Odekon M. Encyclopedia of paleoclimatology and ancient environments

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Triassic (Carnian Molteno Formation), the paleoclimate was

humid with formation of non-calcareous red paleosols (Alfi-

sols) and swamp paleosols (Histosols). By early Jurassic (Het-

tangian-Toarcian Elliot and Clarens Formation), paleosols were

again calcareous and nodular, like dry climate soils (Aridisols).

Other indications of Triassic and Jurassic paleoclimates

come from sedimentary indicators such as coals, evaporites,

laterites, and bauxites (Parrish, 1998). These indicators of

wet, dry and tropical climates, respectively, confirm paleocli-

matic restorations based on plants and soils. Such linkage is

not surprising considering that most coals form in Histosol

soils, many evaporites in salid Aridisols, and many bauxites

in Oxisols (Retallack, 2001a). One exception to the use of coal

as a paleoclimatic indicator is a lack of coal anywhere in the

world during the entire duration of the early Triassic. This

multimillion-year coal gap may be due to extinction of coal-

forming plants during the Permian-Triassic life crisis (Retal-

lack, 1996), or to some continuing hazard such as methane

degassing, which would have reduced oxygenation of already

near-stagnant wetlands (Sheldon and Retallack, 2002).

Quantitative paleoclimatic proxies

Among a variety of proxies for the ups and downs of green-

house paleoclimate, the best known is the oxygen isotopic

composition of fossil seashells. This approach was first applied

to the study of glacial versus interglacial benthic foraminifera

by Cesare Emiliani, who found low isotopic values (d

O) at

times of warm oceans and high isotopic values at times of cold

oceans. Such isotopic measurements have been extended to

a variety of tropical fossil seashells ranging back in age by

550 million years (Veizer et al., 2000). The variation in isotopic

value reflects increased temperature-dependent evaporation of

the light isotope of oxygen (

O) from the ocean relative to

the heavier common isotope (

O), and its sequestration in

snow and ice of polar ice caps. There is also a salinity effect,

because there is more evaporation of the light isotope than

the heavy one from fresh waters. The net effect is that when

the oxygen isotopic value (d

Oin% of standard notation)

goes down, the ocean is warm, less saline and with small ice

caps. Salinity and ice cap effects are generally disregarded in

using a transfer function for the relationship between tempera-

ture and isotopic composition of modern foraminifera to esti-

mate paleotemperatures of the past (Parrish, 1998). During

the Triassic and Jurassic, however, there was a long-term

increase in isotopic values, which is probably due to greater

facility of the light isotope to emerge at the surface after sub-

duction and crustal recycling in sedimentary pore water (Veizer

et al., 2000). Nevertheless, the oxygen isotopic record is

valuable because it is known in considerable detail, and

because it indicates paleoclimatic volatility (Figure T19). This

record suggests that tropical marine temperatures have varied

considerably, perhaps by as much as 6



C, between relatively

cool periods that were nevertheless warmer than at present,

and peaks of warmth during the earliest Triassic (Griesba-

chian), early-middle Triassic (Spathian), early Late Triassic

(Carnian), earliest Jurassic (Hettangian), early-middle Jurassic

(Toarcian), middle Jurassic (Bathonian-Callovian), and earliest

Cretaceous (Berriasian).

These various peaks of warmth indicated by oxygen isotopic

composition of tropical sea shells are also times of peak atmo-

spheric carbon dioxide revealed by studies of the stomatal

index of Lepidopteris and Ginkgo leaves (Figure T19), thus

supporting a greenhouse mechanism for paleoclimatic change

in the past. The stomatal index is a measure of the number of

plant gas-intake pores (stomates) normalized to the total num-

ber of epidermal cells in well-preserved fossil leaves from

which cuticles can be prepared and examined under the micro-

scope. Plant leaves have fewer stomates when atmospheric car-

bon dioxide is high in order to conserve water transpired

through the same stomatal pores. Quantification of this well-

known plant response to rising atmospheric carbon dioxide

has come from greenhouse studies of living Ginkgo plants, as

well as of Ginkgo leaves conserved in herbaria during the past

century of industrially generated increase in atmospheric car-

bon dioxide (Retallack, 2002). The amounts of atmospheric

carbon dioxide revealed by this method are close to modern

levels (280 ppmV) for Permian time, when glacial ice caps

were as extensive as they are today. This result indicates that

this quantification does not show long-term drift, as is apparent

from the oxygen isotopic record (Figure T19). Earliest Triassic

(Griesbachian), earliest (Hettangian) and middle Jurassic

(Bathonian), and earliest Cretaceous (Berriasian) peaks of car-

bon dioxide abundance are startlingly high ( >2,000 ppmV),

and coincide with mass extinctions of life on land and in the

ocean. There are other peaks of lesser magnitude, and interven-

ing times of atmospheric carbon dioxide about twice modern

levels (ca. 500 ppmV).

Paleosols also can provide proxies for paleoclimatic mean

annual precipitation and mean annual temperature. Mean an-

nual precipitation, for example, can be inferred from the rela-

tionship observed in modern calcareous soils of increased

depth to the horizon of carbonate nodules in increasingly

humid climates. For paleosols, compaction correction needs

to be made to the depth of the calcic horizon in a paleosol in

order to restore it to the original thickness before burial by

overlying sediments (Retallack, 2001a). Mean annual precipita-

tion and temperature also can be approximated by the degree to

which non-calcareous, clayey, subsurface horizons of paleosols

have been leached of alkaline earth (Ca, Mg) and alkali ele-

ments (K, Na). Both methods have been used to estimate paleo-

climatic parameters of paleosols (Sheldon et al., 2002), but not

Figure T19 Ups and downs of the early Mesozoic greenhouse revealed

by biogenic carbonate oxygen isotopic composition as a proxy for

paleotemperature, and by stomatal index of fossil plants as a proxy for

atmospheric carbon dioxide concentration (data from Veizer et al.,

2000; Retallack, 2002). The oxygen isotopic data is plotted in reverse

order so that peaks are warm and valleys are cold, and its long-term

trend is in part an artefact of crustal recycling.

TRIASSIC-JURASSIC CLIMATES 965

widely applied to Triassic and Jurassic paleosols. The Triassic

and Jurassic paleosol record of the Colorado Plateau looks

especially promising in this regard (Retallack, 1997b), be-

cause early middle Triassic paleosols (Anisian, Moenkopi

Formation) are weakly calcareous, as are early late Triassic

paleosols (Carnian Chinle Formation), but later Triassic paleo-

sols (Carnian-Norian Chinle Formation) become increasingly

calcareous with shallow calcic horizons. Earliest Jurassic paleo-

sols (Hettangian Kayenta Formation) are weakly calcareous,

but those few paleosols known from the early-middle Jurassic

(Pliensbachian-Bajocian Navajo Sandstone – a desert dune

sequence) are dolomitic and shallowly calcareous. Weakly cal-

careous paleosols are again found in the later middle Jurassic

(Bathonian-Callovian Curtis and Summerville Formations),

but calcareous paleosols are abundant in the late Jurassic

(Tithonian-Kimmeridgian Morrison Formation). These obser-

vations suggest that global warm spikes were also wet spikes

in the paleosol record. Warm-wet paleoclimatic spikes alternat-

ing with cold-dry paleoclimates can also be inferred from vege-

tation modeling (Beerling and Woodward, 2001) and biome

reconstruction (Figure T18 ). Global thermal maxima of the

Triassic and Jurassic were times when cool temperate vegeta-

tion retreated to higher latitudes and altitudes, temperate and

tropical shrublands covered smaller areas, and summerwet sub-

tropical and rain forests covered larger areas, compared with

global thermal minima.

Yet another quantitative approach has been computer

modeling of paleoclimate using modern general circulation

models adjusted for ancient paleotopography, and for estimates

of soil and vegetation albedo and other ecosystem properties.

These models have predicted extreme seasonal paleoclimates

for the continental interiors and megamonsoons for peritropi-

cal margins of the northern supercontinent Laurussia and

the southern supercontinent Gondwana. Such paleoclimatic

extremes are not compatible with the fossil plants, animals,

and soils found there (Retallack, 1997b; Retallack and Krull,

1999). Some, but not all, of these problems are mitigated by

running the computer models with elevated atmospheric car-

bon dioxide (Wilson et al., 1994), thus giving theoretical con-

firmation of the observed link between paleotemperature and

greenhouse gases.

Causes of Triassic and Jurassic global paleoclimatic

variation

Triassic and Jurassic records are effective tests for theories of

global climatic variation because they allow exploration of

greater paleoclimatic extremes than found during the Cenozoic,

yet with animals and plants of more familiar kinds than found

in the Paleozoic and Precambrian. The distinctive paleogeo-

graphic arrangement of a meridional supercontinent of Pangea,

barely including the poles (Figure T18), is a striking difference

from the scattered continents today, with one polar and several

northern peripolar continents.

The supercontinent Pangea has been blamed for climatic

cooling by virtue of its large size and drawdown of atmospheric

carbon dioxide by soil formation and weathering, which pro-

ceeds by consumption of carbonic acid. In contrast, greenhouse

pulses have been related to times of active sea floor spreading

when mid-ocean ridges were high and displaced ocean water to

form epicontinental seas. This idea was first advanced to

explain long-term paleoclimatic change. The Permian icehouse

was caused by assembly of Pangea and the Mesozoic green-

house by Pangean breakup and dispersal, followed by a

Cenozoic ice-house initiated by Northern Hemisphere aggrega-

tion of Asia and North America (Fischer, 1984). Detailed

paleoclimatic records now available (Figure T19) show paleo-

climatic volatility on shorter timescales than continental drift,

so this mechanism is not the whole answer.

Transient greenhouse warming has also been linked to vol-

canic activity, especially to continental flood basalts which

coincide with observed warm spikes (Figure T19): earliest Trias-

sic (Griesbachian Siberian Traps), earliest Jurassic (Hettangian

Newark and Brito-Arctic volcanics), early-middle Jurassic

(Toarcian Ferrar-Karoo basalts) and middle Jurassic (Bathonian

Antarctic basalts), and earliest Cretaceous (Berriasian Parana-

Etendeka traps). Global cooling, or “volcanic winter,” can be

produced by eruptions of ashy particles and sulfate aerosols

into the atmosphere, as first observed by Benjamin Franklin

during Europe’s “year without a summer,” 1783, following the

eruption of the Icelandic volcano Laki. Large flood basalts were

not explosive eruptions like those of Laki, Pinatubo and Mt.

Helens. Instead, flood basalts were low-viscosity lavas, which

flowed without large explosions over large areas. Both kinds

of eruptions also release enormous amounts of carbon dioxide

and water vapor, which are potent greenhouse gases, their effects

are felt after particles and aerosols are scrubbed from the atmo-

sphere by rain. Unfortunately, the currently estimated gas release

from even these enormous eruptions is too spread out in time

and too small (Wignall, 2001) to have created the likely variation

in atmospheric carbon dioxide (Figure T19).

Impact of large asteroids and comets also create global

paleoclimatic havoc (Rampino, 2002), including short-term

impact-winter from dust and aerosols cast aloft by crater exca-

vation. This dust is more strongly acidic than for volcanic erup-

tions as it includes sulfur oxide gases that are rained out as

sulfuric acid. Sulfuric acid may be very abundant if asteroids

or comets vaporize gypsum or other sulfur-bearing rocks within

the impact crater, as at the Chicxulub crater, Yucatan, Mexico.

Another strong acid formed during impact is nitric acid, gener-

ated by impact shocking, and oxidation and hydration of atmo-

spheric nitrogen. Both sulfuric and nitric acids are suspected

to have created strongly leached boundary beds, and to

have extinguished many shelled marine organisms at the

Cretaceous-Tertiary boundary. These days to years of strong

acid rain were followed by several tens of thousands of years

of paleoclimatic warming, due to elevated greenhouse gases

including water vapor and carbon dioxide. This carbon dioxide

could have come from multiple sources including the death and

decay of many kinds of organisms from the trauma of impact

and impact winter, and the burning of forests browned by

acid rain. Although scenarios for impact perturbation of life

and paleoclimate have been most thoroughly explored for

the Cretaceous-Tertiary boundary, large impacts coincide with

other mass extinctions and global warm spikes during the

Triassic and Jurassic (Rampino, 2002), including the early-late

Triassic (Spathian) Araingahua Crater (40 km), the late Triassic

(Norian) Manicougan (100 km) and Rochechouart (23 km)

Craters, the middle Jurassic (Bajocian) Puchezh-Katunki Crater

(80 km), and the earliest Cretaceous (Berriasian) Morokweng

Crater (>70 km). Like flood basalts, however, the magnitude

of the perturbations does not currently correlate closely with

impact magnitude, as judged from crater size, or other evidence

of impact (Rampino, 2002).

966 TRIASSIC-JURASSIC CLIMATES

Another important cause of global paleoclimatic perturba-

tion is methane hydrate destabilization, which could exacerbate

impact and volcanic warming to truly catastrophic proportions.

Methane created by microbial decay of organic matter deep

within sediments and permafrost accumulates within subsur-

face heat traps maintained by geothermal gradients beneath

cold oceans and permafrost. The methane is hydrated within

ice crystals and is physically stable, but can be released by

volcanic eruption, meteorite impact, submarine landslides, or

general warming of overlying sediments or ice. This form of

methane has a very distinctive isotopic composition, which

has been used as a tracer for massive releases of methane,

now well documented for the earliest Eocene (Ypresian), early

Cretaceous (Aptian), early-middle Jurassic (Pliensbachian)

and earliest Triassic (Griesbachian). The global warming and

unusual abundance of

C at these times is attributed to

thousands of gigatons of methane released to the atmosphere,

where the isotopically light methane was oxidized to iso-

topically light carbon dioxide within 7–24 years, and this in

turn fixed within isotopically light organic carbon and carbo-

nate. Methane itself is a powerful greenhouse gas, 50 times

more effective than carbon dioxide. Atmospheric pollution

by methanogenic hydrocarbons has had catastrophic effects

in the past, because some documented releases were also

times of mass extinction (Hesselbo et al., 2000; Krull and

Retallack, 2000).

If volcanoes, impacts, and methane release are the principal

agents of global warming, the principal agent of global cooling

is life itself, particularly its photosynthetic ability to reduce car-

bon within the bodies of organisms and to bury much carbon

within sediments. Root respiration and soil animals also pro-

mote consumption of atmospheric carbon dioxide, as carbonic

acid, the principal agent of silicate weathering that releases

alkali (Na, K) and alkaline earth (Ca, Mg) elements, which fuel

plant growth (Berner and Kothavala, 2001). Mountain uplift

has also been proposed to cool the planet by means of silicate

weathering (Raymo and Ruddiman, 1992). The mountains

themselves do not promote cooling because silicate weathering

is reduced under alpine meadows and ice fields compared with

lowland forests and grasslands. Nevertheless, glacial loess from

montane glaciers and desert dust from montane rain shadows

are important sources of nutrients for soil formation and silicate

weathering, which could have climatic ramifications. Changing

ocean currents and thermohaline circulation also have been

thought to have a cooling effect (Broecker, 1997), but these

are also mechanisms for increasing biological productivity of

surface waters and creating organic matter, which then sinks

and is buried in oceanic sediments. Most marine productivity

and carbon sequestration is nearshore and fueled by leakage of

nutrients from soils and ecosystems on land. Transient global

warming events caused by geological catastrophes were prob-

ably curtailed largely by biotic recovery and carbon sequestra-

tion in soils, lakes and nearshore marine environments.

Carbon sequestration by Triassic and Jurassic terrestrial eco-

systems was probably different from that in both modern and

Permian ecosystems for purely evolutionary reasons. Unlike

modern ecosystems, there were no grasses or grasslands during

the Triassic and Jurassic (Retallack, 2001b). Dry continental

interiors supported dry shrublands of small-leaved cycadeoids,

conifers and seed ferns in soils lacking the fine cracking and

carbon-rich surface horizons of grassland soils. An array of rain

forest, summerwet tropical forest, winter-wet subtropical to

temperate woodland, warm temperate forest, and cool tempe-

rate woodlands formed similar soils to these plant formations

today (Figure T18). No tundra and taiga vegetation was known

in polar regions, but the montane equivalents alpine fellfield

and krummholz were probably present, because there is evi-

dence from fossil soils that plants had evolved into tundra

and taiga vegetation as old as Carboniferous and Permian,

respectively (Retallack, 2001a). Triassic and Jurassic ecosys-

tems, unlike Permian ecosystems, had enormous herbivores

and hordes of social insects. Early to middle Triassic therapsids

were as large as rhinos, and late Triassic and Jurassic dinosaurs

exceeded the bulk of large elephants. Ants and termites are

known as body fossils only as old as Cretaceous, but fossil

nests suggest comparable insects as old as Triassic (Hasiotis

and Dubiel, 1995). Dinosaurs and social insects would have

consumed large amounts of vegetation, and could have contrib-

uted to general global warmth during the colder and drier

phases of early Mesozoic climatic oscillations (Retallack,

1997b). Climate and life on Earth have long been intimately

interrelated because of their complimentary dependence on

greenhouse gases.

Gregory J. Retallack

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Cross-references

Arid Climates and Indicators

Atmospheric Evolution, Earth

Bolide Impacts and Climate

Carbon Cycle

Carbon Isotope Variations over Geologic Time

Climate Change, Causes

Coal Beds, Origin and Climate

Continental Sediments

Evaporites

Flood Basalts, Climatic Implications

“Greenhouse” (warm) Climates

Laterite

Mass Extinctions, Role of Climate

Methane Hydrates, Carbon Cycling, and Environmental Change

Mountain Uplift and Climate Change

Monsoons, Pre-Quaternary

Oxygen Isotopes

Paleobotany

Paleoclimate Modeling, Pre-Quaternary

Paleotemperatures and Proxy Reconstructions

Paleosols, Pre-Quaternary

Paleo-Precipitation Indicators

Plate Tectonics and Climate change

Pre-Quaternary Milankovitch Cycles and Climate Variability

Sedimentary Indicators of Climate Change

Volcanic Eruptions and Climate Change

Weathering and Climate

968 TRIASSIC-JURASSIC CLIMATES

URANIUM-SERIES DATING

The discovery of natural radioactivity at the beginning of the

twentieth century fundamentally changed our understanding

of the physical and biological history of the Earth. Before then,

the antiquity of the Earth was a very debatable issue and thus

any process that required a long time, such as Darwin’s theory

of evolution, was very much in doubt. A number of clever

approaches, based on estimates of time required for certain

physical processes, were used to constrain the age of the Earth,

such as Lord Kelvin’s estimates based on the assumption of a

molten Earth cooling to its present state. All of these estimates

greatly underestimated the age of the Earth. The discovery of

natural radioactivity (Becquerel, 1896) and the fact that the

rate at which a particular radioactive nuclide decays is constant

opened the way to obtaining absolute dates.

Many light elements, such as carbon and potassium have

isotopes that decay in a single step to a stable isotope of

another element. For example,

C decays to

N and

decays to

Ar. All isotopes of elements heavier than bismuth

are radioactive and have decay schemes that involve multiple

stages of decay to intermediate isotopes. One of these elements,

uranium, has two primary isotopes,

238

U and

235

234

U, is an

intermediate daughter of

238

U. The decay of

238

Uto

207

involves 18 intermediate nuclides, while the decay of

235

207

Pb involves 14 intermediate nuclides. U decay-series

nuclides that are currently utilized for dating and their respec-

tive half-lives are shown in Figure U1.

Principle of radiometric dating

For a simple decay scheme, such as

Rb decaying to

Sr, where

the radioactive isotope decays to a stable isotope, the amount of

daughter isotopes, D, that is present after a time, t, has elapsed is

related to the amount of parent isotope, P, at present as follows:

D ¼D

þPðe

 1Þð1Þ

where D is the total number of atoms of the daughter isotope

at present, D

is the number of daughter isotopes present at

the beginning (since t time has passed, for example the age

of a rock if one is dating a rock), P is the number of atoms

of the parent isotope at present, l is the decay constant and

t is the age. As stated previously, the decay constant for a given

radioactive isotope is constant and is related to the half-life

(the amount of time required for a given number of atoms of

a given radioactive isotope to be reduced by half due to decay)

according to:

1=2

ln 2

0:693

ð2Þ

where t

1/2

is the half-life and l is the decay constant for a

particular radioactive isotope.

The range of events that are datable using a given parent-

daughter isotope system depends on the half-life of the parent

isotope. Typically, we cannot date events that are older than

about seven half-lives of a given system. Conversely, to date

young events we need to utilize isotope pairs with a short half-

life parent. A number of the U-series nuclides have short

half-lives that are useful to date recent events of human interest,

such as past climate and environmental change and also, archeo-

logical and volcanic processes.

For radioactive isotopes that decay to another radioactive

isotope, such as U-series nuclides, the number of an immediate

daughter isotope that remains after time t is related to the num-

ber of atoms of the immediate parent:

D ¼

 l

ðe

l

 e

l

ÞþD

l

ð3Þ

where D is the number of atoms of daughter isotope at present, D

is the number of atoms of the daughter isotope present initially, P

is the number of atoms of the parent isotope present initially and

and l

are the decay constants of the parent isotope and the

daughter isotope, respectively. The last term accounts for the decay

of the daughter nuclide since time t. Expressions can be written

describing the relationship between one nuclide and another

nuclide at any stage in a given decay scheme.

After about seven half-lives of the shorter-lived nuclide

of any two nuclides in a given series, the two nuclides reach

a state called “secular equilibrium.” When secular equilibrium

is reached, the activity (the number of atoms of a nuclide multi-

plied by its decay constant) of both nuclides is equal, i.e.,

¼ 1

In other words, their atomic ratio is constant from then on,

equal to l

, and the age can no longer be determined.

Methods of U-series isotope analysis

The utility of U decay series isotopes for dating was recognized

very early (Bateman, 1910) but the application was limited by

technical difficulties related to separation chemistry and mea-

surements of very small amounts of daughter nuclides. For

example, in most rocks the abundance of

230

Th, the daughter

238

U through

234

U (see Figure U1), is about one atom for

every million atoms of uranium. Moreover, the ionization effi-

ciency of Th in thermal ionization mass spectrometers (TIMS),

the mass spectrometers used until recently for such measure-

ments, is typically about 2  10

–4

. In other words, about two

ions (only ions are detected by TIMS) are produced for every

10,000 atoms loaded into the mass spectrometer.

The theoretical limit to precision of data (the counting statis-

tics) is related to the number of ions detected as follows:

Uncertainty ð1sÞ¼

ﬃﬃﬃﬃ

where N is the number of atoms detected. Thus, the combined

low abundance and low ion yield for many of the nuclides of

interest made the application very difficult.

Early measurements, requiring large samples (thus low tem-

poral resolution) were performed by alpha or gamma counting

(e.g., Cherdyntsev, 1956; Broecker, 1963; Ku, 1965). Advances

in separation chemistry and mass spectrometry using TIMS

(e.g., Chen et al., 1986; Edwards et al., 1987 for measurements

of uranium and thorium isotopes, respectively; and Pickett

et al., 1994 for measurement of protactinium by mass spectro-

metry) have made it possible to obtain high-precision U-series

data for dating and isotope tracer application.

Application of U-series dating

A requirement for the application of U-series nuclides for tracer

and dating purposes is that the process of interest has to frac-

tionate one nuclide from another partially or completely.

The most widely used application of U-series dating has

been for dating corals (Edwards et al., 1987) and speleothems

(e.g., Polyak and Asmerom, 2001) for studying the timing

of past climate change. These materials are ideal for dating

because uranium under oxidizing conditions, which is the case

in surface oceans and caves, is preferentially dissolved in

water, while thorium is not soluble. Carbonate, calcite (CaCO

)

in speleothems and aragonite in corals (also CaCO

, but differ-

ent crystal form), can contain up to 4 parts per million (ppm) U

and negligible amounts of the daughter

230

Th nuclide, making

them readily and precisely datable using the

238

U–

234

U–

230

series (Edwards et al., 1987). Assuming the sample did not

contain

230

Th initially:

230

238



¼ð1  e

l

230

 l

234



234

1;000



ð1  e

ðl

234

l

230

Þt

where

230

238



is the activity ratio, t, is the age of the sample and d

234

Uis

equal to:

234

238

 1



1;000

234

238

U is the atomic ratio. Uranium and thorium are chemi-

cally separated from carbonates and the

238

234

U, and

230

are measured in a mass spectrometer. The age, t, is calculated

iteratively. Any event or process that is associated with the for-

mation of carbonate can be dated using this method.

Similar equations are written for dating volcanic rocks (Allè-

gre, 1968), except that in the case of volcanic rocks, one cannot

assume zero initial

230

Th. For volcanic rocks, the most versatile

application of U-series nuclides has been in tracing sources of

magma, constraining the time-scale of magma formation and

evolution and providing relative timing of these events.

One of the novel applications for chronology using U-series

nuclides has been combining

238

U-

234

U-

230

Th and

235

U-

231

chronology on the same sample. If the ages are concordant

(i.e., the two systems give the same age) it gives added validity

Figure U1 Schematic diagram of the U-series decay chain, showing some of the nuclides utilized in chronology and their approximate half lives.

byr, Billions of years; kyr, thousands of years.

970 URANIUM-SERIES DATING

to the age (e.g., Edwards et al., 1997). In another approach, the

presence of an excess or deficiency in daughter activity,

i.e., daughter/parent ratio 6¼ 1, puts a maximum limit on the

age of the material. For example, the presence of excess

210

Pb, which has a half-life of 22.3 years, indicates that an

object is no older than around 150 years.

210

Pb is widely used

for dating recent sediments in environmental research.

The utility of U-series isotopes for chronology and tracer

work has been recognized for a long time, beginning after the

discovery of radioactivity. However, implementation of the sys-

tem was hindered by technical difficulties. Recent advances in

mass spectrometry enable realization of the full potential of the

system. One of the most exciting areas of application of this

method that benefits from these technical gains is the study

of past climate change, in particular dating annually banded

speleothems (Asmerom and Polyak, 2004). These specimens

often contain a variety of proxies of past climate change,

including isotopic, chemical and physical proxies.

Recent advances in mass spectrometry, whereby a plasma

source is combined with a magnetic analyzer and multiple

detectors that consist of either a combination of Faraday cups and

multiple electron multipliers, or Faraday cups and multiple chan-

neltrons, stand to revolutionize the field. These high-resolution

multi-collector inductively-coupled plasma mass spectrometers

(MC-ICPMS) are capable of greater ion yields, up to an order of

magnitude and greater , as compared with TIMS (Robinson et al.,

2002), allowing for high spatial and temporal resolution.

Yemane Asmerom

Bibliography

Allègre, C.J., 1968.

230

Th dating of volcanic rocks: a comment: Earth

Planet. Sci. Lett., 5, 209–210.

Asmerom, Y., and Polyak, J.V., 2004. Comment on Betancourt et al.

(2002): A test of annual resolution in stalagmites using tree rings.

Quaternary Res., 61,119–121.

Bateman, H., 1910. The solution of a system of differential equations

occurring in the theory of radioactive transformations. Proc. Camb.

Philos. Soc., 15, 423.

Becquerel, H., 1896. Sur les radiations invisibles emises par phosphores-

cence. Comptes Rendus, 122, 420.

Broecker, W.S., 1963. A preliminary evaluation of the uranium series

in equilibrium as a tool for absolute age determination on marine

carbonates. J. Geophys. Res., 68, 2817–2834.

Chen, J.H., Edwards, R.L., and Wasserburg, G.L., 1986.

238

234

and

232

Th in seawater, Earth Planet. Sci. Lett., 80, 241–251.

Cherdyntsev, V.V., 1956. Determination of the absolute age of the Paleo-

lithic. Sovietski Arkheologi, 25,64–86.

Edwards, R.L., Chen, J.H., and Wasserburg, G.L., 1987.

238

U-

234

U-

230

Th-

232

Th systematics and the precise measurement of

time over the past 500,000 y. Earth Planet. Sci. Lett., 81, 175–192.

Edwards, R.L., Cheng, H., Murrell, M.T., and Goldstein, S.J., 1997.

Protactinium-231 dating of carbonates by thermal ionization mass

spectrometry: Implications for Quaternary climate change, Science,

276, 782–786.

Ku, T.-L., 1965. An evaluation of the

234

238

U method as a tool for

dating pelagic sediments. J. Geophys. Res., 70, 3457–3474.

Pickett, D.A., Murrell, M.T., and Williams, R.W., 1994. Determination

of femtogram quantities of protactinium in geologic samples by thermal

ionization mass spectrometry. Anal. Chem., 66, 1044–1049.

Polyak, V.J., and Asmerom, Y., 2001. Late Holocene climate and cultural

changes in the southwestern United States, Science, 294, 148–151.

Robinson, L.F., Henderson, G.M., and Slowey, N.C., 2002. U-Th dating

of marine isotope stage 7 in Bahamas slope sediments. Earth Planet.

Sci. Lett., 196, 175–187.

Cross-references

Coral and Coral Reefs

Dating, Radiometric Methods

Isotope Fractionation

Speleothems

URANIUM-SERIES DATING 971

VARVED SEDIMENTS

Introduction

Varved sediments are sequences of sedimentary laminations

deposited within a single year. A varve (Swedish: varv, layer)

is a pair or set of laminae formed during different seasons

within a year (varves – seasonal rhythmites – annually lami-

nated sediments). Varves have been described in glacial, lacus-

trine and marine environments. Clastic varves in glacial

lacustrine environments, i.e., glacial varves, are classic exam-

ples of varved sediments, but seasonal, rhythmic changes in

biogenic production (e.g., diatoms) or water chemistry (e.g.,

carbonate content) also give rise to varve formation. In addi-

tion, annual layers in ice cores are varves, and chemical varves

may be formed in precipitates and evaporates, such as stalag-

mites and gypsum beds.

Glacial varves

Varved clays are commonly found in Quaternary deposits in

glaciated areas, but they have also been identified in deposits

of older glaciations, such as the Carboniferous, Ordovician,

and the late and early Proterozoic.

Glacial varves reflect the seasonal fluctuation of glacial

meltwater flow from continental ice sheets or smaller glaciers

(Figure V1). During the intensive melting period in spring

and summer, coarse grained silt and sand is transported and

deposited at the bottom of a water body, whereas finer particles

settle out slowly, the finest clay particles only doing so during

the next winter when the water bodies are frozen. The bedding

in the coarse summer part of a varve is graded and the

boundary between the fine winter layer and the next summer

layer is sharp. Varves of this type are diatactic. They are formed

in fresh water, whereas symmictic glacial varves that are

formed in salt or brackish water are more complicated to

identify because fine suspended particles group together into

larger grains and settle rapidly with coarser grains. The varves

that are formed close to a melting glacier may be up to several

tens of centimeters thick, whereas varves deposited distally

may be less than a millimeter thick. Varve thickness also

varies from year to year due to variations in climatic and local

factors, e.g., the intensity of melting, the material supply and

the velocity of meltwater flow. Variations in the thickness of

varves within a sequence form a basis for correlating

sequences, in addition to their color, grain size and chemical

composition.

Nonglacial varve s

A single annual unit – a varve – is composed of two or more

layers (laminae) that can be distinguished on the basis of their

thickness, composition and texture. In most cases, these indivi-

dual layers are identified as representing certain seasons or

even short-term events within seasons, e.g., spring flood

allochthonous discharge or vernal precipitation of calcite crys-

tals (Figure V2). The annual structure of the sediment is only

preserved if there are no post-depositional disturbances of the

sediment surface, such as bioturbation or water turbulence in

the sediment-water interface.

In different lake environments worldwide, various character-

istic limnological and hydrological processes govern the com-

position and structure of varves and the laminae they contain.

They usually appear as calcareous (Brunskill, 1969), ferrogenic

(Anthony, 1977; Renberg, 1981), biogenic (Saarnisto et al.,

1977), clastic (Sturm, 1979) or clastic-organic (Renberg and

Segerström, 1981; Ojala et al., 2000)(Figure V3) laminae cou-

plets. The majority of identified varved lake records are located

either in North America (Anderson et al., 1985; Lamoureux

et al., 2001) or in Europe, particularly in the Alpine areas, in

Fennoscandia, and in number of maar lakes (Petterson, 1996;

Zolitschka, 1998; Brauer et al., 2000; Ojala et al., 2000). How-

ever, they are also found in many other localities, such as Japan

(Fukusawa, 1999) and Africa (Johnson et al., 2001).

Varved-clay chronology

Varved-clay chronology is based upon correlations between

short varve series (50–200 years, but rarely as much as 1,000

years) from site to site in the general direction of ice-margin

retreat. The original principle of De Geer (1884, 1940)was

adopted in Finland by Sauramo (1918). The Swedish glacial

varve chronology has been connected to the present-day through

studies on postglacial and recent, varved deltaic sediments in

the Ångermanälven river in northern Sweden (Cato, 1998).

The Swedish chronology covers 13,300 years, including the

deglaciation from Scania to the Ångermanälven river area.

The Swedish varve chronology has been constantly revised

since De Geer’s original work, and its correlation with the

Greenland ice core records suggests that nearly 900 (875) varves

have not been recognized in the Holocene part of the Swedish

varve chronology (Andrén et al., 1999). The Finnish varve

chronology covers nearly 2,800 years (Sauramo, 1929),

13,000–10,200 year, when the ice margin retreated from the

Gulf of Finland to the Gulf of Bothnia. It is a floating chronol-

ogy that has been connected to the calendar year time scale by

reference to characteristic varves formed in connection with the

drainage of the Baltic Ice Lake and the dating of this event

(e.g., Strömberg, 1990).

Attempts to construct a regional deglaciation chronology using

varved sediments outside Sweden and Finland have been rare. In

North America, Antevs (1928) worked out a varve chronology

based on 4,000 varves in Connecticut Valley, which was later

extended by ca. 1,000 glacial varves by Ridge and Toll (1999).

Glacial varves have also been investigated in a number of loca-

lities in Patagonia and other parts of northern Europe (Estonia,

NW Russia).

Varved sediment sequences have been widely applied to

reconstruct a continuous timescale of the sediment (varve

chronology) and to determine the rates of annual accumulation

of sediments and of any particular substance or organism in the

sedimentary basin, as expressed in spatial and temporal units

(influx). The longest known varved lake sediment sequence,

Lago Grande di Monticchio (southern Italy) (Brauer et al., 2000),

extends back to more than 100 ka interrupted by several tephra

layers, but continuous records typically cover ca. 10,000 years

or less. Among the most studied records in Europe are those

from Lake Holzmaar (western Germany) covering 14,000 years

(Zolitschka, 1998), Lake Gościąż (central Poland) covering

13,000 years (Goslar, 1998) and Lake Kassjön (northern Sweden)

covering 6,500 years (Renberg and Segerström, 1981). However,

some of these chronological sequences are floating in time (i.e.,

not fixed to the present day deposits). Moreover, even the best

quality varves cannot provide a completely accurate chronology

as they have been accumulated in a naturally variable environ-

ment. In most cases, and in the long-term, an acceptable margin

Figure V2 A simplified sketch of the formation of clastic-organic varves

in temperate areas.

Figure V3 Images of Lake Nautaja

rvi (central Finland) clastic-organic

nonglacial varves at ca.

AD 200 taken from the fresh sediment surface

(left) and thin-section using a back scattered mode of surface scanning

electron microscope (right).

Figure V1 Varved glacial clays in Tampere, southern Finland, which

were deposited in the early Holocene Baltic basin. (Photo by J. J.

Sederholm, 1907.)

974 VARVED SEDIMENTS

of varve counting error can be expected to be around 2%,

and at best it yields an accuracy of 1%, as in Lakes Kassjön

and Nautajärvi (central Finland) (Ojala and Saarinen, 2002).

Corresponding chronological accuracy has been postulated for

the Holocene part of Greenland ice core records (Alley, 2000).

Nonglacial Holocene varves have also been applied in assign-

ing ages in relative sediment dating methods, such as paleo-

magnetic reference curves and widely distributed tephra layers.

In Fennoscandia, Ojala and Saarinen (2002) and Snowball

and Sandgren (2002) dated the features of paleomagnetic

secular variation curves using the varve chronology of nearly

10,000-year-long sections of clastic-organic varves, whereas

Zillén et al. (2002) assigned varve ages to several tephra layers

deposited in southern Sweden in the Mid-Holocene.

Recently, the application of varve records has transformed

from providing stratigraphical and chronological tools to

being high-resolution indicato rs of the paleoenvironment,

and in particular, they are being used as a tool to study the

effects of climate forcing during the Holocene. Advancements

in sediment sampling (e.g., epoxy-embedding, thin section-

ing) combined with digital image analysis have provided an

effective means to record and study the physical properties

of varves with a high temporal resolution. Proglacial lakes

in the Canadi an Arctic are perhaps the most carefull y studied

environments (Hardy et al., 1996; Hughen et al., 2000;

Lamoureux et a l., 2001). Physical varve data (varve thickness,

laminae thickness) have been shown to be strongly correlated

with summer temperatures through their effects on the inten-

sity of snowmelt runoff and the discharge of suspended sedi-

ment. In Fennoscand ia, fluc tuatio ns of clastic-organic varves

have been correlated with winter climate components, namely

temperature and the amount of precipitated snow in the catch-

ment. Based on such connections, it is possible to reconstruct

past climatic conditions, thereby providing important kno wl-

edge about the global climate system in order to better predict

the future.

Matti Saarnisto and Antti E. K. Ojala

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Andrén, T., Björck, J., and Johnsen, S., 1999. Correlation of the Swedish

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Anthony, R.S., 1977. Iron-rich rhythmically laminated sediments in

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Brauer, A., Mingram, J., Frank, U., Günter, C., Schettler, G., Wulf, S.,

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oscillations during the Early Weichselian recorded at Lago Grande di

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Brunskill, G.J., 1969. Fayetteville Green Lake, New York. II. Precipitation

and sedimentation of calcite in a meromictic lake with laminated sedi-

ments. Limnol. Oceanogr., 14, 830–847.

Cato, I., 1998. Ragnar Lidéns post-glacial varve chronology from the

Ångermanälven valley, northern Sweden. Sveriges Geologisk Under-

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Cross-references

Continental Sediments

Dating, Magnetostratigraphy

Glacial Sediments

Ice Cores, Antarctica and Greenland

Lacustrine Sediments

Tephrochronology

VARVED SEDIMENTS 975