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4
Deciphering Past Climates—
Reconciling Models and Observations
Forecasts of climate change for the next century are based on general
circulation models (GCMs) that have been developed and tuned primar-
ily using twentieth century records, but also with some input based on
the understanding of the climate system from the more recent geologi-
cal past (e.g., the Last Glacial Maximum and the Holocene). In part, this
reflects the high levels of radiometric calibration and temporal resolution
(subannual to submillennial) offered by near-time paleoclimate archives,
which are capable of identifying the typically nonlinear components of
the climate system—characterized by rapid response times—that are
relevant to human society. A critical prerequisite for accurate forecasts of
future regional and global climate changes based on GCMs, however, is
the requirement that these models use parameters that are relevant to the
future we seek to better understand. In this context, the recent climate
archive captures only a small part of the known range of climate phe-
nomena, since it has been derived from a time dominated by ice dynamics
at both poles. Furthermore, the magnitude of forcing that the planet is
now experiencing exceeds any that has occurred during the past ~30 mil -
lion years. As GCMs are transformed into Earth System Models for the
Intergovernmental Panel on Climate Change (IPCC) Fourth Assessment
Report, they will encompass vastly more system physics, and deep-time
climate studies will offer modelers the only real-world scenarios for test -
ing the full climate response to the large increases in greenhouse gas levels
that are projected.
As the climate system departs from the conditions captured by these
well-studied near-time climate analogues, it is necessary for the scientific
81
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82 UNDERSTANDING EARTH’S DEEP PAST
community’s efforts to expand to capture the full range of variability and
climate-forcing feedbacks of the global climate system, in particular for the
past “extreme climate events” and warmer Earth intervals that may serve
as analogues for future climate. Full testing of climate models for these
time periods will require evaluation of feedback processes within models,
enhanced spatial resolution, and longer simulations to better characterize
climate model variability. All of these requirements, especially those for
resolution and variability, will require significant computational resources.
For deep-time climate systems, the representation of paleogeographic
boundary conditions can be a much greater source of uncertainty than it is
for simulations based on modern geography. Furthermore, discrepancies
between model outputs and paleoclimate observations may indicate the
existence of additional processes, feedbacks, and/or sensitivities that
are not present in the model or expressed in the modern climate sys-
tem. For example, the exceptionally warm high latitudes during all past
warm periods—whether transient or long term—cannot be reproduced by
models without invoking unreasonable CO2 levels, revealing the inability
of current models to fully capture the processes and feedbacks governing
heat transport and retention or the processes that might generate heat in
the polar regions under elevated atmospheric greenhouse gas levels (Covey
and Barron, 1988; Rind and Chandler, 1991; Covey, 1991; Sloan and Pollard,
1998; Bice et al., 2006; Huber, 2008; Kump and Pollard, 2008; Spicer et al.,
2008; Zachos et al., 2008). Thus, model development, which is based on
improving specific processes and climate feedbacks and, in turn, evalu-
ating the impact of these improvements on model simulations, depends
on the availability of spatially resolved, robust, deep-time paleoclimate
reconstructions of appropriate geochronological resolution and constraint.
In addition, the utility of paleoclimate proxies for climate reconstruction
and data-model comparisons relies on the proxies being sufficiently well
preserved and the existence of an adequate understanding of the under-
lying processes, sensitivities, and uncertainties captured by these proxies.
Recent paleoclimate studies of deep-time successions have docu-
mented the potential of the older part of the geological record to reveal
long-duration archives of forcings, responses, and long-term (centuries to
tens of millennia) feedbacks that are of magnitudes and/or durations not
captured by Pleistocene and Holocene paleoclimate records. Constraining
the nature (e.g., rates, phasing between proxies) and origin (forcings) of
climatic shifts, particularly rapid and/or transient events across climate
thresholds from the deep-time record, will be greatly enhanced where
orbital-scale cycles can be identified and resolved in the rock record
(Box 4.1, Figures 4.1 and 4.2). Indeed, millennial to seasonal signals—at
times calibrated to the orbital timescale—have been extracted from the
sedimentary record spanning hundreds of millions of years (e.g., Feldman
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DECIPHERING PAST CLIMATES—RECONCILING MODELS AND OBSERVATIONS
BOX 4.1
Determining Time in the Geological Record
One of the biggest challenges in using the deep-time record for under-
standing Earth systems is determining the rates of processes and dating
when specific events occurred. Determining rates requires very precise time
control, particularly if the processes being studied occur at an ecological
timescale (1 year to several centuries). One method for such precise time
control is by using annually layered sediments in ancient anoxic or hyper-
saline basins, as long as a few age control points are present. Examples
include laminated sediments from the Pleistocene Santa Barbara Basin
(Figure 4.1a), Eocene sediments of the North Sea (Figure 4.1b) (Schiøler
et al., 2007), black shale sequences in the Cretaceous North Atlantic
(Figure 4.1c), and the Permian Castile Formation of West Texas (Anderson,
1982). Quite highly resolved relative timescales can also be achieved
using cyclic sequences, with resolutions of a few centuries to several tens
of millennia, based on the identification of distinct orbital periods (Figure
4.1d; Figure 4.2). Both annual layers and orbitally tuned records can reveal
ecological dynamics, snapshots of natural variability at different parts of
Earth history, and the duration of threshold shifts in ecosystems.
Orbital cycles can sometimes be tied to astronomically tuned time-
scales during the past 40 million years to provide excellent absolute
timescales. One example is the tuning of orbital cycles in glacial events
during the Oligocene using combined geochemical and sediment prop-
erty cycles tied to an astronomically calibrated timescale (Pälike et al.,
2006a,b; see Figure 4.2). Orbital cycles have been recognized far back
in the Phanerozoic sedimentary record and, together with high-resolution
U-Pb dating, offer the potential to reconstruct Earth system dynamics in
great detail (Erwin 2006; Davydov et al., 2010).
et al., 1993; Olsen and Kent, 1996; Eriksson and Simpson, 2000; Loope et
al., 2001, 2004; Ivany et al., 2004; Wagner et al., 2004; Elrick and Hinnov,
2007; Jahren and Sternberg, 2008; Kennedy et al., 2009). The ability to
precisely and accurately quantify geological time has improved dra-
matically with recent advances in radiometric dating and interlaboratory
cross-calibration (e.g., the EARTHTIME initiative) permitting unprec-
edented temporal resolution (e.g., ID-TIMS [isotope dilution-thermal
ionization mass spectrometry] uranium-lead [U-Pb] ages with analytical
uncertainties of ≤0.01 percent; Ramezani et al., 2007). Some recent radio-
metric calibrations of the sedimentary record integrate astrochronology,
providing Milankovitch-scale resolution through long intervals of time
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84 UNDERSTANDING EARTH’S DEEP PAST
a. Annual sediment layers
Holocene Santa Barbara Basin, California
140
0 100
50
Depth (cm)
b. Laminated sediment layers
Early Eocene, North Sea, North Atlantic
140
0 100
50
Depth (cm)
c. Laminated sediment layers
Cretaceous, Demerara Rise, North Atlantic
140
0 100
50
Depth (cm)
d. Orbitally-paced color cycles
Latest Paleocene, Florida Continental Slope, North Atlantic
140
0 100
50
Depth (cm)
FIGURE 4.1 Sediment cores can be detailed recorders of time: (a) Core sample
(Core MV1012-TC-3) from the Holocene of Santa Barbara Basin, California, dis-
playing well-developed millimeter-scale laminations related to annual cycles in
biological productivity and sediment runoff. (b) Eocene laminated sediments from
the North Sea (Nini-3 well), potentially offering very high resolution records of
climate and ecosystem variability. (c) Similar laminations in a drill core through
Cretaceous black shales from the tropical North Atlantic (from ODP Site 1259).
(d) Orbital cycles in sediment color paced by the precession cycle (~21 kyr) from
a drill core in the Paleocene of the western North Atlantic (from ODP Site 1051;
Norris and Röhl, 1999).
SOURCE: ODP core images courtesy Integrated Ocean Drilling Program Science
Services.
(e.g., Kuiper et al., 2008; Davydov et al., 2010). Furthermore, integration
of orbital-stratigraphic approaches with bio-, magneto-, cyclo-, and/or
chemostratigraphy has successfully placed high-resolution temporal con -
straints on past events (e.g., Olsen and Kent, 1996; Sageman et al., 2006;
Westerhold et al., 2008; Adams et al., 2010; Galeotti et al., 2010). Several
epochs and stages of the Phanerozoic have been fully orbitally tuned,
presenting the possibility of geochronological resolution at 104- to 105-year
scales (e.g., Hinnov and Ogg, 2007). Many of these records, however, await
radiometric calibration to the absolute timescale.
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DECIPHERING PAST CLIMATES—RECONCILING MODELS AND OBSERVATIONS
FIGURE 4.2 Astronomically tuned climate record from the Oligocene of the Cen-
tral Pacific. Top two panels show the astronomical calculation of the 41,000-year
cycle in the tilt of Earth’s spin axis (known as obliquity), and the 405,000-year
cycle in eccentricity (thin black line). In the middle panel is the carbon isotope
record of the deep Pacific which displays a well developed ~400,000-year cycle
that closely matches the calculated astronomical cycle (dashed green line). ETP is
the calculated cycle expected if the sediment record integrated the combined ec -
centricity, tilt and precession cycles. The geomagnetic polarity timescale is shown
on the bottom, here calibrated to the astronomical cycles.
SOURCE: Modified from Pälike et al. (2006b).
The potential of the deep-time paleoclimate record to provide unique
insight into scientific understanding of the climate system’s response to
greenhouse gas forcing, however, is underdeveloped. To maximize this
potential, the community is presented with three primary challenges:
• To determine precise chronologies for existing and new geo -
logical archives of paleoclimate interest—where feasible at the temporal
resolutions that are possible in the Pleistocene and Holocene—through
continued improvements in the precision and accuracy of geochronologi-
cal techniques applicable to the sedimentary record (Ar-Ar [argon-argon]
and U-Pb), and the development of novel radiometric approaches such as
U-Pb dating of carbonates (Rasbury and Cole, 2009) and rhenium-osmium
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86 UNDERSTANDING EARTH’S DEEP PAST
(Re-Os) dating of black shales (Ravizza and Turekian, 1989; Selby and
Creaser, 2005; McArthur et al., 2008a).
• To develop within such geochronological and/or orbitally tuned
frameworks, marine and terrestrial—ultimately linked—time series of
high temporal resolution and spatial density and distribution. For terres -
trial records that are notoriously geographically fragmented and poorly
resolved, this will rely in large part on systematically acquired, targeted
continental drilling of continuous and highly resolved records.
• To obtain proxies of a broad range of surface and atmospheric
conditions (paleotemperatures, pCO2, pO2, contents of other greenhouse
gases, paleoprecipitation, seasonality, relative humidity)—of higher preci-
sion and accuracy than currently available—through some combination of
proxy refinement, proxy development, and multiproxy studies.
CLIMATE MODEL CAPABILITIES AND LIMITATIONS
Current paleoclimate model capabilities (Box 4.2) include the appli-
cation of fully coupled three-dimensional models to past climates. These
are the same models that are used to study the present climate state and
future changes to Earth’s climate, and they are the models that provide
the modeling foundation for the IPCC periodic assessments. For many
years, these models included only the physical aspects of the atmosphere,
dynamic ocean, land, and sea ice components of the climate system. More
recently, however, these models have begun to include coupling to a
dynamic ice sheet model and prognostic components for biogeochemistry,
atmospheric chemistry, dynamic vegetation, and ecology. Many models
can even provide calculations of the isotopic composition of precipita -
tion, making direct comparisons with δ18O marine and terrestrial proxies
possible (Roche et al., 2006; Poulsen et al., 2007a, 2010; Zhou et al., 2008).
Thus, global climate models have evolved from physical climate system
models to more comprehensive Earth system models that permit more
realistic coupling between the physical climate system and the biosphere
(e.g., Slingo et al., 2009; Cadule et al., 2010).
In terms of the mean state, climate system models are able to realisti -
cally capture many characteristics of the current climate, such as observed
equator-to-pole thermal gradients, large-scale spatial distribution of pre -
cipitation patterns, and various aspects of regional climate variability
(e.g., El Niño-Southern Oscillation, Pacific Decadal Oscillation). Although
the more comprehensive Earth system models offer many advantages,
many aspects of regional-scale climates are still not captured accurately
(although see Sewall and Sloan [2006], Thrasher and Sloan [2009], for
exceptions). Simulating regional-scale phenomena requires the existence
of high-resolution paleoclimate boundary data, which may not exist for
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DECIPHERING PAST CLIMATES—RECONCILING MODELS AND OBSERVATIONS
BOX 4.2
Climate Models
Climate models are numerical representations of the climate system that
provide a means to study the processes that determine Earth’s climate state.
Over the past 50 years, climate models have evolved to include a hierarchy
of approaches to representing the climate system (Figure 4.3).
Geochemical box models are used to study the temporal evolution of
quantities such as atmospheric CO2 and oxygen, ocean stable isotopes,
and other geochemical variables. These models are based on theoretical
expressions of the sources and sinks of a range of geochemical properties,
providing global mean information on timescales of tens of thousands to
millions of years.
Earth system models of intermediate complexity (EMICs) extend the box
model concept to include spatial resolution and are useful tools to study
Earth processes for timescales exceeding 10,000 years. Typically, these
models include a detailed two- or three-dimensional ocean model coupled
to simplified one- or two-dimensional atmospheric models. The energy
balance atmospheric model predicts the geographic distribution of surface
temperature and other energetic atmospheric quantities. The ocean model
includes a marine biogeochemical component that simulates the chemi-
cal state of the ocean. The horizontal spatial resolution of these models
is ~500 km. These models include detailed physical and biogeochemical
processes that are often missing in the more complex models. However,
their major limitation is that the atmospheric components are highly tuned
to the present-day world, and they cannot incorporate realistic mechani-
cal and thermodynamic forcing of the atmosphere on the ocean. Overall,
these models are of value to look at transient climate change, such as the
long-term fate of pCO2 over a few hundred thousand years.
Global climate models (GCMs) are the most comprehensive models for
studying the climate system (IPCC, Fourth Assessment Report, Chapter 8,
2007). These models are usually composed of three-dimensional represen-
tations of the atmosphere, ocean, sea ice, and land systems. These systems
are dynamically coupled and allow for feedbacks among the various com-
ponents. Recently, these fully coupled Earth System Models have begun
to include other processes such as atmospheric chemistry, terrestrial and
marine biogeochemistry, and ecological models. With recent increases in
computational power, atmospheric GCMs are now simulating the climate
system on spatial scales of 50 to 75 km (Kim et al., 2008). Lower resolution
versions of fully coupled GCMs (spatial resolutions of ~100 to 400 km)
can be run for thousands of years (Liu et al., 2010), and this is especially
important for deep-time climate research since the equilibrium time for
the oceans is ~3,000 years (e.g., Kiehl and Shields, 2005). With continued
continued
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88 UNDERSTANDING EARTH’S DEEP PAST
BOX 4.2 Continued
advances in computing as well as increases in the availability of massively
parallel supercomputers, it is likely that multimillennial simulations using
coarser-resolution GCMs (~400-km grid spacing) will soon be possible.
Such a hierarchy of climate models is essential for studying climate
change on a wide range of timescales from decades to millions of years.
Information from the more computationally expensive GCMs can be used
as input for the EMICs, which can then be run for hundreds of thousands
of years. Information from the GCMs or EMICs can also be used to better
represent climate feedbacks in geochemical box models.
Spatial Resolution (kilometers)
104
Box Models
EMICS
103
GCMs
102
100 101 102 103 104 105 106 107
Typical Length of Temporal Simulation (Years)
FIGURE 4.3 Graphical representation of the ranges of spatial resolution
and simulation run times for the major categories of climate model.
such spatial resolutions (e.g., ~50 km). Additional challenges exist for the
modeling of ancient climates, which are characterized by different paleo-
geography, paleotopography, atmospheric pCO2, solar luminosity levels,
and other boundary conditions. Yet, the community’s confidence in the
ability of GCMs to forecast future regional and global climate changes
may be unfounded if these models cannot simulate past climates that dif-
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DECIPHERING PAST CLIMATES—RECONCILING MODELS AND OBSERVATIONS
fer as much from the present as the future is likely to differ from current
conditions.
At this time, climate models of past periods use certain parameter-
izations defined for the present-day global climate system, necessar-
ily requiring questionable assumptions about the relevance of present-
day conditions to the older parts of the geological record. For example,
assumptions concerning plant physiology, biome composition, and sur-
face distribution based on present-day vegetation are not likely to have
been applicable to times prior to the evolution of angiosperms (~130
million years ago [Ma]) and the expansion of grasses (~34 Ma). Similarly,
the emission of precursor gases that form atmospheric aerosols is linked
to current understanding of atmospheric chemistry because proxies for
paleoemissions of gases that could create aerosols, in particular for those
that can affect cloud properties such as biologically mediated dimethyl
sulfide (Henriksson et al., 2000; Kump and Pollard, 2008), do not exist at
this time. Thus, the definition of boundary conditions inferred from the
geological proxy record and the parameterization of physical processes
in warmer world models intrinsically come with significant uncertainty.
Key boundary conditions that must be prescribed for deep-time or
warm world models include the paleogeography of land areas, past veg-
etation distributions, paleotopography, and ice sheet extent. Coupling to
ocean models additionally requires knowledge of paleoeustasy in order
to specify the distribution of shallow seas and the paleobathymetry for
the deep oceans:
• Global paleogeography is an important boundary condition for
constraining elevation models of continental topography and oceanic
bathymetry, the geography of oceanic gateways and shallow continental
(epeiric) seas that influence oceanic circulation, ocean heat transport, and
climatic conditions (e.g., Crowley and Burke, 1998). The relative posi-
tions of the continents are well known back to ~200 to 180 Ma, when the
oldest extant ocean floor was formed, but the uncertainty of deeper time
paleogeographic reconstructions increases dramatically going farther back
in time because of the absence of a seafloor record (Ziegler et al., 1983;
Scotese, 2004; Blakey, 2008). Paleolatitudes can often be resolved to within
about ±5°, which is somewhat coarser than the geographic resolution of
the global climate models used for recent IPCC simulations of future cli -
mate. However, a greater concern is that the lack of geological record of
intervening ocean basins, particularly prior to the Cretaceous, means that
longitudinal control is not possible.
• Paleotopography is an important boundary condition for pre-
dicting stationary wave patterns in the atmosphere, convective effects
associated with uplifted regions, and their impact on the distribution of
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90 UNDERSTANDING EARTH’S DEEP PAST
precipitation. Promising new approaches to reconstructing paleoaltimetry
have been developed in recent years. The application of the fossil leaf
stomata index to paleoaltimetry is based on the established relationship
between leaf stomata frequency and ambient pCO2 and the predictable,
globally conserved decrease in pCO2 with altitude (McElwain, 2004). The
uncertainty, which can reach levels as large as the potential height of the
surface (Peppe et al., 2010), is determined by the uncertainty in the CO2
concentration in air as a function of time. The oxygen isotope compositions
of pedogenic minerals (Rowley and Currie, 2006; Forest, 2007; Sahagian
and Proussevitch, 2007) and the hydrogen isotope composition of n-alkanes
from epicuticular plant waxes preserved in lacustrine deposits (Polissar et
al., 2009) may be sensitive proxies of surface water and precipitation compo-
sitions and, in turn, paleoelevation through isotope-altitude relationships.
As this proxy is based on systematic trends in the distribution and isotopic
composition of modern precipitation with climate and topography, the
uncertainty in estimates is dependent on the degree to which the isotopic
composition of paleoprecipitation is faithfully recorded by the authigenic
(formed in situ) minerals (Blisniuk and Stern, 2005). The recently developed
“clumped-isotope” carbonate paleothermometer (see discussion below)
shows good potential for paleoelevation reconstruction, using assumed
temperature lapse rates with elevation (Ghosh et al., 2006; Quade et al.,
2007). Recent studies have demonstrated that surface uplift influences the
regional climate and isotope distribution and thereby severely complicates
paleoaltimetry interpretations (e.g., Ehlers and Poulsen, 2009; Poulsen et al.,
2010). Therefore, ultimately, the accurate reconstruction of paleotopography
requires some degree of iteration between modeling and proxy methods
since topographic relief strongly affects regional climate patterns, influenc-
ing all of the proxy-based estimates (Ehlers and Poulsen, 2009).
• Paleobathymetry has been reconstructed back to the late Jurassic
through well-known age-depth relationships for oceanic crust (Parsons
and Sclater, 1977; Stein and Stein, 1992), but paleobathymetry for older
parts of the record is far more challenging to constrain because of the loss
of seafloor through subduction. The development of oceanic plateaus and
oceanic swells, and uncertainty in how the rate of ocean floor production
has varied over time (Rowley, 2002; Stein and Stein, 1997), further chal-
lenges reconstructions of paleobathymetry and hence sea level change
(Kominz, 1984). Besides eustasy, uncertainties in estimating changes in the
rate of ocean floor production also impact estimates of mantle outgassing
that have been used to drive carbon cycle models and delineate the evolu-
tion of atmospheric pCO2, pO2, and CH4 through the Phanerozoic (Berner,
2006, 2009; Beerling et al., 2009). This uncertainty is significant considering
that for the time intervals for which seafloor is preserved (≤180 Ma), the rate
of change of seafloor production remains a debated issue (Rowley, 2002).
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DECIPHERING PAST CLIMATES—RECONCILING MODELS AND OBSERVATIONS
• The Earth’s vegetation contributes to, and is affected by, a variety
of important climatic feedbacks and is thus an important boundary condi -
tion for paleoclimate modeling. Changes in vegetative land cover directly
influence albedo and Earth surface radiation (Betts, 2000; Chase et al.,
2000; Pielke et al., 2002; Marland et al., 2003; Horton et al., 2010). Changes
in terrestrial vegetation also lead to changes in evapotranspiration and the
hydrological cycle (Shukla and Mintz, 1982; Rind et al., 1990; Baldocchi
et al., 2000; Alpert et al., 2006), with attendant indirect effects on radia -
tive fluxes and atmospheric chemistry as a function of changes in cloud
cover and water vapor mass in the atmosphere (Elliott and Angell, 1997;
Hennessy et al., 1997). In addition, vegetative land cover both influences
and is influenced by soil moisture content, and changes in soil moisture
content can have significant climatic effects through shifts in the relative
influences of latent heat flux versus sensible heating (Alpert et al., 2006;
Niyogi and Xue, 2006). Vegetation-climate feedbacks have not yet been
fully incorporated into GCMs because of the difficulty of parameterizing
the complex, nonlinear interactions that range from cellular scale in
physiological approaches to regional or global scale in biome and physical
approaches to plant definition (Alpert et al., 2006). GCMs incorporating
vegetation-climate feedbacks, however, generally yield amplified climate
responses such as higher climate sensitivities (up to 5.5°C per CO2 dou-
bling; Cox et al., 2000), larger amplitudes of paleoglacioeustasy (Horton
et al., 2010), and greater high-latitude amplification of warming (DeConto
et al., 1999) relative to models lacking such feedbacks. The deep-time geo-
logical record offers several large-scale “natural experiments” (see Chap-
ter 2) from which insights regarding the operation and scaling of these
vegetation-climate feedbacks can be gleaned (Peteet, 2000; Parmesan and
Yohe, 2003; Horton et al., 2010). The knowledge of the composition and
spatial distribution of vegetation on a global scale, prior to the evolution
of angiosperms (Early Cretaceous) and grasslands (Cenozoic), however,
must be further developed. A far more coordinated effort is needed to
expand scientific understanding of fossil plant physiological mechanisms
and to synthesize disparate paleobotanical data into more comprehen -
sive and temporally constrained compilations that can be used to refine
dynamic vegetation models for climate modeling.
Further improvements in scientific knowledge of these physical and
ecological boundary conditions will require more detailed analysis of paleo-
magnetic, paleoclimatic, paleotectonic, and paleontologic data (Van der
Voo, 1993; Parrish, 1998; Kiessling et al., 1999), as well as development of
more sophisticated geodynamic models. For example, there is growing evi-
dence for a systematic shallow bias (5-10°) in paleomagnetic data from the
sedimentary record (Tauxe and Kent, 2004), increasing the paleolatitudinal
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DECIPHERING PAST CLIMATES—RECONCILING MODELS AND OBSERVATIONS
BOX 4.3
Miocene Climate Change—CO2 or Ocean Trigger?
The Miocene was a pivotal time in the current icehouse, marking the
transition between fluctuating glacial conditions and rapid expansion of the
Antarctic ice sheet under a cooler climate (Zachos et al., 2001a; Shevenell
et al., 2004). Variations in oceanic heat transport and atmospheric vapor
transport, governed by changes in deep-ocean circulation and oceanic
gateways as well as decreasing pCO2, have been implicated as drivers
of this late Cenozoic global cooling (Raymo, 1994; Flower and Kennett,
1995; Holbourn et al., 2007). Recent coupled ice sheet-climate model
simulations suggest that the Miocene climate transition was largely CO2-
induced, involving the crossing of a threshold pCO2 of 400 parts per million
by volume (ppmv) (Langebroek et al., 2008). Overall, estimates of pCO2
during the Miocene range from 140 to more than 700 ppmv (Table 4.1),
illustrating the critical need for improved CO2 proxy methods (Tong et al.,
2009). Moreover, contrasting estimates of atmospheric CO2 exist for the
transient period of mid-Miocene warmth (~17 Ma; compare marine-based
estimates of Pagani et al. [2005], and vascular plant-based estimates of
Kürschner et al. [2008]) confound the precise relationship between tem-
perature and CO2.
TABLE 4.1 Estimates of Atmospheric CO2 Levels for the Middle
Miocene
CO2
Reference (ppmv) Method Uncertainties
δ11B
Pearson and 140 to ~20% at low CO2, very
Marine
Palmer (2000) 300 high at >500 ppmv
Marine δ13C
Pagani et al. 180 to ~20% near modern
(1999) 290 ocean conditions,
(alkenone carbonate);
infinite at CO2 ≥ 2,000
marine δ18O
ppmv
Royer et al. 300 to Leaf stomatal indices ~20% at low CO2, very
(2001b) 450 high at >500 ppmv
Kürschner et al. 300 to Stomatal frequency ~20% at low CO2, very
(2008) 600 data from tree species high at >500 ppmv
Cerling (1991) <700 Paleosol carbonate 50-100%
δ13C
SOURCE: Modified from Tong et al. (2009); uncertainty estimates based on Hansen
et al. (2008); Pagani et al. (2005); Royer et al. (2001b); and Cerling (1991).
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96 UNDERSTANDING EARTH’S DEEP PAST
calculated for much of the past 400 million years. For the Cenozoic, boron
isotopes of foraminiferal calcite (assumed to mirror that of δ11B of borate
in seawater) have been used as a proxy for seawater pH and pCO2 (e.g.,
Sanyal et al., 1997; Spivack and You, 1997; Foster, 2008; Pearson et al.,
2009). Where the δ11B proxy has been tested against alkenone-based pCO2,
there is a good agreement (e.g., Seki et al., 2010). In contrast, estimates
of Mesozoic and Paleozoic pCO2 are based largely on soil carbonate and
goethite paleo-CO2 proxies. The carbonate-based paleo-pCO2 barometer
(Cerling, 1991, 1992), however, is particularly sensitive to soil CO2, a
parameter that is challenging to constrain in fossil soils (Ekart et al., 1999;
Breecker et al., 2009) and varies with soil moisture and productivity. Soil-
formed iron oxides (goethite, gibbsite) provide a complementary min-
eral paleobarometer, because they form in soils that do not accumulate
carbonate and the carbon isotope composition and content in their ferric
carbonate component (Fe(CO3)OH) are typically diagenetically robust
(Yapp and Poths, 1992; Schroeder and Melear, 1999;Yapp, 2004; Tabor and
Yapp, 2005). Despite the large uncertainties associated with mineral-based
pCO2 estimates, comparison of mineral- and plant-based estimates are
important because of their complementary differences in sensitivity—
plant-based proxies lose sensitivity above 800 to 1,000 ppmv, whereas
mineral-based proxies are more sensitive above 1,000 ppmv (Royer et al.,
2001a)—and the lack of extant plant calibrations for stomatal index-based
estimates in pre-Cretaceous intervals.
Marine Temperatures
A major challenge in determining ancient climate sensitivity is the
need for robust estimates of ocean temperatures and latitudinal ocean
temperature gradients. This need is particularly great for climate recon-
structions of the pre-Cretaceous, for which we presently have only binary
“icehouse or greenhouse” reconstructions (Figure 1.2) and highly inter-
pretive and largely unpublished climate syntheses (e.g., PALEOMAP
Project1). Several proxies have been employed to reconstruct ancient
sea surface and deep-water temperatures, each with its strengths and
limitations. Measuring multiple proxies not only provides refined SST
reconstructions but also offers constraints on taxonomic (vital), environ -
mental, and diagenetic influences that might affect each proxy. For several
decades, the gold standard for reconstructing sea surface and bottom tem-
peratures for the post-Jurassic has been from δ18O values of foraminifera
(e.g., Shackleton, 1987; Zachos et al., 2001a), whereas for older geological
intervals, it has been the δ18O values of metazoan skeletal calcites (Veizer
1 See http://www.scotese.com/climate.htm.
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DECIPHERING PAST CLIMATES—RECONCILING MODELS AND OBSERVATIONS
et al., 1999; Grossman et al., 2008) and conodont apatites (Joachimski et al.,
2006; Buggisch et al., 2008). There is, however, growing appreciation for
the susceptibility of fossil carbonate to diagenesis and thus overprinting
of the precipitation (seawater) temperature signal in its δ18O values (Box
4.4) (Schrag et al., 1995; Pearson et al., 2001, 2007; Came et al., 2007; Gross -
man et al., 2008; Kozdon et al., 2009; Cochran et al., 2010). Furthermore,
carbonate δ18O values are influenced by both temperature and seawater
δ18O, which in turn has varied through time due to varying ice sheet δ18O
composition, their effects of waxing and waning ice sheets on seawater
δ18O, and surface net evaporation balance, in particular in the broad, shal -
low epicontinental seas of the pre-Cenozoic (Holmden et al., 1998; Veizer
et al., 1999; Panchuk et al., 2006). The compound influence of changing
seawater δ18O and temperature on carbonate δ18O values for periods that
were not ice-free and the possible effects of seawater alkalinity on carbon -
ate δ18O (Spero et al., 1997; Zeebe, 1999) have led to increased focus on
independent temperature proxy methods.
The Mg/Ca ratios of planktonic foraminiferal calcite have been shown
to be sensitive to temperature, increasing exponentially with increasing
SSTs, providing an SST paleothermometer (e.g., Elderfield and Ganssen,
2000; Lear et al., 2000). Application of the Mg/Ca temperature proxy to
deep-sea sediments has challenged climate change paradigms by docu-
menting dynamic glacial-interglacial temperature variation in Pleistocene
tropical oceans and linkages to extratropical climates, and refining phasing
between changes in atmospheric CO2, surface and deep-ocean tempera-
ture changes, and Antarctic glaciation (Lea et al., 2000; Zachos et al., 2008).
Mg/Ca paleothermometry, however, requires species-specific calibra-
tion to account for environmental and vital (taxonomic) effects and may
be limited by salinity, pH, and/or carbonate ion effects on magnesium
partitioning in foraminiferal calcite (Dekens et al., 2002, 2008; Lear et al.,
2002, 2008; Russell et al., 2004; Coxall et al., 2005; Ferguson et al., 2008;
Hoogakker et al., 2009). Importantly, recent multiproxy reconstructions of
deep-sea sediments document good agreement between alkenone-based
and Mg/Ca temperature estimates (Bard, 2001; Dekens et al., 2008). In
addition to deconvolving the temperature and ice volume signals in the
isotopic record, integrating foraminiferal δ18O values and Mg/Ca ratios
has permitted the reconstruction of oceanic surface salinity distributions
and their effect on oceanic heat transport (e.g., Schmidt et al., 2004). Efforts
to apply the Mg/Ca temperature proxy to pre-Cenozoic biotic calcites
(e.g., mollusks) have been limited by the influence of vital effects in these
calcifying organisms (Immenhauser et al., 2005), variation in seawater
alkalinity and [CO32–], and the need to account for likely secular variation
in seawater Mg/Ca and ocean saturation state.
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98 UNDERSTANDING EARTH’S DEEP PAST
BOX 4.4
Proxies and the “Tropical Climate Paradox”
Proxy evidence for mid- to high-latitude warmth in both the marine
(Crowley and Zachos, 2000; Zachos et al., 2001a) and the continental
(Greenwood and Wing, 1995; Wolfe, 1995; Markwick, 1998) realms of
the Early Eocene (57-50 Ma) are consistent with global warmth and a
high-CO2 atmosphere (Pearson and Palmer, 2000; Pagani et al., 2005).
However, marine isotope data from low paleolatitudes initially indicated
little or no ocean surface warming, and perhaps even cooling, in the Eocene
tropics (Shackleton and Boersma, 1981; Barron, 1987), leading to what was
termed the “cool tropics paradox” (Barron, 1987). Subsequently, scientists
have revisited the low-latitude planktonic foraminiferal record (Pearson
et al., 2001, 2007; Norris et al., 2002; Kozdon et al., 2009), arguing that
diagenesis on the cold deep seafloor has imparted a cool overprint signal
to the oxygen isotopic composition of warm-water planktonic foraminifera,
biasing the record in the tropics toward cold temperatures.
The tropical climate paradox may be fully resolved with new estimates
of tropical SST for the Eocene derived from pristine foraminifera and a new
organic molecule proxy, TEX86,a that indicate temperatures 5-10°C warmer
than previous reconstructions (Pearson et al., 2007) (Figure 4.5). Such
paleotemperature reconstructions further challenge the tropical thermo-
static regulation hypothesis presented in Chapter 2 and imply that the
cooling trend of the Eocene was primarily a high-latitude phenomenon
with little effect on the tropics, where climate remained warm and stable
(Pearson et al., 2007).
a Tetraether index of 86 carbon atoms; paleothermometer based on the composi-
tion of membrane lipids of marine picoplankton.
Several additional proxy methods show promise for paleothermometry
but are still in the development and testing phases:
• The calcium isotope (δ44Ca) composition of well-preserved fora-
minifera may provide an independent paleotemperature proxy and a test
of the reliability of the Mg/Ca proxy (e.g., Nägler et al., 2000), although
the complex calcium isotope fractionation behavior, ancient seawater
δ44Ca, and possible vital effects during biocalcification are not yet fully
understood (Gussone et al., 2009). Limited δ44Ca data for Cretaceous rudist
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DECIPHERING PAST CLIMATES—RECONCILING MODELS AND OBSERVATIONS
Olig.
35
late
40
middle
Age (Ma)
45
Eocene
50
early
55
Paleocene
late
modern
60 SST range
e.
10
5 15 20 25 30 35
Temperature (o C)
FIGURE 4.5 Tropical paleotemperature estimates based on δ18O data
from unaltered benthic and planktonic foraminifera (red circles and
squares, respectively) and from TEX86 data (yellow circles) from onshore
drill core samples from Tanzania. These are plotted for comparison with
benthic (open circles) and planktonic (green circles) foraminifera δ18O
data from ODP cores collected from the tropical Pacific Ocean, inferred
to have been diagenetically altered and thereby indicating anomalously
low paleotemperature estimates.
SOURCE: Pearson et al. (2007).
calcite demonstrate the potential for extending this temperature proxy to
pre-Cenozoic calcitic macrofauna (Immenhauser et al., 2005). Similarly,
the magnesium isotope (δ26Mg) composition of aragonite corals shows
promise as a seawater paleothermometer (Wang et al., 2008).
• Recent studies have documented that the clumping (ordering) of
carbon and oxygen isotopes into bonds in biogenic and abiotic carbon-
ates is temperature dependent and an independent record of the fluid
δ18O in which the carbonate precipitated or stabilized during diagenesis
(“clumped isotope method”; Came et al., 2007; Eiler, 2007; Tripati et al.,
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100 UNDERSTANDING EARTH’S DEEP PAST
2010). The carbonate clumped isotope proxy holds considerable promise
for constraining marine and continental (e.g., speleothems, vertebrate
bioapatite) paleotemperatures, although kinetic and diagenetic effects
must be better understood (Affek et al., 2008; Came et al., 2008; Eagle et
al., 2010).
• Biomarker proxies permit reconstruction of paleo-SSTs that are
independent of mineral-based proxy estimates, and include UK37’ (Brassell
et al., 1986), which is based on the relative abundances of C37 alkenones
photosynthesized by marine green algae, and the more novel biomarker,
TEX86, which is based on the relative abundances of C86 tetraether lipids
that form in the water column by archaeal microbiota (Schouten et al., 2002;
Eglinton and Eglinton, 2008). The low diagenetic susceptibility of these
biomarkers and the preservation of tetraether lipids in sediments as old as
the Cretaceous have provided new insight into the climate dynamics of the
recent icehouse (Haug et al., 2004; Kienast et al., 2006; Martrat et al., 2007;
Dekens et al., 2008) through Cretaceous and Eocene greenhouse periods,
including contributing to the resolution of the Cretaceous cool tropics
paradox (see Box 4.4) (Pearson et al., 2007; Schouten et al., 2007). Integra -
tion of multiple proxies greatly increases the range of paleotemperature
sensitivity and probably also increases the precision of estimates because
of the variable sensitivity of different proxies. For example, the Mg/Ca
proxy is least sensitive at low temperatures, whereas the Uk’37’ method is
least sensitive at high temperatures.
INDICATORS OF REGIONAL CLIMATES
Climate models and paleoclimate archives indicate that one of the
larger impacts of global warming is likely to be regional changes in
continental temperatures and precipitation. It is thus imperative to con-
strain past temporal and geographic variability in continental climate—in
particular for periods of abrupt and/or major climate transitions and
for climates that were warmer than the present day—in order to better
understand how regional climates may change in the future. For continen-
tal settings, deep-time paleoclimate reconstructions require a multiproxy
approach involving comparable proxies.
Estimating Continental Temperatures
There are numerous and reasonably well-developed proxies for esti-
mating continental paleotemperatures from lacustrine, coastal, and ter-
restrial deposits. Fossil plant leaves and pollen have been the major
source of continental paleotemperature estimates, because the composi -
tion and physiological properties of plant communities change rapidly
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DECIPHERING PAST CLIMATES—RECONCILING MODELS AND OBSERVATIONS
with temperature (and relative humidity) changes. For flowering plants
(angiosperms), the shapes (style of leaf margin) and sizes (physiognomy)
of fossil leaves have been shown to vary with mean annual temperature
(MAT) (Wolfe, 1993; Wilf, 1997; Kowalski and Dilcher, 2003; Royer et al.,
2005) and have been applied throughout the Cenozoic, including for the
PETM (Wing et al., 2005). Plant-based continental temperature estimates
for periods that predate the evolution of angiosperms (pre-Cretaceous),
however, are missing. Estimates of paleocontinental temperatures also
have been interpreted from pollen distributions in lake sediments, through
comparison with the temperature sensitivity of modern plant communi -
ties (Overpeck et al., 1985). The pollen distribution approach, however,
becomes significantly less certain on longer timescales due to increased
differences among ancient and modern plant species and communities.
Mineral-based isotopic paleothermometry offers an independent set
of proxies that are not restricted stratigraphically to the post-Jurassic, and
for which interpretations are not limited by lack of calibration to extant
plants. For example, the δ18O values of pedogenic carbonates and δ18O
and δD values of hydroxylated clay minerals (kaolinite and smectite)
and iron oxides (goethite and hematite) from fossil soils are being used
to estimate paleotemperatures in ancient soils hundreds of millions of
years old (Dworkin et al., 2005; Tabor and Montañez, 2005), although this
approach requires assumptions regarding the stable isotope composition
of meteoric water. This limitation can be overcome for soil carbonates by
application of the carbonate clumped isotope thermometer to paleosol
carbonates (Passey et al., 2010) and through oxygen isotope–mineral pair
thermometry between coexisting pedogenic clays and iron oxides (Tabor,
2007). Both methods allow for paleosoil temperature estimates that are
independent of the soil water δ18O in which the minerals formed. Con-
ventional δ18O and clumped isotope analysis of vertebrate bioapatites (and
δ18O analysis of body scales of some freshwater fish) offer yet another
independent proxy of continental MAT and have been used to reconstruct
MAT geographic patterns for deep-time warm periods (e.g., Koch et al.,
2003; Fricke and Wing, 2004) and to constrain body temperatures of extant
and extinct vertebrates (e.g., Barrick and Kohn, 2001; Eagle et al., 2010).
In addition to isotopic paleothermometry, the major element composi-
tions of paleosols have been used for MAT reconstructions as far back as the
Paleozoic, based on transfer functions calibrated using modern soils and
associated mean annual temperatures (Sheldon et al., 2002; Retallack, 2005).
Applications of this proxy yield estimates of Cenozoic paleo-MAT that are
consistent with fossil leaf morphology-based paleotemperature estimates
(Sheldon, 2009). For all of these mineral-based proxies, the accuracy and
uncertainty of paleotemperature estimates are largely dependent on using
appropriate fossil material that has been minimally altered by diagenesis.
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102 UNDERSTANDING EARTH’S DEEP PAST
Ultimately, reconstructions of regional variation in continental tem -
peratures will require further calibration studies of existing proxy methods
along with the development of new higher-resolution proxies and con-
tinued development of spatially highly resolved multiproxy datasets. One
promising new direction is the use of biomarkers preserved in lacustrine
and marginal marine sediments that appear to be diagenetically relatively
robust. The presence of picoplankton members of the Archaea in ancient
lake deposits opens up the possibility of using the relative abundances
of their membrane lipids (TEX86) as a paleothermometer of surface water
(Eglinton and Eglinton, 2008). Microbial lipid patterns in modern soils also
show potential as a biomarker paleotemperature proxy, but the influence
of other soil parameters on their abundances (e.g., pH) requires further
calibration studies.
Estimating Regional Hydroclimates
The understanding of regional patterns for continental paleoclimate
changes in the younger part of the record comes primarily from a wealth
of paleolacustrine records and a rapidly increasing speleothem paleo -
climate database. Lacustrine records offer the continuity and temporal
resolution potential to provide key high-resolution sedimentological, geo-
chemical, and paleontologic time series for reconstructing paleocontinental
regional climate change. The paleo-water balance of some ancient lakes,
such as those in the western United States, also provides strong signals of
regional changes in effective moisture (e.g., Benson et al., 2003). Speleothem
physical and geochemical proxies are proving to be powerful recorders of
changes in regional air temperature and effective moisture (e.g., Oster et al.,
2009; Wagner et al., 2010). Notably, precisely dated (U-series) speleothem
records are revealing rapid—century to subdecadal—climate anomalies,
often involving inter- and intrahemispheric atmospheric teleconnections
(Wang et al., 2001, 2005; Yuan et al., 2004). However, the bulk of these con-
tinental records are from the late Cenozoic, and thus have not been used
to reconstruct regional hydroclimate patterns during warm Earth climates.
Reconstructing regional patterns in relative humidity and precipita-
tion is far more challenging in deep-time records because of the overall
lower levels of temporal and spatial resolution, stratigraphic continu-
ity, and geochemical susceptibility to diagenesis, although several new
approaches are being evaluated. For much of the pre-Neogene, scientific
understanding of climate regimes is based on low spatial and tempo-
ral resolution global syntheses of published databases (Ziegler et al.,
2003; Boucot et al., 2004). The morphological characteristics of ancient
soils and their bulk geochemical composition have climatic significance
because the intensity of pedogenesis is dominantly related to precipita -
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DECIPHERING PAST CLIMATES—RECONCILING MODELS AND OBSERVATIONS
tion patterns and surface temperature. Quantitative proxies for esti -
mating mean annual precipitation have been developed that use the
iron content in pedogenic Fe-Mn nodules, the depth to the pedogenic
carbonate horizon, or the chemical composition of particular paleosol
horizons (known as the Chemical Index of Alteration, CIA), all of which
are based on empirical relationships derived from modern soils (Stiles et
al., 2001; Sheldon et al., 2002; Retallack, 2005; Sheldon and Tabor, 2009).
These approaches have been applied to a wide age range of Phanerozoic
paleosols (e.g., Driese et al., 2005; Prochnow et al., 2006), including to the
PETM where these soil proxy-based precipitation estimates indicate a tran-
sient drying in western North America associated with transient global
warming (Kraus and Riggins, 2007). Where multiproxy records permit,
CIA-based estimates of mean annual precipitation are consistent with
independent paleobotanical estimates. The measured δ18O compositions
of ancient soil-formed minerals (phyllosilicates, carbonates, iron oxides,
and sphaerosiderites) have been shown to be reliable proxies of soil-water
δ18O and, in turn, δ18O precipitation at a given paleolatitude after consid-
eration of evaporative or nonequilibrium effects (Stern et al., 1997; Yapp,
2000; Vitali et al., 2002; Ufnar et al., 2004; Tabor and Montañez, 2005). The
fact that these minerals form in equilibrium with ambient hydrological
conditions means that they provide a sensitive record, where formed, of
shifts in seasonality and precipitation rates, allowing them to be used to
evaluate the hydrological cycle in past greenhouse worlds and periods of
icehouse-to-greenhouse transition.
In the same way that marine and lacustrine biomarkers have been
used as quantitative paleothermometers, the hydrogen isotope ratios ( δD)
of individual lipids in fossil plant tissues show great potential for recon -
structing paleocontinental hydrological conditions. Leaf wax n-alkanes are
some of the most abundant lipid molecules biosynthesized by terrestrial
plants (Eglinton and Hamilton, 1967), containing C-bound hydrogen that
is geologically stable (Schimmelmann et al., 1999, 2006). Plant n-alkane
δD values have been shown to correlate with local meteoric water δD
(Sternberg, 1988), further modified by isotope enrichment in leaf water via
transpiration and soil water evaporation (Sachse et al., 2006). Given the
control of relative humidity on these processes, fossil leaf wax n-alkanes
are being explored as a paleoaridity proxy (e.g., Liu and Huang, 2005;
Pagani et al., 2006; Smith and Freeman, 2006). Scientific understanding
of the role of climate and plant physiology on compound-specific δD
systematics is still evolving (Chikaraishi and Naraoka, 2003; Pedentchouk
et al., 2008; Diefendorf et al., 2010), requiring further empirical study
of plant-water-deuterium systematics before this proxy can be applied
straightforwardly to ancient continental systems. Further development
and refinement of emerging proxies such as the aforementioned mineral
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104 UNDERSTANDING EARTH’S DEEP PAST
stable isotope and biomarker approaches, along with improved spatial
and temporal resolution of terrestrial proxy records, are fundamental to
successful calibration and testing of climate models for a future warmer
Earth using deep-time analogues.
INDICATORS OF OCEANIC PH AND REDOX
Recent advances in instrumentation (e.g., multicollector inductively
coupled plasma mass spectrometry, nanoscale secondary ion mass spec-
trometry) coupled with development and modern calibration of geo-
chemical and isotopic proxies that span the periodic table have greatly
expanded the range of paleoceanographic proxies of oceanic redox, alka-
linity, and pH. Studies over the past decade have documented the linear
or exponential relationships between trace element ratios (Mg/Ca, Cd/
Ca, Zn/Ca, U/Ca) and stable (O, C, B) isotopic compositions of carbonate-
bearing fossil fauna and changes in seawater carbonate content [CO 32–]
and pH (Lea et al., 1999; Marchitto et al., 2000; Russell et al., 2004). For
example, the U/Ca ratios of certain planktonic foraminifera genera have
been used to determine variations in seawater carbonate ion content over
the last glacial cycle that track atmospheric pCO2 variations archived in
polar ice cores (Russell et al., 1996, 2004).
The timing and geographic extent of past events of oceanic hypoxia
and anoxia can be resolved through the integration of sulfur isotopes
of reduced (pyrite) and oxidized minerals (carbonate-associated sulfate
in carbonates) and fossil organic matter, and the abundance and stable
isotopic composition of heavy metals (e.g., Fe, U, and Mo isotopes). The
abundance of redox-sensitive transition elements (V, Mo, Fe, Cr) and their
partitioning between various mineral phases in organic-rich deposits
have provided much insight into the origin of O2-deficient waters in pre-
Cenozoic marine basins—in particular in past greenhouse worlds (e.g.,
Sageman et al., 2003; Meyers et al., 2005; McArthur et al., 2008b; Lyons
et al., 2009; Algeo et al., 2010). Carbonate-associated sulfate data coupled
with pyrite sulfur isotope data from Precambrian oceanic deposits have
refined scientific understanding of the oxygenation of Earth’s early atmo -
sphere (Kah et al., 2004). Recently, the utility of molybdenum (δ97Mo),
uranium (δ238U), and iron (δ56Fe) isotopes of various components in
organic-rich black shales has been demonstrated as a sensitive proxy of
oceanic O2 levels, revealing the protracted oxygenation of Earth’s early
ocean in the Proterozoic (e.g., Anbar and Knoll, 2002; Arnold et al., 2004),
and elucidating the global expansion of oceanic anoxia during past warm
periods (e.g., Jenkyns et al., 2007; Gordon et al., 2009; Duan et al., 2010;
Montoya-Pino et al., 2010).
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DECIPHERING PAST CLIMATES—RECONCILING MODELS AND OBSERVATIONS
Synergy of Observations and Models
A synergistic approach combining observations and modeling pro-
vides an optimal strategy for answering critical questions regarding how
Earth’s climate system has responded to varying levels of greenhouse
gases and other forcing factors. Model simulations provide a global picture
of the state of the climate system and also a window into how various pro-
cesses operate to maintain a given climate state. Disparities between simu-
lated climate variables (e.g., surface temperatures, precipitation, ocean
circulation) and proxy observations of these state variables pose questions
regarding how much of the disparity is due to model biases or deficien-
cies and how much is due to observational bias. An example of data bias
is related to simulated tropical and subtropical SSTs during warm Paleo -
gene climates, which for years were high compared to proxy data. Recent
recognition and correction of problems with the proxy data (see Box 4.4)
have not brought models and data into greater agreement. An example
of model bias is related to simulated high-latitude surface temperatures
in warm climate regimes, which have been too low compared to proxy
data. Here, continued improvement and development of new innovative
observational techniques have strengthened the conclusion that models
are challenged to simulate such high polar surface temperatures. This dis-
parity has led to active model exploration of feedback processes that may
operate in warm greenhouse climates but are not revealed by data-model
studies of Earth’s more recent glacial state. Thus, it is to the benefit of both
observational and modeling communities to work in close collaboration
through real-time data-model comparison studies. Disparities between
models and observations represent synergistic research opportunities.
