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is preserved in the d

N values of organic matter from algae and

from land plants (Figure O28, bottom) because the limited avail-

ability of nitrogen discourages much isotopic discrimination.

Denitrification occurs during dysoxic conditions in lakes and

oceans and strongly discriminates against

N to yield N

enriched in

N and residual nitrate enriched in

N. These condi-

tions accompany the strong stratification that seasonally happens

in many lakes and is a permanent feature of some oceanic areas.

Subsequent uptake of the residual nitrate by algae leads to

organic matter with high (12–14%) d

N values in the sediment

records of these settings. Conversely, good ventilation of the

water column is recorded as relatively low (3–4%) d

Nvalues.

Nitrogen-fixing organisms such as cyanobacteria are auto-

trophs that convert dissolved N

to bioavailable NH

when

anoxia exists within the photic zone of lakes and oceans. Their

activity leads to lower d

N values in sedimented organic mat-

ter. Evidence of this special environmental condition exists in

sediments deposited between 10 to 3.3 ka in Lake Bosumtwi,

Ghana, where it records a period of strong stratification in this

lake (Talbot and Johannessen, 1992). Suitable conditions for

cyanobacterial activity also commonly occur in eutrophic lakes.

Sediments deposited in central Florida lakes display shifts to

low (1to3%) d

N values as eutrophication of these lakes

increased (Brenner et al., 1999). The remarkably low (4to

0%) d

N values of Cretaceous black shales (Rau et al.,

1987) and of Pliocene-Pleistocene sapropels in the Mediterra-

nean Sea (Struck et al., 2001) imply strong salinity stratifica-

tion of these marine settings and consequent enhancement of

cyanobacterial primary production. As such, these sediment

records indicate periods of wetter paleoclimatic conditions that

were maintained for multi-thousand-year periods.

Sediment d

N values are also sometimes useful in recon-

structing algal paleoproductivity histories, especially in lakes

where phosphorus is often more biolimiting than nitrogen. In

nitrogen-replete environments, biota discriminate in favor of

N, and sediment d

N values are relatively low. During times

of amplified phosphorus delivery, either from human activities

or from climate-enhanced weathering of regolith, increased algal

utilization of bioavailable nitrogen and consequent decreased

discrimination in favor of

N appear as shifts to larger d

values in sediments (Meyers and Lallier-Vergès, 1999).

Biomarker molecules

Biomarker molecules are compounds that characterize biotic

sources and that retain their source information after burial

in sediments (Table O2). Biomarkers in sediments and rocks

are usually diagenetically stabilized derivatives of specific

precursor compounds, yet the carbon skeletons of the derived

molecules preserve the source-distinctive structures of their

precursor compounds. Although biomarkers typically con-

stitute less than 1% of sediment organic matter, they yield

important paleoenvironmental information that bulk organic

geochemical proxies cannot provide.

Hydrocarbons

Because of their low susceptibility to microbial degradation

relative to other types of organic matter, saturated hydrocar-

bons are generally robust recorders of the origins of organic

matter in sediments. The hydrocarbon compositions of many

aquatic algae and photosynthetic bacteria are dominated by

the C

n-alkane (Cranwell et al., 1987). Vascular land plants

contain large proportions of C

, and C

n-alkanes

in their waxy coatings, and the biomarker constitutions are

indicative of different kinds of plants. The n-alkane distribu-

tions of grasses and herbs are dominated by C

, birches by

, and most trees by C

(Schwark et al., 2002).

Fatty acids

Fatty acids in lake sediments typically originate from multiple

sources. The n-C

and n-C

alkanoic and alkenoic acids

are ubiquitous components of biota, whereas the even-chain

–C

n-alkanoic acids originate principally from the waxy

coatings of land plants (Cranwell et al., 1987). The long-chain

n-acid components of organic matter in sediments therefore

represent transport of land-derived debris to the depositional

setting. Unsaturated n-C

and n-C

acids are major constitu-

ents of the lipids of algae, yet they are rapidly degraded by

microbes during and after sedimentation. The microbes contri-

bute their own distinctive biomarker acids to the sediments. For

example, normal and anteiso C

alkanoic acids have been

used as indicators of microbial biomass in lake sediments

(Goossens et al., 1989), and they represent in situ production

of secondary lipids at the expense of primary organic matter.

Figure O28 Generalized elemental and carbon isotopic (top) and

nitrogen and carbon isotopic (bottom) compositions of organic matter

from marine and freshwater algae, C

and C

land plants, and

cyanobacteria.

660 ORGANIC GEOCHEMICAL PROXIES

Unlike the shorter chain-length fatty acids, the longer chain-

length ones survive in sediments and provide evidence of their

plant origins.

Fatty alcohols

Fatty alcohols have source patterns roughly similar to those of

fatty acids. Epicuticular waxes of land plants contain even-

chain n-alkanols from C

to C

(Rieley et al., 1991). How-

ever, unlike fatty acids, n-alkanol distributions dominated by

the C

to C

components generally indicate organic matter

derived from aquatic algae and bacteria and not biota in general

(Volkman et al., 1999). Individual species of plants can have

distinctive chain-length patterns that deviate from these gener-

alized distributions.

Alkenones

Long-chain (n-C

, n-C

) methyl and ethyl alkenones

are distinctive biomarkers of marine coccolithophores. Their

abundance in sedimentary records can be used as a paleopro-

ductivity proxy for these algae, but their special usefulness is

as paleotemperature proxies. Alkenone distributions usually

consist of a series of homologs having a mixture of 1, 2, 3,

and 4 double bonds that is temperature-sensitive. The algae

produce proportionally more polyunsaturated homologs as

water temperatures decreases in order to maintain cell-wall

fluidity. Alkenones also appear in lake sediment records (Zink

et al., 2001), but their utility as a paleoenvironmental proxy

is not yet established because their origin is unknown.

Sterols and their derivatives

Sterol compositions distinguish aquatic contributions of

organic matter in sediments from land-derived material. Cho-

lesterol (C

) is the dominant algal sterol, whereas b-sitosterol

) is the major sterol in emergent water plants (Nishimura

and Koyama, 1977). Land-plant leaf waxes contain C

and

but not C

sterols (Rieley et al., 1991). Ternary plots of

the C

, and C

sterol compositions of algae, vascular

plants, soils, and sediments identify ecological patterns (Huang

and Meinschein, 1979) that can be applied to paleoenviron-

mental interpretations. Specific sterols are particularly useful

to paleoenvironmental reconstructions. Stronger wind-induced

upwelling is indicated from increased amounts of dinosterol

and C

4-methylstanols, which are dinoflagellate bio-

markers, in glacial-age sediments in the equatorial Pacific

(Prahl et al., 1989).

Pigments

Plants have evolved a variety of photosynthetic pigments that

can serve as biomarkers. Chlorophyll a, the most common pig-

ment, is non-specific, but aquatic plants and microbes produce

distinctive secondary pigments (Table O2) that facilitate photo-

synthesis in water. Human-induced eutrophication of Esthwaite

Water in the English Lake District is recorded as progressive

increases in the concentrations of myxoxanthophyll and oscil-

laxanthin, two carotenoids indicative of cyanophytes and hence

conditions of elevated lake productivity (Griffiths, 1978). Climate-

induced strong surface stratification of the Mediterranean

Sea during times of Pliocene sapropel formation is implied by

Table O2 Examples of important biomarker molecules and their significance to organic geochemical reconstructions of paleoenvironmental and

paleoclimatic histories

Biomarker molecules Source/ significance

Hydrocarbons

and C

n-alkanes Vascular land plant waxes; indicates land-plant organic matter contributions and community compositions

and C

n-alkanes Aquatic algae; indicates aquatic organic matter

7- and 8-methyl-branched C

alkanes Cyanobacteria, indicates aquatic nitrogen fixation

Pristane Phytol sidechain of chlorophyll; indicates zooplankton-based food chains

Phytane Phytol sidechain of chlorophyll; methanogens; halophilic bacteria; indicates anoxic conditions

Botryococcane Mat-forming freshwater alga Botryococcus braunii

Fichtelite Conifer resins; indicates reducing depositional conditions

Retene Conifer resins; indicates oxidizing depositional conditions

Fatty acids

, and C

n-alkanoic acids Vascular land plant waxes

Iso and anteiso branched acids Bacteria; evidence of microbial organic matter

Fatty alcohols

, and C

n-alkanols Vascular land plant waxes

Alkenones

, and C

alkenones Marine haptophyte algae; proxies for sea surface paleotemperatures

Sterols and steroids

Dinosterol and derivatives Dinoflagellate-derived organic matter

4-methyl sterols Dinoflagellate-derived organic matter

Brassicasterol Diatom-derived organic matter

sterols Land-plant organic matter

sterols Marine or freshwater algal organic matter

Pigments

Fucoxanthin, diadinoxanthin Diatom-derived organic matter

Peridinin Dinoflagellate-derived organic matter

Myxoxanthophyll, oscillaxanthin Cyanobacteria; indicates aquatic nitrogen fixation

Isorenieratene derivatives Green sulfur bacteria; anoxic conditions in photic zone

Lignin-derived phenols

Cinnamyl/ vanillyl ratio Proportions of woody and non-woody land-plant tissues

Syringyl/ vanillyl ratio Proportions of gymnosperm and angiosperm land-plant organic matter

ORGANIC GEOCHEMICAL PROXIES 661

isorenieratene derivatives in the sapropel layers but not outside

them (Passier et al., 1999). Isorenieratene is a carotenoid that is

diagnostic for green sulfur bacteria, which are photosynthetic

obligate anaerobes, and their presence means that a near-

surface anoxic zone intruded into the photic zone.

Lignin derivatives

Lignins are phenolic polymers that are synthesized by higher

plants to construct parts of their vascular and structural sys-

tems. Nearly all vascular plants grow on land, and therefore

the lignin fraction of sediment organic matter largely records

the contribution and preservation of land-plant materials. Gym-

nosperms and angiosperms synthesize distinctive types of lig-

nin components. Past changes in continental vegetation can

therefore be inferred from the kinds of lignin found in sediment

records. Molecular analysis of lignin typically begins with

an aggressive oxidation step to break down the biopolymer

into various phenolic monomeric fragments. The oxidation

fragments of lignin have been grouped by Hedges and

Mann (1979) as a guide to identifying their plant sources

(Table O2). A measure of the relative contributions of woody

and non-woody land-plant tissues is given by the C/V (cinnamyl/

vanillyl) ratio, which is the sum of p-coumaric acid plus ferulic

acid concentrations divided by the sum of the three vanillyl

phenols. Gymnosperm sources of land-plant residues are distin-

guished from angiosperm sources by the S/V (syringyl/vanillyl)

ratio, which is the sum of the amounts of the three syringyl phenols

divided by the sum of the vanillyl phenols.

Compound-specific isotopic compositions

Isotopic analysis of individual biomarker compounds provides a

powerful source of paleoclimatic and paleoenvironmental infor-

mation. Compound-specific stable isotope analysis improves

identification of the biological origin of specific components

within the complex mixture of materials that constitutes sedi-

ment organic matter.

Compound-specific d

C values

Studies of the d

C values of biomarkers have been especially

important in tracing the evolution of C

plants. A shift to

less negative d

C values of individual plant-wax C

to C

n-alkanes and C

to C

n-alkanols enabled Freeman and Colar-

usso (2001) to identify a dramatic expansion of C

vegetation on

the Indian subcontinent between 8 and 6 Ma that has persisted to

modern times. This paleoecological change reflects uplift of the

T ibeta n Plateau and associated growth of regional aridity.

The distinctive isotopic compositions of biomarkers from C

and C

plants also trace Holocene environmental changes.

Sediments of Lake Baikal, Siberia, exhibit a 3% shift to more

negative bulk organic matter d

C values that might record a

shift from glacial-age C

to post-glacial C

plants. Brincat et al.

(2000) show from d

C values of individual C

to C

n-alkanes in the sediment record that C

plants have remained

the dominant vegetation throughout this period. Instead of

recording a floral transition, the bulk organic carbon isotopic

shift indicates increased fluvial delivery of isotopically light

soil-derived carbon to the lake in response to a change to

locally wetter post-glacial climate.

Applications of multiple biomarker compound-specific car-

bon isotopic measurements are especially fruitful for detailed

paleoenvironmental reconstructions. Filley et al. (2001) utilize

the d

C values of biomarkers from vascular plants (lignin

derivatives, C

and C

n-alkanes, C

n-alkanol), cyanobac-

teria (7- and 8-methylheptadecanes), and phytoplankton (C

n-alkane) to reconstruct the delivery of sedimentary organic

matter in Mud Lake, Florida, over the past five millennia.

Originally a land-plant dominated marsh, the lake evolved into

a cyanobacteria-dominated system as the regional climate

became more humid and the local water table rose.

Compound-specific dD values

Because the hydrogen in most organic matter originates from

water, the D/H ratio in aquatic biomarker compounds is a sen-

sitive paleohydrologic proxy. The dD of lake water can be

reconstructed within 10% from the D/H ratios of sterols

derived from phytoplankton (Sauer et al., 2001). The D/H

ratios of the C

n-alkane that is a biomarker for Sphagnum

were used by Xie et al. (2000) to reconstruct a multi-century

paleoclimatic record for Bolton Fell Moss, Cumbria, England.

The dD values increase during times of warmer climate and

decrease when climate was cooler as rates of evaporative distil-

lation of bog waters increased or decreased. Similarly, Huang

et al. (2002) employed dD variations in n-hexadecanoic acid,

which is produced principally by lake algae, to reconstruct

the post-glacial temperature record preserved in the sediments

of Crooked Pond, Massachusetts.

Summary

Sediment organic matter provides a variety of elemental, iso-

topic, and molecular proxies to reconstruct past climates

and ancient environments. C/N ratios identify the general

origin of organic matter from aquatic or land plants. Bulk

and compound-specific d

C values distinguish proportions

of C

and C

plant organic matter and thereby identify

periods of dry climate. Nitrogen cycling is sensitive to various

environmental conditions that leave distinctive impacts on d

values. Biomarker molecules identify organic matter contribu-

tions from specific sources, which record different environmen-

tal settings, and they provide information about depositional

conditions from their alterations from precursor compounds.

Philip A. Meyers

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N) signatures of sedimented organic

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205–221.

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of C

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in the sediment of a stratified, freshwater lake. Org. Geochem., 14,

27–41.

Griffiths, M., 1978. Specific blue-green algal carotenoids in sediments of

Esthwaite water. Limnol. Oceanogr., 23, 777–784.

662 ORGANIC GEOCHEMICAL PROXIES

Hedges, J.I., and Mann, D.C., 1979. The characterization of plant tissues by

their lignin oxidation products. Geochim. Cosmochim. Acta, 43,

1803–1807.

Huang, W.-Y., and Meinschein, W.G., 1979. Sterols as ecological indica-

tors. Geochim. Cosmochim. Acta, 43, 739–745.

Huang, Y., Shuman, B., Wang, Y., and Webb, T., 2002. Hydrogen isotope

ratios of palmitic acid in lacustrine sediments record late Quaternary cli-

mate variations. Geology, 30, 1103–1106.

Meyers, P.A., and Lallier-Vergès, E., 1999. Lacustrine sedimentary organic

matter records of Late Quaternary paleoclimates. J. Paleolimnol., 21,

345–372.

Meyers, P.A., Leenheer, M.J., and Bourbonniere, R.A., 1995. Diagenesis of

vascular plant organic matter components during burial in lake sedi-

ments. Aquat. Geochem., 1,35–52.

Nishimura, M., and Koyama, T., 1977. The occurrence of stanols in various

living organisms and the behavior of sterols in contemporary sediments.

Geochim. Cosmochim. Acta, 41, 379–385.

Passier, H.F., Bosch, H.J., Nijenhuis, I.A., Lourens, L.J., Böttcher, M.E.,

Leenders, A., Sinninghe Damsté, J.S., de Lange, G.J., and de

Leeuw, J.W., 1999. Sulphidic Mediterranean surface waters during Plio-

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Cross-references

Alkenones

Carbon Isotopes, Stable

Deuterium, Deuterium Excess

Geochemical Proxies (Non-Isotopic)

Nitrogen Isotopes

Paleoceanography

Paleoclimate Proxies, an Introduction

Paleolimnology

Sapropels

OSTRACODES*

Introduction

Ostracodes (also ostracods, Ostracoda) are a class of bivalved,

aquatic Crustacea that secrete a small (0.1 to > 2-mm long) cal-

citic shell (the carapace) that is commonly fossilized. Ostracodes

are used widely in paleoclimatology and the reconstruction of

marine and non-marine paleoenvironments because of their

small size, ecological sensitivity, well-known biology and shell

chemistry, long stratigraphic range (Ordovician-present), and

occurrence in sediments from lakes, estuaries, bays, and oceans.

There are an estimated 33,000 described living and extinct

ostracode species divided into two subclasses, the Myodocopa

and Podocopa, and 10 orders distinguished from one another on

the basis of appendages and carapace features such as size, shape,

muscle scar pattern, pore patterns, and hinge structure. Many

Myodocopids are pelagic in habitat; some are non-calcitic and

are rarely fossilized. The Podocopa includesthe order Podocopida,

(Figures O29 and O30), which are the most commonly fossilized

ostracode group and the most applicable to paleoclimatology.

Morphology and ecology

The carapace consists of two valves articulated along the dorsal

margin, encompasses the living animal, and forms part of the

organism’s cuticle, which is secreted by the epidermis. Ostra-

codes grow by molting (ecdysis), during which podocopids pro-

duce 8–9 molts or instars. Both adult and juvenile carapaces and

valves are commonly fossilized. Most ostracode genera and spe-

cies can be identified on the basis of external and internal cara-

pace features, although examination of soft parts is necessary

to identify species in some groups and recent genetic studies

have added considerably to understanding ostracode taxonomy.

Most species are free-living (either benthonic or pelagic),

move by crawling or swimming, reproduce sexually (a few are

parthenogenic), and disperse passively (there is no planktonic

larval stage). They inhabit a full range of aquatic environments

from small ephemeral pools of water, groundwater, to all types

of lakes, estuaries, bays, rivers, and oceanic environments. Spe-

cies are adapted to specific microhabitat conditions and physical

and chemical factors controlling species distribution include

temperature, salinity, light availability, solute chemistry, alkali-

nity, oxygen and carbon dioxide levels, and food resources.

Application to paleoclimatology and

paleoenvironments

The ecological sensitivity of living species has made ostracodes

useful for paleoenvironmental reconstruction since the nine-

teenth Century, especially in studies of Quaternary sediments

where living species are often found as fossils. Quantitative

analyses of faunal assemblages and indicator species have

been used extensively to reconstruct water solute composition

and regional atmospheric conditions in lakes, salinity and sub-

merged aquatic vegetation in estuaries, ocean temperatures on

continental shelves, dissolved oxygen on continental slopes,

and deep water masses in abyssal environments.

During the past decade, the application of ostracode shell

chemistry to contemporary issues surrounding climate varia-

bility has accelerated rapidly, especially in paleolimnology

OSTRACODES 663

because ostracodes are usually the only calcitic microfossil

group. With greatly improved understanding of the factors that

control the trace element (magnesium/calcium, strontium/

calcium ratios) and stable isotopic (oxygen and carbon) com-

position of the carapace, ostracode researchers have been able

to quantify temporal changes in lake and ocean temperature,

atmospheric temperature and precipitation, estuarine and bay

salinity, and other climate-related parameters (see review

papers in Holmes and Chivas, 2002). For example, in semi-arid

to subhumid, hydrologically-closed lakes, precipitation

influences the salinity, solute chemistry and

O ratios of

the water. Chivas et al. (1993) applied these relationships to

infer salinity and evaporation changes in Lake Keilambete,

Australia during the late Quaternary using a d

O isotopic

record and Sr/Ca ratios for several ostracode species. In deep

lakes, ostracode shell chemistry can provide a record of atmo-

spheric temperature. Von Grafenstein et al. (1999) recon-

structed decadal scale changes in atmospheric temperature

from the d

O record of Candona from the Ammersee,

Germany for the important glacial, deglacial and Holocene

Figure O29 Morphology of Podocopid ostracode Hemicythere villosa (courtesy of David J. Horne).

664 OSTRACODES

interval, and showed a remarkable correspondence to the paleo-

temperature record from Greenland ice cores. Wansard (1996)

also reconstructed summer temperatures from Lake Banyoles,

Spain for the interval 30,000–6,000 years ago using Mg/Ca

ratios in Cyprideis.

Ostracode shell chemistry has also been applied to shallow

water paleoclimate records. In the Gulf of Carpentaria, northern

Australia, DeDeckker et al. (1988) reconstructed the late Qua-

ternary salinity changes caused by changing sea level from

Mg/Ca, Sr/Ca and

Sr/

Sr ratios of Cyprideis. Dwyer et al.

(2001) showed that salinity influences Mg/Ca ratios in Florida

Bay, which allowed them to reconstruct paleosalinity history

controlled by climate processes such as the El Nino-Southern

Oscillation. In Chesapeake Bay, eastern North America, water

temperature controls the Mg/Ca ratios in Loxoconcha and

Cronin et al. (2003) reconstructed a 2,200-year bay temperature

record. In the deep sea, studies by Dwyer et al. (1995, 2000),

Cronin et al. (2000), and Didié and Bauch (2002) provide

evidence from ostraocde shell chemistry for Pliocene and

Quaternary variability in deep-sea temperatures, associating

them with orbital and suborbital control of deep-water formation.

Thomas M. Cronin

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Research. Washington, USA: American Geophysical Union, pp. 279–299.

Dwyer, G.S., and Cronin, T.M., 2001. Ostracode shell chemistry as a paleo-

salinity proxy in Florida Bay. In Wardlaw, B. (ed.), Bulletins of

American Paleontology No. 361, pp. 249–276.

Dwyer, G.S., Cronin, T.M., Baker, P.A., Raymo, M.E., Buzas, J., and

Corrège, T., 1995. Late Pliocene and Quaternary bottom water

temperature change in the deep North Atlantic. Science, 270,

1347–1351.

Dwyer, G.S., Cronin, T.M., Baker, P.A., and Rodriguez-Lazaro, J., 2000.

Changes in North Atlantic deep-sea temperature during climatic fluctua-

tions of the last 25,000 years based on ostracode Mg/Ca ratios.

Geochem. Geophys. Geosyst., 1(12): (doi:10.1029/ 2000GC000046).

Holmes, J.A., and Chivas, A.R. (eds.), 2002. The Ostracoda: Applications

in Quaternary Research. Washington, USA: American Geophysical

Union, 313pp.

Von Grafenstein, U., Erlenkeuser, H., Brauer, A., Jouzel, J., and Johnsen. S.J.,

1999. A mid-European decadal isotope-climate record from 15,500 to

5,000 years

BP. Science, 284,1654–1657.

Wansard, G., 1996. Quantificaton of paleotemperature changes during

isotopic stage 2 in the La Draga continental sequence (NE Spain)

based on the Mg /Ca ratio of freshwater ostracods. Quaternary Sci.

Rev., 15, 237–245.

Cross-references

Geochemical Proxies (non-isotopic)

Oxygen Isotopes

Paleolimnology

Paleotemperatures and Proxy Reconstructions

Strontium Isotopes

OUTWASH PLAINS

Outwash plains are extensive areas of glaciofluvial sediments

deposited in the proglacial zone. They occur in topographically

unconfined locations, such as the margins of ice sheets or pied-

mont glaciers, distinct from the valley trains or valley bottom

deposits that occur in front of valley glaciers in confined settings.

The Icelandic term “Sandur” (plural: Sandar) is often used.

Sediment is deposited primarily by braided meltwater streams

that migrate across the surface of the outwash plain. The whole

plain is rarely covered all at onceexcept during extremejökulhlaup

floods. During floods, previously deposited sediments may be

reworked and redeposited. Outwash deposits can retain sedimen-

tological signatures of long periods of discharge from the glacier,

making it possible to judge the relative significance of ablation

dominated discharge and discharge related to extraordinary events

such as subglacial volcanic eruptions as in Iceland. Outwash

plains that are periodically inundated by floods exhibit distinctive

sediment structures.

Outwash plains can aggrade to thicknesses of several hundred

meters and bury the snout of the glacier from which the meltwater

and sediment are derived. Melting of buried ice can then lead to

the formation of kettles and pitted outwash surfaces, and rapid

retreat of the ice margin can leave a distinctive ice-contact

slope at the head of the outwash plain. Fluctuations in meltwater

discharge or supply of sediment can cause changes in the rate

of aggradation, and incision can occur creating terraces in the

outwash surface.

Close to the glacier, sediments are often coarse grained and

poorly sorted, with sub-angular clasts. Farther from the glacier

sediments are characteristically finer grained and better sorted

with more rounded clasts.

Figure O30 Scanning electron photo of deep sea ostracode Bradleya.

Left valve, 1.0 mm long.

OUTWASH PLAINS 665

Sedimentary structures within the outwash can reveal details

of flow regimes and sedimentary processes,can help to reconstruct

paleodischarges, and can help to unravel the dynamic history of

the glacier from which they are derived Benn and Erans (1998)

provide a useful review of this topic.

Peter G. Knight

Bibliography

Benn, D.I., and Evans, D.J.A., 1998. Glaciers and Glaciation. London,

UK: Arnold, 734pp.

Cross-references

Glacial Geomorphology

Glacial Sediments

Kettles

OXYGEN ISOTOPES

Natural occurrence

Oxygen has three stable isotopes with atomic mass numbers

of 16, 17, and 18 (

O and

O), which occur naturally in

relative proportions of 99.76%, 0.04%, and 0.2%, respectively.

Because

O and

O fractionate (to a first order) proportionally

with respect to

O, and because of the very small abundance of

O, the isotopic ratio

O(1/500) is most commonly

measured in water (e.g., groundwater, ice), carbonate sediments

(CaCO

), oxygen gas (O

), carbon dioxide (CO

), and organic

matter, in which the

O ratio varies by more than 100%.

The

O ratio of a sample is measured by mass spec-

trometry on CO

or O

. The isotopic composition is usually

expressed relative to a standard (which depends on the mea-

sured material) and expressed with the d notation:

O ¼

OÞ

Sample

OÞ

Standard

 1

ð1;000; per milÞð1Þ

Atmospheric fractionation and local temperature

control of th e isotopic composition

Isotopic fractionation occurs during water evaporation and con-

densation (processes 1, 2 and 3 in Figure O31). At equilibrium,

the

O ratio in the condensed phase (liquid or solid) is

higher by approximately 10 per mil as compared to the vapor

phase. In the hydrological cycle, water is mostly fractionated

in the atmosphere during the net transport of vapor between

low to high latitudes, ocean to continents, or low to high eleva-

tions, as a result of successive condensation–evaporation stages

(Dansgaard, 1964). At each stage, the isotopic enrichment of

the condensate depletes the remaining vapor in

O as com-

pared to

O(Figure O32, and process 3 in Figure O31), and

this results in more depleted precipitation at the next con-

densation stage. In this process of distillation, the isotopic

depletion of the vapor is primarily related to its exhaustion

(quantity of remaining vapor compared to the initial quantity).

Multivariate analyses of the

O ratio in modern precipita-

tion have shown that the main parameters controlling this

depletion are the local temperature, the quantity of precipita-

tion, the distance to the coast, and the elevation (Figure O33).

The corresponding effects have been coined “temperature

effect,”“amount effect,”“continental effect,” and “altitude

Figure O31 Range of the isotopic composition of oxygen at the surface, and relationships between the different compositions.

666 OXYGEN ISOTOPES

effect.” This isotopic depletion in the atmosphere determines

the composition of the water reservoirs on the Earth, from

the ocean to the cryosphere. See Mook ( 2001) for a thorough

exposition of the environmental isotopes.

When minerals like carbonate, phosphate, and silica, which

constitute sediments, form in water, their

O ratios are

fractionated with respect to the water ratio, a fractionation

mainly controlled by temperature (Epstein et al., 1953). Hence,

fossil water, mainly ice and groundwater, as well as sedi-

ments, represent potential archives of information regarding

past climates and environments.

Precipitation and the isotopic thermometer

The first primary standard for water was the Standard Mean

Ocean Water (SMOW; Craig, 1961a,b). Because of problems

with its definition, the International Atomic Energy Agency

(IAEA) defined the Vienna Standard Mean Ocean Water

(V-SMOW), with an isotopic composition very close to

SMOW, and prepared in a large amount such that it is still

distributed to laboratories today. Its absolute

O atomic

ratio is 2,005.20 (  0.45)  10

6

(Baertschi, 1976). This com-

position has been chosen to be very close to the oceanic

composition. As such, most precipitation has a much lower

O ratio (Figure O31 and O33), and secondary, more

depleted, standards are used for their measurements. IAEA

now recommends that isotopic measurements on water be

expressed with respect to V-SMOW, and normalized on a

two-standard scale, V-SMOW and SLAP (Standard Ligh t

Antarctic Precipitation), assuming a value d

O ¼55.5%

for S LAP ( Coplen, 1995).

The usual measurement technique is to equilibrate a small

amount of CO

gas with each water sample, and to determine

the isotopic composition of this CO

by mass spectrometry.

This technique requires only a few milliliters of water. The

Figure O32 Isotopic evolution of a vapor mass, modeled as a

Rayleigh-type distillation: an initial vapor mass is exhausted by

successive condensation stages, for instance during an advection to

higher latitude. At each stage, the

O ratio of the condensate is

enriched by 10% compared to the vapor ratio, which depletes the

remaining vapor. F is the fraction of remaining vapor.

Figure O33 d

O – surface temperature relationship for modern precipitation (annual averages) (data from the GNIP network Rozanski et al., 1993),

augmented by some polar data (cf. Figure O34). The linear correlation is calculated for points with an annual temperature <20



C(N = 243 points,

r = 0.987). The insert shows the annual d

O – precipitation relationship (“amount effect”) for stations with annual temperature >20



OXYGEN ISOTOPES 667

sample composition is measured with respect to a reference sam-

ple, which has been previously calibrated against V-SMOW.

Both samples are measured under the same conditions, so that

a correction for the gas–water fractionation does not have to

be applied. This technique usually achieves a precision better

than 0.1%, but with lower accuracy, as shown from interlabora-

tory comparisons organized by IAEA. Other techniques have

been developed, like water hydrolysis and laser spectrometry.

The isotopic composition of precipitation is imprinted, more

or less directly, in many continental records like glacier ice,

snow, groundwaters, tree rings, freshwater carbonates, spe-

leothems, etc. The discussion below focuses on ice, but may

be extended in a similar way to the other archives (see below).

Isotopic measurements on precipitation have been shown

to correlate strongly with surface temperature at mid-to-high

latitudes, and with the amount of precipitation at low latitudes,

the so-called “amount effect” (Figure O33 ; Dansgaard, 1964;

Rozanski et al., 1993). The isotope-temperature relationship

can be described with the vapor exhaustion, as shown by a sim-

ple distillation model (batch distillation, also called “Rayleigh

distillation”: Figure O32; Dansgaard, 1964). The modeling of

this relationship can be improved by considering other second-

ary parameters, like the proportion of different water phases in

the clouds (vapor, liquid, and solid water), evaporation of

the falling droplets (especially for the “amount effect”); and

kinetic fractionation during the formation of snow (Jouzel

and Merlivat, 1984).

The observed correlation between d

O and temperature

(Figure O33) is strongest for annual means, and in polar

regions (Figure O34). It is based on different stations, and thus

a spatial gradient Dd

O/DT is defined, where D is a variation

over space. This gradient is traditionally used to interpret isoto-

pic variations (e.g., in the ice) in terms of past temperature, and

has been called the “isotopic paleothermometer” (Figure O34).

Yet, a temporal gradient Dd

O/DT must be used to interpret

past isotopic variations, where D is a variation, at one location,

over the timescale of the considered climatic variations. In recent

years, estimates of temperature variations in Central Greenland

have been achieved independently of d

O. They have shown

that, over the last glacial period, the isotopic paleothermometer

underestimates past temperature variations by a factor of two;

that is, the temporal gradient is about half the spatial one (Cuffey

et al., 1995; Jouzel, 1999). So far, there is no evidence for such

a strong bias in East Antarctica records (Jouzel et al., 2003).

The difficulty in interpreting past d

O variations thus consists

in determining the relevant temporal d

O–temperature gradient,

because a number of climatic parameters can cause this gradient

to differ from the spatial gradient. Some of these parameters

are briefly described below; a thorough review can be found in

Jouzel et al. (1997).

The origin of the temperature-isotope relationship in preci-

pitation has to be determined for each case. The d

Oof

the condensate is determined by both the vapor d

O and the

local temperature at the condensation level, T

. Over the year,

temperatures at the surface (T

) and at the condensation level

) may well correlate during precipitation events, parti-

cularly in polar regions. For Antarctica, a linear relationship

¼ 0.67  T

– 1.2 has been obse rved (J ouz el and Merl ivat,

1984). This justifies the use of the surface temperature T

rather than T

, in t he paleothermometer, and alternatively

may help explain some bias of this thermometer if the correla-

tion between Ts and Tc has not held through climatic changes.

The d

O of the vapor, before condensation, mainly depends

on the fraction of remaining vapor (F in Figure O32). To a

first order approximation, as shown by simple modeling,

this fraction is somehow determined by the difference between

the temperature of the remote vapor source and the tempera-

ture at the location of precipitation. This is so because the

temperature governs the quantity of vapor the air can hold

(specific humidity). Thus, interpreting a d

O record in terms

of local temperature changes only requires that the remote

(source) temperature has remained constant in time. For

instance, Aristarain et al. (1986) analyzed snow from the

Antarctic Peninsula and showed that, given that the precipita-

tion originates in the Southern Ocean, a covariation of the local

and oceanic temperatures can explain a temporal gradient

Figure O34 Isotopic paleothermometer for polar stations: data from Dahe et al. (1994) for Antarctica, and Johnsen et al. (1989) for Greenland.

668 OXYGEN ISOTOPES

O/DT lower by a factor of two than that of the observed

spatial slope. A way to check the temperature stability of the

precipitation source is provided by the deuterium-excess, a

second-order isotopic parameter measured in precipitation

(Jouzel and Merlivat, 1984). To a first order, the isotopic

composition of hydrogen in the water (dD, see Deuterium,

deuterium excess) is linearly related to the composition of

oxygen, d

O, because both vary similarly with temperature;

thus, combining d

O and dD gives access to other climatic

parameters.

It is also important to consider that each isotopic measure

from an archive results from the accumulation of many pre-

cipitation events, and is thus a weighted average, usually

over several months or years. If the accumulation is not evenly

distributed over the year, the temperature record may be

biased towards one particular season. For Central Greenland,

modeling studies have shown that the contribution of the warm

season to the yearly accumulation may have been more impor-

tant during glacial periods than today (Krinner et al., 1997).

This would bias the glacial d

O signal towards warmer

temperature, and may explain why the temporal gradient

O/DT is lower than the spatial gradient (cf. supra).

At low latitudes, the d

O depends less on temperature

and more on the quantity of precipitation (Figure O33), which

is called the “amount effect.” The isotopic signal may be

interpreted in terms of hydrologic changes, especially in the

monsoon area.

The isotopic record of carbonate sediments

Carbonate isotopic fractionation and the

temperature scale

Carbonate is a widely distributed sediment, usually of biogenic

origin, and for this reason it is also the most frequently used to

retrieve a climatic signal routinely from the

O composition.

The climatic interpretation of the carbonate d

O signal is based

on the pioneering work done by Harold C. Urey, Samuel

Epstein, Cesare Emiliani, and their colleagues at the University

of Chicago. They developed the concept of paleothermometry,

and defined the first isotopic reference, a belemnite from the

Peedee formation in South Carolina, called PDB. The isotopic

ratio is measured with a mass spectrometer directly on the

degassed by reaction with pure H

(McCrea, 1950).

For PDB, this isotopic ratio is 2,067.2  10

6

(Figure O31 ).

From measurements of the isotopic fractionation between water

and carbonate, they demonstrated the potential of fossil carbo-

nate to record environmental parameters (Urey, 1947 ; McCrea,

1950), especially temperature: newly formed carbonate records

the isotopic composition of the water, accounting for an isoto-

pic fractionation that is temperature-dependent. Epstein et al.

(1953) calibrated a temperature equation on living mollusks,

based on the isotopic composition of the carbonate (d

) and

of the water (d

t ¼ 16:5  4:3 ðd

 d

Þþ0:14 ðd

 d

ð2Þ

with t in



C and the d in per mil with respect to VPDB. Apply-

ing this equation to a long marine record, Emiliani (1955 )

pioneered the use of d

O in paleoceanography. Due to a short-

age in PDB, IAEA defined the Vienna Peedee Belemnite

(V-PDB) reference with respect to the standard of the National

Bureau of Standard NBS-19 as: d

NBS19 / VPDB

¼2.2%.

The reference material currently distributed to laboratories is

NBS-19.

Isotopic records and stratigraphy for the last

million years

The paleotemperature equation (Equation 2) is applied to carbo-

nate shells secreted by planktonic or benthic micro-organisms

like foraminifera, coccolithophoridae, which are found in mar-

ine sediments retrieved by deep sea coring. It is also applied to

other organic carbonate found in marine as well as freshwater

environments, like corals, mollusks, etc. (Figure O31, pro-

cesses 4 and 5), and to inorganic carbonates like speleothems.

Comparison between inorganically and organically precipitated

calcite has shown a fair agreement for the isotopic fractiona-

tion of oxygen, suggesting that, to a first order, organic carbo-

nate precipitates close to equilibrium with the surrounding

water. Yet, the temperature dependence of this fractionation

(Equation 2) has been recalibrated several times on different

species to account for small departures from equilibrium

(Wefer and Berger, 1991; Bemis et al., 1998). This temperature

dependence is approximately 0.25 per mil of d

per



C, or 4



per mil of d

. Equation 2 also shows that, for a constant tempera-

ture, a 1 per mil variation of d

is equivalent to a 4



C change.

Over the last million years, which have been dominated by

glacial cycles, both temperature and sea water composition

have changed, with similar effects on the isotopic record.

The marine glacial cooling of a few degrees has increased

fractionation by around 1 per mil, while stocking up water

in wide ice sheets has also enriched d

(and thus d

)by

1 per mil. This amplification of global climatic changes

explains the similarity of isotopic records retrieved from dif-

ferent oceanic regions, which underlies the so-called oxygen

isotope stratigraphy: periodic d

O up-and-downs are matched

between records and numbered, with odd and even num-

bers corresponding to warm and cold periods, respectively

(Figure O35). The SPECMAP group (Imbrie et al., 1984)

compiled a global d

O signal by retaining the most signifi-

cant excursions of different records. The final objective of

SPECMAP was to produce a common temporal framework

in order to date other records. The absolute dating of this fra-

mework has been a difficult task. Magnetic reversals detected

in the sediments are widely used, the youngest one being the

Brunhes-Matuyama about 780,000 years ago. Sea level mar-

kers like coral terraces are also used, since sea level variations

are imprinted in the marine isotopic signal. Spectral analysis

of the isotopic records have proved their origin in the varia-

tions of the Earth orbit (Hays et al., 1976), confirming the

astronomical theory of glacial cycles. A dating technique

derived from this confirmation, called “orbital tuning,” con-

sists in tuning the age of a record to the orbital configura-

tions, assuming some relationships between them.

The marine d

O signal: temperature and ice volume

The marine carbonates record both variations of water tem-

perature and of water d

O. The latter is conceptually divided

into a global and a local component. The global component

affects the quantity of

O of the ocean as a whole, on a

timescale longer than the ocean mixing time so that it is

expected to be homogeneous. This is due to fluctuations of

the ice sheet volume, which extract or release isotopically

depleted water from or to the oceans (Figure O31). On time-

scales longer than hundred million years, an imbalance in the

O oceanic budget is also possible (Lécuyer and Allemand,

1999). The local component affects the spatial distribution

O. For the surface, this is mainly due to hydrological

OXYGEN ISOTOPES 669