Odekon M. Encyclopedia of paleoclimatology and ancient environments

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et al., 1992). The dominance of the 100-kyr periodicity in

glacial-interglacial oscillations of the last million years, on the

other hand, is not easily accounted for by the astronomical

theory of climate since eccentricity variations of the Earth’s

orbit lead to only minor changes in the insolation budget.

Non-linear responses to insolation changes have been pro-

posed – amongst several other possible mechanisms – to deal

with the puzzling 100-kyr oscillations of the Earth’ s climate

(Imbrie et al., 1993).

2. Analysis of phase relationships provides information about

the response time of different oceanic and climate compo-

nents (i.e., sea surface and deep water hydrology (T, S), sur-

face productivity, wind forcing, etc.) and their causal

relationships (e.g., leads/lags indicating which changes

might be forcing others).

3. Analysis of geographic distribution of the response provides

information about which regions play critical roles in

climate changes.

Figure S26 SPECMAP analysis of cold-season SST (Sea Surface Temperature) variations at 17 Atlantic sites in the precession, obliquity and

eccentricity frequency bands. Cyclic changes in SST are analyzed relative to d

O oscillations. Phase-relationships between these two signals are

displayed as arrows following the so-called “phase wheel” representation, with d

O being set at the “zero phase” (¼ vertical). The leads / lags (in

ka) between d

O and SST are also indicated next to each arrow (a negative value indicates that d

O changes lag SST changes, whereas a positive

value indicates that d

O variations lead SST changes). The length of each arrow is proportional to the coherent amplitude of the SST cyclic

change (see the 1



C reference scale shown in each panel) (figure modified from Imbrie et al., 1989).

914 SPECMAP

Figure S26 shows the amplitude of Atlantic Sea Surface Tem-

perature (SST) oscillations at 17 sites in the precession, obli-

quity and eccentricity bands, and their phase relationships

relative to d

O changes. The largest SST changes (longest

arrows) occur generally near 50



N in the North Atlantic and

reflect oscillations of the Gulf Stream and North Atlantic polar

front in response to changes affecting the nearby continental ice

sheets. In this area, SPECMAP results indicate large amplitude

SST changes in the obliquity band (Figure S26, lower left

panel), a result consistent with the strong influence of Earth’s

axis tilt oscillations on high-latitude insolation changes. South

of about 45



N, the obliquity forcing drops down, as shown

by the weak amplitude of SST changes. In contrast, the preces-

sion signal keeps its strength closer to the low-latitudes

(Figure S26, upper left panel). Major climatic features at low-

latitudes – such as monsoons – generally show strong preces-

sion variability.

In the eccentricity band, there is a suggestion of a phase shift

across the equator: in the South Atlantic, temperature changes

generally lead oxygen isotopes, whereas in the North Atlantic

the SST changes either are in phase or lag d

Ochanges

(Figure S26, lower right panel). This early response of Southern

Hemisphere SST suggests a drawback in Milankovitch’s theory

of climate changes. Milankovitch proposed that insolation

changes in the high-latitudes of the Northern Hemisphere are

the key to controlling glacial-interglacial variations. SPECMAP

data show that the insolation over the Southern Hemisphere

(and inter-hemispheric heat transport) has to be taken into

account in scenarios dealing with astronomical control of global

Earth’s climate.

Conclusion

The SPECMAP project launched in the 1980s radically

improved our understanding of how the Ice Age climate

responded to changes in latitudinal and seasonal distribution

of solar energy controlled by variations in the orbital para-

meters of the Earth. The development of a high-resolution time

frame for the upper Pleistocene made it possible to compare

climatic proxy records from sedimentary series at different

oceanic sites to the geographic pattern of insolation changes

with an unprecedented time precision.

The SPECMAP group also proved to be extremely effective

in pointing out some weaknesses of the astronomical theory of

climate (i.e., the puzzling 100-kyr cycle that dominates global

climate changes over the last million years) and called for more

complex mechanisms than those that were originally proposed

by pioneers such as Milutin Milankovitch. For instance, the

fact that Southern Hemisphere sea surface temperatures seem

to lead changes in Northern Hemisphere ice sheets showed

the limits of Milankovitch’s hypothesis which puts emphasis

on insolation changes at about 65



N. Amongst important

key players for climate changes that were not taken into

account by astronomical theory of climate are greenhouse gases

such as CO

. Long records of atmospheric composition

obtained from bubbles trapped in Antarctica and Greenland

ice cores indicate that the pCO2 of the atmosphere showed

large amplitude variations associated with glacial-interglacial

global climate oscillations. During the last glacial terminations,

the atmospheric pCO

rise led the melting of ice sheets by

a few thousand years (Broecker and Henderson, 1998). This

clearly indicates that atmospheric pCO

, which is ultimately

controlled by changes taking place in the ocean, is an impor-

tant primary driver of the Earth’s climate and that Ice Age

oscillations cannot be fully explained through insolation con-

trol on Northern Hemisphere ice sheets.

Franck C. Bassinot

Bibliography

Baksi, A.K., Hsu, V., McWilliams, M.O., and Farrer, E., 1992.

Ar/

dating of the Brunhes-Matuyama geomagnetic field reversal. Science,

256, 356–257.

Bassinot, F.C., Labeyrie, L.D., Vincent, E., Quidelleur, X., Shackleton, N.J.,

and Lancelot, Y., 1994. The astronomical theory of climate and the age of

the Brunhes-Matuyama magnetic reversal. Earth Planet. Sci. Lett., 126,

91–108.

Broecker, W.S., 1987. The biggest chill. Nat. Hist., 96,74–82.

Broecker, W.S., and Henderson, G., 1998. The sequence of

events surrounding Termination II and their implications for the

cause of glacial-interglacial CO2 changes. Paleoceanography, 13,

352–364.

Hays, J.D., Imbrie, J., and Shackleton, N.J., 1976. Variations in the Earth’s

orbit: Pacemaker of the Ice Ages. Science, 194, 1121–1132.

Imbrie, J., and Imbrie, J.Z., 1980. Modelling the climatic response to orbi-

tal variations. Science, 207, 942–953.

Imbrie,J.,Hays,J.D.,Martinson,D.G.,McIntyre,A.,Mix,A.C.,Morley,J.J.,

Pisias, N.G., Prell, W.L., and Shackleton, N.J., 1984. The orbital theory of

Pleistocene climate: Support from a revised chronology of the marine

O record. In Berger, A., et al., (eds), Milankovitch and Climate. Hing-

ham, MA: D. Reidel, pp. 269–305.

Imbrie, J., McIntyre, A., and Mix, A., 1989. Oceanic response to orbital

forcing in the late Quaternary: Observational and experimental strate-

gies. In Berger, A., et al., (eds), Climate and Geo-Sciences. Dordrecht:

Kluwer. pp. 121–164.

Imbrie, J., Boyle, E.A., Clemens, S.C., Duffy, A., Howard, W.R., Kukla, G.,

Kutzbach, J., Martinson, D.G., McIntyre, A., Mix, A.C., Molfino, B.,

Morley, J.J., Peterson, L.C., Pisias, N.G., Prell, W.L., Raymo, M.E.,

Shackleton, N.J., and Toggweiler, J.R., 1992. On the structure and origin

of major glaciation cycles. 1. Linear responses to Milankovitch forcing.

Paleoceanography, 7,701–738.

Imbrie, J., Boyle, E.A., Clemens, S.C., Duffy, A., Howard, W.R., Kukla, G.,

Kutzbach, J., Martinson, D.G., McIntyre, A., Mix, A.C., Molfino, B.,

Morley, J.J., Peterson, L.C., Pisias, N.G., Prell, W.L., Raymo, M.E.,

Shackleton, N.J., and Toggweiler, J.R., 1993. On the structure and origin

of major glaciation cycles. 2. The 100,000-year cycle. Paleoceanography,

8,699–735.

Martinson, D.G., Pisias, N.G., Hays, J.D., Imbrie, J., Moore, T.C., and

Shackleton, N.J., 1987. Age dating and the orbital theory of the ice

ages: Development of a high-resolution 0 to 300,000 year chronostrati-

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sea chronology for the last 750,000 years. Earth Planet. Sci. Lett., 53,

279–295.

Pisias, N.G., Martinson, D.G., Moore, T.C., Shackleton, N.J., Prell, W.,

Hays, J., and Boden, G., 1984. High resolution stratigraphic correlation

of benthic oxygen isotopic records spanning the last 300,000 years.

Mar. Geol., 56,119–136.

Raymo, M.E., Ruddiman, W.F., Backman, J., Clement, B.M., and

Martinson, D.G., 1989. Late Pliocene variations in Northern

Hemisphere ice sheets and North Atlantic deep water circulation.

Paleoceanography, 4, 413–446.

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Backman, J., 1989. Pleistocene evolution: Northern Hemisphere ice

sheets and North Atlantic Ocean. Paleoceanography, 4, 353–412.

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R. Soc. London A, 357, 1731–2007.

Cross-references

Astronomical Theory of Climate Change

Carbon Dioxide and Methane, Quaternary Variations

SPECMAP 915

Glaciations, Quaternary

History of Paleoclimatology

Ice cores, Antarctica and Greenland

Obliquity

Oxygen Isotopes

Precession, Climatic

Pleistocene Climates

Quaternary Climate Transitions and Cycles

Time-Series Analysis of Paleoclimate Data

SPELEOTHEMS

Speleothems, from the Greek spelaion (cave) and thema

(deposit), are chemical deposits formed in caves. Although

close to 200 minerals have been identified as secondary miner-

als in caves, the bulk of formations that occur in caves consist

of carbonates. Dripstones, such as stalagmites (Figure S27a)

and stalactites (Figure S27b), and flowstones are types of spe-

leothems. Stalactites grow downward from the ceiling, while

stalagmites are their counterparts that grow upwards from the

floor. Flowstones, as the name implies, form from flowing

water. There are numerous varieties of speleothems that form

in the cave environment but carbonate dripstones and flowstones

are the most common and voluminous.

The chemical evolution that results in the formation of

speleothem calcite (CaCO

) starts by water absorbing carbon

dioxide (CO

), from the overlying soil to form carbonic acid

) (Reaction 1). CO

levels in soil can reach up to two

orders of magnitude higher than the atmospheric value of

about 370 ppm due to plant respiration and decay of organic

matter (Hendy, 1971). As the carbonic acid containing water

percolates down, it dissolves calcium carbonate (Reaction 2).

Once the carbonate-rich water is exposed to the CO

-poor cave

atmosphere, it releases CO

, resulting in the precipitation of

calcite or aragonite (Reaction 3).

O þ CO

¼ H

þH

ð1Þ

þCaCO

¼ Ca

2þ

þ2HCO



ð2Þ

2þ

þ2HCO



¼ CaCO

þCO

þH

O ð3Þ

Speleothems are typically made of calcite (CaCO

), and to a

lesser extent of aragonite, another polymorph (same chemical

composition but different crystal type) of CaCO

. Calcite

forms hexagonal crystals, while aragonite forms orthorhombic

crystals; which form crystallizes depends on the degree of

saturation of the solution. Highly saturated solutions in warm

climates tend to crystallize aragonite, this is especially true in

caves formed from dolomite [CaMg(CO

)

], because high Mg

concentration from the dissolution of dolomite inhibits calcite

formation. The occurrence of alternating calcite and aragonite

in stalagmites has been interpreted to indicate differences in

the severity of aridity in the region where the speleothems are

collected (Denniston et al., 2000a; Polyak and Asmerom, 2001).

Growth layering in speleothems

Speleothem growth can vary widely, depending on the cave

environment, and in some cases it can preserve annual-scale

layering, especially in stalagmites, where growth layering is

simple surface addition. Figure S28 shows a photomicro-

graph of annual growth banding in a stalagmite from a

cave close to Carlsbad Caver, New Mexico. Each layer, around

20 microns thick, represents a single year’s growth. Annual

banding in speleothems was recognized very early (e.g.,

Broecker et al., 1960), and has been reported from a variety

of settings (e.g., Genty and Quinif, 1996; Brook et al., 1999;

Polyak and Asmerom, 2001; Hou et al., 2003). Annual banding

Figure S27 (a) Stalactites hanging from a low ceiling, Slaughter Canyon Cave, Carlsbad Caverns National Park, New Mexico, USA (photo: Yemane

Asmerom). (b) Giant Stalagmites, Carlsbad Caverns, Carlsbad Caverns National Park, New Mexico, USA (photo: Yemane Asmerom).

916 SPELEOTHEMS

can also be expressed in a variety of other ways, including

luminescence banding of organic matter (Baker et al., 1993),

and in variation of trace elements (e.g., Fairchild et al., 2001;

Huang et al., 2001).

Climate and environmental changes from

speleothem data

Spleleothems contain a wide-range of chemical and physical

proxies of past climate in continents that may be resolved at

an annual scale, given annual-scale banding and the ability to

date them using uranium-series (U-series) isotopic techniques.

The method is based on measurement of the decay isotopes

of uranium, especially the decay of

238

Uto

234

U, which in turn

decays to 230-Thorium. Uranium is soluble in the solution that

forms speleothems, while thorium is not. Thus, the amount

230

Th is a measure of the age of the sample and the concen-

tration of

238

U and

234

U. With the development of modern

mass spectrometric techniques for measuring U-series isotopes

(Chen et al., 1986; Edwards et al., 1987), it is possible to obtain

very precise ages on small amounts, using less than 200 mg of

speleothem powder, yielding typical age uncertainties of less

than 0.5% (2–s) (Dorale et al., 2001).

One of the earliest areas of studies of past climate was look-

ing at the isotopic variations of oxygen (O) and carbon (C)

of speleothem calcite. Hendy (1971) outlined the methodo-

logy and rationale for the approach. Oxygen has two major

(

O) and one minor (

O) isotopes, while carbon has

two stable isotopes,

C and

C. The O isotope variations

of speleothems reflect processes active in the source and

atmosphere meteoric water and the fractionation (one isotope

favored another) that occurs when calcite crystallized from drip

water. If the O isotope composition of the water is known,

either from fluids trapped in the crystal (Matthews et al.,

2001) or known independently, and, if the calcite formed in

equilibrium with the water, then changes in O isotopes may

be related to changes in temperature. In deep enough caves,

this generally reflects the mean annual temperature. Variations

in O isotopic composition may also indicate changes in the

moisture source of the meteoric water from which the drip

water was derived. Such information can give insight into

large-scale changes in ocean-atmospheric interactions, such as

monsoon intensity (e.g., Wang et al., 2001).

The carbon isotopic composition of speleothems primary

reflects: (a) the C isotopic composition of the soil CO

, derived

from oxidation of local organic matter that dissolves into the

drip water; (b) the C isotopic composition of the bedrock,

which is constant for a given site; and (c) isotopic fractiona-

tion during calcite formation (Hendy, 1971 ). Carbon isotopes

in speleothems have been used to observe the change between

plants that utilize the C

photosynthetic pathway and plants

that use the C

pathway. C

plants are adapted to an arid, high

temperature climate, although it has been suggested that they

also evolved to deal with lower atmospheric CO

in the

Cenozoic (Cerling et al., 1998). C

plants have a stronger

preference for

C over

C, as compared to C

plants. Thus,

variations in

C ratios may indicate variations in C

and C

plant cover (e.g., Dorale et al., 1998; Denniston et al., 2000b).

A number of other studies have used variations in the

growth banding of annually-banded speleothems (Asmerom

and Polyak, 2004) as an indicator of changes in precipitation

through time. For example, Polyak and Asmerom (2001)

reconstructed climate change over the past 4,000 years for

the southwestern USA and correlated it with cultural changes

over the same time period. Changes in mineralogy of spe-

leothems (aragonite = arid, calcite = wet) have been used to

study the relative intensity of the monsoon in Nepal (Denniston

et al., 2000a).

The above is a small sampling of the rich-variety of climatic,

cultural and environmental studies that are being conducted

using speleothem isotopic, chemical and physical data. Above

all, speleothems are some of the most beautiful sculptures of

nature, worthy of the highest stewardship. A great deal can be

learned using speleothems and every effort should be put

towards preserving these natural wonders.

Yemane Asmerom

Bibliography

Asmerom, Y., and Polyak, J.V., 2004. Comment on Betancourt et al. 2002:

“A test of annual resolution in stalagmites using tree rings.” Quaternary

Res., 61,119–121.

Baker, A., Smart, P.L., Edwards, R.L., and Richards, D.A., 1993, Annual

growth banding in a cave stalagmite. Nature 364, 518–520.

Broecker, W.S., Olson, E.A., and Orr, P.C., 1960. Radiocarbon measure-

ments and annual rings in cave formations. Nature, 185 ,93–94.

Brook, G.A., Rafter, M.A., Railsback, L.B., Sheen, S., and Lundberg, J.,

1999. A high-resolution proxy record of rainfall and ENSO since

AD1550 from layering in stalagmites from Anjohibe Cave, Madagascar.

The Holocene, 9, 695–705.

Cerling, T.E., Ehleringer, J.R., and Harris, J.M., 1998. Carbon dioxide star-

vation, the development of C-4 ecosystems, and mammalian evolution.

Philos. Trans. R. Soc. Lond. Ser. B Biol. Sci., 353, 159–170.

Chen, J.H., Edwards, R.L., and Wasserburg, G.L., 1986.

238

234

U and

232

Th in seawater. Earth Planet. Sci. Lett., 80, 241–251.

Denniston, R.F., González, L.A., Asmerom, Y., Sharma, R.H., and Reagan,

M.K., 2000a. Speleothem evidence for changes in Indian summer

monsoon precipitation over the last 2,300 years. Quaternary Res., 53,

196–202.

Denniston, R.F., González, L.A., Asmerom, Y., and Reagan, M.K., 2000b.

Speleothem records of early and late Holocene vegetation dynamics in

the Ozark Highlands, USA. Quaternary Int., 67,21–27.

Dorale, J.A., Edwards, R.L., Ito, E., and González, L.A., 1998. Climate and

vegetation history of the midcontinent from 75 to 25 ka. A speleothem

record from Crevice Cave, Missouri, USA. Science, 282, 1871–1874.

Figure S28 A photomicrograph of annual banding, representing about

17 years of growth in an aragonite stalagmite from Carlsbad Caverns,

Carlsbad National Park, New Mexico USA (photo: Victor Polyak and

Yemane Asmerom).

SPELEOTHEMS 917

Dorale, J.A., Edwards, R.L., Alexander, E.C. Jr., Shen, C.C., Richards, D.A.,

and Cheng, H., 2001. Uranium-series dating of speleothems: Current

techniques, limits, and applications. In Mylroic, J., and Sasowsky, I.D.

(eds.), St udi es o f C av e S ed im en ts. N ew Yor k: Kl uw er / P le nu m.

Edwards, R.L., Chen, J.H., and Wasserburg, G.L., 1987.

238

U-

234

U-

230

Th-

232

Th systematics and the precise measurement of time

over the past 500,000 y. Earth Planet. Sci. Lett., 81, 175–192.

Fairchild, I.J., Baker, A., Borsato, A., Frisia, S., Hinton, R.W., McDermott, F.,

and Tooth, A.F., 2001. Annual t o su b-ann ual resolu tion of multip le tr ace -

ele ment t rend s i n spele ot hems. J. Ge ol . S o c. L ondon , 158, 831–841.

Genty, D., and Quinif, Y., 1996. Annually laminated sequences in the

internal structure of some Belgian stalagmites – importance for

paleoclimatology. J. Sediment. Res., 66, 275–288.

Hendy, C.H., 1971. The isotopic geochemistry of speleothems, I. The

calculation of the effects of different modes of formation on the isotopic

composition of speleothems and their applicability as paleoclimatic

indicators. Geochemica et Cosmochimica Acta, 35, 801– 824.

Hou, J.Z., Tan, M., Cheng, H., and Liu, T.S., 2003. Stable isotope records

of plant cover change and monsoon variation in the past 2,200 years:

Evidence from laminated stalagmites in Beijing, China. Boreas, 32 ,

304– 313.

Huang, Y., Fa ir child, I.J., Borsato, A., Frisia, S., C assidy, N .J., Mc De rmott, F.,

and H awkeswor th, C .J ., 2001. Se as onal variations in Sr, Mg, a nd P in

modern spe le othe ms (Grotta di Erne sto, I ta ly). Chem. G eol., 175,

429–448.

Matthews, A., Ayalon, A., and Bar-Matthews, M., 2001. D / H ratios of

fluid inclusions of Soreq Cave (Israel) speleothems as a guide to the

eastern Mediterranean meteoric line relationships in the last 120 ky.

Chem. Geol., 166, 183–191.

Polyak, V.J., and Asmerom, Y., 2001. Late Holocene climate and cultural

changes in the southwestern United States. Science, 294, 148–151.

Wang, Y.G., Cheng, H., Edwards, R.L., An, Z.S., Wu, J.Y., Shen, C.-C.,

and Dorale, J.A., 2001. A high-resolution absolute-dated late

Pleistocene monsoon record from Hulu Cave, China. Science, 294,

2345– 2348.

Cross- refere nces

Carbon Isotopes, Stable

Isotope Fractionation

Mineral Indicators of Past Climates

Oxygen Isotopes

Uranium-Series Dating

STABLE ISOTOPE ANALYSIS

Backgr ound

Atoms of an element always have the same number of electrons

as protons (atomic number), but may differ in the number of

neutrons in the nucleus, producing slight differences in atomic

mass. These are referred to as isotopes. Few elements consist

of just one isotope; most elements in nature have more than

two isotopes. Whereas most of these atoms are naturally stable,

some isotopes a re natu rally unstable (radiogenic), sponta-

n e o u s l y releasing energy and transforming into more stable

atoms. Stable isotopes with low atomic number – the light

stable isotopes – are the major focus of paleoclimate research.

These isotopes include those of oxygen (

O), hydrogen

(

H = D), carbon (

C), nitrogen (

N), and sulfur

(

S).

Notati on

Stable isotopes with lower atomic mass are typically more

abundant than their larger counterparts ( Table S2 ), and their

abundances in natural materials are reported in relative terms

as the ratio of the heavier (rare) and lighter (abundant) isotopes:

R ¼

Abundance of rare isotope

Abundance of abundant isotope

ð1 Þ

Absolute stable isotope ratios are, however, rarely reported

because current analytical instrumentation has a relatively low

sensitivity that makes it unsuitable for detecting very small

natural variations of stable isotopes. Absolute isotope ratios

generally possess poor reproducibility. In paleoclimatic studies,

changes in stable isotope ratios are therefore more relevant

than the absolute isotope ratios. Thus, it is preferable to use

an international reference material with which to compare the

samples. This standard can provide information about changes

in the ratio of rare and abundant isotopes in the sample relative

to the ratio in the standard. Since these variations are less than

one and are reported to at least five or six significant figures,

the delta ( d ) notation is used to represent parts per thousand

(per mil) deviation of the isotope ratio of the sample relative

to that of the reference material (Equation ( 2)).

A= B

Sample

Standard

 1



1; 000 ð2 Þ

where R is the isotope ratio, and the values of d are given in per

mil ( %). Equation ( 2 ) indicates that an enrichment of the rare

(heavy) isotope in the sample relative to the reference material

(standard) produces higher d values. Positive d values, then,

result when R

Sample

 R

Standard

, whereas negative d values

result when R

Sample

< R

Standard

Stand ards

Table S3 lists the internationally accepted reference standards

and their absolute isotope ratios. By definition (Equation ( 2)),

the d value of an absolute standard is 0%. Since absolute

reference materials are currently either exhausted or avail-

able in small quantities, laboratories routinely employ internal

standards that are calibrated against the absolute standards.

These calibrations are needed to convert d values obtained

from internal standards into d values based on absolute stan-

dards (Craig, 1957).

Table S2 Abundance of light stable isotopes

Element Isotope Abundance (%)

Hydrogen

H 99.985

H=D

0.015

Carbon

C 98.98

C 1.11

Nitrogen

N 99.63

N 0.37

Oxygen

O 99.959

O 0.037

O 0.204

Sulfur

S 95.00

S 0.76

S 4.22

H is also called deuterium (D). Although

O and

S are stable isotopes,

they are not currently used for paleoclimatic studies because reproducible

analyses of their abundances in natural materials have only been recently

achieved.

918 STABLE ISOTOPE ANALYSIS

Fraction ation effects

Despite having the same number of electrons and protons, iso-

topes of the same element behave differently during chemical

and physical processes which results in a dissimilar distribu-

tion of isotopes among reactants and products. This partitioning

of isotopes is referred to as isotope fractionation. There are

fundamentally two phenomena that cause isotope fractionation:

lower mobility of heavier isotope molecules and lower binding

energies of lighter isotope molecules. The former produces

lower diffusion rates relative to those of their lighter counter-

parts. Lower mobility also produces a reduction in collision fre-

quency with other molecules during chemical reactions, leading

to slower rates of chemical reaction for heavier molecules.

In addition to higher mobility, faster reaction rates occur with

lighter isotope molecules due to their lower binding energies.

These two phenomena (i.e., mobility and binding energy)

cause, for example, heavier water molecules (

O and

O) to evaporate at a slower rate than their lighter coun-

terpart (

O).

The partitioning of isotopes between substances A and B

(either two different chemical compounds or two phases of

the same compound) can be represented as:

AB

ð3 Þ

where R is the isotope ratio, and a is referred to as the

fractionation factor. The partitioning of isotopes between sub-

stances A and B as represented by Equation ( 3 ) can be equated

to the more familiar equilibrium constant ( K ) of a reaction by

the following equation:

AB

¼ K

1= n

ð4 Þ

where n is the number of atoms exchanged. The magnitude of

the fractionation factor depends on temperature, as indicated by

the pioneer work of H. Urey ( 1947). A general expression that

relates temperature and a . is of the form:

AB

¼ N e

M =T

ð5 Þ

where N and M are temperature-independent coefficients, and

T is temperature. Theoretical, semi-empirical, experimental,

and observational calculations of the magnitude of a with tem-

perature exist for different chemical reactions, isotope

exchange reactions, and phase changes of compounds. These

relationships are typically expressed as the power series solu-

tion of the natural logarithm of Equation ( 5 ):

ln a

AB

¼ C

ð6 Þ

where C

, C

are constants, and T is temperature

(in Kelvin).

Urey ( 1947) was the first to predict the paleoclimatic

potential of the dependence of the partitioning of isotopes on

temperature. His pioneering work focused on thermodynamic

equations that describe the partitioning of isotopes between

water and calcium carbonate. Based on these equations, Urey

argued that the stable-oxygen isotopic composition of carbo-

nate minerals provides information about past temperatures. To

assess the temperature-dependent relationship between isotope

ratios of calcium carbonate and water, mollusks were grown

in water of known isotopic composition at different tem-

peratures (Epstein et al., 1953). The obtained experimental rela-

tionship was in close agreement with Urey ’ s theoretical

determinations, providing support for the idea that stable-

oxygen isotopes can be used for paleoclimatologic research.

Probably the first most significant use of stable isotopes

in paleoclimatology was the work of Emiliani ( 1955). He deter-

mined the oxygen-isotopic composition ( d

O) of foraminiferal

shells from cores retrieved around different ocean basins and

found numerous fluctuations in ( d

O) values that he equated

with glacial / interglacial cycles. As indicated by Equa-

tions ( 3), ( 5), and ( 6), d

O values of foraminifera depend on

seawater d

O values and temperature. Paleotemperatures can

only be determined when seawater d

O values are known.

Emiliani ( 1955) realized that seawater d

O values have chan-

ged through time due to different volumes of continental ice

caps. Employing simple assumptions, Emiliani estimated that

glacial-related shifts in d

O values of seawater produce less

than 20% of the changes observed in foraminifera d

O values,

and the remaining effect was due to temperature. However,

subsequent work by Shackleton and Opdyke ( 1973) strongly

suggests that changes in d

O values of seawater are the main

parameter controlling d

O values of foraminifera.

Although Urey’ s pioneer work showed that stable isotopes

can be used as temperature proxies, the work by Emiliani

( 1955) and Shackleton and Opdyke ( 1973) illustrate the most

important factors constraining paleoenvironmental interpreta-

tions based on stable isotope data. These factors include

temperature and the isotopic composition of water.

Stab le isoto pes in pal eothermom etry

If the isotopic composition of water is known or is assumed

to have remained constant over the time period of interest, past

changes in temperature can be estimated from stable isotope

data. This approach has primarily been employed on carbonate

substrates precipitated in marine settings. For instance, changes

in carbonate d

O values of Holocene corals and marine

bivalves have been interpreted to indicate changes in seawater

temperature, where a decrease in water temperature results

in elevated shell d

O values. Corals in particular offer one

of the highest resolution (seasonal to century-scale) paleocli-

matic archives of the tropics, commonly encompassing the

last 400 yr. A strong relationship ( r

= 0.92) has been found

between sea-surface temperature and the oxygen isotopic com-

position of corals (Quinn et al., 1996; Fairbanks et al., 1997).

Although the relationship between temperature and the extent

O-enrichment in carbonates (Equations ( 3 )–( 5 )) has been

established at equilibrium conditions (e.g., Epstein et al.,

1953), d

O values of modem biogenic carbonate in marine

settings exhibit a systematic offset from the values expected at

equilibrium. For instance, direct measurements of sea-surface

Table S3 Absolute reference material

Standard Isotope ratio Abundance

V-SMOW D/ H 1.5575  10

–4

V-PDB

C 1.1237  10

–3

Air

N 3.677  10

–3

V-SMOW

O 2.0052  10

–3

V-PDB

O 2.0672  10

–3

CDT

S 4.5005  10

–2

Note: V-SMOW, Vienna-Standard Meteoric Ocean Water; V-PDB, Vienna-

PeeDee Belemnite; and CDT, Canyon Diablo Troilite.

STABLE ISOTOPE ANALYSIS 919

temperature and d

O values of corals indicate a significant (up

to 0.5%) deviation from the value expected at equilibrium; this

difference points to a kinetic and/or biological effect during

carbonate secretion (“vital effect”). Fortunately, the isotope off-

set resulting from the biological secretion of carbonate appears

to be constant among specimens of the same species, but differ-

ent among individuals of different taxa. Constant vital isotope

effects in corals and bivalves make their d

O values sensitive

indicators for paleoclimate because a simple correction factor

can be applied to convert biogenic carbonate d

O values into

temperature.

Whereas the isotopic composition of seawater can be

assumed to remain constant (0–1%) in certain locations for

at least the late Holocene, the isotopic composition of water

on land is extremely variable (from about 0% to  35 % for

O). As a result, it is important to distinguish the effect of

temperature from changing d

O values of water to make accu-

rate paleoclimatic reconstructions. An extreme situation occurs

when the range of variability of water d

O values far exceeds

the temperature effect on shell d

O values. Under this sce-

nario, it still might be possible to obtain paleotemperature

assessments from changes in shell d

O values as a result of

the strong covariance found between surface temperature and

the isotopic composition of meteoric water.

Stable isotope ratios in meteor ic water

The basis for the use of hydrogen (dD) and oxygen isotopic

composition (d

O) of meteoric water comes from observations

of modern precipitation in mid- and high-latitudes showing a

strong (r

= 0.78) correspondence between temperature and

values of dD and d

O of meteoric water. This strong relation-

ship has been found to occur on seasonal, interannual, and

multiannual comparisons (Dansgaard, 1964; Rozanski et al.,

1992; Fricke and O’Neil, 1999).

The isotopic composition of rainwater chiefly depends on

two processes: evaporation and condensation. During evapora-

tion, the lighter isotopes are preferentially incorporated into the

gas phase as a result of their higher mobility. Most evaporation

in the world occurs at low- and mid-latitudes, and to a lesser

extent from vegetation and lakes. During condensation, the

lighter isotopes preferentially remain in the vapor phase due

to their higher vapor pressure relative to that of their heavier

counterparts. Removal of heavier isotopes during condensation

enriches the lighter isotopes in the cloud. Subsequent conden-

sation events produce a further enrichment of lighter isotopes

in the cloud, resulting in progressively lower d

O and dD

values of rainwater. This progressive depletion of heavier iso-

topes due to partial rainout is observed when air masses travel

landward and poleward. For instance, rainwater at maritime

and coastal sites exhibits higher d

O and dD values than rain-

water falling in continental interiors (Figure S29). Similarly,

rainwater of tropical regions shows higher d

O and dD values

in relation to rainwater falling at higher latitudes (Figure S30).

Isotope data of meteoric water collected around the world

verifies the gradual decrease in d

O and dD values by showing

a gradient towards lower d

O and dD values of rainwater with

increasing latitude and continentality.

The extent of incorporation of heavier isotopes in conden-

sates (i.e., fractionation between gas and liquid phase) depends

on the prevailing temperature at the cloud base (Gat, 1996).

This dependence appears to be the source of the strong relation-

ship found between temperature and the isotopic composition

of rainwater in mid-latitudes and polar regions (Figure S31).

The dependence of d

O values of meteoric water on

temperature has been used in terrestrial records, primarily in

ice-core studies, to reconstruct past changes in temperature.

In fact, the covariance between temperature and d

O values

has independently been confirmed via bore-hole measurements

in Central Greenland (Johnsen et al., 1995), validating the

accuracy of paleotemperature estimations deduced from isotope

data. Using d

O data, Greenland ice records have revealed

exceptionally large and abrupt changes in climate during the

most recent glacial interval (Dansgaard et al., 1993). These

results suggest that climate is highly variable, showing very

rapid transitions from “cold” to “warm” modes. In Antarctica,

temperature estimates inferred from isotope data reveal a dra-

matic 10



C temperature variation between glacial and inter-

glacial intervals. Similar to Greenland records, the conversion

Figure S29 Changes in isotopic composition of air masses as they travel inside continents. Values depicted in the figure correspond to oxygen

isotope ratios.

920 STABLE ISOTOPE ANALYSIS

of temperature from isotope data is based on modem calibra-

tion studies of Antarctic snow. It is important to indicate, how-

ever, that these modem calibrations might not be directly

applicable to snow or ice accumulated during glacial intervals

because atmospheric and oceanic processes were different from

those operating today. As a result, the isotopic composition of

water in clouds, seawater, and rainwater was likely different

in glacial times with respect to modern (interglacial) intervals.

Although the strong relationship between surface tempera-

ture and d

O values of meteoric water in temperate and polar

regions can be used to reconstruct past temperatures, there are

a number of confounding factors that also control the isotopic

composition of rainwater at a given locality. One of these fac-

tors is the so-called “amount effect” (Dansgaard, 1964). Large

rain events tend to be associated with an anomalous enrichment

of lighter isotopes, leading to lower d

O and dD values

of meteoric water. This “amount effect,” characteristic of tropi-

cal settings, typically overprints any temperature-dependent

isotope fractionation. As a result, tropical regions exhibit a

weak relationship between temperature and the isotopic com-

position of meteoric water.

Another confounding factor is related to the complexity of

cloud formation.

Isotopic and thermal equilibrium must exist during conden-

sation in order to use the isotopic composition of meteoric

water as a temperature proxy. Non-equilibrium conditions exist

in clouds during ice formation or during super saturation

events. Moreover, evaporation during travel from the cloud to

the ground could result in the removal of lighter isotopes. This

evaporation process can also weaken the correlation between

temperature and the isotopic composition of meteoric water.

Another potential significant source of error in correlating

temperature and the isotopic composition of meteoric water

is the mixture of moisture from different sources. For instance,

precipitation in North America results from the interplay of air

masses coming from the Gulf of Mexico, the Pacific Ocean,

and the Arctic region. Because of different evaporation/

condensation histories, moisture derived from these masses

has distinct isotopic composition. Consequently, the super-

imposition of moisture from different sources weakens the

correlation between temperature and isotope data of meteoric

water. This weakening is further enhanced when evapo-

transpiration returns significant amounts of moisture to the

atmosphere. A typical example of this process includes the

Amazon Basin, where about 30% of the precipitation originates

from significant evapo-transpiration of rain forests (Gat, 1996).

Although the interplay of different air masses and evapo-

transpiration affects the relationship between mid-latitude tem-

perature and rainwater d

O values, isotope data from sites

where these processes are pronounced can be used to assess

the intensity of these processes. In these places, isotope data

can consequently be used to assess variations in atmospheric

circulation that either shift storm tracks or enhance/suppress

evapo-transpiration.

Evaporation/precipitation ratios

The partitioning of stable isotopes that occurs during evapora-

tion has been used as an indicator of the balance of evaporation

and precipitation (E/P) in a region. The best systems to assess

E/P ratios include relatively closed aquatic systems, where the

main source of water is meteoric water. In this setting, the oxy-

gen and hydrogen isotopic composition of the closed aquatic

system corresponds to the isotopic composition of meteoric

water plus the fractionation effect related to evaporation. One

of the substrates employed to reconstruct past E/P ratios

include carbonates precipitated in lakes. Since evaporation

favors the incorporation of lighter isotopes in the gas phase,

carbonates in closed aquatic systems experiencing high eva-

poration rates (high E/P ratios) exhibit elevated d

O values.

The carbonates analyzed for E/P studies primarily come from

ostracod shells and lacustrine marls. Isotope analyses on bulk

marls in lakes are relatively easy to perform, allowing the pos-

sibility of performing high-resolution studies. The disadvantage

of lacustrine marl deposits is that bulk carbonates represent a

heterogeneous mixture of authigenic, biogenic, and detrital

components. Moreover, precipitation of these components

could occur under the influence of discrete processes operating

at different places, lake depths, and time. In contrast, isotope

analyses of biogenic carbonates offer an opportunity to mini-

mize uncertainties related to spatial and temporal heterogeneity.

Whereas lacustrine marls tend to precipitate at steady-state

conditions, a number of studies have documented a systematic

offset between d

O values of ostracod carbonate relative to

expected theoretical calcite d

O values formed at equilibrium

(Epstein et al., 1953). This offset (“vital effect”) appears to

be constant within a given species, but it is variable among

Figure S31 Variation of temperature and the oxygen isotopic

composition (d

O) of meteoric water.

Figure S30 Variation of latitude and the oxygen isotopic

composition (d

O) of meteoric water. Notice that lower d

O values

occur in polar regions.

STABLE ISOTOPE ANALYSIS 921

species, yielding values as low as 0.5 % to as high as 2%.

Because the offset of ostracod d

O values from those expected

at equilibrium is constant, ostracod isotope data have success-

fully been used to reconstruct past E / P ratios. For exam-

ple, ostracod isotope data from lakes in the Yucatan Peninsula

indicate that the collapse of the Mayan culture coincided

with a period of extensive aridity in Meso-America (Hodell

et al., 1995).

As indicated before, assessments of E / P ratios from carbo-

nate d

O values can be achieved in closed-basin lakes where

rainfall is the main component in the isotopic hydrological

balance of lakes. Most lakes, however, are not closed because

outlets and groundwater inflow and outflow exist. In these

cases, additional information regarding the hydrological isotope

balance of the basin is needed to perform accurate interpretations

of past changes in E / P ratios from carbonate isotope data.

Although groundwater fluxes are difficult to estimate among

the different hydrological components, coarse assessments of

these fluxes can be achieved from surveys of the isotopic compo-

sition of lake-, stream-, and rain-water during the dry and wet

seasons. These isotopic data combined with information about

rainfall amount, stream flow, and runoff can be used to estimate

the influence of groundwater systems into a lake. Although these

studies provide only a coarse indication of groundwater fluxes,

they are useful to constrain any paleoclimatic interpretation from

authigenic and biogenic carbonates in lakes.

A special case is open-basin lakes with short residence times

(i.e., low ratios of lake / watershed area) because these lakes

tend to exhibit d

O values of lake water close to those of

recharging water. Isotopic analyses of lacustrine carbonates

formed in these lakes can then provide information about tem-

poral changes in the watershed ’ s hydrology. Although recharge

varies from lake to lake, it reflects the balance of precipita-

tion, runoff, evaporation, and groundwater inflow and outflow

(Talbot, 1990).

A potential source of uncertainty in assessing P / E ratios is

the effect of temperature. Because d

O values of carbonates

decrease by about 0.23% per + 1



C, changes in P / E can only

be detected in lake carbonates when the effect of evaporation

is larger than that of temperature, which is the case for lakes

with relatively large surface area and small volume. In contrast,

seasonal changes in P / E ratios in large lakes produce a rela-

tively small range of variability in lake water d

O values. This

small variability in d

O values can be comparable to the range

caused by seasonal changes in temperature, thereby making

carbonate d

O values unsuitable for evaluating past P / E ratios

unless an independent temperature proxy is utilized.

In addition to lakes, inference of E / P ratios from carbonate

isotope data has successfully been employed in corals to assess

changes in the balance of evaporation and precipitation in

regions where seawater d

O values are sensitive to changes

in the balance of precipitation and evaporation. Lindsley et al.

( 1994), for instance, found decadal changes in d

O values

in 300 year coral record from Gulf Chiriqui, Panama.

Because the oxygen isotopic composition of seawater at

this site is sensitive to rainfall amount, these changes were

interpreted to reflect decadal changes in the strength or position

of the lntertropical Convergence Zone.

Carb on isotope s of organi c matter

There are two main approaches employed in reconstructing

paleoenvironments using stable carbon isotopes. The first

approach consists in determining changes in the distribution

of two isotopically distinct groups of vascular plants: C

and

. In this approach, increased abundance of C

plants reflects

increased aridity. The second approach is to infer changes in

humidity from the carbon isotopic composition of tree-rings.

In this approach, elevated d

C values indicate decreased

humidity.

and C

plants possess different photosynthetic mechan-

isms that produce distinct carbon isotope signatures. The C

photosynthetic pathway is an efficient CO

concentrating

mechanism (Leegood, 1999) that provides C

plants with

higher water use efficiency. As a result, most C

grasses

are found today in hot ( > 22



C) and relatively dry places

(Ehleringer, 1997; Collatz et al., 1998). C

plants are common

in the monocot families Cyperaceae (sedges) and Poaceae

(grasses). Almost 50% of all grasses are C

, inhabiting tropical

to temperate regions with abundant warm-season precipitation.

The C

photosynthetic pathway also causes C

plants to show

C values that vary from 9 % to 16%. In contrast, C

plants exhibit d

C values that range from 21% to 30%.

The nonoverlapping d

C values of these two classes of plants

produce a distinct isotopic signature that can be identified in

past environments, provided that sufficient terrestrial organic

matter persists in the geologic record. Changes in the distribu-

tion of C

and C

plant communities have been inferred from

stable carbon isotopic composition of soil organic matter, lake

deposits, and coastal and marine sediments. In this context,

increased abundance of C

plants inferred from elevated d

values has been interpreted to correspond to periods of

increased aridity. In addition to widespread aridity, it has been

argued that C

grasses can out-compete the more abundant C

plants under low concentrations of atmospheric carbon dioxide,

even under relatively wet conditions. The role of aridity and

concentrations on expansions of C

plants in the geologic

record is still unresolved.

Another application of stable carbon isotopes in paleocli-

mate research consists of inferring changes in relative humidity

from changes in the stable carbon isotopic composition of

tree-rings. This approach relies on the extent of the biological

fractionation effect that occurs during photosynthesis by C

trees. The carbon-isotopic composition of a C

plant can be

expressed by the following equation (Farquhar et al., 1988):

Plant

¼ d

Air

 a ðb  aÞC

ð7Þ

where a is the fractionation effect of 4.4% that occurs during

the diffusion of CO

through stomata, and b is the biological

fractionation effect of 28% that results during CO

fixation.

and C

, represent the concentration of CO

inside and out-

side of the leaf, respectively. d

Air

is constant for relatively

long time-scales because the troposphere is relatively well

mixed. Consequently, the only variable actively determining

Plant

in Equation ( 7 ) is the C

/ C

ratio. Field observations

and laboratory experiments indicate that different plant species

could exhibit different C

ratios under the same envir-

onmental conditions, pointing to a species-specific effect. As

indicated before, the natural range of d

C values for C

plants

is between 21% and 30%, which translates to C

ratios

of about 0.4–0.8. Elevated d

Plant

values occur with rela-

tively low C

ratios that result when stomata are partially

closed. Experimental studies indicate that plants tend to close

their stomata under low values of relative humidity in order

to suppress water loss associated with transpiration, resulting

in elevated d

C values (Farquhar and Richards, 1984). In

922 STABLE ISOTOPE ANALYSIS

addition to humidity, other factors, including light conditions,

nutrient levels, and plant maturity, among others, also influence

the extent of discrimination against

C that occurs during

photosynthesis in C

plants. Despite these complexities, field

studies (e.g., Hemming et al., 1998; Edwards et al., 2000) have

concluded that a strong correlation exists between relative

humidity and d

C values in angiosperm and gymnosperm

trees. This covariance is the basis for the use of carbon-stable

isotope data in paleoclimate research. The carbon-isotopic

composition of tree-rings offer the best substrate for paleocli-

mate research because a number of confounding factors that

affect d

C values in trees are constrained. On the other hand,

bulk organic matter appears to be an unreliable substrate

because the range of d

C values caused by changes in relative

humidity is on the same order of magnitude as that caused by

species-specific effects, diagenesis, and admixture of organic

matter from different sources.

Summary

Temperature estimations from stable isotope data fall into two

categories: (a) those based on fractionation effects related to

carbonate formation from water with known or estimated

O values, and (b) those based on the modem relationship

between the isotopic composition of meteoric water and tem-

perature in temperate and polar regions. The isotopic composi-

tion of meteoric water in these regions can be assessed from

that of a number of substrates including ice, fluid inclusions,

groundwater, pedogenic minerals, bones, and organic matter,

among others.

Assessments of water balance inferred from stable isotope

data can be achieved either through d

O and/or dD at sites

sensitive to changes in precipitation and evaporation (such

as closed-basin lakes) or via changes in C

and C

plant com-

munities at sites sensitive to the colonization of these taxa.

Germán Mora

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Cross-references

Carbon Isotopes, Stable

Coral and Coral Reefs

Dendroclimatology

Deuterium, Deuterium Excess

Foraminifera

Ice Cores, Antarctica and Greenland

Nitrogen Isotopes

Ocean Paleotemperatures

Ostracodes

Oxygen Isotopes

Paleoclimate Proxies, An Introduction

Paleolimnology

Paleoprecipitation Indicators

Paleotemperatures and Proxy Reconstructions

Sea-Surface Temperatures

Strontium Isotopes

Sulfur Isotopes

STRONTIUM ISOTOPES

Strontium isoto pe systematics

The element with atomic number 38, strontium, has four

isotopes 84, 86, 87 and 88 with approximate proportions of

0.56: 9.87: 7.04: 82.53 (Faure, 1986). Geologically short-lived

nuclides (e.g.,

Sr) are not considered in this contribution. The

STRONTIUM ISOTOPES 923