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Phil. Trans. R. Soc. A (2011) 369, 933–956
doi:10.1098/rsta.2010.0317
Are there pre-Quaternary geological analogues
for a future greenhouse warming?
BY ALAN M. HAYWOOD1, *, ANDY RIDGWELL2 , DANIEL J. LUNT2 ,
DANIEL J. HILL3 , MATTHEW J. POUND1 , HARRY J. DOWSETT4 ,
AISLING M. DOLAN1 , JANE E. FRANCIS1 AND MARK WILLIAMS3,5
1 School
of Earth and Environment, Earth and Environment Building,
University of Leeds, Woodhouse Lane, Leeds LS2 9JT, UK
2 School of Geographical Sciences, University of Bristol, University Road,
Bristol BS8 1SS, UK
3 British Geological Survey, Keyworth, Nottingham NG12 5GG, UK
4 Eastern Geology and Paleoclimate Science Center, USGS 926A National
Center Reston, VA 20192, USA
5 Deaprtment of Geology, University of Leicester, University Road,
Leicester LE1 7RH, UK
Given the inherent uncertainties in predicting how climate and environments will respond
to anthropogenic emissions of greenhouse gases, it would be beneficial to society if science
could identify geological analogues to the human race’s current grand climate experiment.
This has been a focus of the geological and palaeoclimate communities over the last 30
years, with many scientific papers claiming that intervals in Earth history can be used
as an analogue for future climate change. Using a coupled ocean–atmosphere modelling
approach, we test this assertion for the most probable pre-Quaternary candidates of the
last 100 million years: the Mid- and Late Cretaceous, the Palaeocene–Eocene Thermal
Maximum (PETM), the Early Eocene, as well as warm intervals within the Miocene and
Pliocene epochs. These intervals fail as true direct analogues since they either represent
equilibrium climate states to a long-term CO2 forcing—whereas anthropogenic emissions
of greenhouse gases provide a progressive (transient) forcing on climate—or the sensitivity
of the climate system itself to CO2 was different. While no close geological analogue exists,
past warm intervals in Earth history provide a unique opportunity to investigate processes
that operated during warm (high CO2 ) climate states. Palaeoclimate and environmental
reconstruction/modelling are facilitating the assessment and calculation of the response
of global temperatures to increasing CO2 concentrations in the longer term (multiple
centuries); this is now referred to as the Earth System Sensitivity, which is critical in
identifying CO2 thresholds in the atmosphere that must not be crossed to avoid dangerous
levels of climate change in the long term. Palaeoclimatology also provides a unique and
independent way to evaluate the qualities of climate and Earth system models used to
predict future climate.
Keywords: palaeoclimate; proxies; climate models; climate sensitivity; Earth System
Sensitivity; analogue
*Author for correspondence ([email protected]).
One contribution of 13 to a Theme Issue ‘The Anthropocene: a new epoch of geological time?’.
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This journal is © 2011 The Royal Society
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A. M. Haywood et al.
1. Introduction
(a) Extending uniformitarianism
Uniformitarianism, the observation that, fundamentally, the same geological
processes that operate today also operated in the distant past is a key principle of
modern geology. The methodological significance of the principle is summarized
in the statement the present is the key to the past. Uniformitarianism originated
with the work of the geologist James Hutton, but was refined by John Playfair
and popularized by Charles Lyell’s Principles of geology [1]. Can, though,
the principle be reversed and extended so that the past becomes the key
to the future?
The search for geological analogues for future climate change is not new,
but what is new is a realization that, given current and projected levels of
greenhouse gases in the atmosphere, we must travel far into the geological
past to find intervals of time in which greenhouse gases, as well as global
temperatures, were comparable to what climate models predict will occur by the
end of the twenty-first century [2]. Trace gas records from ice cores indicate that
atmospheric concentrations of CO2 are already higher than at any time during
the last 800 000 years [3–5]. Evidence from new alkenone-based, boron isotopebased and stomatal density-based CO2 proxy data indicate that the current
(2010) concentration of CO2 (390 ppmv recorded in 2009; data from the National
Oceanic and Atmospheric Administration’s Earth System Research Laboratory,
Mauna Loa Observatory) in the atmosphere may not have been reached in
the last 3 million years. A doubling of pre-industrial concentrations of CO2 to
560 ppmv, as expected later this century, has not been reconstructed for the last
20 million years [6–11].
In 1991, Thomas J. Crowley published what the authors consider to be
a landmark paper in the Journal of Climate [12]: he examined, for the first
time, suggestions made by Kellogg [13], Hansen & Lebedeff [14], Budyko &
Sedunov [15] and Zubakov & Borzenkova [16] that warm intervals in Earth
history could be used as a frame of reference or even possible analogues for future
atmospheric CO2 -induced warming. Existing climate-model predictions of the
climatological pattern and response times predicted a doubling of atmospheric
CO2 were compared with early, and in many ways limited, palaeoclimate
modelling case studies available at the time from the Pleistocene (Early Holocene
and Last Interglacial) and pre-Pleistocene (Early Pliocene, Early Eocene and
Mid-Cretaceous). For various reasons, many of which we will directly or indirectly
re-examine in this paper, it was concluded that there may be no satisfactory
geological analogue and therefore that future discussions on geological analogues
should be ‘restricted to the study of processes operating in the climate system, and
that continued use of the term for past warm periods be abandoned’. Yet, there
are many examples of scientific papers published since 1991 that continue to refer
to time periods in Earth history as being geological analogues, or at least partial
analogues for the current anthropogenic climate perturbation (e.g. [17,18]). It
is interesting to consider why this is the case. Is it because the science has
progressed since Crowley [12], and that clear examples of past warm intervals
as analogues for future greenhouse-gas-induced warming are now available? Or
is it because Crowley’s [12] analysis was limited by the modelling tools and case
studies available at the time?
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In this paper, we revisit the pre-Quaternary case studies examined by Crowley
to explore the possibility of geological analogues for future greenhouse warming
(from the oldest to the youngest). We do this by using outputs from fully
coupled ocean–atmosphere climate models that are able to overcome some
of the limitations of the early climate models used by Crowley [12], and
through reference to the published literature. We consider a new potential
analogue that has arisen since 1991, the Palaeocene–Eocene Thermal Maximum
(PETM), and conclude with a discussion on the relevance and usefulness of
an examination of Earth history for future climate change. We answer the
question are there any satisfactory pre-Quaternary geological analogues for a
future greenhouse warming?
To prevent our investigation of analogues becoming unmanageable, we have
established criteria that restrict the number of case studies we must discuss.
For the pre-Quaternary, we have only selected intervals when atmospheric CO2
concentrations and global mean temperatures were higher than pre-industrial
values, as we consider this a pre-requisite for a time interval considered as
a potential analogue. We have also restricted our discussions to intervals of
time when unaltered marine sediments are present. This ensures that we have
primary information available from both the marine and terrestrial realms to
support the published palaeoclimate reconstructions and to constrain our climate
simulations (i.e. back to 100 million years ago (Ma)). We have also attempted
to examine palaeoclimates that may have been in equilibrium with ambient CO2
concentrations as well as aberrant events when greenhouse-gas forcing and climate
change was transient. Combined, these criteria dictate a focus on greenhouse
states of the Cretaceous and Eocene, the PETM as well as warm intervals within
the Miocene and Pliocene epochs.
2. Predictions for future greenhouse-gas-induced warming
(a) Predictions of policy-relevant climate change
Before we begin our exploration of geological case studies described above, it is
first necessary to briefly review the latest model predictions for ‘climate policy
relevant’ future climate change (i.e. end of this century) so that reconstructions
and model predictions of past climates can be placed in the appropriate context.
The Intergovernmental Panel on Climate Change (IPCC) summary [19] of
post twenty-first century climate change will also be considered. Climate-model
predictions forced with different CO2 emissions scenarios (Special Report on
Emissions Scenarios (SRES) B1, A1T, B2, A1B, A2 and A1FI), presented in
Working Group I of the IPCC [19], project global average surface warming
and sea-level rise at the end of the twenty-first century (temperature change
in ◦ C at 2090–2099 relative to 1980–1999) to range from 1.1–6.4◦ C and 0.18–
0.59 m, respectively [20]. The precise nature of regional climate responses to CO2
emissions depends on the selection of a CO2 emission scenario, but certain climate
trends are seen to be independent of the choice of scenario. Warming is expected
to be greatest over land and at most high northern latitudes, and least over the
Southern Ocean (near Antarctica) and the northern North Atlantic. Sea ice is
projected to shrink in both the Arctic and Antarctic under all SRES scenarios.
In some projections, Arctic late-summer sea ice disappears almost entirely by the
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A. M. Haywood et al.
latter part of the twenty-first century. Based on a range of models, it is probable
that future tropical cyclones (typhoons and hurricanes) will become more intense,
with larger peak wind speeds and more heavy precipitation associated with
ongoing increases of tropical sea-surface temperatures (SSTs). Extra-tropical
storm tracks are projected to move poleward, with consequent changes in wind,
precipitation and temperature patterns. Increases in precipitation are very likely
at high latitudes, while decreases are likely over most subtropical land regions [20].
(b) Climate predictions beyond the twenty-first century
Climate change may continue to be induced by emissions of CO2 long after
these emissions have been stabilized, owing to the inertia of the climate system.
This makes examination of climate trends after 2100, along with any similarities
of consistencies between such trends during warm intervals in Earth history,
necessary and advantageous [19–21]. If radiative forcing were to be stabilized,
keeping all the radiative forcing agents constant at B1 or A1B levels in 2100,
model experiments show that a further increase in global average temperature
of about 0.5◦ C would still be expected by 2200. In addition, thermal expansion
alone would lead to 0.3–0.8 m of sea-level rise by 2300 (relative to 1980–1999).
Thermal expansion would continue for many centuries, owing to the time required
to transport heat into the deep ocean. Both past and future anthropogenic CO2
emissions will continue to contribute to warming and sea-level rise for more than
a millennium, owing to the time scales required for the removal of this gas from
the atmosphere [20].
(c) HADCM3 and predictions of equilibrium global climate with instantaneous
CO2 forcing
Model predictions of the climate response to a doubling of atmospheric CO2
have been performed using instantaneous forcing and transient forcing-type
experiments. In the former, atmospheric CO2 partial pressure is instantaneously
increased from a pre-industrial value, with the simulation then run to equilibrium.
In the latter, forcing is gradually introduced by increasing CO2 concentrations
in each year of the simulation, which is a more realistic approach. However,
palaeoclimate experiments are normally conducted using the instantaneous
forcing approach. Therefore, to maintain consistency in our experimental design,
we present results of a doubled CO2 experiment using instantaneous forcing.
The future and palaeoclimate simulations presented in this paper were all
performed using the Hadley Centre for Climate Prediction and Research fully
coupled ocean–atmosphere climate model v. 3 (HADCM3 and HADCM3L). The
particulars of this model are well documented (e.g. [22,23]). HADCM3 and
HADCM3L were one of the first coupled atmosphere–ocean climate models that
required no flux corrections to be made, even for simulations of 1000 years or
more [24]. The general circulation model consists of a linked atmospheric model,
ocean model and sea-ice model. In HADCM3 and HADCM3L, the horizontal
resolution of the atmosphere model is 2.5◦ in latitude by 3.75◦ in longitude.
This gives a grid spacing at the equator of 278 km in the north–south direction
and 417 km east–west, and is approximately comparable to a T42 spectral
model resolution. The atmospheric model consists of 19 layers. The spatial
resolution over the ocean in HADCM3 is 1.25◦ × 1.25◦ and the model has 20
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layers, while in HADCM3L, the horizontal resolution of the ocean is the same as
the atmosphere. The atmospheric model has a time step of 30 min and includes
a radiation scheme that can represent the effects of minor trace gases [25]. A
parametrization of simple background aerosol climatology is also included [26].
The convection scheme is that of Gregory et al. [27]. A land-surface scheme
includes the representation of the freezing and melting of soil moisture. The
representation of evaporation includes the dependence of stomatal resistance on
temperature, vapour pressure and CO2 concentration [28].
The ocean model includes the use of the Gent–McWilliams mixing scheme [29].
There is no explicit horizontal tracer diffusion in the model. The horizontal
resolution allows the use of a smaller coefficient of horizontal momentum viscosity
leading to an improved simulation of ocean velocities. The sea-ice model is a
simple thermodynamic scheme and contains parametrizations of ice drift and
leads (Polynyas; [30]).
3. Geological case studies
(a) The Cretaceous and Early Eocene
We begin our analysis of potential pre-Quaternary analogues by examining some
of the most extreme and long-term warm events in the last 100 million years,
specifically the so-called ‘greenhouse intervals’ of the Cretaceous and Early
Eocene (100–50 Ma). Reconstructing atmospheric CO2 this far back in time is
extremely challenging and fraught with uncertainties. However, an overview of
all available Cretaceous and Early Eocene CO2 proxy data, along with massbalance calculations, indicates a picture of higher ambient CO2 concentrations,
albeit with significant variability during this interval. This variability is probably
present within a single geological stage, such as the Maastrichtian. For example,
available estimates for palaeo CO2 during the Maastrichtian from the d13 C
fractionation of pedogenic carbonates [31–33] indicate that CO2 may have ranged
from pre-industrial concentrations (or even lower) to approximately 1500 ppmv.
This dynamic range encompasses any SRES emissions scenario for CO2 and
therefore the Cretaceous is often referred to as a potential analogue for future
climate change (e.g. [17]).
We compare the response of precipitation and evaporation (and precipitation
minus evaporation) predicted by HADCM3L for the Late Cretaceous
(Maastrichtian stage) and present-day, both with higher than pre-industrial
atmospheric CO2 concentrations (figure 1a–d). In a control Maastrichtian,
simulation with a CO2 concentration twice that of the pre-industrial level
(2 × PAL), the average global annual mean precipitation rate equals 3.32 mm d−1 .
In our future climate-change experiment, with 2 × PAL CO2 , the average global
mean precipitation rate equals 2.95 mm d−1 . Initially, this appears to support the
utility of the Late Cretaceous as an analogue for the future, since in both cases,
the model predicts an enhanced global precipitation rate compared with the preindustrial simulation (2.83 mm d−1 global mean average total precipitation rate).
However, total precipitation rates in the Late Cretaceous are greater than the
future climate-change experiment. A Late Cretaceous simulation using 1 × PAL
CO2 reveals that even in the absence of higher than pre-industrial levels of CO2 ,
the global annual mean precipitation rate is still greater than the pre-industrial
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A. M. Haywood et al.
(a)
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Figure 1. HADCM3 predictions for a future climate-change experiment and the Maastrichtian
experiment: (a) annual mean evaporation rate (mm d−1 ) for 2 × PAL CO2 experiment; (b) annual
mean evaporation rate (mm d−1 ) for 2 × PAL CO2 Maastrichtian experiment; (c) difference in
annual mean precipitation minus evaporation rate (P – E; mm d−1 ) between a 2 × PAL CO2
experiment and a pre-industrial experiment; (d) difference in annual mean precipitation minus
evaporation rate (P – E; mm d−1 ) between a 2 × PAL CO2 Maastrichtian experiment and a preindustrial CO2 Maastrichtian experiment; (e,f ) prescribed bathymetry (m) in HADCM3 for the
modern and Maastrichtian simulations. Note the increased area of shallow marine seas (epeiric
seas) in the Maastrichtian case. (Online version in colour.)
or even future climate-change experiment by approximately 10 per cent and
5 per cent, respectively. More significantly, the magnitude of the response in
total precipitation rate to a doubling of CO2 in the future climate-change and
Cretaceous scenarios also differs (4.2% and 7.1%, respectively).
This altered intensity of precipitation-rate response between the future and
Cretaceous experiments, with or without an increase in CO2 to 2 × PAL, can
be ascribed to a number of factors. Firstly, in the Cretaceous scenario, there is
no ice at either pole, hence the high-latitude, subpolar and temperate regions
are warmer than the present-day and thus evaporation and precipitation are
enhanced. Continental position and the area of epeiric seas, which are currently
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Analogues of future warming
939
at an all-time Phanerozoic low [34], also appear to play a role. Owing to the
continental configuration during the Maastrichtian stage, the area of epeiric
seas is greatly expanded compared to today (figure 1e,f ). Given that tropical
temperatures are high, this has the effect of changing the amount of water
evaporated from the oceans in the tropics, introducing a large additional flux
of water vapour in the Late Cretaceous scenario that cannot be matched in
the pre-industrial or future climate-change experiments (figure 1a–d). Equivalent
simulations of the Early Eocene using HADCM3L demonstrate a similar pattern
of precipitation-rate enhancement, even in the absence of higher CO2 [35].
Greenhouse worlds of the Cretaceous and Eocene represent equilibrium climate
states to long-term CO2 forcing (rather than a transient CO2 forcing), and the
sensitivity of the climate system to CO2 during these intervals was apparently
different, which is very significant for their utility as analogues.
(b) The Palaeocene–Eocene Thermal Maximum and carbon cycling
A substantial transient warming of the Earth’s surface occurred approximately
55.5 Ma (the PETM); a range of temperature (and hydrological cycling) proxies
were recorded [36] and synchronous with a pronounced negative carbon isotopic
(d13 C) excursion [37]. This isotopic excursion has been observed globally: in
marine cores from the Arctic through Equatorial Pacific to the Southern
Ocean [38,39], in shelf sediments [40], as well as in those from terrestrial
deposits [39,41], and occurs in both biogenic carbonates and organic matter
[37]. To explain the event’s global nature, as well as large (ca 4‰) magnitude
compared with general Cenozoic variability [42], requires a release of isotopically
light carbon to the ocean and/or atmosphere. This, in turn, strongly supports a
primary driving role for increased atmospheric greenhouse-gas concentrations in
the observed warming and hence highlights the PETM as a potential analogue
for future climate change. Of particular future relevance is whether methane
(CH4 ) hydrate destabilization was involved, particularly, in terms of what
triggered it and how strong the feedbacks between global warming and CH4
release are. The PETM also represents a past case study into the impacts of
acidification of the surface ocean on calcifying organisms such as foraminifera and
coccolithophores [43]. However, to draw out any analogue with the future, we need
to constrain the source, amount and fate of the carbon release, as these bear on
the nature of carbon-cycle feedbacks operating, the degree of ocean acidification
and climate sensitivity.
The recorded magnitude of the excursion alone can, in theory, be used to infer
the carbon release. The relationship between carbon release (DMnew ) and observed
excursion (Dd13 C) can be approximated by
DMnew =
Minitial
,
(d13 Cnew /Dd13 C) − 1
(3.1)
where Minitial is the initial inventory of carbon in the readily exchangeable
surface reservoirs (atmosphere, ocean and terrestrial biosphere) and d13 Cnew is
the isotopic composition of the carbon source. However, two main problems are
encountered in this. Firstly, the full magnitude of the carbon isotopic excursion
is rarely (if at all) recorded in marine carbonates owing to concurrent dissolution
(a consequence of ocean acidification and reduced carbonate preservation [44]),
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A. M. Haywood et al.
152 000 PgC @ –5
mantle CO2
20 000
18 000
carbon input required (PgC)
16 000
14 000
12 000
organic matter
10 000
8000
thermogenic methane
6000
4000
2000
0
biogenic methane
–10
–20
–30
–40
–50
–60
–70
d13C of source ( )
Figure 2. The carbon mass input required to produce a −4‰ carbon isotope excursion within
an exogenic (atmosphere + ocean + terrestrial biosphere) carbon reservoir of 38 000 Pg of carbon
(PgC), as a function of the carbon isotopic composition of the source (equation (3.1)). This
approximation assumes a mean isotopic composition of the exogenic reservoir prior to carbon input
of 0‰. The estimated isotopic values of major sources follow Maslin & Thomas [52]. Adapted from
Dunkley Jones et al. [37].
while records contained within terrestrial plants may be affected by changes in
atmospheric CO2 concentrations, local temperature and precipitation (e.g. [39]).
The second problem is the multiplicity of potential sources of carbon for the
PETM event, as different carbon sources have different isotopic signatures.
Sources hypothesized to date include: comet impact [45], volcanic carbon release
associated with the North Atlantic Igneous Province [46,47], large-scale peat
fires [48], weathering of organic-rich desiccated epeiric sea sediments [49] and CH4
release, either thermogenic CH4 associated with North Atlantic volcanism [50], or
biogenic CH4 from destabilized methane hydrate within ocean sediments [51]. The
sensitivity of the total carbon release to the source (equation (3.1)) is illustrated
in figure 2.
A second independent constraint can be placed on the total carbon release by
interpreting the variation across the event in the relative abundance of calcium
carbonate (CaCO3 ) mineral particles in deep-sea sediments. The basis for this
is the tendency for CaCO3 to dissolve and not be preserved in accumulating
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sediments as CO2 is added to the ocean and sea water becomes increasingly
acidified [53]. Estimates for the total carbon release that best explains observed
changes in the CaCO3 content of marine sediments are somewhat model
dependent (depending on e.g. how much CaCO3 is assumed to be present in
sediments globally immediately prior to the event) and range from 3000 [54]
to 6800 Pg of carbon (PgC) [55]. This would imply either a dominant biogenic
CH4 source (3000 PgC at −50‰) or a dominant organic matter (or thermogenic
CH4 ) contribution (6800 PgC at −22‰). This calculation is also influenced by
the assumed time scale of carbon release [55]. As the duration of carbon release
approaches and exceeds the carbonate buffering time scale of the ocean (ca
2–8 kyr [53]), more total carbon is required to create the same degree of CaCO3
dissolution in deep-sea sediments. The much longer (ca 100 kyr) carbon isotopic
response time of the system dictates that slower, larger releases must have a less
isotopically depleted source. Other effects such as changes in ocean circulation [56]
and mixing of the sediments by bioturbation [44] can also affect the calculation.
The time scale of the carbon release at the PETM is particularly important in
the degree to which the PETM can provide analogue insights into anthropogenic
change. For instance, it is notable that marine organisms have experienced
substantial secular variability in their environment over geological time scales,
with surface ocean, pH conditions likely to have been some approximately
0.6–0.7 pH units lower during the Cretaceous and Jurassic compared with
modern [43,57], a pH change comparable to or exceeding projected future ocean
acidification [58]. Yet, calcareous plankton originated, diversified and proliferated
during this time. However, this is a false comparison as not only can individuals
and ecosystems adapt and evolve to changing climate and ocean geochemistry on
‘long’ (million year) time scales, but also the carbon saturation state of the ocean,
which determines the thermodynamic ‘ease’ by which marine organisms can
produce carbonate shells and skeletons, is generally well regulated on geological
time scales [57]. In contrast, only events involving geologically abrupt (less than
10 kyr) releases of CO2 can (temporarily) overwhelm the buffering afforded by
marine sediments to produce a coupled decline in both pH and saturation state
(as is occurring now).
Current estimates of the time scale of carbon release at the PETM are of a
carbon input that was essentially complete within 10 kyr [36,59], with much of this
potentially released within just a few kiloyears [54]. Placing further constraints
on the rate of release and warming at the onset of the PETM is particularly
difficult because changes in climate (and weathering) and carbonate dissolution
associated with the onset of the PETM strongly affect sedimentation rates and
hence available age models. However, a main release taking place as rapid as over
1 kyr (or less) cannot be ruled out geologically, although most of the hypothesized
carbon sources would be unable to deliver sufficient carbon to the ocean or
atmosphere this rapidly.
In conclusion, while the PETM represents a greenhouse-gas-driven global
warming and ocean acidification transient event, with a total carbon release
comparable to available fossil-fuel reserves, the rate of climatic and ocean
geochemical changes is likely to have been an order of magnitude slower. Hence,
it would be prudent to take the observed biotic impacts and ecosystem disruption
at the PETM as a lower estimate of potential future impacts rather than a
direct analogue.
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(c) Miocene warm intervals
There are a number of intervals within the Miocene epoch that are generally
considered to have been warmer than the present-day (i.e. the Mid-Miocene
Climatic Optimum and the Late Miocene; [60–63]). Continental position and
topography were changing during the epoch, developing towards a modern
configuration/condition by the Late Miocene. These changes/developments
included the presence of a Central American Seaway (CAS), a marine
encroachment from the south into Argentina, a large extension of Eurasia into the
Arctic Sea to approximately 80◦ N, the large Pannonian Lake in central Europe
and a wider Indonesian Seaway [64].
The Miocene represented a crucial epoch of uplift and the generation of arid
regions [65,66]. The uplift of the Himalayas from a relatively low Tibetan Plateau
(1–3 km) in the Late Oligocene to an average height of 4–5 km in the Late Miocene
(approx. 9 Ma) had effects on global atmospheric circulation, weathering rates
and the Asian Monsoon [65,67–70]. The Andes may have been at half their
modern height at 10.7 Ma (approx. 1800 m) and have since been uplifting at
0.2–0.3 mm yr−1 [71]. The Rocky Mountains of western North America are a
product of several orogenic events, the most recent of which was the Laramide
Orogeny that is dated to the Late Cretaceous to Palaeocene [72]. Subsequent
to this major event, the Colorado Plateau has been uplifted by nearly 2 km
since the Cretaceous [73]. Estimates of the exact timing of the uplift and the
rate are still unresolved, but recent work focusing on the Colorado Plateau
suggests that a change in the dynamic topography of 400–1100 m has occurred
in the last 30 million years [74,75]. The Alps in the Early to Mid-Miocene were
merely islands between the Paratethys and western Tethys Seas, being at an
estimated height of less than 1800 m, then major uplift occurred after 14 Ma until
present [76].
CO2 levels for the Miocene have been estimated using boron isotopes
[77], alkenones [8], stomatal indices [6,78], pedogenic carbonate [33] and
GEOCARB mass-balance modelling [79]. These techniques estimate CO2 to
range between Last Glacial Maximum values and approximately 2 × PAL
[6,8,77–79], although pedogenic carbonates used to estimate CO2 go as high as
1170 ppmv at 10 Ma [33]. It is important to note that this dynamic range is not
supported by all CO2 proxies. Most notably, alkenone-based CO2 reconstructions
for the Miocene reach approximately pre-industrial concentrations by the
beginning of the epoch and maintain this concentration throughout, even
through notable warm intervals such as the Mid-Miocene Climatic Optimum,
which was arguably the warmest phase in Earth history over the last 15
million years. This has led some to suggest that Miocene climate change
was not related to atmospheric CO2 variations [8,80,81], which, if true,
would automatically disqualify the epoch for consideration as an analogue.
However, this is far from clear, since newer CO2 estimates, such as those
derived from stomatal indices, are consistent with the coevolution of Miocene
climate and CO2 [6]. The apparent discrepancy/inconsistencies between CO2
proxies for the Miocene require resolution before warm intervals in the
Miocene can be fully explored in terms of analogue issues (also see caveats
associated with Pliocene interglacials in §3d that are relevant to the Miocene
as well).
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(d) Pliocene warm intervals
Most recent climate-model predictions [21,82] suggest that during warm
‘interglacials’ of the Pliocene epoch, global annual mean temperatures were 2–3◦ C
higher than during the pre-industrial interval. During these warm interglacials,
sea levels were higher than today (estimated to be 10–30+ m), meaning that
global ice volume was reduced [83]. There were large fluctuations in ice cover
on Greenland and West Antarctica, and during the interglacials, they may
have been largely free of ice [84–86]. Some ice may also have been lost from
around the margins of East Antarctica, especially in the Aurora and Wilkes subglacial basins [87]. Coniferous forests replaced tundra in the high latitudes of the
Northern Hemisphere [88], and the Arctic Ocean may have been seasonally free of
sea ice [89–92]. The most recent and detailed estimates of the CO2 concentrations
in the atmosphere range between 280 and 450 ppmv [7,9]. The Mid-Pliocene warm
period (mPWP; 3.26–3.025 Ma BP; time scale of Lisiecki & Raymo [93]) is a
particularly well-documented interval of warmth during the Pliocene, with global
datasets of multi-proxy SSTs, bottom water temperatures, vegetation cover,
topography and ice volume readily available as boundary conditions for global
climate models [94,95].
Warm intervals of the Pliocene, and particularly the mPWP, are attractive
targets to examine analogue issues, and many parallels have been drawn between
the predicted Pliocene and twenty-first century climate anomalies from presentday, particularly in terms of (i) the change in annual mean global temperature
[96,97], (ii) changing meridional surface-temperature profiles showing a strong
polar amplification of the warming [88,90,98,99], (iii) changing precipitation
patterns and storm tracks [100], and even (iv) hurricane intensity and El Niño
Southern Oscillation (ENSO)-event frequency/extra-tropical teleconnections
[18,101]. This attraction is made more intense by the fact that the continents
had essentially reached their modern position, and the greater confidence in the
modern analogue techniques used in palaeoclimate reconstructions, because the
biotic assemblages and relationships are much more similar to modern than in
previous warm intervals [88,102].
However, caveats remain in using warm periods of the Pliocene as a guide
to future climate and environmental change. Although the continents had
reached their modern position, there may still have been significant differences
in topography that would have a bearing on the local and regional, perhaps
even global, sensitivity of the climate system to higher than pre-industrial
concentrations of CO2 in the atmosphere. For example, there is uncertainty
surrounding the precise nature and timing of the evolution of the western
Cordillera of North and South America (e.g. [103,104]). The uplift of the Andes
and Rockies through the Pliocene had significant effects on local and regional
climate and environments in affecting atmospheric circulation (figure 3a).
A number of critical ocean gateways/sills and throughflows also evolved during
the Pliocene (Bering Strait, the Greenland–Scotland Ridge/Indonesian through
flow and the CAS). Arguably the most significant of these is the CAS, which
shoaled and finally closed during the Pliocene, cutting off the tropical connection
between the Atlantic/Caribbean Sea and the Eastern Equatorial Pacific. The
precise timing of this closure, and whether or not it was a simple history of
closure or more a complex history with numerous breaching events, remains an
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Figure 3. HADCM3 sensitivity tests for the mPWP: (a) difference in annual mean surface air
temperature (SAT; ◦ C) between two experiments, one with the Western Cordillera of North and
South America at modern height, the other reduced by 50%; (b) difference in annual mean surface
air temperature between an open and closed Central American Seaway experiment; (c) difference
in annual mean surface air temperature between a 560 and 280 ppmv CO2 simulation; (d) difference
in annual mean surface air temperature between a future climate-change experiment (560 ppmv
CO2 ) and an experiment for the pre-industrial era (280 ppmv CO2 ); and (e) difference between (c)
and (d). (Online version in colour.)
area of ongoing research (e.g. [105]). What is more certain, from an ocean and
coupled ocean–atmosphere modelling perspective, is the significant impact that
an open CAS would have had on the meridional overturning circulation and
global climate during the Pliocene. In HADCM3, the response of opening the
CAS is bipolar with a strong cooling in the Northern Hemisphere matched by an
equal warming of the Southern Hemisphere ([106]; figure 3b). This is linked to
a dramatic change (weakening) in the meridional overturning circulation in the
Atlantic with an open CAS. Such a response has been noted in a number of other
models of various complexities (e.g. [107]). If the CAS was open during much of
the Pliocene, the regional effects of higher than pre-industrial concentrations of
CO2 could then be expressed differently to the future climate scenario when the
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meridional overturning circulation is more vigorous. This is a further example
of how the sensitivity of the climate system to CO2 forcing in the past may
have changed.
An obvious solution would be to examine only the interglacials of the later
Pliocene when we can be confident that the CAS had closed. Sensitivity
experiments with HADCM3 indicate a very similar anomaly in mean annual
temperature to a doubling of CO2 in both Pliocene and pre-industrial scenarios
(with both pre-industrial and Pliocene simulations initialized with pre-industrial
and 2 × PAL CO2 ; figure 3c,d). However, there are subtle regional variations in
the pattern of mean annual warming (figure 3e). For example, the amount of
SST warming in response to a doubling of CO2 in the Nordic Seas is subdued
in the Pliocene case. The reason is intuitive: even though both Pliocene and
pre-industrial 560 ppmv simulations are compared with control simulations using
pre-industrial levels of CO2 , the Pliocene 280 ppmv scenario is still warmer than
the pre-industrial equivalent in this region because of the altered high-latitude
vegetation and ice-sheet coverage. This reduces the amount of sea ice in the
Pliocene 280 ppmv simulation (compared with the pre-industrial experiment),
meaning that when the additional CO2 forcing is applied, the response in the
Arctic cannot be strong, since there is less sea ice in the Pliocene case compared
with the pre-industrial scenario in the first place; the sea-ice albedo feedback
mechanism cannot work as efficiently in the Pliocene scenario.
In the subpolar North Atlantic, the opposite is found, with the Pliocene
showing a higher sensitivity to CO2 increases (figure 3e). This results from the
changes in the Rocky Mountains that alter atmospheric and ocean circulation
patterns [97]. Increases in the subpolar zonal winds increase export of Labrador
Sea surface waters and air masses into the North Atlantic. These seem to be
particularly sensitive, as this one basin contains both the regional sea-ice winter
and summer limits, which both retreat in response to CO2 increases. This high
sensitivity is largely confined to the immediate Labrador Sea area in the modern.
However, with greater export in both the atmosphere and ocean, the Pliocene
North Atlantic also exhibits this increased sensitivity.
A more subtle example is available from North America. Here, the difference in
CO2 sensitivity between the Pliocene and pre-industrial adopts a dipole pattern,
with the Pliocene showing greater warming in the north and less warming in
the south of the continent (figure 3e). This type of dipole pattern is classically
observed over North America during an El Niño event. While this is unlikely to
be the sole cause of the dipole, regression analyses confirm that teleconnections
with ENSO over North America are expressed more clearly in a Pliocene world
than a pre-industrial world owing to boundary condition changes in the Pliocene,
specifically the lower Western Cordillera of North America ([101]; S. G. Bonham
2010, private communication). Therefore, it is logical to expect this trend to
continue and perhaps to intensify with a doubling of CO2 .
The differences in climate sensitivity (Charney Sensitivity) described above
are all based on analyses of differences in the global mean response. For warming
events during a pre-Quaternary interval to be considered a good analogue for the
future, climate sensitivity through the seasons must also be the same. Analysis
of HADCM3-predicted seasonal differences between the Pliocene and modern
sensitivity to 2 × PAL of CO2 indicates regional differences of more than ± 10◦ C
(figure 4).
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Figure 4. Difference in the seasonal sensitivity of surface air temperature (◦ C) predicted by
HADCM3 between Pliocene simulations running with 280 and 560 ppmv CO2 and pre-industrial
simulations running with 280 and 560 ppmv CO2 . (a) December, January, February; (b) June,
July, August; (c) March, April, May; (d) September, October, November. These plots demonstrate
the difference in the Charney-based climate sensitivity between the Pliocene and modern climate
systems. (Online version in colour.)
Warm interglacials of the Pliocene have many advantages over other time
intervals discussed as an analogue for future climate change. Even so, Pliocene
interglacials presumably reflected long-term equilibrium to a given ambient CO2
level, whereas current greenhouse-gas emissions provide a rapid transient forcing
on climate. Warm events during the Pliocene are a more robust choice than earlier
warm intervals of the pre-Quaternary, yet important unresolved questions over
the nature of topography and critical ocean gateways must be resolved through
further research. HADCM3 experiments indicate that while the global climate
sensitivity of the Pliocene and modern worlds is almost identical, the regional
and seasonal expression of the climate sensitivity may have been different.
4. A focus on understanding Earth System Sensitivity and on
climate-model evaluation
(a) Understanding and calculating Earth System Sensitivity
The concept of climate sensitivity has been discussed extensively over the last 30
years, in particular, since the 1979 National Research Council report [108]. It is
defined as the increase in global temperature owing to a doubling of CO2 after ‘fast
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feedbacks’, typically acting on time scales of years to decades, in the atmosphere
and upper ocean have had time to equilibrate [109]. These fast feedbacks
correspond to the physics available in climate models ca 1980 (e.g. [110]), and
include, for example, water vapour, snow albedo, sea-ice albedo and clouds.
This sensitivity (Charney Sensitivity) remains a very useful benchmark for
comparing different climate models in idealized circumstances, and has been one
of the central concepts used by the IPCC in their assessments and scenarios
for future climate change. The IPCC summarizes studies that have aimed to
characterize the uncertainty in Charney Sensitivity (e.g. [111–113]) by stating
that ‘[Charney Sensitivity] is likely to lie in the range 2◦ C to 4.5◦ C, with a
most likely value of about 3◦ C’. More recent studies have agreed with this
assessment (e.g. [114]). These estimates of climate sensitivity have also been used
to determine the probable impacts, both environmental and social/economic, of
various CO2 stabilization scenarios (e.g. [115]), or the level of greenhouse-gas
emissions consistent with stabilization of the global mean temperature below a
certain value (e.g. [116]).
However, owing to insufficient understanding of key mechanisms, a political
focus on the next few decades and a lack of the necessary computational resources,
these estimates of climate sensitivity have all ignored important feedbacks in
the Earth system, which act on longer time scales when the palaeoclimate case
studies discussed here become more relevant (typically multiple centuries). For
example, it is well known that vegetation changes can provide an important
positive feedback on climate change, both in future projections (e.g. [117]) and
in past climates (e.g. [118]). Similarly, ice sheets are a key player in long-term
climate change, as in the continental-scale ice coverage of North America and
northern Eurasia during the Last Glacial Maximum, 21 000 years ago [119].
Other processes such as dust–climate interactions, atmospheric chemistry and
non-CO2 greenhouse gases are also potentially important, but neglected in most
current estimates of climate sensitivity owing to a lack of understanding of key
mechanisms.
The fact that these additional feedbacks will modify the long-term climate
sensitivity is now beginning to be appreciated in the climate-science community
(e.g. [120]), and Lunt et al. [21] proposed a new term—that of ‘Earth System
Sensitivity’ (ESS) as a long-term and more complete version of Charney
Sensitivity. Lunt et al. [21] demonstrated that a combined palaeoclimate
modelling and data approach can be used to investigate the concept and
significance of ESS. Provided that the CO2 forcing and palaeoenvironmental
boundary conditions (e.g. vegetation and ice sheets) are sufficiently known, they
can be prescribed in a climate model, allowing relatively short integration times to
reach equilibrium. The model can therefore be used to provide a global estimate
of long-term equilibrium surface temperature for the given CO2 forcing. This
temperature estimate can be assessed by additional proxy data (e.g. SST) for
those locations where it is available. The initial study of Lunt et al. [21], focusing
on the mPWP of the Pliocene epoch, showed that the approach above gave an
estimate of ESS 30–50% greater than the traditional Charney Sensitivity.
However, estimates of ESS are far from complete, and no study has attempted
to explore the uncertainty or sensitivity of the ESS. Since stabilization targets
are often defined in terms of absolute temperature thresholds, it is critical that
the difference between ESS and Charney Sensitivity be properly quantified and
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explored; this can only be achieved through reference to warm intervals in Earth
history. Continental position and topography remain significant uncertainties
in using the past to calculate ESS, since the effects of such changes must
be quantified and removed from any calculation of ESS. Therefore, it is
sensible to focus our efforts to calculate ESS on the warm intervals in Earth
history when such changes were minimal, such as interglacial events of the
latter part of the Pliocene epoch. Given the enormous amount of careful
research and available proxy climate and environmental data, the mPWP is an
ideal choice.
(b) Climate and Earth system model evaluation
Along with the assessment of ESS, the use of climate archives to confront
climate and Earth system models highlights the value of palaeoclimatology.
Climate-model predictions for the future cannot be confronted with observations.
How well the models simulate modern climate is routinely evaluated (e.g. [121]),
but it is possible to make a case that this does not provide a truly independent test
of the models since observational datasets are used in the tuning of the model’s
performance, and is implicit in the development of parametrization schemes
included in all models. Therefore, an evaluation purely on the basis of observed
climate could potentially yield an overly optimistic assessment of the predictive
abilities of the models. Interestingly, there is frequently palaeoclimatic data in
regions of the world where modern observations to test climate models are rare
or absent (i.e. the Arctic and Antarctic), and this had led to some of the most
surprising and persistent discrepancies between model predictions and geological
reconstructions, for example, continental interior temperatures and high-latitude
temperatures during the Cretaceous and Eocene greenhouses being too cold in
the models (e.g. [122]). A pattern is emerging where the models generally appear
to underestimate the magnitude of climate and environmental change compared
with the data, which if true is cause for concern.
Model/data comparisons are not a panacea for pre-Quaternary palaeoclimatology; in fact, it is extremely challenging and technical work and demands
the full quantification of uncertainty within the data as well as model estimates.
In the pre-Quaternary, this is possible to achieve, but it has rarely been
attempted. To be done well, such work requires a synergy, understanding and
trust between modeller and data collector that remains rare in the community.
To prove a model wrong, we must first prove that a model prediction lay
clearly outside the envelope of data uncertainty, and then we must know that
the error is due to the physics of the model itself, rather than a prescribed
boundary condition. While the climate signal (anomaly to present) is large in
pre-Quaternary climates, the further back in time, the larger the uncertainties
become. Furthermore, the data community should appreciate that even if
problems are identified in a single palaeoclimate simulation, this does not prove
that the underlying physics of the models are flawed. A model can only be as
good as the geological boundary conditions that drive it, and when data/model
discrepancies occur, it is quite normal for the problem to be traced to an initial
geological condition (e.g. CO2 , palaeogeography/topography/bathymetry). This
is the real test and challenge to the usefulness of data/model comparisons in the
pre-Quaternary; can the geological community provide sufficiently robust and
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constrained boundary conditions to the climate modellers so that the ‘blame’ for
data/model mismatches can unequivocally be placed at the foot of the models
and their physics?
Those who also model pre-Quaternary climates now recognize that more must
be done to quantify uncertainty. The Quaternary modelling community has
devoted time and effort to quantifying uncertainty in model predictions, but
this is only in its infancy for the pre-Quaternary (e.g. [94,123]). Examples of
multi-model ensembles (MMEs) and intercomparisons and perturbed physics
ensembles (PPEs) are available in Quaternary Science (e.g. [124,125]), but with
the exception of the Pliocene Model Intercomparison Project (PlioMIP; [94]) and
Quantifying Uncertainty in Climate Model Predictions for the Pliocene (PlioQUMP; [123]) are absent in the pre-Quaternary. Without MMEs and PPEs,
combined with ensembles exploring the uncertainty generated by geological
boundary conditions for the pre-Quaternary, it is not possible to assess the true
predictive abilities of models.
5. Conclusions
Great societal benefit would come from the identification of a true pre-Quaternary
geological analogue for future climate change. This would enable science to
understand how regional climates and environments will respond to climate
change during this century. It would provide a natural laboratory for the
evaluation of numerical models of climate and the Earth system that we rely
on to produce accurate predictions of future change. To be an analogue, a past
warm interval in Earth history must be a result of increased concentrations of
greenhouse gases in the atmosphere (compared with the pre-industrial era); the
climate sensitivity of the Earth must also be demonstrably the same as predictions
for the future, and the regional or global pattern of climate and environmental
change in the past must not be modified (amplified/dampened) owing to changes
in continental position or orography. The past warm interval must also be a
response to a transient greenhouse-gas forcing rather than being a result of a
long-term equilibrium condition to greenhouse-gas concentrations.
Through the application of coupled ocean–atmosphere models, the most
probable candidates (warm intervals) with the pre-Quaternary record of the last
100 million years were tested, and we conclude that there is no satisfactory
pre-Quaternary geological analogue to a future greenhouse warming. While
Earth history fails to provide a true and direct analogue, it does provide many
examples of how Earth responded to variations in atmospheric greenhouse-gas
concentrations in the longer term (multiple centuries), and therefore illuminates
the sensitivity of the Earth system, rather than simply the atmosphere and surface
oceans in isolation, to greenhouse gases. Given the long residence time of CO2
in the atmosphere, this is essential if we are to constrain reliable scenarios for
emission reductions that enable us to avoid dangerous levels of climate change
beyond 2100. Warm climates in Earth history also provide a tractable way of
evaluating the predictive abilities of the current generation of climate and Earth
system models. This is essential if we are to have confidence in their predictions
of change in the future and for science to be able to place additional uncertainty
limits on the predictions from such complex models.
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A. M. Haywood et al.
A.M.H. acknowledges the Leverhulme Trust for the award of a Philip Leverhulme Prize (2008).
A.R., D.J.L, A.M.H. and J.E.F. conducted this research as part of the NERC grants entitled
‘Dynamics of the Palaeocene–Eocene Thermal Maximum’, and NERC AFI-grant entitled ‘Terminal
Cretaceous climate change and biotic response in Antarctica’. H.J.D. acknowledges the US
Department of Interior and the US Geological Survey global change programme for funding. The
paper forms part of the US Geological Survey PRISM project (Pliocene Research Interpretations
and Synoptic Mapping). D.J.L. acknowledges Research Councils UK for the award of an RCUK
fellowship. A.M.D. acknowledges NERC for the provision of a doctoral training grant.
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