Odekon M. Encyclopedia of paleoclimatology and ancient environments

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tephra layers, and tephrochronologies have become established

for several volcanic regions (Turney and Lowe, 2001). Tephro-

chronology enables correlation between lakes, providing the

tephra layers have distinctive chemical characteristics that

enable them to be tied to a particular eruption. Paleomagnetic

properties including field reversals and secular variations have

been used in some localities to give chronologies, but the wide-

spread use of paleomagnetism as an independent dating tool for

recent lake sediments is not common. Magnetic measurements

have been more extensively used to correlate cores from the same

basin and to trace catchment erosion (Maher and Thompson,

1999). In recent studies, the use of spheroidal carbonaceous

particles (SCPs) from high temperature combustion processes

have been useful in giving a marker horizon, providing the his-

tory of industrialization in the region is known (Smol, 2002).

Virtually all published paleolimnological studies use a chronol-

ogy derived from radiometric methods. Various radionuclides

with different half-lives are routinely used, and choice is depen-

dent on the temporal scale of the study and the materials present

within the sediments. Recent paleolimnology relies heavily on

210

Pb that has a half life of 22 years and is suitable for dating

the last 150–200 years. It is usually partnered with

137

Cs, as this

artificial radionuclide has distinctive peaks in the early 1960s

and 1986, relating to the testing of atomic weapons and the

Chernobyl nuclear accident, respectively. Studies of the last

40,000 years make extensive use of

C measurements to pro-

duce a chronological framework. Radiocarbon measurements

from lake sediments can be made on bulk organic material, char-

coal, or plant macrofossils. The latter are preferred as terrestrial

plants are more likely to be in equilibrium with the atmosphere

and therefore give a more reliable date. Some lakes in karst terrains

are difficult to date using

C because lake biota can incorporate

old dissolved carbonate leading to

C dates that are too old. An

alternative method is U/Th dating on carbonates or occasionally

on silicates. U/Th has great potential in the dating of lake

sediments up to 250,000 years old and is increasingly used by

paleolimnologists.

Paleolimnology and climate change

Lakes have been referred to as “natural archives,” since like

libraries they contain information that can be “read ” if the “lan-

guage” of the particular proxy indicator is known. To continue

the literary analogy, the evidence contained in lake sediments is

like an incomplete collection of books, with some pages miss-

ing, written in an unfamiliar language, and all with different

perspectives on the question under consideration. Therefore,

interpreting different environmental indicators is problematic,

as they are rarely responsive to a single variable, they are

frequently incomplete reflections of former ecosystems, and

are often prone to diagenesis or other corruption. Moreover,

few indicators can be used to directly reconstruct former

climate variables such as precipitation or temperature, and a

surrogate measure of climate such as lake water salinity is often

used. With these caveats this discussion will now briefly con-

sider the types of evidence used by paleolimnologists. It is

not possible to review the full range of indicators used here,

but some of the most common techniques will be introduced.

The water balance of a closed basin lake is determined by

precipitation inputs, evaporation losses, and the net exchange

with groundwater. Open systems also lose water through fluvial

outflows. In theory then, it is possible to reconstruct precipita-

tion minus evaporation (P-E) from the lake water volume,

usually calculated from the lake level, providing groundwater

flux and any outflow component can be estimated. Fossil shor-

elines provide the only absolute method of reconstructing lake

level, but paleolimnological indicators can at least show rela-

tive changes in P-E. An indirect measure of P-E is offered by

salinity changes, since, as the volume contracts in a closed

basin, the salt concentration should increase, assuming other

variables are constant. This relationship is complicated by sev-

eral factors, including changes in brine evolution since different

salts precipitate at different stages of concentration, (see Cohen,

2003), and hysteresis as salt precipitates are redissolved during

lake water dilution.

Diatoms, unicellular algae with silica shells (frustules), are one

of the most ubiquitous autochthonous components preserved in

lake sediments and have a great number of uses in palaeolimnol-

ogy (Stoermer and Smol, 1999). Diatoms are sensitive to changes

in pH, nutrient levels, and salinity. The latter variable is especially

useful in paleolimnological studies of climate change from closed

basin lakes where salinity can approximate P-E providing local

hydrological factors are understood. In the salt lakes of East

Africa, diatoms have revealed long records of salinity change

linked to climate fluctuations. One of the best comes from Lake

Abhé, a closed basin lake in the Afar depression of Djibouti, where

changes in the diatom community are indicative of monsoon

strength and orbital forcing of palaeoclimate during the past

70,000 years (see Gasse et al., 1997). The interpretation of climate

from diatom records is indirect and has to consider a range of

taphonomic and diagenetic changes (Gasse et al., 1997). Similar

data can be obtained using ostracods, millimeter-sized animals

with carbonate shells or carapaces. Ostracods are often abundant

in carbonate-rich lakes where statistical inference models can, as

for diatoms, be used to predict past salinities from their assemblage

composition (Holmes, 2001). Often, ostracods are found in sedi-

ments devoid of diatoms and vice versa. Ostracods also provide

an alternative geochemical approach to salinity reconstruction,

from the ratio of Ca/Sr and Mg/Sr in their carapaces.

Further insights into P-E can be obtained using oxygen iso-

topes. In closed basin lakes, oxygen isotope analysis, the ratio

between

O and

O levels (d

O), can be used to infer shifts

in P-E (Cohen, 2003). In its simplest form, the lighter oxygen

is preferentially evaporated under relatively dry climates, leav-

ing waters enriched in the heavier isotope. Source area, water

temperature, and precipitation amount all influence the d

of lake waters, and a modeling approach is required to get

quantitative climate estimates (Gat, 1995). The isotope ratios

are derived from authigenic carbonates and the shells of organ-

isms such as molluscs and ostracods. Until recently, this

method was not applicable to the large number of lakes where

carbonate was absent, but it is now possible to get reliable

measurements from silica such as diatoms, and plant cellulose

(see discussion in Cohen, 2003). More development work is

necessary to establish fractionation processes and sources of

error in oxygen isotope analysis from non-carbonate hosts but

results do appear promising.

Distinguishing between the major climate variables of preci-

pitation and temperature using paleolimnological indicators is

difficult but can be achievable in some situations. Pollen analy-

sis is one of the best known methods of environmental recon-

struction, including quantitative estimates of climate variables,

and it is usually abundant in anaerobic lake sediments (Bennett

and Willis, 2001). In some cases it is possible to calibrate pol-

len records to reconstruct former precipitation and temperature

values using transfer functions (Bonnefille et al., 1990). It is

not always possible to confidently identify the contribution of

PALEOLIMNOLOGY 741

each of these variables, or to exclude the influence of edaphic

conditions, ecological competition, and human agency on

vegetation. Another widely used method of temperature recon-

struction comes from chironomids (midges). These are extre-

mely good indicators of temperatures, at least in high latitude

regions. Fossil chironomid larvae in lake sediments can be used

to reconstruct paleotemperatures providing temperature is the

dominant variable in the region being investigated (Brooks

and Birks, 2001). The greater abundance and ubiquity of chir-

onomid remains in lake sediments gives them an advantage

over coleoptera, although beetle remains have also yielded

excellent temperature reconstruction.

Changing nutrient cycles in lakes can be reconstructed by

paleolimnologists to better understand ecosystem changes and

also to give insights into climate processes. Carbon isotope analy-

sis (

C, usually d

PDB

) of bulk lake sediments measures

changes in lake productivity and shifts in the source of organic

materials to the lakes. On Mt. Kenya, Street-Perrott et al. (1997),

demonstrated how more positive d

C values occurred in the bulk

organic matter of glacial age sediments because of materials

derived from C

plants. Plants using the C

and CAM photosyn-

thetic pathway are adapted to moisture and CO

stress, as would

occur under relatively arid climates with low pCO

. Conversely,

in Holocene sediments, the d

C values became more negative as

plants using the C

pathway increased, replacing the C

species;

thereby indicating greater moisture and CO

availability. The ana-

lysis has been taken further by measuring compound specific iso-

tope ratios, enabling the distinction of materials originating from

algae, aquatic macrophytes, and terrestrial plants (Cohen, 2003).

At the glacial-interglacial scale, paleolimnological research has

shown that changes in the biogeochemical cycling of carbon is

coupled with major climate shifts.

The cycling of nitrogen within lakes is also of interest to

paleolimnologists given the importance of this element to aqua-

tic productivity, where along with phosphorus it is often a lim-

iting nutrient. Analysis of the

Nto

N ratio (d

AIR

) helps

determine the origin of organic materials (i.e., algae or higher

plants), N-fixation and limitation, NH

volatilisation under

high pH, flushing rates of soil N, dissolved inorganic N abun-

dance, and the stratification regime (Talbot, 2001). Nitrogen

isotopes have been successfully used by paleolimnologists,

with one of the best examples being from Lake Bosumtwi,

Ghana. Here a multidisciplinary study of bulk organic proper-

ties, d

C and d

N has shown limnological changes linked to

regional climatic variability (Cohen, 2003).

Biological and geochemical indicators are most common in

paleolimnological studies of climate change. Physical parameters

can also be used as a proxy of catchment changes that may be

indirectly linked to climate. Probably the most useful methods in

this respect are a cluster of related magnetic techniques that can

discriminate between particular minerals that are indicative of pro-

venance. Some of these minerals are derived from topsoil, others

from subsoil, and some from weathering of bedrock. It is possible

to identify the nature and rate of the material being eroded and also

the area supplying the major sources of sediment in catchments of

mixed geology. These data are useful in a range of environmental

management applications as well as for sediment tracing experi-

ments (Maher and Thompson, 1999).

Conclusions and perspective

Paleolimnology has provided limnologists with time series data

beyond the scope of ecological studies, and has answered ques-

tions about lake ontogeny and ecosystem response to climate that

would not have otherwise been possible. Furthermore, some of the

longest, highest resolution sequences of environmental change are

from the sediments of lakes. Paleolimnological investigations of

climate change have been conducted worldwide, but a particular

contribution has been made in low latitude regions where other

hosts of paleoclimate data are either limited or have proven diffi-

cult to exploit.

Over the last decade, the number of publications in paleolim-

nology has risen enormously (i.e., in the Journal of Paleolimnol-

ogy). Recent developments include high resolution automated

techniques, greater accuracy, and more focused scientific ques-

tions. There is also a trend toward analysis of specific sedimentary

components (e.g., biomarkers) rather than bulk samples, thereby

achieving greater precision in reconstructions. Numerical model-

ing has further improved the importance and application of paleo-

limnology to climate change.

Philip Barker

Bibliography

Bennett, K.D., and Willis, K.J., 2001. Pollen. In Smol, J.P., Birks, H.J.B.,

and Last, W.M. (eds.), Tracking Environmental Change Using Lake

Sediments. Terrestrial, Algal and Siliceous Indicators. Developments

in Paleoenvironmental Research. Dordrecht, The Netherlands: Kluwer,

pp. 5–32.

Bonnefille, R., Roeland, J.C., and Guiot, J., 1990. Temperature and

rainfall estimates for the past ca.40,000 years in equatorial Africa.

Nature, 346, 347–349.

Brooks, S.J., and Birks, H.J.B., 2001. Chironomid-inferred air temperatures

from Lateglacial and Holocene sites in north-west Europe: Progress

and problems. Quat. Sci. Rev., 20, 1723–1741.

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

Systems. New York: Oxford University Press, p. 500.

Dean, W., Anderson, R., Bradbury, J.P., and Anderson, D., 2002. A

1500-year record of climatic and environmental change in Elk Lake,

Minnesota. I. Varve thickness and gray-scale density. J. Paleolimnol.,

27, 287–299.

Gasse, F., Barker, P., Gell, P.A., Fritz, S.C., and Chalié, F., 1997.

Diatom-inferred salinity in palaeolakes: An indirect tracer of climatic

change. Quaternary Sci. Rev., 16, 547–563.

Gat, J.R., 1995. Stable isotopes of fresh and saline lakes. In Lerman, L.,

Imboden, D.M., and Gat, J.R. (eds.), Physics and Chemistry of Lakes.

Berlin: Springer, p. 334.

Holmes, J.A., 2001. Ostracoda. In Smol, J.P., Birks, H.J.B., and Last, W.M.

(eds.), Tracking Environmental Change Using Lake Sediments.

Zoological Indicators, vol 4. Dordrecht, The Netherlands: Kluwer,

pp. 125–151.

Lamoureux, S., 2001. Varve chronology techniques. In Last, W.M., and

Smol, J.P. (eds.), Tracking Environmental Change Using Lake

Sediments. Basin Analysis, Coring, and Chronological Techniques.

Developments in Paleoenvironmental Research. Dordrecht, The Nether-

lands: Kluwer, pp. 247–260.

Lotter, A.F., 1998. The recent eutrophication of Baldeggersee (Switzerland)

as assessed by fossil diatom assemblages. Holocene, 8, 395–405.

Maher, B.A., and Thompson, R., 1999. Quaternary Climates, Environ-

ments and Magnetism. Cambridge: Cambridge University Press, p. 328.

Smol, J.P., 2002. Pollution of Lakes and Rivers. A Paleoenvironmental

Perspective. London: Arnold, p. 280.

Stoermer, E.F., and Smol, J.P., 1999. The Diatoms: Applications for the

Environmental and Earth Sciences. Cambridge: Cambridge University

Press, p. 469.

Street-Perrott, F.A., Huang, Y., Perrott, R.A., Eglinton, G., Barker, P.,

Ben Khelifa, L., Harkness, D.A., and Olago, D. 1997. Impact of

lower atmospheric CO

on tropical mountain ecosystems: Carbon-

isotope evidence. Science, 278, 1422–1426.

Talbot, M.R., 2001. Nitrogen isotopes in palaeolimnology. In Last, W.M.,

and Smol, J.P. (eds.), Tracking Environmental Change Using Lake

Sediments. Physical and Geochemical methods. Dordrecht, The

Netherlands: Kluwer, Developments in Paleoenvironmental Research.

pp. 401–439.

742 PALEOLIMNOLOGY

Turney, C.S.M., and Lowe, J.J., 2001. Tephrochronology. In Last, W.M.,

and Smol, J.P. (eds), Tracking Environmental Change Using Lake

Sediments. Basin Analysis, Coring, and Chronological Techniques.

Developments in Paleoenvironmental Research. Dordrecht, The

Netherlands: Kluwer, pp. 451–471.

Cross-references

Beetles as Quaternary and Late Tertiary Climate Indicators

Carbon Isotopes, Stable

Dating, Magnetostratigraphy

Dating, Radiometric Methods

Diatoms

Glacial Megalakes

Lacustrine Sediments

Lake Level Fluctuations

Nitrogen Isotopes

Ostracodes

Oxygen Isotopes

Palynology

Pollen Analysis

Radiocarbon Dating

Tephrochronology

Uranium-Series Dating

Varved Sediments

PALEO-OCEAN pH

Before discussing past variations in the pH of the oceans, it

is first necessary to understand what controls the present distri-

bution of pH in seawater. Along with alkalinity, total dissolved

inorganic carbon (S

DIC

) and the partial pressure of carbon

dioxide (pCO

), pH is one of the four key variables of the sea-

water dissolved carbonate system. If any two of these four vari-

ables are specified, then the values of the other two are fixed.

For those areas of the modern oceans that are in equilibrium

with the CO

content of the atmosphere, their pH is set at a

value of ~8.2 (depending on the temperature of the surface

ocean). An increase in atmospheric CO

(for a given value of

alkalinity or S

DIC

) will result in a decrease in the pH of surface

waters. For example, increasing atmospheric pCO

from the

pre-industrial value of 280 ppmv (parts per million by volume)

to its present day value of ~370 ppmv would result in a

decrease in pH from 8.18 to 8.08. Similarly, the rise in atmo-

spheric CO

values from ~180 ppmv at the height of the last

glacial interval to 280 ppmv would have resulted in a decrease

in pH from 8.32 to 8.18.

Higher pH values are also seen in areas of the modern

oceans where there is significant uptake of dissolved CO

biological activity and downwelling of the surface waters

before equilibrium with the atmosphere can be re-established.

In the modern oceans, this process is mainly operating in

the North Atlantic Ocean and in the southern oceans around

Antarctica, and results in these areas acting as sinks for atmo-

spheric CO

. Deeper waters have higher CO

concentrations

and lower pH than surface waters (see below). Hence, areas

of the oceans where there are strong upwelling currents (mainly

the eastern equatorial Pacific and the Arabian Sea) have much

lower pH values than would otherwise be expected. These

areas of the oceans act as the Earth’s largest natural source

of CO

to the atmosphere.

The pH of the oceans also varies with depth. In the surface

oceans, phytoplankton incorporate dissolved CO

to form

organic carbon and, in the case of coccoliths, they also form

calcium carbonate shells. The former process acts to increase

the pH of the surface water, while the latter acts to decrease

it. In either case, the pH of the surface waters generally remains

buffered at a constant value by exchange of CO

with the atmo-

sphere. Once the phytoplankton and other organisms die, their

remains sink through the water column. The organic matter

is rapidly oxidized to CO

by microbiological activity, such

that the pH of seawater decreases below the mixed layer, to

reach a minimum value of between ~7.6 and 7.8 at depths

of ~200–600 m. The precise depth and magnitude of the pH

minimum depend on the rate of supply of organic matter from

the surface (essentially a function of primary productivity) and

the physical oceanography of the water column. At even

greater depths, first aragonite shells and then calcite shells start

to undergo dissolution, resulting in an increase in pH, such that

the deepest waters of the ocean have similar pH to surface

waters. The net result of all these processes is a water column

pH profile similar to that shown in Figure P25.

It is apparent, therefore, that if we were able to generate a

record of past variations in the pH of the oceans, we could

glean much information about processes such as the history

of atmospheric CO

, the spatial distribution of oceanic CO

sinks and sources, and the spatial and temporal history of ocea-

nic primary production. Recent developments in boron isotope

geochemistry have provided us with the first opportunity

to address these problems.

In seawater, boron is a conservative element with an oceanic

residence time of 20 million years. Although the boron/sali-

nity ratio is constant throughout the oceans, dissolved boron

can exist as B(OH)

or B(OH)

–

, with the former species being

dominant at lower pH and vice versa. The boron isotope com-

position of seawater is also constant throughout the oceans at a

B value of +40%. However, the boron isotope systematics

are such that the light isotope,

B, is preferentially partitioned

into B(OH)

–

. The magnitude of this isotope fractionation is

temperature dependent, but is of the order of 20 % at 20



The relative proportions of B(OH)

and B(OH)

–

are pH depen-

dent, but the boron isotope fractionation between them is con-

stant, so the absolute d

B value of both species is also pH

dependent (Figure P26).

Empirical studies and culture experiments have shown that

when foraminifera form their calcite shells they incorporate

Figure P25 Schematic pH-depth profile in the ocean.

PALEO-OCEAN pH 743

10 ppm boron. This boron is exclusively derived from the

dissolved B(OH)

–

in seawater and it is incorporated without

significant fractionation of the d

B value of this species

(Sanyal et al., 1996). Hence, measurements of the d

B of for-

aminiferal calcite can be used to reconstruct the pH of the

waters from which they precipitated their shells. As with mea-

surements of any other paleo-oceanographic proxy, care must

be taken to ensure that the shells have not been altered (e.g.,

recrystallized or partially dissolved) since they were first

formed, and the foraminiferal calcite must be cleansed of any

potential contaminant phases (such as clays and coccoliths)

before it is analyzed. In addition, the analytical methods

required for determining the d

B value of the small quantities

of boron present (1 ng) in the typical sample sizes available

(20–50 shells) are not simple and much effort is required to

obtain reliable data. Nevertheless, important results have

emerged from this technique.

The utility of this approach in reconstructing the past pH

of the oceans was provided by a study of the boron isotope

composition of well-preserved foraminifera from sedim-

ents recovered from the Ontong-Java Plateau in the western

equatorial Pacific by the Ocean Drilling Project (ODP) (Palmer

et al., 1998). Samples of sediment deposited at five different

times from 85 thousand years ago (ka) to 15.7 million years

ago (Ma) were hand picked to separate individual species

of foraminifera. This is important as different species of for-

aminifera precipitate their shells at different depths within

the water column. For example, Globigerinoides sacculifer

precipitates its shell within the mixed layer of the surface ocean

that is in approximate equilibrium with atmospheric CO

whereas Globorotalia crassaformis calcifies its test at depths

of 200–400 m. Foraminifera from deeper waters would be

expected to record d

B values indicating that they precipitated

their shells in lower pH water than those that lived near the

surface. Indeed, Figure P27 demonstrates that in the five time

slices considered the foraminifera recorded pH values that

yielded similar pH-depth profiles to that observed in the

modern ocean at this site.

Figure P26 Boron isotope composition of dissolved boron species

in seawater.

Figure P27 pH depth profiles for various time slices in the Neogene (ellipses define pH values calculated from the d

B of foraminifera; solid lines

are present day pH-depth curve).

744 PALEO-OCEAN pH

This technique of reconstructing pH-depth profiles was

applied to foraminifera from the middle Eocene (43 Ma)

(Pearson and Palmer, 1999). At that time, the global climate

was considerably warmer than today, such that polar ice caps

were either absent entirely, or very much smaller than those

of today. The warmer climate of the Eocene has been variously

ascribed to higher atmospheric CO

levels or different ocean

circulation patterns that resulted from the position of the conti-

nents at that time. Although boron has a very long oceanic resi-

dence time, which means that its isotopic composition will

not change over periods of thousands of years, it is likely that

there will have been changes in the d

B value of seawater over

millions of years. Hence, any changes that are observed in the

B of ancient foraminiferal calcite could arise from changes

in the boron isotope ratio of seawater as well as changes in

the pH of the seawater. Fortunately, this problem can be

addressed by analyzing the d

B of foraminifera that precipi-

tated their shells at different depths. As noted above, the

decrease in pH from surface to intermediate depth waters is a

result of the oxidation of organic carbon. This process utilizes

dissolved oxygen in the water column and thus provides two

constraints on the minimum pH that can be calculated for dee-

per waters. Firstly, the pH cannot have been so low as to inhibit

the calcification of the deepest dwelling foraminifera. Second,

the change in pH from surface to deep waters cannot have been

so great as to completely deplete the seawater of dissolved oxy-

gen, as foraminifera cannot live in anoxic waters. Using these

constraints Pearson and Palmer (1999) were able to calculate

that the d

B of mid-Eocene waters lay between +41 and +38

(best estimate of +39.4) and that the pH of surface waters

was between 7.91 and 8.33, with a best estimate of 8.05. This

in turn implied that the CO

of the atmosphere at that time

was similar to that of today or only slightly higher.

By using the boron isotope composition of surface dwelling

foraminifera from the past 60 million years, Pearson and

Palmer (2000) were able to provide a record of surface water

pH over this time period (Figure P28a). This study showed

that surface water pH values were roughly similar to present

day values over the past 20 Ma, but that they were considerably

lower during the late Paleocene thermal maximum (LPTM) and

for much of the Eocene. By making certain assumptions con-

cerning the past alkalinity of the oceans, it was also possible

to reconstruct a record of atmospheric CO

levels since this

time (Figure P28b). As might be expected, the uncertainties

in the calculated pH and CO

values become greater further

back in time. This uncertainty arises from some of the assump-

tions that have to be made concerning the past d

B-pH and

alkalinity of the oceans, but the increased uncertainties in pH

and CO

also arises from the fact that the d

B relationship illu-

strated in Figure P26 is not linear. Hence, at low pH values

the same analytical error in measured d

B as was observed at

higher pH results in a larger error in the calculated pH and

Figure P28 (a) Sea surface pH for the past 60 Ma; (b) Atmospheric CO

for the past 60 Ma.

Figure P29 (a) pH of surface waters from the western equatorial

Pacific Ocean; (b) pCO

of surface waters from the

western equatorial Pacific Ocean – solid line gives contemporaneous

atmospheric CO

levels.

PALEO-OCEAN pH 745

values. Despite these uncertainties, this study lends sup-

port to the theory that the very warm climate of the LPTM

was due to high atmospheric CO

levels that also resulted in

low surface water pH values.

More recently, Palmer and Pearson (2003) measured the

B values of G. sacculifer to determine the change in pH

and CO

of surface waters from the western equatorial Pacific

(Figure P29a). By comparison with the record of atmospheric

preserved in ice cores, they were able to show that surface

waters from this area were in equilibrium with the atmosphere

during most of the Holocene and at the height of the last glacia-

tion, but that they had much higher pCO

values during the tran-

sition from glacial to interglacial periods (Figure P29b). This

study suggests that increased upwelling in the eastern equatorial

Pacific may have played an important role in the increase in

atmospheric CO

levels from glacial to interglacial times.

Although still very much in its infancy, it is apparent that

the reconstruction of the paleo-pH of the oceans is a powerful

tool with which to investigate some of the fundamental

processes that have controlled the Earth’s climate. With

improvements in analytical techniques, it can be anticipated

that we will be able to reconstruct increasingly detailed records

of the spatial and temporal variation of seawater pH in the

ancient oceans.

Martin R. Palmer

Bibliography

Palmer, M.R., and Pearson, P.N., 2003. A 23,000 year record of surface

water pH and pCO2 in the western equatorial Pacific. Science, 300

(5618), 480–482.

Palmer, M.R., Pearson, P.N., and Cobb, S.J., 1998. Reconstructing past

ocean pH-depth profiles. Science, 282, 1468–1471.

Pearson, P.N., and Palmer, M.R., 1999. Middle Eocene seawater pH and

atmospheric carbon dioxide concentrations. Science, 284, 1824–1826.

Pearson, P.N., and Palmer, M.R., 2000. Atmospheric carbon dioxide con-

centrations over the past 60 million years. Nature, 406, 695–699.

Sanyal, A., Hemming, N.G., Broecker, W.S., Lea, D.W., Spero, H.J., and

Hanson, G.N., 1996. Oceanic pH control on the boron isotope compo-

sition of foraminifera: Evidence from culture experiments. Palaeocea-

nography, 11, 513–517.

Cross-references

Atmospheric Evolution, Earth

Carbon Dioxide and Methane, Quaternary Variations

Carbon Dioxide, Dissolved (Ocean)

Carbon Isotope Variations Over Geologic Time

Foraminifera

Marine Carbon Geochemistry

Paleocene-Eocene Thermal Maximum

Paleogene Climates

Quaternary Climate Transitions and Cycles

Stable Isotope Analysis

PALEO-PRECIPITATION INDICATORS

Introduction

Precipitation is one of the most important elements of the

climate and the variability of precipitation affects much of life

on Earth. Instrumental records of precipitation (and related mea-

surements such as riverine discharge, lake-level change, out-

going longwave radiation, and others) exist for many parts of

the Earth, but rarely exceed one century in duration. In order

to extend knowledge of precipitation variability further back

in time, the use of paleo-precipitation indicators, also known

as “proxies,” is required. There are no perfect proxies for past

precipitation amount and Mann (2002) has argued persuasively

that accurate paleoclimatic reconstructions are best achieved

through application of multi-proxy approaches.

Natural archives of paleo-precipitation

The most important paleoclimatic archives are historical docu-

mentary records, lacustrine and marine sediments, mountain and

continental glacial ice, tree rings, reef corals, speleothems (cave

deposits), and loess (e.g., Bradley, 1999). Each archive has its

own characteristic rate of formation and fidelity of preservation;

both factors constrain the ultimate temporal resolution of sam-

pling. For example, snow may accumulate on a mountaintop at

a rate of centimeters to meters each year (a potential record of

precipitation amount on a sub-annual timescale), but some of this

snow may be subsequently lost by sublimation, melting, or wind

erosion. The snow that remains is eventually buried, compacted,

and recrystallized to ice; a portion of it may flow laterally under

the force of gravity. All of these processes, which collectively

may be called “diagenesis” (changes happening to materials after

their initial formation), introduce complications for the interpre-

tation of proxy measurements from glacial ice cores (for a thor-

ough, but non-technical account of all aspects of ice cores see

Alley, 2002). Similar diagenetic alteration may also impact proxy

measurements made on sediments, corals, and speleothems.

Lacustrine indicators of paleo-precipitation

Lakes, especially lakes with no outlets, may effectively behave

as very large rain gauges. Precipitation (plus runoff into the lake,

minus any outflow from the lake) is recorded as changes in

lake level, changes that are also dependent upon temperature-

controlled evaporation, although several other climatic elements

(e.g., humidity and wind speed) also influence evaporation rates.

There are many published studies that attempt to quantitatively

reconstruct precipitation history from a history of lake level (e.g.,

Blodgett et al., 1997; Cross et al., 2001). Lacustrine sediments

contain many components that are useful in reconstructing lake

level or related variables such as salinity. Pollen grains and spores

commonly accumulate in lakes and wetlands. The analysis of pol-

len in sediments, the field of palynology, enables a regional recon-

struction of the vegetation cover in the watershed of the lake (or

beyond). Plant growth is dependent on precipitation as well as

other variables such as temperature, light, and nutrients (atmo-

spheric carbon dioxide). Because each may vary independently,

multivariate statistical analysis is an important methodology that

can sometimes permit the deduction of past precipitation. As a

simple example, the ratio of non-arboreal pollen to arboreal pollen

is sometimes utilized as a qualitative indicator of aridity (e.g.,

Maley, 1996). Diatoms are unicellular algae that inhabit rivers,

lakes, and the ocean. Some diatom taxa are planktonic and some

are benthic (shallow water). The ratioof benthic-to-planktonic dia-

toms in a sediment sample is often used as a qualitative measure of

lake level (e.g., Gasse et al., 1997). Some diatom taxaare restricted

to specific ranges of salinity, hence can also be useful indicators of

lacustrine water balance (precipitation minus evaporation). The

mineralogy, trace element composition, and especially the stable

oxygen isotopic composition of calcium carbonate components

of lacustrine sediments can all be useful measures of water bal-

ance. Most often, the carbonate components used in such studies

are the shells of organisms such as ostracodes, gastropods, or

746 PALEO-PRECIPITATION INDICATORS

bivalves (e.g., Schwalb, 2003). Less ideally, the isotopic analysis

can be performed on the bulk carbonate fraction that may consist

of shell fragments, carbonate precipitated by macrophytes or

microorganisms, or apparently abiogenic carbonate precipitates.

When calcium carbonate is absent, stable oxygen isotopic analysis

of siliceous microfossils, such as diatoms, is also possible (e.g.,

Hu and Shemesh, 2003). Stable oxygen or hydrogen isotopic

studies are increasingly being performed on specific organic com-

pounds found in lacustrine sediments, such as cellulose, n-alkanes,

or palmitic acid (Huang et al., 2004). The stable isotopic comp-

osition of all of these compounds is dependent not only upon the

precipitation amount, but also upon the isotopic composition of

the lake water (and in some cases the temperature of formation

of the compound). This, in turn, is determined by the isotopic

composition of the source precipitation and inflow as well as the

amount of evaporative concentration of lake water. In many cases,

variation of the isotopic composition of the source water (preci-

pitation and inflow) is the dominant source of variation. The lack

of knowledge of the temporal variability of the isotopic composi-

tion of precipitation and inflow is a common limit to quantitative

application of this method, although in exceptional cases the iso-

topic composition of source precipitation may be known (e.g.,

Seltzer et al., 2000).

Marine sedimentary indicators of paleo-precipitation

Most open-marine (pelagic) sediments contain few indicators

of paleo-precipitation. However, in marginal marine environ-

ments, sediments derived from adjacent land masses can be

useful indicators of terrestrial runoff (precipitation minus eva-

poration). For example, color reflectance (Peterson et al.,

2000) and Ti/Al ratios (Haug et al., 2001) of marine sediments

from the Cariaco Basin just offshore of Venezuela were used to

deduce the runoff from the adjacent continent. These studies

demonstrated that northern South America was relatively dry

during the Little Ice Age, the Younger Dryas, and the Last Gla-

cial Maximum. DeMenocal et al. ( 2000 ) concluded that a rapid

increase of wind-blown lacustrine diatoms and other terrestrial

sediments in marine sediments off the coast of northwest Africa

indicated an abrupt drying in the Sahara Desert at about 4,000

BP. Pollen grains can also be transported by wind or runoff

and deposited in marine sediments where they can provide

important information about past precipitation amounts on the

adjacent continent. In the marine sediments of the Amazon

fan, analyses of pollen (Haberle and Maslin, 1999), iron-oxide

minerals (Harris and Mix, 1999), oxygen isotopes in marine

calcareous organisms (Maslin and Burns, 2000), and organic

biomarkers of rain forest vegetation (Kastner and Goni, 2003)

have all been related to past precipitation amount in the adja-

cent Amazon basin.

Ice cores as recorders of snow accumulation

Ice cores record many aspects of past climates. It is usually possi-

ble to detect annual layers in the upper portions of ice cores. These

annual layers are due to summer melting and re-freezing (in

high latitude ice cores) or to dry-season dust accumulation

(in low-latitude ice cores). The annual precipitation amount can

be calculated (as water equivalent) from the thickness of the

annual layer and the snow/ice density. Unfortunately, wind ero-

sion, melting, sublimation, and lateral flow are often significant

and such direct deduction of past precipitation amount is limited

to the upper portion of ice cores. In tropical ice cores, it is possible

to reconstruct paleo-precipitation from the oxygen (or hydrogen)

stable isotopic composition of the ice (Hoffmann et al., 2003).This

method of reconstruction (although not universally accepted)

relies upon the observation that there is a significant correlation

between precipitationamount and d

O(ordD) of the precipitation

(e.g., Rozanski et al., 1993) in tropical latitudes and that this rela-

tionship has held through time.

Glacial mass balance (accumulation minus ablation), like the

water balance of lakes, is dependent upon precipitation amount

(as well as temperature). Reconstructions of snowline elevation

in arid mountain regions can provide useful estimates of pastlevels

of precipitation (Kull and Grosjean, 2000).

Speleothems as recorders of paleo-precipitation

Speleothems are calcium carbonate deposits, including stalag-

mites and stalactites, that are deposited from carbon dioxide

dissolved in groundwater that seeps into caves. These deposits

are increasingly being used for high-resolution reconstruction

of precipitation. Rates of growth of the cave deposits may be

related to precipitation amount, with more rapid deposition of

calcium carbonate occurring in wetter periods. In rare cases,

cave deposits may form annual layers, whose thickness can

be related to annual precipitation amount. Oxygen isotopic

composition of the speleothem calcite is determined by its tem-

perature of formation and the isotopic composition of ground-

water. If mean annual temperature changes are relatively

small through time, then changes in d

O of the calcite are lar-

gely controlled by changes in the isotopic composition of

groundwater, which is most often nearly the same as the isoto-

pic composition of regional precipitation. As mentioned pre-

viously, in the tropics, d

O of rainwater is correlated with

precipitation amount, thus, within accepted bounds, the isoto-

pic composition of tropical cave deposits is a reliable indicator

of past precipitation (e.g., Fleitmann et al., 2003). Because cave

deposits can be precisely dated by the U/Th method, this

method of precipitation reconstruction is of great utility for

the past few hundred thousand years. Furthermore, if the

hydrogen isotopic composition of fluid inclusions in the cave

deposit is measured along with the oxygen isotopic composi-

tion of the host calcite, paleotemperatures can be directly calcu-

lated (e.g., Schwarcz and Yonge, 1983).

Other recorders of paleo-precipitation

Many other proxies have been used as recorders of paleo-

precipitation. In semi-arid environments, tree growth is particu-

larly dependent upon water availability and annual extension

(tree-ring width) may be a reliable indicator of past precipitation

amount (e.g., Villalba et al., 1998). The magnetic susceptibility of

loess sequences in China has been utilized as a paleo-precipitation

indicator (Maher and Thompson, 1995). The oxygen isotopic

compositionof the calcium carbonateprecipitated by reef-building

corals is controlled by both temperature and isotopic composi-

tion of surface seawater, itself partly dependent on precipitation

amount, thus if temperature can be independently determined

(for example, using the strontium/calcium ratio of the coral arago-

nite), then it may be possible to reconstruct paleo-precipitation

(McCulloch et al., 1994).

Paul A. Baker

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Cross-references

Dendroclimatology

Diatoms

Ice Cores, Mountain Glaciers

Lacustrine Sediments

Lake Level Fluctuations

Loess Deposits

Ostracodes

Paleoclimate Proxies, an Introduction

Paleosols, Pre-Quaternary

Pollen Analysis

Stable Isotope Analysis

Speleothems

Uranium-Series Dating

PALEOSOLS, PRE-QUATERNARY

Introduction and definitions

Paleopedology is the study of the genesis, properties, climate, and

landscape records of fossil soils, or paleosols. Pre-Quaternary

paleosols are buried, “fossilized” soil horizons, which are older

than 2 Ma; they are also commonly lithified, i.e., converted from

soil into rock by the geological processes of burial diagenesis

(Retallack, 1991). Unlike the more rigid definition used by soil

scientists, who consider a soil to be a body of geologic material

that supports plant life, geoscientists interpret pre-Quaternary

paleosols as any former subaerially exposed surfaces or layers

of earth material that have been affected by physical, chemical,

and biological weathering processes. However, evidence for

the presence of plants is not considered a prerequisite for

identification of a fossil soil (Retallack, 2001). This broader

geological definition allows for greater flexibility in examining

parts of Earth history during which terrestrial plants were

not present.

Paleosol features

Impetus for the study of pre-Quaternary paleosols largely origi-

nated with Dr. Gregory Retallack at the University of Oregon,

who has published a widely used textbook on paleopedology

(Retallack, 2001). In a 1985 Geological Society of America Pen-

rose Conference on paleosols (organized by Dr. Retallack), geos-

cientists, for the first time, began to think about soils of the

past set in the backdrop of “deep” geologic time. Recognition

and interpretation of pre-Quaternary paleosols requires the use

of multidisciplinary approaches that involve integration of field

morphological, microscopic, and geochemical data, as well as an

appreciation for evolutionary changes in terrestrial biota and flora

that have occurred over time (Mora and Driese, 1999). Field mor-

phological features such as plant root traces (in post-Silurian

paleosols), soil horizons, and soil structures (especially peds or

natural soil aggregates), are very diagnostic for pre-Quaternary

paleosol identification (Figure P30a;Retallack,1988). Micromor-

phological (thin section) features are commonly well preserved in

748 PALEOSOLS, PRE-QUATERNARY

paleosols and include root and soil animal traces, peds, and con-

centrations (or evidence for removal) of soil constituents such as

clays, carbonates, Fe-Mn oxides, etc. (Figure P30b;Drieseand

Foreman, 1992; Mora and Driese, 1999). Geochemical patterns

related to weathering and element translocation, including mole-

cular ratios of oxides or elements (e.g., Figure P30c; Retallack,

2001), concentration ratios (molecular ratios normalized to an

assumed immobile element such as Ti; Driese and Foreman,

1992), and mass balance (residual enrichment, volume change

during weathering, and mass transport, normalized to an immobile

element such as Ti or Zr; Driese et al., 2000), provide supporting

evidence for interpretations of paleosols. Stable isotopes, espe-

cially d

C values of paleosol organic matter and pedogenic carbo-

nate, are useful in identifying terrestrial sources of organic matter

and ecosystem types (Cerling, 1991). Most plant communities

during the Paleozoic and Mesozoic eras were C3 (Calvin cycle),

with d

C values of soil organic matter averaging 26% PDB

and pedogenic carbonate values ranging from5to12%

PDB (Mora et al., 1996; Mora and Driese, 1999). The d

Ovalues

of pedogenic carbonate, if not diagenetically altered, can be used

as proxies for meteoric water compositions and paleotemperatures

(Cerling, 1984), but most values measured for Paleozoic pedo-

genic carbonates have been apparently reset during burial dia-

genesis (Mora and Driese, 1999).

Problems in recognition and interpretation of paleoclimates

from pre-Quaternary paleosols include extensive burial diage-

netic alteration (Retallack, 1991). Common diagenetic alteration

processes include physical compaction and consequent modifica-

tions of original soil thickness and morphology, oxidation of

soil organic matter, burial gleization (Fe loss, largely through

microbial reduction), color modifications (intensification) related

to dehydration and recrystallization of hydrous mineral phases

(such as FeOOH), recrystallization of soil smectites to illites

(or even to metamorphic mineral assemblages; Table P3), and

exchange of oxygen isotopes between burial fluids and pedo-

genic carbonates and clays (Mora and Driese, 1999). Mack et al.

(1993) proposed an alternative paleosol classification, because of

burial diagenetic alteration and associated lack of preservation of

Figure P30 Pennington Formation (325 Ma, late Mississippian), Tennessee, USA paleosol. (a) Field photograph: A (surface), Bss (slickensided

B) and Bkss (slickensided B with pedogenic carbonate) refer to paleosol horizons. Note angular blocky ped structure in outcrop (scale card is

15 cm long). (b) Thin-section photomicrograph of paleosol shown in (a), which exhibits oriented bright clay microfabric related to seasonal

cycles of wetting and drying and associated clay shrinking and swelling, as well as dark-colored Fe-Mn oxide nodule (cross-polarized light).

), clay accumulation (Al

/ SiO

), Fe

accumulation (Fe

/Al

), and salinization (Na

O/K

O). Note carbonate and Fe accumulation in deeper portion of paleosol, but relatively

uniform clay content and salinization.

Table P3 Diagenetic alteration of paleo-Vertisols (high clay-content paleosols with extensive shrink-swell features such as slickensides) compared

with a modern soil analog (Houston Black series), as a function of increasing burial depth, temperature, and time (Driese et al., 2000; Rye and

Holland, 2000)

Name, Geologic age Burial depth (km) Burial temperature (



C) Clay mineral assemblage wt% K

Houston Black (modern) Surface 22 (mean annual temperature). Na-smectite 1.29

Pennington (325 Ma) 2–360–90 Illite ( + kaolinite + chlorite) 4.61

Hekpoort (2.25 Ga) 12–15 350 Sericite – muscovite – chlorite 9.50

PALEOSOLS, PRE-QUATERNARY 749

features necessary for rigorous application of USDA Soil Taxon-

omy (Figure P31). Evolutionary changes in terrestrial plant and

animal communities over time also create difficulties in paleosol

interpretations because of attendant changes in morphological

features such as root traces, which are: (a) not present in Ordovi-

cian paleosols; (b) rhizomatous and fine in Silurian paleosols;

and (c) larger, deeper, and more “modern” in Devonian and

post-Devonian paleosols (Driese and Mora, 2001).

Paleosols as paleoclimate proxies research

Current research on pre-Quaternary paleosols no longer con-

cerns simple identification of paleosols in the geologic record,

but instead focuses on the use of paleosols as proxies for

reconstructions of paleoclimate, paleolandscape evolution, and

paleoatmospheric chemistry. Paleoclimate reconstructions pri-

marily emphasize the determination of estimates of mean annual

precipitation (MAP) by using a variety of proxy measures, includ-

ing: (a) the depth to pedogenic carbonate (Bkss) horizon used by

Caudill et al. (1996) to estimate mean annual paleoprecipitation

(MAP) of 648 141 mm yr

1

(after Retallack, 2001;

Figure P30a), (b) Fe content of pedogenic Fe-Mn nodules and con-

cretions (Figures P30b and P32a), which was used by Stiles et al.

(2001) to estimate MAP of 989 mm based on wt% Fe content;

(Figures P30c and P32b; e.g., Chemical Index of Alteration minus

Potash or CIA-K, Sheldon et al., 2002); and (d) total element

mass-flux and mass-balance calculations (Stiles et al., 2003).

Paleolandscape reconstructions in which topography or hydro-

logy are important variables (interpretation of paleocatenas)

are commonly conducted at both local and more regional scales.

Reconstructions of Phanerozoic pCO

employ the CO

-carbonate

paleobarometer of Cerling (1991) to interpret Paleozoic (Mora

et al., 1996; Mora and Driese, 1999) and Mesozoic paleoatmo-

spheres. This technique utilizes the d

C values measured from

pedogenic carbonates, measurements of d

C values of paleosol

organic matter (or an estimate based on marine proxy records),

and assumptions of soil productivity to infer paleoatmospheric

pCO

. These results are in good agreement with pCO

estimates

based on long-term mass-balance carbon models (Figure P33).

Studies of oxygen levels of Precambrian atmospheres focus on

Figure P31 Classification for paleosols (Mack et al., 1993).

Figure P32 Geochemical climate proxies derived from pre-Quaternary paleosols. (a) Paleoprecipitation (MAP) for mid-Mississippian (Maccrady),

late Mississippian (Pennington), and early Permian (Dunkard) Appalachian basin paleo-vertisols based on measured Fe content of paleosol Fe-Mn

bodules and regression equation of Stiles et al. (2001) for total Fe content of Fe-Mn nodules in modern soil analogs in Texas, USA. Estimated MAP

ranges from 850 to 1,310 mm yr

1

. (From Stiles et al., 2001, used with permission). (b) Mean annual precipitation (MAP) for six Appalachian basin

paleo-Vertisols derived from bulk chemistry and regression equation of Sheldon et al. (2002) that relates MAP to Chemical Index of Alteration

minus Potash (CIA-K = molar ratio of Al

/ (Al

+Na

O + CaO)  100. P = 14.265(CIA-K)37.632, where P = MAP in mm yr

1

= 0.73), which

increases by 400–500 mm MAP during latest Mississippian time in southeastern Kentucky, USA.

750 PALEOSOLS, PRE-QUATERNARY