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factors so that part of the variability observed in the proxy

quantity is not related to changes in the parameter of interest.

Comprehensive analysis of the relationships between proxies

and a series of observed parameters can offer some insight into

how to remove part of the undesired influence of other factors

from the reconstruction.

Given the complexity involved in developing skillful trans-

fer functions as well as in identifying and correcting for poten-

tial sources of error, a common strategy is to reconstruct the

same parameter of interest using different proxies. Such ana-

lyses are known as multi-proxy reconstructions (Fischer and

Wefer, 1999).

Types of proxies:

The systematic use of proxies in quantitative reconstructions of

past oceanic environments originated in the second half of the

twentieth century. Since then, a large number of proxy techni-

ques have been established and more are constantly being

developed. Proxies can be grouped in six broad categories,

based on the type of direct measurement (Fischer and Wefer,

1999). These are listed below, together with brief descriptions

of the main variables of interest associated with each proxy.

The following chapters on paleotemperature, paleoproductivity

and paleocirculation present in more detail the proxies relevant

to each of these fields. Comprehensive discussions of oceano-

graphic proxies can be found in the literature (Bradley, 1999;

Fischer and Wefer, 1999; Henderson, 2002).

 Microfossil assemblages. The relative abundance of plank-

tonic and benthic species of foraminifera, coccoliths, radi-

olaria, diatoms and other organisms can be used to

estimate past ocean temperature, productivity and sea ice

distribution. This proxy type was used for the CLIMAP

project, which produced the first global distribution of

sea surface temperature for the Last Glacial Maximum

(CLIMAP – Project Members, 1976).

 Stable isotopes are based on the ratio between different iso-

topes of an element. The ratios are usually standardized by a

reference value and named after the heavier isotope. The

ratio between

O and

O, for example, is represented by

O. Isotope readings are retrieved mainly from foramini-

fera skeletons (tests), organic matter or other sources (e.g.,

water molecules in continental ice sheets). The amount of

O incorporated by organisms like foraminifera and corals

increases as temperature decreases. Continental ice is rela-

tively depleted in

O compared to sea water. This makes

O a proxy for both temperature and the extent of conti-

nental ice sheets. d

B is used as a proxy for pH. Productiv-

ity, nutrient concentration and past circulation can be

reconstructed from d

N and d

C is taken up with

slight preference to

C during photosynthesis). Together

with microfossil assemblages, d

O and d

C are the paleo-

ceanographic proxies with the most widespread use.

 Radiogenic isotopes. The different solubilities of uranium

and two products of its naturally occurring decay, thorium

(Th) and protactinium (Pa) can be used to estimate the rate

of deep water flow and the flux of particles from the water

column to the sediments. This flux can also be used as a pro-

ductivity estimate.

C preserved in organic matter is used to

estimate the age difference between near surface and deep

waters and hence, ventilation rates.

 Biogenic compounds. The concentrations of some com-

pounds, mainly organic carbon, calcium carbonate and opal,

are used as estimates of past productivity. Calcium carbon-

ate is also an indicator of the calcite compensation depth.

Alkenones are long chained organic molecules resistant to

degradation. The alkenones produced by some coccolitho-

phors can have two or three double bonds in their structure.

The ratio between molecules with two and three double

bonds reflect the temperature at the time of synthesis.

 Elements. The concentrations and ratios of certain elements

in the sediment, organic remains, tests and corals are also

used as proxies. The ratios of strontium to calcium (Sr/Ca)

and magnesium to calcium (Mg/Ca) in biologically precipi-

tated marine carbonates are temperature dependent. The cad-

mium to calcium (Cd/Ca) ratio is used for nutrient

reconstructions. Barium concentration and the Ba/Ca ratio

are proxies for productivity and alkalinity, respectively.

 Sedimentology. Grain size distribution can provide qualita-

tive information about bottom current speeds and act as an

indicator of ice rafted debris. Information about past tides

can be inferred from layered sediments called rhythmites.

The mineralogy of the sediments can be used to establish

source areas and direction of transport for both water and

wind borne sediments.

Conspicuously absent from the quantities of interest listed

above is salinity, a fundamental parameter which influences

many aspects of the ocean environment. There is, at the

moment, no independent proxy for this quantity. As salinity

also influences d

O, an indirect measurement can be obtained

by using independent estimates of temperature (alkenones, for

example) to remove the temperature signal from existing d

series. Attempts have also been made to infer salinity from

microfossil assemblages. Unfortunately, both approaches gen-

erate errors of 1 psu, very large compared to the range of

salinity variability in the oceans (Fischer and Wefer, 1999).

Reconstructions (Models)

The theoretical approach to paleoceanography uses quantitative

ocean and climate models to reconstruct paleoconditions of the

ocean and interpret observations. A number of different ocean

and coupled climate models have been used in the past and

we will give a very short overview of the hierarchy and use

of these models in the broad research area of paleoceanography.

It is, of course, impossible to build an exact model of the cli-

mate system; for the ocean alone, the position and momentum

of approximately 5  10

molecules would have to be calcu-

lated at each instant of time. As an alternative, the ocean, atmo-

sphere, sea ice, or land surface is split into discrete

macroscopic elements with measurable characteristics such as

temperature, density, velocity, etc. The state of and exchange

between these discrete elements follow physical laws and can

therefore be determined with numerical models. Small scale

processes within each element can also influence the large scale

pattern and therefore have to be parametrized. Existing ocean

(and climate) models differ in regards to:

 their temporal and spatial resolution (resolution is defined as

the spatial scale which defines the boundary between pro-

cesses that are resolved by the model and those which are

parametrized)

 the nature of processes which are resolved (for example,

some models include biogeochemical cycles whereas other

models resolve physical processes only)

PALEOCEANOGRAPHY 691

 the number of subsystems taken into account (for example,

an ocean model needs boundary conditions at the surface

which can either be provided by data, or by an atmosphere

model which is physically interacting with the ocean model)

As the different subsystems (ocean, atmosphere, sea ice, conti-

nental ice sheets, vegetation, etc.) of the climate system interact

with each other in a complex manner and on a very broad range

of timescales, climate modelling reduces to the process of iden-

tifying isolable subsystems and processes that are relevant to

the problem at hand. While identifying these subsystems and

processes, the researcher has also to keep in mind that these

processes have to be suitable to be simulated by limited math-

ematical models and will provide results in a reasonable com-

putational time (Crowley and North, 1991).

The simplest class of climate models includes one-dimensional

Energy Balance Models which were first developed by Budyko

(1969) and Sellers (1969). It is interesting to note that these

simple models yield two stable solutions under present day

boundary conditions: the present day climate and a completely

frozen Earth (also called “Snowball Earth”).

Today’s physical ocean models can be classified based on

the following categories:

 geography (regional models, global models)

 physics (hydrodynamic, thermodynamic or hydro-

thermodynamic models)

 surface approximation (free surface, rigid lid)

 vertical discretization (fixed level, isopycnal, sigma-

coordinate, semi-spectral)

 density variation (barotropic, baroclinic)

Because the boundary data for paleoclimatic simulations tend to

contain large uncertainties, global models are better suited for

this research area than regional models. The most comprehen-

sive results are given by global ocean general circulation models

(OGCMs). These models can be driven with reconstructed data

specifying the boundary conditions. For example, numerous

modelling studies restored the ocean surface characteristics to

the CLIMAP data set (CLIMAP – Project Members, 1976)for

simulations of the Last Glacial Maximum.

A better approach than using ocean-only models is to use

coupled ocean-atmosphere models; by computing the surface

boundary conditions, one can bypass the data problem.

However, coupled ocean-atmosphere GCMs often need flux

adjustments. Flux adjustments balance surface fluxes at the

ocean-atmosphere interface to avoid a numerical drift of the

coupled system. As flux adjustments have been “tuned” to

the present day climate, the use of these adjustments to simu-

late past climates is not very reliable. However, some recent

studies use coupled atmosphere-ocean GCMs which do not

need artificial flux adjustments (e.g., LeGrande et al., 2006 ).

Some studies use simple atmosphere models which still pro-

vide reasonable boundary conditions for the ocean model

(Weaver et al., 2001). Other climate subsystems may also have

an important influence on the state of the ocean (continental ice

sheets, sea ice and land surface processes for example). The

class of Earth System Models has been developed recently

and comprise all models which take into account more than

two subsystems of the climate system. Today, Earth System

Models are widely used for paleoceanographic and paleocli-

mate simulations. There is also growing evidence that geo-

chemical interactions between the subsystems are more

important than initially thought (carbon cycle, nitrogen cycle,

methane, etc.) and some initial attemps in integrating these

cycles in Earth System Climate models for paleoclimatic simu-

lations have been made (Crucifix, 2005). The modelling com-

munity is continuously integrating new processes and

subsystems in their models to obtain a better representation of

the climate system dynamics.

Combining models and proxy data

The interpretation of paleoproxy data is an ongoing challenge

for paleoclimate scientists. As a striking example, one could

compare the studies of Clark et al. (2002) and Bond et al.

(2001). Both papers are highly regarded, yet draw opposite

conclusions from the atmospheric D

C record. Whereas Bond

et al. (2001) relate the variability of atmospheric D

Cto

changes in solar radiation, Clark et al. (2002) interpret the same

type of record as a signature of variability in the thermohaline

circulation and ocean heat transport. On the other hand, model

simulations of past climates depend strongly on boundary con-

ditions, assumptions, and the model used for the study. The

simulated climate for a certain time span can be radically differ-

ent depending on the model and boundary conditions used.

Interpretation of measured paleoclimate data is thus urgently

needed through collaboration between modellers and observers.

One possible approach to combine proxy data with climate

models is that of “data assimilation.” Data assimilation

involves the construction of a field that accommodates best

the information obtained from paleoproxies with the physical

(and dynamical) constraints of the climate system using

coupled climate models (e.g., Paul and Schäfer-Neth, 2005).

Another possible approach is to incorporate paleoproxy data

(e.g., d

O, deuterium excess, D

C, d

Be, etc.) as prog-

nostic active tracers in climate models. Perturbations (such as

meltwater events or changing solar activity) and other climate

states (e.g., the Last Glacial Maximum) can then be simulated

and the behaviour of these simulated proxies can be compared

to observed proxy data obtained from ice cores, marine sedi-

ments and other records. This has been done with uncoupled

ocean (or atmosphere) general circulation models (e.g.,

Schmidt, 1999; Werner et al., 2000; Butzin et al., 2005), vege-

tation models (Kaplan et al., 2002) and continental ice sheet

models (e.g., Clarke et al., 2005). However, the importance of

interactions between atmosphere, oceans and other systems

such as the biosphere and the cryosphere point to the necessity

of using coupled models. To date, there have been only a

few studies simulating paleoproxy data with either coupled

ocean-atmosphere GCMs or Earth System Models (e.g.,

Stocker and Wright, 1996 ; Meissner et al., 2003; Roche et al.,

2004; Crucifix, 2005; LeGrande et al., 2006).

A third way to bridge the gap between the modelling and

proxy data approach is to find locations of proxy data records

of special interest with the use of climate models. A simulation

including prognostic paleoproxy tracers can determine the geo-

graphical region of greatest impact on a given paleoproxy data

during a given climate event.

In conclusion, large amounts of paleoproxy data have been

retrieved from various types of archives, but attempts to use

numerical models for verification and interpretation of this data

are sparse. The science of using three-dimensional climate

models to interpret paleo records is still in its infancy.

Processes

In this Section, we give a short overview of the processes and

boundary conditions which might have influenced the oceans

in the past.

692 PALEOCEANOGRAPHY

Paleotides

The history of tides over geological time is associated with the

evolution of the Earth-Moon system and the shape of ocean

basins. Tidal currents generate friction at the bottom of the

ocean resulting in the transfer of energy and angular momen-

tum associated with the Earth’s rotation to the Moon’s orbital

motion. This process has been gradually slowing the Earth’s

spin and increasing the radius of the Moon’s orbit. According

to some estimates, at 620 million years ago (Ma), days were

approximately 22 h long and the Earth-Moon distance was

about 96% of its present day value (Williams, 2000). The main

proxy for tidal (and Earth-Moon system) changes over periods

of millions of years is based on the analysis of tidal rhythmites,

laminated sediments whose deposition is associated with tidal

currents (Williams, 2000).

Tidal dissipation depends strongly on the shape of ocean

basins. As tectonic processes cause the basins to change, the

effects of tidal friction should also change. There are many

indications that this is the case. For example, the present rates

of dissipation appear to be higher than the average rates over

the planet’s history (Kagan, 1997; Gills and Ray, 1999).

Over the Quaternary, the shape of ocean basins was altered

due to changes in sea level and the presence or absence of

ice shelves. Numerical models show that in the Labrador Sea,

tidal amplitudes during the last glacial were about twice as

large compared to present day conditions. This has led to the

suggestion that these higher tides could have destabilized float-

ing ice shelves and caused Heinrich events (Arbic et al., 2004).

Another connection between tides and climate relates to a

millennial tidal cycle. It has been proposed that very high tides,

occurring every 1,800 years, can cause increased mixing

and cooling of the sea surface. This cooling would be related

to abrupt climate change observed with similar periodicity

(Keeling and Whorf, 2000). Some researchers contest this

hypothesis, arguing that tidal forcing at these frequencies is

very weak (Munk et al., 2001).

Radiation

Incoming short wave radiation from the Sun is the ultimate

source of energy for ocean dynamics, the hydrological cycle

and life in the oceans. Geographic and seasonal variations in

the intensity of insolation result in temperature and pressure

gradients which have an important influence on ocean

dynamics and the climate system. Although the ocean surface

circulation is mostly wind driven, the winds themselves are

the result of the uneven distribution of energy on Earth. Den-

sity driven currents on the other hand depend strongly on tem-

perature (and thus indirectly on energy distribution) and

salinity gradients (which result from precipitation/evaporation

patterns and are thus closely linked to the hydrological cycle).

Incoming solar radiation changes over a range of very dif-

ferent timescales. Firstly, the solar luminosity has gradually

increased throughout the Earth’s lifetime. It is estimated that,

during the early days of our planet, the solar luminosity was

25–30% weaker than today’s value. According to model

results, such a reduction in incoming shortwave radiation

should result in a completely frozen planet (“Snowball Earth”).

However, even if evidence exists for snowball Earth conditions

in early Earth’s climate history, there is also counter evidence

that during long periods of time the paleoceans were ice-free.

Thus, it is inferred that, at these times, greenhouse gas concen-

trations must have been higher than present day levels in order

to prevent the system from slipping into an icehouse state (see

Faint young Sun paradox, this volume).

On shorter timescales (order of tens to hundred thousands of

years), the incoming solar radiation is modulated by changes in

the Earth’s orbital parameters which describe the character of

the Earth’s orbit around the Sun. These three parameters

change continuously, causing a variation in the total amount

of solar radiation received on Earth as well as the seasonal

and latitudinal distribution of insolation. The first parameter,

called eccentricity, describes the degree of ellipticity in the

Earth’s orbit around the Sun and hence the shape of its orbit.

The characteristic periods of changes in eccentricity are

95,000, 131,000, 413,000, and 2,100,000 years. Eccentricity

is the only parameter which modulates the total global amount

of solar energy received at the top of the atmosphere. The

change in Earth’s axial tilt through time is described by obli-

quity and has a distinct period of 41,000 years. Finally, the pre-

cession of the equinoxes combines the axial precession

(wobbling of the axis) and the precession of the ellipse (rota-

tion of the elliptical shape of Earth’s orbit) and consists of a

strong cycle with a 23,000 year period and a weaker one with

a 19,000 year period (Ruddiman, 2000). Precession and obli-

quity do not alter the total amount of solar radiation received,

however, they change the distribution of incoming radiation

by latitude and by season. All the frequencies of these para-

meters can be found in climatological records of the paleocean.

Thus, orbital parameters play an important role in driving the

climate system and ocean dynamics.

Finally, short-term variability in solar luminosity (which is

correlated with changes in the number of sunspots visible on

the Sun’s surface) acts on timescales of decades to millenia.

Over the last hundred years, the global mean temperature has

followed a trend similar to the sunspot record. Some climate

scientists have hypothesized that periods of global cooling

(e.g., the Little Ice Age (1560–1850)) have been partly

caused by a minimum in sunspots (Sporer sunspot minimum

(1460–1550) and Maunder sunspot minimum (1645–1715)).

Ocean basin changes

The shape of the ocean basins sets the boundary conditions for

the ocean circulation. Different continent configurations result

in different flow patterns. Understanding how the boundaries

influenced past circulation patterns might offer insight into

the dynamics and other important processes of the present

day oceans. For example, the existence of an unobstructed

low latitude passage in the Tethys Ocean (160–14 Ma) has

been associated with increased poleward heat transport

(Hotinski and Toggweiler, 2003). Of course, both reconstruc-

tions and simulations of ocean currents this far into the past

are subject to many uncertainties. For example, even the exis-

tence of a prominent, large scale feature such as the Tethys cir-

cumequatorial current is still not unequivocally accepted

(Barron and Peterson, 1989; Bush, 1997; Poulsen et al., 1998;

Hotinski and Toggweiler, 2003).

A number of proxy and modelling studies show that the

opening and closing of passages between two basins can have

a large impact on ocean circulation and climate. The closing

of the isthmus of Panama which is the gap between the North

and South American continents at 5–3 Ma is thought to have

intensified the Gulf Stream and the associated northward

transport of heat and salt into the North Atlantic (Haug and

Tiedemann, 1998). The closure of the isthmus of Panama has

PALEOCEANOGRAPHY 693

also been associated with the onset of Northern Hemisphere

glaciation (Keigwin, 1982; Lear et al., 2003; Mudelsee and

Raymo, 2005) as well as with changes in water mass properties

and the overturning circulation of the North Pacific (Motoi

et al., 2005). The opening of the Drake passage between

Antarctica and South America (28–33 Ma) and subsequent

establishment of the Antarctic Circumpolar Current is asso-

ciated with the glaciation of the Antarctic continent (Kennett,

1977; Toggweiler and Bjornsson, 2000). It has been proposed

that the relative stability of the Atlantic Meridional Overturning

Circulation (MOC) during the Holocene is related to the open

connection between the Arctic and Pacific Oceans provided

by the Bering Strait (De Boer and Nof, 2004). According to

this scenario, low salinity anomalies in the North Atlantic make

the MOC unstable when the strait is blocked (as during the last

glacial). On the other hand, the flux of low salinity water from

the Pacific into the Arctic through a wide open Bering Strait

during conditions with higher sea levels (last interglacial,

115–130 ka BP) has been linked to the more unstable

MOC (Shaffer and Bendsten, 1994).

Oscillations in sea level result in changes of the surface’s

land to ocean ratio which in turn influences the planetary

albedo. According to model results, alteration of the planet’s

albedo caused by a sea level drop of approximately 400 m dur-

ing the late Ordovician (455–445 Ma) was one of the factors

that could have driven the climate into a cooler state (Herrmann

et al., 2004). By submerging vegetated areas or making new

land available to plants, changes in sea level can also impact

carbon storage. The amount of carbon present on shelves inun-

dated by the rise in sea level since the Last Glacial Maximum

appears to be equivalent to the increase in the atmospheric

stock of carbon during the same period (Montenegro et al.,

2006).

Ice

A gradual cooling over the last 55 million years led to the pre-

sence of extensive continental ice sheets in both the Southern

and Northern Hemisphere. Intensification of Northern Hemi-

sphere glaciation in Eurasia and North America occurred

between 2.7 and 2.5 Ma. With increasing terrestrial ice sheets,

sediment records show an intensification of climate oscillations

between extreme cold glacial maxima and interglacial warm per-

iods on timescales related to the orbital parameters. The bound-

ary conditions for paleoceans varied dramatically between

glacial and interglacial periods. The formation and melting of

extensive ice sheets in North America and Eurasia, important

variations in atmospheric CO

concentrations as well as shifts

in the extent of sea ice changed the radiation balance, salinity

distribution, dynamical forcing and productivity of the oceans.

Continental ice sheets

Continental ice sheet growth and decay during glacial cycles

may have influenced the ocean circulation in several ways:

 Sea level changes due to the storage of freshwater on conti-

nents and bedrock depression may have led to different cir-

culation patterns in the ocean.

 Extraction of freshwater from the ocean to build up conti-

nental ice sheets led to changes in the global average and

distribution of ocean salinity. These changes in salinity

(and therefore density) may in turn have influenced the cir-

culation and heat transport (Meissner and Gerdes, 2002).

 Changes in land surface albedo altered the local radiation

balance over continental ice sheets, which in turn must have

changed sea surface temperatures and circulation.

 Elevation changes in regions of continental ice sheets could

also have caused changes in atmospheric circulation

(Lehman, 1993) by affecting the atmosphere’ s stationary

wave pattern (Jackson, 2000) and therefore the dynamic

forcing of the ocean circulation.

 Changes in global atmospheric temperatures resulting from

the presence or absence of ice sheets led to changes in ocean

temperatures which may in turn have affected the circulation

and heat transport.

Climate variability on millenial time scales has been more

important during glacial periods than during interglacials.

Dansgaard-Oeschger temperature oscillations, which seem to

be part of slow-cooling cycles occurring every 10–15 kyr

(Dansgaard et al., 1993), have been related in several studies

to changes in the strength of the meridional overturning

(MOC, see Thermohaline Circulation, this volume). Local sea

surface salinity perturbations at high latitudes due to meltwater

and/or iceberg discharges (Heinrich Events) may have wea-

kened the MOC (e.g., Rahmstorf, 1995 and references therein)

leading to abrupt climate change (Broecker, 1994; Manabe and

Stouffer, 1995).

Sea ice

Sea ice regulates exchanges of heat and freshwater between

ocean and atmosphere and can change sea surface salinity

through melting or freezing (brine rejection). It insulates the

relatively warm ocean water from the cold polar atmosphere,

changes the surface albedo (and therefore the local radiation

balance) dramatically and influences evaporation and therefore

local cloud cover and precipitation. By changing the surface

characteristics of the oceans, sea ice plays an important

role in deep water formation and meridional heat transport in

the ocean. The positive ice-albedo climate feedback might have

led to a “run away” icehouse feedback in the Precambrian

(“Snowball Earth,” Hoffman et al., 1998).

Conclusion

Paleoceanography is an exciting research area that unifies spe-

cialists with a broad range of backgrounds. Scientists in each of

the two classical schools of paleoceanography (reconstructions

through proxy data and theoretical quantitative analysis using

models) are more and more exchanging expertise and working

together for a better understanding of the past oceanographic

environment. However, inherent problems associated with the

poor spatial and temporal resolution of proxy data as well as

uncertainties related to the proxy data themselves make their

interpretation difficult and sometimes impossible. At the same

time, models are limited by resolved processes, resolution and

the quality of boundary conditions. Overall, there is much pro-

gress to be made in both fields. In spite of these problems,

the geological data as well as model simulations provide a

substantial set of results which gives us some insight on how

the ocean and the whole climate system functions. This knowl-

edge is of ultimate importance to understand and predict future

climate changes due to anthropogenic perturbations.

Katrin J. Meissner, Alvaro Montenegro, and Chris Avis

694 PALEOCEANOGRAPHY

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PALEOCEANOGRAPHY 695

Cross-references

Alkenones

Astronomical Theory of Climate Change

Atmospheric Circulation During the Last Glacial Maximum

CLIMAP

Dating, Radiometric Methods

Eolian Dust, Marine Sediments

Faint Young Sun Paradox

Geochemical Proxies (Non-Isotopic)

Ice-Rafted Debris (IRD)

Little Ice Age

Marine Biogenic Sediments

Maunder Minimum

Ocean Paleotemperatures

Organic Geochemical Proxies

Paleocean Modeling

Paleoclimate Proxies, an Introduction

Plate Tectonics and Climate Change

Quaternary Climate Transitions and Cycles

Snowball Earth Hypothesis

Thermohaline Circulation

PALEOCENE-EOCENE THERMAL MAXIMUM

Approximately 55.5 Ma ago in marine and terrestrial sections all

around the globe, there was an exceptionally large carbon iso-

tope (d

C) excursion (CIE) which is unique in the Tertiary for

its size and rapidity of its onset. This event, which was first

recognized in multiple cores by Kennett and Stott (1991), is con-

temporaneous with a significant warm anomaly (and oxygen iso-

tope anomaly) and has thus been named the Paleocene-Eocene

Thermal Maximum (PETM) (Figure P8). Due to the singular

nature of this perturbation and the ecological, faunal and physi-

cal gradients across it, the base of this event is now used to

define the boundary of the Paleocene-Eocene transition. The pre-

vious definition of the P/E boundary was significantly later (at

the base of the Ypresian section in Europe), and thus placed

the CIE in the late Paleocene. Previously authors had therefore

referred to this event as the Late (or Latest) Paleocene Thermal

Maximum (LPTM). Some authors have since referred to it

as the Initial Eocene Thermal Maximum (IETM). The CIE is

contemporaneous with the Clarkforkian/W asatchian transition,

long accepted as the Paleocene-Eocene boundary in the North

American Land Mammal Series.

Full characterization of the event has been handicapped by

the number of incomplete sections and frequent hiatuses found

near the Paleocene-Eocene transition. The Global Stratigraphic

Section and Point (GSSP) for this transition has been defined to

be the Dababiya Section, 35 km south of Luxor, on the right

bank of the Nile Valley. The event lies within Nano-Plankton

zone 9 (NP-9) and reverse-polarity magnetic period Chron 24r.

The dating of the PETM is estimated to be around 55.5 Ma,

based on the chronology of (1996) using interpolation between

well-dated ash layers.

Background

The Paleocene was a time of large shifts in the global environ-

ment following the disruptions that occurred at the Cretaceous-

Tertiary boundary (65 Ma). Evolutionary increases in diversity,

and increasing ocean productivity marked this epoch, leading

to the most positive carbon isotope values in the Tertiary.

Temperatures were generally warm, and the polar regions (such

as Ellesmere Island) were replete with sub-tropical fauna and

flora (crocodiles, ferns etc.). There is no evidence of extensive

glacial ice, though evidence for absence of sea ice in high

northern latitudes is not as convincing. Tropical temperature

estimates do not show as much (if any) change compared to

present, but tropical proxies are known to suffer more from

diagenetic alteration and the possibly complicating factors

relating to the tropical hydrological cycle.

Focusing more specifically on the 2 Ma late Paleocene to

early Eocene transitional period, global temperatures increased

(to an estimated 2–4



C warmer than present day), deep ocean

temperatures increased by approximately 3



C and d

C values

Figure P8 A compilation of d

O values in marine carbonate sediments

since the Cretaceous showing the PETM anomaly in the context

of Cenozoic variability (after Zachos et al., 2001).

696 PALEOCENE-EOCENE THERMAL MAXIMUM

decreased over this long term (Figure P9; Aubry et al., 1998,

and references therein). This led to the Eocene Hyperthermal

(or early Eocene climatic optimum), the warmest period in

the Cenozoic. The collision of the Indian and Asian plates

continued throughout this period although a land bridge may

have already existed in the late Maastrichtian period. At some

point during the late Paleocene, the island arc volcanism at

this plate boundary would have ceased. However, the opening

of the North Atlantic was perhaps accompanied by increased

volcanism along the spreading ridge there. It is likely that a

land bridge between North America and Europe was estab-

lished during this period due to rapid extrusion of lava flows

in the Faroe-Rockall area. In the Gulf of Alaska accretionary

prism large amounts of sediment were being sequestered (peak-

ing between 65 and 55 Ma) (Hudson and Magoon, 2002).

These tectonic processes may have affected climate through

the carbon cycle by removing sinks of organic carbon, global

temperatures through a limiting of volcanic aerosol input,

sea level and possibly ocean circulation through the closing

of important gateways (in particular the Tethys ocean and access

to the Northern Atlantic). The vegetation over this period was

predominantly tropical and sub-tropical with climatic boundaries

significantly further poleward than at present. Wetlands have

been estimated to be significantly more extensive (Sloan et al.,

1992, 1999). Similarities in deep and intermediate water masses

point to a single, southern ocean source for deep water for most

of this period (Corfield and Norris, 1998).

levels have been estimated to be relatively high based

on geochemical cycling models, but there are many conflicting

lines of evidence from paleosol d

C, or leaf-based stomatal

density proxies, leaving a consensus elusive (Berner, 1994;

Pearson and Palmer, 2001; Royer et al., 2001; Sinha and

Stott, 1994). The levels of other radiatively active constituents

(CH

, volcanic aerosols) are even more highly uncertain.

Dustiness appears to have decreased significantly across this

interval, possibly indicating a long-term decrease in winds

(Rea, 1994). However, the link between changes in radiative

forcing and the long term warming, although much discussed,

remains speculative.

It must be made clear that changes over the whole

Paleocene-Eocene boundary interval (54–57 Ma) were large

and significant, but not necessarily connected with the much

shorter duration PETM event to which we now turn.

Characterization of the PETM event

The principle and unique definer of the PETM is the carbon iso-

tope excursion (Figure P9). Since this was first recognized it has

been found in many deep ocean cores and terrestrial sequences.

In the deep ocean the excursion is 3–4% and has a very rapid

onset (<10,000 years) (Zachos et al., 2000). The initial carbon

isotope excursions appear to have been stepped and extremely

rapid, occuring in bursts of less than 1 kyr over this interval

(Bains et al., 1999; Röhl et al., 2000). The terrestrial excursion is

Figure P9 Detailed benthic foraminiferal carbon and oxygen isotopic series across the PETM event (Zachos et al., 2000) from ODP sites 690

(Maud Rise, South Atlantic), 865 (Allison Guyot, Equatorial Pacific), 525 and 527 (Walvis Ridge, South Atlantic).

PALEOCENE-EOCENE THERMAL MAXIMUM 697

higher (up to 7%) measured in both paleosol carbonates and mam-

malian tooth enamel (hydroxyapatite) (Koch et al., 1992).

Contemporaneous with this excursion, though possibly slightly

lagged, is a large increase in deep ocean temperatures (up to

5–7



C) as indicated by d

O values in carbonate sediments. Iso-

tope signals from upper ocean-dwelling planktonic foraminifera

also register a similar change, indicating that the warming

occurred over the whole ocean column, particularly at higher lati-

tudes. Tropical temperatures, although more uncertain, may have

also increased by 1–2



C (Bralower et al., 1997). The deep water

changes have been hypothesized to have been due to a switch from

high latitude deep water production to a tropical source of warm

saline bottom water (WSBW), but unequivocal evidence for this

is not yet available (Kennett and Stott, 1991).

There also seems to have been a widespread (but not univer-

sal) increase in carbonate dissolution at this point, complicatin-

gattempts to find continuous sequences across the event. There

was an increase in kaolinite clay deposition in coastal locations

at the CIE, which indicates increased weathering, or at least,

increased transport of weathered material to the ocean (Robert

and Kennett, 1994). Some coastal regions saw an increase in the

deposition of black shales, indicative of increased productivity.

There are widespread discontinuities at the Paleocene-Eocene

transition, and evidence of massive slope failures at Blakes

Nose, in particular, in the North Atlantic (Katz et al., 1999).

There is a coincident benthic foraminiferal extinction (BFE)

event at the onset of the carbon isotope excursion, which

saw approximately 50% of the extant benthic species disappear

at this time (Thomas and Shackleton, 1996). This extinction

has been related to a decrease in dissolved O

, increased pH

(consistent with the increased carbonate dissolution) and large

temperature changes. In addition to foraminifera, other benthic

species groups (ostracodes, molluscs) also underwent significant

turnover. In particular, there was a widespread expansion of exo-

tic planktonic taxa and dinoflagellates. On land, the CIE marked

the final appearance of a number of archaic mammal orders, and

the onset of a more “cosmopolitan” range of mammalian fauna

in both North America and Eurasia. Large-scale ecological turn-

over is also indicated at this point.

The duration of any climate event deep in the geological

record is particularly difficult to assess. The most useful meth-

ods rely upon estimates of Milankovitch cycle variability

(mostly precessional cycles 20 Kyr) in d

C records which

can be counted across the event. Other estimates are based on

interstellar dust accumulation rates (as measured by

He/

ratios). Best estimates for the duration of the event as deter-

mined by the d

C excursion are between 90–200 Kyr. These

estimates are consistent with present day estimates for the resi-

dence time of perturbations to the carbon cycle. Climate

perturbations would be shorter than this timescale since the

equilibrium of atmospheric and oceanic carbon happens much

more quickly than the recycling time for carbon isotopes.

Mechanisms of PETM climate change

The size and rapidity of the PETM CIE serve to usefully elim-

inate many candidates for a primary cause. A perturbation of

up to 4% in the global carbon cycle implies a very large input

of depleted carbon. Given reasonable estimates of the isotopic

signature of terrestrial carbon, the necessary mass of carbon

from any source can be estimated. For instance, if the land bio-

sphere were to provide the source, three times the total carbon

content of the current biosphere would need to be added, and

since no disruption on this kind of scale is observed at this

time, this source can be ruled out. Similarly, the amount of vol-

canism (900 present) required would have certainly left other

geochemical traces (Dickens et al., 1995). An intriguing alter-

native is a carbonaceous extraterrestrial bollide impact, for

which there is little evidence as yet (Kent et al., 2003).

The most compelling theory to date for the source of this light

carbon comes from the large scale dissociation of methane from

gas hydrates sequestered on the continental shelf (Dickens et al.,

1995). The methane is produced by methanogenic bacteria that

live off the organic material in sediment in predominantly anae-

robic situations. Methane hydrates (or clathrates) are stable only

when cold and under pressure. If the oceans warm, or pressure is

reduced, the methane gas can be released. The carbon in such

hydrates at present has an approximate value of -60%,and

approximately 1,500–2,000Gt(1Gt=10

g) of carbon would

be required to match the CIE. This amount is significantly less

than the present-day known reserves of gas hydrates in the

ocean. The maximum in d

C in the mid-Paleocene may be an

indication that significantly more methane was sequestered by

the late Paleocene. The methane hydrate hypothesis implies car-

bon additions to the global ocean that are consistent with the

widespread carbonate dissolution at this time.

The main issue with this hypothesis is how the dissociation

was triggered. Was it caused by a long-term ocean warming

that crossed a threshold, a massive slope failure, or an abrupt

change of ocean circulation, possibly related to a switch to

WSBW formation (Bice and Marotzke, 2002; Katz et al.,

1999; Kennett and Stott, 1991)? Was the warming actually

caused by the greenhouse effect related to the CH

(and its oxi-

dation product CO

)? All of these possibilities have been

tackled in the numerous modeling studies related to the PETM.

Modeling

A hierarchy of models from simple box models of the car-

bon cycle to full atmospheric and oceanic general circulation

models have been applied to the PETM. Box models of the

Paleocene carbon cycle have been used to constrain the total

amount of depleted carbon required to match the CIE at about

1,500–2,000 Gt of carbon (Dickens, 2001). These models

have also been used to assess processes that would eventually

remove the excess carbon from the systems by increased

organic carbon burial, increased weathering, or a recharging

of the methane ‘capacitor’. None of these mechanisms have

yet been eliminated as possibilities.

Inputs from methane would have important effects on atmo-

spheric chemistry if emissions went directly into the atmo-

sphere, leading to a reduced oxidation capacity and increased

lifetimes for methane, other hydrocarbons (such as isoprene)

and CO (Sloan et al., 1999). The radiative effects of these ele-

vated methane (and subsequently CO

) concentrations are

important and have been estimated to be sufficient to have

been the cause of the contemporaneous warming (Schmidt

and Shindell, 2003). However, uncertainties in timing, back-

ground climate state, and Paleocene background emissions

make this a tentative conclusion at best. Even more speculative

is a possibly increased radiative role for polar stratospheric

clouds (due to increased stratospheric water vapour associated

with increased methane oxidation in the stratosphere) (Sloan

and Pollard, 1998).

The use of full GCMs has necessarily been limited by the lack

of well constrained boundary conditions (sea surface tempera-

tures (SSTs), vegetation distributions, continental configura-

tions, etc.); however a series of interesting experiments that

698 PALEOCENE-EOCENE THERMAL MAXIMUM

may have implications for the PETM have been published. Simu-

lations using estimates of the varying equator-pole temperature

gradient as a forcing function may give insight into the climatic

consequences of those changes (Huber and Sloan, 2001). How-

ever, models that attempt to predict the SST changes (by using

increased greenhouse gases for instance) are not necessarily con-

sistent with these reconstructions (Sloan and Rea, 1995). In par-

ticular, whether the main westerly wind belts strengthen or not

at the PETM is model dependent (Schmidt and Shindell, 2003;

Sloan and Rea, 1995). Ocean-only model results, although lim-

ited by the lack of an interactive atmosphere, have been used

to: (a) estimate whether hydrates could be destabilized in suffi-

cient quantities by a shift in deep water circulation, or alternately,

(b) whether radiatively-forced climate changes are sufficient

to match the changes in ocean structure reflected in the benthic

and planktonic isotope data (Bice and Marotzke, 2001, 2002).

No definitive conclusions have yet been drawn, but no study

has yet been able to demonstrate the existence of WSBW.

Particular problems for ocean models are the depth of important

sills (in the North Atlantic, the Walvis ridge and the Isthmus

of Panama, for instance) and finding consistent surface fresh-

water and heat fluxes. Other ubiquitous problems with GCM

simulations of warm climates are also relevant here (e.g.,

too cold continental interiors and apparently too large tropical

warming).

Future studies are likely to yield further insights into possi-

ble mechanisms as dynamical oceans are interactively coupled

to atmospheric models for this period, and carbon cycle models

start to take the methane component (including atmospheric

chemistry and subsequent radiative forcing) into account. Some

work is also required to better define the vegetation distribution

(which has consequence for the albedo, chemical emissions and

deposition of other chemical species (such as O

) which impact

the oxidation capacity of the atmosphere (Beerling, 2000)) and

also river drainage systems and wetland extent.

Mesozoic analogs

It has been hypothesized that similar CIE have occurred in the

past. While none of the other excursions in the Cenozoic period

have a documented size and rapidity of the PETM, excursions

associated with ocean anoxic events in the Cretaceous, Jurassic,

or possibly at the end of the Neoproterozoic glaciations do

have similarities. However, the quality of data for these earlier

events is significantly poorer, and so these possible analogs

will probably remain ambiguous for some time (Hesselbo

et al., 2000; Jahren et al., 2001).

Relevance

Estimates of exogenous carbon inputs during the PETM are

comparable to the rates of fossil fuel-derived emissions at

the present day. Therefore understanding the fate of the carbon

during this episode may lead to insights into the eventual fate

of anthropogenic carbon.

The PETM is also one of the first short-term climatic events

that have been closely studied prior to the Quaternary. It is

clear that the high resolution study of deep-time climate events

is likely to be an increasingly fruitful line of research in under-

standing the earth system (Zachos et al., 2000). The singular

example of the PETM has already led to re-assessments of

the role of methane hydrates in the carbon cycle, the time-

scales for carbon recycling, and the importance of atmospheric

chemistry in paleo- (and present day) climate, and it is likely

that further insights will be forthcoming as data become more

precise and widespread.

Gavin A. Schmidt

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Cross-references

Carbon cycle

Carbon isotope variations over geologic time

Carbon isotopes, stable

Methane hydrates, carbon cycling, and environmental change

Oxygen isotopes

Paleoclimate modeling, pre-Quaternary

Paleogene climates

Plate tectonics and climate change

PALEOCLIMATE MODELING, PRE-QUATERNARY

Introduction

Pre-Quaternary paleoclimate modeling is the science of simu-

lating climate change over the earliest 99.9% of Earth’s history

using numerical models of the climate system. The record of

pre-Quaternary climate documents enormous climate swings

with global ice-covered and ice-free end members. A number

of climate factors shaped Quaternary and pre-Quaternary paleo-

climate on timescales ranging from hundreds of thousands to

several years. However, in contrast to Quaternary paleoclimate,

tectonically-controlled processes (including continental drift,

orogenies, and fluctuations in the long-term carbon cycle) and

solar evolution were also important controls on pre-Quaternary

climate. Understanding how these factors have influenced cli-

mate change through the Precambrian, Paleozoic, Mesozoic,

and Cenozoic eras, and how they are expressed in the geologi-

cal record is a principal objective of pre-Quaternary paleocli-

mate modeling.

The utility of pre-Quaternary paleoclimate modeling has been

severalfold. Pre-Quaternary climate modeling has contributed to

ocean, atmosphere, and geological sciences by: (a) identifying

mechanisms of climate change; (b) quantifying the climate re-

sponse to a variety of forcing factors; (c) recognizing climatic

feedbacks that amplify/dampen climatic forcings; (d) identify-

ing limitations of climate proxies from the geological record;

and (e) demonstrating strengths/ shortcomings in numerical

climate models. Pre-Quaternary paleoclimate modeling has

made fundamental contributions to climate science. Among

these, the recognition that the Earth’s climate has experienced

enormous variability on a multitude of temporal and spatial

scales as a result of complex interactions between multiple

climate factors is one of the most profound.

Methodology

Today, “paleoclimate models” are synonymous with numerical

climate models implemented to study past climates. Prior to the

wide availability of numerical climate models, paleoclimate

models also referred to conceptual models constructed using

physically-based rules about the general circulation of the

atmosphere and ocean (e.g., Ziegler et al., 1977; Parrish et al.,

1982). Numerical climate models are mathematical expressions

of the theoretical laws that govern the climate system, approxi-

mated and written in computer code for fast, efficient computa-

tion. The model domain, frequently global in extent, is

discretized into an array of horizontal and vertical grid cells.

Not all models incorporate the full suite of governing equat-

ions in three dimensions. In fact, a hierarchy of numerical cli-

mate models has been used to study past climates. Numerical

climate models can range in complexity from zero-dimensional

energy balance models (EBMs) that simply express the conser-

vation of energy, to three-dimensional general circulation mod-

els (GCMs) of the ocean and atmosphere that predict fluid flow

on a rotating sphere heated by solar radiation. Many aspects of

the climate system cannot be explicitly calculated, frequently

because specific phenomena develop and act on sub-grid cell

scale (i.e., over distances that are much smaller than repre-

sented by a model grid cell). Clouds provide an instructive

example. Cloud motions vary over a horizontal scale of tens

to hundreds of meters. Yet, the horizontal resolution of a

GCM is typically tens to hundreds of kilometers. In order to

represent clouds in GCMs, cloud processes are approximated

or parameterized. Many of the differences between common

types of numerical climate models (e.g., GCMs) arise from dif-

ferent parameterizations. To date, there are several different

classes of GCMs, including atmosphere-only GCMs, ocean-

only GCMs, coupled atmosphere-ocean GCMs, and GCMs

coupled to models of the biosphere, cryosphere, and litho-

sphere. Climate models also differ in their horizontal and verti-

cal resolution. As computational speed and efficiency has

increased, so has the complexity and resolution of numerical

700 PALEOCLIMATE MODELING, PRE-QUATERNARY