Odekon M. Encyclopedia of paleoclimatology and ancient environments

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SSTand Mg/Ca (Mitsuguchi et al., 1996), and also SSTand U/Ca

(Min et al., 1995), although such correlations are rarely constant

within an individual coral or different corals from the same loca-

tion and have not received widespread utility in paleoenviron-

mental reconstruction. Similar measurements have also been

made on bivalve shells and fish otoliths. All may be subject to vital

effects (less predictable physiological variations associated with

the life cycle of an organism) and micro-environmental (habitat)

variations.

Measurement techniques

While early work on temperature reconstruction from scleracti-

nian coral Sr/Ca ratios had a precision of just 3



C, close to

the seasonal SST range of their tropical habitat, the use of thermal

ionization mass spectrometry (TIMS) and subsequently (sector

field) inductively coupled plasma atomic emission spectrophoto-

metry (ICP-ES) has led to an order of magnitude improvement

in precision. Similarly, advances in combining laser ablation

technology with inductively coupled plasma-mass spectrometry

(ICP-MS) offer the opportunity to measure multiple elemental

ratios simultaneously and rapidly on individual foraminifera or

calcitic coccolith platelets.

Molecular source indicators

Organisms can regulate the production of constituent lipids,

dependent on variable prevailing environmental conditions

such as temperature, light, salinity and nutrient availability.

The preservation of such compounds in the molecular sedimen-

tary record provides a sink for potentially reconstructing envir-

onmental and climatic variations. Therefore, the molecular

composition of sedimentary organic matter can reveal detailed

information about its production, transport and preservation.

Research since the 1970s has sought to identify and characterize

a complex suite of sedimentary organic compounds diagnostic of

their biological origins, and to relate these compounds to likely

sources of OM (e.g., Brassell et al., 1986;Repetaetal.,1992).

Biological marker compounds,orsimplybiomarkers, are extra-

ctable sedimentary lipids with relatively specific biological

origins that can provide equally specific organic matter source

information (Eglinton and Calvin, 1967). They are typically

less sensitive to degradation than the average for bulk organic

matter, and may follow an established pattern of alteration

during diagenesis, resulting in predictable lipid end-products

(“geolipids”). Several hundred biomarkers have been identified

in marine sediments, of which only those most commonly inves-

tigated in sedimentary records with potential to reconstruct

paleoclimate are discussed here; the long-chain C

alkenones,

n-alkanes, n-alkanols, n-alkanoic acids, sterols and pigments.

Several factors are prerequisites for the utility of any speci-

fic biomarker. The survival of the biomarker signal during sedi-

mentation is crucial to its use. Unfortunately, some of the most

diagnostic biomarkers, such as carotenoids, are frequently the

most labile. Alternatively, a biomarker may be ubiquitous or too

widespread in the sedimentary record, such as long chain

n-alkanes, to provide detailed diagnostic information about a spe-

cific organism, but may still be informative about a particular

environment. Thus, it is important for the utility of a biomarker

that it should originate from a specific organism or type of organ-

ism, from a restricted environment, with modern analogous organ-

isms. It is preferable that a biomarker occurs widely, thus

providing information about a variety of sedimentary regimes,

and over an extended period of geological time. Finally, it is also

preferable that the biomarker be generated and buried during a

short time period, preserving a time-dependent signal useful in

high-resolution environmental reconstructions.

Variable lipid contributions from vascular land plants or

algae can be distinguished according to consistent differences

in molecular chain-length. For example, there is significant

variation in the distribution of n-alkanes between organisms:

aquatic algae are dominated by the presence of lower (shorter

chain) n-alkanes, such as C

and C

(e.g., Blumer et al., 1971;

Simoneit, 1977; Cranwell et al., 1987), whereas the leaf waxes

of higher plants are typically dominated by longer chain C

and C

n-alkanes in the sedimentary record (e.g., Eglinton

and Hamilton, 1967; Rieley et al., 1991). The distributions of

long-chain n-alkanols and n-alkanoic acids can be similarly

related to long-term variations in organic matter supply.

The waxy coatings of land-plant leaves, flowers and pollen

contain C

and C

long-chain n-alkanoic acids, whereas

and C

n-acids are dominant in algae. A ratio of

these source indicators has been used to identify changes in

the terrigenous versus aquatic organic matter contribution (e.g.,

Bourbonniere and Meyers, 1996). Those properties of vascular

plants that protect them against environmental change, such

as desiccation, similarly favor the survival of longer chain

n-alkanes during microbial degradation, eolian transportation

over long distances (e.g., Gagosian and Peltzer, 1986) and sink-

ing in aquatic environments. Conversely, lake and marine

sediment trap studies of settling particles indicate that the con-

tribution of short-chain biomarkers from algal material is

selectivelydepletedduringsinking,presumablybecausesuch

material is fresher and more labile (Meyers and Eadie, 1993).

Combined with palynological information, the lipid data can

provide important paleoclimatic evidence regarding wind

strength and direction and changing e nvironmental conditions

in source areas such as the Chinese Loess Plateau or Sahara.

Studies have shown a correspondence between the flux of

higher plant n-alkan es and n-alkanols in sediments and cli-

matic severity, especially during cold glacial intervals (e.g.,

Poynter et al., 1989;Pelejeroetal.,1999; Figure G9).

Sterol biomarkers

A number of sterols (tetracyclic alcohols) have been exploited as

paleoenvironmental proxies, differentiated by the presence/

absence of double bonds, positions of methyl groups and length

of the C

branched sidechain (Volkman, 1986; de Leeuw and

Baas, 1986). Individual sterols have been assigned specific organ-

isms of origin, such as dinosterol (dinoflagellates), brassicasterol

(diatoms) and cholesterol (Figure G10a). In addition, algal and

vascular plant contributions have been differentiated by the pre-

sence or absence of particular sterols, for example on the basis that

leaf waxes contain C

and C

,butnotC

sterols (Rieley et al.,

1991; Figure G10b). However, sedimentary sterols are practically

ubiquitous, and many sources in marine and terrestrial environ-

ments remain uncharacterized (Volkman, 1986). The distribution

of sterol biomarkers appears most informative when there are suc-

cessive variations in abundance and precursor/product relation-

ship within a single sedimentary sequence (e.g., Higginson and

Altabet, 2004).

Alkenones

The most widely applied biomarkers for paleoceanographic

reconstruction are temperature-sensitive lipids comprising the

series of C

–C

di-, tri- and tetra-unsaturated methyl and

GEOCHEMICAL PROXIES (NON-ISOTOPIC) 347

ethyl ketones (“alkenones”; Figure G11a,b). First identified

in sediments from Walvis Ridge off southwest Africa during

DSDP Leg 40 (Boon et al., 1978), they are now known to

occur in marine sediments from all oceans (e.g., Brassell,

1980; Marlowe et al., 1984). Their detection in sediments of

Cretaceous age confirmed their production over an exten-

ded period of geological time (Farrimond et al., 1986). The

compounds were initially traced to the abundant marine unicel-

lular coccolithophorid, Emiliania huxleyi (Volkman et al., 1980),

However, they have subsequently been identified in other

members of the Gephyrocapsaceae algal family (e.g., Gephyro-

capsa oceanica; Marlowe et al., 1984; Conte et al., 1998).

Because alkenones are specific biomarkers of coccolithophorids,

quantification of the abundance of the long-chain alkenones in

deep-sea sediments is used as a proxy for changing paleopro-

ductivity (e.g., Prahl et al., 1988; Madureira et al., 1997). More

significantly, the distribution of marine sedimentary alkenones

appears to be related to growth temperature of their source algae,

and thus climate (Brassell and Eglinton, 1984). Subsequent

testing of this hypothesis in laboratory cultures grown at differ-

ent temperatures (e.g., Brassell et al., 1986 and references

therein; Conte et al., 1998) has verified the temperature depen-

dence in the biosynthetic production of alkenone unsaturation.

This may be related to the ability of such aquatic organisms to

change the molecular composition of their lipid bilayer to main-

tain a constant cell membrane fluidity under a range of thermal

conditions, although this remains to be proven.

Values of U

, defined as the alkenone unsaturation index,

are calculated from the relative concentrations of the C

37:2

37:3

and C

37:4

alkenones according to the formula (Brassell

et al., 1986; Prahl and Wakeham, 1987):

¼½C

37:2

½C

37:4

= ½C

37:2

þ½C

37:3

þ½C

37:4

ðÞð2Þ

These values appear to increase linearly with growth tem-

perature, and may be converted to ambient water temperatures

through a calibration equation according to the prevailing

oceanic setting (e.g., Prahl and Wakeham, 1987; Prahl et al.,

Figure G9 d

O isotopic composition of G. ruber, U

sea surface temperature (SST), n-nonacosane, and n-hexacosan-1-ol abundances in a core

from the South China Sea covering the last 140 kyr (after Pelejero et al., 1999; reprinted from Marine Geology, 156, Pelejero C., Grimalt J.O.,

Sarnthein M., Wang L., and J.-A. Flores, Molecular biomarker record of sea surface temperature and climatic change in the South China Sea during

the last 140,000 years, pp. 109–121, Copyright (1999), with permission from Elsevier).

348 GEOCHEMICAL PROXIES (NON-ISOTOPIC)

1988; Figure G12). Where the C

37:4

alkenone is absent, such as

in warm tropical waters, the U

index, which excludes this

component, is more commonly used (Müller et al., 1998; cf.

Figure G13a–c). The index relates to the general chemical

property that the number of carbon-carbon double bonds in

related organic molecules varies inversely with their melting

points. As a consequence, values of relative abundance of the

di- and tri-unsaturated alkenones for bottom sediments in equa-

torial regions are close to unity. Conversely, the C

37:3

alkenone

dominates in sediments from beneath the much cooler Southern

Ocean (Sikes and Volkman, 1993 ). Since the algae generating

the alkenones synthesize their lipid constituents within the

photic zone, their rapid response to changes in water tempera-

tures produces relative abundances in bottom sediments that

approximate near-surface temperatures at the time of maximum

biosynthesis during spring blooming (e.g., Herbert et al., 1998).

The algal source of alkenones discovered in lacustrine sedi-

ments has not been similarly resolved, thereby precluding any

specific paleoenvironmental reconstruction from such records.

Reconstruction of sea surface temperature is informative

about changing ocean circulation, upwelling intensity (wind

strength) and near-surface stratification. The alkenone method

conveniently avoids biases inherent to temperature calibrations

based upon variable biological groups since production has

been shown to be specific to haptophytes (Marlowe et al.,

1984), and the production ratio/temperature relationships of

the two most common individuals (Emiliania huxleyi and

Gephyrocapsa oceanica) are statistically inseparable (Conte

et al., 1998). Although alkenones experience degradation, espe-

cially under oxic conditions (Prahl et al., 1989), this seems to

affect all alkenones roughly comparably, and has little signifi-

cant effect on U

values (Marlowe et al., 1984; Rontani et al.,

1997; Hoefs et al., 1998). Site-specific biases in temperature

reconstruction have been associated with alkenone incorpora-

tion, differential preservation of the chemical species, and

region-specific inconsistencies of growth seasonality. Further-

more, alkenones can be produced at various depths (and tem-

peratures) by populations of a single species (Conte et al.,

1998) and there may be temperature-independent differences

in alkenone unsaturation rates due to regional nutrient varia-

tions affecting coccolithophorid growth depth. Nevertheless,

in areas and times in which E. huxleyi has been dominant

(since first occurrence at 268 ka), such as in the North Atlantic,

paleotemperature and paleoclimatic reconstructions have

Figure G10 (a) Chemical structure of cholesterol. (b) C

, and C

sterol source characteristics (compiled from Meyers, 1997; reprinted

from Organic Geochemistry, 27, Meyers P.A., Organic geochemical

proxies of paleoceanographic, paleolimnologic, and paleoclimatic

processes, pp. 213–250, Copyright (1997), with permission from

Elsevier).

Figure G11 Molecular structures of (a) C

37:2

alkenone; (b) C

37:3

alkenone.

Figure G12 Relationships between U

index and annual mean SST

(0 m) for surface sediments from the global ocean, 60



S60



N. The

global core-top calibration ðU

¼ 0:033SST þ 0:044Þ is equivalent to

that obtained from cultures of E. huxleyi by Prahl and Wakeham (1987)

(reprinted from Geochimica et Cosmochimica Acta, 62, Mu

ller P.J.,

Kirst G., Ruhland G., Von Storch I., and A. Rosell-Mele

, Calibration of the

alkenone paleotemperature index U

based on core-tops from

the eastern South Atlantic and the global ocean (60



N–60



S),

pp. 1757–1772, Copyright (1998) with permission from Elsevier).

GEOCHEMICAL PROXIES (NON-ISOTOPIC) 349

proven valuable (e.g., Weaver et al., 1999) since they are

otherwise independent of past changes in sea water salinity

and ice volume.

Pigments

Changes in past higher plant and algal populations can be

recorded in pigment compositions preserved in sediments since

such molecules are essential for photosynthesis. Some organic

compounds exhibit specificity to certain types of plants and

have therefore been used as paleoenvironmental biomarkers.

The role of pigments is to absorb (light) energy at certain

wavelengths, converting it into chemical energy. They are

found on the chloroplast membranes and are physically

arranged to maximize light absorption. Since each pigment

may only be active within a narrow wavelength range, several

kinds of pigments are often found together. Photosynthetic

pigments can be grouped into three classes: chlorophyll,

carotenoid and xanthophyll. Each contains a chemical group

capable of selective light absorption (resulting in characteristic

coloration) known as a chromophore, which typically involves

conjugated C=C double bonds (each separated from the other

by a single bond). Combined with oxygen-containing func-

tional groups that are common to many, these sites render pig-

ments susceptible to diagenetic alteration to more stable

sedimentary end-products.

Chlorophylls are the principal photosynthetic pigments com-

mon to all plants. They contain a central tetrapyrrole ring, and

differ according to the nature of the attached alkyl sidechain.

The most common, chlorophyll a, has an ester-linked diterpe-

noid alcohol sidechain, phytol. Diagenesis affects the amount

and types of pigments preserved in sediments. Specifically,

chlorophyll is rare in marine sediments, especially where there

is a deep water column and similar pigments are degraded if they

are not rapidly incorporated into sediments. Even after incor-

poration, pigments continue to undergo diagenesis. Sedimentary

studies have led to the establishment of predictable conversion

routes for chlorophyll pigments, including de-metallization to

form pheophytins, ester hydrolysis to form pheophorbides and

transesterification to form the more resistant steryl chlorin esters

(Keely and Brereton, 1986). These products have been given the

collective name “chlorins” (Higginson, 2000).

Chlorophylls absorb solar energy most efficiently at longer

wavelengths of the visible spectrum, wavelengths that are also

readily absorbed by water. Therefore, aquatic plants have evolved

carotenoid accessory pigments to expand energy capture to a

wider range of wavelengths useful for photosynthesis. Carotenoid

pigments such as b-carotene (Figure G14a) are composed of

two six-carbon rings separated by a conjugated carbon chain.

The xanthophylls such as fucoxanthin (Figure G14b)arealso

classified as secondary pigments, since they do not transfer energy

directly from sunlight to the photosynthetic pathway, but involve

chlorophyll as an intermediary. These secondary pigments are pre-

sent in different amounts in different plants/algae and therefore act

as discrete source identifiers for sedimentary organic matter.

Figure G13 Representative partial gas chromatograms of the relative

abundance of C

37:2

(2), C

37:3

(3), and C

37:4

(4) long-chain alkenones in

surface sediments from (a) Norwegian Sea (72



N), (b) North

Atlantic Ocean (54



N), and (c) tropical Atlantic Ocean (4



(after Rosell-Mele

, 1994).

Figure G14 (a) Chemical structure of b-carotene. (b) Chemical structure of fucoxanthin.

350 GEOCHEMICAL PROXIES (NON-ISOTOPIC)

Combined sedimentary lipid/pigment studies can yield

important information regarding the history of regional phyto-

plankton community structure. For example, pigment and lipid

analyses of a core of marine sediments recovered from the

northeast Arabian Sea revealed that four biomarker com-

pounds, dinosterol, alkenones, brassicasterol and chlorins,

attributed to dinoflagellate, prymnesiophyte, diatom, and chlor-

ophyll production, respectively, exhibited concordant concen-

tration maxima coincidental with total organic carbon peaks

over the past 200 kyr (Schubert et al., 1998; Figure G15).

The similarity of the individual biomarker distributions and

good correlations with total chlorins indicated a stable phyto-

plankton community structure driven by nutrient upwelling.

A lack of correlation with records of long-chain n-alkanes

and n-alcohols illustrates the independence of the terrestrial

organic contribution.

Techniques

The analysis of organic geochemical proxies is heavily reliant

upon liquid and gas chromatographic techniques. The term

chromatography arises from the Greek words for color and

writing and accordingly involves the separation of components

of a mixture (such as colored pigments) by the differential dis-

tribution of those components between a mobile and a station-

ary phase. Mobile phases can be gas, liquid or supercritical

fluids; stationary phases can be solids or liquids adsorbed to a

solid. Separation depends on the differential affinities of solutes

for a mobile and stationary phase.

Gas chromatography (GC) developed rapidly in the 1960s after

the epic pioneering work of James and Martin (1952). For the

modern analysis of lipids, sufficient quantities of lipids to quantify

terrestrial or aquatic inputs, or to determine U

values, are typi-

cally extracted from <1 g of freeze-dried, well-homogenized bulk

Figure G15 Abundance data for paleoproductivity biomarkers (organic carbon, C

alkenones, chlorins, dinosterol and brassicasterol) plotted

against age for Arabian Sea sediment core 94KL (reprinted with permission from Nature, 394, Schubert C.J., Villanueva J., Calvert S.E., Cowie G.L.,

von Rad U., Schulz H., Berner U., and H. Erlenkeuser, Stable phytoplankton community structure in the Arabian Sea over the past 200,000 years,

pp. 563–566, Copyright (1998)).

GEOCHEMICAL PROXIES (NON-ISOTOPIC) 351

sediment by (Soxhlet) solvent extraction. This process can be

expedited using an accelerated solvent extractor (ASE) at high

temperature and pressure using small solvent volumes for a short

duration (e.g., >100



C, >1,000 psi, 20 min per sample). Organic

extracts are typically evaporated under a N

stream and then re-

diluted (e.g., with toluene) and spiked with recovery standards.

Each sample is vaporized and injected onto the head of the chro-

matographic column, which contains a stationary phase adsorbed

onto the surface of an inert solid, for example chemically bonded

to the inside of a fused silica tube. The sample is transported

through the column in its vapor phase by the flow of an inert, gas-

eous mobile phase such as helium, separated into its components

according to volatility under a temperature gradient program.

The response of each component is precisely quantified in a detec-

tor such as an FID (flame ionization detector). For alkenones, the

analytical procedure typically has a reproducibility of better than

0:005 U

units, corresponding to a temperature uncertainty of

0.2



High Pressure Liquid Chromatography (HPLC) was devel-

oped in the mid-1970s following a variety of older techniques

for chemical separations, including open-column chromatogra-

phy, paper chromatography and thin-layer chromatography,

begun by Tswett (1905 ). HPLC quickly improved with the

development of column packing materials and the additional

convenience of on-line detectors. It is now routinely employed

for the separation, identification and quantification of pigments

and sterols. It utilizes a liquid mobile phase to separate the

components of a mixture (dissolved in a solvent), which are

forced to flow through a chromatographic column under high

pressure. The mixture is resolved into its components in the

column according to their interaction with the stationary phase

(column packing material). As each component emerges from

the column, it is carried in the mobile phase to a flow cell in

which detection occurs. Typical detectors include fixed/vari-

able wavelength and diode array UV/Visible, and fluorescence

detectors. Each component is identified according to its charac-

teristic emission spectrum response.

Summary

Geochemical proxy indicators for environmental reconstruction

rely on both organic and inorganic sources. They may reveal

the history of temperature, vegetation, wind, aridity, nutrients,

ocean c irculation, and plankton trophic structure. Analysis

of organic matter involves both identification and quantifica-

tion of its bulk and molecular composition. Complementary

information is derived from inorganic proxies reliant upon

biogenic/lithogenic and metal/Ca ratios in sediment and

skeletal material. Geochemical proxy interpretations may be

confounded by post-depositional diagenesis, especially in

environments exposed to non-stable temperature, oxidation

or biological activity. The lability of geochemical compo-

nents typically renders them more sus ceptible to diagenesis.

Abundance records are susceptible to systematic bias from

preservation effect s su ch as dissolution or water column

depth. Ratio proxies may overcome such difficulties, but their

link with the primary signal (such as paleoproduction, nutrients

or temperature) is typically more tenuous and usually wholly

empirical. Although ideally the interpretation of a proxy record

relies upon the laws of thermodynamics, most paleoceanogra-

phic data are at least partly governed by poorly understood bio-

logical processes. Because further biological processes following

the production of geochemical markers by organisms can create

additional biases in the sedimentary record, multiple proxy

studies are favored to identify the primary signals of interest.

A multi-proxy approach, such as combining analysis of alke-

nones and planktonic Mg/Ca for SST reconstruction, may com-

pensate for post-depositional effects and improve the

interpretation of paleoclimates and ancient environments.

Matthew J. Higginson

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Cross-references

Alkenones

Carbon Cycle

Coccoliths

Coral and Coral Reefs

Foraminifera

Ocean Paleotemperatures

Organic Geochemical Proxies

Oxygen Isotopes

Paleoclimate Proxies, an Introduction

Phosphorus Cycle

GLACIAL ERRATICS

A glacial erratic is a glacially transported clast of non-local pro-

venance. Some erratics are noteworthy because of their large

size or exotic bedrock type. Erratics were pivotal to the devel-

opment of the Ice Age theory, popularized by Louis Agassiz.

They continue to be used to determine the limits and flow paths

of past glaciers. Erratics have also been useful in mapping ore

minerals.

Saussure was the first to use the term erratic (in 1779) for gran-

ite boulders overlying limestone in the western Alps (Flint, 1971).

Like others, he believed they were transported by water. In 1787,

Benard Friederich Kuhn was the first to interpret that they were

carried by more extensive glaciers (Imbrie and Imbrie, 1979).

James Hutton reached the same conclusion in 1795. Inhabitants

of the Alps recognized that they were carried by past glaciers

(e.g., Jean-Pierre Perraudin in 1815). Both Ignaz Venetz (from

1816–1829) and Jean de Charpentier (from 1829–1833) further

developed this hypothesis after conversations with Perraudin. Jens

Esmark in Norway (in 1824) and Reinhard Bernhardi (in 1832)

in Germany both noted erratics and argued that they indicated

glaciation of northern Europe. Jean de Charpentier introduced

Louis Agassiz to erratics in 1836. In 1837, Agassiz presented an

expanded version of Charpentier’s ideas. Agassiz proposed that

a great Ice Age had once gripped Europe. During the next 30 years,

Agassiz championed the Ice Age. Upon landing in North America

in 1848, Agassiz found erratics there also. Louis Agassiz is buried

in Boston’s Mount Auburn Cemetery and his grave is marked by

an erratic of gray granite.

Glacial erratics are commonly found in till and on end mor-

aines. They may also be found overlying bare bedrock. In

alpine glaciers, large erratics can be carried on the surface of

the ice onto which they fall during landslides. In ice sheets

such as those that covered North America and Europe 20,000

years ago, erratics must have been eroded near their beds and

then transported in the ice or on the ice surface. The exact

processes by which the very largest erratics become incorpo-

rated and transported great distances is not completely under-

stood, but erratic paths indicate that transport may occur in

stages during more than one glaciation.

Erratics range in size from pebbles to megaliths of well over

100 tons. Many erratics have clearly traveled more than 500 km,

but none appears to have traveled farther than 1,200 km (Flint,

1971). Erratics form boulder fans or erratic trains, which are

fan-shaped distributions of boulders downstream from the out-

crop of origin. Some of the most famous indicator erratics in

northern Europe include the Rapakavi granites and copper of

Finland, and the porphyries of Sweden. In North America, some

of the most famous include the Alberta erratic trains, the Sioux

Quartzite of South Dakota and Minnesota, and native copper

from the Lake Superior region. Diamonds are found as erratics

in glacial till of the western Great Lakes. A diamond-bearing

kimberlite pipe in the Canadian Northwest Territory was discov-

ered by tracking indicator minerals in erratics (Kjarsgaard and

Levinson, 2002). In Finland, numerous studies have located ores

by mapping and tracing erratic trains.

Patrick M. Colgan

Bibliography

Flint, R.F., 1971. Glacial and Quaternary Geology. New York: Wiley,

892pp.

Imbrie, J., and Imbrie, K.P., 1979. Ice Ages: Solving the Mystery.

Cambridge, MA: Harvard University Press, 224pp.

Kjarsgaard, B.A., and Levinson, A.A., 2002. Diamonds in Canada. Gems

& Gemology, 38, 208–238.

Cross-references

Glacial Geomorphology

Glacial Sediments

History of Paleoclimatology

Moraines

Mountain Glaciers

Tills and Tillites

GLACIAL EUSTASY

Introduction

Glacioeustasy is defined as global sea-level changes resulting

from terrestrial ice-volume changes. Sea level is defined as

the distance of the sea surface relative to the center of the

Earth, which closely coincides with the geoid – a surface of

equal potential gravitational energy. A more accurate definition

would take into consideration undulations due to continents or

mountain ranges, or ocean currents. Two factors control sea

level. The first is the volume of water in the ocean, which is

largely dependent on glacial eustasy and to a lesser extent on

the volume of water in terrestrial aquifers, as well as thermal

354 GLACIAL ERRATICS

expansion and contraction of seawater. The second are changes

in the shape and volume of the ocean basin, which is controlled

by continental collision, spreading rates of mid-ocean ridges,

and the formation of large oceanic igneous provinces. Of

these mechanisms, only glacial-eustatic changes are both large

(>10 m) and rapid (<<1 Myr).

As sea level either rises or falls eustatically, flooded portions

of the crust experience water loading or unloading, so that

observed fluctuations in water depth relative to the ocean floor

(termed apparent sea-level (ASL) changes; Pekar et al., 2002a)

exceed the original eustatic changes by a factor of 1.48

(Kominz and Pekar, 2001). Changes in ASL are solely of

eustatic origin, but they differ in amplitude because their refer-

ence frame is different. The concept of ASL change also differs

from that of water-depth change at passive continental margins

and other sedimentary basins, because water depth at such loca-

tions is influenced by sediment accumulation and by subsi-

dence unrelated to changes in water loading (Pekar et al.,

2003). The concept of relative sea-level change, as widely used

in the stratigraphic literature, implicitly includes the effects of

sediment loading (e.g., Posamentier et al., 1988). In the litera-

ture today, many of the sea-level estimates are not strictly

eustatic, in that these estimates include water loading. For

example, the sea-level changes observed at small islands in

the ocean (e.g., Barbados; Fairbanks, 1989) are ASL changes

because the crust that underlies such islands is regionally

loaded by water. The ice-volume-related changes in sea level

obtained from such locations for calibration against oxygen iso-

tope records need to be multiplied by 0.68 to obtain eustatic

changes (Pekar et al., 2003).

Today most of the Earth’s ice is stored in the East Antarctica

Ice Sheet (26 km

) with lesser volumes in the ice sheets of

Greenland (2.7 km

) and West Antarctica (3 km

). Portions

of the ice sheet that are below sea level do not contribute to

sea-level rise given that they already displace ocean water, as

in the example of the West Antarctic Ice Sheet, which mainly

resides below sea level. The portion of the East and West

Antarctic Ice Sheets above sea level contains the water equiva-

lent of a 66 m sea-level rise. The Greenland Ice Sheet lacks

large floating ice shelves and contains the water equivalent of

a 7 m sea-level rise. Ice stored by mountain or alpine glaciers

(0.2 km

) is minor in comparison and would be equivalent

to less than a meter of eustatic change. If the total stored

ice volume today were to melt, this would result in a eustatic rise

equivalent to 73 m. In contrast, as recent as 20 ka B.P., sea

level was 120 m lower than today because of large continental

ice sheets in the Northern Hemisphere (e.g., Fairbanks, 1989).

Methods for estimating glacial eustasy

Within historical time scales (10

–10

yr), tide gauges have

provided estimates of local sea-level changes. Although these

sea-level changes at any particular site contain a glacial-eustatic

signal, they are also controlled by expansion of sea-surface

waters due to global warming and by local crustal movement

of the area. The latter is controlled by uplift or subsidence on tec-

tonically active margins (e.g., California), subsidence due to

sediment loading, thermal cooling of the crust, and uplift from

isostatic adjustments after the retreat of an ice sheet. By examin-

ing tide gauges globally and removing local effects, the global

mean sea-level rise is estimated to be 1 to 2 mm yr

–1

(e.g.,

Houghton et al., 2001). However, it is uncertain how much of this

signal is due to the expansion of water as the oceans warm.

Measuring sea-level changes from Holocene and late Pleis-

tocene records have been accomplished mainly by examining

records from shoreline emergent reefs (e.g., Fairbanks and

Matthews, 1978; Fairbanks, 1989; Chappell et al., 1996) and

salt-water marshes (e.g., Scott and Medilio, 1978; Goodbred

et al., 1998). In the case of emergent reefs, the species Acropora

palmata was used as a sea-level indicator as this coral typically

lives on reef crests in 5 m of water or less. Glacial-eustatic

changes were estimated by making assumptions of uplift rates

of the reef. Sea-level estimates based on salt marsh studies have

used benthic foraminiferal and plant zonations to provide sea-

level estimates for the Holocene with accuracy to 10 cm.

Furthermore, modeling of the quantity of ice volume that existed

during the Last Glacial Maximum (Peltier, 1998) has provided

an additional means to constrain glacial-eustatic change.

For pre-Pliocene times, methods for constraining glacioeus-

tasy have relied on stratigraphic records from passive margins

and d

O records from both planktonic and benthic foraminiferal

shells (called tests) obtained from deep-sea cores. Although stra-

tigraphic records have been especially useful in constraining

pre-Pliocene eustatic changes (e.g., Vail et al., 1977; Haq et al.,

1987;Milleretal.,1998; Kominz and Pekar, 2001), a triad of

processes controls these records: eustasy, sediment supply, and

local tectonics. Evaluation of eustatic changes from stratigraphic

records requires estimates of the effects of sediment compaction

as well as subsidence due to flexural sediment loading of the

margin and thermal contraction of the crust (e.g., Steckler and

Watts, 1978; Steckler et al., 1999; Kominz and Pekar, 2001).

Various methods have been used to estimate eustatic changes

from continental margins. In 1977, the Exxon Production

Research Company (EPR) demonstrated the relationship of the

stratigraphic record to sea-level change by estimating sea level

from calculations of sediment aggradation on passive margins

using seismic reflection profiles and core material (Vail et al.,

1977; Haq et al., 1987). Nevertheless, controversies have arisen

over the specific methodologies used and the proprietary nature

of the data (Christie-Blick, 1991; Christie-Blick and Driscoll,

1995). However, non-proprietary studies have verified both the

number and timing of many of the sea-level changes identified

in the EPR records (e.g., Eberli, 2000; Miller et al., 1996;Pekar

and Miller , 1996; Miller et al., 1998), although the amplitudes

appear to be overestimated by up to a factor of two for many

of them (e.g., Miocene: Miller et al., 1997; Oligocene: Miller

et al., 1985; Pekar et al., 2002a; Eocene: Browning et al.,

1996; Pekar et al., 2005; Cretaceous: Miller et al., 1999).

Recent studies have modeled the effects of compaction,

sediment loading, and crustal cooling in order to isolate a

eustatic signal from stratigraphic records (i.e., backstripping).

Kominz et al. (1996) provided constraints on Eocene through

Miocene times using one-dimensional backstripping. These

methods provided quantitative constraints on the subsidence

due to compaction, but not on the rates of subsidence due to

flexure loading and thermal contraction. More recently, Kominz

and Pekar (2001), using two-dimensional flexural backstripping,

have provided perhaps the best-constrained estimates of pre-

Pliocene sea-level changes (between 34 and 23 Ma), which they

attribute mainly to ice-volume changes in Antarctica.

In cases when flexural backstripping is not used, stratigra-

phers have relied on correlating sequence boundaries (e.g., water

depth lowering) to other sites globally as well as to oxygen iso-

tope increases (inferred glacioeustatic lowerings) in order to indi-

cate a eustatic control (e.g., Browning et al., 1996; Pekar and

Miller, 1996; Oleg et al., 1999). Oxygen isotopes obtained from

GLACIAL EUSTASY 355

marine carbonates vary as a function of ice-volume changes, as

well as the temperature and precipitation/evaporation history

(i.e., salinity effect) of the seawater in which the carbonate

formed. Thus, glacial-eustatic estimates derived from d

records (e.g., Miller et al., 1985, 1991, 1998; Abreu and Ander-

son, 1998) require an evaluation of the temperature signal and

the “salinity effect”. Studies using coupled Mg/Ca ratios and

O measurements have begun to provide estimates of bot-

tom-water temperature for the Quaternary (Martin et al., 2002),

which have permitted evaluation of the d

O of the seawater

and thus placed constraints on ice volume. This technique has

been extended through most of the Cenozoic, with limited appli-

cation in older records due to diagenesis affecting the original

seawater value of the calcite (e.g., Lear et al., 2000;Martinetal.,

2002; Billups et al., 2002). Although questions have been raised

about possible oceanic Mg/Ca ratio changes through time (Bill-

ups and Schrag, 2003) and isotopic species offset, bottom-water

temperature estimates have begun to allow the isotopic proxy for

continental ice volume to be evaluated for these older times.

Finally, direct evidence of ice sheet extent, such as ice rafted

debris (IRD), has demonstrated the presence of continental ice

sheets beginning no later than the base of the Oligocene (Breza

and Wise, 1992). However, IRD provide evidence for existence

of an ice sheet only when glaciers reach the coastline.

History of glacial eustasy

Holocene and Pleistocene glacial-eustatic changes

During the past 18,000 years, the collapse of large ice sheets in

the Northern Hemisphere as well as ice volume reductions

in Antarctica have resulted in a sea-level rise of about 120 m

(e.g., Fairbanks, 1989). The decay of the ice sheets occurred

in several steps, with two meltwater pulses occurring at

12 ka and 9.5 ka, and rates of sea-level rise ranging from

24 to 28 meters per 1,000 years (Fairbanks, 1989). With the

final demise of the Northern Hemisphere ice sheet around

8 ka B.P., the rate of sea-level rise has slowed continuously

since about 8 ka (Peltier, 1999), with the current eustatic rise

due to ice melting estimated at 1mmyr

–1

The late Pleistocene contains the largest glacial-eustatic

changes of the past 200 Ma with a number of amplitudes

>100 m. In d

O records, these glacial cycles have followed a

sawtooth shape for the past 800 ka with a periodicity in phase

with the eccentricity cycle (the “100-kyr cycle”) of Milankovitch

(Shackleton and Opdyke, 1973). Beyond 800 ka, d

Orecords

indicate that the obliquity cycle (41-kyr cycle) dominated the

record through the Pliocene, coupled with smaller glacial-eustatic

amplitudes. Large scale Northern Hemisphere glaciation was

initiated at approximately 2.4 Ma (Raymo et al., 1989), as evi-

denced by d

O excursions concomitant with the occurrence of

IRD in the North Atlantic, providing unequivocal evidence for

proximal continental ice sheets (Shackleton et al., 1984). Between

2.4 and 2.7 Ma, evidence for glacial-eustatic changes less than

half those of the late Pleistocene are implied by d

O values

(Raymo et al., 1989;Jansenetal.,1990; Raymo, 1994), which

are thought to have originated from Antarctic and limited North-

ern Hemisphere ice growth. An early Pliocene warming has been

suggested to have resulted in at least a partial collapse of the

Antarctic Ice Sheet (Raymo et al., 1989;Heartyetal.,1999).

Eustatic estimates also indicate that sea level was up to 40 m

higher, supporting the notion of a significant deglaciation (Haq

et al., 1987; Dowsett and Cronin, 1990). However, this is still

controversial, with further studies required (Denton et al., 1984).

Pre-Pliocene glacia l-eustatic changes

Unlike Pleistocene glacial-eustatic changes that were the result of

both Northern and Southern Hemisphere ice sheets, pre-

Pliocene glacial eustasy is mainly the result of ice-volume changes

in Antarctica. Most of our understanding of glacial-eustatic

changes for the Tertiary comes from the Ocean Drilling Program

drilling of shallow marine strata (Bahamas, Eberli, 2000;Miller

et al., 1996; Miller et al., 1998) and deep-sea d

O records (e.g.,

Miller et al., 1987; Miller et al., 1991;Zachosetal.,2001). Miller

et al. (1991) noted that covariance between subtropical planktonic

and benthic foraminiferal d

Orecordspluslarged

O amplitudes

(<0.5%) with high d

Ovalues(<2%) in benthic d

Orecords

from the Miocene and Oligocene implied the growth of ice sheets.

Subsequent stratigraphic studies showed that sequence boundary

development and d

O events correlate well, indicating a global

mechanism (e.g., Eberli et al., 2000; Miller et al., 1996; Pekar

and Miller, 1996;Milleretal.,1998). Furthermore, estimates of

eustatic amplitudes used paired Mg/Ca ratios and d

O ratios to

place constraints on ice-volume changes and thus glacial-eustatic

changes for the past 50 Ma (Lear et al., 2000;Billupsetal.,

2002; Pekar et al., 2002b). Results from stratigraphic records

(Miller et al., 1998) indicate that glacial-eustatic changes ranged

up to perhaps 40 m ASL during the early Miocene (21–16 Ma).

In contrast, using calibrated d

O records, estimates of ice volume

ranged from 125% to 25% of the present-day Antarctic Ice Sheet

during glacial maxima and glacial minima, respectively (Pekar

and DeConto, 2006).

The Oligocene is considered the start of the Icehouse World

of the Cenozoic (Miller et al., 1991) and contains perhaps the

best estimates of sea level for a pre-Pliocene time. Kominz

and Pekar (2001) and Pekar and Kominz (2001) estimated

glacial-eustatic changes for the period between 23 and 34 Ma

by combining two-dimensional flexural backstripping with

two-dimensional paleoslope modeling of foraminiferal biofacies

and lithofacies of the New Jersey Coastal Plain (Figure G16).

Results indicate ASL amplitudes ranging from 20  10 to

80  20 m, with the largest sea-level change occurring during

the earliest Oligocene (33.5 Ma). The size of this change implies

that the Antarctic Ice Sheet was small or absent at the end of

the Eocene, but grew rapidly (<100 kyr; Zachos et al., 1996)

to nearly the present size of the East Antarctic Ice Sheet, or

perhaps up to 30% larger (Kominz and Pekar, 2001;Lear

et al., 2004; Coxall et al., 2005; Pekar and Christie-Blick,

2007)(Figure G17).

The late Oligocene Epoch has presented a climatic conun-

drum because of low d

O values from deep-sea records, which

suggest bottom water warming, coupled with at least partial

collapse of the Antarctic Ice Sheet. However, Antarctic drilling

results (e.g., Cape Roberts Project), indicate that East Antarc-

tica was heavily glaciated during the late Oligocene (Naish

et al., 2001; Roberts et al., 2003). Additionally, high d

values from Weddell Sea Sites 689 and 690 (2.9 and 3.3%)

are consistent with a fully-glaciated East Antarctic Continent

(EAC) and cold bottom water temperatures (2.0–2.5



C) (Pekar

et al., 2005a). This paradox can be explained by spatial restric-

tion of deep waters produced near the Antarctic Continent

(proto-AABW), coupled with the presence of a large ice sheet

covering most of East Antarctica. Bottom waters elsewhere

may have been influenced by a combination of proto-AABW

and warmer, more saline deep waters (Pekar et al., 2006).

Glacial-eustatic estimates at higher frequency timescales during

the Oligocene have been developed by calibrating amplitudes

356 GLACIAL EUSTASY