EA39CH17-McInerney ARI 29 March 2011 21:21
The Paleocene-Eocene
Thermal Maximum:
A Perturbation of Carbon
Cycle, Climate, and Biosphere
with Implications for the Future
Francesca A. McInerney1,∗ and Scott L. Wing2
1Department of Earth and Planetary Sciences, Northwestern University,
Evanston, Illinois 60208; email: cesca@earth.northwestern.edu
2Department of Paleobiology, National Museum of Natural History, Smithsonian Institution,
Washington, DC 20013; email: wings@si.edu
Annu. Rev. Earth Planet. Sci. 2011. 39:489–516
First published online as a Review in Advance on
March 1, 2011
The Annual Review of Earth and Planetary Sciences is
online at earth.annualreviews.org
This article’s doi:
10.1146/annurev-earth-040610-133431
Copyright c© 2011 by Annual Reviews.
All rights reserved
0084-6597/11/0530-0489$20.00
∗Formerly Francesca A. Smith.
Keywords
PETM, global warming, paleoclimate, paleoecology, carbon cycle
Abstract
During the Paleocene-Eocene Thermal Maximum (PETM), ∼56 Mya,
thousands of petagrams of carbon were released into the ocean-atmosphere
system with attendant changes in the carbon cycle, climate, ocean chem-
istry, and marine and continental ecosystems. The period of carbon release
is thought to have lasted<20 ka, the duration of thewhole eventwas∼200 ka,
and the global temperature increase was 5–8◦C. Terrestrial and marine or-
ganisms experienced large shifts in geographic ranges, rapid evolution, and
changes in trophic ecology, but few groups suffered major extinctions with
the exception of benthic foraminifera. The PETMprovides valuable insights
into the carbon cycle, climate system, and biotic responses to environmental
change that are relevant to long-term future global changes.
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INTRODUCTION AND BRIEF HISTORY OF STUDY
Almost 20 years ago Kennett & Stott (1991) published an account of rapid shifts in stable carbon
and oxygen isotope ratios observed in species-specific foraminiferal carbonate fromOceanDrilling
Program (ODP) Site 690B off the coast of Antarctica. Key features of the event were a rapid onset
(which they estimated to be ∼6 ka), a decline in oxygen isotope ratios (indicating warming of
3–4◦Cin surfacewater and∼6◦Cindeepwater), a negative shift in the δ13Cof benthic (∼−2!) and
planktic (∼−4!) forams, influx of a diverse surface water microfossil assemblage with subtropical
affinities, an increase in kaolinite, and elimination or even reversal of depth gradients in both
δ13C and δ18O. All these changes were synchronous with an already-known extinction of benthic
foraminifera (Thomas 1989).
Following the paper by Kennett and Stott on what has come to be known as the Paleocene-
Eocene Thermal Maximum (PETM), Koch et al. (1992) measured the negative carbon isotopic
excursion (CIE) in pedogenic carbonate and mammalian tooth enamel derived from continental
rocks, and they recognized that the CIE coincided exactly with the largest mammalian faunal
turnover of the Cenozoic. The paper by Koch et al. (1992) established the PETM as a global
event that could be detected in the continental realm and showed how the CIE could be used to
correlate marine and continental sections.
Recognition that the PETMwas a global environmental shift associatedwith profound biologi-
cal consequences attracted additional study, as did the problem of formalizing a Paleocene-Eocene
boundary that could be used in both terrestrial and marine sections (e.g., Aubry 2000). By the end
of the 1990s more than 50 papers had been published on the PETM, and the number has exceeded
400 as we write (Figure 1, Supplemental Figures 1 and 2). (To view the Supplemental figures,
table, and accompanying references, follow the Supplemental Materials link from the Annual
Reviews home page at http://www.annualreviews.org.) The literature on the PETM is now too
extensive to review comprehensively, so we focus on changes in carbon cycle, paleoclimate, and
biotas during the PETM.
AGE AND DURATION OF THE PALEOCENE-EOCENE
THERMAL MAXIMUM
The Paleocene-Eocene boundary has been defined by the placement of a Global Boundary Stra-
totype Section and Point (GSSP) in the Dababiya Quarry section of Egypt and is most commonly
recognized by the base of the CIE (Luterbacher et al. 2000). The age of the CIE onset has been
estimated recently as 56.011–56.293 Ma, using radiometric dates of marine ash layers and orbital
tuning of marine sediments (Westerhold et al. 2009). Zircons from the upper part of the CIE are
estimated to be 56.09 ± 0.03 Ma, determined from chemical abrasion–thermal ionization mass
spectrometry ( Jaramillo et al. 2010). This is consistent with an onset at approximately 56.3 Ma.
Prior to the establishment of the Paleocene-Eocene boundary, this same event was known as the
Latest Paleocene Thermal Maximum (LPTM); following the boundary designation it was also
termed the Initial Eocene Thermal Maximum (IETM). PETM is now widely used.
Two main approaches have been used to estimate the duration of the PETM: astronomical
cyclostratigraphy and extraterrestrial 3He fluxes. Orbital timescales suggest a total duration of
150–220 ka (Aziz et al. 2008; Ro¨hl et al. 2003, 2007), whereas 3He suggests 120–220 ka (Farley &
Eltgroth 2003, Murphy et al. 2010). The onset of the CIE—that is, the time from the last samples
with pre-CIE carbon isotope composition to the most depleted values—has been estimated to
take place in less than 10 ka (Zachos et al. 2005), but the dissolution of uppermost Paleocene
and lowermost Eocene carbonate associated with the CIE in many deep marine sections makes
490 McInerney ·Wing
Supplemental Material
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www.annualreviews.org • Paleocene-Eocene Thermal Maximum 491
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it difficult to estimate the duration of the onset (Murphy et al. 2010). The onset of the CIE in
continental sections, which are not affected by carbonate dissolution, occupies 4–7 m of section,
equivalent to 8–23 ka based on calculated sedimentation rates (Aziz et al. 2008, Bowen et al. 2001,
Magioncalda et al. 2004).
A central debate about the PETM timescale concerns the duration of the “recovery” from
the CIE, during which carbon isotope composition returned almost to latest Paleocene values.
Sediments of the recovery phase are roughly twice as thick as those from the “body” of the CIE
(sensu Bowen et al. 2006) at ODP Site 690 on Maud Rise (Bains et al. 1999), but in many other
marine and terrestrial records the recovery phase is shorter relative to the body of the excursion
(Bowen et al. 2001, Giusberti et al. 2007, Zachos et al. 2005). In the most recent estimate of the
duration of the CIE, based on 3He calibration of depositional rates at ODP Site 1266 on Walvis
Ridge, the body (or core) of the excursion is∼113 ka and the total recovery is∼83 ka, with∼33 ka
in an early rapid phase (I) and∼50 ka in a subsequent gradual phase (II) (Murphy et al. 2010). The
1266 record, when calibrated to time using orbital cyclostratigraphy, has a body of ∼59 ka and
a total recovery interval of ∼112 ka (Ro¨hl et al. 2007). The difference in CIE time estimates for
deep marine sediments probably arises from variation in sedimentation rates caused by carbonate
dissolution at the onset of the CIE and enhanced carbonate deposition during the recovery phase
(Kelly et al. 2005, 2010; Murphy et al. 2010). The one continental section that has been calibrated
to time with orbital cycles (Polecat Bench,Wyoming) gives estimates of∼115 ka for the body and
∼42 ka for what appears to be the rapid recovery (phase I) (Aziz et al. 2008). Because a long body
and short recovery to the CIE are seen in 3He-calibrated deep marine sections as well as shallow
marine and continental sediments, they are likely better estimates of the true duration of the CIE
and its phases.
ENVIRONMENTAL CHANGE DURING THE PALEOCENE-EOCENE
THERMAL MAXIMUM
The PETM is defined by a global temperature rise that was initially inferred from a>1! negative
excursion in δ18O of benthic foraminifera, indicating a deep-water temperature increase of ∼5◦C
(Kennett & Stott 1991, Thomas & Shackleton 1996, Zachos et al. 2001). A similar magnitude
of temperature increase has also been inferred from Mg/Ca ratios (Tripati & Elderfield 2005,
Zachos et al. 2003). Planktonic foraminiferal δ18O and Mg/Ca ratios are also generally consistent
with ∼5◦C of warming, though estimates reach up 8◦C of warming at Wilson Lake, New Jersey
(Zachos et al. 2006), and ODP Site 690, Maud Rise (Kennett & Stott 1991, Thomas et al. 2002).
Although there is spatial variation in the amount of warming associated with the PETM, it is
significant that temperature increase in theArctic (LomonosovRidge) determined frombiomarker
paleothermometers (TEX86 ′ and GDGT) is ∼5–8◦C (Sluijs et al. 2006, Weijers et al. 2007). The
absence of greater warming at polar latitudes during the PETM implies an absence of the ice-
albedo feedback loop that would have generated more warming at high latitudes if sea or land ice
had been present in the late Paleocene (Sluijs et al. 2006). A warming of ∼5–7◦C is also inferred
from leaf margin analysis and oxygen isotope composition of fish scales, mammalian tooth enamel,
and soil nodules in the fully continental Bighorn Basin, Wyoming (Bowen et al. 2001, Fricke &
Wing 2004, Fricke et al. 1998, Koch et al. 2003, Wing et al. 2005).
Modeling of the PETM, as with equable climates of the early Cenozoic in general, has tended
to produce temperatures in polar regions that are too cool (e.g., Huber et al. 2003, Shellito& Sloan
2006) even if the atmosphere is stipulated to have 16 times the preindustrial concentration of CO2
in the atmosphere (Winguth et al. 2010). This persistent problem with modeling a PETM climate
that has a low equator-to-pole temperature gradient and warmwinters in mid-latitude continental
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interiorsmay require consideration of the effect ofmethane or other trace greenhouse gases (Sloan
et al. 1992) and clouds (Abbot & Tziperman 2009). Suggestions that tropical temperatures were
%35◦C (Huber 2008), allowing for a steeper latitudinal gradient, are unlikely to be born out
because terrestrial floras from the tropics become more rather than less diverse during the PETM
( Jaramillo et al. 2006, 2010).
In addition to warming, there is widespread evidence for changes in runoff, often interpreted
as indicating increased precipitation on the continents during the PETM. High abundance of
kaolinite, a product of weathering in warm, wet climates, has been observed near the Paleocene-
Eocene boundary in many marginal marine sections from Antarctica to eastern North America
to North Africa and Pakistan (Bolle & Adatte 2001, Gibson et al. 2000, Robert & Kennett 1994).
The influx of kaolinite, however, likely reflects erosion of previously formed clay, rather than
weathering and erosion during the geologically brief PETM (Schmitz et al. 2001, Thiry &Dupuis
2000). The sharp increase in eutrophic dinoflagellate populations in coastal marine sections during
the PETM has also been cited as evidence for enhanced terrestrial input (Crouch et al. 2001,
2003), which has in turn been equatedwith increased precipitation on land. In fact, the relationship
between sediment production and precipitation is not positive andmonotonic but rather unimodal,
with the highest erosion and sediment loads occurring under highly seasonal precipitation of
∼300–600 mm year−1 (Cecil & DuLong 2003). Thus more terrestrial runoff in nearshore marine
sediments is likely an indicator of higher seasonality or episodicity rather thangreater precipitation.
More episodic precipitation could explain enhanced export of terrestrial organic matter to the
Arctic Ocean basin during the PETM (Boucsein & Stein 2009, Pagani et al. 2006b) as well as
higher iron content of marine carbonates in New Zealand (Villasante-Marcos et al. 2009).
Proxy data indicate higher variability of precipitation in a number of regions during the PETM.
PETM sections in the Pyrenees of northern Spain (paleolatitude ∼35◦N) include the distinctive
and widespread Claret Conglomerate, a megafan deposit typical of areas with highly intermittent,
violent rainfall (Schmitz & Pujalte 2007, Schmitz et al. 2001). Palynological assemblages from
the same area indicate a loss of taxodiaceous conifers, which prefer wetter habitats (Schmitz et al.
2001). Recent evidence from silicon isotopes in paleosols from the Paris Basin, farther to the north,
fails to indicate enhanced chemical weathering during the PETM, consistent with no increase in
total precipitation (Rad et al. 2009). PETM marine deposits in Austria show a strong increase in
siliciclastic deposition but remain dominated by smectite clays that indicate weathering under a
seasonal subtropical climate; this has been interpreted as an increase in seasonality of precipitation
leading to greater erosion and delivery of terrestrial material to the marine environment (Egger
et al. 2005). In North America, paleosol horizons of PETM age in the Bighorn Basin ofWyoming
(paleolatitude ∼49◦N) indicate less chemical weathering (Kraus & Riggins 2007), and changes in
ichnofossil assemblages are consistent with drier or better-drained soils (Smith et al. 2008a,b). In
the same area, decrease in leaf size from latest Paleocene to early PETM floras reflects an increase
in water stress (Wing et al. 2005, 2009). A more arid climate is indicated by paleosols from the
southern Rocky Mountains (paleolatitude ∼44◦N) during the PETM (Bowen & Bowen 2008).
Although general circulation models of PETM climate predict 10–20% higher precipitation
in many continental areas (e.g., Huber & Sloan 1999, Winguth et al. 2010), available proxy
data (mostly mid-latitude northern hemisphere) are more consistent with increased seasonality of
precipitation, or more intense storms, rather than wetter climates. Higher precipitation is cited
as a factor in increasing the rate of silicate weathering on the continents, which in turn has been
proposed as a major cause of long-term drawdown of CO2 from the atmosphere during recovery
from the CIE (but see Bowen & Zachos 2010). We suggest that if higher precipitation were a
major factor in higher rates of silicate weathering, it did not occur in the mid-latitude regions that
are currently best sampled.
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CARBON CYCLE CHANGES AT THE PALEOCENE-EOCENE
THERMAL MAXIMUM
Two observations indicate amassive release of carbon at the PETM: the large, global negativeCIE
(Figure 2) and the extensive dissolution of deep-ocean carbonates. The negative shift in carbon
isotope (δ13C) values shows that the carbon released was depleted in 13C relative to the exogenic
reservoir (ocean + atmosphere + biomass) and was likely organic carbon because organisms
discriminate against 13C during biosynthesis. The geologically rapid onset of the CIE (<20 ka)
implicates the addition of 13C-depleted (organic) carbon rather than a reduction in organic carbon
burial (100 ka timesale; Kump& Arthur 1999). The sudden rise in the calcite compensation depth
(CCD) (below which no calcite accumulates) at the onset of the CIE indicates that the ocean
absorbed a large amount of CO2, reducing its pH and carbonate ion content [CO32−] (Zachos
et al. 2005). Although the large carbon release at the PETM is accepted, the source and mass of
carbon are still actively debated. Five sources are being discussed:
! Methane clathrates: Methane clathrates are icy solids consisting of methane surrounded by
water molecules. They are stable in deep-sea sediments, but they can be destabilized by
increasing temperature caused by changes in ocean circulation (Dickens et al. 1995, 1997)
or by decreasing pressure resulting from slope failure (Katz et al. 1999). This microbial
methane has a δ13C value of ∼−60! (Kvenvolden 1993).
! Wildfires: Burning of the extensive peat and coal deposited during the Paleocene (δ13C
of ∼−22!) could have resulted from increasing atmospheric O2, dryer climates, and/or
uplift of coal basins (Kurtz et al. 2003). However, no increase in combustion byproducts was
observed in cores from either the Atlantic or Pacific (Moore & Kurtz 2008).
! Thermogenic methane: Injection of magma into organic-rich sediments would have caused
the explosive release of thermogenicmethane (δ13C of∼−30!) fromCretaceous-Paleocene
mudstones in the North Atlantic (Svensen et al. 2004, 2010; Westerhold et al. 2009).
! Drying epicontinental seas: Tectonically driven isolation of an epicontinental seaway would
have led to rapid (<20 ka) desiccation and oxidation of organic matter (δ13C of ∼−22!)
(Higgins& Schrag 2006). Though vast areas of central Asia were covered by shallow seaways
in the Paleocene-Eocene, none are known to have dried up coincident with the PETM
(Gavrilov et al. 2003).
! Permafrost: During the Paleogene, Antarctica did not support a large ice cap and may have
stored vast quantities of carbon as permafrost and peat that could have been rapidly thawed
and oxidized, releasing carbon (δ13C of ∼−30!) (DeConto et al. 2010).
How can these hypotheses be evaluated? Each source has a characteristic δ13C value; thus the mass
of carbon required to create the CIE can be calculated for each source (see sidebar, Mass Balance
Calculations). These masses can be compared with the amount of carbon required to create the
observed carbonate dissolution. In addition, these carbon releases must also be consistent with
the observed warming given our understanding of the greenhouse effect. In spite of progress, the
magnitude of the CIE, the amount of carbonate dissolution, and atmospheric pCO2 levels remain
difficult to quantify, as discussed below.
Mass Balance Accounting
Mass balance calculations require an estimate of the magnitude of the CIE in the exogenic carbon
pool, but different archives vary widely.We compiled 165 carbon isotope records from the PETM,
which show ameanCIE of−4.7±1.5 (± standard deviation) for terrestrial records and−2.8±1.3
for marine records (Table 1). Even within each realm archives differ. Comparison of median
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MASS BALANCE CALCUATIONS
Steady-state isotopic mass balance is essentially instantaneous mixing. The final mass and isotopic ratio of carbon
is the result of the initial mass and isotopic composition plus the mass and isotopic composition of added carbon.
To calculate the mass of carbon added during the PETM, use the following equations:
(Mfinal)× (δ13Cfinal) = (Minitial)× (δ13Cinitial)+ (Madded)× (δ13Cadded). (1)
M is mass of carbon and δ13C is the carbon isotope ratio of each pool (initial, added, and final). Carbon isotope
ratios are reported in delta notation as
δ13C = [(Rsample/Rstandard)− 1] in per thousand or permil (!) units, (2)
R = 13C/12C, (3)
Mfinal =Minitial +Madded, (4)
CIE = δ13Cfinal − δ13Cinitial. (5)
By combining Equations 1, 4, and 5, the mass of carbon added can be calculated as
Madded = −CIE×Minitial/(δ13Cfinal − δ13Cadded). (6)
In Figure 3, we calculate the mass added for each reservoir and CIE using estimates for the initial mass and carbon
isotope ratio of the Paleocene surface reservoir as 50,000 Pg C (1 Pg = 1 Gt = 1015 g) (Archer 2007) and −2.5!
(Killops & Killops 2005).
values shows that the smallest CIEs are in marine carbonates; intermediate CIEs are in marine
and terrestrial organic matter; and the largest CIEs are in terrestrial plant lipids, mammalian tooth
enamel, and pedogenic carbonate nodules (Table 1, Figure 2). A Kruskal-Wallis test rejects the
null hypothesis that the CIEs for all groups are drawn from a single population (p < 0.0001);
the marine and terrestrial records are also significantly different (Wilcoxon-Mann-Whitney test,
p < 0.0001).
Which archive is most reliable? Marine carbonates are subject to severe dissolution during
the PETM. Benthic foram and bulk carbonate δ13C records clearly show greater dissolution and
truncation of the CIE with increasing depth (McCarren et al. 2008, Zachos et al. 2005). The
δ13C of bulk organic matter in marine records is influenced by productivity and terrestrial input
(Hilting et al. 2008); the δ13C of bulk organic matter in the terrestrial realm is affected by soil
respiration (Wynn 2007) and many other factors (e.g., Diefendorf et al. 2010, Smith et al. 2007).
Soil carbonate δ13C values vary with soil properties, atmospheric pCO2, rate of carbon input, and
rate of soil respiration (Bowen et al. 2004, Sheldon & Tabor 2009).
Terrestrial plants take up CO2 directly from the atmosphere and therefore could provide
the most reliable measure of the CIE. However, environmental factors such as water availability
and intrinsic features such as plant functional type (e.g., conifer versus angiosperm) influence
carbon isotope fractionation by C3 plants (Diefendorf et al. 2010). Bowen et al. (2004) suggested
increased water availability caused by an enhanced hydrologic cycle could have increased carbon
isotope fractionation. Pagani et al. (2006b) argued that this was unlikely in the cool, wet Arctic,
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Table 1 Summary of 165 records of the carbon isotopic excursion at the Paleocene-Eocene Thermal Maximum
Archive n Minimum Maximum Average
Standard
Deviation Median
Benthic forams 36 −0.6 −5.1 −2.5 1.0 −2.6
Planktic forams 36 −0.7 −4.4 −2.7 1.0 −2.6
Bulk carbonate 33 −1.1 −4.9 −2.7 1.1 −2.4
Bulk marine organic matter 11 −1.2 −8.0 −4.1 2.2 −3.5
Algal lipid 1 −3.5 −3.5 −3.5 n/a −3.5
Soil carbonate 14 −2.5 −7.3 −5.5 1.7 −6.3
Plant lipids 16 −2.8 −7.6 −5.1 1.3 −5.0
Bulk soil organic matter 14 −2.2 −4.6 −3.5 0.6 −3.5
Tooth enamel 4 −4.3 −5.2 −4.8 0.4 −4.9
All marine 117 −0.6 −8.0 −2.8 1.3 −2.6
All terrestrial 48 −2.2 −7.6 −4.7 1.5 −4.6
where leaf wax n-alkanes record a CIE of −5!. Smith et al. (2007) observed a negative CIE of
−5! in leaf wax n-alkanes from the Bighorn Basin but pointed out that different chain-length
n-alkanes recorded different CIEs, possibly indicating a shift from mixed conifer-angiosperm
forest to purely angiosperm forest that amplified the CIE. Schouten et al. (2007) showed that
plant functional type influences the magnitude of the CIE: Conifer terpenoids record a CIE
of −3!, whereas angiosperm terpenoids record −6!. Diefendorf et al. (2010) quantified the
isotopic effects of both precipitation and plant type. They concluded that the isotopic effect of a
shift from conifers to angiosperms would have been largely counteracted by an overall decrease in
plant fractionation caused by higher water use efficiency in response to higher water stress (Kraus
& Riggins 2007, Wing et al. 2005).
Marine carbonates most likely underestimate the CIE; marine bulk organic matter and soil
carbonates may overestimate it. McCarren et al. (2008) proposed the CIE was no smaller than
−3.5!, as recorded in benthic forams from the shallowest, least truncated record onWalvis Ridge
that best reflects the excursion in deep-ocean dissolved inorganic carbon, the largest reservoir of
carbon in the exogenic pool.McCarren et al. (2008) also advocated that the CIEwas no larger than
−5!, the value from leaf-wax n-alkanes at the same site (Hasegawa et al. 2006). Diefendorf et al.
(2010) produced a considered estimate of the atmospheric CIE (−4.6!) by accounting for the
effects of both plant functional group and climate on fractionation. The releases associated with a
CIE of −4.6! are 4,300 Pg C for methane clathrates, 10,000 Pg C for thermogenic methane or
permafrost, and 15,400 Pg C for wildfires or epicontinental seas (Figure 3).
−−−−−−−−−−−−−−−−−−−−−−−−−−−−−−−−−−−−−−−−−−−−−−−−−−−−−−−−−−−−−−−−−−−−−−−→
Figure 2
Measurements of the negative carbon isotopic excursion (CIE) at the Paleocene-Eocene Thermal Maximum
(PETM) grouped by source material (n = 165; see Supplemental Material for references). (a) Each record
is characterized by the pre-PETM mean, the most negative data point within the CIE (or average for that
sample level), and the post-PETM mean (blue circles represent marine records; green circles represent
terrestrial records). The dashed lines are the generalized shape of the CIE connecting the means of the
plotted values (black circles). All x-axes represent 10!. (b) The CIE for each archive is shown in the
box-and-whisker plots. The horizontal line represents the median; the box identifies the upper and lower
quartile and contains 50% of the data; the whiskers identify the range; open circles are outliers.
496 McInerney ·Wing
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–23–32 –29 –26
5–5 –3 –1 1 3
δ13C (‰) 
δ13C (‰) δ13C (‰) δ13C (‰) δ13C (‰)
δ13C (‰) δ13C (‰) δ13C (‰) δ13C (‰) 
5–5 –3 –1 1 3 5–5 –3 –1 1 3 –24–33 –30 –27 –27–36 –33 –30
Bulk soil 
organic matter
Pre-PETM
CIE
Post-PETM
Soil carbonate 
–8–17 –14 –11 –27–36 –33 –30
Pre-PETM
CIE
Post-PETM
Plant lipids 
–9–18 –15 –12
Tooth enamel 
Benthic 
foraminiferans
Planktonic 
foraminiferans Bulk carbonate 
Bulk marine
organic matter Algal lipid 
0
–1
–2
–3
–4
–5
–6
–7
–8
–9
Be
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Bu
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Bu
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Al
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n=36 n=36 n=33 n=11 n=1 n=14 n=14n=16 n=4
a
b
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(‰
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–8–2 –3 –4 –5 –6 –7
35,000
0
5,000
10,000
15,000
20,000
25,000
30,000
Carbon isotope excursion (‰)
M
as
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 (P
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C)
Wild!re/
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seas
Thermogenic
methane/
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Methane
hydrate
–22‰
–30‰
–60‰
1,500
3,000
4,500
6,000
7,500
9,000
10,500
0
Marine range of median CIEs
Terrestrial range of median CIEs
A
tm
ospheric pCO
2  increase (ppm
 CO
2 )
Figure 3
Mass balance estimates (see sidebar, Mass Balance Calculations) of mass of carbon released from different
sources over a range of absolute values for the carbon isotopic excursion (CIE). Carbon sources: terrestrial
wildfires or oxidation of organic matter from epicontinental seas (δ13C = −22); thermogenic methane or
thawed permafrost (δ13C = −30); methane hydrate (δ13C = −60). Area shaded blue indicates marine range
of median CIEs. Area shaded green represents terrestrial ranges for median CIEs (Table 1). Open circles
represent masses for CIE of −4.6! (Diefendorf et al. 2010). Dashed horizontal line indicates minimum
carbon release of 6,800 Pg C from modeling of carbonate dissolution allowing for variable bioturbation
(Panchuk et al. 2008). Dotted horizontal line represents maximum mass initial release of 3,000 Pg C
estimated by modeling carbonate dissolution while allowing for changes in ocean circulation (Zeebe et al.
2009). Second y-axis converts the mass of carbon released to an increase in atmospheric CO2 (1 Pg
C = 0.3 ppm CO2) derived from the Grid ENabled Integrated Earth system (GENIE) model for late
Paleocene ocean alkalinity, dissolved inorganic carbon, and seafloor carbonate content (Panchuk et al. 2008).
Carbonate Dissolution Accounting
The carbon release can be constrained by the CO2 required to cause the observed dissolution of
marine carbonates. On the basis of a depth transect in the Southeast Atlantic at Walvis Ridge,
Zachos et al. (2005) estimated that the CCD shoaled by >2 km and that 4,500 Pg C would be
required to cause similar CCD shoaling globally. However, the CCD rose by only ∼500 m in the
Pacific (Colosimo et al. 2006).
Two recent studies take a spatially explicit approach to modeling the observed pattern of dis-
solution in different ocean basins. Panchuk et al. (2008) coupled a three-dimensional ocean and
atmosphere circulation model (GENIE-1) to models of the marine carbon cycle and seafloor
sedimentation. Importantly, their sediment model allowed for predictions of wt% CaCO3
with variable bioturbation, which mixes layers of sediment, increasing buffering capacity per
unit area of seafloor. Bioturbation appears to have been reduced in the Atlantic during the
PETM (Dickens 2000, Ridgwell 2007). Panchuk et al. (2008) estimated a minimum release of
6,800 Pg C over 10 ka, assuming no bioturbation in the Atlantic and deep Pacific during the
PETM.
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Zeebe et al. (2009) used the long-term ocean-atmosphere-sediment carbon cycle reservoir
model (LOSCAR). To explain the observed difference in magnitude of shoaling in different ocean
basins, they assumed an increase in formation of North Pacific deep water (NPDW) led to more
corrosive deepwaters in the Atlantic and that a portion of carbonwas directly injected into the deep
Atlantic. A maximum initial injection of 3,000 Pg C over 5 ka explains the observed dissolution in
both the Pacific and the Atlantic. To model the shape of the excursion, they followed the initial
pulse with a sustained release of 1,480 Pg C over 50 ka.
Although these studies differ in many details, three main factors cause the different estimates of
the size of the carbon pulse. First, Panchuk et al. (2008) calculated circulation in the atmosphere
and ocean using physical principles, whereas Zeebe et al. (2009) prescribed circulation. Second,
Panchuk et al. (2008) calculated the pre-PETMCCD in their model and obtained a deeper Pacific
CCD than did Zeebe et al. (2009), who estimated their pre-PETM CCD from sediment cores.
The deeper pre-PETM Pacific CCD obtained by Panchuk et al. requires more carbon to reach
the observed PETM CCD level. Third, Panchuk et al. (2008) explored the role of bioturbation,
whereas Zeebe et al. (2009) examined the role of increased NPDW formation in the differential
dissolution in thePacific and theAtlantic.Discerningwhich approachmost accurately reconstructs
carbonate dissolution at the PETMwill require a better understanding of the relative importance
of bioturbation versus NPDW formation in modulating deep-ocean acidification.
Climate Sensitivity Accounting
Climate sensitivity refers to the temperature increase associated with a doubling of atmospheric
CO2 and all the associated physical, chemical, and biological climate feedbacks. Estimates range
from 1.5◦C to 4.5◦Cper doubling (IPCC 2007), which would imply between 1.1 and 3.3 doublings
of pCO2 to cause 5◦CofPETMwarming. Proxy estimates for Paleocene pCO2 are scarce and range
widely, from 200 ppm to 2,800 ppm (Breecker et al. 2010, Hilting et al. 2008, Pearson & Palmer
2000, Royer 2006, Royer et al. 2001). Assuming pre-PETM pCO2 values of 600–2,800 ppm,
1.1 to 3.3 doublings would equate to an increase of 700–25,000 ppm CO2 or the addition of
2,300–83,000 Pg C (see Figure 3; see also Pagani et al. 2006a).
Causes of the Paleocene-Eocene Thermal Maximum
Ideally, carbonate dissolution and climate sensitivity constraints would converge on a carbon
release that, when compared with mass balance calculations based on the CIE, would identify the
likely source(s) of carbon. Modeling of carbonate dissolution yields nonoverlapping solutions of
less than of 3,000 Pg C initially (Zeebe et al. 2009) or more than 6,800 Pg C (Panchuk et al.
2008). Consideration of climate sensitivity leads to an estimate of >5,400 Pg C (Pagani et al.
2006b). Our current best estimate for the CIE of −4.6! (Diefendorf et al. 2010) yields a carbon
release of 4,300 Pg C for methane clathrates and >10,000 Pg C for all the other sources. Thus,
modeling of carbonate dissolution by Panchuk et al. (2008) and climate sensitivity (Pagani et al.
2006a) indicates that methane clathrates could not have been the source, whereas the carbonate
modeling of Zeebe et al. (2009) suggests methane clathrates were the only viable source. Zeebe
et al. (2009) recognize that their modeled carbon release would be insufficient to cause the full
warming and argue for other unknown climate feedbacks or forcings as the cause of the remainder
of the warming.
Atmospheric CO2 concentrations during the latest Paleocene are critical to estimating the
amount of carbon required to generate PETMwarming. Pagani et al. (2006b) assumed aminimum
pCO2 of 600 ppm in the Paleocene because high pCO2 is required to simulate warm poles in
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GCMs. Many proxies for pCO2, however, suggest values of ∼400 ppm (Breecker et al. 2010,
Hilting et al. 2008, Pearson & Palmer 2000, Royer 2006, Royer et al. 2001). If these proxies are
correct, two doublings (to 1,600 ppm) would yield the addition of 1,200 ppm CO2 or 4,000 Pg C
and a warming of 6◦C, assuming a standard climate sensitivity of 3◦C per doubling (Figure 3).
Methane clathrates would then be a viable source from a mass balance perspective. However, a
strong argument against methane clathrates is that the mass of carbon held in Paleocene clathrates
was likely too small (∼1,000 Pg C) to produce the observed CIE (Buffett & Archer 2004). Mass
balance estimates of the carbon required frommethane clathrates have increased from 1,200 Pg C
(Dickens et al. 1997) to 4,300 Pg C as the estimated CIE has increased from −2.5! (Dickens
et al. 1997) to−4.6! (Diefendorf et al. 2010). If the estimate of Buffet & Archer (2004) is correct,
a different source or combination of sources would be required.
Early work described the CIE as a single release with an asymptotic return to more positive
δ13C values (Dickens et al. 1997), but recent marine and continental records show a ∼100-ka
interval during which isotope values are low but stable, referred to as the body of the CIE (Bowen
et al. 2006, Sluijs et al. 2007a) or the core (Murphy et al. 2010). The body of the CIE could indicate
an alternate stable state of the carbon cycle (Bowen et al. 2006) or the continued release of light
carbon (Zeebe et al. 2009) from one or more sources. Evidence is mounting for a ∼4◦C warming
prior to the onset of the CIE, which also suggests multiple releases and possibly multiple sources
of carbon (Secord et al. 2010, Sluijs et al. 2006).
Thermogenic methane and Antarctic permafrost/peat have intermediate isotopic signatures
and require 10,000 Pg C for a CIE of−4.6!. This mass could explain both the observed warming
and the modeled carbonate dissolution according to Panchuk et al. (2008). The organic rich
rocks in the North Atlantic Volcanic Province, the source of thermogenic methane, would have
held vast amounts of carbon (Svensen et al. 2004). Sufficiently large carbon stocks (potentially
>10,000 Pg C) also could have been stored on Antarctica because of greater land area and longer
accumulation time in the Paleocene (DeConto et al. 2010). Both sources are capable of multiple
and prolonged releases of carbon. Although debate is ongoing, the Antarctic permafrost/peat
hypothesis has the power to explain not only the PETM but also the magnitude and pacing of the
subsequent hyperthermals that occur in the following ∼3 Ma (ETM2 and ETM3). These warm
periods coincide with high eccentricity and obliquity capable of melting Antarctic permafrost
(DeConto et al. 2010). In addition, the recharge of the Antarctic soil carbon reservoir would
have helped the recovery phase and prepared the system for the next release. Each subsequent
release would be expected to be smaller, as observed, because of the overall depleted carbon stocks
(DeConto et al. 2010). Therefore, Antarctic peat and permafrost present intriguing potential
sources for the carbon released at the PETM and subsequent hyperthermals.
Natural Carbon Sequestration: The Recovery Phase
Following the body of the CIE, which lasted ∼113 ka, carbon isotope values and temperatures
returned to near pre-PETM values over∼83 ka (Murphy et al. 2010). This recovery phase repre-
sents natural carbon sequestration and negative feedbacks that curtailed the CO2-induced warm-
ing. Three mechanisms have been proposed for the drawdown of CO2: increased C storage in the
terrestrial biosphere (Beerling 2000, Bowen & Zachos 2010), increased marine export production
(Bains et al. 2000, but see Torfstein et al. 2010), and enhanced silicate weathering on land with
increased preservation of carbonates in the ocean (Kelly et al. 2005, 2010; Ravizza et al. 2001;
Torfstein et al. 2010). Bowen & Zachos (2010) point out that the initial carbon isotope recovery
is quite rapid, occurring over ∼30 ka, and therefore likely required a process that preferentially
sequestered 13C-depleted carbon, such as the burial of organic matter in soils and peats.
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BIOTIC EFFECTS OF THE PALEOCENE-EOCENE
THERMAL MAXIMUM
Marine Ecosystems
The effects of the PETM onmarine ecosystems have been documented in more than 50 cores and
outcrop sections from near the North Pole through the tropics to offshore Antarctica, from open
oceans to shallow coastal seas, and from almost every major ocean basin and continental margin
(Figure 2). In spite of the geographic extent and environmental spread of marine records, no
macroinvertebrate fossil assemblages have been described from the PETM (Ivany & Sessa 2010),
so PETM effects on marine ecosystems are seen entirely through microfossils.
Even before the Paleocene-Eocene transitionwas recognized as amajor perturbation of climate
and carbon cycle it was known as the interval of the largest benthic foraminiferal extinction in the
past 90Ma, removing 30–50%of benthic foraminiferal diversity (Thomas 1989, 1998, 2003, 2007;
Tjalsma & Lohmann 1983). This is often referred to as the benthic foram extinction or benthic
extinction event, and it marks the rapid transition from the Cretaceous-Paleocene fauna to the
Eocene fauna—one of only three major shifts in benthic foram faunas since the mid-Mesozoic
(Thomas 1998, 2007).
The extinction, which begins near the onset of the CIE, was most severe at middle bathyal
and greater depths, but shallower faunas also underwent rapid changes in composition as well as
some extinction (Alegret et al. 2009a,b; Aref & Youssef 2004; Thomas 1998). Throughout the
CIE, benthic foram faunas at all depths have low diversity and high dominance, and they are
composed largely of small, thin-walled or agglutinated (noncalcareous) species (Thomas 2007).
Detailed studies of the benthic foram extinction at Alamedilla and Zumaia in Spain show that the
extinction was rapid but not ecologically instantaneous (Alegret et al. 2009a,b).
The benthic foram extinction has been attributed to increased corrosivity of deep waters, lower
oxygen levels, changes in food supply, and higher temperatures (Thomas 2003, 2007). There is
evidence for change in all these environmental variables in at least some parts of the ocean, but
with the exception of temperature, none of them affected the sea floor globally, favoring the
idea that the temperature increase of 4–5◦C was the ultimate driver of extinction (Alegret et al.
2009a,b; Sluijs et al. 2007b; Thomas 2007). How warming would have driven extinction is less
clear: Thomas (2007) proposed that higher temperatures increased the metabolic rates of benthic
forams, the microbes they consumed, and their metazoan predators, with cascading effects on
species at each trophic level.
In contrast to the marked extinction of benthic foraminifera, other groups of benthic and
planktic microfossils show little or no extinction at the PETM. Ostracodes are the only other
group of deep benthic organisms with detailed records across the PETM, and their response
ranges from decreased diversity and abundance (Steineck & Thomas 1996), to little change in
diversity or relative abundance distribution (Webb et al. 2009), to rapid changes in ostracode
species composition (Morsi et al. 2011, Speijer & Morsi 2002).
Planktic forams shifted geographic ranges during the PETM. The tropical genus Morozovella
occured at high latitude sites just before and during the early part of the PETM (Thomas &
Shackleton 1996), and an increase in warm climate forams is seen in many middle and lower
latitude sites either just before or at the onset of the PETM (Arenillas et al. 1999, Berggren &
Ouda 2003, Lu & Keller 1995, Pardo et al. 1999).
Forams also exhibit rapid evolutionary change in form during the PETM. In the Pacific, plank-
tic forams increased in size while benthic forams decreased in size (Kaiho et al. 1996, Petrizzo
2007). Body-size increase in planktic forams may reflect the success of larger individuals with
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greater capacity for photosymbionts in the oligotrophic surface waters created by thermal stratifi-
cation (Kaiho et al. 1996). Benthic dwarfing may have been mediated by the effect of temperature
on bottom-water oxygenation, with smaller forams favored in less oxic conditions (Kaiho et al.
2006), though higher temperatures may also have increased metabolic rates, causing food to be-
come limiting (Alegret et al. 2010). Rapid diversification and morphological evolution of planktic
forams have also been observed in the Pacific, with PETM“excursion taxa” having distinctivemor-
phology that evolved in <10 ka and persisted only through the PETM (Kelly et al. 1996, 1998).
Detailed morphometric and stable isotopic analyses suggest that ancestral populations temporar-
ily collapsed. Furthermore, in some lineages, the descendants may have originated suddenly from
peripherally isolated populations, whereas other lineages may have undergone sympatric specia-
tion (Kelly et al. 1998). As with increasing size, morphological specializations of the tests may be
related to supporting photosymbionts.
Larger forams were a major reef-forming group in shallow tropical seas during the late
Paleocene (Hottinger 1998, Scheibner & Speijer 2008, Scheibner et al. 2005), perhaps because
their algal symbioses are more resistant to high temperatures (Hallock 2000). With the PETM
there was a major turnover in larger foram community composition (Pujalte et al. 2009): Species
thought to be the most specialized for oligotrophic conditions were eliminated (Scheibner et al.
2005). This is consistent with a transient increase in productivity in shallow parts of Tethys (Bolle
& Adatte 2001, Crouch et al. 2001, Egger et al. 2005, Ravizza et al. 2001, Scheibner & Speijer
2008, Speijer & Morsi 2002).
Nannofossils, the calcareous platelets of unicellularmarine algae (e.g., Coccolithophorida), also
show significant but geographically heterogeneous changes during the PETM (Agnini et al. 2007;
Bown&Pearson 2009; Bralower 2002;Gibbs et al. 2006b; Jiang&Wise 2007;Monechi et al. 2000;
Raffi et al. 2005, 2009). The main factors influencing nannoplankton were probably temperature
and nutrient availability (Gibbs et al. 2010). Nannofossils favoring oligotrophic conditions became
more abundant at open ocean sites (Bralower 2002, Gibbs et al. 2006b, Tremolada & Bralower
2004), but eutrophic nannofossils increased in some marginal seas (Agnini et al. 2007, Angori
et al. 2007, Gibbs et al. 2006b). Off Antarctica, warm-water nannofossils replaced cold-water taxa
(Angori et al. 2007), and at lower latitudes cool-water taxa were eliminated (Bown&Pearson 2009,
Mutterlose et al. 2007). Rates of species extinction and origination were extremely high globally
among nannofossils, especially near the onset of the CIE, with the PETM exhibiting the highest
rates of turnover for the Cenozoic (Gibbs et al. 2006a).
One of the most salient events among marine plankton during the PETM is the large increase
in abundance and geographic range of the heterotrophic dinoflagellate Apectodinium (Bujak &
Brinkhuis 1998; Crouch et al. 2001, 2003; Sluijs et al. 2005, 2006; for a review see Sluijs et al.
2007a). This lineage, which had been largely restricted to low latitudes in the late Paleocene,
expanded to polar latitudes in both hemispheres during the PETM, attaining abundances of
>40% of dinocysts in many samples. Apectodinium was able to spread to higher latitudes because
of higher temperature, but the abundance of this heterotroph may be related to higher nutrient
levels near the continents (Sluijs et al. 2007b).
Terrestrial Ecosystems
PETM sections containing terrestrial fossils and biomarkers have been described from the Arctic
Ocean,NorthAmerica, northernSouthAmerica,westernEurope, central and southAsia, east Asia,
and New Zealand (Figure 1, Supplemental Table 1). North America has the greatest number
and geographic extent of continental and continental margin PETM sites, including sites from the
Arctic to the southern shore of the continent, and from intramontane and foreland basins in the
502 McInerney ·Wing
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Rocky Mountains to marginal marine settings of the Gulf and Atlantic Coastal Plains. Given the
heterogeneity of terrestrial climates and biotas, it is not surprising that faunal and floral responses
to the PETM differ by continent and region.
For more than a century it has been recognized that a burst of mammalian first appearances
occurs at the base of the Eocene, including the earliest even-toed ungulates (Artiodactyla), odd-
toed ungulates (Perissodactyla), primates (not including Plesiadapiformes), and Hyaenodontidae
(Gingerich 1989, 2003, 2006; Hooker 1998; Rose 1981 and references therein). Koch et al. (1992,
1995) showed that the first appearances of these taxa in the northern Rocky Mountains coincided
closely with the onset of theCIE and suggested that the newmammals dispersed toNorth America
across high latitude land bridges while the climate was warm. This faunal turnover is sometimes
referred to as the mammalian dispersal event and is seen in European as well as North American
faunas (Hooker 1998).
Fine stratigraphic resolution of mammalian faunal change and direct correlation with the CIE
have been achieved at Polecat Bench in the Bighorn Basin (Bowen et al. 2001; Gingerich 1989,
2001). The body of the CIE is 40–45 m thick in this section (Bains et al. 2003, Bowen et al. 2001,
Magioncalda et al. 2004). The onset of the CIE here is in rocks producing Clarkforkian (late
Paleocene) mammals, but within a few meters of section, just below the level where the isotopic
composition of pedogenic carbonate nodules reaches its nadir, the mammalMeniscotherium and a
member of the Perissodactyla appear in the record (Gingerich 2003, 2006). A few meters higher,
in the lowest part of the body of the CIE, the rest of the mammalian immigrant taxa appear (Wa0
zone of Gingerich 2003 andGingerich 2006, earliest Eocene). TheWa0 immigrants are abundant
from their first appearances, retain their abundance for millions of years subsequent to the PETM,
and cause a major shift in mammalian trophic adaptations and body-size distributions (Clyde
& Gingerich 1998). The Wa0 fauna has an unusually high abundance of digging mammals—
taeniodonts and palaeanodonts (Gingerich 1989).
The Polecat Bench section has been calibrated to time by correlating the shape of the CIE
measured in paleosol carbonate nodules with the orbitally calibrated CIE in bulk carbonate from
ODP Site 690 (Bains et al. 2003), by interpolating stratigraphic thickness between magnetochron
boundaries (Secord et al. 2006), and by cyclostratigraphy using paleosol color (Aziz et al. 2008).
These methods give sedimentation rates of 0.3–0.5 m ka−1, indicating the onset of the CIE was
10–15 ka before the first mammalian immigrants (includingMeniscotherium) and 18–27 ka before
the lowest Wa0 fauna.
The source of the PETM mammalian immigrants to North America is not fully known, but
Bowen et al. (2002) showed that the extinct familyHyaenodontidae occurs just below the base of the
CIE in the Hengyang Basin of south China, indicating that this group moved from Asia to North
America. Although not conclusive, mammalian faunas from Inner Mongolia are consistent with
the simultaneous appearance of the Primates, Artiodactyla, and Perissodactyla in Asia and North
America (Bowen et al. 2005). The continent of origin for these three orders is still unknown (Clyde
et al. 2003). Studies by Smith et al. (2006) and Beard (2008) purport to fix the direction of range
expansion in primates by correlating their earliest occurrences in Europe,North America, andAsia
using the shape of the CIE, but this approach is hampered by differences in the shape of the CIE
in each area that are caused by local changes in depositional rate and by stratigraphic uncertainty
in the level of first occurrence of primates resulting from the small number of specimens.
Mammals were not the only vertebrate group that underwent major compositional change at
the PETM. At least seven species of lizards described from the PETM in Wyoming were recent
immigrants to the region, and six of the seven belong to groups whose living members are found
only in frost-free areas of the Americas (Smith 2009).With one possible exception, lizard dispersal
during the PETMwas intracontinental rather than intercontinental. Turtles also have several first
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occurrences during or immediately after the PETM in Wyoming, including lineages with both
Asian and American distributions (Bourque et al. 2008, Holroyd et al. 2001).
In addition to the dispersal event, mammalian faunas of theWa0 zone are distinctive because of
their small body size, with 11 lineages, both immigrants and natives, known from individuals about
half the body weight of their pre- and/or post-PETM relatives (Chester et al. 2010; Gingerich
1989, 2006; Strait 2001). In the southeastern Bighorn Basin, dense sampling of perissodactyls
shows that body size decreased dramatically from their first occurrence near the onset of the
CIE to the middle of the body of the CIE, which in marine records is the warmest part of the
PETM (Secord et al. 2008). The rapid “dwarfing” could be related to increasing temperature
and/or to a decrease in the protein content of plants as a consequence of the reduced production
of the photosynthetic protein Rubisco under high atmospheric CO2 levels (Gingerich 2006).
However, the latter hypothesis does not appear to account for the dwarfing of omnivorous or
faunivorous mammals, unless change in climate and atmospheric composition also lowered net
primary productivity (Chester et al. 2010), which is directly correlated with mammalian body size
across modern gradients (Yom-Tov & Geffen 2006). Soil faunas also dwarfed during the CIE,
perhaps because of increasing temperature and/or declining productivity (Smith et al. 2009).
Early paleobotanical studies of the PETM found little change across the Paleocene-Eocene
transition (Wing 1998, Wing & Harrington 2001), but they lacked samples from the PETM.
Once megafloras were found within the CIE they proved to be radically different in composition
from those immediately before and after (Wing et al. 2005). Pre- and post-PETM megafloras
are a mix of deciduous and evergreen broad-leaved lineages (e.g., birches, katsura, walnuts, elms,
and laurels) as well as conifers in the bald cypress family (Wing 1998, Wing et al. 1995). PETM
megafloras lack conifers and are dominated both in diversity and abundance by the bean family,
important in dry tropical and subtropical areas today (Figure 4; Smith et al. 2007; Wing et al.
2005, 2009). The floral shift took place very rapidly, between latest Paleocene samples∼8m below
the onset of the CIE and PETM samples 2–3 m above the onset (Wing et al. 2009). This implies
extirpation of resident floodplain plants and expansion of the ranges of “excursion” taxa into the
Bighorn Basin in<30 ka. Range expansions on the order of 1000 km are probable for some PETM
immigrants (Wing et al. 2005). Megafloras from the middle of the PETM (60–80 ka post onset
of the CIE) show continued dominance by Fabaceae and other families with modern relatives in
dry tropical forest, but during the recovery phase of the CIE, many typical late Paleocene–early
Eocene plants returned (Wing et al. 2009). Late or post-CIE floras were once again dominated
by long-ranging Paleocene-Eocene plants, accompanied by a few diagnostic Eocene forms that
were intercontinental immigrants from Europe or Asia (Wing et al. 2005, 2009). The appearance
in late or post-PETM time of taxa from other continents is also documented in palynofloras of
the adjacent Powder River Basin (Wing et al. 2003).
Although the PETM had a dramatic effect on megafloral composition, palynofloral change in
the same sections is subtle, with the bald cypress, birch, and walnut families abundant throughout
the Paleocene-Eocene interval (Harrington 2001; Wing et al. 2003, 2005). PETM palynofloras
differ from those immediately before and after only in rare occurrences of pollen types from the
Gulf Coastal plain, which otherwise do not occur so far north (Wing et al. 2005). The continued
abundance of long-ranging pollen types during the PETM may reflect low pollen production by
the immigrants and high pollen production by the small remaining populations of Paleocene taxa
(Wing et al. 2005).
In other parts of the world, PETMfloral change is known exclusively from pollen assemblages.
Palynofloras cored from Lomonosov Ridge in the Arctic Ocean show decreased abundance of
conifers and increased abundance of angiosperms during the PETM (Sluijs et al. 2006), a pattern
also seen in biomarkers (Schouten et al. 2007). The Kumara-2 core from New Zealand shows
504 McInerney ·Wing
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46 immigrations, body of CIE
(some dry tropical)
22 local
extirpations
2 persistent taxa
5
extinctions
5 immigrations
22 locally 
extirpated taxa returnN
um
be
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3 C
bu
lk
 C
aC
O
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‰
)
Time (ka)
Body of CIE
Body of CIE
Recovery I Recovery IIOnset
c
b
δ13C from
bulk carbonate
ODP1266
Temperature
from δ18O N. truempyi
ODP1263
Plant species ranges
Wyoming vegetation
Marine isotope data
2001000
2.0
1.5
1.0
0.5
0.0
0
40
20
10
15
20
25
Tem
perature (°C)
Time
Figure 4
Vegetation change and floral turnover through the Paleocene-Eocene Thermal Maximum (PETM) in the Bighorn Basin, Wyoming.
(a) During the body of the carbon isotopic excursion (CIE), vegetation was probably more open because of water stress caused by
seasonal precipitation and higher evapotranspiration related to warmer temperature. Conifers were absent during the body of the CIE
but present immediately before the onset and during the recovery phase. Drawing by Zell Stoltzfus. (b) Diagram showing types of plant
species ranges: blue, 5 Paleocene taxa that went extinct near the onset of the CIE; light green, 22 taxa, many related to living temperate
plants, that were not found during the body of the CIE but were present before and after (e.g., the coniferMetasequoia); orange, 46 taxa
that in Wyoming are known only from the CIE (some related to living dry tropical plants); dark green, 5 Eocene immigrants, most
related to living temperate plants; gray, 2 types of plant that occur throughout: leaves of the palm Sabalites and winged fruits of Deviacer.
(c) Bulk carbonate carbon isotope record from Ocean Drilling Program Site 1266 on Walvis Ridge and timescale from Murphy et al.
(2010). Oxygen isotope ratio of benthic foraminifera (N. truempyi ) from Ocean Drilling Program Site 1263 (McCarren et al. 2008) is
on timescale of Murphy et al. (2010). Correlation to the Bighorn Basin is fromWing et al. (2005).
mostly stable plant communities across the PETM with modest changes that include first occur-
rences of the tropical mangrove palm, Nypa, and a warm-climate eudicot Cupaneidites (Crouch
et al. 2009). Palynofloras from northern South America increase in diversity through the late
Paleocene–early Eocene because of diversification within many groups already present in the re-
gion during the late Paleocene ( Jaramillo 2002, Jaramillo et al. 2006). Recently it has been shown
that this diversity increase begins during the PETM, suggesting that tropical plants were not
adversely affected by rising temperatures ( Jaramillo et al. 2010).
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Harrington & Jaramillo (2007) noted declining terrestrial palynomorph diversity across the
Paleocene-Eocene boundary in the Gulf Coastal Plain of Mississippi, and they suggested the
possibility of elevated extinction among paratropical plants during the PETM. Recent work by
van Roij (2009), however, suggests a lower stratigraphic position for the CIE. Therefore, the 35%
decrease in palynomorph diversity could represent the early Eocene cool interval also seen in
Rocky Mountain floras (Wing et al. 2000) rather than warming during the PETM.
Only one study has directly examined the effect of the PETM on trophic interactions among
terrestrial organisms. Currano et al. (2008) found that fossil leaves from the PETM had suffered
significantly higher rates of insect damage than those in late Paleocene and post-PETM–early
Eocene. They also found that specialized types of insect feedingweremore common.This increase
could have been caused by higher temperature (increasing insect populations andmetabolic rates),
higher CO2 (leading to decreased protein in leaves and higher feeding rates), or both (Currano
et al. 2008).
DISCUSSION
Scientists have been aware of the PETM for two decades, and the event has been the topic
of hundreds of research papers, multiple conferences and symposia, scores of news reports, and
thousands of Internet postings.What havewe learned from the PETMabout connections between
the carbon cycle, climate, and biosphere, and how well has it held up as a model for future
anthropogenic global warming?
Research on the PETMhas resoundingly confirmed several basic ideas about the Earth system.
A release of thousands of petagrams of carbon (an amount comparable to that likely to be released
by human activities) did, as expected, cause global warming and deep-ocean acidification (IPCC
2007). The climatic perturbations from the PETMpersisted for∼200 ka, as anticipated for climate
change driven by fossil fuel emission (Archer et al. 2009). Also as predicted, the changing climatic
and oceanographic conditions drove rapid adjustments in the geographic ranges of most forms of
life, regardless of habitat, trophic level, or prior distribution.
There are also some fundamental unresolved questions inPETMresearch, includinghowmuch
carbon was released from which source(s). After initial wide acceptance of a methane clathrate
source, significant objections have emerged. As detailed above the highly depleted carbon isotope
composition of clathrates makes it difficult to satisfy the requirements of carbonate dissolution,
warming, and the CIE. Furthermore, the Paleocene clathrate reservoir may have been too small to
produce the CIE (Buffett & Archer 2004). Future research into pCO2 changes across the PETM
will better constrain climate sensitivity and help evaluate the potential sources of carbon. The
release of carbon from Antarctic permafrost and peat can explain the observed carbonate dissolu-
tion, warming, and CIE and also provides a cohesive explanation for the magnitude and pacing of
the PETM and subsequent hyperthermals (DeConto et al. 2010). In addition, the rapid recovery
of the CIE suggests the role of the biosphere and soil carbon in carbon sequestration (Bowen &
Zachos 2010). These processes have implications for understanding the potential climate impacts
of permafrost thawing and biospheric carbon sequestration in modern climate change (IPCC
2007).
Although interest in the PETM is in large part stimulated and funded because of its parallels
with anthropogenic global warming, differences between the two have come into sharper focus in
recent years. The carbon release that caused the carbon isotope excursion at the PETM probably
took at least 8 ka, roughly 15 times longer than any anticipated anthropogenic carbon release. This
much slower rate of carbon addition translates to much less severe acidification and carbonate
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dissolution in the surface ocean, presumably with less severe consequences for surface-dwelling
organisms during the PETM than in the coming centuries (Ridgwell & Schmidt 2010).
Some aspects of the biotic response to the PETMare predictable, whereas others are surprising.
Many of the biotic effects of the PETM appear to have been driven by increasing temperature,
with no clear effects yet demonstrated from changes in ocean chemistry or atmospheric pCO2.
It is surprising that cool-adapted species already living at higher latitudes before the onset of
the PETM are not known to have experienced major extinctions. Do low extinction levels at
the PETM (except for benthic forams) reflect the existence of cool refugia or the ability of high-
latitude taxa to adapt to warmer climate during the PETM?Another unexpected pattern is that the
highest levels of extinction are observed in benthic forams—a group that is highly dispersable and
tends to have large geographic ranges (Thomas 2007). High extinction levels suggest that changes
in deep benthic environments left few refugia for benthic forams, but relatively low extinction in
other deep-sea taxa point to conditions that were particularly harsh for benthic forams.
This absence of significant extinction in most groups is particularly interesting in light of the
predictions of substantial future extinction with anthropogenic global warming. Projections are
∼35% extinction globally across many groups for >2◦C of warming (Thomas et al. 2004). The
higher rate of environmental change in future projections does not explain the higher extinction
rates in the models, because Thomas et al. (2004) effectively assumed instantaneous dispersal.
Projected future rates of extinctionmay be higher than observed fossil rates because the projections
include species with small geographic ranges (narrow climatic tolerances) that are unlikely to be
detected in the fossil record at all. It is also possible that the ability of species to exist outside their
current climatic envelopes has been underestimated. Low levels of extinction in the face of rapid
environmental change during the Quaternary pose a similar challenge to modeled extinctions
under future greenhouse warming (Botkin et al. 2007).
Rapid evolution is an expected consequence of rapid environmental change (Davis et al. 2005),
and though the selective factors that caused rapid evolution during the PETM are not well under-
stood, dwarfing was common in terrestrial vertebrates (Chester et al. 2010; Gingerich 1989, 2003),
soil-living invertebrates (Smith et al. 2009), and some forams (Kaiho et al. 2006). In both marine
and terrestrial realms, declining productivity/food availability is a potential cause of evolution to
smaller body size (Chester et al. 2010, Kaiho et al. 2006). Even the evolution of larger size in some
planktic forams may relate to lower productivity of surface waters in the open ocean and the need
for algal symbiotes (Kaiho et al. 2006). If there is a generalization to be made about what forces
evolutionary change during the PETM it would seem to be this: A warmer world is a hungrier
world.
SUMMARY POINTS
1. The Paleocene-Eocene Thermal Maximum, which took place ∼56 Mya and lasted for
∼200 ka, stands as the most dramatic geological confirmation of greenhouse theory—
increased CO2 in the atmosphere warmed Earth’s surface.
2. The large release of organic, 13C-depleted carbon caused a global carbon isotopic excur-
sion, widespread deep-ocean acidification, and carbonate dissolution.
3. Carbonwas removed from the exogenic pool on a timescale of∼100 ka, primarily through
silicate weathering and eventual precipitation of carbonate in the ocean and/or uptake
by the biosphere and subsequent burial as organic carbon.
www.annualreviews.org • Paleocene-Eocene Thermal Maximum 507
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4. Warming associated with the carbon release implies approximately two doublings
of atmospheric pCO2 unless climate sensitivity was significantly different during the
Paleogene.
5. Although there was a major extinction of benthic foraminifera, most groups of organisms
did not suffer mass extinction.
6. Geographic distributions of most kinds of organisms were radically rearranged by 5–8◦C
of warming, with tropical forms moving poleward in both marine and terrestrial realms.
7. Rapidmorphological change occurred in bothmarine and terrestrial lineages, suggesting
that organisms adjusted to climate change through evolution as well as dispersal and local
extirpation. Where best understood, these evolutionary changes appear to be responses
to nutrient and/or food limitation.
8. Research on the PETM and other intervals of rapid global change has been driven by the
idea that they provide geological parallels to future anthropogenic warming, but much
remains to be done to gain information that can be acted on.
DISCLOSURE STATEMENT
The authors are not aware of any affiliations, memberships, funding, or financial holdings that
might be perceived as affecting the objectivity of this review.
ACKNOWLEDGMENTS
We thank ourmany colleagues who shared their ideas, comments, and in-press manuscripts freely.
We apologize for the hundreds of papers we were aware of but could not cite. Thank you to Kate
Freeman, Lee Kump, andWill Clyde for thoughtful discussions and comments on themanuscript.
Thanks also to Zell Stoltzfus, who drew the reconstruction, and Brandon Murphy and Heather
McCarren, who provided data for Figure 4.
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Annual Review
of Earth and
Planetary Sciences
Volume 39, 2011 Contents
Plate Tectonics, the Wilson Cycle, and Mantle Plumes: Geodynamics
from the Top
Kevin Burke ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! 1
Early Silicate Earth Differentiation
Guillaume Caro ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! !31
Building and Destroying Continental Mantle
Cin-Ty A. Lee, Peter Luffi, and Emily J. Chin ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! !59
Deep Mantle Seismic Modeling and Imaging
Thorne Lay and Edward J. Garnero ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! !91
Using Time-of-Flight Secondary Ion Mass Spectrometry
to Study Biomarkers
Volker Thiel and Peter Sjo¨vall ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! 125
Hydrogeology and Mechanics of Subduction Zone Forearcs:
Fluid Flow and Pore Pressure
Demian M. Saffer and Harold J. Tobin ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! 157
Soft Tissue Preservation in Terrestrial Mesozoic Vertebrates
Mary Higby Schweitzer ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! 187
The Multiple Origins of Complex Multicellularity
Andrew H. Knoll ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! 217
Paleoecologic Megatrends in Marine Metazoa
Andrew M. Bush and Richard K. Bambach ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! 241
Slow Earthquakes and Nonvolcanic Tremor
Gregory C. Beroza and Satoshi Ide ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! 271
Archean Microbial Mat Communities
Michael M. Tice, Daniel C.O. Thornton, Michael C. Pope,
Thomas D. Olszewski, and Jian Gong ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! 297
Uranium Series Accessory Crystal Dating of Magmatic Processes
Axel K. Schmitt ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! 321
viii
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EA39-FrontMatter ARI 24 March 2011 6:51
A Perspective from Extinct Radionuclides on a Young Stellar Object:
The Sun and Its Accretion Disk
Nicolas Dauphas and Marc Chaussidon ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! 351
Learning to Read the Chemistry of Regolith to Understand the
Critical Zone
Susan L. Brantley and Marina Lebedeva ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! 387
Climate of the Neoproterozoic
R.T. Pierrehumbert, D.S. Abbot, A. Voigt, and D. Koll ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! 417
Optically Stimulated Luminescence Dating of Sediments over the Past
200,000 Years
Edward J. Rhodes ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! 461
The Paleocene-Eocene Thermal Maximum: A Perturbation of Carbon
Cycle, Climate, and Biosphere with Implications for the Future
Francesca A. McInerney and Scott L. Wing ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! 489
Evolution of Grasses and Grassland Ecosystems
Caroline A.E. Stro¨mberg ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! 517
Rates and Mechanisms of Mineral Carbonation in Peridotite:
Natural Processes and Recipes for Enhanced, in situ CO2 Capture
and Storage
Peter B. Kelemen, Juerg Matter, Elisabeth E. Streit, John F. Rudge,
William B. Curry, and Jerzy Blusztajn ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! 545
Ice Age Earth Rotation
Jerry X. Mitrovica and John Wahr ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! 577
Biogeochemistry of Microbial Coal-Bed Methane
Dariusz Stra˛poc´, Maria Mastalerz, Katherine Dawson, Jennifer Macalady,
Amy V. Callaghan, Boris Wawrik, Courtney Turich, and Matthew Ashby ! ! ! ! ! ! ! ! ! ! 617
Indexes
Cumulative Index of Contributing Authors, Volumes 29–39 ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! 657
Cumulative Index of Chapter Titles, Volumes 29–39 ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! 661
Errata
An online log of corrections to Annual Review of Earth and Planetary Sciences articles
may be found at http://earth.annualreviews.org
Contents ix
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