Odekon M. Encyclopedia of paleoclimatology and ancient environments

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annual to multidecadal modes of Labrador climate variability inferred

from tree rings. Clim. Dyn., 20, 219–228.

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Fritts, H.C., 1990. Methods of calibration, verification, and reconstruction:

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drecht: Kluwer Academic, pp. 163–165

Fritts, H.C., 1991. Reconstructing Large-Scale Climatic Patterns from

Tree-Ring Data: a Diagnostic Analysis. Tucson, AZ, USA: University

of Arizona Press.

Fritts, H.C., Blasing, T.J., Hayden, B.P., and Kutzbach, J.E., 1971. Multi-

variate techniques for specifying tree growth and climate relationships

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Fritts, H.C., and Shaskin, A.V., 1994. Modeling tree-ring structure as

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changes of a tree. For. Sci., 1(4), 271–6.

Gates, D.M., 1968. Transpiration and leaf temperature. Ann. Rev. Plant

Physiol., 19,211–238.

Giddings, J.L., Jr., 1938. Recent tree-ring work in Alaska. Tree-Ring Bull.,

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Giddings, J.L., Jr., 1943. Some climatic aspects of tree growth in Alaska.

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models. In, Proceedings of the Seventh Conference on Probability and

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Guttman, L., 1954. Some necessary conditions for common-factor analysis.

Psychometrika, 19, 149–161.

Hughes, M.K., Leggett, P., Milsom, S.J., and Hibbert, F.A., 1978. Dendro-

chronology of oak in North Wales, U.K., Tree-Ring Bull., 38,15–24.

Hughes, M.K., Kelly, P.M., Pilcher, J.R., and LaMarche, V.C., Jr. (eds.).

1982. Climate From Tree Rings. Cambridge, UK: Cambridge Univer-

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Jacoby, G.C., Cook, E.R., and Ulan, L.D., 1985. Reconstructed summer

degree days in central Alaska and northwester Canada since 1524.

Quaternary Res.. 23,18–26.

Jacoby, G.C., and D’Arrigo, R.D., 1989. Reconstructed northern hemi-

sphere annual temperature since 1671 based on high-latitude tree-ring

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Salinger, M.J., Palmer, J.G., Jones, P.D., and Briffa, K.R., 1994. Recon-

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Stockton, C.W., 1976. Long-term streamflow reconstruction in the upper

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Cross-references

Dating, Dendrochronology

Medieval Warm Period

Paleoclimate Proxies, an Introduction

Transfer Functions

DESERT VARNISH AS A PALEOCLIMATE PROXY

Desert varnish, a paper-thin accretion on rock surfaces

(Figure D34), greatly alters the appearance of bare rock faces.

Almost any lithology can host varnish, but the surface must

remain stable for thousands of years in order to accumulate var-

nish in most desert environments. Rock varnish is the better

term because this same rock coating forms on rocks in all

environments, for example, alpine, Antarctic, Arctic, perigla-

cial, stream, temperate, and tropical environments. The term

desert varnish is most often used in arid regions (Dorn, 1998).

Environmental changes influence varnish. Lichens and many

fungi, for example, secrete biological acids that destroy desert var-

nish by dissolving the manganese and iron oxides. Where rocks

exist in a desert pavement or in other settings such as an opened

rock crevice, local environmental conditions play the key role in

varnish development. However, where rock surfaces are not

greatly influenced by the local microenvironment, varnish layers,

called visual microlaminations (VML), record changes in the arid-

ity of an area. Varnishes seen in ultra thin sections (<5 mm) with a

light microscope reveal orange-yellow and black layers.

In these boulder-top positions, wetter climates favor bacter-

ial enhancement of manganese (Figure D35), producing the

black layers. Drier climates with more alkaline dust foster

development of orange-yellow (Mn-poor) layers that record

arid periods (Dorn, 1990). In places where VMLs have

been calibrated by radiometric dating methods, they can yield

DESERT VARNISH AS A PALEOCLIMATE PROXY 275

correlated ages for rock surfaces (Liu, 2003, 2006), as revealed

in a recent test: “... results of the blind test provide convincing

evidence that varnish microstratigraphy is a valid dating tool

to estimate surface exposure ages ” (Marston, 2003, p. 197).

Based on a decade of detailed analyses of over 10,000 microse-

dimentary basins, Liu ( 2003, 2006) found that black varnish

layers correspond with Heinrich events, the Younger Dryas

( Figure D36 ), and wet events during the Holocene.

The wet periods recorded in rock varnish, however, do vary

in timing regionally. Cremaschi ( 1996), for example, found a

mid-Holocene Mn-rich layer in Tunisian varnish correlating

with a regional wet pulse in North Africa. When regionally cali-

brated, varnish microlaminations provide archaeologists and

geomorphologists a powerful Quaternary tool (Liu and

Broecker, 2008), because they reveal both climatic change and

a time signal ( Figure D37). The potential of VML extends

beyond the Quaternary in circumstances where varnish is

Figure D35 Budding bacteria concentrate manganese in rock varnish, as seen here through electron microscopy and energy dispersive chemical

analyses of Negev Desert varnish.

Figure D36 Calibration of varnish microlaminations from the basin

and range of western North America (based on data in Liu, 2003),

where ka is thousands of years.

Figure D34 Backscattered electron microscope view of brighter rock

varnish formed on granodiorite (feldspar and quartz minerals) at Kitt

Peak, Arizona. The varnish is only about 0.05 mm thick. Note the

distinct boundary between varnish and rock, indicating that varnish is

an accretion.

276 DESERT VARNISH AS A PALEOCLIMATE PROXY

preserved in sedimentary environments; investigators can infer

from VML that the environment probably fluctuated between

arid and less arid environments (Dorn and Dickinson, 1989).

There are other paleoclimatic indicators found in associa-

tion with desert varnish. When other types of rock coatings

(Dorn, 1998) alternate with desert varnish, the change in coating

type often has implications for environmental change. Changes

in the trace element geochemistry of varnish, for example,

changes in carbon isotopes and lead (Dorn, 1998), also record

environmental changes. Still other indicators include micromor-

phological changes in varnish layering that provide an indicator

of the dustiness of the depositional setting. Because desert var-

nish is a weathering feature, sensitive to climatic change, changes

in its physical and chemical structure provides researchers oppor-

tunities to understand paleoclimatic changes in arid regions.

Ronald I. Dorn

Bibliography

Cremaschi, M., 1996. The desert varnish in the Messak Sattafet (Fezzan,

Libryan Sahara), age, archaeological context and paleo-environmental

implication. Geoarchaeology, 11, 393–421.

Figure D37 Rock varnish record of the latest Pleistocene wet events in Death Valley, California and its possible correlation with GISP2 d

O ice core

record in Greenland (below), where the graphic is modified from Liu (2006). In fast accumulating varnishes, such as this sample from Death Valley,

there are three wet phases in the Younger Dryas event (WP0), four wet phases in Heinrich event 1 (WP1), and two wet phases between them

(WP0þ) that can discriminate different events in the terminal Pleistocene.

DESERT VARNISH AS A PALEOCLIMATE PROXY 277

Dorn, R.I., 1990. Quaternary alkalinity fluctuations recorded in rock var-

nish microlaminations on western U.S.A. volcanics. Palaeogeogr.

Palaeoclimatol. Palaeoecol., 76, 291–310.

Dorn, R.I., 1998. Rock Coatings. Amsterdam: Elsevier, 429pp.

Dorn, R.I., and Dickinson, W.R., 1989. First paleoenvironmental interpreta-

tion of a pre-Quaternary rock varnish site, Davidson Canyon, south

Arizona. Geology, 17, 1029–1031.

Liu, T., 2003. Blind testing of rock varnish microstratigraphy as a chrono-

metric indicator: Results on late Quaternary lava flows in the Mojave

Desert, California. Geomorphology, 53, 209–234.

Liu, T., 2006. VML Dating Lab, http://www.rmldatinglab.com/ (accessed

June 24, 2006).

Liu, T., and Broecker, W.S., 2008. Rock varnish microlamination dating

of late quaternary geomorphic features in the drylands of western

USA. Geomorphology, 93, 501–523.

Marston, R.A., 2003. Editorial note. Geomorphology, 53, 197.

Cross-references

Dating, Radiometric Methods

Heinrich Events

Mineral Indicators of Past Climates

Weathering and Climate

Younger Dryas

DEUTERIUM, DEUTERIUM EXCESS

Measurements

Deuterium is the heavy stable isotope of hydrogen,

HorD,witha

natural occurrence on Earth (

H isotopic ratio, hereafter D/H)

of 153 10

6

(Lécuyer et al., 1998). For the purpose of recon-

structing paleoenvironments, it is commonly measured in paleo-

waters (e.g., groundwaters, ice, fluid inclusions), organic matter

(e.g., tree cellulose), and hydroxylated minerals (e.g., clays,hydro-

xides), in which the D/H ratio ranges between approximately

16010

6

and 80 10

6

(Mook, 2001;Jouzel,2003).

The D/H ratio of a sample is measured by mass spectrome-

try on H

(hydrogen gas), after either a total reduction of the

water, or a catalytic equilibration with H

. Both techniques

allow some automation, but the latter requires larger sample

volume (few ml, compared with few ml for the former). A

new laser spectrometric technique is achieving promising

results (Kerstel et al., 1999). The isotopic composition is

usually expressed relative to a standard, currently the Vienna

Standard Mean Ocean Water (V-SMOW), defined in 1966 by

the International Atomic Energy Agency (IAEA), with a D/H

ratio close to 155.75 10

6

. The relative composition is

expressed with the d notation:

dD ¼½ðD=HÞ

sample

=ðD=HÞ

VSMOW

 1ð1; 000; per milÞ

It is recommended that isotopic measurements be expressed

with respect to V-SMOW, and normalized on a two-standard

scale, V-SMOW and SLAP (Standard Light Antarctic Precipita-

tion), assuming a value dD=428% for SLAP (Coplen, 1995).

The different measurement techniques achieve a precision

usually better than 1%, but with a lower accuracy, as shown

by interlaboratory comparisons (Meijer, 1999).

Reconstruction of paleoenvironments

The deuterium composition dD allows similar reconstructions

of paleohydrological and paleoclimatic conditions as can be

achieved with the d

O, both quantities being related by the

ratio of their fractionation factors. Because condensation of

water occurs mainly at equilibrium in the atmosphere, and

because the ratio of their equilibrium fractionation factors is

more or less constant with temperature, the dD and d

O values

of precipitation (and of most surface waters) are linearly related

with a slope close to 8, the so-called “Meteoric Water Line”

(Craig, 1961). In contrast, kinetic conditions prevailing at eva-

poration modify both the slope and the intercept of this line

(Merlivat and Jouzel, 1979). To infer these kinetic conditions,

a second order isotopic parameter, the deuterium excess d,

has been defined as d = dD – 8 d

O (Dansgaard, 1964).

Gilles Delaygue

Bibliography

Coplen, T.B., 1995. Reporting of stable carbon, hydrogen, and oxygen

isotopic abundances. In Proc. Meet. Reference and Intercomparison

Materials for Stable Isotopes of Light Elements, Vienna, IAEA TecDoc

825, pp. 31–34 (available online at www.iaea.org).

Craig, H., 1961. Isotopic variations in meteoric waters. Science, 133,

1702–1703.

Dansgaard, W., 1964. Stable isotopes in precipitation. Tellus, 16, 436–468.

Jouzel, J., 2003. Water stable isotopes: Atmospheric composition and appli-

cations in polar ice core studies. In Holland, H.D., and Turekian, K.K.

(eds.), Treatise on Geochemistry. Amsterdam: Elsevier.

Kerstel, E.R.T., van Trigt, R., Reuss, J., Meijer, H.A.J., and Dam, N., 1999.

Simultaneous determination of the

O, and

isotope abundance ratios in water by means of laser spectrometry. Anal.

Chem., 71, 5297–5303.

Lécuyer, C., Gillet, P., and Robert, F., 1998. The hydrogen isotope compo-

sition of seawater and the global water cycle. Chem. Geol., 145,

249–261.

Meijer, H.A.J., 1999. Isotope ratio analysis on water: A critical look at

developments. In Proc. Meet. New approaches for stable isotope ratio

measurements, Vienna, IAEA TecDoc 1247, pp. 105–112 (available

online at www.iaea.org).

Merlivat, L., and Jouzel, J., 1979. Global climatic interpretation of the

deuterium–oxygen 18 relationship for precipitation. J. Geophys. Res.,

84, 5029–5033.

Mook, W.G. (ed.), 2001. Environmental Isotopes in the Hydrological

Cycle: Principles and Applications. IHP Tech. Doc. 39, UNESCO/

IAEA. (available online at: http:/ /www.iaea.org/ programmes/ ripc/

ih/ volumes/ volumes.htm).

Cross-references

Ice cores, Antarctica and Greenland

Isotope Fractionation

Oxygen Isotopes

Paleoclimate Proxies, an Introduction

Paleohydrology

Paleotemperatures, Proxy Reconstruction

Stable Isotope Analysis

DIAMICTON

Diamicton is a general term used to describe a non-sorted or

poorly sorted, sometimes non-calcareous, terrigenous or marine

sediment containing a wide range of particle sizes derived from

a broad provenance (Harland et al., 1966). Diamictons may

have sorted and stratified subfacies, usually contain “floating”

clasts of a wide provenance, and may have a fine-grained

matrix (Dreimanis and Lundqvist, 1984). The term diamictite

278 DEUTERIUM, DEUTERIUM EXCESS

is used to describe a lithified diamicton. The term, today, also

carries with it the connotation of a glacial origin but diamictons

can be deposited within a variety of geological settings.

The term, diamictite, and, from it diamicton, was first prop-

osed by Flint et al. ( 1960 ) to replace the term symmicite.

(Diamicton comes from the Greek, diamognymi – to mingle thor-

oughly.) Often synonymous with diamicton is the term “till” but

not all diamictons are tills. The term “diamict” has been used

occasionally as a separate term but has become an alternative to

diamicton (Hambrey and Harland, 1981, p. 23).

Tills are, in general, poorly sorted glaciogenic sediments,

with occasional sorted subfacies, that subdivide into lodgment,

flow, and meltout (Dreimanis, 1988). As till research expan-

ded, it became clear that tills form within terrestrial and

marine environments but such origins are not always easily

ascertained.

Diamictons can be, in general, classified as conglomerates

or various forms of cataclastics that contain varying propor-

tions of fine-grained matrix within which “floating” clasts of

a range of sizes and shapes occur (Pettijohn, 1975; Allen,

1982). Some classifications use different terms depending on

the percentage of clasts or the degree of fine-grained matrix

present (Folk, 1954; Schermerhorn, 1966; Pettijohn, 1975);

however, essentially all of these types are diamictons as defined

above. As such, diamictons can be derived from three dominant

sources: glacial, mass movement, or endogenic. All three types

can also be formed in terrestrial and subaqueous environments

and in many instances appear remarkably similar in hand-

specimen. One distinguishing characteristic of glacial diamic-

tons may be the presence of striations on the surfaces of clasts,

but caution must be exercised in using this diagnostic criterion

(Gravenor et al., 1984; Iverson, 1991).

Glacial diamictons represent the largest group of diamictons

both on land and on the ocean floor . Although the three classes of

tills, as noted above, have become accepted, newer research sug-

gests that all tills are formed under stress conditions such that it

may be more appropriate to group lodgment and flow tills as part

of a spectrum of glacial mélange (Ruszczynska-Szenajch, 1983;

Menzies, 1990). Meltout tills are a special group of diamictons that

deposit under relatively low stress conditions such that deformation

is limited (Lawson, 1981); their global extent, however , is limited

(Rappol, 1987;PaulandEyles,1990). In using micromorphology,

it is apparent that the dominant three till (diamicton) types need to

be revisited since most glacial diamictons are products of various

ductile and brittle deformation processes of deposition or emplace-

ment (Hiemstra and Meer, 1997;Meervander,1997). It has been

proposed that these glacial diamictons be renamed “tectomicts,”

indicative of the deformation history undergone in the processes

of deposition (Meer et al., 2003).

John Menzies

Bibliography

Allen, J.R.L., 1982. Sedimentary Structures: Their Character and Physical

asis (2 vols.). Amsterdam: Elsevier Science Publishing, pp. 611–679.

Dreimanis, A., 1988. Tills, their genetic terminology and classification. In

Goldthwait, R.P., and Matsch, C.L. (eds.), Genetic Classification of

Glacigenic Deposits. Rotterdam: A.A. Balkema, pp. 17–83.

Dreimanis, A., and Lundqvist, J., 1984. What should be called till? Striae,

20,5–10.

Flint, R.F., Sanders, J.E., and Rodgers, J., 1960. Diamictite, a substitute

term for symmictite. Geol. Soc. Am. Bull., 71, 1809–1810.

Folk, R.L., 1954. The distinction between grain size and mineral composi-

tion in sedimentary rock nomenclature. J. Geol., 62, 344–359.

Gravenor, G.P., von Brunn, V., and Dreimanis, A., 1984. Nature and

classification of waterlain glaciogenic sediments, exemplified by

Pleistocene, Late Paleozoic and Late Precambrian deposits. Earth Sci.

Rev., 20, 105–166.

Hambrey, M.J., and Harland, W.B. (eds.), 1981. Earth’s Pre-Pleistocene

Glacial Record. Cambridge, UK: Cambridge University Press, 1004pp.

Harland, W.B., Herod, K.N., and Krinsley, D., 1966. The definition and

identification of tills and tillites. Earth Sci. Rev., 2, 225–256.

Hiemstra, J.F., and van der Meer, J.J.M., 1997. Pore-water controlled grain

fracturing as indicator for subglacial shearing in tills. J. Glaciol., 43,

446–454.

Iverson, N.R., 1991. Morphology of glacial striae: Implications for abrasion

of glacier beds and fault surfaces. Geol. Soc. Am. Bull., 103, 1308–1316.

Lawson, D.E., 1981. Distinguishing characteristics of diamictons at the

margin of the Matanuska glacier, Alaska. Ann. Glaciol., 2,78–84.

Meer, van der J.J.M., 1997. Particle and aggregate mobility in till: Micro-

scopic evidence of subglacial processes. Quaternary Sci. Rev., 16,

827–831.

Meer, van der J.J.M., Menzies, J., and Rose, J., 2003. Subglacial till: The

deforming glacier bed. Quaternary Sci. Rev., 22 , 1659–1685.

Menzies, J., 1990. Sand intraclasts within a diamicton mélange, southern

Niagara Peninsula, Ontario, Canada. J. Quaternary Sci., 5, 189–206.

Paul, M.A., and Eyles, N., 1990. Constraints on the preservation of diamict

facies (Melt-out tills) at the margins of stagnant glaciers. Quaternary

Sci. Rev., 9,51–69.

Pettijohn, F.J., 1975. Sedimentary Rocks. New York: Harper Row.

Rappol, M., 1987. Saalian till in The Netherlands: A review. In van der

Meer, J.J.M. (eds.), Tills and Glaciotectonics. Rotterdam: A.A.

Balkema, pp. 3–21.

Ruszczynska-Szenajch, H., 1983. Lodgement tills and syndepositional gla-

citectonic processes Related to subglacial thermal and hydrologic con-

ditions. In Schlüchter, Ch., Rabassa, J., and Evenson, E.B. (eds.), Tills

and Related Deposits: Genesis, Petrology, Application, Stratigraphy.

Rotterdam: A.A. Balkema Publishers, pp. 113–117.

Schermerhorn, L.J.G., 1966. Terminology of mixed coarse-fine sediments.

J. Sediment. Petrol., 36, 831–835.

Cross-references

Glacial Geomorphology

Glaciomarine Sediments

Ice-Rafted Debris (IRD)

Sedimentary Indicators of Climate Change

Tills & Tillites

DIATOMS

Diatoms, a type of unicellular algae, are one of the most abun-

dant groups of phytoplankton and are an essential component

of the oceanic food web. As autotrophs, their primary require-

ments for survival are adequate light and nutrients. Their

primary photosynthetic pigments are chlorophyll a and c,

xanthophylls and fucoxanthin. They are widely distributed in

a variety of marine and freshwater environments, including

open ocean and lake waters, shallow marine and lake sedi-

ments, sea ice, soils and even aerial habitats (such as wet rocks

and bark). Diatoms can exist in a free-floating or planktonic

form, or can be attached to a substrate, including rocks, other

algae or sea ice. Species live both as solitary cells and colo-

nially. One of their most important distinguishing characteris-

tics is that their cell walls are impregnated with opaline

silica. These paired silica valves fit together much like a box,

with the larger epitheca fitting on top of the smaller hypo-

theca. They are held together with overlapping connecting

bands, collectively termed the girdle and, together, these com-

ponents comprise a frustule. These tests, which are intricately

DIATOMS 279

ornamented and taxonomically diagnostic, are often well-

preserved in the fossil record, hence their use in paleoecologic

and stratigraphic studies. They generally range in size from a

few microns to a few hundred or even thousand microns, though

most are around tens of microns in size (see Figures D38–D40).

Taxonomically, diatoms are placed in the class Bacillariophy-

ceae, and are grouped with the Chrysophyceae (includes silico-

flagellates and calcareous nannofossils) and Xanthophyceae in

the phylum Chrysophyta. Two orders are distinguished, the

Centrales, with radial symmetry, and the Pennales, which have

bilateral symmetry. Over 100,000 different species (extinct and

extant) have been described but perhaps only 15,000 of these

are valid species, since many of the taxa described are now

known to be variations of a single species. For example, through

culturing work, workers have documented morphologic varia-

bility at different life cycle stages. A sparse fossil record of

marine diatoms might date as far back as the Jurassic; however,

these reports of Jurassic diatoms are not universally accepted.

Certainly, a diverse flora of marine diatoms extends back to the

lower Cretaceous, suggesting their previousevolution. Lower Cre-

taceous diatoms are known from California, Russia, Antarctica

and the southwest Pacific and Indian Oceans. Freshwater diatoms

are known only as far back as the Paleocene.

Data on past and present abundance and distribution of dia-

tom assemblages are used to reconstruct and document changes

in environment, including those occurring today, for example,

in the acidification and eutrophication of lakes, and pollution

of freshwater and marine systems. Diatoms are particularly

sensitive to a number of different environmental parameters,

such as temperature, salinity, water chemistry (pH, alkalinity,

nutrient concentration) and light levels, so analysis of assem-

blages over time is used as a proxy for these factors. The rapid

regeneration rate of diatoms allows them to respond quickly

to environmental change; their siliceous tests provide long-term

documentation of these changes. Diatom data also provide

important biostratigraphic information and have been used

extensively as a correlation tool. Finally, diatoms have proven

useful in a variety of other settings, such as archeological work,

forensic studies and industrial use of diatomite (sedimentary

rock formed by the accumulation and lithification of diatomac-

eous deposits), primarily for filtration.

Amy Leventer

Cross-references

Lacustrine Sediments

Marine Biogenic Sediments

Paleoclimate Proxies, an Introduction

DINOFLAGELLATES

The ecology of dinoflagellates

Dinoflagellates are microscopic unicellular organisms or pro-

tists belonging to the division of Dinoflagellata (Fensome et al.,

1993). They inhabit most types of aquatic environments,

from lakes to open ocean, and from equatorial to arctic settings.

Figure D38 Asteromphalus sp., Southern Ocean, west of the Antarctic

Peninsula, core NBPø1ø7 KC8.

Figure D39 Thalassiosira antarctica, Southern Ocean, west of the

Antarctic Peninsula, core NBPø1ø7 KC8.

Figure D40 Cocconeis sp., Southern Ocean, west of the Antarctic

Peninsula.

280 DINOFLAGELLATES

When blooming, dinoflagellates can be responsible for “red

tides”, so called because the very large number of cells in the sur-

face water induces a color change. A few dinoflagellate species

produce neurotoxins that may be bioconcentrated by filtering

organisms, notably shellfishes, which then become poisonous and

dangerous for the health of animals feeding on them.

Living dinoflagellates are characterized by two flagella,

which permit swimming motion. Many dinoflagellates are

phototrophic (i.e., undergoing photosynthesis) and form an

important part of the planktonic primary production in lakes

and oceans. Some dinoflagellates are heterotrophic (i.e., feeding

on other organisms), and others are symbionts of marine

invertebrates such as corals, in which they are known as zoox-

anthellae.

The life cycle of dinoflagellates and the cyst stage

Most dinoflagellates have a complex life cycle involving sev-

eral stages, asexual and sexual, motile and non-motile (see

Figure D41). During the course of sexual reproduction, some

species form a diploid cell (i.e., with 2n chromosomes follow-

ing the fusion of the gametes) that is protected within a cyst,

permitting survival of the organism during a dormancy period

of variable length (Fensome et al., 1993).

The motile and cyst stages are characterized by distinct

morphologies. The cell wall or theca of many motile dinofla-

gellates is armored and consists of cellulose plates that form a

distinctive geometry or tabulation, whereas the cysts only bear

traces of the theca tabulation.

The organic-walled dinoflagellate cysts or dinocysts

Many species of dinoflagellates produce fossilizable cysts

that are composed of highly resistant organic matter, similar

in composition to the sporopollenin of pollen grains. The

organic-walled cysts, also known as “dinocysts” and previously

called “hystrichospheres” by paleontologists, are usually well

preserved in sediment. The dinocysts are typically 15–100 µm

in diameter, and are observed in palynological slides prepared

by standard laboratory procedures used for the study of pollen

and spores, which involve treatments with hydrofluoric and

hydrochloric acids. The study of dinocysts is thus considered a

sub-discipline of palynology.

Dinocysts in paleoecology, paleoceanography

and paleoclimatology

Fossil dinocysts are mainly known from marine sediments.

They appear to be particularly abundant along continental

margins (estuaries, continental shelves and slopes, epiconti-

nental seas). Dinocysts were widely used in biostratigraphy

and paleoecology of the Mesozoic and Cenozoic. In the field

of Quaternary paleoceanography and paleoecology, the study

of dinocysts is of growing interest. Because they are very resis-

tant, dinocysts are generally well perserved in sediment despite

dissolution that may affect calcalreous or siliceous biological

remains. Moreover, the development of reference databases

from surface sediment samples (Dale, 1996; Rochon et al.,

1999; Matthiessen and de Vernal, 2001; Zonneveld and Marret,

2003) has led to the documentation of strong relationships

between the distribution of dinocyst assemblages and sea-

surface parameters, including productivity and hydrographical

conditions. Transfer functions derived from these databases

have permitted the quantitative reconstruction of temperature,

salinity, and sea-ice cover extent. For example, hydrographical

maps of the northern North Atlantic during the Last Glacial

Maximum (21 ka) were established using dinocyst data

(de Vernal et al., 2000), and many regional reconstructions

are currently being developed.

Figure D41 Schematic illustration of idealized dinoflagellate life-cycle (adapted from Fensome et al., 1993).

DINOFLAGELLATES 281

The calcareou s dinoflagellates

In addition to organic-walled cysts or dinocysts, some dinofla-

gellate taxa yield calcareous microfossils generally associated

with the cyst stage (Vink et al., 2000). Calcareous dinoflagel-

lates appear to be abundant in oceanic sediments at middle to

low latitudes, where they can be used as tracers of hydro-

graphic conditions in the upper water column.

Anne de Vernal

Bibliography

Dale, B., 1996. Dinoflagellate cyst ecology: Modelling and geological

applications. In Jansonius, J., and McGregor, D.C. (eds.), Palynology:

Principles and Applications. Texas: American Association of Strati-

graphic Palynologists Foundation, College Station, 1249–1275.

de Vernal, A., Hillaire-Marcel, C., Turon, J.L., and Matthiessen, J., 2000.

Sea-surface conditions in middle to high latitudes of the North Atlantic

during the last glacial maximum (LGM): The cold paradigma revisited.

Can. J. Earth Sci., 37, 725–750.

Fensome, R.A., Taylor, F.J.R., Norris, G., Sarjeant, W.A.S., Wharton, D.I.,

and Williams, G.L., 1993. A classification of living and fossil dinofla-

gellates. Micropaleontology, 7,1–351 (Special Publication).

Matthiessen, J., and de Vernal, A. (eds.), 2001. Marine dinoflagellate

cysts and high latitude Quaternary paleoenvironmental reconstructions.

J. Quaternary Sci., 16(7), 595–751.

Rochon, A., de Vernal, A., Turon, J.-L., Matthiessen, J., and Head, M.J.,

1999. Distribution of dinoflagellate cyst assemblages in surface sedi-

ments from the North Atlantic Ocean and adjacent basins and quantita-

tive reconstruction of sea-surface parameters. Special Contribution

Series of the American Association of Stratigraphic Palynologists,

35, Dalas, TX, pp. 1–152.

Vink, A., Zonneveld, K.A.F., and Willems, H., 2000. Distributions of cal-

careous dinoflagellate cysts in surface sediments of the western equator-

ial Atlantic Ocean, and their potential use in paleoceanoography. Mar.

Micropaleontol., 38, 149–180.

Zonneveld, K.A.F., and Marret, F., 2003. Atlas of the global geographic

distribution of modern organic-walled dinoflagellate cysts. Rev. Palaeo-

bot. Palynol. 125,1–200.

Cross-references

Coral and Coral Reefs

Lacustrine Sediments

Marine Biogenic Sediments

Palynology

DOLE EFFECT

In 1935–1936, Dole (Dole, 1935) and Morita (Morita and Titani,

1936) discovered independently that the isotopic composition

of atmospheric oxygen (

) is enriched relative to ocea-

nic water. This isotopic enrichment of molecular oxygen in

the heavier

O was named the Morita-Dole effect or just the

Dole effect (DE in the following). The exact value of the enrich-

ment was badly constrained for a long time after the discovery

of the DE. It was not until 1972 that Kroopnick and Craig

(Kroopnick, 1987) determined the DE with sufficient precision.

Expressed relative to the Vienna Standard of Mean Ocean

Water (V-SMOW) (Craig, 1961), they identified a value for

the modern DE of

Atm

¼ð

OÞ

Atm

=ð

OÞ

VSMOW

 1

¼ 23:5  0:1‰:

However, the reasons for this relatively strong enrichment

of 23.5% were not clear for a long time and still are disputed

to some extent. The first assumption was that photosynthesis

produces an enriched O

flux; however, measurements did not

confirm this hypothesis (Ruben et al., 1941; Dole and Jenks,

1944). Subsequently, various other hypotheses have been formu-

lated: photochemical isotopic exchange between O

and H

(Roake and Dole, 1950) or between O

and CO

in the strato-

sphere (Dole et al., 1954), or, alternatively, preferred consump-

tion of the light

O of dissolved oxygen by marine organisms

(Rakestraw et al., 1951) or by soil bacteria during respiration pro-

cesses (Dole et al., 1947). In each case, the measured enrichment

due to the different processes could not confirm the correspond-

ing hypothesis. Lane and Dole (1956) demonstrated convin-

cingly that respiration processes in the ocean are indeed at least

one important contribution to the DE. They measured a progres-

sive isotopic enrichment of dissolved oxygen in oceanic water

samples due to remineralization of organic material and demon-

strated therewith that the biochemistry of plants must be studied

to determine the origin of the DE.

Bender et al. (1994a) gave a detailed overview of current

knowledge of processes contributing to the DE (see Figure D42

for a scheme of the global

O budget). Photosynthesis and respira-

tion are the principal global O

fluxes. Photosynthetic production

of O

tags the outgoing O

flux with the isotope signal of ambient

water without any further fractionation (Guy et al., 1993). In the

marine biosphere, the water of the euphotic surface layer has an

isotopic composition of approximately 0% globally. In the terres-

trial biosphere, the isotopic signature of water adjacent to photo-

synthetic production is characterized by large geographical and

temporal variability. FollowingFarquhar et al. (1993),Benderet al.

(1994a) opted for a global value of 4.4%. This isotope signal of

ambient water in leaves is significantly enriched as compared to

oceanic surface waters and is mainly responsible for the different

contribution to the DE by the terrestrial and the marine biosphere.

The most important contribution to the high value of the DE

takes place during autotrophic and heterotrophic respiration.

Various respirative O

fluxes (photorespiration, Mehler reac-

tion, cyanide-sensitive and cyanide-insensitive pathways in

dark respiration) fractionate

O against

O differently, but

the mean fractionation is close to 20% in both the terrestrial

and the marine environment. In the ocean, part of the respira-

tion takes place in the deep ocean, isolated from the atmo-

spheric oxygen reservoir. This slightly reduces the overall

marine contribution to the DE due to a fraction of oxygen that

fractionates against an already enriched oxygen pool.

Furthermore, there is a small temperature dependent fractio-

nation during the dissolution of O

in water (0.7%) (Benson

and Krause, 1984). Equally, a small fractionation (0.4%) takes

place in the stratosphere during the photochemical exchange

between CO

and O

. When adding up all these different fractio-

nations, weighted with the corresponding fluxes, Bender et al.

(1994a) computed a global equilibrium of d

Atm

= 20.8%,

underestimating by approximately 3% relative to the observed

value. Besides a number of uncertainties in the considered pro-

cesses, there might also be some processes neglected in this

global budget such as fractionation during the diffusive mixing

of O

in the upper soil column (Angert et al., 2001).

There are a number of applications of the DE (or of

the d

Atm

signal respectively) in paleoclimate studies. Past

Atm

can be reconstructed by analyzing the isotopic compo-

sition of air bubbles enclosed in the deep ice from Greenland

282 DOLE EFFECT

and Antarctica. The d

Atm

signal is globally uniform within

today’s instrumental precision. Its variations in the past can

therefore be used to synchronize ice cores from the Northern

and Southern Hemisphere (Sowers et al., 1991; Bender et al.,

1994b). Furthermore, the d

Atm

signal was shown to be

remarkably in phase with precessional insolation variations

during the last four glacial/interglacial cycles (Petit et al.,

1999). This observation was used to establish a more precise

dating of polar ice cores (Parrenin et al., 2000; Shackleton,

2000). Modeling the DE also offers the potential of quantita-

tively constraining marine and terrestrial productivity in the

past (Leuenberger, 1997; Hoffmann et al., 2003).

Georg Hoffmann

Bibliography

Angert, A., Luz, B., and Yakir, D., 2001. Fractionation of oxygen isotopes

by respiration and diffusion and its implications for the isotopic compo-

sition of atmospheric O

. Global Biogeochem. Cycles, 15(4), 871–880.

Bender, M., Sowers, T., and Labeyrie, L., 1994a. The Dole effect and its

variations during the last 130,000 years as measured in the Vostok ice

core. Global Biogeochem. Cycles, 8, 363–376.

Bender, M., Sowers, T., Dickson, M.-L., Orchado, J., Grootes, P.,

Mayewski, P.A., and Meese, D.A., 1994b. Climate correlations between

Greenland and Antarctica during the past 100,000 years. Nature, 372,

663–666.

Benson, B., and Krause, D., 1984. The concentration and isotopic

fractionation of oxygen dissolved in fresh water and

seawater in equilibrium with the atmosphere. Limnol. Oceanogr., 318,

349–352.

Craig, H., 1961. Standard for reporting concentrations of deuterium and

oxygen-18 in natural water. Science, 133, 1833–1834.

Dole, M., 1935. The relative atomic weight of oxygen in water and in air.

J. Am. Chem. Soc., 57, 2731.

Dole, M., and Jenks, G., 1944. Isotopic composition of photosynthetic

oxygen. Science, 100, 409.

Dole, M., Hawkins, R.C., and Barker, H.A., 1947. Bacterial fractionation of

oxygen isotopes. J. Am. Chem. Soc., 69, 226–228.

Dole, M., Lane, G.A., Rudd, D.P., and Zankelies, D.A., 1954. Isotopic

composition of atmospheric oxygen and nitrogen. Geochim Cosmochim

Acta, 6,65–78.

Figure D42 Scheme of the modern Dole effect according to Bender et al. (1994b). Three different oxygen consuming respirative fluxes (dark

respiration, photorespiration, Mehler reaction) contribute to the total ecosystem respiration, R

eco

, each with a different specific fractionation e.

This oxygen consuming process is exactly balanced by oxygen production during photosynthesis tagged isotopically with the isotopic

composition of leaf water d

Leaf

. The total terrestrial DE is 22.4% with a terrestrial O

flux of 20.4 Pmol yr

1

. Respirative fractionation in the

ocean is slightly smaller than on land mainly due to respiration of already enriched oxygen in the deep and intermediate ocean. The

photosynthetically produced O

flux has an isotopic signal of approximately 0% resulting in an oceanic contribution to the DE of 18.9% with

an oceanic productivity of 12 Pmol yr

1

. The steady state solution of this best estimate is 20.8% compared to an observed DE of 23.5%.

DOLE EFFECT 283

Farquhar, G.D., Lloyd, J., Taylor, J.A., Flanagan, L.B., Syvertsen, J.P.,

Hubick, K.T., Wong, S.C., and Ehleringer, J.R., 1993. Vegetation

effects on the isotope composition of oxygen in atmospheric CO

Nature, 363, 439–443.

Guy, R.D., Fogel, M.L., and Berry, J.A., 1993. Photosynthetic fractionation

of the stable isotopes of oxygen and carbon. Plant Physiol. 101,37–47.

Hoffmann, G., Cuntz, M., Weber, C., Ciais, P., Friedlingstein, P., Heimann,

M., Jouzel, J., Kaduk, J., Maier-Reimer, E., Six, K., and Seibt, U.,

2003. A model of the Earth’s Dole effect. Global Biogeochem. Cycles,

Kroopnick, P.M., 1987. Oxygen 18 in dissolved oxygen. In Oestlund, G.,

Craig, H., Broecker, W.S., and Spencer, D. (eds.), GEOSECS Atlantic,

Pacific and Indian Ocean Expeditions. Washington D.C.: U.S. Govern-

ment Printing Office, pp. 27–182.

Lane, G.A., and Dole, M., 1956. Fractionation of oxygen isotopes during

respiration. Science, 123, 574–576.

Leuenberger, M., 1997. Modeling the signal transfer of seawater d

Otothe

O of atmospheric oxygen using a diagnostic box model for the terres-

trial and marine biosphere. J. Geophys. Res., 102, 26841–26850.

Morita, N., and Titani, T., 1936. Uber den Unterschied in der Isotopenzu-

sammensetzung von Luft- und Wassersauerstoff. Bull. Chem. Soc.

Japan, 11,36–38.

Parrenin, F., Jouzel, J., Waelbroeck, C., Ritz, C., and Barnola, J.-M., 2000.

Dating the Vostok ice core by an inverse method. J. Geophys. Res.

Petit, J.R., Jouzel, J., Raynaud, D., Barkov, N.I., Barnola, J.M., Basile, I.,

Bender, M., Chappellaz, J., Davis, J., Delaygue, G., Delmotte, M.,

Kotlyakov, V.M., Legrand, M., Lipenkov, V., Lorius, C., Pépin, L.,

Ritz, C., Saltzman, E., and Stievenard, M., 1999. 420,000 years of cli-

mate and atmospheric history revealed by the Vostok deep antarctic ice

core. Nature, 399, 429–436.

Rakestraw, N.W., Rudd, D.P., and Dole, M., 1951. Isotopic composition of

oxygen in air disolved in Pacific ocean water as a function of depth.

J. Am. Chem. Soc., 73, 2976.

Roake, W.E., and Dole, M., 1950. Oxygen isotope exchange in the electric

discharge. J. Am. Chem. Soc., 72,36–40.

Ruben, S., Randall, M., Kamen, M., and Hyde, J.L., 1941. Heavy oxygen

(

O) as a tracer in the study of photosynthesis, J. Am. Chem. Soc.,

63, 877–879.

Shackleton, N.J., 2000. The 100,000-Year Ice-Age Cycle Identified and

Found to Lag Temperature, Carbon Dioxide, and Orbital Eccentricity.

Science, 289, 1897–1902.

Sowers, T., Bender, M., Raynaud, D., Korotkevich, Y.S., and Orchado, J.,

1991. The

O of atmospheric O

from air inclusions in the vostok

ice core: Timing of CO

and ice volume changes during the penultimate

deglaciation. Paleoceanography, 6, 679–696.

Cross-references

Ice cores, Antarctica and Greenland

Isotope Fractionation

Oxygen Isotopes

DRUMLINS

A drumlin is a subglacial landform that appears as a distinctive

streamlined hill of sediment in the postglacial landscape. The

term is derived from a Gaelic word meaning a mound or

rounded hill. Drumlins vary significantly in size, shape and

composition, and their origins remain controversial.

Drumlins are streamlined in the direction of ice movement,

with a degree of elongation typically about 3:1 but sometimes

as much as 60:1. At higher elongations drumlins grade into

megaflutes. The up-glacier (stoss) end is typically the steeper

and blunter while the downstream (lee) face is more gently slop-

ing and more sharply pointed in plan form. Drumlins may occur

singly or in “swarms” containing hundreds or thousands of hills.

The largest drumlins reach 50 m in height and may be 20 km

long, while the smallest examples are only 10 m long and grade

into other classes of smaller streamlined features such as flutes.

Small drumlins can be superimposed on larger ones.

The composition of drumlins is variable but typically domi-

nated by massive diamicton usually interpreted as till. Some

drumlins contain a rock core and many include coarse stratified

glaciofluvial deposits. The material is often highly deformed.

Several different origins for drumlins have been proposed

and it is likely that different drumlins are formed in different

ways. Some models suggest that drumlins are remnants of sub-

glacial erosion of previously deposited sediments. It has been

suggested that erosion was accomplished by huge subglacial

“megafloods.” Other models suggest that drumlins are deposi-

tional features reflecting either subglacial lodgment or meltout.

Many theories propose that the key processes are deforma-

tional. A widely accepted model is that drumlins are created

by differential movement of material within a deforming sub-

glacial sediment layer. The drumlins reflect areas of less mobile

sediment while intervening areas reflect more rapidly moving

material. Stratified sediments in some drumlins have been

interpreted as lee side stratification sequences in water-filled

cavities on the downstream end of drumlins. The flow of water

through subglacial sediment under varying pressures, and the

deformation of material in response to both glacial and water

pressure are central to much current thinking about drumlins.

It has been suggested that drumlins form in a narrow belt

only 20–30 km wide close to an ice-sheet margin. As the mar-

gin retreats, the zone migrates, and extensive fields of drumlins

can be created in a time-transgressive sequence. It has also been

suggested that drumlin fields mark the positions of former ice

streams. Although drumlins are possibly the most intensively

studied glacial landform, our understanding of them remains

incomplete. However, it is clear that they have the potential to

provide important information about former ice sheets both

through their geographic distribution and through their sedimen-

tology. Useful reviews are provided by Benn and Evans (1996),

Bennett and Glasser (1996) and Hambrey (1994).

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:

John Wiley and sons, 364pp.

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

296pp.

Cross-references

Basal ice

Diamicton

Glacial geomorphology

Glaciofluvial sediments

Late Quaternary megafloods

Tills & tillites

DURICRUSTS

Duricrusts are near-surface geochemical crusts formed as a

result of low-temperature physicochemical processes operating

within the zone of weathering that lead to the absolute or

relative accumulation of minerals through the replacement

284 DRUMLINS