression plant
compression plant
compression plant
-22.7
-21.0
-23.7
-23.4
-23.6
-22.7
-22.8
-24.3
-22.3
-23.0
-23.4
-23.8
-23.2
-21.3
-22.4
-23.8
-22.1
-23.9
-22.7
-21.2
-22.3
-22.9
-23.5
-21.8
-22.4
-23.5
-22.9
-22.7
-23.0
-22.2
-22.5