Odekon M. Encyclopedia of paleoclimatology and ancient environments

Подождите немного. Документ загружается.

Berger, W.H., Smetacek, V.S., and Wefer, G., 1989. Ocean productivity and

paleoproductivity – an overview. In Berger, W., Smetacek, V.S., and

Wefer, G. (eds.), Productivity in the Oceans: Present and Past. New

York: Wiley, pp. 1–34.

Bishop, J.K.B., 1988. The barite-opal-organic carbon association in oceanic

particulate matter. Nature, 332, 341–343.

Bonn, W.J., Gingele, F.X., Grobe, H., Mackensen, A., and Futterer, D.K.,

1998. Palaeoproductivity at the Antarctic continental margin: Opal

and barium records for the last 400 ka. Palaeogeogr. Palaeoclimatol.

Palaeoecol., 139, 195–211.

Calvert, S.E., 1966. Accumulation of diatomaceous silica in the sediment

of the Gulf of California. Geol. Soc. Am. Bull., 77, 569–572.

Curry, W.B., and Crowley, T.J., 1987. The d

C of equatorial Atlantic

surface waters:Implications for ice age pCO

levels. Paleoceanography,

2, 489–517.

Dehairs, F., Chesselet, R., and Jedwab, J., 1980. Discrete suspended

particles of barite and the barium cycle in the open ocean. Earth Planet.

Sci. Lett., 49(2), 528–550.

Delaney, M.L., 1998. Phosphorous accumulation in marine sediments

and the oceanic phosphorous cycle. Glob. Biogeochem. Cycles, 12,

563–572.

De La Rocha, C., Brzezinski, M., DeNiro, M., and Shemesh, A., 1998.

Silicon-isotope composition of diatoms as an indicator of past oceanic

change. Nature, 395, 680–683.

DeMaster, D.J., 1981. The supply and accumulation of silica in the marine

environment. Geochim. Cosmochim. Acta, 45, 1715–1732.

Deuser, W.G., Ross, E.H., and Anderson, R.F., 1981. Seasonality in the

supply of sediments to the deep Sargasso Sea and implications for

the rapid transfer of matter to the deep ocean. Deep-Sea Res., 28A,

495–505.

Duplessy, J.C., Shackleton, N.J., Matthews, R.K., Prell, W.L., Ruddiman,

W.F., Caralp, M., and Hendy, C., 1984.

C record of benthic

foraminifera in the last interglacial ocean: Implications for the carbon

cycle and the global deep water circulation. Quaternary Res., 21,

225–243.

Dymond, J., Suess, E., and Lyle, M., 1992. Barium in deep-sea sediment:

A geochemical proxy for paleoproductivity. Paleoceanography, 7(2),

163–181.

Eagle, M., Paytan, A., Arrigo, K., van Dijken, G., and Murray, R., 2003.

A comparison between excess barium and barite as indicators of carbon

export. Paleoceanography, 18(1), 1021.

Elderfield, H., and Rickaby, R., 2000. Oceanic Cd/ P ratio and nutrient

utilization inthe glacial Southern Ocean. Nature, 405, 305–310.

Emerson, S., 1985. Organic carbon preservation in marine sediments.

In Sundquist, E.T., and Broecker, W.S. (eds.), The Carbon Cycle and

Atmospheric CO

: Natural Variations Archean to Present. Washington,

DC: American Geophysical Union. pp. 78–87.

Farrell, J.W., Pedersen, T.F., Calvert, S.E., and Nielsen, B., 1995. Glacial-

interglacial changes in nutrient utilization in the equatorial pacific

ocean. Nature, 377, pp. 514–516.

Filippelli, G.M., 2002. The global phosphorus cycle. In Kohn, M.J.,

Rakovan, J., and Hughes, J.M. (eds.), Reviews in Mineralogy and

Geochemistry, Vol. 48. Blacksburg, VA: Mineralogical Society of

America, pp. 391–425.

Finney, B.P., Lyle, M.W., and Heath, G.R., 1988. Sedimentation at

MANOP Site H (eastern equatorial Pacific) over the past 400,000 years:

Climatically induced redox variations and their effect on transition

metal cycling. Paleoceanography, 3, 169–189.

Ganeshram, R., Francois, R., Commeau, J., and Brown-Leger, S., 2003.

An experimental investigation of barite formation in seawater.

Geochemica. Cosmochimica Acta, 67(14), 2599–2605.

Goñi, M.A., Ruttenberg, K.C., and Eglinton, T.I., 1998. A reassessment

of the sources and importance of land-derived organic matter in surface

sediments from the Gulf of Mexico. Geochimica et Cosmochimica

Acta, 62, 3055–3075.

Herguera, J.C., 2000. Faunal perspectives on paleoproductivity. Mar.

Micropaleontol., 40, 131–134.

Hutchins, D.A., and Bruland, K.W., 1998. Iron limited diatom growth

and Si:N uptake ratios in a coastal upwelling regime. Nature, 393,

561–564.

Imbrie, J., and Kipp, N.G., 1971. A new micropaleontological method

for quantitative paleoclimatology: Application to a Late Pleistocene

Caribbean core. In Turekian, K. (ed.), The Late Cenozoic Glacial Ages.

New Haven, CT: Yale University Press, pp. 71–181.

Keil, R.G., Tsamakis, E., Fuh, C.B., Giddings, C.A., and Hedges, J., 1994.

Mineralogical and textural controls on the organic composition

of coastal marine sediments: Hydrodynamic separation using SPLITT-

fractionation. Geochimica et Cosmochimica Acta, 58, 879–893.

Kumar, N., Anderson, R., Mortlock, R., Froelich, P., Kubik, P., Dittrich-

Hannen, B., and Suter, M., 1995. Increased biological productivity

and export production in the glacial Southern Ocean. Nature, 378,

675–680.

Loubere, P., 1999. A multiproxy reconstruction of biological productivity

and oceanography in the eastern equatorial Pacific for the past 30,000

years. Mar. Micropaleontol., 37, 173–198.

Lyle, M., Murray, D., Finney, B., Dymond, J., Robbins, J., and

Brooksforce, K., 1988. The record of Late Pleistocene biogenic

sedimentation in the eastern tropical Pacific Ocean. Paleoceanography,

3,39–59.

Mekik, A.F., Loubere, P., and Archer, D., 2002. Organic carbon flux and

organic carbon to calcite flux ratio Recorded in deep sea carbonates:

Demonstration and a new proxy. Glob. Biogeochem. Cycles, 16, 1052,

doi:10.1029/ 2001GB001634, 2002.

Mix, A.C., 1989. Pleistocene paleoproductivity: Evidence from organic

carbon and foraminiferal species. In Berger, W.H. et al., (eds),

Productivity of the Ocean, Past and Present. New York: Wiley,

pp. 313–340.

Montoya, J.P., 1994. Nitrogen isotope fractionation in the modern ocean:

Implications for the sedimentary record. In Zahn, R., Pedersen, T.F.,

Kaminski, M.A., and Labeyrie L. (eds.), Carbon Cycling in the Glacial

Ocean: Constraints on the Ocean’s Role in Global Change. NATO ASI

Series, vol. 17, Berlin: Springer. pp. 259–279.

Muller, P.J., and Suess, E., 1979. Productivity, sedimentation rate, and

sedimentary organic matter in the oceans. I. Organic carbon preser-

vation. Deep Sea Res., 26, 1347–1362.

Murray, R.W., and Leinen, M., 1996. Scavenged excess Al and its relation-

ship to bulk Ti in biogenic sediment from the central equatorial Pacific

Ocean. Geochimica et Cosmochimica Acta, 60, 3869–3878.

Murray, R., Knowlton, C., Leinen, M., Mix, A., and Polsky, C., 2000.

Export production and carbonate dissolution in the central equa-

torial Pacific Ocean over the past 1 Myr. Paleoceanography, 15(6),

570–592.

Nürnberg, C.C., Bohrmann, G., and Schlüter, M., 1997. Barium accumula-

tion in the Atlantic sector of the Southern Ocean: Results from 190,000-

year record. Paleoceanography, 12, 594–603.

Paytan, A., Kastner, M., and Chavez, F., 1996. Glacial to interglacial

fluctuations in productivity in the equatorial Pacific as indicated by

marine barite. Science, 274, 1355–1357.

Pedersen, T.F., 1983. Increased productivity in the eastern equatorial

Pacific during the last glacial maximum (19,000 to 14,000 yr

B.P.).

Geology, 11,16–19.

Rowe, G.T., 1983. Biomass and production of the deep-sea macrobenthos.

In Rowe, G.T. (ed.), Deep-sea Biology, The Sea, vol. 8. New York:

Wiley, pp. 97–121.

Ruhlemann, C., Frank, M., Hale, W., Mangini, A., Mulitza, S., Muller, P.J.,

and Wefer, G., 1996. Late Quaternary productivity changes in the

western equatorial Atlantic: Evidence from

230

Th-normalized carbonate

and organic carbon accumulation rates. Mar. Geol., 135, 127–152.

Rutsch, H., Mangini, A., Bonai, G., Dittrich-Hanen, B., Kubik, P., Suter,

M., and Segl, M. 1995.

Be and Ba concentrations in West African

sediments trace productivity in the past. Earth Planet. Sci. Lett., 133,

129–143.

Ruttenberg, K.C., and Berner, A.B., 1993. Authigenic apatite formation

and burial in sediments from non-upwelling, continental margin envir-

onments. Geochim. Cosmochim. Acta, 57, 991–1007.

Sarnthein, M., Winn, K., Duplessy, J-C., and Fontugne, M.R., 1988. Global

variations of surface ocean productivity in low and mid latitudes:

Influence on CO

reservoirs of the deep ocean and the atmosphere

during the last 21,000 years. Paleoceanography, 3, 361–399.

Schneider, R.R., Müller, P.J., and Wefer, G., 1994. Late Quaternary paleo-

productivity changes off the Congo deduced from stable carbon

isotopes of planktonic foraminifera. Palaeogeogr. Palaeoclimatol.

Palaeoecol., 110, 255–274.

Schneider, R.R., Price, B., Muller, P.J., Kroon, D., and Alexander, I., 1997.

Monsoon related variations in Zaire (Congo) sediment load and

influence of fluvial silicate supply on marine productivity in the east

equatorial Atlantic during the last 200,000 years. Paleoceanography,

12, 463–481.

650 OCEAN PALEOPRODUCTIVITY

Shackleton, N.J., Hall, M.A., and Shuxi, C., 1983. Carbon isotope data

in core V19–30 confirm reduced carbon dioxide concentrations in

the ice age atmosphere. Nature, 306, 319–322.

Shackleton, N.J., and Pisias, N.G., 1985. Atmospheric carbon dioxide,

orbital forcing, and climate, in The Carbon Cycle and Atmospheric

: Natural Variations Archean to Present, Geophys. Monogr. Ser.,

32, edited by E. Sundquist and W.S. Broecker, pp. 303–317, AGU,

Washington, D.C.

Stoll, H.M., and Schrag, D.P., 2000. Coccolith Sr/ Ca as a new indicator

of coccolithophorid calcification and growth rate. Geochim. Geophys.

Geosys., 1, 1999GC000015.

Tappan, H., 1968. Primary production, isotopes, extinctions and the

atmosphere. Paleogeogr. Paleoclimatol. Paleoecol., 4, 187–210.

Treuger, P., Nelson, D., van Bennekom, A., Demaster, D., Lanaert, A.,

and Queguiner, B., 1995. The silica balance in the world ocean:

A re-estimate. Science, 268, 375–379.

Walter, H.J., Rutgers, M.M., van der Loeff and Francois, R., 1999. Reli-

ability of

231

Pa/

230

Th activity ratio as a tracer for bioproductivity of

the ocean. In Fischer, G., and Wefer, G. (eds.), Use of Proxies

in Paleoceanography: Examples from the South Atlantic. Berlin:

Springer, pp. 393–408.

Cross-references

Animal Proxies, Invertebrates

Carbon Cycle

Carbon Dioxide, Dissolved (Ocean)

Carbon Isotopes, Stable

Carbonate Compensation Depth

Coccoliths

Diatoms

Dinoflagellates

Foraminifera

Geochemical Proxies (Non-isotopic)

Iron and Climate Change

Marine Biogenic Sediments

Marine Carbon Geochemistry

Organic Geochemical Proxies

Paleoclimate Proxies, an Introduction

Paleoceanography

Phosphorus Cycle

Stable Isotope Analysis

Transfer Functions

OCEAN PALEOTEMPERATURES

Ocean paleotemperatures are the principal characteristics of past

ocean scenarios. Over geological time, ocean temperature pat-

terns underwent large changes, responding on long timescales

( 10

yr) to different plate tectonical configurations, favoring

either circum-equatorial flow and promoting a generally warmer

ocean, or circum-polar flow with steep low-to-high latitude tem-

perature gradients. On shorter timescales ( 10

–10

yr), ocean

temperatures respond to insolation changes controlled by Earth

orbital parameters which determine the amount of radiation the

Earth’s surface receives. High-amplitude, short-term temperature

fluctuations are superimposed on decadal-to-millennial time

scales and match the short-term temperature fluctuations

recorded in Greenland ice cores during glacial times. This is a

recent discovery that has immediately focused interest in climate

research because of the potential analogy to human-induced cli-

mate change.

Because ocean temperatures are a key component of the cli-

mate system, deciphering ocean paleotemperatures from “mar-

ine climate archives” reflecting the climatic state of the Earth

continues to be a longstanding aim. Sea surface temperatures

(SSTs) exert an immediate impact on the atmosphere, driving

atmospheric circulation in terms of heat and moisture transport,

and oceanic-atmospheric gas exchange. Moreover, aside from

salinity, temperature determines the density of surface water,

which drives thermohaline deepwater circulation. Thus, paleo-

SSTs allow conclusions to be drawn about oceanic and atmo-

spheric circulation. Deepwater temperatures depend on the loci

where they are formed by sinking of dense surface waters;

paleo-deepwater temperatures reflect the state of thermohaline

circulation.

Different concepts are used to reconstruct water tempera-

tures of past times, including micropaleontological, stable

isotope, biomarker, and geochemical approaches. Several

micropaleontological and geochemical data sets have been cali-

brated with the aim of yielding quantitative paleotemperature

proxies. A variety of studies of past ocean scenarios (e.g., of

the Last Glacial Maximum, Cenozoic warm climates), based

on paleotemperature proxies, has greatly improved our under-

standing of the climate system over the last few decades. In

particular, ocean paleotemperatures may serve as boundary

conditions to initiate and validate climate model calculations

that aim to predict the magnitude of future global warming.

History

Geological and paleontological studies aiming to reconstruct

Earth’s climate are perhaps as old as these disciplines. The old-

est paleoclimate records based on marine fossils from land out-

crops date back to the late eighteenth and early nineteenth

century, when Cretaceous boreal, tropical, and austral pro-

vinces were identified in mollusk, ammonite and other faunas

and changes in biogeographical distribution patterns were inter-

preted in terms of climatic change (e.g., Hutton, 1795; Cuvier,

1817).

The earliest climate records aiming for quantitative tempera-

ture reconstructions from microfossils were based on the iden-

tification of key species in Pleistocene marine sediments;

species that could be related to climate zones (e.g., Phleger

et al., 1953). However, it was not until the second half of the

twentieth Century that an array of studies on the biogeographi-

cal distribution of modern planktonic foraminifera in the world

ocean and in marine sediments became available, the potential

of this species group as a tool for climate studies was discov-

ered (e.g., Bradshaw, 1959; Bé, 1977; and others), and a con-

ceptual framework for quantitative faunal approaches was

found (Imbrie and Kipp, 1971).

Stable isotopes began to flourish as a tool in paleoceanogra-

phy in the mid-twentieth century. Urey (1947) first explored the

potential of oxygen isotope fractionation in natural carbonates

as a paleotemperature indicator. Epstein et al. (1953) estab-

lished the first paleotemperature equation based on oxygen iso-

topes. They found a 0.2% decrease in oxygen isotopic

composition (expressed as d

O) per 1



C temperature increase

for mollusk calcite. Emiliani (1955) published the first “paleo-

temperature records” based on d

O variations of foraminifera

in Caribbean sediment cores, which he interpreted purely in

terms of temperature fluctuations of 6–8



C. However, it was

yet not fully recognized to what extent the oxygen isotopic

composition of the ocean changes through time. Shackleton

(1967) discovered d

O fluctuations of similar amplitude in

benthic foraminifera of the deep Pacific where deepwater is con-

tinuously close to freezing temperature. Here it became evident

that the major portion of the d

O fluctuations observed in deep

sea sediments are a result of waxing and waning continental ice

OCEAN PALEOTEMPERATURES 651

volumes where the light

O isotope is alternately stored away

and released into the ocean. This finding has converted d

O into

a stratigraphical tool, and when combined with independent tem-

perature proxies, into a salinity and ice volume proxy.

Since the mid-1980s, alkenones have become established as

a sea surface temperature (SST) proxy. The use of this proxy is

based on the finding that these lipids, which are produced by

marine algae, show a characteristic unsaturation pattern that is

linked to growth temperature (Marlowe et al., 1984; Brassell

et al., 1986). Over the last 20 years, analytical methods have

improved and data sets for calibration that cover wide ocean

regions have been established for water column, laboratory cul-

tures, and sediment surface samples. These eventually gave rise

to a new proxy that is now widely applied in paleoceanography

(Müller et al., 1998).

Knowledge that magnesium contents in carbonates vary

with temperature dates back to the early part of the twentieth

century (Clarke and Wheeler, 1922). Nevertheless, only the

development of precise analytical methods in the last decade

of the past century enabled utilization of this trace element in

carbonate microfossils as a quantitative proxy for paleotem-

perature (Nürnberg et al., 1996).

Ongoing work concentrates on: (a) achieving a still more pre-

cise calibration of established paleotemperature proxies, in parti-

cular for deepwater temperature reconstructions and near the

“cold end” of temperature transfer functions, (b) exploring the

use of quantitative proxies further back in geological time, and

ing have focused interest on the temperature reconstruction of

former warm periods; (d) Multi-proxy approaches enable com-

parison of different reconstructions and, moreover, help to iden-

tify the influence of different seasons and growth habitats on the

SST signal. In this context, reef corals are particularly promising

in providing SST records with annual, seasonal, and even

monthly time resolution and thus have gained increasing impor-

tance in addressing questions of short-term climate change.

Concepts for paleotemperature reconstructions

Most concepts for reconstructing past water temperatures are

based on microfossils because of their widespread occurrence

in marine sediments. Sea surface waters are inhabited by plank-

tonic microorganisms, some of them bearing shells. In general,

marine phyto- and zooplankton occur everywhere in the open

ocean. Temperatures range from less than –1.8



C in the polar

ice-covered areas up to 30



C in the tropical belt.

As the shells settle to the sea floor and are fossilized in mar-

ine sediments they preserve information about the environment

in which they once lived. Paleoceanographers take advantage

of this ‘climate archive’ to decipher paleotemperatures from

microfossils, either from their faunal (or floral) assemblages

or by analyzing the stable isotope or trace element composition

of the calcareous shells. Most commonly employed microfossil

groups are planktonic foraminifera. In addition, coccoliths,

dinoflagellate cysts, and siliceous microfossils (radiolaria and

diatoms) are used, the latter in particular to reconstruct Southern

Ocean environments, where carbonate preservation is poor.

In contrast, deep-water temperatures are far more uniform

and are generally near the cold end of transfer functions, ran-

ging from close to freezing temperature to approximately



C. This makes them more difficult to assess. Deep-sea

benthic foraminifera and ostracoda (little crustaceans bearing

a calcareous carapace) living at the seafloor similarly record

deepwater temperatures in the isotopic and geochemical com-

position of their shells. Benthic assemblage structures, how-

ever, are much more complex than those of planktonic

foraminifera and are primarily controlled by variables different

from temperature.

Transfer functions based on microfossil species

assemblages

Planktonic species assemblages are primarily sensitive to

temperature, most species showing preference for a certain tem-

perature range to which their vital functions are optimally

adjusted. Therefore, planktonic assemblages display an ocean-

wide, bipolar, roughly latitudinal distribution pattern, reflecting

the occurrence of tropical, subtropical, transitional, subpolar,

and polar water masses (Figure O23). Plankton distribution pat-

terns of former times may thus be used to determine SSTs by

using so-called transfer functions if evolutionary changes did

not affect these patterns – that is, approximately over the last

0.5 million years. Imbrie and Kipp (1971) were the first to

apply multivariate statistical methods to census data of plank-

tonic foraminifera in order to derive quantitative SST estimates.

Their method commonly consists of subjecting relative species

abundances counted in core tops from a specific ocean region

to a principal component (Q mode) analysis in order to discri-

minate factors by correlation between single species. The fac-

tors represent statistical assemblages, characterized by an end

member species, the distribution pattern of which is related to

specific watermass qualities (Figure O23).

Usually 5–6 factors are sufficient to explain > 90% of total

variance in the assemblages. The role of single species grouped

in a factor is weighted by factor loadings. Regression analysis

(linear or non-linear) of the factors with instrumental tempera-

tures above each station provides a set of equations (transfer

functions) which can be used to calculate paleotemperatures

from fossil species abundances. The equations have the shape

of Equation (1)

est

¼ k

þ k

A þ k

B þ k

C þ k

D ð1Þ

Where P is an environmental parameter, ks are predictive coef-

ficients, and A–D are variables. In example (2), by Imbrie and

Kipp (1971), T

is the summer temperature estimate, and the

letters A–D stand for the values of tropical (Figure O24), sub-

tropical, subpolar, and gyre margin assemblages.

¼ 19:7A þ 11:6B þ 2:7C þ 0:3D þ 7:6 ð2Þ

The quality of the calculated temperatures may be assessed by

the measure of communality. This statistical measure, which is

defined as the squared sum of factor loadings calculated for the

fossil assemblage, shows how well the fossil assemblage fits

into the pattern of modern assemblages.

The most impressive example of temperature reconstruc-

tions based on this approach was, without doubt, a global

reconstruction of sea surface temperatures for the Last Glacial

Maximum (LGM), accomplished by the CLIMAP Project

members (1981) in 1976–1981. They applied the “Imbrie and

Kipp (I&K) method” to a worldwide set of 700 glacial sedi-

ment samples where they used planktonic foraminifera, radi-

olaria, and coccolith assemblages as a joint base to produce

SST maps of the glacial ocean.

A more advanced approach to derive paleotemperatures

from faunal assemblages is the Modern Analog Technique

652 OCEAN PALEOTEMPERATURES

(MAT) (Hutson, 1977), relying on the assumption that there is

a direct relationship between the distribution of modern faunas

and the physical properties of the environment. The method

compares any fossil assemblage of planktonic foraminifera or

other plankton with a set of census data from modern sedi-

ments in an ocean region (calibration data set) and transfers

the measured temperatures above the most similar modern sam-

ples (= modern analogs) to the site of the fossil fauna. This

approach employs a measure of faunal dissimilarity or similar-

ity to identify the best modern analogs in the calibration data

set. The most commonly used dissimilarity coefficient is the

squared chord distance (Equation (3)).

dij ¼ Sigma kðpik 1=2-pjk1=2Þ2 ð3Þ

where dij is the squared chord distance between two multivari-

ate samples i and j, and pik is the proportion of species k in

a sample. Commonly, 10 modern analogs are selected. Mea-

sured sea surface temperatures at the various sample sites

selected are averaged to give the paleotemperature. Further

Figure O23 Distribution of core top samples dominated by statistical assemblages defined by Imbrie and Kipp (1971).

Figure O24 Geographical distribution of the tropical planktonic foraminifera group as defined by the multivariate statistical method. Four

dominant species are shown with factor loadings ( 10) (from Imbrie and Kipp, 1971).

OCEAN PALEOTEMPERATURES 653

refined variants of the MAT include SIMMAX (“similarity

maximum”; Pflaumann et al., 1996, 2003), which weights the

modern analog samples according to their geographical dis-

tance to the fossil sample, presuming that closer samples are

more likely to record accurate temperatures. The Revised Ana-

log Technique (RAM, Waelbroeck et al., 1998) pays tribute to

the circumstance that core top samples as modern analogs are

not evenly distributed across the ocean and therefore not all

temperature settings are equally represented. Therefore, RAM

does not use a pre-given number of analogs but defines for

each sample the number to include, according to the dissimilar-

ity. Moreover, artificial (interpolated) modern analogs are intro-

duced into the database to strengthen the weight of poorly

sampled temperature settings.

The use of Artificial Neural Networks (ANN; Malmgren

et al., 2001) presents a more recent innovative approach to

employ planktonic faunal data for SST reconstruction. The

ANN technique is a computer-intensive method, which is based

on an algorithm that has the ability to autonomously “learn” of

a relationship between two groups of numbers (e.g., faunal and

temperature data). The advantage of ANNs is that they have the

ability to overcome problems of fuzzy and nonlinear relation-

ships between sets of input and output variables. Each trained

neural network serves as a unique transfer function, yet at

the same time, this highly nonlinear and recurrent function is

so complicated that it has the ability to simulate a decision

algorithm.

To assess the differential quality of the outlined transfer

functions, they have been applied to the same calibration data

set of Atlantic core tops. Comparison of the resulting SST esti-

mates to measured SST reveals that typical error ranges of

transfer functions based on planktonic foraminifera range

between 1



C (SIMMAX; ANN; RAM) to 1.4



C (I&K)

(Malmgren et al., 2001).

The outlined transfer functions have the following general

requirements in common: (a) The species composition in core

top samples needs to be systematically linked to the environ-

ment of the ambient surface water mass; (b) the ecosystem

has not undergone fundamental evolutionary changes, meaning

that currently occurring species in an ocean basin are the same

as those found in the sediments of the past period under inves-

tigation; (c) the ecological preferences of species have not

changed over time; and (d) differential carbonate dissolution

at the deep-sea floor has not significantly affected the species

composition in the sediment.

Therefore, transfer functions sensu strictu are only valid for

the period of time when all modern species have already

existed. This applies to planktonic foraminifera of the last

0.5–0.7 million years. However, a number of authors have also

tried to extend the use of transfer functions beyond that time

limit. Aiming to reconstruct Miocene to Pliocene SSTs, Hooper

and Funnell (1986) argued that it was legitime to replace two

species out of some 28, two species that only appeared in the

Pleistocene, and the Miocene/Pliocene ancestor species that

probably occupied similar ecological niches. Wang (1994)

developed a transfer function for the reconstruction of Pliocene

SST in the West Pacific. However, these attempts have not

convinced the majority of investigators.

Oxygen isotopes

As with all stable isotopes, fractionation (= partitioning of iso-

topes between substances) of

O vs.

O in carbonates is tem-

perature-dependent. The thermodynamic fractionation between

O and

O that occurs during carbonate precipitation offsets

the d

O values of carbonate minerals relative to seawater

by þ30%. The d

O fractionation is a logarithmic function

of temperature with a slope of –0.2% to –0.25% per degree

Celsius, which means that a 1



C temperature increase results

in an approximately 0.23% decrease in carbonate d

O. Stable

isotopes in carbonates are measured by mass spectrometry by

determining the mass ratio in CO

gained from the reaction

in which the carbonate sample is dissolved with phosphoric

acid. Because

O is very abundant (99.76%) in nature and

O very rare (0.2%), the ratio between both isotopes is deter-

mined relative to a known standard (d notation) (Equation (4)).

Sample

¼ð

Sample



Standard

Þ1;000=

Standard

ð4Þ

In general, carbonate shells of foraminifera, both planktonic

and benthic species, provide a fairly robust d

O record of pre-

dicted equilibrium. Thus, they allow for reconstructions of both

sea surface and deepwater temperatures. However, the d

signal of some species is offset from predicted equilibrium

values by secondary effects related to cell biology, often

referred to as vital effects. Several paleotemperature equations

have been established for foraminiferal d

O, based on calibra-

tions from core-tops (Shackleton, 1974; Ganssen and Kroon,

2000) and laboratory cultures (Erez and Luz, 1983). They fit

the polynomial form of equation (Equation (5))

T ¼ a þ bðd

Calcite

 d

Water

þcðd

Calcite

 d

Water

ð5Þ

Where T = temperature in



C, a = temperature when d

calcite

–

water

=0,b is the slope, and c is the second order polyno-

mial value. Shackleton’s equation (Equation (6)), for example,

predicts a 0.23% decrease with reference to PDB (= Pee Dee

Belemnite calcite) standard in d

Oper1



C increase.

T ¼ 16:9  4:38ðd

Calcite

 d

Water

þ 0:1ðd

Calcite

 d

Water

ð6Þ

Measurements of d

O by mass spectrometry are analytically

highly precise, yielding an error smaller than 0.1%, which

would translate into a temperature error as small as < 0.4–

0.5



C, if only temperature was involved. Because d

Oin

carbonates reflects both temperature fractionation and the

O of ambient seawater where the carbonate has been preci-

pitated, changes in seawater d

O over time also need to be

considered. The global “ice volume effect” depends on the

amount of ice accumulated on continents during the time under

discussion. Where this ice volume effect is known, the portion

of temperature-related fractionation in a d

O carbonate value

can be identified.

In addition, d

O values in the ocean vary regionally with

changes in the evaporation–precipitation balance, and locally

with river and meltwater runoff. Since this “water mass effect”

is difficult to assess in past ocean scenarios, it introduces a

major portion of uncertainty into paleotemperature estimates

from d

O. Spatial and temporal variations of seawater d

thus limit the use d

O in carbonates as a temperature proxy.

However, d

O is still widely used to estimate temperatures of

warm climates in the absence of major ice sheets and low-to-

high-latitude climatic gradients much smoother than in the Pleis-

tocene; for example, for early Paleogene and Cretaceous times.

654 OCEAN PALEOTEMPERATURES

In these sediment records, the impact of diagenetic alteration on

O may have been significant. Pearson et al. (2001) demon-

strated from well-preserved planktonic foraminifera that low-

latitude SSTs during the Late Cretaceous and Eocene epochs

were as warm as 28–32



C rather than 15–23



C, as previously

estimated. Their study also revealed a strong impact of diagenetic

alteration on d

O of old fossil carbonates.

Since many sites in the Southern Ocean contain virtually no

carbonate, the use of d

O from diatoms (marine algae bearing

opal shells) has also been tested. Unfortunately, the systematics

of oxygen isotopes in opal appears to be considerably more

complex than for carbonates (Juillet-Leclerc and Labeyrie,

1986; Shemesh et al., 1992).

An innovative stable isotope approach currently under

investigation is d

Ca, based on the temperature-related fractio-

nation between

Ca and

Ca isotopes. Note that the fractiona-

tion of Ca isotopes is inverse to that of oxygen isotopes. In

culturing experiments with the planktonic foraminifer Globi-

gerinoides sacculifer, a1



C increase in temperature resulted

in a 0.24% increase of d

Ca (Nägler et al., 2000).

Magnesium-calcium ratio

In the ocean, magnesium may largely behave as a conservative

tracer. This means that it occurs in a fixed ratio to calcium, dis-

playing a long residence time. This concept, however, has been

questioned by Martin et al. (2002), who suggested some

glacial-to-interglacial changes in the oceanic Mg balance.

In calcite shells of planktonic foraminifera, this trace element

occurs at very low concentrations (about 0.5–5mmolmol

1

being directly incorporated during calcite precipitation. Cronblad

and Malmgren (1981) published the first study on Quaternary

climate cycles based on Mg in planktonic foraminifera. Newly

developed high-precision analytical devices (Atomic Absorp-

tion Spectrophotometry (AAS); Inductively Coupled Plasma

Spectrometry (ICP-MS; ICP-AS; ICP-AES)) facilitated high-

precision Mg/Ca analyses of foraminiferal calcite and per-

mitted its potential as a quantitative paleotemperature indicator

to be explored. Precise measurements require a complex clean-

ing process of foraminifera shells, including physical and che-

mical steps where extraneous phases of Mg are removed, but

the Mg content of the shells must not be affected.

The underlying concept for using Mg/Ca as a temperature

proxy is based on the endothermic substitution of Mg in calcite,

favoring the Mg substitution at higher temperatures. This rela-

tionship is not linear but underlies an exponential function. For

inorganic calcite precipitates, the van’ t Hoff equation predicts

an exponential Mg/Ca increase by 3% per degree Celsius. For-

aminifera shells in general contain 5–10 times less Mg than inor-

ganic calcite precipitates from seawater. The reason for this shift

is still unknown. Possibly, lower Mg or higher Ca somehow is

favorable for the internal-cell carbon cycle. On the other hand,

foraminifera show a stronger response of shell Mg (on average,

9% increase in Mg per degree centigrade Celsius) to temperature

than predicted by thermodynamics (Figure O25; Lea et al.,

1999). This makes the Mg/Ca ratio of both planktonic and

benthic foraminifera shells a very sensitive tool for reconstruct-

ing surface and deep-water temperatures.

The Mg substitution in calcite is different for the different

foraminifera species, which necessitates a temperature calibra-

tion of each single species. Calibrations based on cultured

Figure O25 Comparison of exponential fits of Mg /Ca in planktonic foraminifera with temperature from cultered and core top planktonic

foraminifera (from Lea et al., 1999).

OCEAN PALEOTEMPERATURES 655

planktonic foraminifera under controlled temperature condi-

tions and from core tops fit equations of the form

Mg=Caðmmol mol

1

Þ¼bemT ð7Þ

where b is the pre-exponential constant (relevant for the abso-

lute temperature, m the exponential constant (relevant for the

temperature difference) and T is the temperature (Lea et al.,

1999) (Equation (7)). Parallel Mg/Ca measurements on differ-

ent planktonic foraminifera species that calcify their shells pre-

ferentially at different water depths may be used to reconstruct

vertical temperature profiles of the water column according to

calcification depth.

It is generally more difficult to assess the temperature of

cold deep-water masses than warm temperature masses because

temperature differences in deep-water masses are generally

small. Moreover, the calibration of the slope between tempera-

ture and Mg/Ca is still quite uncertain near its “cold end.” To

address questions on the evolution of deep-water temperatures,

still more precise calibrations of benthic species are required

(Martin et al., 2002). Nevertheless, the potential of Mg/Ca in

benthic foraminifera has been demonstrated in a recent study

of Lear et al. (2000). They suggested an ocean-wide gradual

decrease in deep-water temperature of 12



C over Cenozoic

times on the basis of a mixed benthic Mg/Ca record

(Figure O26).

Apart from foraminifera, ostracods have also proven useful

indicators for bottom water temperatures (Dwyer et al., 1995;

Cronin et al., 1996). As opposed to foraminifera, ostracod

Mg/Ca ratios are fitted to temperatures through a linear relation-

ship. Dwyer et al. (1995) revealed that a temperature drop of 3



in the deep North Atlantic Ocean occurred together with major

Northern Hemisphere glaciation during the late Pliocene.

Differential dissolution results in an unwelcome secondary

effect on Mg/Ca ratios in carbonate shells because Mg is prefer-

entially removed from the shells, thus falsely suggesting cold

paleotemperatures. A major complication arises from the fact

that foraminifera shells are composed of carbonate layers con-

taining different Mg concentrations, which are differentially sus-

ceptible to dissolution. Also salinity, pH, and, in the long term,

changes in the marine Mg reservoir may affect Mg/Ca ratio.

Strontium represents a further trace element with potential

for temperature reconstructions, yet has been little exploited

(especially in coral records). Its distribution in planktonic

foraminifera is much more uniform (1.2–1.6 mmol mol

1

) than

Mg, with a slope of approximately 5% per 1



Long-chain alkenones

Apart from shells, organic remains of marine plankton are

also preserved in marine sediments. Proxies based on the

composition of organic matter are particularly useful for sedi-

ments barren of carbonate, which occur below the carbonate

compensation depth. Biomarkers are organic substances in the

sediment that preserve sufficient structural characteristics to

allow for conclusions on the original biogenic compound, even

when modified. Lipids in the cell membranes of phytoplankton

may vary in response to environmental stress in molecular

chain length and unsaturation and thus record a paleotempera-

ture signal. Notably, unsaturated long-chain ketones (= alke-

nones) shows a distinct dependency on growth temperature.

Alkenones are produced by a few phyto-plankton species

belonging to the class of Prymnesiophyceae; in Late Quatern-

ary sediments, they are mostly produced by the Coccolitho-

phorid Emiliana huxleyi. Their physiological function in the

living cells still is unknown, yet it is suspected that the increas-

ing alkenone unsaturation with cooler temperatures compen-

sates for increasing viscosity of the cell membrane (Prahl and

Wakeham, 1987). With decreasing temperature, the concentra-

tion of 3-fold unsaturated (C37:3) and 4-fold unsaturated

(C37:4) alkenones increases. Unsaturation of alkenones is mea-

sured by gas chromatography peak area integration. Unsatura-

tion indices U

and U

(Equations (8) and ( 9)) express the

ratio of (C37:4) (C37:3), and (C37:2) ketones or between 3-

and 2-fold unsaturated (C:37) ketones.

¼½C37:2½C37:4=½C37:2 þC37:3 þC37:4ð8Þ

¼½C37:2=½C37:2 þ C37:3ð9Þ

Indeed, the unsaturation indices adopt high values in low lati-

tude warm waters and low values in high latitude cold waters.

The role of 4-fold unsaturated alkenones, which only occur

at high latitudes, is still debated. Some authors suspect that

they are related to low salinity rather than to low SST. Calibra-

tions of the U

index versus growth temperature have been

Figure O26 Gradual decrease of bottom water temperatures over Cenozoic times as deduced from Mg /Ca of mixed benthic foraminifera species

(Lear et al., 2000).

656 OCEAN PALEOTEMPERATURES

established based on culturing experiments (Prahl and Wakeham,

1987; Prahl et al., 1988), particulate organic matter from

sediment trap samples obtained in the water column (Herbert,

2001), and sediment surface samples (Müller et al., 1998).

The calibrations generally fit linear regressions of U

with

temperature in the form of the example (Equation (10)) by

Müller et al.

¼ 0:033SST þ 0:044 ð10Þ

The global calibration of core top samples by Müller et al.

(1998) covers the ocean with 456 samples between 60



N and



S and a temperature range of 1–29



C(Figure O27; Equa-

tion (10)). The regression does not significantly differ from that

obtained by Prahl et al. (1988, 1993) based on culturing experi-

ments. The best overall fit has been shown for average annual

temperatures at 10 m water depth (Müller et al., 1998). How-

ever, growth depth may vary regionally, as shown for the sub-

tropical gyres, where coccolithophorids tend to live in deeper

layers (references in Herbert, 2001).

Strictly speaking, these calibrations are valid for the Late

Pleistocene back to 250,000 y

BP, when Emiliana huxleyi made

its first appearance in marine sediments. However, alkenones

produced earlier by other Prymnesiophyts may have preserved

a comparable temperature signal. Herbert and Schuffert (1998)

analyzed alkenones in Miocene to Pliocene samples from ODP

site 958 and found a 5



C temperature drop off Northwest Africa

in the Late Pliocene. They argued that as long as alkenones per-

formed a similar (though as yet unknown) biological function,

calibrations established for the Late Pleistocene should similarly

apply to older alkenone records. Recently, alkenones were also

identified in Lower Cretaceous sediments, representing the old-

est records of these compounds (Brassell and Dumitescu, 2003).

Summary

As a primary expression of the Earth is past climatic stages,

paleotemperatures are a central issue in paleoceanography.

Most concepts used to reconstruct ocean paleotemperatures

rely on marine microfossils, either on plankton assemblages

or on the isotopic or trace metal composition of carbonate

shells. Table O1 summarizes the main approaches currently

used in paleoceanography to reconstruct paleotemperatures.

Mara Weinelt

Figure O27 Global calibration of U

index versus temperature by

ller et al. (1998).

Table O1 Summary of main approaches currently used to reconstruct ocean paleotemperatures and environment and time period where they are

applicable

Approach/ method Reconstruction of Applicable time

Transfer functions based on plankton assemblages

I&K method SST 500,000–700,000 y

MAT /SIMMAX /RAM SST 500,000–700,000 yBP

ANN SST 500,000–700,000 yBP

Long-chain alkenone

Unsaturation indices SST 250,000 y

BP Neogene /Cretaceous?

Planktonic foraminifera Temperature of upper water column

(depending on calcification depth of single species)

Lower Cretaceous to present

Siliceous microfossils SST

Benthic foraminifera Deep-water temperature Unlimited, where d

O of seawater

can be assessed

Mg/ Ca

Planktonic foraminifera Temperature of upper water column

(depending on calcification depth of single species)

Lower Cretaceous to present

Benthic foraminifera Deep-water temperature In principle, unlimited

Ostracods Deep-water temperature In principle, unlimited

Reef corals

Sr/ Ca Tropical shallow water temperatures Historical times, in particular

O Tropical shallow water temperatures Historical times, in particular

OCEAN PALEOTEMPERATURES 657

Bibliography

Bé, A.W.H., 1977. An ecological, zoogeographic and taxonomic review

of recent planktonic foraminifera (chapter 1). In Ramsay, A.T.S. (ed.),

Oceanic Micropaleontology, vol. 1. London: Academic Press,

pp. 1–100.

Bradshaw, J.S., 1959. Ecology of living planktonic foraminifera in the

North and Equatorial Pacific Ocean. Cushman Found. Foramniferal

Res. Contr., 10,25–64.

Brassell, S.C., and Dumitrescu, M., (2003). Alkenones in early Aptian

organic-rich sediments from Shatsky Rise, ODP Leg 198. Geophys.

Res. Abstr., 5, 13106.

Brassell, S.C., Eglington, J., Marlowe, I.T., Pflaumann, U., and Sarnthein,

M., 1986. Molecular stratigraphy: A new tool for climatic assessment.

Nature, 320, 129–133.

Clark, F.W., and Wheeler, W.C., 1922. The inorganic constituents of marine

invertebrates. US Geological Survey Professional Paper 124, 55pp.

CLIMAP, 1981. Seasonal reconstructions of the Earth’s surface at the last

glacial maximum. The Geological Society of America Map and Chart

Series MC–36.

Cronblad, H.G., and Malmgren, B.A., 1981. Climatically controlled varia-

tion of Sr and in Quaternary planktonic foraminifera. Nature, 291,

61–64.

Cronin, T.M., Raymo, M.E., and Kyle, K.P., (1996). Pliocene (3.2–2.4 Ma)

ostracode faunal cycles and deep ocean circulation, North Atlantic

ocean. Geology, 24, 695–698

Cuvier, M., 1817. Essay on the theory of the Earth. Edinburgh, New York:

Arno Press (reprinted 1978).

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

Correge, T., 1995. North Atlantic deepwater temperature change during

Late Pliocene and Late Quaternary Climate cycles. Science, 270,

1347–1351.

Emiliani, C., 1955. Pleistocene temperatures. J. Geol., 63, 538–578.

Epstein, S., Buchsbaum, R., Lowenstam, H.A., and Urey, H.C., 1953.

Revised carbonate-water isotopic temperature scale. Geol. Soc. Am.

Bull., 64, 1315–1325.

Erez, J., and Luz, B., 1983. Experimental paleotemperature equation for

planktonic foraminifera. Geochim. Cosmochim. Acta, 47, 1025–1031.

Ganssen, G.M., and Kroon, D., 2000. The isotopic signature of planktonic

foraminifera from NE Atlantic surface sediments: implications for the

reconstruction of past oceanic conditions. J. Geol. Soc., London

, 157,

693–699.

Hemleben, C., Spindler, M., and Anderson, O.R., 1989. Modern Planktonic

Foraminifera. New York: Springer, 363pp.

Herbert, T.D., 2001. Review of alkenone calibrations (culture, water col-

umn, and sediments). Geochem. Geophys. Geosyst., 2(2), doi:

2000GC000055.

Herbert, T.D., and Schuffert, J.D., 1998. Alkenone unsaturation estimates

of Late Miocene through Late Pliocene sea surface temperatures at Site

958. In Proceedings of the Ocean Drilling Program, Scientific Results,

159T,17–21.

Hooper, P.W., and Funnell, 1986. Palaeoceanographic significance of Late

Miocene to early Pliocene planktonic foraminifers at DSDP Site 609.

Initial Reports of the Deep Sea Drilling Project, 94, 925–934.

Hutson, W.H., 1977. Transfer functions under no-analog conditions:

Experiments with Indian Ocean planktonic foraminifera, Quaternary

Res., 8, 355–367.

Hutton, J., 1795. Theory of the Earth. Edinburgh: Creech, vol. 1: 620pp,

vol. 2: 576pp.

Imbrie, J., and Kipp, N.G., 1971. A new micropaleontological method

quantitative paleoclimatology: Application to a late Pleistocene Carib-

bean core. In Turekian, K. (ed.), The Late Cenozoic Ice Ages.

New Haven: Yale University Press, pp. 71–181.

Juillet-Leclerc, A.D., and Labeyrie, L.D., 1986. Temperature dependance of

the oxygen isotope fractionation between diatom silica and water. Earth

Planet. Sci. Lett., 84,69–74.

Lea, D.W., Mashiotta, T.A., and Spero, H.J., 1999. Controls on magnesium

and strontium uptake in planktonic foraminifera determined by live cul-

turing. Geochim. Cosmochim. Acta, 63, 2369–2379.

Lear, C., Elderfield, H., and Wilson, P.A., 2000. Cenozoic deep-sea tem-

peratures and global ice-volumes from Mg/ Ca in benthic foraminiferal

calcite. Science, 287, 269–272.

Malmgren, B.A., Kucera, M., Nyberg, J., and Waelbroeck, C., 2001. Com-

parison of statistical and artificial neural network techniques for

estimating past sea-surface temperatures from planktonic foraminifer

census data. Paleoceanography, 16, 520–530.

Marlowe, I.T., Brassell, S.C., Eglinton, G., and Green, J.C., 1984. Long-

chain unsaturated ketones and esters in living algae and marine sedi-

ments. Org. Geochem., 6, 135–141.

Martin, P.A., Lea, D.W., Rosenthal, Y., Shackleton, N.J., Sarnthein, M., and

Papenfuß, T., 2002. Quaternary deep sea temperature histories derived

from benthic foraminiferal Mg/ Ca. Earth Planet. Sci. Lett., 198,

193–209.

Müller, P.J., Kirst, G., Ruhland, G., von Storch, I., and Rosell-Melé, A.,

1998. Calibration of the alkenone paleotemperature index Uk’37 based

on core-tops from the eastern South Atlantic and the global ocean

(60



N–60



S). Geochim. Cosmochim. Acta, 62, 1757–1772.

Nägler, T.F., Eisenhauer, A., Müller, A., Hemleben, C., and Kramers, J.,

2000. The d

Ca temperature calibration on fossil and cultured Globi-

gerinoides sacculifer: new tool for reconstruction of past sea surface

temperatures. Geochem., Geophys., Geosys., 1, doi: 2000GC000091.

Nürnberg, D., Bijma, J., and Hemleben, C., 1996. Assessing the reliability

of magnesium in foraminiferal calcite as a proxy for water mass tem-

peratures. Geochimica et Cosmochimica Acta, 60, 803–814.

Pearson, P.N., Ditchfield, P.W., Singano, J., Harcourt-Brown, K.G.,

Nicholas, C.J., Olsson, R.K., Shackleton, N.J., and Hall, M.A., 2001.

Warm tropical sea surface temperatures in the late Cretaceous and

Eocene epochs. Nature, 413, 481–487.

Pflaumann, U., Duprat, J., Pujol, C., and Labeyrie, L.D., 1996. A modern

analog technique to deduce Atlantic sea surface temperatures from

planktonic foraminifera in deep-sea sediments, Paleoceanography, 11,

15–35.

Pflaumann, U., Sarnthein, M., Chapman, M., Duprat, J., Huels, M., Kiefer,

T., Maslin, M., Schulz, H., van Kreveld, S., Vogelsang, E., and Weinelt,

M., 2003. The Glacial North Atlantic: Sea-surface conditions recon-

structed by GLAMAP-2000. Paleoceanography, 18, 1065.

Phleger, F.B., Parker, F.L., and Peirson, J.F., 1953. North Atlantic core for-

aminifera. Report of the Swedish Deep-Sea Expedition, 1947–1948,

7(1), 3–122.

Prahl, F.G., and Wakeham, S.G., 1987. Calibration of unsaturation patterns

in long-chain ketone compositions for paleotemperature assessment.

Nature, 330, 367–369.

Prahl, F.G., Muehlhausen, L.A., and Zahnle, D.L., 1988. Further evaluation

of long-chain alkenones as indicators of paleoceanographic conditions.

Geochim. Cosmochim. Acta, 52, 2303–2310.

Prahl, F.G., Collier, R.B., Dymond, J., Lyle, M., and Sparrow, M.A., 1993.

A biomarker perspective on Prymnesiophyte productivity in the

Northeast Pacific Ocean. Deep Sea Res. I, 40, 2071–2076.

Shackleton, N.J., 1967. Oxygen isotope analysis and Pleistocene tempera-

tures re-assessed. Nature, 215,15–17.

Shackleton, N.J., 1974. Attainment of isotopic equilibrium between ocean

water and the benthonic foraminifera genus Uvigerina: Isotopic

changes in the ocean during the last glacial. Cent. Nat. Rech. Colloq.

Int., 219, 203–209.

Shemesh, A., Charles, C.D., and Fairbanks, R.G., 1992. Oxygen isotopes in

biogenic silica – global changes in ocean temperature and isotopic com-

position. Science, 256

, 1434–1436.

Urey, H.C, 1947. The thermodynamic properties of isotopic substances. J.

Chem. Soc., 562–581.

Waelbroeck, C., Labeyrie, L., Duplessy, J.-C., Guiot, J., Labracherie, M.,

Leclaire, H., and Duprat, J., 1998. Improving past surface temperature

estimates based on planktonic fossil faunas. Paleoceanography, 13,

272–282.

Wang, L., 1994. Sea surface temperature history of the low latitude western

Pacific during the last 5.3 million years. Palaeogeogr. Palaeoclimatol.

Palaeoecol., 108, 379–436.

Cross-references

Alkenones

Cenozoic Climate Change

CLIMAP

Coccoliths

COHMAP

Coral and Coral Reefs

Cretaceous Warm Climates

Diatoms

658 OCEAN PALEOTEMPERATURES

Dinoflagellates

Foraminifera

Geochemical Proxies (Non-Isotopic)

Heat Transport, Oceanic and Atmospheric

Ice Cores

Isotope Fractionation

Last Glacial Maximum

Marine Biogenic Sediments

Miocene Climate

Organic Geochemical Proxies

Ostracodes

Oxygen Isotopes

Paleotemperatures and Proxy Reconstructions

Pleistocene Climates

Radiolaria

Thermohaline Circulation

Transfer Functions

ORGANIC GEOCHEMICAL PROXIES

Introduction

Organic matter is produced principally by photosynthetic

plants, bacteria, and archea. The amounts and kinds of lipids,

carbohydrates, proteins and other biochemical components

have varied as the amounts and kinds of plants and microbes

have changed over time. Incorporation of organic matter into

the sediments of lakes and oceans, which are nature’s archives,

typically blends components from many sources. Biochemical

components of biota are altered to become the geochemical

constituents of sediments, first by metabolic utilization by

microbes and other biota and later by oxidation-reduction pro-

cesses. Despite these very early diagenetic changes, sedimen-

tary organic matter retains important information about its

origins, and it provides equally important information about

how it was delivered and deposited.

Different environments support different biological commu-

nities, with the result that geochemical components indicative

of changes in the communities exist in sediment horizons

deposited at different times. These proxies of paleoclimatic,

paleolimnologic, and paleoceanographic histories are provided

by the elemental, isotopic, and molecular compositions of sedi-

mentary organic matter.

Organic matter C/ N ratios

The proportions of sedimentary organic matter that originate

from aquatic as opposed to land sources can be distinguished

from compositional differences between algae and vascular

land plants. Fresh organic matter from aquatic algae is

protein-rich and cellulose-poor and yields molar C/N values

that are commonly between 4 and 9. Vascular land plants,

which are protein-poor and cellulose-rich, create organic matter

that usually has C/N ratios of 20 and greater.

Partial degradation of organic matter during early diagenesis

sometimes modifies the initial elemental compositions and hence

C/N ratios of organic matter in sediments. C/N ratios of fresh

wood samples are generally higher than those of wood that has

been buried in sediments (Meyers et al., 1995). This change

reflects selective degradation of carbon-rich sugars and lipids in

the buried wood. In contrast, the C/N ratio of algal organic mat-

ter can increase during sinking and early sedimentation as nitro-

gen-rich proteins are selectively degraded. When diagenetically

elevated C/N ratios of marine organic matter are found in sedi-

ment records, they imply periods of enhanced aquatic paleopro-

ductivity, such as occur in oceanic upwelling systems (Twichell

et al., 2002). In most depositional settings, changes in the ele-

mental composition of sedimentary organic matter are minor

and do not erase the distinctive C/N differences between the

organic matter derived from aquatic algae and land plants.

Organic matter stable isotopic compositions

The carbon and nitrogen stable isotopic compositions of bulk

sedimentary organic matter are widely used to reconstruct

paleoenvironmental conditions by assessing sources of organic

matter in sediment records and past changes in the availability

of nutrients in the surface waters of lakes and oceans.

Organic d

C values

Carbon isotopic ratios are useful to distinguish between marine

and continental plant sources of sedimentary organic matter

and to identify organic matter from different types of land

plants. The carbon isotopic compositions of organic matter

reflect principally the dynamics of photosynthetic carbon

assimilation and the

C ratio of the carbon source. Most

plants assimilate carbon using the C

pathway, which discrimi-

nates against

C to produce a d

C shift of about 20% from

the source isotopic composition. Some plants use the C

path-

way, which creates an isotope shift of about 7%. All C

plants are specialized for warm, dry, land environments.

The distinctive d

CvaluesofC

and C

plants can be com-

bined with the characteristic C/N values of algal and land-plant

tissues to identify the major sources of organic matter in sedi-

ments (Figure O28, top). Deviations from the generalized d

and C/N values occur, which represent natural variations in bio-

chemical compositions, diagenetic modifications of initial com-

positions, or valuable evidence of paleoenvironmental changes.

The amount of inorganic carbon that is available for assim-

ilation can affect the d

C values of photosynthesized organic

matter. Periods of increased primary productivity in lakes

and oceans are commonly recorded as increases in the d

values of the organic matter that becomes buried in the sedi-

ment record. Production of algal organic matter selectively

removes

C from photic zone dissolved inorganic carbon reser-

voirs. As availability of the inorganic carbon becomes less,

the

C ratio of the remaining inorganic carbon increases

and leads to less negative d

C values of subsequently pro-

duced organic matter. However, if inorganic carbon concentra-

tions are elevated, then algal photosynthesis selectively

assimilates

C from the replete carbon reservoir to produce

organic matter with more negative d

C values. Cretace-

ous black shales, which record elevated algal production under

aCO

-rich atmosphere and have d

C values of about 28%

(Rau et al., 1987), are a classic example of this kind of

environment.

N values

Nitrogen is a biolimiting nutrient in aquatic systems, and it has

multiple oxidation states that have different isotopic composi-

tions. The d

N values of sediment organic matter consequently

reflect the paleoenvironmental processes that affect nitrogen bio-

geochemical cycling. The d

N value of the dissolved NO



that

is commonly assimilated by algae is typically 5–10%,whereas

the value of atmospheric N

made accessible to land plants by

nitrogen-fixing soil bacteria is about 0%. This isotopic difference

ORGANIC GEOCHEMICAL PROXIES 659