By the year 2400, it is predicted that humans will have released about
5,000 gigatonnes of carbon (Gt C) to the atmosphere since the start
of the industrial revolution if fossil-fuel emissions continue una-
bated and carbon-sequestration efforts remain at current levels1. This
anthropogenic carbon input, predominantly carbon dioxide (CO2),
would eventually return to the geosphere through the deposition
of calcium carbonate and organic matter2. Over the coming mil-
lennium, however, most would accumulate in the atmosphere and
ocean. Even if only 60% accumulated in the atmosphere, the par-
tial pressure of CO2 (pCO2) would rise to 1,800 parts per million by
volume (p.p.m.v.) (Fig. 1). A greater portion entering the ocean would
decrease the atmospheric burden but with a consequence: significantly
lower pH and carbonate ion concentrations of ocean surface layers1
(Fig. 1).
A marked increase in atmospheric pCO2 would increase mean global
temperature, thereby affecting atmospheric and oceanic circulation,
precipitation patterns and intensity, the coverage and thickness of sea
ice, and continental ice-sheet stability. However, forecasting the tim-
ing and magnitude of these responses is challenging because they can
be nonlinear. Of particular concern are potential positive feedbacks
that could amplify increases in the concentrations of greenhouse gases
— water, CO2, methane and nitrous oxide (N2O) — effectively esca-
lating climate sensitivity to initial anthropogenic carbon input3. For
example, ocean surface warming and freshwater discharge at high lati-
tudes could slow the exchange of shallow and deep water in the ocean,
impeding both abiotic and biotic removal of anthropogenic carbon
from the atmosphere. Potential negative feedbacks are also garnering
great interest. As a possible counterbalance to decreased density of
surface water on a warmer Earth, stronger zonal winds might increase
ocean overturning (see page 286).
Observations of modern and Holocene (the past 10,000 years
or so) climates have provided essential constraints for understand-
ing climate dynamics and a baseline for predicting future responses
to carbon input. But such observations can provide only limited
insight into the response of climate to massive, rapid input of CO2.
To evaluate climate theories more thoroughly, particularly with regard
to feedbacks and climate sensitivity to pCO2, it is desirable to study
samples obtained when CO2 concentrations were high (approaching
or exceeding 1,800 p.p.m.v.) and to make observations for intervals
longer than those of ocean overturning and carbon cycling (more
than 1,000 years)4. Earth scientists have therefore turned increasingly
to ancient time intervals, particularly those in which pCO2 was much
higher than now, in which pCO2 changed rapidly, or both. Recent
reconstructions of Earth’s history have considerably improved our
knowledge of known ‘greenhouse’ periods and have uncovered several
previously unknown episodes of rapid emissions of greenhouse gases
and abrupt warming.
Cenozoic greenhouse climates
The Cenozoic era, the last 65 million years of Earth’s history, provides
an ideal backdrop from which to understand relationships between
carbon cycling and climate. In contrast to the present day, much of the
early Cenozoic was characterized by noticeably higher concentrations
of greenhouse gases, as well as a much warmer mean global temperature
and poles with little or no ice5,6 (Fig. 2). The extreme case is the Early
Eocene Climatic Optimum (EECO), 51–53 million years ago, when
pCO2 was high and global temperature reached a long-term maximum.
Only over the past 34 million years have CO2 concentrations been
low, temperatures relatively cool, and the poles glaciated. This long-
term shift in Earth’s climatic state resulted, in part, from differences in
volcanic emissions, which were particularly high during parts of the
Palaeocene and Eocene epochs (about 40–60 million years ago) but
have diminished since then. Changes in chemical weathering of silicate
rocks were also important7. On long timescales, this process sequesters
CO2, preventing concentrations from rising too high or from falling
too low. As the atmospheric CO2 concentration rises, temperature and
precipitation increase and thereby enhance chemical weathering; as the
concentration declines, temperature and precipitation decrease, slowing
weathering. Whereas other processes (such as the oxidation and burial
of organic carbon) change CO2 concentrations, the negative weathering
feedback loop maintains Earth’s climate within a habitable range over
millions of years and longer7.
On shorter timescales, atmospheric CO2 concentration and tem-
perature can change rapidly, as demonstrated by a series of events dur-
ing the early Cenozoic known as hyperthermals. These were relatively
brief intervals (less than a few tens of thousands of years) of extreme
global warmth and massive carbon addition but with widely differing
scales of forcing and response. During the most prominent and best-stud-
ied hyperthermal, the Palaeocene–Eocene Thermal Maximum (PETM;
about 55 million years ago), the global temperature increased by more
than 5 °C in less than 10,000 years6 (Fig. 3). At about the same time, more
than 2,000 Gt C as CO2 — comparable in magnitude to that which could
occur over the coming centuries — entered the atmosphere and ocean.
Evidence for this carbon release is found in sedimentary records
across the event. This includes a rapid and pronounced decrease in the
13C/12C ratio of carbonate and organic carbon across the globe (that
is, a negative carbon isotope excursion) and a prominent drop in the
carbonate content of marine sediment deposited at several thousands
of metres water depth (that is, a deep-sea dissolution horizon)8. The
first observation indicates injection into the atmosphere or ocean of
a very large mass of 13C-depleted carbon, affecting the composition of
the global carbon cycle. The second observation is a telltale signature
of ocean acidification. The entire event lasted less than 170,000 years.
Given the residence time of carbon (the average time a carbon atom
spends in the ocean; about 100,000 years), this is consistent with a fast
An early Cenozoic perspective on greenhouse
warming and carbon-cycle dynamics
James C. Zachos, Gerald R. Dickens & Richard E. Zeebe
Past episodes of greenhouse warming provide insight into the coupling of climate and the carbon cycle and
thus may help to predict the consequences of unabated carbon emissions in the future.
279
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ocean overturning and increased surface temperatures should have
decreased the flow of dissolved oxygen to deep water. Several direct
lines of evidence, such as laminated sediment in cores from the Car-
ibbean and central Arctic regions, suggest that dissolved oxygen did
indeed decrease across the PETM. Moreover, the PETM coincided with
a major extinction of benthic foraminiferans, with widespread oxygen
deficiency in the ocean as a possible cause17.
With such ocean conditions, greater preservation and burial of solid
organic carbon in deep-sea sediments might be predicted, effectively
countering the decreased carbon flux from surface waters. However, this
has not been documented. Two largely unexplored processes involving
the microbial decomposition of organic carbon, both functioning as
additional positive feedbacks, might operate during times of massive
carbon input and rapid warming. Carbonate dissolution in the deep
ocean decreases sedimentation rates, exposing organic carbon at or near
the sea floor for a longer duration, and warming of deep waters will
accelerate overall microbial activity and the consumption of organic
carbon. Future investigations might therefore focus specifically on the
evidence for changes in ocean overturning, oxygen deficiency and the
burial of organic carbon.
The positive feedbacks of greatest concern for understanding overall
global warming may be those that could release hundreds to thousands
of gigatonnes of carbon after initial warming11–13. The large masses of
organic carbon stored in soils (for example, as peat) or sediments of shal-
low aquatic systems (for example, wetlands, bogs and swamps) represent
a potential carbon input, should regions that were humid become drier.
Rapid desiccation or fire could release carbon from these reservoirs at
rates faster than carbon uptake by similar environments elsewhere. By
contrast, regions that once were dry might emit methane as they become
wetter18. Methane might also enter the ocean or atmosphere through the
–1
0
1
2
3
4
5
0 10 20 30 40 50 60
Miocene Oligocene
4
0
8
12Antarctic ice sheets
Full scale and permanent
Partial or ephemeral
PETM
(ETM1)
4,000
0
1,000
2,000
3,000
5,000
Boron
Alkenones
0a
b
10 20 30 40 50 60
Anthropogenic peak (5,000 Gt C)
ETM2
PalaeoceneEocene
Mid-Eocene
Climatic Optimum
Early Eocene
Climatic Optimum
Nahcolite
Trona
Northern Hemisphere ice sheets
?
Plio-
cene
Pleistocene
Ic
e-
fr
ee
te
m
pe
ra
tu
re
(
°C
)
A
tm
os
ph
er
ic
C
O
2,
p
C
O
2
(p
.p
.m
.v
.)

Age (millions of years ago)
Mid-Miocene
Climatic Optimum
δ18
O
(‰
)
CO2 proxies
Figure 2 | Evolution of atmospheric CO2 levels and global climate over
the past 65 million years. a, Cenozoic pCO2 for the period 0 to 65 million
years ago. Data are a compilation of marine (see ref. 5 for original sources)
and lacustrine24 proxy records. The dashed horizontal line represents the
maximum pCO2 for the Neogene (Miocene to present) and the minimum
pCO2 for the early Eocene (1,125 p.p.m.v.), as constrained by calculations of
equilibrium with Na–CO3 mineral phases (vertical bars, where the length
of the bars indicates the range of pCO2 over which the mineral phases are
stable) that are found in Neogene and early Eocene lacustrine deposits24.
The vertical distance between the upper and lower coloured lines shows the
range of uncertainty for the alkenone and boron proxies. b, The climate for
the same period (0 to 65 million years ago). The climate curve is a stacked
deep-sea benthic foraminiferal oxygen-isotope curve based on records from
Deep Sea Drilling Project and Ocean Drilling Program sites6, updated with
high-resolution records for the interval spanning the middle Eocene to
the middle Miocene25–27. Because the temporal and spatial distribution of
records used in the stack are uneven, resulting in some biasing, the raw data
were smoothed by using a five-point running mean. The δ18O temperature
scale, on the right axis, was computed on the assumption of an ice-free
ocean; it therefore applies only to the time preceding the onset of large-scale
glaciation on Antarctica (about 35 million years ago). The figure clearly
shows the 2-million-year-long Early Eocene Climatic Optimum and the
more transient Mid-Eocene Climatic Optimum, and the very short-lived
early Eocene hyperthermals such as the PETM (also known as Eocene
Thermal Maximum 1, ETM1) and Eocene Thermal Maximum 2 (ETM2;
also known as ELMO). ‰, parts per thousand.
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NATURE|Vol 451|17 January 2008 YEAR OF PLANET EARTH FEATURE

dissociation of gas hydrate in marine sediment. This feedback would
probably take several thousands of years to initiate because heat must
be propagated by ocean advection to water depths at which hydrates
can form (more than 1 km in the early Cenozoic), and then by diffusion
into sediments. However, the amount of methane that could be liber-
ated is enormous, and after gas hydrate dissociation was initiated, the
flux might proceed rapidly as overpressured pore waters triggered fluid
expulsion or sediment slides on the sea floor11.
These potential carbon-cycle feedbacks for amplifying warmth are not
fully understood. In fact, demonstrating that such feedbacks have oper-
ated in the past remains a major challenge. The abrupt negative carbon
isotope excursions that mark the hyperthermals and attest to a massive
input of isotopically depleted carbon cannot be used alone to identify the
source, especially if more than one source existed. Records of geochemical
or physical fingerprints, such as hopanoids from methanotrophs or char-
coal from wildfires, would help18. Constraining the rate and mass of car-
bon released, for example by quantifying changes in ocean carbonate
chemistry, is also essential for identifying sources19.
The PETM and other hyperthermals should also provide insight into
the longer-term response of the carbon cycle to massive inputs of car-
bon, including the primary negative feedbacks that temporarily and
permanently sequester carbon. Various simulations of the long-term
fate of anthropogenic carbon emissions show consistent results. After
the cessation of emissions, and a peak in atmospheric pCO2, the ocean
steadily absorbs much of the carbon, although with a decrease in pH and
carbonate-ion concentration (Fig. 1). The carbonate-ion concentration
is restored by dissolution of carbonate on the sea floor within several
thousand years, but dissolved inorganic carbon and alkalinity remain
high for tens of thousands of years afterwards. As a consequence, atmos-
pheric pCO2 does not return to pre-anthropogenic values but stabilizes at
levels at least 50% higher than before the carbon injection (Fig. 1).
Marine-sediment records that span the PETM show features consist-
ent with this pattern. The initial release of carbon, as represented by the
carbon isotope excursion, is accompanied by widespread and significant
dissolution of seafloor carbonate and a net deficit in deep-sea carbon-
ate accumulation (Fig. 3). This is followed by an increase in carbonate
accumulation at many locations, presumably reflecting a recovery of car-
bonate ion concentration8. Interestingly, carbonate accumulation during
this recovery phase seems greater than before carbon injection, suggest-
ing carbonate oversaturation. Although no detailed reconstructions of
pCO2 are available for the PETM, surface temperatures remain warm
for thousands of years after the input of carbon seems to cease. Thus,
at first glance, observations of the PETM support the theory about the
long-term fate of fossil-fuel CO2. The carbonate ‘overshoot’ represents a
negative feedback, probably through enhanced silicate weathering and
delivery of dissolved calcium and bicarbonate to the ocean. Gauging
the sensitivity of this effect will enable the establishment of constraints
on long-term forecasts for the carbon cycle following anthropogenic
carbon emissions.
Climate sensitivity
Early Cenozoic climate has received considerable interest because
the response of climate to a broad range of high atmospheric values
of pCO2 (probably 1,000 to more than 2,000 p.p.m.v.) can be examined.
One feature common to all greenhouse periods, whether transient or
long-lived, is exceptionally warm poles15. In the more extreme cases,
the EECO and PETM, high-latitude temperatures were substantially
higher than can be simulated by models without unreasonably high pCO2
(refs 20, 21). Somehow, models are not precisely simulating processes
critical to poleward heat transport, albedo, or polar heat retention at
higher greenhouse gas levels. Modified ocean heat transport has been
investigated and found to be incapable of transporting heat fast enough
to compensate for polar heat loss22. In contrast, polar stratospheric
clouds, which might have been more extensive during the greenhouse
intervals because of higher concentrations of methane in the atmos-
phere, seem to be very effective at trapping heat20. Similarly, non-CO2
greenhouse gases, which are usually neglected, may have had a major
role. Recent theoretical and experimental studies indicate that, under
high pCO2, background concentrations of trace gases such as methane
and N2O should be higher because of greater production under warmer
and wetter conditions (that is, more extensive wetlands) and because
of lower rates of oxidation in the atmosphere (resulting from lower
emissions of volatile organic compounds by plants)21. Collectively, such
physical and biochemical feedbacks would tend to enhance the sensitiv-
ity of climate to changes in CO2 and might explain the unusual polar
warmth of the early Cenozoic.
Another prominent feature of the transient greenhouse episodes,
specifically the PETM, are marked shifts in the distribution and inten-
sity of precipitation, as inferred from fossil vegetation and other proxy
data. Most regions, particularly in middle to high latitudes, experienced
a shift towards wetter climates. However, the response on a regional
scale was far more complex. For example, recent studies show that some
regions, such as the western interior of North America, became drier at
the onset of the PETM, whereas other regions, such as western Europe,
experienced increased extreme precipitation events and massive flood-
ing23. These palaeo-observations imply a high degree of sensitivity in the
hydrological cycle to extreme changes in pCO2 and temperature. Addi-
tional documentation of precipitation changes for climatically sensi-
tive regions during Eocene greenhouse episodes could prove useful for
assessing how well models simulate extremes in climate.
56.055.555.054.554.0
Age (millions of years ago)
3.0a
b
c
2.0
1.0
0
–1.0
–2.0
δ13
C
(
‰
)
δ18
O
(
‰
)
1.0
–1.0
–0.5
0
0.5
14
12
10
8
Te
m
pe
ra
tu
re
(
°C
)
0
20
40
60
80
100
C
aC
O
3
(%
)
1262 (4.8 km)
1263 (2.6 km)
South Atlantic
(water depth)
690
865
525
527
Southern Ocean
Central Pacific
South Atlantic
Figure 3 | Low-resolution marine stable-isotope records of the PETM and
the carbon isotope excursion, together with the seafloor sediment CaCO3
record. The carbon isotope (a) and oxygen isotope (b) records are based
on benthic foraminiferal records (see ref. 6 for original sources), and the
CaCO3 records (c) are from drill holes in the South Atlantic
8. Panel b also
shows inferred temperatures. Ocean drilling site locations are indicated in
the keys. The decrease in sedimentary CaCO3 reflects increased dissolution
and indicates a severe decrease in seawater pH (that is, ocean acidification).
The base of the CaCO3 dissolution horizon is below the onset of the carbon
isotope excursion because most of the carbonate dissolution involved
uppermost Palaeocene sediments that were deposited before the event
(chemical erosion). Panels a and b adapted, with permission, from ref. 6.
282
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Outlook for the future
If fossil-fuel emissions continue unabated, in less than 300 years pCO2
will reach about 1,800 p.p.m.v., a level not present on Earth for roughly
50 million years. Both the magnitude and the rate of rise complicate the
goal of accurately forecasting how the climate will respond. Foremost
among the challenges that must be overcome to achieve this goal is the
development of a deeper understanding of the complex interactions that
link the climate system with the biogeochemical cycles, specifically the
role of positive and negative feedbacks. The occurrence of past green-
house warming events provides one opportunity to test theory about the
physical and biogeochemical interactions in rapidly shifting systems.
There are of course limitations on which facets of theory and models
can be tested given uncertainties in proxies and the limited spatial and
temporal resolution of palaeorecords. Nevertheless, the past greenhouse
events provide glimpses of the future. Until the most salient features of
these events, for example the global patterns of carbonate deposition or
the extreme polar warmth, can be replicated with dynamical models,
forecasts of climate beyond the next century (that is, under extreme
greenhouse gas levels) should be viewed with caution, and efforts to
comprehend the underlying physics and biogeochemistry of the cou-
pling between climate and the carbon cycle should be hastened. ■
James C. Zachos is in the Department of Earth and Planetary Sciences,
University of California at Santa Cruz, Santa Cruz, California 95060,
USA. Gerald R. Dickens is in the Department of Earth Sciences, Rice
University, Houston, Texas 77005, USA. Richard E. Zeebe is at the School
of Ocean and Earth Science and Technology, University of Hawaii at
Manoa, 1000 Pope Road, MSB 504, Honolulu, Hawaii 96822, USA.
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Author Information Reprints and permissions information is available at
npg.nature.com/reprints. Correspondence should be addressed to J.C.Z. and R.E.Z.
(jzachos@es.ucsc.edu; zeebe@hawaii.edu).
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