Odekon M. Encyclopedia of paleoclimatology and ancient environments

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measuring Fe contents and Fe lossesduring weatheringof Precam-

brian paleosols. These investigations have been further expanded

to include other redox-sensitive elements, as well as the ratio of

oxygen demand to demand for CO

during weathering (Rye and

Holland, 2000). Other research directions include systematic stu-

dies of root diameter, depth, and density in paleosols, using root

traces, and relating these changes to development of soil morphol-

ogy (Driese and Mora, 2001), widespread deposition of black

shales, and carbon sequestration.

New research directions on pre-quaternary paleosols

Study of pre-Quaternary paleosols appears to be heading in

many future directions. One important research area involves

studies of soil chronosequences (a suite of related soils in

which all soil-forming factors are held constant except for

the time duration of pedogenesis) in order to interpret the

time significance of paleosols in the geologic record (Stiles

et al., 2003). Because paleosols represent unconformities

within otherwise conformable stratigraphic successions, the

ability to resolve time in paleosols is extremely important in

terms of interpreting earth history. U-Pb age-dating of pedo-

genic carbonates offers a promise of improved chronology of

pre-Quaternary paleosols, whose ages place them beyond the

limits of

C or other dating methods commonly used for Qua-

ternary paleosols. Applications of paleosol research to interpre-

tations of sequence stratigraphy will help to connect time and

sea level-base level relationships between marine and terrestrial

stratigraphic sections. Careful documentation of continental

traces of invertebrate animals (paleoichnological studies), espe-

cially those relating soil animals and their relationships to soil

forming-processes in the past, is yet another area that is cur-

rently developing.

Steven G. Driese

Bibliography

Caudill, M.R., Driese, S.G., and Mora, C.I., 1996. Preservation of a paleo-

Vertisol and an estimate of Late Mississippian paleoprecipitation. Jour.

Sed. Res., 66,58–70.

Cerling, T.E., 1984. The stable isotopic composition of modern soil carbonate

and its relationship to climate. Earth Planet. Sci. Letters, 71, 229–240.

Cerling, T.E., 1991. Carbon dioxide in the atmosphere: Evidence from Cen-

ozoic and Mesozoic paleosols. Am. J. Sci., 291, 377–400.

Driese, S.G., and Foreman, J.L., 1992. Paleopedology and paleoclimatic

implications of Late Ordovician vertic paleosols, southern Appala-

chians. J. Sed. Petrol., 62,71–83.

Driese, S.G., and Mora, C.I., 2001. Diversification of Siluro-Devonian

plant traces in paleosols and influence on estimates of paleoatmospheric

levels. In Gensel, P.G., and Edwards, D. (eds.), Plants invade

the Land: Evolutionary and Environmental Perspectives. New York:

Columbia University Press, pp. 237–253.

Driese, S.G., Mora, C.I., Stiles, C.A., Joeckel, R.M., and Nordt, L.C., 2000.

Mass-balance reconstruction of a modern Vertisol: Implications for

interpretations of geochemistry and burial alteration of paleo Vertisols.

Geoderma, 95, 179–204.

Mack, G.H., James, W.C., and Monger, H.C., 1993. Classification of paleo-

sols. Geol. Soc. Am. Bull., 105, 129–136.

Mora, C.I., Driese, S.G., and Colarusso, L.A., 1996. Middle to Late Paleo-

zoic atmospheric CO

from soil carbonate and organic matter. Science,

271, 1105–1107.

Mora, C.I., and Driese, S.G., 1999. Palaeoclimatic significance and stable

carbon isotopes of Palaeozoic red bed paleosols, Appalachian Basin,

USA and Canada. In Thiry , M., and Simon-Coinçon, R. (eds.), Palaeo-

weathering, Palaeosurfaces and related Continental Deposits. International

Association of Sedimentologists Special Publication No. 27, pp. 61–84.

Retallack, G.J., 1988. Field recognition of paleosols. In Reinhardt, J., and

Sigleo, W.R. (eds.), Paleosols and Weathering Through Geologic Time.

Boulder, CO: Geological Society of America. GSA Special Paper 216,

pp. 1–20.

Retallack, G.J., 1991. Untangling the effects of burial alteration and ancient

soil formation. Ann. Rev. Earth Planet. Sci., 19, 183–206.

Retallack, G.J., 2001. Soils of the Past: An Introduction to Paleopedology,

(2nd ed.). Oxford: Blackwell Science, 416pp.

Rye, R., and Holland, H.D., 2000. Geology and geochemistry of paleosols

developed on the Hekpoort Basalt, Pretoria Group, South Africa. Am. J.

Sci., 300,85–141.

Sheldon, N.D., Retallack, G.J., and Tanaka, S., 2002. Geochemical climo-

functions from North American soils and application to paleosols

across the Eocene-Oligocene Boundary in Oregon. J. Geol., 110, 687–696.

Stiles, C.A., Mora, C.I., and Driese, S.G., 2001. Pedogenic iron-manganese

nodules in Vertisols: A new proxy for paleoprecipitation? Geology, 29,

943–946.

Stiles, C.A., Mora, C.I., Driese, S.G., and Robinson, A.C., 2003. Distinguish-

ing climate and time in the soil record: Mass-balance trends in Vertisols

from the Texas Gulf Coastal Prairie. Geology, 31, 331–334.

Cross-references

Atmospheric Evolution, Earth

Carbon isotope Variations Over Geologic Time

Carbon Isotopes, Stable

Early Paleozoic Climates (Cambrian-Devonian)

Geochemical Proxies (non-isotopic)

Late Paleozoic Paleoclimates (Carboniferous-Permian)

Mesozoic Climates

Mineral Indicators of Past Climates

Oxygen Isotopes

Paleoclimate Proxies, an Introduction

Paleo-Precipitation Indicators

Paleosols–Quaternary

Stable Isotope Analysis

Figure P33 Estimates of middle to late Paleozoic atmospheric CO

levels (expressed as times present atmospheric level) derived from

Appalachian basin paleosols (closed symbols from Mora et al., 1996;

other sources (patterned symbols) of pedogenic carbonate-based

estimates cited in Driese and Mora, 2001), calculated using the soil

carbonate carbon isotope paleobarometer (Cerling, 1991). Other

proxy-based estimates include stomatal density (open symbols) and

carbon mass-balance model (line; see discussion in Driese and

Mora (2001)).

PALEOSOLS, PRE-QUATERNARY 751

PALEOSOLS, QUATERNARY

Introduction

Paleosols, the remains of former soil covers of the Earth now

buried under younger deposits, are an important testimony of the

environmental conditions of ancient epochs. Paleosols occur in

sediment sequences of various origins, such as alluvial sediments

of river terraces, and in shallow marine (lagoon), volcanic, and

even glacial deposits. However, the most widespread and best-

preserved paleosols occur in loess deposits. The buried soils are

detected on all continents except for the Antarctic. The first half

of the twentieth century was the period of upbuilding data about

paleosols in loess sequences. It was established that paleosols

contain very important information on chronostratigraphy and

paleogeography.

Paleosols as a chronostratigraphic component

of the Pleistocene

The horizons of paleosols are widely used as key horizons for

the stratigraphic subdivision of continental sedimentary series.

Frequently located in the periglacial zone and occupying an inter-

mediate position between regions of moraine distribution on the

one hand, and marine deposits on the other hand, horizons of

buried soils represent an important link in a system of general cor-

relation of events in glacial, periglacial, and maritime areas.

The reliability of usage of paleopedologic data has increased,

especially with the development of methods of absolute and rela-

tive dating: radiocarbon, thermoluminescent, paleomagnetic,

amino acid, and paleontological (palynological and paleofaunis-

tic). The use of paleosols as components of the stratigraphic record

has increased, especially after paleosols morphotypical features of

different ages were studied (Figure P34).

Paleosols in various sediment sections were once compound

units of soil covers of different ages. The possibility of their use

as a reliable tool for chronostratigraphic subdivision and correla-

tion is based upon the fact that each of the interglacial epochs of

the Quaternary soil formation was characterized by intrinsic fea-

tures of pedogenesis, combinations of paleosols in automorphic

conditions, and specific features of their geographic distribution.

These include, for example, soil complexes that belong to the

late Pleistocene (Sangamon) interglacial in North America, to

the Eemian interglacial in West Europe, to the Mikulino (East

Europe), and to the Kazantsevo interglacial in Siberia. Such

paleosols were named ‘‘soil stratigraphic units’’ or ‘‘pedostratigra-

phic units’’ (Catt, 1990). For these, R. Morrison (1967) has sug-

gested the term ‘‘geosol.’’ Not only interglacial soils can be used

as stratigraphic markers. Interstadial Pleistocene soils having

regional names (e.g., Kesselt, Gleina, Stilfrid B, and Mende F in

West Europe; the Bryansk soil in East Europe; Farmdale in North

America) with established genetic connections and age control

based on radiocarbon dating (32–24 ky

BP) can also be used for

distant correlations.

Figure P34 The Pleistocene sequence of the East European Plain.

752 PALEOSOLS, QUATERNARY

Diagnostic methods of paleosol genesis

The main aim of paleopedological studies is to reconstruct genetic

paleosol profiles and to identify the soil formation factors. The

structure of a soil profile can therefore be used for paleopedologic

reconstructions.

Consequently, analysis of a paleosol in the field starts with

a detailed description of its genetic profile on well-cleaned

walls of natural exposures – those of open pits, ravines, river

banks, marine coasts, and others.

Paleosols contain much more information on the relevant pro-

cesses, which occurred at different times and affected not only the

formation of the soil profiles, as such, but also their preservation.

We can distinguish the following main processes responsible for

the formation of the buried soils observed in loess-paleosol for-

mations (where the best preserved paleosols are found):

1. Processes of soil formation, which took place during periods

of intensive pedogenesis in warm semicycles of glacial-

interglacial cycles. Their direction and intensity changed in

response to different environmental-climatic conditions.

2. Processes specific to transitional stages to cold semicycles and

to glacial ones. These specific processes, which took place at

the end of or at the completion of soil formation during warm

semicycles, greatly affected the preservation of soil of profiles,

especially their upper parts that were marginal to the overlap-

ping loess horizons. Cryogenic processes that contributed to

distortion of soil horizons include processes of sedimentation

intensification, when intermediate deposits transitional from

fossil soils to loess horizons were formed between loess and

soil, and removal and re-deposition processes that affected

materials in the upper parts of fossil soils located on slopes (for-

mation of so-called pedosediments).

3. Pedometamorphic (after Gerasimov, 1971) or diagenetic

processes that transformed the buried fossil soils. As a result

of these processes, fossil soils lost their indices of low stabi-

lity and acquired new secondary ones.

In studying the genesis of fossil soils, the greatest attention is

naturally paid to indices of pedogenesis stages, which reflect con-

ditions extant when the soils proper were formed. These processes

lie at the base of paleopedologic reconstructions. However, study

of the different forms of disturbances of the soil profile and of ped-

ometamorphic changes in their properties is also very important.

The set of analytic methods that we use in our studies

depends on the preservation of the most stable indices of soil

formation. They include the following:

Analytical methods for studying the mineral componen ts

of soils

Data from chemical and physical-chemical analyses, including

particle-size (granulometric) analysis, total chemical composition,

clay minerals, and others can be used for diagnostic purposes

(Figure P35a and P36a). In a number of cases, however,

when fossil soils have been subjected to the impact of secondary

Figure P35 (a) Analytical data of Mezin paleosol complex, Mezin section, Dnieper basin; (b) Clay illuviation coating in Bt horizon.

Figure P36 (a) Analytical data of the Bryansk interstadial paleosol, Bryansk section, Dnieper basin; (b) Ooid type of microstructure (humus horizon).

PALEOSOLS, QUATERNARY 753

cryogenesis and fracturing, their genetic horizons may appear to

be mixed, and the data of the distribution of mineral components

along their profiles may be greatly distorted.

Analytical methods for investigating organic matter

The composition and characteristics of organic matter in the

soil are the most important diagnostic criteria of soil formation

processes. Soils show differences in their humus content depend-

ing on natural climatic conditions. The fractional composition of

soils, the properties of their humic acids, and the regularities of

humus distribution along this profile all change as well.

When the soil is buried, the amount of organic matter it

contains decreases noticeably in spite of the fact that the soil

preserves its dark color. Analyses of radiocarbon data on the

ages of different fractions of humus indicate that fractions of

the humic acids proper and the humic acids of humins are

the most of stable and the least subject to pedometamorphic

transformation (Gerasimov and Chichagova, 1971). Numerous

studies have confirmed the hypothesis, first advanced by I. P.

Gerasimov, who suggested that the correlation of the main

groups of humic matters (humic and fulvic acids, humins) in

fossil soils generally reflects two main directions in the forma-

tion of paleosols: the forest and steppe types.

The data on some properties of humic acids (optical density,

stability to electrolytes) are also used as diagnostic indicators

of former pedogenesis. This was demonstrated for the first time

in the study of humic acids in soils of the Bryansk interval and

of the Mikulino interglacial (Morozova and Chichagova, 1979).

Micromorphological method

This method, developed by W. L. Kubiena (1970), is widely

used for the study of fossil soils in different countries (Fedoroff,

1971; Bronger, 1972; Smolikova, 1972; Konecka-Betley, 1979;

Morozova, 1981; Matviishina, 1982; Mucher and Morozova,

1983; Catt, 1990; and others). The micromorphological method

permits us to study those details of the soil profile that contain

the basic information both on processes of the old pedogenesis,

and on those that resulted from soil fossilization and that disturbed

its overall structure.

As micromorphological diagnostics of soils, the following are

used: data about microstructure (aggregation and voids), grade

of pedality (degree of aggregation; Bullock et al., 1985), ground-

mass (b-fabric), pedofeatures, and other features (Figure P35b

and P36b).

To define paleosols, the paleoactualistic principle is applied:

paleosol properties are compared to the properties of modern soils,

and their genesis is inferred on the basis of their commonality.

Micromorphological research using these materials is presented

in the works of Kubiena (1970) and Bullock et al. (1985). Genetic

types of paleosols are named according to International (FAO)

system and its correlation with the Russian system of soil classifi-

cation (Gerasimova et al., 1996).

Evolution of soil formation in the Quaternary period

Quaternary types of soil formation epochs

The difference in heat and moisture supply between phases of the

climatic macrocycles of the Pleistocene controlled succession of

soil formation types, and various patterns of soil zonality, under

changing conditions of sedimentation and lithogenesis through

macrocycles.

Interglacial soil formation. This soil type developed during

the thermal optimum of the climatic macrocycle. Interglacial

paleosols either corresponded by their genesis to modern soils in

the considered territory or were formed against the background

of a warmer and wetter climate. At temperate latitudes, over the

vast spaces of large watersheds, a latitudinal soil zonality was

developed. In the late and middle Pleistocene, cambisols, luvic

cambisols, stagnic podzol-luvisols (pseudogleys), luvisols, phaeo-

zems, and chernozems were formed. During the early Pleistocene

warm epoch (Cromer), paleosols with subtropical features of

pedogenesis were formed.

Interstadial soil formation. During the cold semicycle,

drastic landscape-climatic changes occurred. These were

accompanied by climatic oscillations of heat and moisture

supply, both of a stadial and interstadial nature, as reflected in

the processes of soil formation and sedimentation.

Interstadial soil formation was characterized by two main

types of paleosols. The first was typified by predominance of

paleosols of the cryo-gley genesis (gelic gleysols). For example,

in the Bryansk (Farmdale) epoch, the gleyezems were widespread

in west and middle Europe, on the southwestern margins of the

East European Plain, and in some regions of North America.

Gleyezems were characterized by the development of gleyzation

processes (i.e., water logged, anaerobic conditions), weakly devel-

oped pedality, and massive microstructures. Paleosols with aggre-

gates of ooid type have been identified in the central regions of the

East European Plain.

The second type differed from the interglacial paleosols by

greater uniformity in the manifestation of soil-forming processes,

with humus accumulation predominant among them. Pedality

was moderately developed, with aggregation and plasma clay-

humus. The latitudinal zonality was then essentially simplified,

becoming nearly hyperzonal. Interstadial soils of early glacial

stages (such as the Krutitsa soil of the Mezin complex) belong

to that type.

Pleniglacial soil formation corresponded to the maximum of

the cold phase of a semicycle. During this time, permafrost was

widespread beyond the ice-sheet in the periglacial zone. Sedi-

mentation was dominant over pedogenesis. Overall, the biogeo-

chemical processes of rock breakdown were overwhelmed, and

specific synlithogenic paleosols (loesses) were formed. These

processes can be compared to those characteristic for the gray

soils of semideserts, to some extent. Very low-intensity soil for-

mation occurred during slightly wetter phases; it was mani-

fested in the processes of gleyization, segregation of ferric

hydroxides, and discrete humification.

General features of the pedogenesis in the Quaternary

period

The nature of soil formation did not repeat itself, but changed

consistently due to the overall trend of climatic changes. In gen-

eral, interglacial soil formation within the middle latitudes of the

Northern Hemisphere (the present-day temperate belt) changed

from subtropical (Mediterranean) in the early stages of the Pleisto-

cene up to subboreal (boreal) at its termination. These changes

have been established based on genetic types of the interglacial

paleosols known from the loess sections of North America,

Europe, and Asia (West Siberia, and China).

Evolution of soil formation processes in the Quaternary

and their climatic interpretation (based on the example

of the central East European Plain)

At present, soddy podzols and gray forest soils in combination

with leached chernozems of the meadow steppe occur in the

central regions of the East European Plain. The main cli-

matic parameters of this territory are: t

= –10



C, t

=+18



754 PALEOSOLS, QUATERNARY

=+8



C, P

=550–600 mm, where t

= average January

temperature, t

= average July temperature, t

= annual mean tem-

perature, and P

= annual mean precipitation.

The loess-soil series in this region are the most studied

(Figure P37). Here, it is possible to observe changes in interglacial

soil formation during the last million years. The earliest known

interglacial soils are the paleosols of two climatic macro-

cycles prior to the Brunhes-Matuyama (780 ky

BP) paleomagnetic

boundary.

During the period when the oldest fossil soils studied were

formed (the Balashovskaya paleosol, in the Matuyama zone),

the type of pedogenesis with well-expressed subtropical fea-

tures dominated. For this time, two types of paleosol profiles

are established. One of them corresponded to wet subtropical

conditions due to a texture-differentiated structure of the profile

and presence of braunlehm plasma (containing fragments of

illustrated coatings in finely dispersed clay; Kubiena, 1970)in

the Bt horizon. In the second type, the features of rubefication

and ferruginization (reddening of soil due to iron oxide accu-

mulation) of the soil plasma are established.

During the Okatovo interglacial, when the Rzhaksa soil

formed, phaeozems appeared in this region (Figure P34). The

Vorona (Muchkap) interglacial (the late Cromer) had set in

after the Don glaciation. Polygenetic soils from less than two

contiguous paleosols were detected in the soil cover. The soils

of the main stage of soil formation were represented by phaeo-

zems in combination with luvic cambisols. For this stage, the

following climatic parameters (average values) have been

reconstructed based on soil types: t



C, t

= +23



= +12



C, P

= 800–900 mm.

During the main phase of the Inzhavino epoch of soil for-

mation, corresponding to the Likhvin (Holstein) interglacial,

pedogenesis occurred under conditions of the subboreal forest

zone. The forest soils that had a genetically differentiated

profile of eluvial-illuvial type, with traces of surface gleyization,

were widespread (luvic cambisols, with participation of stagnic

Figure P37 Quaternary evolution of central part of East European Plain. Modern soil type analogs: I – gelic gleysols, gelic podzoluvisols;

II – chernozem-like sols of cold steppe and forest steppe; III – luvisols; IV – luvisols with participation cambisols; V–VI – cambisols with participation

of luvisols and stagnic podzoluvisols; VII – phaeozems with participation of cambisols. Area of distribution of modern soil type analogs: 1 – upstream

basin of Lena river; 2, 5, 7, 9 – intermountain depressions of Central Altay; 3 – northern part of West and Central Europe; 4 –North of western Siberia;

6 – upstream basin of Visla river; 8 – upstream basins of Rhine and Danube rivers; 10 – intermountain depression of Danube midstream.

PALEOSOLS, QUATERNARY 755

podzoluvisols). The modern analogs of these fossil soils are now

widespread in West Europe, in the upper reaches of the Rhine

and Danube rivers. The reconstructed climatic parameters for

the Likhvin interglacial based on the paleopedologic data are:



C, t

=+21



C, t

=10



C, P

= –700 to 750 mm.

During the main phase of the Kamenka (Demnitz) intergla-

cial period of soil formation, the forest soils dominated. These

soils can be attributed to luvic cambisols and cambisols. Simi-

lar forest soils are now widespread in West Europe and in

the south of central Europe, in particular, in the upper reaches

of the Vistula River. For the Kamenka interglacial, the

following climatic parameters are reconstructed: t

= –1



= +21



C, t

=+9



C, P

= 600–700 mm.

Apparently, it is possible to consider the tundra-gley soil or

gelic gleysols as a modern analog of soils of the Romny inter-

val (Dnieper glaciation; Figure P32). The soils have a profile

A1g-Cg. Near the upper contact of the Romny soil, the

festoon-like cryogenic deformations are described. Their ana-

logs can be found in similar formations of the modern active

layer of permafrost in the western Yamal Peninsula. It is possi-

ble to assume that the modern analogs of the Romny intersta-

dial paleosols lay within the limits of the West Siberian

tundra. The climatic parameters of this interstadial based

on paleopedologic data average: t

= –26



C, t

= +13



= –6



C, P

= 250–350 mm.

At the base of the late Pleistocene loess-soil series, the poly-

genetic Mezin complex occurs. Its early phase (Salyn) corre-

sponds to the Mikulino (Eemian, Sangamon) interglacial,

the late phase to the Krutitsy (Brorup) interstadial. The soils

of the Mikulino (Eemian) interglacial were formed in condi-

tions corresponding to the forest zone of the subboreal belt.

They are represented by paleosols with the distinct texture-

differentiated profile A-E-Bt-C. The Bt horizon is characterized

by abundance of illuviation coatings, testifying to participation

of inwashing processes (illuviation of clay) in the soil forma-

tion. As modern analogs of the Mikulino interglacial soils,

one can consider luvisols, which comprise the present soil

cover in the north of central and middle Europe. Climatic

parameters of the Mikulino interglacial based on the paleope-

dological data are as follows: t

=+1



C, t

= +17



=+9



C, P

= 600–650 mm. For this interglacial, small-

scale maps of the soil cover of Europe and of the East

European Plain were compiled (Figure P38).

Soils of the Krutitsy interstadial had a genetic profile A1-

AB-C. The chernozem-like soils under the meadow-steppe

vegetation or open larch woodlands in the intermountain

depressions of Central Altai can be considered as their modern

analogs. The reconstructed climatic parameters are: t

= –7



= +15



C, t

= –0.5



C, P

= 220–450 mm.

It is necessary to note that the interglacial soil formation of

the middle and, probably early Pleistocene, apparently included

the concluding phases of interstadials of subsequent glacia-

tions, when the soils, similar to those of the Krutitsy intersta-

dial, were formed.

In the concluding phase of the Bryansk interstadial (corre-

lated with the middle Valdai, Denekamp interstadial) soil for-

mation took place under the conditions of the polar belt. The

soils are identified as tundra-gley (gelic gleysols), with charac-

teristic features of these paleosols being strong gleyzation of

the entire profile (Ag-Bg-Cg), and a particular ooid aggre-

gation. An area of modern analogs of these paleosols lies

within the upper Lena River basin. The reconstructed climatic

Figure P38 Mikulino interglacial soil map of East European Plain (central and south part); 1 – podzoluvisols; 2 – luvisols of broadleaf forests;

3 – luvic chernozems and gleyic chernozem of forest-steppe; 4 – chernozems of grass steppe; 5 – kastanozems of grass steppe.

756 PALEOSOLS, QUATERNARY

parameters of the Bryansk interstadial (average values) are:

= –23



C, t

= +13



C, t

= –6



C, P

= 150–250 mm.

The peak of the late Valdai glaciation (20–18 ky

BP), corre-

sponding to the Last Glacial Maximum, was the main climatic

minimum of the Cenozoic (Velichko, 1973). At this time, a vast

periglacial zone stretched. Beyond the Scandinavian Ice Sheet, a

considerable portion of this zone (up to 47–48%) was occupied

by permafrost (Velichko, 1982). In this region, peculiar synlitho-

genic paleosols of pleniglacial type occurred. These were chara-

cterized by slow humus accumulation, microaggregation, and

accumulation and re-distribution of secondary carbonates without

their leaching. The reconstructed climatic parameters inferred

from paleocryogenic structures of this time are: t

= –25 to



C, t

=+12



C, t

= –9



C, P

=250–300 mm.

At the Pleistocene/Holocene transition, the permafrost

boundary retreated to 54–55



N. Within the loess periglacial

regions, thermokarst processes of the ground ice melting

took place, along with formation of a specific complex of land-

forms – the relic cryogenic microrelief (RCM) (Velichko,

1973). RCM strongly influenced the structure of the Holocene

soil cover. Within the range of the light-gray and soddy podzo-

lic soils of the central Russian Plain, the small-contour differen-

tiation of soils is represented by the combination of the eroded

hummocky soils in place of polygonal blocks, and the meadow

soils in the troughs and depressions in place of the melted out

ice-veins of the Yaroslavl’ cryogenic stage. The meadow soils

were formed during the optimum phase of the Holocene and

were preserved in the form of the relic second humus horizon

in genetic profiles of modern soddy-podzolic and gray forest

soils (Velichko et al., 1996).

Conclusion

During the Quaternary, abrupt changes took place in the direc-

tion and intensity of pedogenesis, corresponding to glacial-

interglacial climate cycles and general trend. According to

these changes, three main types of soil-formation epochs are

distinguished: interglacial, interstadial, and pleniglacial (gla-

cial). Within the macrocycles, all three types of pedogenesis

occur. Due to the general trend of climate change, the soil

formation process changed from subtropical type in the early

phases of the Pleistocene to polar types in its final stages.

Tatiana D. Morozova and Andrei A. Velichko

Bibliography

Bronger, A., 1972. Zur Mikromorphologie und Palaoboden aus Loss in

Karpatenbecken. In St. Kowalinsky (ed.), Soil Micromorphology. Wars-

zawa, Poland: Naukowe, pp. 607–615.

Bullock, P., Fedoroff, N., Jongerius, A., Stoops, G., and Tursina, T.V.,

1985. Handbook for Soil Thin Sections Description. Wolverhampton,

UK: Waine Research Publications, 152pp.

Catt, J. (ed.), 1990. Paleopedological manual. Quaternary Int., 6,1–95.

Fedoroff, N., 1971. The usefulness of micromorphology in paleopedology.

In Vaalon, D.R. (ed.), Paleopedology. Origin, Nature and Dating

of Paleosols. Jerusalem, Israel: Israel University Press, pp. 156–160.

Gerasimov, I.P., 1971. Nature and originality of paleosol. In Vaalon D.R.

(ed.), Paleopedology. Origin, Nature and Dating of Paleosols.

Jerusalem, Israel: Israel University Press, pp. 15–27.

Gerasimov, I.P., and Chichagova, O.A., 1971. Nekotorie voprosy

radiouglerodnogo datirovaniya gumusa. In Pochvovedenie, vol. 10.

pp. 3–11.

Gerasimova, M.I., Gubin, S.V., and Shoba, S.A., 1996. Soil of Russia and

Adjacent Countries. In Suiedema R., (ed.), Geography and Micromor-

phology. Moskow-Wagenningen, 204pp.

Konecka-Betley, K., 1979. Reliktowe procesy glebotworcze w glebach

wspolczesnych wytworrzonych z gleny zwalowej. Zeszyty Naukowe

Szkoly Glownej Gospodarstwa Wiejskiego Akademii Rolniszej

w Warszawie. Rolnictwo, 18, pp. 78–97.

Kubiena, W.L., 1970. Micromorphological Features of Soil Geography.

New Brunswick, NJ: Rutgers University Press, 255pp.

Matviishina, Zh.N., 1982. Micromorfology of Pleistocene Soils of the

Ukraine. Kiev, Ukraine: Naukova Dumka, p. 174.

Morozova, T.D., 1981. Razvitie Pochvennogo Pokrova Evropy v Pozdnem

Pleistotsene. Moscow, Russia: Nauka, p. 231.

Morozova, T.D., and Chichagova, O.A., 1979. Peculiarities of organoc mat-

ter of fossil soils in connection with genesis and diagenesis.

In: Vassoevich, N., (ed.), Organicheskoe veshestvo v sovremennyh I isko-

poemyh osadkah (sedikakhity), Moscow, Russia: Izd. Mosk. Univ., pp.

68–69.

Morrison, R.B., 1967. Principles soil of Quaternary stratigraphic. In

Morrison, R.B., and Wright, H.E. (eds.), Quaternary Soils. Proc. VII

INQUA Congress, vol. 9, Denvez, Colozado, USA. pp. 1–69.

Mucher, H.I., and Morozova, T.D., 1983. The application of soil micro-

morphology in Quaternary geology and geomorphology. In Bullock

and Murphy, G.P. (eds.), Soil Micromorphology, vol. 1. Berkamsted,

UK: A.B. Acad. Publ., pp. 151–194.

Smolikova, L., 1972. The significance of soil micromorphology for

the solutions of soil evolution in the geology of quaternary. In St.

Kowalinsky, (ed.), Soil Micromorphology. Warszawa, Poland: Naukowe,

pp.543–557.

Velichko, A.A., 1973. Prirodnyi Process v Pleistitsene. Moskow, Russia:

Nauka, p. 265.

Velichko, A.A., 1982. Main features of the last climatic macrocycle and the

present state of environments. In Gerasimov, I.P., and Velichko, A.A.

(eds.), Paleogeography of Europe during the last hundred thousand

years (Atlas-Monograph). Moscow, Russia: Nauka Press, pp. 139–144.

Velichko, A.A., Morozova, T.D., Nechaev, V.P., Poroshnyakova, O.M.,

1996. Paleogenesis, Soil Cover and Agriculture. Moscow, Russia:

Nauka Press, p. 151.

Cross-references

Eemian (Sangamonian) Interglacial

Encyclopedia of Soil Science and Technology

Interstadials

Last Glacial Maximum

Loess Deposits

Pleistocene Climates

Quaternary Climate Transitions and Cycles

Wisconsinan (Weichselian, Würm) Glaciation

PALEOTEMPERATURES AND PROXY

RECONSTRUCTIONS

Introduction

Earth’s global temperature is mainly a function of energy derived

from incoming solar radiation (insolation) of varying wave-

lengths, with a very minor energy input coming from geothermal

processes. About 28% of the total incoming insolation is reflected

back to space by clouds, particles and molecules in the atmo-

sphere, about 25% is absorbed in the atmosphere by ozone, clouds

and water vapor, and about 47% is absorbed at the Earth’ssurface.

Of this 47%, about 18% is reflected to back to space, while the

remainder heats the atmosphere through latentheating(conversion

of water to water vapor), convection, and absorption of infrared

radiation by greenhouse gases. This trapping of solar energy by

the atmosphere is why the Earth’s mean annual temperature is

33



C warmer than it would be otherwise and the main reason

there is concern about future global warming caused by anthropo-

genic greenhouse gas emissions.

PALEOTEMPERATURES AND PROXY RECONSTRUCTIONS 757

Earth’s temperature, however, is not stable. It is governed

by many external and internal processes operating over different

spatial and temporal scales. Changes in the output of energy by

the Sun, in the seasonal and geographic distribution of insola-

tion due to changes in Earth’s orbital configuration (axial tilt,

precession and the eccentricity of its orbit), and in physical

(wind-blown dust, volcanic emissions) and chemical (radia-

tively-active gas concentrations) characteristics of the atmosphere,

as well as tectonic processes altering the continental distribution

and oceanic circulation, all serve to influence temperature. For

example, the oxygen isotopic record of ocean foraminifera

(Figure P39) provides a vivid picture of the progressive cooling

in climate over the past 100 million years, due mainly to the

influence of long-term changes in ocean basin and continental

configurations, atmospheric composition and other factors (Miller

et al., 1987).

Historical and instrumental temperature records are far too

short to understand these and other complex patterns of Earth’s

temperature variability over different timescales, or to serve as

a baseline against which to evaluate current global tempera-

ture trends and human influence on climate. Therefore, one of

the most actively researched topics in paleoclimatology is the

reconstruction and interpretation of past atmospheric and ocean

temperatures. This entry deals with general aspects of Earth’s

temperature history, mainly atmospheric temperature; see also

Ocean paleotemperatures for methods of estimating oceanic

temperatures.

Methods of paleotemperature reconstruction

Paleoclimatologists use a wide variety of temperature “proxies”

preserved in natural archives such as tree rings, ocean and lake

sediments, glaciers and ice sheets (ice cores), speleothems (preci-

pitated calcite), ancient soils (paleosols), corals and other minera-

lized organisms, and geological and geomorphological features

to estimate past changes in temperature (Table P4). Paleotempera-

ture reconstructions can be qualitative or quantitative, depending

on the proxy method used; however, all methods carry a certain

error bar due to analytical uncertainty, incomplete knowledge of

processes governing organic and biogeochemical systems, and

the influence of factors other than temperature. Uncertainties not-

withstanding, paleotemperature estimates have improved mark-

edly in their accuracy and reliability over the past few decades,

and some of the more important applications to reconstruct past

atmospheric temperatures are discussed here.

Tree rings

The study of tree rings, or dendroclimatology (Fritts, 1976;

Schweingruber, 1988), relies on the influence that temperature

and precipitation have on the growth of trees and is most useful

in the study of climate variability over the past millennium.

Tree rings consist of couplets of alternating light and dark

layers of earlywood and latewood growth, respectively, form-

ing seasonal growth patterns that vary year-to-year depending

on environmental conditions, soil, the age of the tree, and other

factors. Tree ring width, density (a function of several aspects

of cell and wood morphology), and the isotopic (carbon, oxy-

gen and hydrogen) composition of the wood are commonly

used to reconstruct climate history especially in middle to

high latitudes. Summer temperature is a particularly important

influence on latewood density in boreal trees. The success of

dendroclimatology hinges on careful selection of sites and trees,

sophisticated statistical analyses to calibrate tree ring charac-

teristics to climatic parameters, extensive cross-calibration of

multiple trees for chronological control, and an understanding

of the ecology and growth of various tree genera. For example,

ring widths have been used to produce annual temperatures in

northern parts of North America for the past few centuries

(D’Arrigo and Jacoby, 1992). Longer temperature records are

available from carbon isotopes in the case of a 2,000-year

record of temperature developed from conifer trees from the

western United States (Stuiver and Braziunas, 1987).

In recent years, researchers have produced extensive net-

works of tree ring records of temperature and other climate para-

meters such as sea level pressure in North America, Europe, and

Asia (Briffa and Matthews, 2002; Cook et al., 2002), with more

Figure P39 Progressive cooling of global climate is illustrated by

changes in the isotopic composition of foraminifera from deep-sea

cores. More negative d

O values characterize the Cretaceous

and Eocene periods of global warmth and minimal polar ice volume;

more positive values during the Oligocene through Quaternary

signify gradual cooling and the build-up of polar ice caps. Modified

from Miller et al. (1987); Cretaceous data from Huber et al. (2002).

Courtesy of Kenneth G. Miller, Rutgers University, Brian Huber,

Smithsonian Institution.

758 PALEOTEMPERATURES AND PROXY RECONSTRUCTIONS

limited data from the Southern Hemisphere. By combining tree

ring data with other temperature proxy records (i.e., ice cores,

lakes) and splicing the proxy records with instrumental records

dating back to the nineteenth century, mean annual hemispheric

and global temperature anomalies and climate variability can be

estimated for the past 2,000 years (Mann and Jones, 2003,

Figure P40). These studies provide convincing evidence that

Northern Hemisphere twentieth century temperatures are among

the warmest in the last 1,000 years, although the density of South-

ern Hemisphere sites is too sparse to make conclusions about that

region. These and other studies also suggest that the interannual

and decadal climate variability during the twentieth century exhi-

bits patterns consistent with the hypothesis that anthropogenic

greenhouse gases have influenced global climate. They also

demonstrate the importance of integrating temperature records

derived from multiple proxy indicators to achieve hemispheric

and global scale atmospheric temperature reconstructions.

Ice core studies of ice sheets and glaciers

Water trapped as ice in glaciers and continental ice sheets pro-

vides a wealth of information on temperature history of high

latitudes and high elevation regions because the isotopic com-

position of the water is heavily dependent on the atmospheric

temperature at time of precipitation. Ice cores from Greenland

and Antarctica recovered paleotemperature records dating back

about 110,000 and 700,000 years, respectively. Of the dozens of

paleoclimate proxies recovered from ice cores, the most widely

used to reconstruct past atmospheric temperatures are the oxygen

(

O ratio) and hydrogen (deuterium =

H) isotopic composi-

tionof the ice. Due to the lower vapor pressures of water molecules

containing heavy isotopes of H (D) and O (

O), water condensed

from vapor is preferentially enriched in heavier isotopes. At cooler

temperatures, there is progressively less enrichment in conden-

sation. The relationship between d

O of precipitation and atmo-

spheric temperature has been expressed in a number of equations

depending on geographic location and other climatic factors

(Jouzel et al., 1987; Rozanski et al., 1993; Bradley, 1999).

Complications with the

O-temperature relationship arise due

to changing moisture source and seasonal variability in precipita-

tion, which can also influence the oxygen isotopic composition

of the water. These processes are particularly important over

longer timescales and researchers must devise means to sort them

out to obtain paleotemperature curves. One approach is to exploit

the relationship between differential fractionation of hydrogen

and oxygen during precipitation, known as the deuterium excess,

or d. During precipitation, a global relationship exists between

hydrogen and oxygen expressed as dD=8d

O+10.The

value + 10 represents an offset between the hydrogen and oxygen

Table P4 Summary of methods used to reconstruct atmospheric temperatures

Paleoclimate archive

Proxy method for temperature Comments and complications Reference

Ice Sheets & Glaciers

(ice cores)

Oxygen and hydrogen isotopes Changes in moisture source and seasonal

precipitation

Jouzel et al. (1997);

Grootes et al. (1993)

Deuterium excess Bradley (1999)

Melt layers Summer only

Lake Sediments and

Peats

Insects (mainly beetles) Elias and Coope (1994)

Fossil pollen Prentice et al. (1991),

Guiot et al. (1989)

Ostracode shell isotopic composition Von Grafenstein et al. (1999)

Ostracode shell trace element

composition

Holmes and Chivas (2002)

Diatoms

Speleothems (stalagtites,

stalagmites, vein

calcite)

Oxygen isotopes Temperature effects during calcite (aragonite)

deposition; composition of regional precipitation

McDermott et al. (2001),

Winograd et al. (1992),

see text

Tree rings

(dendroclimatology)

Ring density and ring width Usually <1,000 years; standardization due to growth Fritts (1976), Schweingruber

(1988)

Carbon, oxygen, hydrogen isotopic

composition

Precipitation and other factors important; complex

hydrological and biological processes during wood

growth

Briffa and Matthews (2002),

Mann et al. (1999),

Mann and Jones (2003)

Geological &

Geomorphological

features

Paleosols (fossil soils) Rainfall has strong influence Retallak (1990)

Glacial sediments High latitudes, elevations

Periglacial sediments Only near glacial margins Washburn (1979)

Lithology & mineralogical

associations

Qualitative Parrish (1998)

Other

Tree lines, elevation

changes

Pollen, macroflora from packrat

middens

Discontinuous, mainly glacial-interglacial timescales Betancourt et al. (1990)

Tree lines, geographic

changes

Pollen Mainly used in Arctic region

Plant macrofossils Plant associations, leaf morphology

General Paleoclimatology texts with discussion of paleotemperature reconstruction include Crowley and North (1991), Parrish (1998), Bradley (1999),

Cronin (1999).

These are either comprehensive texts on the proxy method or important applications.

PALEOTEMPERATURES AND PROXY RECONSTRUCTIONS 759

isotopes and is known as d. As atmospheric conditions vary over

time, the value for d also changes and by measuring both d

and dD in ice cores, a measure of d through time is obtained.

The influence of kinetic processes during both evaporation and

condensation, changing moisture sources and temperature of the

water are complex and need to be sorted out for application of d

to the study of paleotemperature.

Despite these and other complexities, some of the most

important discoveries about Earth’s climate history have come

from oxygen and hydrogen isotopic records from Greenland,

Antarctica, and low latitude alpine glaciers. At the Summit

region of Greenland (mean annual temperature = 31



C), two

large research projects carried out under the auspices of the

European research program called GRIP (Greenland Ice Core

Project), and its American counterpart, GISP II (Greenland Ice

Sheet Project II), have provided excellent isotopic records

of past temperature variability (Dansgaard et al., 1993; Grootes

et al., 1993) that have become a standard against which all late

Quaternary climate records are compared. Figure P41 depicts

the GRIP oxygen isotopic record for the past 100,000 years,

indicating extreme temperature-related isotopic stabi lity during

the Holocene (10,000 years to present) and the existence of more

than 20 millennial-scale climatic oscillations between 100,000

and 15,000 years ago. These events, known as Dansgaard-

Oeschger events, reflect extreme climate instability during the

last glacial period, and they are also recognized in the GISP

II core and many paleoclimate records from around the world.

Since the Greenland ice core isotopic fluctuations were first

described, the interpretation of the temporal variability of the

Greenland d

O-temperature relationship has been improved

based on borehole calibration and other means (Cuffey et al.,

1995; Johnsen et al., 1995; Jouzel et al., 1997) and suggest

a glacial to interglacial temperature difference of as much

as 15



Speleothems and precipitated calcite

Speleothems are rocks deposited when calcite or aragonite precipi-

tate from groundwater, often forming stalactites and stalagmites.

Vein calcite is a type of low temperature (34



C) hydrothermal

deposit precipitated from groundwater. Speleothems and vein

calcite, when sampled at fine sub-millimeter-scale intervals, can

yield important paleoclimate information because oxygen fractio-

nation that occurs during deposition of the calcitic rock is depen-

dent on surface temperature. However, precipitation of source

water is itself temperature dependent and also varies due to kinetic

processes, such that the interpretation of speleothem oxygen iso-

tope records must be carried out carefully, on a case-by-case basis.

For example, in an excellent centennial-scale speleothem record

from Ireland, McDermott et al. (2001) were able to detect rapid

climate shifts in the regional climate, which they attributed to both

the effects of temperature and changes in water vapor sources.

In a series of studies of vein calcite at Devils Hole (DH),

Nevada, Winograd, Landwehr and colleagues (Winograd et al.,

1992, 1997; Landwehr et al., 1997; Landwehr and Winograd,

2001) provided an important, but controversial late Quaternary

paleotemperature record reflecting mainly eastern Pacific sea

surface and western United States atmospheric temperatures

(Figure P42). The DH record is unique because, in contrast to

marine records covering the past 400,000 years, DH is directly

dated by the uranium-series methods, and thus yielded a test of

the Milankovitch theory of climate (see Astronomical theory of

climate change). The DH temperature record shows remarkable

Figure P40 Reconstructed and measured global mean temperatures averaged from tree rings, polar and alpine ice cores, and lake and estuarine

sediments for the last 1,800 years. ( a) Reconstructed temperatures combining Northern and Southern Hemisphere records. Alternative temperature

reconstructions based on different weighting methods given in Mann and Jones (2003) and references therein. (b) Hemispheric temperature

difference based on same reconstructed and instrumental temperature curves. From Mann and Jones (2003); Courtesy of Michael E. Mann,

Pennsylvania State University.

760 PALEOTEMPERATURES AND PROXY RECONSTRUCTIONS