Odekon M. Encyclopedia of paleoclimatology and ancient environments

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Interstadials during the Weichselian

Many different interstadials are defined in Europe based on

palynology. The interstadials that occurred during, for

instance, the Weichselian Glac ial have been named after the

place where the interstadial was described for the first time.

Many of these Weichselian interstadial stratotypes have

been defined in NW Europe and are described by Behre

(1989), Zagwijn (1989), and de Beaulieu and Reille ( 1992)

(Table I4). Dansgaard et al. (1993) described 24 interstadial

phases in the GRIP oxygen isotope record from the Greenland

ice, characterized by a sudden i ncrease and gradually decrease

of oxygen isotope values. Recent evidence points to a total of

25 interstadial phases that are resolved in the Greenland ice

core (NG RIP Members, 2004). These so-called Dansgaard-

Oeschger interstadial events can be correlated with low per-

centages of the planktonic cold water foraminifera Neoglobo-

quadrina pachyderma in high-resolution records from the

North Atlantic (Bond et al., 1993).

Wim Z. Hoek

Bibliography

Behre, K.E., 1989. Biostratigraphy of the last glacial period in Europe.

Quaternary Sci. Rev., 8,25–44.

Bond, G., Broecker, W., Johnsen, S., McManus, J., Labeyrie, L., Jouzel, J.,

and Bonani, G., 1993. Correlations between climate records from North

Atlantic sediments and Greenland ice. Nature, 365, 143–147.

Dansgaard, W., Johnsen, S.J., Clausen, H.B., Dahl-Jensen, D., Gun-

destrup, N.S., Hammer, C.U., Hvi dberg, C.S., Steffensen, J.P.,

Sveinbjörnsdottir, A.E., Jouzel, J., and Bond, G., 1993. Evidence

for general instability of past climate from a 250-kyr ice-core record.

Nature, 364, 218–220.

de Beaulieu, J.L., and Reille, M., 1992. The last climatic cycle at La

Grande Pile (Vosges, France): A new pollen profile. Quaternary

Science Reviews, 11, 431–438.

Jessen, K., and Milthers, V., 1928. Stratigraphical and palaeontological

studies of interglacial freshwater deposits in Jutland and north-west

Germany. Danmarks Geologiske Undersøgelse II, 48, 380pp.

Gibbard, P.L., and West, R.G., 2000. Quaternary chronostratigraphy: The

nomenclature of terrestrial sequences. Boreas, 29, 329–336.

Lowe, J.J., and Walker, M.J.C., 1997. Reconstructing Quaternary Environ-

ments. London, UK: Longman, 446pp.

North Greenland Ice Core Project (NGRIP) Members, 2004. High-

resolution record of Northern Hemisphere climate extending into the

last interglacial period. Nature, 431, 147–151.

Zagwijn, W.H., 1989. Vegetation and climate during warmer intervals in

the Late Pleistocene of western and central Europe. Quaternary Int.,

3–4,57–67.

Cross-references

Bølling-Allerød Interstadial

Millennial Climate Variability

Dansgaard-Oeschger Cycles

Ice Cores, Antarctica and Greenland

Oxygen Isotopes

Palynology

Pollen Analysis

IRON AND CLIMATE CHANGE

Iron availability limits the ability of phytoplankton in some

parts of the surface ocean to photosynthesize biomass from dis-

solved CO

in seawater. An important source of iron to phyto-

plankton is the deposition of dust from continents. The

sensitivity of dust deposition to climatic factors such as aridity

and windiness makes iron a possible link between climate and

the CO

concentration in the atmosphere. Iron climate linkages

and their biogeochemical ramifications are explored further in a

recent article by Jickels et al. (2005).

The element iron is abundant in the core and crust of the

Earth. In the early Earth, the oceans contained high concentra-

tions of dissolved iron in the form of Fe

. However, as the

oceans became oxidized by increasing O

concentrations in

the atmosphere, iron oxidized to the relatively insoluble state

, and deposited into sedimentary structures known as

Banded Iron Formations. Except for anoxic regions such as in

sediments or isolated basins such as the Black Sea, dissolved

iron concentrations have remained low in the oceans ever since.

However, iron is not absent in the oxic ocean. Fe

is stabi-

lized in the dissolved form by complexing organic compounds

called chelators or siderophores (Rue and Bruland, 1995).

These compounds appear to be biogenic, but little is known

of their structure, lifetime, or origins. Similar compounds are

secreted in soil and freshwater systems to enable plants, fungi,

and bacteria to acquire iron. A vast majority of the iron in sea-

water is found to be complexed to siderophores. In addition,

the precipitation of Fe

into solid form does not result in its

immediate removal by sinking when the particles generated

are very small. More than half of the Fe

in seawater is found

in colloidal form, too small to sink but potentially too large to

be biologically useful. In particular, the diffusion coefficient of

colloidal complexed iron will be slower than that for truly dis-

solved iron, and its availability for biological uptake will be

correspondingly less (Wu et al., 2001).

In spite of its scarcity, iron is an essential co-factor in enzy-

matic systems such as the light-harvesting photosynthesis

systems, and in nitrogenase, the enzyme which fixes NH

from

molecular N

(Falkowskietal.,1998). Because of its low solubi-

lity in oxic seawater, iron has been shown to limit the production

of biomass in several regions of the surface ocean. These regions,

historically called High Nutrient Low Chlorophyll (HNLC)

regions, include the equatorial Pacific, the North Pacific, and

the Southern Ocean. When the iron concentration in the surface

ocean is artificially increased in these areas, the iron stimu-

lates increases in chlorophyll concentration, photosynthetic

efficiency, and phytoplankton growth rates (Coale et al.,

1996). These iron fertilization experiments constitute one of

the most dramatic breakthroughs in biological oceanography

in several decades.

Table I4 Interstadials during the Weichselian in Europe: Denmark (DK),

The Netherlands (NL), Germany (D) France (F), Great Britain (UK),

compared to GRIP Summit glacial interstadials IS as defined by

Dansgaard et al. (1993)

Terrestrial sequences Europe IS

Bølling-Allerød (DK), Windermere (UK) 1

Denekamp (NL) 8

Hengelo (NL) 12

Moershoofd (NL) 13

Glinde (D) 14

Oerel (D) 16

Odderade (D) St. Germain II (F) 21

Brørup (DK), St. Germain I (F), Chelford (UK) 23

Amersfoort (NL)

478 IRON AND CLIMATE CHANGE

The link between phytoplankton and the CO

concentration

of the atmosphere occurs via the biological pump in the ocean.

Photosynthesis produces particles that can leave their source

waters by sinking. The downward net flux of biological mate-

rial leaves the surface ocean depleted in nutrients such as

2

and PO

3

, as well as dissolved CO

.CO

gas dissolves

into surface seawater, or evaporates into the atmosphere,

according to its concentration in seawater or air, and its solubi-

lity (mostly controlled by temperature). Depletion of CO

from

surface waters by biological activity therefore draws down the

concentration of the atmosphere. In the present-day ocean,

Southern Ocean surface waters contain the majority of the poten-

tially biologically usable nutrients NO

2

and PO

3

. In addition,

because the Southern Ocean is a conduit to the largest oceanic

carbon reservoir – the deep ocean, atmospheric pCO

in models

is more sensitive to surface chemistry changes in the Southern

Ocean than other oceanographic regions. Therefore, the search

for an explanation for lower glacial pCO

values has centered

on deposition of iron onto the Southern Ocean.

There are arguments for and against the Southern Ocean fer-

tilization hypothesis. On the positive side, the deposition of

dust and Fe in Antarctic ice cores was indeed higher during

glacial times, by more than an order of magnitude. Most of this

dust appears to originate from a specific region in Patagonia,

but dust deposition increased globally by a factor of two or

more. The decline in Antarctic dust deposition is also one of

the earliest indicators of the coming deglaciation (suggesting

that dust decrease could be an ultimate cause for the other

aspects of the deglaciation) (Broecker and Henderson, 1998).

On the other hand are arguments against the Southern Ocean

fertilization hypothesis. Several models of iron geochemistry

in the ocean conclude that the dominant iron source to the

Southern Ocean surface is from below, in the upwelling water,

and that increasing the negligible deposition flux will have little

effect (Archer and Johnson, 2000). Fertilization experiments in

the Southern Ocean have not resulted in increased sinking

of organic matter into deep waters, but rather an increased stand-

ing biomass stock in surface waters as the added iron is effi-

ciently recycled (Boyd et al., 2000). Trace element tracers for

sea surface nutrient concentration in the Southern Ocean have

not shown a substantial drawdown during glacial time (Boyle,

1992). Furthermore, the pCO

sensitivity to the Southern Ocean

decreases with increasing model complexity and realism, such

that even a complete drawdown of Southern Ocean nutrients

would not be enough to generate glacial pCO

values (Archer

et al., 2000). Finally, an increase in the biological pump would

result in a decrease in the O

concentration of the deep ocean,

which is not apparent in the sedimentary record.

With these counter-arguments in mind, the community is

also investigating the possibility that an increased glacial iron

supply might stimulate an increase in nitrogen fixation, result-

ing in an increase in the nitrate inventory of the ocean

(Broecker and Henderson, 1998). A release from nitrogen lim-

itation might result in a stronger biological pump, although

phosphorus availability could limit this effect. In addition, iron

fertilization is being explored as a potential fix for rising CO

concentrations resulting from fossil fuel consumption.

David Archer

Bibliography

Archer, D.E., and Johnson, K., 2000. A model of the iron cycle in the

ocean. Global Biogeochem. Cycles, 14, 269–279.

Archer, D., Eshel, G., Winguth, A., Broecker, W.S., Pierrehumbert, R.T.,

Tobis, M., and Jacob, R., 2000. Atmospheric pCO

sensitivity to the

biological pump in the ocean. Global Biogeochem. Cycles, 14,

1219–1230.

Boyd, P.W., Watson, A.J., Law, C.S., Abraham, E.R., Trull, T., Murdoch, R.,

Bakker, D.C.E., Bowie, A.R., Buesseler, K.O., Chang, H., Charette, M.,

Croot, P., Downing, K., Frew, R., Gall, M., Hadfield, M., Hall, J.,

Harvey, M., Jameson, G., LaRoche, J., Liddicoat, M., Ling, R.,

Maldonado, M.T., McKay, R.M., Nodder, S., Pickmere, S., Pridmore, R.,

Rintous, S., Safi, K., Sutton, P., Strzepek, R., Tannaberger, K., Turner, S.,

Waite, A., and Zeldis, J., 2000. A mesoscale phytoplankton bloom

in the polar Southern Ocean stimulated by iron fertilization. Nature, 407,

695–702.

Boyle, E.A., 1992. Cadmium and d

C paleochemical ocean distributions

during the Stage 2 glacial maximum. Annu. Rev. Earth Planet. Sci.,

20, 245–287.

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

ing termination II and their implications for the cause of glacial-interglacial

changes. Paleoceanography, 13,352–364.

Coale, K.H., Johnson, K.S., Fitzwater, S.E., Gordon, R.M., Tanner, S.,

Chavez, F.P., Ferioli, L., Sakamoto, C., Rogers, P., Mil lero, F.,

Steinberg, P., Nightingale, P., Cooper, D., Cochlan, W.P., Landry, M.

R., Consta ntinou, J., Rollwagen, G., Trasvina, A., and Kudela, R.,

1996. A massive phytoplan kton bloom induced by an ecosystem-scale

iro n fertilization experiment in the equatorial Pacific Ocean. Nature,

383,495–501.

Falkowski, P.G., Barber, R.T., and Smetacek, V., 1998. Biogeochemical

controls and feedbacks on ocean primary production. Science, 281,

200–206.

Jickels, T.D., 2005. Global iron connections between desert dust, ocean

biogeochemistry, and climate. Science, 308,67–71.

Rue, E.L., and Bruland, K.W., 1995. Complexation of iron(III) by natural

organic ligands in the Central North Pacific as determined by a new

competitive ligand equilibration/ absorptive cathodic stripping voltam-

metric method. Mar. Chem., 50,117–138.

Wu, J., Boyle, E., Sunda, W., and Wen, L.-S., 2001. Soluble and colloidal

iron in the oligotrophic North Atlantic and North Pacific. Science, 293,

847–849.

Cross-references

Banded Iron Formations and the Early Atmosphere

Marine Carbon Geochemistry

ISOTOPE FRACTIONATION

Isotopes are atoms whose nuclei contain the same number of pro-

tons but a different number of neutrons. The term ‘isotope’ is

derived from a Greek word meaning equal place: the various iso-

topes of an element occupy the same position in the periodic table.

Isotopes of an element can be either stable or unstable (radio-

genic). Differences in chemical and physical properties arising

from variations in atomic mass of an element are called ‘isotope

effects’. The partitioning of isotopes between two substances or

two phases of the same substance with different proportions of iso-

topes is called ‘isotope fractionation’ (Hoefs, 1997; Criss, 1999).

The isotopic composition of materials contains informa-

tion that can be used, for example, to reconstruct the palaeo-

environment and to understand the hydrological cycle and

biochemical pathways.

History

In 1913, isotopes were discovered by J. J. Thomson who found

that the element neon has two different kinds of atoms with

atomic weights 20 and 22. A few years later, 212 of the 287

ISOTOPE FRACTIONATION 479

naturally occurring isotopes were known. The reason for the

existence of isotopes was, however, unclear until the discovery

of the neutron by Chadwick in 1932. The seminal papers by

Urey (1947) and Bigeleisen and Mayer (1947) on the thermo-

dynamic properties of isotopic substances provided the basis

for the utilization of stable isotopes in geology, geochemistry,

biogeochemistry, paleoceanography and elsewhere.

Notation, defini tions

Isotopic ratios, R, are measures of the relative abundances of

isotopes of an element, E; they are usually arranged so that

the lightest stable isotope (index a), which is often also the

most abundant isotope, appears in the denominator:

R =

(stable boron isotopes are an exception:

B is about four times

more abundant than

B). The three stable oxygen isotopes

O, and

O will be used to explain the notation and basic

definitions.

The oxygen isotope with mass number 16 contains 8 pro-

tons and 8 neutrons and is denoted by

O. It is by far the most

abundant (99.76%) of the three stable oxygen isotopes. For

example, the ratio of

Oto

O in water, or the ratio of

O in carbon dioxide can be written as:

O½

;

O½þC

O½

O½þC

O½

The isotopic composition, d, of a sample, determined by

mass spectrometric methods, is measured with respect to a

standard (std):

sample

std

 1



 1;000

where b is the atomic mass of the isotope and the factor 1,000 con-

verts the d value to per mil. The standards (cf. Hoefs, 1997) used

for stable oxygen isotopes are V-SMOW (Vienna-Standard Mean

Ocean Water;

V-SMOW

·10

= 2,005.20  0.43) and V-PDB

(Vienna-Pee-Dee Belemnite,

V-PDB

·10

= 2,067.1  2.1).

The fractionation factor, a, is defined as the ratio between

the isotopic ratio in compound X and that in compound Y:

ðXYÞ

þ 1 ;000

It is a measure of the partitioning of isotopes between two

or more phases in response to an isotope effect.

Whereas the d value is the result of the whole history of the

sample, the a value is characteristic for, say, an equilibrium

between X and Y or for a process that leads from X to Y. Since

a values are usually very close to 1.0, the e notation is com-

monly used to express isotope fractionations in per mil (%):

ðXYÞ

 1



 1;000

Isotope effects

The most important isotope effects arise from differences in: (a)

random mean velocities and (b) vibrational frequencies of

molecules.

a) Random mean velocity effects. In (local) thermodynamic

equilibrium, the mean kinetic energy is equal for all molecules

and determined by temperature alone:

kin;j



T ¼ E

thermal

where m

and v

rms;j

ﬃﬃﬃﬃﬃﬃﬃﬃﬃﬃ

are mass and random mean square (rms) velocity of molecule j,

is Boltzmann’s constant, and T is absolute temperature. At a

given temperature, the rms velocity varies with the mass of the

molecule: light molecules are faster than heavy molecules (Gra-

ham’s law of diffusion):

rms;1

rms;2

ﬃﬃﬃﬃﬃﬃﬃﬃﬃﬃﬃﬃﬃﬃ

As a consequence, molecules containing the light isotope

diffuse faster. The ratio of the diffusion coefficients D

for

the light (mass m

) and the heavy molecule (mass m

)ofa

gas that diffuses through air (mass m

= 29) is given by

ðabÞ

ﬃﬃﬃﬃﬃﬃﬃﬃﬃﬃﬃﬃﬃﬃﬃﬃﬃﬃﬃﬃﬃﬃﬃﬃﬃﬃﬃﬃﬃﬃﬃﬃﬃﬃﬃ

þ m

 m

þ m

For

and

, one obtains

diff: CO

in air

ﬃﬃﬃﬃﬃﬃﬃﬃﬃﬃﬃﬃﬃﬃﬃﬃﬃﬃﬃﬃﬃﬃﬃﬃﬃﬃﬃﬃﬃﬃﬃﬃﬃ

45 þ 29

45  29

44  29

44 þ 29

 0:9956

The relation does not apply, however , for diffusion of gases

through liquids where the ratio of the diffusion coefficients is clo-

ser to one (

diff. CO

in water

 0.9993). The variation of the rms

velocity with mass also leads to differences in evaporation.

b) Effects due to molecular vibrations. The zero-point

energy of vibrations in molecules, E

, is mass dependent. For

diatomic molecules E

= hv/2, where h is Planck’s constant

and v is the (vibration) frequency. The ratio of frequencies v

and v

for two diatomic molecules of different isotopic compo-

sition is approximately given by v

/v =(µ/µ

)

1/2

where µ = m

· m

/(m

+ m

) is the reduced mass and m

and m

are the

masses of the two atoms involved.

Example: compare the frequencies for

O(v, µ = 16 · 16/

(16 + 16) = 8),

O(v

, µ

=17·16/(17+16) 8.2424), and

O(v

, µ

=18·16/(18+16) 8.4706): v

/v  0.9852,

/v 0.9718, i.e., the strength of the isotope effect increases with

the mass difference and thus it is not surprising that the resulting

isotope fractionation is also ‘mass-dependent’, as a rule.

Isotope fractionation

The difference in isotopic composition between various com-

pounds or between various phases of a single compound is

called isotope fractionation. It is noted that even the largest iso-

tope effect may not cause fractionation if the reaction goes to

completion, i.e., a quantitative reaction in which the reactant

is completely transformed into the product. However, an isoto-

pic fractionation will always be observed when a reaction has

an isotope effect and the formation of the product is not quan-

titative (Hayes, 1982).

Equilibrium isotope fractionation

Isotope fractionation can occur in equilibrium, for example,

between the various isotopic forms of carbon dioxide, bicarbonate,

and carbonate ions (Zeebe and Wolf-Gladrow, 2001). As a rule:

“The heavy isotope goes preferentially to the chemical compound

in which the element is bound most strongly.” (Bigeleisen, 1965).

For example, in the temperature range between 0 and 25



C, the

480 ISOTOPE FRACTIONATION

CO2

in seawater is 8–10% lower (“isotopically lighter”) than

HCO3



The fractionation factors for equilibrium fractionation invol-

ving three isotopes with masses m

< m

usually scale such

thata

2=1

¼ a

3=1

,whereb =(1/m

– 1/m

)/(1/m

– 1/m

)(Young

et al., 2002). For the three stable oxygen isotopes:

a =(

0.529

Nonequilibrium (“kinetic”) isotope fractionation

Nonequilibrium effects are associated with incomplete or uni-

directional processes such as evaporation, kinetic isotope

effects in chemical reactions, diffusion, or metabolic effects.

Kinetic isotope effects in chemical reactions occur when reac-

tion rates for compounds containing light or heavy isotopes

are different, which is almost always the case and thus isotope

fractionation can be expected. If the reservoir of reactants is

finite, the isotope effect associated with the reaction will not

only yield a product of different isotopic composition but will

lead to a change in the isotopic composition of the reservoir

as well (Rayleigh distillation; Bigeleisen and Wolfsberg,

1958). Kinetic isotope effects can quantitatively be understood

on the basis of the ‘transition state theory’ (Bigeleisen and

Wolfsberg, 1958). The fractionation factors for kinetic fractio-

nation involving three isotopes with masses m

< m

often scale such that a

2=1

¼ a

3=1

, where b =l

)/l

) (Young et al., 2002). For the three stable oxygen isotopes:

a =(

0.515

“Mass-independent” fractionation

A mass-independent fractionation of stable oxygen isotopes

was reported (for the first time) in 1973 in meteorites and in

1983 in laboratory studies (Thiemens and Heidenreich, 1983).

These studies show differences in the isotopic compositions

for certain compounds that contain

O only (Substance X)

and Compounds X and compound Y with fractionation

factors

a and

a (between X and Y) that are almost equal,

i.e.,

a 

a. These fractionation processes are called ‘mass-

independent’. They may occur at low pressures and, for exam-

ple, play an essential role for stratospheric ozone. The resulting

unusual composition of atmospheric oxygen has been used in

the so-called ‘triple-isotope method’ to derive estimates of

the biosphere productivity (Luz et al., 1999). A theoretical

explanation based on an extension of the Rice, Ramsperger,

Kassel, Marcus (RRKM) theory has been developed only

recently (Gao and Marcus, 2001). A mass-independent isotope

effect has also been observed during thermal decomposition of

carbonates in vacuo (Miller et al., 2002).

Dieter A. Wolf-Gladrow and Richard E. Zeebe

Bibliography

Bigeleisen, J., 1965. Chemistry of isotopes. Science, 147, 463–471.

Bigeleisen, J. and M.G. Mayer. Calculation of equilibrium constants for

isotopic exchange reactions. J. Chem. Phus., 15: 261–267, 1947.

Bigeleisen, J., and Wolfsberg, M., 1958. Theoretical and experimental

aspects of isotope effects in chemical kinetics. Adv. Chem. Phys., 1,

15–76.

Criss, R.E., 1999. Principles of Stable Isotope Distribution. New York:

Oxford University Press, 264pp.

Gao, Y.Q., and Marcus, R.A., 2001. Strange and unconventional isotope

effects in ozone formation. Science, 293, 259–263.

Hayes, J.M., 1982. Fractionation et al.: An introduction to isotopic

measurement and terminology. Spectra, 8(4), 3–8.

Hoefs, J., 1997. Stable Isotope Geochemistry, (4th edn.). Heidelberg:

Springer.

Luz, B., et al., 1999. Triple-isotope composition of atmospheric oxygen as

a tracer of biosphere productivity. Nature, 400, 547–550.

Miller, et al., 2002. Mass-independent fractionation of oxygen isotopes

during thermal decomposition of carbonates. Proc. Natl. Acad. Sci.,

99(17), 10988–10993.

Thiemens, M.H., and Heidenreich, J.E. III., 1983. The mass-independent

fractionation of oxygen: A novel isotope effect and its possible cosmo-

chemical implications. Science, 219, 1073–1075.

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

J. Chem. Soc., 562–581.

Young, et al., 2002. Kinetic and equilibrium mass-dependent isotope

fractionation laws in nature and their geochemical and cosmochemical

significance. Goechim. Cosmochim. Acta, 66(6), 1095–1104.

Zeebe, R.E., and Wolf-Gladrow, D., 2001. CO

in Seawater: Equilibrium,

Kinetics, Isotopes. Amsterdam: Elsevier.

Cross-references

Carbon Isotopes, Stable

Oxygen Isotopes

Paleotemperatures and Proxy Reconstructions

Stable Isotope Analysis

Strontium Isotopes

Sulfur Isotopes

ISOTOPE FRACTIONATION 481

KAMES

A kame is an ice contact glaciofluvial landform. It is composed

of sediment deposited by water in contact with glacier ice

either at the side of, or on the surface of a glacier. The sediment

fills of supraglacial streams and lakes, ice marginal channels

and ice-dammed lakes often form flat-topped hills or terraces

after the ice has disappeared. The term kame is a Scottish word

to describe such flat-topped hills. Kames are closely related to

eskers, and transitional forms between the two occur. Generally

kames are less continuous and are associated with supraglacial

or ice-marginal water, whereas eskers are more continuous and

are associated with subglacial water. The presence of flat or

gently sloping surfaces, representing fluvial surfaces, is some-

times said to be a diagnostic characteristic of kames. However,

postglacial collapse following the loss of ice support can ser-

iously disturb kame morphology.

Kames generally consist of well-sorted, stratified sediments.

If deposited by glacial streams they are dominated by sands

and gravels. If deposited in lakes they are characterized by clay,

silt and fine sand. Kames can occur as individual hills, as pla-

teau areas or as terraces. At the smallest end of the spectrum of

kame forms are supraglacial crevasse fillings. Where many

kames occur together within an area the term kame field is

sometimes used. Kames are often associated with kettle holes

to produce “kame and kettle” topography. Kame landscapes

tend to reflect environments where ice retreat was accompanied

by abundant meltwater and sediment supply.

The term “kame” is often applied to mean “ice contact” in

compound terms such as kame terrace, kame delta or kame

plateau. Terraces generally form where streams flow along

margins of glaciers, and therefore they lie parallel to the ice mar-

gin where the margin lies on a reverse slope. So-called “flights”

of terraces at different levels may be created as a glacier surface

reduces in elevation during ice retreat. Kame deltas, also known

as delta moraines or simply ice-contact deltas, form where gla-

cial streams emerge from the glacier into water.

As they are created in ice-contact situations kames are prone

to collapse as ice retreats. Their sedimentary architecture

therefore reflects not only the processes of deposition but also

the processes of structural collapse. Supraglacially-derived sedi-

ments are likely to be more disturbed than ice marginal deposits.

Survival of extensive kames indicates that an area was domi-

nated by down-wasting and meltout of ice after any phase of

extensive proglacial fluvial erosion, which would have removed

the kames. Benn and Evans (1998), Bennett and Glasser (1996)

and Hambrey (1994) provide useful reviews of this topic.

Peter G. Knight

Bibliography

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

UK: Arnold, 734pp.

Bennett, M.R., and Glasser, N.F., 1996. Glacial Geology. Chichester, UK:

Wiley, 364pp.

Hambrey, M.J., 1994. Glacial Environments, London, UK: UCL Press,

296pp.

Cross-references

Eskers

Glacial Geomorphology

Glacial Sediments

Glaciofluvial Sediments

Kettles

KETTLES

A kettle, or kettle hole, is an enclosed depression within glacial

sediments, caused by the melting of ice that was buried within

the sediments. The ice may be a remnant of the glacier margin

detached by ablation, or an ice block deposited from a glacial

stream. Kettles are characteristic of proglacial outwash plains

and areas of stagnant ice. They very often fill with water to

form kettle lakes. Kettles vary in size and may be just a few

meters or several hundred meters in diameter.

Some kettles form from the melting of completely buried ice

blocks, others from the melting of partially covered ice blocks,

and the morphology of the holes created reflects their origin. Where

ice is completely buried, as may be the case when stagnant glacier

ice is blanketed by supraglacial meltout sediment or where ice

blocks in a proglacial flood are deeply buried in glacifluvial depos-

its, kettles form by subsidence and may initially be very steep sided.

Where ice blocks are partially exposed, as may be the case where

ice blocks are deposited on a sandur plane during a short proglacial

flood, the debris content of the ice block and the depth of burial

have a significant impact on the form of the kettle. Debris released

from the melting ice block may create a ridge around the edge of the

kettle or may completely fill the kettle with debris, creating a

mound r ather than a hollow (Maizels, 1992) . D i ff er en t f or m s t h a t

have been described include“rimmed,”“crater,” and “till-fill”

kettles. Where an outwash plain is marked by a large number of ket-

tles the term “pitted plain” is sometimes used. Many ice blocks are

deposited on outwash plains during glacial floods, and the distribu-

tion of kettles on an outwash plain can be used to reconstruct the

routing and magnitude of flow during flood events.

Kettles may be completely filled with sediment if deposition

occurs after the ice has completely melted, but will still be recog-

nizable in the sedimentary record. The kettle will be represented

by a downfaulted block, and the amount of downfaulting at dif-

ferent levels in the profile will indicate the depth above which

sediment was deposited after the ice completely melted. Benn

and Evans (1998) and Bennett and Glasser (1996) provide useful

reviews of this topic.

Peter G. Knight

Bibliography

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

UK: Arnold, 734pp.

Bennett, M.R., and Glasser, N.F., 1996. Glacial Geology, Chichester, UK:

Wiley, 364pp.

Maizels, J., 1992. Boulder ring structures produced during Jökulhlaup

flows. Geogr. Ann., 74A,21–33.

Cross-references

Glacial Geomorphology

Glacial Sediments

Kames

Outwash Plains

484 KETTLES

LACUSTRINE SEDIMENTS

Introduction

Concerns about global change and its potential effects on

human activities, including economic, social, and political

affairs, are exerting great pressure on science to produce mod-

els and predictions for future climatic variability. In this regard,

the key to the future is a proper understanding of the past. In

principle, data sources for past climatic variations are instru-

mental meteorological and historical data. However, meteor-

ological time series are rather short and thus limited for

interpretation (Bradley, 1999). Historical data are available for

the last couple of centuries, prior to the instrumental record,

and these enable a more or less accurate determination of nat-

ural climatic and anthropogenic changes. These data, however,

become scarce and unreliable for longer timescales and thus

need to be complemented by information from indirect archives

for a better interpretation of the past. On the continents, lacus-

trine sediments provide an excellent archive for the extension

of instrumental and historical data of environmental and cli-

matic changes on a high-resolution decadal or even annual

timescale.

Lacustrine sediments contain a wide variety of information

indirectly reflecting environmental conditions. They can be

analyzed by sedimentological, geochemical, geophysical, and

biological methods combined within the field of paleolimno-

logy (Berglund, 1986; Last and Smol, 2001; Cohen, 2003).

The obtained data are generally called “paleoclimate proxy

data” or “proxies” and are used to evaluate the variability of

past climatic conditions as well as human influences on the

lake system and on the catchment basin. This is possible

because lacustrine sediments record many environmental pro-

cesses that occur in the water column and in the drainage area

of the lake.

Due to this availability of multi-proxy data sets in combina-

tion with high resolution and accurate dating, lacustrine sedi-

ments often provide a more comprehensive environmental

interpretation than other continuous continental records like

tree rings, peat bogs, and speleothems. Discontinuous records

like fluvial, colluvial, and loess deposits or soils are prone to

hiatuses and therefore are only suitable for environmental

reconstructions under special circumstances, for short time

intervals, or for certain scientific questions.

The formation of lacustrine sediments is controlled by atmo-

spheric processes (climate) but also by the geology of the

catchment basin, which is connected to the lake system through

(climatically-controlled) hydrology (Figure L1). Climate and

geology also control the formation of soils and plant cover in

the catchment as well as water chemistry and plankton commu-

nities in the lake (Wetzel, 2001). As geology does not change

significantly through the life span of most lakes, variations

recorded in the sedimentary record are almost certainly related

to climate variability, which itself is controlled by solar forcing.

However, in the most recent past anthropogenic influences

have become increasingly important.

Occurrence

Lakes are standing bodies of water that accumulate sediments

from the surrounding environment and from their water col-

umn. Lakes are relatively young features (10

years) that

have predominantly been formed during the last cold period

of the Pleistocene by glacial activities (Hutchinson, 1957;

Gierlowski-Kordesch and Kelts, 1994). Sediments deposited

in these lakes generally provide excellent archives for a large

variety of Late Glacial and Holocene paleoenvironmental

investigations. Lakes can also be found in volcanic regions

(crater lakes). In some rare occurrences, lakes of tectonic origin

(e.g., along the Great African Rift System) or impact crater

lakes exist, all of which may contain much longer sediment

records than glacial lakes, covering several glacial/interglacial

cycles or even the entire Quaternary (up to 10

years).

Formation of lacustrine sediments

The formation of lacustrine sediments depends on climatic fac-

tors and therefore, under natural conditions, the dominating

regional climate controls the type of lacustrine sedimentation

(clastic, biogenic, or evaporitic). Lacustrine sediments are the

result of three different mechanisms (Håkanson and Jansson,

1983):

 The flux of mineral particles through fluvial input from the

catchment area or by atmospheric deposition forms clastic

sediments, which to a certain extent may contain organic

particles, e.g., leaves or pieces of wood carried into the lake

by runoff.

 Biological productivity in the lake generates organic or bio-

genic sediments as the dead organic matter sinks and accu-

mulates at the lake bottom.

 Chemical precipitation of minerals from the water column

forms evaporitic sediments.

In general, sediment components can be classified into two

groups: those produced in the lake (autochthonous compo-

nents) and those formed in the catchment basin (allochthonous

components), which are transported to the lake by overland

flow, rivers and streams, or by eolian activities. Although there

are clear constraints for the environmental conditions necessary

for the formation of clastic, organic, or evaporitic sediments,

lacustrine sediments usually are the result of a combination of

these processes.

Clastic lacustrine sediments

Clastic (detrital) sediments predominate under cold climatic

conditions, e.g., at high latitudes or in alpine regions. Such

clastic or minerogenic sediments are typical of proglacial lakes.

Intensive physical weathering and the lack of a dense vegeta-

tion cover provide high amounts of minerogenic detritus, which

can easily be eroded and transported to the lake. In general, the

sediment transfer from the catchment area into the lake is cor-

related with the amount of runoff. In regions with a continental

climate, runoff is governed by the melting of snow and ice

through solar insolation during the summer. Under oceanic

climatic conditions, runoff is controlled either by the melting

of snow and ice through advective heat transport associated

with increased rainfall or solely by precipitation. Lakes with

clastic deposits are generally poor in nutrients (oligotrophic)

inhibiting high organic productivity.

Organic lacustrine sediments

Under temperate or warm, humid climatic conditions, the catch-

ment area is vegetated, reducing the availability and the

transport capability of clastic material into the lake. Chemical

weathering prevails and releases nutrients from the bedrock

that are incorporated into plant organic matter, buffered in soils,

or washed out and transported as dissolved ions into the lake.

This results in higher nutrient levels in the lake (mesotrophic

to eutrophic conditions) and the related increase of organic pro-

ductivity leads to the formation of organic or biogenic sedi-

ments. Algal blooms start with diatoms in spring, are

followed by green and blue-green algae during summer, and

are sometimes terminated by a second diatom bloom in fall.

While green and blue-green algae are easily decomposed dur-

ing deposition, siliceous diatom frustules are much more resis-

tant to dissolution and are generally well preserved in the

sedimentary record. In addition to this autochthonous sedimen-

tary component, an allochthonous component may occur dur-

ing winter: as the plant cover is reduced and thus soils

become more susceptible to erosion, runoff increases and trans-

ports minerogenic as well as organic debris of littoral and ter-

restrial origin to the lake. The latter two components are

washed into the central part of the lake basin during runoff

events related to fall or winter rains.

If carbonates are present in the catchment area, autochtho-

nous calcite precipitation adds an additional component to the

organic sediment fraction. Carbonaceous rocks are dissolved

and the solutes are transferred into the lake. In these hard water

Figure L1 Simplified illustration of forcing factors and parameters as well as processes that lead to the formation of lacustrine sediments.

486 LACUSTRINE SEDIMENTS

lakes, calcite crystals are produced by autochthonous biochem-

ical precipitation: Ca

and HCO



in the lake water stay in

solution only until the saturation point is reached. This is partly

achieved by the seasonal temperature increase in the surface

water layer during the summer insolation maximum. As CaCO

is less soluble with increasing temperature, calcite may precipi-

tate due to this effect alone. However, even more important is

the temperature-related increase of photosynthetic activities in

the water column through phytoplankton blooms. As a result

of enhanced photosynthesis, CO

is withdrawn from the lake

water and thus the pH rises up to pH 9 with a consequent

reduction in CaCO

solubility. This combination of temperature

increase and phytoplankton activity contributes to additional

precipitation of calcite crystals. Under such conditions, the

shells of calcifying organisms (e.g., ostracods, mollusks) may

also be a significant sediment constituent.

Evaporitic lacustrine sediments

Under arid or semiarid climatic conditions, evaporitic sedi-

ments are formed as the salinity and the pH of a lake increase

through enhanced evaporation of the lake water. This results

in the saturation of specific mineral compounds (salts), which

then precipitate out of the lake’s water column. In addition

to calcite, which can precipitate either biogeochemically in

mid-latitudes or physicochemically under semi-arid climatic

conditions, evaporitic sediments are also composed of calcium

sulfate (gypsum) and sodium chloride (halite), both precipitat-

ing only under arid climates.

Modification of lacustrine sedimentation

Several processes may alter lacustrine sedimentary records:

 Especially in deep lakes with steep sidewalls, turbidity cur-

rents may transport large amounts of sediment from the lake-

shores into the central basin. Turbidites may be associated

with erosion at their base, which causes a hiatus in the sedi-

mentary record.

 In shallow parts of lake basins or if strong winds induce

deep-reaching water movements, currents can cause resus-

pension of fine-grained sediments and redeposition some-

where else.

 Bioturbation caused by the burrowing activities of benthic

fauna mixes the uppermost sediment layers. The effects of

bioturbation are reduced if anoxic conditions prevent the

presence of higher organisms at the lake floor. For example,

organic rich muds (sapropels) cause anoxic conditions near

the sediment/water interface due to oxygen consumption

during the decomposition of organic matter.

 In the course of early diagenesis the growth of iron- and

manganese-minerals like vivianite, siderite, pyrite, or rhodo-

chrosite may also disturb the original layering of the soft

sediments.

If one or more of these processes are or have been present in a

lake, they can introduce hiatuses and mixing into the record

and make high-resolution investigations difficult or impossible.

Sedimentation rates

The rapid reaction of the local hydrological regime to environ-

mental change is coupled with changes in sedimentation rates

(mm yr

1

) and accumulation rates (g cm

2

1

). Lacustrine sedi-

mentation rates are in the range of 0.3 mm to several mm per

year, which is up to three magnitudes higher than in marine

sediment records. As a consequence, this high sedimentation rate

reduces the length of the time window that might be studied to ca.

years, whereas marine records reach up to 10

years.

Therefore, lacustrine sediments facilitate stratigraphic studies

with centennial to decadal resolution. If annually laminated

(varved) sediments are preserved, annual to even seasonal time

resolution is possible (Zolitschka, 2003). Thus, different kinds

of abrupt changes can be detected by the wide variety of applic-

able sedimentological, geochemical, geophysical, and biological

analytical methods (Last and Smol, 2001). Lately, multi-proxy

investigations of lake sediments became more and more impor-

tant because they provide several independent lines of evidence

for a common observation – environmental change as the re-

sponse to climate change and human impact – and therefore

improve the reliability of the interpretation.

Dating

Time control is one of the most important issues in the study of

lacustrine sediments. Without an accurate timescale a compari-

son of the sample proxy data with historical and archaeological

data and with other lacustrine archives as well as with other

proxy records on a local, regional, or global scale is impossible.

Among all sedimentary records, lake sediments provide the best

time control, as different dating methods can be applied to the

same cores. Usually, radiometric dating methods are carried

out depending on the decay rates of radiogenic isotopes, e.g.,

C dating for the last ca. 50,000 years and

137

Cs and

210

Pb for

the last 50 and 200 years, respectively. Radiometric measure-

ments are related to time by the half-life of the isotope and

through models of production and deposition of radiogenic iso-

topes. The time is provided in isotopic years and needs to be

calibrated into calendar years by comparison with monitoring

data of radioactive fallout (

137

Cs), by using depositional models

(

210

Pb), or with the help of dendrochronology (

C).

In many regions, the annual climatic cycle represents the

strongest observed cyclicity. Under permanent or seasonal

anoxic conditions near the lake bottom, this cyclicity can be

preserved as annually laminated or varved sediments. However,

varved sediments are not the rule because processes like biotur-

bation in well-oxygenated lakes disturb this cyclical layering.

Incremental dating can be applied (Zolitschka, 2003) to well-

preserved annually laminated sediments. This method is based

on cyclical accumulation of biological or lithological material

with time, leading to the formation of organic (biogenic) or

clastic varves. The major control mechanisms of this cyclicity

are the seasonal climatic variations that produce an “internal

clock.” Such a varve chronology provides either an absolute

chronology in calendar years, if the varves continue to the pre-

sent, or a floating chronology, if the record is not annually

laminated to the top of the sequence.

In addition to these “absolute” dating methods, stratigraphic

methods are often applied to provide further time control and to

facilitate correlation between neighboring sites of investigation.

Most common is the biostratigraphic method of pollen analysis

(palynology) where distinct changes in the regional pattern

of vegetation are used. For a larger regional scale, paleomagnetic

studies are also sensible and might even reveal global variations

of inclination and declination changes of the Earth’s magnetic

field. For very long records, reversals or event-like excursions

of the polarity of the Earth’s magnetic field can be used as a stra-

tigraphic tool (magnetostratigraphy). If a lake is situated in a vol-

canic region, tephra layers are another possibility to elaborate an

independent time-frame through tephrochronology. Lacustrine

LACUSTRINE SEDIMENTS 487

sediments have the advantage that one record provides several

different dating methods, which allows significant improvement

of the overall dating accuracy.

Proxy data recorded in lacustrin e sediments

Since meteorological parameters are not directly recorded

in lacustrine sediments, indirect information has to be used

for climatic reconstruction. These so-called proxy data are

either quantitative or descriptive information about sedimento-

logical, geophysical, geochemical, or biological processes that

are dominantly controlled by climatic conditions and thus pro-

vide a good approximation of climatic variability. Three major

groups of proxy data are used with lake sediments to recon-

struct past climatic and related environmental changes:

 Biological proxies (e.g., pollen, diatoms, chironomids) are

calibrated using regional training sets, which are statistically

transformed into quantitatively reconstructed values of tem-

perature, pH, salinity, or total phosphorus concentration.

 Physicochemical proxies (e.g., isotopes) can be analyzed

relatively fast and easily compared with biological proxies.

Carbon isotopes of bulk organic matter provide information

about changes in the trophic condition of a lake while oxy-

gen isotopes of authigenic carbonates or calcareous skele-

tons of microfossils like ostracods provide quantitative

information that can be linked to temperature and precipita-

tion on a global scale. Furthermore, as water residence times

in lakes are rather short, isotope records provide data about

climatic changes almost without any time lags whereas lag

effects of decades or more are typical for biological systems.

 Sedimentary proxies (e.g., flux rates, grain sizes, elemental

composition, magnetic properties, clay minerals) provide

qualitative information about major environmental changes

in the lake and the catchment basin. These data are usually

more complex than biological proxies and therefore much

more difficult to interpret.

If several different proxy data are used for a combined interpre-

tation (multi-proxy approach), the overall significance of the

results is improved considerably.

Modern calibration

The key to an improved interpretation of lacustrine sediment

records is a proper understanding of present-day processes

leading to the formation of such deposits. The primary climatic

signal can be modified and transformed in several ways. For

this reason it is necessary to establish correlations between run-

off events, sediment sources, and sediment deposition as well

as between climatic parameters and biological sedimentary

proxies, e.g., vegetation (pollen), diatoms, or chironomids. Cor-

relations need also to be established between climatic parameters

and stable isotopes or between sediment texture, structure, and

composition. The period of monitoring should ideally exceed sev-

eral years to obtain site-specific calibration data sets and transfer

functions valid for larger regions. These can then be used to recon-

struct quantitative variations of climatic variables like tempera-

ture or precipitation, which are crucial for the evaluation and

refining of global to regional climate models.

Conclusions

Evidence of past and present climate variability as well as more

recent anthropogenic influences on lakes and their catchment

basins are recorded in lacustrine sediments. The major difficulties

in employing these natural archives for paleoenvironmental and

paleoclimatic reconstructions are in establishing a precise and reli-

able timescale and interpreting the record properly. Ideally, both

obstacles can be circumvented if lake sediments are studied,

obtaining multi-proxy data on multiple-dated timescales. Success-

fully applied, lacustrine sedimentary archives provide a unique

opportunity to study the past dynamics of environmental systems

on local, regional, and even global scales. This geological perspec-

tive is necessary in order to generate realistic projections of future

environmental changes that might, at least in part, result from

human activities.

Bernd Zolitschka and Dirk Enters

Bibliography

Berglund, B.E. (ed.), 1986. Handbook of Holocene Palaeoecology and

Palaeohydrology. Chichester & New York: Wiley, 869pp.

Bradley, R.S., 1999. Paleoclimatology – Reconstructing Climates of the

Quaternary. San Diego, CA: Harcourt Academic Press. International

Geophysics Series, 64, 613pp.

Cohen, A.S., 2003. Paleolimnology: The History and Evolution of Lake

Systems. Oxford: Oxford University Press, 528pp.

Gierlowski-Kordesch, E., and Kelts, K.(eds.), 1994. Global Geological Record

of Lake Basins. Cambridge, UK: Cambridge University Press, 427pp.

Håkanson, L., and Jansson, M., 1983. Principles of Lake Sedimentology.

Berlin, Heidelberg: Springer-Verlag, 316pp.

Hutchinson, G.E., 1957. A Treatise on Limnology. I. Geography, Physics

and Chemistry. New York: Wiley.

Last, W.M., and Smol, J.P. (eds.), 2001. Tracking Environmental Change

Using Lake Sediments. Developments in Paleoenvironmental Research,

vol. 1–4. Dordrecht: Kluwer.

Wetzel, R.G., 2001. Limnology: Lake and River Ecosystems. San Diego,

CA: Academic Press, 1006pp.

Zolitschka, B., 2003. Dating based on freshwater and marine laminated

sediments. In Mackay, A., Battarbee, R., Birks, J., and Oldfield, F.

(eds.), Global Change in the Holocene: Approaches to Reconstructing

Fine-resolution Climate Change. London: Edward Arnold, 92–106.

Cross-references

Carbon Isotopes, Stable

Climate Forcing

Climate Variability and Change, Last 1,000 Years

Continental Sediments

Cyclic Sedimentation (Cyclothems)

Dating, Biostratigraphic Methods

Dating, Magnetostratigraphy

Dating, Radiometric Methods

Diatoms

Evaporites

Geochemical Proxies (Non-Isotopic)

Holocene Climates

Lake Level Fluctuations

Mineral Indicators of Past Climates

Ostracodes

Oxygen Isotopes

Paleoclimate Proxies, an Introduction

Paleohydrology

Paleolimnology

Palynology

Pleistocene Climates

Pollen Analysis

Proglacial Lacustrine Sediments

Quaternary Climate Transitions and Cycles

Radiocarbon Dating

Sapropels

Sedimentary Indicators of Climate Change

Stable Isotope Analysis

Tephrochronology

Transfer Functions

Varved Sediments

488 LACUSTRINE SEDIMENTS