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3
Climate Transitions, Tipping Points, and
the Point of No Return
Because of the extended timescale—several centuries—necessary for
climate to adjust to an increase in atmospheric CO2, the current icehouse
climate is out of equilibrium with long-term CO2 forcing (Hansen et al.,
2008). As the planet continues to warm, it may be approaching a critical
climate threshold beyond which rapid (decadal-scale) and potentially
catastrophic changes may occur that are not anticipated—because of
complex feedback dynamics and existing computational limitations—by
climate models that are tuned to modern conditions. This chapter focuses
on the insights that can be gleaned from the deep-time geological archive
of climate change concerning such thresholds, with particular focus on
the major societal questions noted in Chapter 1: How soon, abrupt, and
dramatic will climate change be, and how long will the new climate states
persist?
Climate modeling efforts and the geological record provide plenty
of evidence for climate system thresholds, or “tipping points” (Box 3.1),
beyond which rapid changes can occur without any additional forcing
(Hansen et al., 2008; Lenton et al., 2008). Components of the climate system
that are particularly vulnerable to being forced by increasing atmospheric
CO2 across a threshold into a new state include the loss of Arctic summer
sea ice, the stability of the Greenland and West Antarctic ice sheets, the
vigor of the meridional overturning circulation in the North Atlantic and
around Antarctica, the extent of Amazon and boreal forests, and the vari-
ability of the El Niño-Southern Oscillation (ENSO) (Lenton et al., 2008).
The changes in state across such “tipping points” are typically accelerated
relative to the apparent rate of forcing, are accompanied by large-scale
63
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64 UNDERSTANDING EARTH’S DEEP PAST
BOX 3.1
Tipping Points and the Point of No Return
There are sound theoretical reasons to think that tipping points across
climatic thresholds exist (Gladwell, 2000; NRC, 2002). Examples of
threshold behaviors include thermohaline circulation modifications, ice
sheet instabilities, sea ice instabilities, soil-moisture feedbacks, and the
onset of high-latitude convection and associated high-level cloud forcing.
Hansen et al. (2008) introduced the term “tipping element” to describe
subcontinental-scale subsystems of the Earth system that are susceptible
to being forced into a new state by small perturbations. Tipping level—the
magnitude of climate forcing beyond which, if sustained, abrupt climate
change will eventually occur—is differentiated from “point of no return.”
If the tipping level is exceeded for only a brief period of time, the original
state of the system can be restored. More persistent forcing can push the
system to the “point of no return,” where a reduction of the forcing below
the tipping level is ineffective in halting the climate shift (Figure 3.1). This
irreversibility of the system response is referred to as hysteresis (NRC, 2002).
FIGURE 3.1 Equilibrium states of a “system” (valleys) in response to
gradual anthropogenic CO2 forcing (progressing from dark to light blue).
The curvature of the valley is inversely proportional to the system’s
response time (τ) to small perturbations. A threshold is reached when
the valley becomes shallower and finally vanishes causing the ball to
abruptly roll to a new state (to the left).
SOURCE: Lenton et al. (2008), ©National Academy of Sciences, U.S.A.
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65
CLIMATE TRANSITIONS, TIPPING POINTS, AND THE POINT OF NO RETURN
impacts on ecological systems, and typically involve hysteresis (Lenton
et al., 2008).
ICEHOUSE-GREENHOUSE TRANSITIONS
The following sections describe four periods of past climate change—
icehouse-to-greenhouse or greenhouse-to-icehouse transitions—that were
driven by slow (long-term) climate forcing across a critical threshold that
led to abrupt and highly variable climate responses, as examples of what
can be gleaned from the deep-time geological record of climate change
and the scientific challenges that persist.
Initiation of the Cenozoic Icehouse
The early Cenozoic greenhouse Earth was plunged from a pro-
tracted state of warmth into its current glacial state 33.7 million years
ago (Ma), at the Eocene-Oligocene boundary. The transition from a
relatively deglaciated climate state to one in which the Antarctic ice
sheet grew to between 40 and 160 percent of its modern size occurred
within ~200,000 to 300,000 years (Coxall et al., 2005; Liu et al., 2009b). A
long-term decrease in CO2, commencing after the Early Eocene Climate
Optimum at 52 Ma, has been proposed as the main cause of this cool -
ing trend (Box 3.2) (Edmond and Huh, 2003; Kent and Muttoni, 2008).
A CO2 decrease through yet another apparent threshold (from as high
as 415 ppmv [parts per million by volume] in the early Pliocene to
~280 ppmv; Pagani et al., 2010; Seki et al., 2010) most probably accounted
for the initiation and growth of northern hemisphere ice sheets at around
3 Ma (DeConto et al., 2008; Lunt et al., 2008).
All of the elements of a tipping point climate transition are recorded
by this greenhouse-to-icehouse turnover (Kump, 2009). As the climate sys-
tem reorganized itself, it experienced an overshoot (the Oi-1 climate event)
into a deep glacial, which was colder and with larger ice sheets than would
be sustained during the less extreme conditions of the glaciated Oligocene
(Zachos et al., 1996). The calcium carbonate compensation depth in the
oceans deepened substantially in two 40-thousand-year (ky) long steps
(separated by 200 ky) that occurred synchronously with the stepwise
onset of major permanent ice sheets in Antarctica (Coxall et al., 2005). This
instability in the climate system persisted for ~200 to 300 ky (Zachos et
al., 2001b) and caused major changes in ocean and atmosphericic circula -
tion with widespread effects on most marine and terrestrial ecosystems
(Pearson et al., 2008). Such a characteristic response of a homeostatic feed-
back system implies an underlying dynamic that still remains to be fully
understood but could result from changes in, and the interplay between,
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66 UNDERSTANDING EARTH’S DEEP PAST
BOX 3.2
Separating the Influence of Ice Volume and Temperature
The gradual cooling from the hothouse of the Early Eocene Climate Opti-
mum (52 Ma) to the onset of Oligocene glaciation in Antarctica (~34 Ma)
was first inferred from a long-term global trend of increasing benthic
foraminiferal δ18O values (Figure 3.2). The temperature of the deepest water
in the oceans—an indication of global climate—was at least 10°C higher in
the early Eocene than it is today. The cooling trend was disrupted several
times by transient warming events in the Eocene and also by an abrupt shift
toward heavier isotopic values (~1 to 1.5‰ [parts per thousand] increase in
δ18O in all records) at the Eocene-Oligocene boundary (referred to as the
“Oi-1 overshoot”; Zachos et al., 2001b), with this transient cooling a result
of some combination of rapid East Antarctic ice sheet growth and global
cooling (Zachos et al., 2001a; Coxall et al., 2005). Marine carbon isotope
compositions and CaCO3 accumulation rates also exhibit the distinctive
“overshoot,” suggesting teleconnections between the southern hemisphere
high latitudes and the tropical ocean (Coxall et al., 2005).
A number of additional proxies have been used to separate, or decon-
volve, the effects of ice sheet growth from cooling—sequence stratigraphy
to assess sea level change (e.g., Kominz and Pekar, 2001); marine geo-
chemical proxies of temperature, including Mg/Ca ratios of foraminiferal
calcite (e.g., Lear et al., 2000; Katz et al., 2008) and biomarkers (spores
and pollen) in marine sediments (e.g., Liu et al., 2010); as well as terres-
trial climate reconstructions based on oxygen isotopes in teeth and bones
(e.g., Zanazzi and Kohn, 2008). The latest assessments indicate that the
greenhouse-to-icehouse transition occurred in a series of steps with increas-
ing influence of ice volume (Lear et al., 2008) and that cooling preceded
ice sheet expansion, with maximum ice sheet size perhaps as much as 15
percent greater than today’s Antarctic ice sheet (Pälike et al., 2006a; Liu
et al., 2009b). A threshold was likely reached through a combination of
orbitally driven changes in summer insolation and declining atmospheric
CO2 levels (DeConto and Pollard, 2003), although oceanic gateway open-
ing and the thermal isolation of Antarctica may have played a role (Barker
et al., 2007; Jovane et al., 2007).
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67
CLIMATE TRANSITIONS, TIPPING POINTS, AND THE POINT OF NO RETURN
A 10 20 30 40 50 60
0
5,000
CO 2 proxies
Boron
4,000 Alkenones
Atmospheric CO 2, pCO 2 (p.p.m.v.)
Nahcolite
Trona
3,000
Anthropogenic peak (5,000 Gt C)
2,000
1,000
0
B –1
Partial or ephemeral
Full scale and permanent
12
Antarctic ice sheets
0
Ice-free temperature (°C)
Northern Hemisphere ice sheets
8
?
1 Early Eocene
Climatic Optimum
δ18O (‰)
4
2
ETM2
Mid-Eocene PETM
Climatic Optimum (ETM1) 0
3
Mid-Miocene
Climatic Optimum
4
Pleistocene
5 Plio- Palaeocene
Miocene Oligocene Eocene
cene
10 20 30 40 50 60
0
Age (millions of years ago)
FIGURE 3.2 Relationship between atmospheric CO2 (A) and climate (B)
through the Cenozoic. The upper panel shows reconstructed pCO2 from
marine and lacustrine proxy records; the dashed line is maximum pCO2
for the Neogene estimated by equilibrium calculations using lacustrine
mineral phases (Lowenstein and Demicco, 2006). The climate curve in
the lower panel is a composite of deep-sea benthic foraminiferal oxygen-
isotope records, smoothed using a five-point running mean (Zachos et
al., 2001a, 2008). The temperature scale on the right axis was calculated
for an “ice-free ocean,” and is thus applicable solely to the pre-Oligocene
portion of record.
SOURCE: Zachos et al. (2008), reprinted by permission of Macmillan
Publishers Ltd.
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68 UNDERSTANDING EARTH’S DEEP PAST
global silicate weathering rates, the global burial rates of marine CaCO 3
and siliceous plankton, atmospheric CO2 levels, and ice sheet growth and
ablation paced by changes in Earth’s orbit (Coxall et al., 2005; Zachos
and Kump, 2005; Pälike et al., 2006a).
The CO2 threshold behavior exhibited by the Eocene-Oligocene onset
of Antarctic glaciation and the Neogene initiation of the Greenland ice sheet
suggests that multiple equilibrium states exist in the climate system. To the
extent that the ice sheet climate system exhibits hysteresis, the CO2 threshold
identified for the cooling path may be substantially lower than that for the
reverse warming path. Simulations of the modern climate system (DeConto
and Pollard, 2003) and empirical proxy records (Pearson et al., 2009) have
suggested a substantial delay (up to several millennia) in ice sheet response
to increased atmospheric CO2 due to hysteresis. Such studies indicate that
polar ice sheet decay may require CO2 levels well above those that existed
during the initiation of Cenozoic glaciation (Pollard and DeConto, 2005).
Recent evidence, however, indicates the potential for subdecadal response
times of ice sheets and much more rapid melting (Das et al., 2008; van de Wal
et al., 2008). The response of the Neogene polar ice sheets to the atmospheric
CO2 levels during the Middle Miocene climatic optimum (~500 ppmv at
16 Ma; Küerschner et al., 2008) and the early Pliocene (up to 415 ppmv
at 4.5 Ma; Pagani et al., 2010)—values not too different from modern (2010)
concentrations—warrants further exploration to resolve the uncertainties
in ice sheet response times to global warming. If CO2 forcing is sustained
at levels through the point of no return, then rapid meltdown of glaciers
can be anticipated in the future even if carbon emissions to the atmosphere
ultimately decrease (Hansen et al., 2008).
If hysteresis is characteristic of ice sheet melting dynamics, then such a
delay in ice sheet response to elevated CO2 guarantees a future transition
into a warm world that is abrupt, extreme, and with possibly irreversible
catastrophic effects (Hansen et al., 2008; Kump, 2009). Presumably, the
long-term processes that drove the climate system into the glacial state
during the Cenozoic (enhanced silicate weathering and mountain build -
ing, reduced subduction of carbonates and volcanism, and thus low atmo-
spheric CO2 levels) will persist through the anthropogenic perturbation,
so it is reasonable to anticipate that the climate—following the current
transient warming—will cool over the subsequent few tens of millennia.
Eventually, conditions for the reinitiation of the Antarctic and Greenland
ice sheets will be achieved, but these may require atmospheric pCO2 levels
similar to preindustrial values and a favorable orbital state (Berger et al.,
2003; Pollard and DeConto, 2005). Thus, the trip “forward to the past” may
be quite prolonged, perhaps approaching the evolutionary timescales of
species, including Homo sapiens.
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69
CLIMATE TRANSITIONS, TIPPING POINTS, AND THE POINT OF NO RETURN
Paleocene-Eocene Thermal Maximum (PETM)
One of the best-known examples of an ancient global warming event,
with potential parallels to the near future, is the Paleocene-Eocene Thermal
Maximum (PETM). This abrupt climate change occurred at ~56 Ma with
repeated, rapid (millennial-scale), massive releases of “fossil” carbon
and major disruption of the carbon cycle (Kennett and Stott, 1991, 1995;
Dickens et al., 1995; Zachos et al., 2003). The oxygen isotopic compositions
of planktonic and benthic foraminifera record rapid warming of ~5°C in
tropical surface and deep oceans, and as much as 9°C warming at the poles,
that persisted for ~170 ky (Sluijs et al., 2006; Zachos et al., 2006; Röhl et
al., 2007) (Figure 3.3). Greenhouse gas-forced global warming was accom -
panied by extreme changes in hydroclimate and accelerated weathering
(Bowen et al., 2004; Pagani et al., 2006; Schmitz and Pujalte, 2007), deep-
ocean acidification (Zachos et al., 2005), and possible widespread oceanic
hypoxia (Thomas, 2007; Zachos et al., 2008). Whereas regional climates
in the mid- to high latitudes became wetter and were characterized by
increased extreme precipitation events, other regions, such as the western
interior of North America, became more arid (Schmitz and Pujalte, 2007).
With this intense climate change came ecological disruption, including
the immigration of modern mammalian orders (including primates) into
North America, large-scale floral and faunal ecosystem migration (e.g., see
Box 2.8), and widespread extinctions of benthic foraminifera in the deep
ocean (Thomas and Shackleton, 1996; Bains et al., 1999; Wing et al., 2003,
2005). Carbon isotope records indicate that although the onset occurred
within a few millennia, the recovery was much slower, taking well over
100 ky (Figure 3.3).
Dissociation (melting) of methane hydrates as their stability field
crossed a threshold, triggered by a warming trend in the early Eocene
(Dickens et al., 1995), is the most widely cited source of fossil carbon for
the PETM. However, methane’s isotopically light carbon requires that
less carbon (~2000 petagram [Pg]) be added to account for the observed
isotopic excursion than required by some models to account for the degree
of inferred seafloor carbonate dissolution (Zachos et al., 2005; Panchuk et
al., 2008). Some other suggested hypotheses to account for the abundant
fossil carbon include sustained burning of accumulated Paleocene ter-
restrial organic peats and coals (Kurtz et al., 2003; Huber, 2008), although
conclusive evidence in support of this hypothesis is lacking (Moore and
Kurtz, 2008); increased terrestrial methane cycling (Pancost et al., 2007),
although this may not generate a whole-system isotopic shift; desiccation
and oxidation of organic matter in large epicontinental seaways (Higgins
and Schrag, 2006), although the paleogeographic changes and their timing
remain poorly resolved; and more speculatively, the impact of a volatile-
rich comet (Cramer and Kent, 2005), although others have argued that the
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70 UNDERSTANDING EARTH’S DEEP PAST
A 3.0
2.0
1.0
δ13C ( ‰)
0
–1.0 Southern Ocean
690
Central Paci c
–2.0 865
B –1.0 14
South Atlantic
525
527
Temperature (°C)
–0.5 12
δ18O ( ‰)
0 10
0.5 8
1.0
C 100
80
60
CaCO 3 (%)
40
South Atlantic
(water depth)
20 1262 (4.8 km)
1263 (2.6 km)
0
54.0 54.5 55.0 55.5 56.0
Age (millions of years ago)
FIGURE 3.3 Marine stable isotope and seafloor sediment CaCO3 records compiled
using several ocean drilling sites for the PETM, a hyperthermal with some parallels
to modern greenhouse gas-driven global change. (A) δ13C time series developed
from benthic foraminifera illustrating ~2.5 part per thousand (‰) excursion at
~55 Ma. (B) δ18O time series and inferred temperatures record the prolonged
period of ocean warming (~70-80 ky) and its large magnitude. There may have
been several events of greenhouse gas release during the PETM that produced
the large, abrupt changes in ocean temperatures. (C) Record of seafloor calcium
carbonate content from the South Atlantic documents the significant reduction
due to dissolution and deep-ocean acidification during the PETM. The apparent
onset of CaCO3 dissolution prior to the onset of the carbon isotope excursion
reflects the extensive dissolution of uppermost Paleocene sediments by acidic
waters during the PETM.
SOURCE: Zachos et al. (2008), reprinted by permission of Macmillan Publishers
Ltd.
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71
CLIMATE TRANSITIONS, TIPPING POINTS, AND THE POINT OF NO RETURN
putative cometary particles were actually produced by bacteria (Kopp et
al., 2007; Lippert and Zachos, 2007; Schumann et al., 2008). Clearly, there is
no single fully satisfactory source to account for the carbon, and multiple
carbon releases may have occurred in response to an initial warming. A
likely trigger for the initial warming during the PETM is igneous intru -
sion into organic-rich sediments of the North Atlantic, which generated
thermogenic methane and CO2 (Svensen et al., 2004; Storey et al., 2007).
Notably, the sudden release of carbon into the atmosphere-ocean system
occurred at rates that vastly exceeded typical rates in Earth history, activat-
ing components of the climate system that can be triggered by accelerated
warming. The PETM serves as an important base level showing the effect
on the biosphere of a rapid rate of addition of fossil carbon to the atmo-
sphere (~3,000 to 4,500 Pg—on the order of that anticipated if we burn
through all fossil fuels)—yet dwarfed by the present rate of ~1 percent per
year CO2 increase in Earth’s atmosphere (Zeebe et al., 2009).
Transient warming episodes, such as the PETM, were a recurring
phenomenon of the early Eocene warm world. Many of the short-lived
hyperthermals were associated with abrupt and extreme climate change,
an accelerated hydrological cycle, and ocean acidification (Nicolo et al.,
2007; Stap et al., 2009) (Box 3.3). Short-term positive feedbacks active dur-
ing the hyperthermals magnified the climatic effects of the initial carbon
influx. Climate amelioration with each transient warming event would
have been substantially delayed as the rates of short-term feedbacks far
outpaced the negative feedbacks (e.g., weathering) capable of restoring the
global carbon cycle to a steady state (Zachos et al., 2008; Zeebe et al., 2009).
The PETM, and other hyperthermals of the early Cenozoic, occurred
when the Earth was virtually ice-free. This is certainly significantly dif -
ferent from modern and near-future conditions, which are expected to
maintain unipolar glaciation at a minimum. The ice sheets of the Neo -
proterozoic Snowball Earth and the Late Paleozoic Ice Age were far more
extensive than those of the Cenozoic icehouse, recording repeated major
glacial-interglacial transitions and including terminal epic deglaciation.
Despite substantially different land mass-height distributions, ocean circu-
lation patterns, and marine and terrestrial ecosystems from those of today,
the geological record of these deglaciations—specifically the repeated
major transitions between glaciations and glacial minima including their
terminal epic deglaciations—provide the only “icehouse” perspective of
the response of the climate system and ecosystems to perturbation beyond
the range archived in the more recent glacial records.
The Late Paleozoic Deglaciation
Much of the scientific understanding of feedbacks and thresholds in
the current glacial climate system, and their influence on the biosphere,
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72 UNDERSTANDING EARTH’S DEEP PAST
BOX 3.3
Deep-Time Insights into Ocean Acidification
The early Cenozoic hyperthermals, and in particular the PETM, provide
a natural laboratory to study climate sensitivity to pCO2, the interplay of
short- and long-term feedbacks in the climate system, and ocean acidifica-
tion under magnitudes of atmospheric pCO2 increases that are comparable
to present and projected future increases. Clear evidence of deep-ocean
acidification exists for the PETM (Zachos et al., 2005; Zeebe and Zachos,
2007), with corrosive waters completely dissolving calcium carbonate on
the Atlantic seafloor at water depths below 2.5 km; today, this calcium car-
bonate compensation depth (CCD) occurs below 4 km in most ocean basins.
Whether surface waters became undersaturated is less clear since car-
bonate producers such as planktonic foraminifera and coccolithophorids
persisted during the event. However, whereas shallow water reefs composed
of corals, calcareous red and green algae, and larger benthic foraminifera
were abundant prior to the PETM, metazoan reefs nearly vanished between
56 million and 55 million years ago (Scheibner and Speijer, 2008; Kiessling
and Simpson, 2010). Widespread coral reefs did not reappear until the
middle Eocene, at ~49 Ma. It appears that a combination of persistent
warming from the late Paleocene to early Eocene, punctuated by deep-
ocean acidification at the PETM, defined a threshold for coral-algal reefs
that led to rapid loss and only gradual recovery. Notably, the lack of
evidence for surface water acidification probably indicates that the rate
of carbon addition was slower—perhaps by an order of magnitude—than
projected fossil fuel emission rates under the least optimistic scenarios for
the future (e.g., the A1 family of scenarios considered by IPCC [2007])
which, in box models, generates surface and deep-water acidification
(Zeebe et al., 2008, 2009).
Over the past two centuries, the ocean has absorbed 40 percent of
anthropogenic CO2 emissions (Zeebe et al., 2008). If fossil fuel emissions
continue unabated and minimal development is put into carbon seques-
tration technologies, by the time humans burn through estimated fossil
fuel reserves (at ~A.D. 2300 to A.D. 2400), ~5,000 gigatonnes of carbon
will have been released to the atmosphere (Zachos et al., 2008). Because
the rate of anthropogenic carbon input to the atmosphere greatly exceeds
the mixing time of the oceans (1,000-1,500 years), CO2 will build up in
the atmosphere (perhaps to ~2,000 ppmv) and the surface ocean (Kump,
2002; Zachos et al., 2008). What could be in store for this millennium? As
the ocean continues to absorb CO2, carbonate ion (CO32–) concentration
will fall leading to decreases in surface water pH and saturation states, a
condition that is already apparent and will continue over the next century
(Figure 3.4). Acidic surface waters are expected to massively affect ocean
ecosystems, including the widespread loss of coral reefs. With time, acidic
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73
CLIMATE TRANSITIONS, TIPPING POINTS, AND THE POINT OF NO RETURN
25
PETM Scenario
Release (Pg C y )
-1
Business as Usual
20
← 5000 Pg C
15
10
0 Pg C
5
← 300
0
6
Ω surf , Calcite
4
2
PETM Scenario
Business as Usual
0
0 2000 4000 6000 8000 10000
Year
FIGURE 3.4 Initial carbon pulse for the PETM (red curves), estimated to
be 3,000 Pg carbon using published carbon isotope and observed deep-
sea carbonate dissolution records, and a carbon cycle model (LOSCAR;
Zeebe et al., 2008, 2009). The magnitude of the input carbon mass was
inferred from carbonate dissolution records, with the δ13C of the carbon
pulse (≤ –50‰) constrained by requiring the model outcome to match
observed deep-sea δ13C records. The model assumes a large initial input
of carbon over 5 ky, followed by further smaller pulses and a low continu-
ous carbon release (an additional 1,500 Pg) throughout the PETM main
event. Changes in calcite saturation in the surface ocean (lower diagram)
are estimated for the PETM (red curve) and for the future (black curve),
based on the inferred magnitude of the carbon pulse to atmosphere.
SOURCE: Courtesy of R.E. Zeebe, personal communication (2010).
continued
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74 UNDERSTANDING EARTH’S DEEP PAST
BOX 3.3 Continued
water will penetrate to the deep ocean where it will dissolve carbonate
sediments and begin to be neutralized. In response, the saturation horizon
of the deep ocean will shoal on a decadally observable timescale. Calci-
fication rates of corals will slow noticeably and may become negligible in
the next 100-150 years. At first, the rise in the saturation horizon will be
slow, but as the area of seafloor above the saturation horizon declines (fol-
lowing the seafloor hypsometric curve), the shoaling rate will accelerate,
bringing it to as shallow as the depth of the shelf-slope break (~130 m) in
the next several centuries. At this point, barrier reefs, having long since lost
their reef-building biota, will erode through dissolution and disintegrate.
This history of carbonate dissolution will result in a carbonate-poor layer
in the deep ocean, much like sediments associated with past hyperthermals
such as the early Cenozoic PETM.
has been elucidated by studies of climate transitions of the past few mil-
lion years—in particular the moderate-scale glacial-interglacial fluctua -
tions of the Pleistocene. The demise of the Late Paleozoic Ice Age (LPIA;
between 290 and 260 Ma) provides an opportunity to evaluate climate
stability and climate-biota interactions during a major climate transition
coupled to changing CO2 contents. For example, climate models indicate
that climate-driven biome changes at high latitudes may have factored
strongly in controlling LPIA glacial-interglacial changes (e.g., Horton et
al., 2010). For the final stages of this protracted ice age, covariance between
shifts in pCO2 and continental and marine surface temperatures inferred
from isotopic proxies of soil-formed minerals and marine fossil brachio -
pods, and ice sheet extent reconstructed from southern Gondwanan
glacigenic deposits, indicates a strong linkage of pCO2-climate-ice-mass
dynamics that is consistent with greenhouse gas forcing (Montañez et
al., 2007). A pattern of progressively more extensive and long-lived ice
sheets through the Late Carboniferous (340 to 310 Ma; Fielding et al.,
2008) was reversed in the Early Permian—under rising atmospheric CO2
levels—as climate ameliorated and conditions shifted toward a protracted
greenhouse climate state (Montañez et al., 2007). The trend of gradually
increasing surface temperatures and increasing atmospheric pCO2 is
punctuated by larger but shorter-term fluctuations associated with each
discrete glaciation. Surface temperatures and CO2 levels never returned to
the earliest Permian minima associated with the apex of Gondwanan con -
tinental ice sheets. Intermittent warmings—characterized by CO 2 levels
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CLIMATE TRANSITIONS, TIPPING POINTS, AND THE POINT OF NO RETURN
above the simulated threshold for glaciation (Horton et al., 2007; Horton
and Poulsen, 2009)—heralded the more permanent change to an ice-free
world to come. This pattern of episodic changes in atmospheric CO 2 and
surface temperatures in step with transient glaciations, superimposed
on a longer-term warming trend during the demise of the Late Paleozoic
Ice Age, shares characteristics in common with the Eocene-Oligocene
greenhouse-to-icehouse transition (Coxall et al., 2005; Liu et al., 2009a,b).
The similarity in behavior of these very different transitions suggests that
turnovers in climate states are most probably characterized by large-scale
episodic change. Future study of the deep-time record of the Earth’s last
epic deglaciation should shed light on how the cryosphere, hydrosphere,
chemosphere, and biosphere responded to such episodic change under
rising CO2 levels.
Deglaciation During the Neoproterozoic
A phase of rapid global warming is recorded in the late Neoprotero -
zoic (~635 Ma), abruptly terminating what was probably the longest-lived
(~135 Ma; Macdonald et al., 2010) and coldest icehouse period of Earth
history, where at times ice sheets extended to sea level in equatorial
latitudes—a climate state popularly referred to as the “Snowball Earth”
(Hoffman et al., 1998). Carbon isotope trends provide evidence for sub -
stantial, but poorly understood, disruption of the carbon cycle during the
ice age itself, including the possibility that the high albedo of a global-scale
ice sheet dominated climate (Hoffman et al., 1998). The terminal deglacia-
tion in the Neoproterozoic offers an intriguing deep-time archive of how
major changes in long-term processes that regulate climate, such as silicate
weathering and carbon burial and productivity, have been triggered when
a threshold in the climate system has been reached through CO2 forcing.
Abrupt and rapid increase in CO2 at the end of the Neoproterozoic gla-
ciations is recorded by the presence of thin calcium carbonate deposits,
interpreted to have been deposited on a millennial timescale, immediately
overlying Neoproterozoic glacial sediments across the globe (Kennedy et
al., 1998; Hoffman and Schrag, 2002). These carbonate deposits are a physi-
cal record of rapid release of CO2 to the atmosphere calculated at a rate
of ~1 percent CO2 increase per year (Kennedy et al., 2001), similar to the
current rate of CO2 increase of 0.8 to 1 percent per year (IPCC, 2007). While
the analogy is imperfect because of the very different biosphere and conti-
nental configuration in the Neoproterozoic, deglaciation under this strong
greenhouse gas forcing imparted a record unique from that of subsequent
deglaciations. Most notably, the abrupt transition to greenhouse condi-
tions associated with this complete deglaciation appears to have involved
a dominating rapid-warming feedback (Fairchild and Kennedy, 2007) that
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76 UNDERSTANDING EARTH’S DEEP PAST
involved the massive release of CO2 (Hoffman and Schrag, 2002) and/or
the destabilization of methane clathrates and the release of methane gas
from permafrost and marine reservoirs (Kennedy et al., 2001, 2009). It
appears that the dominance by warming feedbacks during the termina-
tion of this ice age precluded the damping effects of other feedbacks that
governed climate oscillations during Phanerozoic ice ages. Consequently,
the late Neoproterozoic deglaciation provides an excellent example of the
long-term feedbacks that can be triggered—and likely accelerated—when
a threshold in a strongly forced climate system is reached.
HOW LONG WILL THE GREENHOUSE LAST?
The potential for climatic consequences with severe impact on humans
resulting from the buildup of fossil fuel CO2 has inevitably resulted in
questions not only of “How bad will it get and how fast?” but also “How
long will it last?” The answers to these questions depend heavily on the
global warming potential of the greenhouse gas release, which accounts
not only for its immediate impact on the planetary radiative energy bal-
ance but also on the longevity of the greenhouse gas in the atmosphere
(Archer et al., 2009).
Economic forecasts suggest that conventional fossil fuel resources
will largely be used up in the next 200-400 years, leading to atmospheric
CO2 levels that could reach ~2,000 ppmv by A.D. 2300 to 2400 (Marland
et al., 2002; Caldeira and Wickett, 2003). However, models of the global
carbon cycle and the geologic record both show that CO2 produced from
fossil fuels and other reservoirs will continue to impact global climate
and atmospheric chemistry for tens to hundreds of thousands of years.
Although CO2 produced by fossil fuel burning is taken out of the atmo-
sphere within decades of its production, the oceans, soils, and vegetation
continue to exchange greenhouse gases back into the atmosphere for far
longer. Greenhouse gases continue to affect climate and ocean acidity
until they are buried as organic matter or converted to mineral forms of
inorganic carbon through rock weathering (Box 3.4).
Simple box models (Figure 3.6) have been used to make long-term
projections of future climate to capture the “recovery” from the fossil
fuel-induced greenhouse state (Walker and Kasting, 1992; Archer, 2005).
Although box model calculations should not be considered definitive,
they do suggest that the fossil fuel perturbation may interfere with the
natural glacial-interglacial oscillation driven by predictable changes in
Earth’s orbit (Berger et al., 2003), perhaps forestalling the onset of the
next northern hemisphere “ice age” by tens of thousands of years. A more
convincing exposition of the central question of “how long” requires more
comprehensive models. Scientific confidence in those models will be high
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CLIMATE TRANSITIONS, TIPPING POINTS, AND THE POINT OF NO RETURN
BOX 3.4
CO2 Sweepers and Sinks in the Earth System
The carbon fluxes in and out of the surface and sedimentary reservoirs
over geological timescales are finely balanced, providing a planetary ther-
mostat that regulates Earth’s surface temperature. Initially, newly released
CO2 (e.g., from the combustion of hydrocarbons) interacts and equilibrates
with Earth’s surface reservoirs of carbon on human timescales (decades to
centuries). However, natural “sinks” for anthropogenic CO2 exist only on
much longer timescales, and it is therefore possible to perturb climate for
tens to hundreds of thousands of years (Figure 3.5). Transient (annual to
century-scale) uptake by the terrestrial biosphere (including soils) is easily
saturated within decades of the CO2 increase, and therefore this compo-
nent can switch from a sink to a source of atmospheric CO2 (Friedlingstein
et al., 2006). Most (60 to 80 percent) CO2 is ultimately absorbed by the
surface ocean, because of its efficiency as a sweeper of atmospheric CO2,
and is neutralized by reactions with calcium carbonate in the deep sea at
timescales of oceanic mixing (1,000 to 1,500 years). The ocean’s ability
to sequester CO2 decreases as it is acidified and the oceanic carbon buffer
is depleted. The remaining CO2 in the atmosphere is sufficient to impact
climate for thousands of years longer while awaiting sweeping by the
“ultimate” CO2 sink of the rock weathering cycle at timescales of tens to
hundreds of thousands of years (Zeebe and Caldeira, 2008; Archer et al.,
2009). Lessons from past hyperthermals suggest that the removal of green-
house gases by weathering may be intensified in a warmer world but will
still take more than 100,000 years to return to background values for an
event the size of the Paleocene-Eocene Thermal Maximum.
In the context of the timescales of interaction with these carbon sinks,
the mean lifetime of fossil fuel CO2 in the atmosphere is calculated to be
12,000 to 14,000 years (Archer et al., 1997, 2009), which is in marked
contrast to the two to three orders of magnitude shorter lifetimes commonly
cited by other studies (e.g., IPCC, 1995, 2001). In addition, the equilibra-
tion timescale for a pulse of CO2 emission to the atmosphere, such as the
current release by fossil fuel burning, scales up with the magnitude of the
CO2 release. “The result has been an erroneous conclusion, throughout
much of the popular treatment of the issue of climate change, that global
warming will be a century-timescale phenomenon” (Archer et al., 2009,
p. 121).
continued
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78
Uptake by the oceans
Projected pCO2 (based on
“business as usual“ and all
Dissolution of sea-floor sediments
fossil fuel combusted)
1200 1600 2000
Weathering of carbonate rocks
800
Silicate Projected natural
weathering CO2 variability
Ice Core pCO2
400
Atmospheric CO2 (ppm)
100,000 BP 10,000 BP Today 10,000 100,000 1,000,000
Past (years) Future (years)
BOX 3.4 Continued
FIGURE 3.5 Graphic portrayal of the CO2 “lifetime” assuming nonlinear CO2 uptake kinetics by
various short (decades to millennia) and long-term (104 to 105 y) surface and sedimentary carbon
reservoirs. Projected natural CO2 variability assumes 100 ky orbital control.
SOURCE: Modified from Walker and Kasting (1992); B.B. Sageman, personal communication.
UNDERSTANDING EARTH’S DEEP PAST
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79
CLIMATE TRANSITIONS, TIPPING POINTS, AND THE POINT OF NO RETURN
1800
5000 Gton C
2000 Gton C
1400
1000 Gton C
350 Gton C
pCO2
1000
600
200
0 20 40 60 80 100
Time, kyr
FIGURE 3.6 Schematic showing the predicted long-term response over 100,000
years of atmospheric CO2, including ocean temperature feedbacks, to a range of
possible fossil fuel emissions totals. The 100,000-year simulations include silicate
weathering (solid lines) and the 35,000-year simulations include seafloor CaCO 3
dissolution (dashed lines). These models highlight how the timescale of carbon
uptake becomes extended as the event unfolds. Fast processes such as ocean
uptake and biomass growth, with high transfer rates but limited capacity, lose
their potency, while slower processes, such as seafloor carbonate dissolution and
rock weathering, come to dominate.
SOURCE: Modified from Archer (2005).
only if they can be evaluated against observation. The historical record,
and even the expanse of the Quaternary climate record, contains nothing
comparable.
Observations and modeling of the past carbon cycle perturbations
provide a basis for projecting future conditions under the full range of
fossil fuel burning scenarios, including the most pessimistic “business-
as-usual” eventuality. Along this trajectory, atmospheric CO2 levels will
rise as long as fossil fuel burning continues (with ultimate input of ~5,000
GtC), rising to levels perhaps as high as 1,600-2,000 parts per million (i.e.,
five to seven times the preindustrial level) (Figure 3.6). The geological
record of past hyperthermal events, including the PETM, suggests that
severe global warming under such magnitudes of carbon emissions will
persist for 20,000 to 40,000 years. Carbon cycle models indicate that even
after 100,000 years, the anthropogenic perturbation to the carbon cycle
will still be important, especially if the total amount of carbon emitted is
large. Consequently, Milankovitch forcings that have so dominated the
pacing and extent of climate variations, and especially ice sheets, over the
last 2 million years will—as they were prior to the onset of the current
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80 UNDERSTANDING EARTH’S DEEP PAST
glacial state in the Oligocene—serve as only a minor modulator of high-
latitude climate variability because climate change will be muted under
such elevated atmospheric pCO2 levels. The Greenland ice cap could dis-
appear in the first few coming millennia, and if CO2 levels rise more than
two to four times present levels, the West Antarctic ice cap could collapse
(Naish et al., 2009), although this conclusion is highly sensitive to orbital
configuration and model parameterizations (Pollard and DeConto, 2005).
By any measure, exploitation of much of the fossil fuel reservoir over only
300 years will clearly leave a far longer lasting legacy.
