Odekon M. Encyclopedia of paleoclimatology and ancient environments

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throughout 1783. This haze is estimated to have extended to an

altitude of at least 16 km, resulting in substantial dimming of

the Sun.

during several of the summer months of the year 1783, when the

effect of the Sun’s rays to heat the Earth in these northern regions

should have been greatest, there existed a constant fog over all

Europe ....

this fog was of a permanent nature; it was dry, and the rays of the

sun seemed to have little effect towards dissipating it. (Benjamin

Franklin, 1784)

The eruption and the accompanying haze have been implicated

in the subsequent high temperatures recorded in July, 1783, and

the successive Northern Hemisphere cold winters of 1783/

1784 to 1785/1786 (for a more complete description, see Grattan,

2005). It is likely that the high July temperatures resulted

from the development of anticyclones over central or northern

Europe, allowing warm air to flow from southern Europe

(Thordarson and Self, 2004). Whether or not these changes in

large-scale weather systems were influenced by the Laki erup-

tion is not known. The winters of 1783/1784 to 1785/1786

were unusually cold, with an average temperature drop of

1.5 C over Europe and North America (Thordarson and Self,

2004). This temperature drop is consistent with the radiative

forcing predicted by the injection of SO

into the stratospheric

aerosol layer.

A full understanding of the Laki flood basalt eruption is

important for two reasons. Firstly, such eruptions are not

uncommon. On Iceland, for example, there have been four such

eruptions in the last 1,200 years, Laki merely being the latest.

Similar (but often larger) lava flows have built the Icelandic

crust, dating back more than 12 million years, and similar erup-

tions may occur in other basaltic provinces associated with

mantle plumes (e.g., Afar, Ethiopia; Hawaii). Secondly, the

eruption provides important information about much larger

events that have occurred in the geological past, and which

have been implicated in mass extinction events.

Columbia river flood basalt province

The Miocene flood basalts of the Columbia River Province in

northwest USA are among the best studied on Earth. The total

volume of the province is of the order of 175,000 km

, and

many of the individual flows have eruptive volumes in excess

of 1,000 km

. Most the province was erupted between about

16.5 and 14.5 Ma. It is estimated that the eruption of the Roza

flow (ca. 1,300 km

of basalt; ~14.7 Ma) produced a phenom-

enal mass of aerosols: approximately 12,000 Mt of SO

(equivalent to 23,000 Mt of H

), 680 Mt of Cl and

1690 Mt of F (Thordarson and Self, 1996). The effect of this

aerosol loading on the global climate is likely to have been sub-

stantial, but with our current understanding of the duration and

eruption intensity of large flood basalts, it is difficult to quan-

tify. Estimates of the duration of large flood basalt eruptions

vary from a few weeks (Shaw and Swanson, 1970) to a few

years (Self et al., 1996). Given the eruption profile of Laki,

where 80% of the magma was released during the first 3 months

of the eruption, it is possible that the bulk of the Roza magma

was also extruded within a relatively short time (months?),

even if the eruption subsequently continued to be active for

several years. If this is the case, then a substantial proportion of

the SO

could have been injected into the stratosphere atop vig-

orous fire fountains and Plinian-style eruption columns (Stothers

et al., 1986). Repeated SO

and dust injection, over a period of

several years, may also have occurred. Substantial cooling of at

least the Northern Hemisphere troposphere would have ensued,

but given that the response of the global climate to such (cumu-

lative) mass loading is likely to be strongly non-linear, the precise

outcome is difficult to predict. Immediate and dramatic cooling

(for at least the duration of the eruption) would have been likely;

transient “volcanic winters” could have ensued, and acid rain

would have cause local environmental devastation.

Even more difficult to predict is the global response to succes-

sive flood basalt eruptions. The Columbia River Province com-

prises some 300 flow units, mostly erupted over a period of

about 2 Ma. This gives an average interval between each flood

basalt eruption of 7,000 years. The repose time between indi-

vidual eruptions is clearly sufficient for the climate to recover,

at least from the cooling effects of sulfate injection, although

Thordarson and Self (1996) have suggested that the eruption

of the Columbia River flood basalts was a contributory

factor in the Miocene global cooling event that began at

about 15–14.5 Ma.

Not only did the flows release large quantities of SO

however, but they also vented prodigious masses of CO

which, being a greenhouse gas, could have caused global

warming. Are these masses significant? Estimates of the CO

content of basaltic magma vary dramatically, because the gas

has a low solubility in the melt at near-surface conditions,

and is rapidly exsolved. However, based on pre-eruptive values

of 7,000 ppm CO

in basalts from ocean island volcanoes

(Bureau et al., 1999), and assuming that 99% of the CO

is lost

to air, then 1 m

of basalt will release a little over 18 kg of CO

Using these figures, the total amount of CO

released from the

Laki eruption was of the order of 270 Mt (or 73 Mt carbon),

and from the Roza flow, a little over 23,000 Mt (6,300 Mt C).

As the residence time of CO

in the atmosphere is much greater

than that of SO

, knowledge of the eruption profile is less impor-

tant than it is for SO

The current anthropogenic output is approximately

7,000 Mt C yr

1

. The mass erupted from Laki was insignificant,

on a global scale. The flux from the Roza flow was likely to

have been enhanced by burning vegetation, etc., but even

assuming that the bulk of the CO

was degassed in the early

stages of the eruption, the flux rates were not substantially

greater than today’s anthropogenic output. Note that total C

release from the whole of the eruptive portion of the Columbia

River basalt province was on the order of 800,000 Mt, a see-

mingly vast amount, but given that the eruption occurred over

a period about 2 million years, the average flux was only

0.4 Mt C yr

1

, a fraction of the current anthropogenic output.

Flood basalts and mass extinctions

The “ big three” mass extinctions of the last 300 Ma all coincide

(within the precision of current dating techniques) with major

flood basalt events. These are the Siberian Traps with the

end-Permian extinction at 250 Ma; the Central Atlantic Mag-

matic Province with the end-Triassic extinction (200 Ma); and

the Deccan Traps with the end-Cretaceous (K-T) extinction

(65 Ma; Figure F1). The likelihood of these associations being

pure chance is remote: about 1 in 2,000 (White and Saunders,

2005), and several other volcanic provinces – either continental

or oceanic – are contemporaneous with other mass extinction

events and other evidence of significant changes in ocean chem-

istry (for example, formation of black shales during periods

of oceanic anoxia at 90 Ma; Kerr, 1998). Both the Siberian and

336 FLOOD BASALTS: CLIMATIC IMPLICATIONS

Deccan Traps are substantially larger than the Columbia River

Province, with eruptive volumes of the order of 3.5 million km

(Reichow et al., 2002) in the case of the former, and around

0.75 million km

in the case of the Deccan Traps. The total erup-

tion duration may have been as short as 0.5 Ma, but is likely to

have been longer than this; there is considerable uncertainty

about the duration of volcanic activity in both provinces. Assum-

ing this “worst case” scenario (0.5 Ma), and using the same

volatile concentrations as in the Roza flow, produces average

fluxes of 20 and 34 Mt C yr

1

for the Deccan and Siberan Traps,

respectively, and 38 and 66 Mt SO

1

for the two provinces. At

face value it is difficult to see how these fluxes could lead directly

to global ecosystem collapse; they are too low. Magmatic fluxes

are far from regular, however, and successive flows may occur

sufficiently frequently that SO

could accumulate in the

atmosphere over several years or decades. This would lead to

substantial climate forcing and, initially, tropospheric cooling.

Meanwhile, simultaneous release of CO

could, on a longer time-

scale, lead to an increase in global temperatures. This in turn

could trigger the catastrophic release of carbon from terrestrial

and marine methane hydrates, as indicated by the isotopically

light carbon signature found in several horizons associated with

mass extinction and flood basalt events (Wignall, 2001).

Conclusions

From the historical observations of the 1783 Laki event, it is

evident that even a relatively small basalt eruption is capable

of causing climate change on a hemispheric scale. This is in

addition to the local devastation caused by acid rain and release

of a range of toxic gases. The height of the convection column

associated with basaltic eruptions is less than with major explo-

sive events, but injection of gas and dust into the stratosphere is

likely given the high magmatic flux rates. Unlike explosive

eruptions, flood basalts can erupt for several months if not years,

leading to prolonged discharge of volatiles. The largest docu-

mented flood basalts indicate release of prodigious volumes of

gas, especially SO

, into the atmosphere, leading to the likeli-

hood of substantial, short-term cooling of the lower atmosphere.

Simultaneous release of carbon dioxide may cause longer-term

global warming if the excess volcanic CO

cannot be removed

(e.g., by biological activity) sufficiently quickly, and if flood

basalt eruptions occur in rapid succession. This may trigger

release of methane hydrates and initiate collapse of global

ecosystems. Detailed modeling of the effects of flood basalt

eruptions on global climate is required to ascertain the full

significance of volcanic forcing.

Andrew D. Saunders

Bibliography

Bureau, H., Métrich, N., Semet, M.P., and Staudacher, T., 1999. Fluid-

Magma decoupling in a hot-spot volcano. Geophys. Res. Lett., 26,

3501–3504.

Coffin, M.F., and Eldholm, O., 1994. Large igneous provinces: Crustal struc-

ture, dimensions, and external consequences. Rev. Geophys., 32,1–36.

Courtillot, V., 1999. Evolutionary Catastrophes: The Science of Mass

Extinction. Cambridge: Cambridge University Press, 173pp.

Grattan, J., 2005. Pollution and Paradigms: Lessons from Icelandic volcan-

ism for continental flood basalt studies. Lithos, 79, 343–353.

IPCC, 2001. Climate Change 2001: The Scientific Basis. Houghton, J.T.,

Ding, Y., Griggs, D.J., Noguer, M., van der Linden, P.J., Dai, X.,

Maskell K., and Johnson C.A. (eds.), Intergovernmental Panel on Cli-

mate Change. Cambridge, UK: Cambridge University Press.

Kerr, A.C., 1998. Oceanic plateau formation: A cause of mass extinction

and black shale deposition around the Cenomanian-Turonian boundary?

J. Geol. Soc. London, 155, 619–626.

McLean, D.M., 1985. Deccan traps mantle degassing in the terminal

Cretaceous marine extinctions. Cret. Res., 6, 235–259.

Figure F1 Extinction rate versus time (multiple-interval marine genera: modified from Sepkoski, 1996) compared with eruption ages of

continental flood basalt provinces. Three of the most severe extinctions, the P-Tr, the Tr-J, and the K-T, correspond with eruption of the

Siberian Traps, Central Atlantic Magmatic Province, and Deccan Traps, respectively. Oceanic plateaus may also have had profound environmental

consequences (Kerr, 1998)(S: Sorachi Plateau, Japan; K: Kerguelen Plateau; OJ: Ontong Java Plateau; C: Caribbean-Colombian) (modified from

White (2002) and White and Saunders (2005)).

FLOOD BASALTS: CLIMATIC IMPLICATIONS 337

Oppenheimer, C., 2002. Limited global change due to the largest known

Quaternary eruption, Toba, ~74 kyr BP. Quaternary Sci. Rev., 21,

1593–1609.

Reichow, M.K., Saunders, A.D., White, R.V., Pringle, M.S.,

Al’Mukhamedov, A.I., Medvedev, A., and Korda, N., 2002. New

Ar/

Ar data on basalts from the West Siberian Basin: Extent of the

Siberian flood basalt province doubled. Science, 296, 1846–1849.

Self, S., Thordarson, T., Keszthelyi, L., Walker, G.P.L.,Murphy, M.T.,

Long, P.E., and Finnemore, S., 1996. A new model for the emplacement

of Columbia River basalt as large, inflated, pahoehoe lava sheets.

Geophys. Res. Letts., 23, 2689–2692.

Sepkoski, J.J., 1996. Patterns of Phanerozoic extinction: A perspective

from global data bases. In Walliser, O.H. (ed.), Global Events and Event

Stratigraphy. Berlin: Springer, pp. 35–51.

Shaw, H.R., and Swanson, D.A., 1970. Eruption and flow rates of flood

basalt. In Gilmore, E.H., and Stradling, D.F. (eds.), Proceedings of the

Second Columbia River Basalt Symposium. Cheney: East Washington

State College Press, 271–299.

Stothers, R.B., Woolf, J.A., Self, S., and Rampino, M.R., 1986. Basaltic

fissure eruptions, plume heights, and atmospheric aerosols. Geophys.

Res. Letts., 13, 725–728.

Thordarson, Th., and Self, S., 1996. Sulfur, chlorine, and fluorine degas-

sing and atmospheric loading by the Roza eruption, Columbia River

Basalt Group, Washington, USA. J. Volcanol. Geotherm. Res., 74,

49–73.

Thordarson, Th., and Self, S., 2004. Atmospheric and environmental effects

of the 1783–1784 Laki eruption: A review and reassessment.

J. Geophys. Res., 108, 1 (20030116), 29pp.

Thordarson, Th., Self, S., Óskarsson, N., and Hulsebosch, T., 1996. Sulfur,

chlorine, and fluorine degassing and atmospheric loading by the

1783–1784 AD Laki (Skaftár Fires) eruption in Iceland. Bull. Volcanol.,

58, 205–222.

White, R.V., 2002. Earth’s biggest ‘whodunnit’: unravelling the clues in the

case of the end-Permian mass extinction. Philos. Trans. R. Soc. Lond,

360, 2963–2985.

White, R.V., and Saunders, A.D., 2005. Volcanism, impact and mass

extinctions: Incredible or credible coincidences? Lithos, 79, 299–316.

Wignall, P.B., 2001. Large igneous provinces and mass extinctions. Earth

Sci. Rev., 53,1–33.

Cross-references

Aerosol (Mineral)

Atmospheric Evolution, Earth

Carbon Isotopes, Stable

Climate Forcing

Cretaceous/ Tertiary (K-T) Boundary Impact, Climate Effects

Evolution and Climate Change

Methane Hydrates, Carbon Cycling, and Environmental Change

Volcanic Eruptions and Climate Change

FORAMINIFERA

Foraminifera are single-celled eukaryotic organisms that live

in both the marine and fresh water environment. They range in

size from 100 mm–15 cm in length. Foraminifera (often referred

to as forams) are classified primarily by the composition and

morphology of their tests (shells). Tests can be made of organic

compounds, sand grains, and other particles cemented together

(agglutinated), or secreted calcium carbonate (CaCO

). Many

groups are commonly made of a number of chambers, added dur-

ing growth (Figure F2). The arrangement of these chambers and

the position and shape of apertures are important for taxonomic

classification.

Foraminifera exhibit both benthic (live on (epifaunal)or

in (infaunal) the substrate) and planktonic (passive floating)

lifestyles. There are approximately 40 species of planktonic

foraminifera in the ocean today. This comprises about 1% of

the extant species of foraminifera. In the open marine environ-

ment, benthic foraminifera are usually much more diverse yet

occur in lower abundances than planktonic foraminifera.

The geologic history of fossil foraminifera begins in the ear-

liest Cambrian. Foraminifers probably existed as cells without

fossilizeable coverings long before that. In the Jurassic, forami-

nifers adapted to a planktonic mode of life.

Foraminifers were probably first recorded by the Greek his-

torian Herodotus in the fifth century

BC in his observations of

large disc-shaped fossils (now known as the benthic foramini-

fer Nummulites gizehensis) found in the building stones of

the Egyptian pyramids (Haynes, 1981). Prior to the advent of the

microscope, seventeenth century naturalists observed foraminifers

with hand lenses and often classified them as gastropods or

cephalopods because many possess a coiled chamber arrange-

ment. As microscopy advanced, more detail in the structure of

foraminiferal tests was observed and they were recognized as

single-celled organisms.

Foraminiferal studies advanced with the advent of the Chal-

lenger expedition of 1872–1876. Brady (1884) illustrated a

number of foraminifers from this first oceanographic cruise and

our understanding of foraminiferal biogeography began to take

shape. During the early to mid twentieth century, the petroleum

industry’s need for age control and environmental interpretation

led to great advances in micropaleontology, more in general, and

much more data on foraminifers. The next major advance was the

development of scanning electron microscopy, which enabled

detailed analysis of test wall ultrastructure for the first time.

The Deep Sea Drilling Project and its successor, the Ocean Dril-

ling Program, have provided a wealth of quantitative data on both

temporal and geographic distribution of benthic and planktonic

foraminifers.

Applications in paleoclimatology

Foraminifers have been widely used in several areas of paleo-

climatology. Calcium carbonate tests are thought to be secreted

in equilibrium with sea water at the time of formation. For this

reason, foraminifers have been the most widely used organisms

for shell chemistry (Mg:Ca) and stable isotope geochemistry

O, d

C). Their small size, abundance, and distribution

make them ideal environmental indicators. With the knowledge

Figure F2 SEM image of the planktonic foraminifers Globigerina

bulloides d’Orbigny (left) and Dentoglobigerina altispira Cushman and

Jarvis (right) showing apertures (a) and chambers (b).

338 FORAMINIFERA

that different species of planktonic foraminifers live at different

depths and reach maximum abundance at different times of

the year, geochemical analyses of foram tests can provide a

wealth of information on the composition and structure of

the water column. With benthic foraminifers providing informa-

tion on bottom waters, elaborate paleoenvironmental and paleo-

climatic information can be obtained from samples containing

foraminifers.

Quantitative analysis of assemblages of planktonic fora-

minifers has been widely used to reconstruct past sea-surface

temperatures using a transfer function or nearest analog

approach. Transfer functions range from simple (Ericson and

Wollin, 1956) to complex factor analytic approaches pioneered

by Imbrie and Kipp (1971). This later technique formed the basis

of the CLIMAP Last Glacial Maximum reconstruction (CLI-

MAP, 1981) and has been modified and used to reconstruct

paleoclimate conditions during other time periods back to the

mid-Pliocene warm period (Dowsett et al., 1999). With the

recognition that planktonic foraminiferal distribution on the sea

floor closely matches overlying oceanic conditions, total faunal

analysis has become a mainstay of paleoenvironmental and

paleoclimatic reconstruction.

The temporal framework upon which paleoclimate recon-

structions are based is of fundamental importance. The same

characteristics that make planktonic foraminifers so useful for

geochemical analysis-high abundance, small size, facies inde-

pendence, ease of identification and high preservation

potential-make foraminifers ideal for biochronological (includ-

ing biostratigraphy and correlation of marine sediments)

analyses. For any pre-Pleistocene study, foraminiferal biochro-

nology is usually still one of the leading tools in terms of accu-

racy and utility. Planktonic foraminifer zonations have been

developed by a number of workers and individual biotic events

(first and last appearances in a temporal sense) have been dated

by calibration to magnetostratigraphy (Berggren et al., 1985,

1995; Dowsett, 1989).

In summary, foraminifers are the original multi-tool for paleo-

climatology. The quantitative distribution of foraminifers in both

temporal and spatial dimensions provides information on age

of sample and any number of aspects of the paleoceanographic

environment. When combined with geochemical and isotopic

analysis of foraminifer tests, no other fossil group has provided

so much to the field of paleoclimatology.

Harry J. Dowsett

Bibliography

Berggren, W.A., Kent, D.V., and Van Couvering, J.A., 1985. Neogene

geochronology and chronostratigraphy. In Snelling, N.J. (ed.), The

Chronology of the Geological Record. London: Geological Society of

London, Memoir 10, pp.211–260.

Berggren, W.A., Kent, D.V., Swisher, C.C., and Aubry, M.-P., 1995. A

revised Cenozoic geochronology and chronostratigraphy . In Berggren, W.A.,

Kent, D.V., Aubry, M.-P., and Hardenbol , J. (eds.), Geochronology , Time

Scales and Global Stratigraphic Correlation. Tulsa: Society for Sedimentary

Geology , Special Publication 54, pp.129–212.

Brady, H.B., 1884. Report on the foraminifera dredged by H.M.S. Challen-

ger, during the years 1873–1876. Report of the scientific results of the

voyage of H.M.S. Challenger, 1873–1876, Zoology, 9,1–814.

CLIMAP, 1981. Seasonal reconstruction of the Earth’s surface at the last

glacial maximum. Geological Society of America Map and Chart Series

MC-36, 1–18.

Dowsett, H.J., 1989. Application of the graphic correlation method to

Pliocene marine sequences. Mar. Micropaleontol., 35, 279–292.

Dowsett, H.J., Barron, J.A., Poore, R.Z., Thompson, R.S., Cronin, T.M.,

Ishman, S.E., and Willard, D.A., 1999. Middle Pliocene paleoenviron-

mental reconstruction: PRISM2. U.S. Geological Survey Open File

Report 99–535.

Ericson, D.B., and Wollin, G., 1956. Correlation of six cores from the equa-

torial Atlantic and the Caribbean. Deep Sea Res., 3, 104–125.

Haynes, J.R. 1981. Foraminifera. New York: Wiley, 433pp.

Imbrie, J., and Kipp, N.G., 1971. A new micropaleontological method

for quantitative paleoclimatology: Applications to a late Pleistocene Car-

ibbean core. In Turekian, K.K. (ed.), The Late Cenozoic Glacial Ages.

New Haven: Yale Univ. Press, pp.71–181.

Cross-references

Carbon Isotopes, Stable

CLIMAP

Dating, Biostratigraphic Methods

Dating, Magnetostratigraphy

Deep Sea Drilling Project (DSDP)

Ocean Drilling Program (ODP)

Oxygen Isotopes

Transfer Functions

FORAMINIFERA 339

GEOCHEMICAL PROXIES (NON-ISOTOPIC)

Introduction

The utility of geochemical proxies requires chemical analysis

of geological samples. Since there is no means of directly mea-

suring past climates and environments, proxy indicators are

employed as indirect measures of the main processes. Informa-

tion for environmental reconstruction comes from both organic

and inorganic sources, even though organic matter (OM) typi-

cally constitutes a minor fraction of sediments. These proxies

may reveal details about paleotemperature, phytoplankton com-

munity structure, vegetation history, dust provenance, nutrient

cycles and availability, ocean circulation and paleoredox condi-

tions (Figure G1).

Technical terms and concepts

Since sedimentary organic matter is the residue of past biota, its

type and quantity reflects environmental conditions affecting

past ecosystems. Analysis of organic matter involves identify-

ing both its general source from bulk elemental compositions,

stable isotopes and Rock-Eval pyrolysis data, and its detailed

origins from analyses of biomarker molecular compositions

(see below). Changes in organic composition can be interpreted

as proxies for varying sea-level, ocean currents, temperature

and continental climates, while changes in abundance may sig-

nify variable paleoproduction or preservation conditions.

Inorganic proxies are typically measured from constituents of

the hard parts of previously living biota such as calcitic plankton

and coral. The rate of incorporation of metal ions as compared with

the principal component calcium varies according to environmental

conditions. Therefore, proxies for past temperature and nutrients

are derived from metal ratios such as Mg/Ca, Sr/Ca and Cd/Ca.

Complementary information from multiple proxies can be obtai-

ned during a single suite of analyses for major and minor metal

abundances.

The longevity of a particular record has a profound effect

upon its constituent components. Geochemically, these compo-

nents can be classified as either labile (readily undergoing

change or breakdown), or refractory (resistant to change, espe-

cially to heat). As a deposit ages, more refractory components

become relatively enriched, while labile components are scar-

cely preserved. Diagenesis causes both the concentration and

composition of sedimentary organic matter to diverge from

the original biologically synthesized material. The lability of

geochemical components renders them more susceptible to

diagenesis. Ultimately, the geochemical composition of a paleo-

environmental record may look a lot different from the time

of its creation or original deposition. Comparison of multiple

geochemical proxies helps to compensate for the effects of

diagenetic alteration and improves their interpretation for a

more complete paleoenvironmental reconstruction.

Significance of geochemical proxies

The faithful reconstruction of paleoclimates and paleoenviron-

ments requires the availability of reliable records. For place-

ment in a temporal framework, the age of a particular record

must be known, or assumptions regarding its continuity must

be made. While “geochemical fossils” exist and are readily

exploited for their paleoenvironmental data, fossil records are

typically incomplete. Furthermore, the preservation of these

records is rarely uniform. For this reason, continental (terres-

trial) records of paleoenvironmental change (e.g., coal, evapor-

ites, loess) have generally been regarded as discontinuous or a

palimpsest, written and rewritten throughout the history of geo-

logic time. Sedimentary records from lakes and other terrestrial

depositional environments are typically Holocene (post-glacial) in

age, although some low-latitude lakes (e.g., Lake Tanganyika) and

cave deposits (e.g., speleothems) have yielded longer records.

Conversely, vast sediment deposits underlying the world’soceans

may have been accumulating quasi-periodically for millions of

years. For this reason, the majority of geochemical proxies are

derived from aquatic-environment sediments, though many tech-

niques may be universally applicable. Deep-sea sedimentation

rates are typically a few centimeters per 1,000 years and with

few exceptions reveal only decadal to millennial-scale changes

in paleoclimate. Conversely, lacustrine records may offer short

and regional records of paleoenvironmental change at a seasonal

or annual scale.

Bulk organic geochemistry

Organic molecules are created in the biosphere by biological pro-

cesses. The geosphere (lithosphere, hydrosphere and atmosphere)

provides the substrate for these biological processes and is the ulti-

mate sink for residual organic molecules. Organic geochemistry

reveals the origin and fate of organic molecules and their utility

as geochemical and paleoenvironmental proxies as they are pre-

served in the geosphere. Since both production and preservation

of organic matter are affected by environmental change, bulk

sedimentary organic matter and a large number of biomarker com-

pounds, together with patterns of compound classes and their iso-

topic compositions, have enabled organic geochemists to deduce

much about the environmental conditions that impacted eco-

systems at different times and places in the geological past.

The primary source of organic matter in aquatic systems

is planktonic photosynthetic production. Large variations in

the rate of production are effected by nutrient availability

in surface waters. Availability is controlled by nutrient (nitrate,

phosphate) transport from land to lakes and oceans, and the

efficiency of nutrient recycling within these systems. Animals

and microbes may use such primary organic matter as a source

of carbon and nutrients for significant secondary production of

reworked, degraded organic matter with its own distinctive

geochemical signature. Detritus from land plants and a small

amount of organic matter from animals may make additional

contributions to sedimentary organic content. Therefore, deter-

minations of the relative importance of algal productivity, land-

plant productivity and transport processes are a critical part of

paleoenvironmental reconstruction.

The information obtained from organic geochemical proxies

is diverse. Mass accumulation rates of organic carbon reveal

the biological paleoproductivity and the environmental condi-

tions that influenced the preservation of this organic matter.

The composition of organic matter can be informative about

aquatic and continental vegetation, which in turn depends on

climatic factors such as temperature, moisture availability and

wind strength. Quantities and isotopic contents of sedimentary

organic matter (carbon, nitrogen) provide detailed histories of

elemental cycling at both local and global scales. Environmental

(e.g., sea surface) temperatures can be deduced from the sedi-

mentary molecular residues of organisms that lived at such

times in the past. The typically small eolian (wind-derived)

fraction of an organic mixture may contain distinctive compo-

nents that record atmospheric transport processes. Comparison

of multiple indicators (the “multi-proxy” approach) typically

leads to better, more confident interpretations.

Preservation of the organic signal

Of the proxy indicators that depend on fluxes, bulk organic

carbon accumulation measurements are perhaps the most widely

used and common. Organic matter is intrinsically unstable in oxi-

dizing environments and degrades in the photic zone, during par-

ticulate sinking and subsequent sediment bioturbation. Many

attempts to relate total sedimentary organic input to surface net

export have been made since the ground-breaking work of Suess

(1980) and Berger et al. (1987). It has been shown that after remi-

neralization, rarely more than a few percent remains to form the

organic geochemical sedimentary record. The fraction of carbon

buried is primarily a function of water depth, deep water oxygena-

tion and linear sedimentation rate (Müller and Suess, 1979). Away

from potential allochthonous (e.g., downslope) sources and under

oxygenated states, environmental conditions cause the organic

matter flux (amount per unit time) surviving degradation to

decrease with increasing water column (depth) due to microbial

oxidation (Figure G2). Degradation rates are greatest for fresh

organic matter, but decline with depth as the residual detritus con-

tains progressively more resistant organic matter. If a significant

proportion of the accumulating sediment constitutes lithogenic

particles, then this additional flux (ballast) may improve preserva-

tion of the organic flux to the seafloor. An important bias in the

source of sedimentary carbon can affect the record of low (export)

productivity zones, such as the inner ocean, where lateral advec-

tion from the continental shelf is often significant.

Former rates of primary production can be estimated from

the organic carbon content of marine sediments according to

the productivity-depth relationship and knowledge of the mass

accumulation rate. For example, Stein (1986) proposed

Figure G1 Summary of principal geochemical proxies and derived information used for environmental reconstruction.

342 GEOCHEMICAL PROXIES (NON-ISOTOPIC)

Paleoproductivity; gC m

2

1

¼ 5:31 ½C

org

ðDBDÞ

0:71

 SR

0:07

 WD

0:45

ð1Þ

where C

org

is weight percentage of marine organic carbon,

DBD is dry bulk density in g cm

3

, SR is sedimentation

rate in cm kyr

1

, and WD is water depth in meters at time of

deposition.

The oxidationof organic matter and animal respiration removes

dissolved oxygen in aquatic environments. Where this rate of

removal exceeds the rate of replenishment (from the atmosphere),

water column anoxia may develop. Under conditions of high pri-

mary production or strong water-column stratification, organic

matter fluxes do not appear to decrease with depth (Karl and

Knauer, 1991) and sedimentary preservation may be improved

(Figure G3).

The rate of delivery of organic matter to the surface layer of

sediments effectively determines the amount of benthic altera-

tion in the bioturbated layer. Benthic oxidation will initially uti-

lize interstitial oxygen, but as conditions become progressively

more anoxic, nitrate, oxides of manganese, iron and eventually

sulfate reduction will occur. These reduced species may be later

interpreted as geochemical proxies of anoxia.

Physical sorting during transport and sinking of organic mat-

ter can create compositional artifacts since different components

are commonly associated with different particle sizes. This com-

plicates the interpretation of co-occurring sedimentary proxies,

which may not share the same temporal or spatial origins.

For example, C/N ratios of OM can be affected by hydrodynamic

sorting: they are generally lower in fine-grained than in coarse-

grained sediments, which typically contain more intact land-

plant debris. Fine sediments may contain more clay with greater

potential for ammonium adsorption. This highlights the problems

of bulk analysis, especially where low organic content results in

samples with a significant inorganic nitrogen fraction. Molecu-

lar-level analyses are not immune to such biasing of the source

information they yield; organic molecules with radiocarbon ages

several thousand years older than co-occurring calcitic plankton

remains have been found in drift-deposit sediments from the

Bermuda Rise (Ohkouchi et al., 2002). Furthermore, geo-lipids

typically constitute a small fraction of the residual (sedimentary)

organic matter, making insidious abundance variations more com-

mon. Although biomarkers are relatively resistant to microbial

degradation, they may still be subject to diagenetic overprinting.

Analysis of bulk organic matter

Petrographic analysis of sedimentary records can reveal the origin

of organic matter, its mode of transport and degree of degradation

through microscopic study. Unstructured, amorphous organic

flakes and intact debris of algae such as cell walls constitute

autochthonous organic matter. Allochthonous material includes

variably-preserved lignin and cellulose debris and soil-derived

organic material. Spores, pollen and the residue of forest fires

transportedbywindformtheeolian organic fraction. By determin-

ing the petrographic composition of sedimentary organic matter, it

is possible to supplement information provided by other techni-

ques concerning changes in origins of organic matter over time

and sedimentary intervals where diagenesis may have altered or

biased the bulk geochemical indicators.

Figure G3 Distribution of total organic carbon in surface sediments of

the Arabian Sea as a function of water depth. The shaded area denotes

the approximate depth range of the present oxygen minimum zone

(OMZ) (reprinted from Organic Geochemistry, 31, Schulte S.,

Mangelsdorf K., and J. Rullko

tter, Organic matter preservation on the

Pakistan continental margin as revealed by biomarker geochemistry,

Figure G2 Concentrations of organic carbon on sinking particles

collected by sediment traps at various depths in the ocean (figure

compiled from Meyers, 1997; reprinted from Organic Geochemistry, 27,

Meyers P.A., Organic geochemical proxies of paleoceanographic,

paleolimnologic, and paleoclimatic processes, pp. 213–250, Copyright

(1997), with permission from Elsevier).

GEOCHEMICAL PROXIES (NON-ISOTOPIC) 343

Algal and terrestrial plant origins of sedimentary organic

matter can be distinguished on the basis of its elemental C/N

ratio. Reliable source information can be derived in moderately

preserved organic matter since algae typically have atomic C/N

ratios of 4–10, while vascular plants (that is, those composed in

part of vascular tissue, including all flowering plants and the

higher cryptogamous plants) have C/N ratios of 20 (Meyers,

1994). This difference relates to the proteinaceous nature of

algal matter compared with the abundance of cellulose in (vas-

cular) plants. Diagenesis has the potential to alter the C/N

ratio, causing it to increase during preferential degradation of

proteinaceous components of algal OM, but decrease during

carbon remineralization due to preferential absorption of

ammonium(-nitrogen) by clay minerals (Müller, 1977).

Rock-Eval pyrolysis

Quantification of hydrocarbons generated by progressive heating

of sediment samples from 200



Cto600



C, related to the origin

of OM, generates three distinct signals related to the quantity of

hydrocarbons, which are gaseous (S

), volatile (S

) and released

by thermal cracking of kerogen (S

). Total organic carbon is the

sum of the residual and pyrolyzed organic carbon (S

) content.

From these results are derived: (a) the Hydrogen Index (HI), the

hydrocarbon potential of the total OM (mg HC/g C

org

), a proxy

for the H/C of OM; and (b) the Oxygen Index (OI), the amount

of oxygen (mg CO

/g C

org

), a proxy for the O/C ratio (Espitalié

et al., 1977). Values are typically plotted against each other on a

van Krevelen plot and grouped according to three main types of

organic matter: (a) rich in HC content, derived from microbial bio-

mass or waxy coatings of plants; (b) moderately rich in HC, origi-

nating from algae; (c) poor in hydrocarbons, rich in carbohydrates,

typical of woody plant material (Figure G4). Values may be

affected by organic matter oxidation, particularly after deposition.

Oxidation causes a reduction in total organic carbon (TOC)

concentration, C/N ratio and HI values while oxygen content

increases, thereby reducing HI/OI and causing Type I or II organic

matter to resemble Type III.

Bulk inorganic geochemistry

Bulk geochemical proxies such as opal, carbonate and quartz

dust have been traditionally employed to attempt to reconstruct

paleoclimate. However, more recent advances in understanding

the constraints on delivery and preservation of these sedimen-

tary components limit the usefulness of the data obtained. An

alternative technique that overcomes many of the problems of

bulk organic matter analysis involves the application of rain

ratios between the organic and inorganic components of the

sediment record. Berger and Keir (1984) proposed that ratios

of organic carbon:carbonate and organic carbon:opal may

increase with export production. The ratio of detrital to bio-

genic content may also be employed to estimate paleoproduc-

tivity. The concentrations of certain inorganic proxies, such

as aluminum and potassium, have been measured to represent

the major part of the detrital sediment content. Found mostly

in clays, their sediment depth profile is considered typical

for the bulk detrital elements. Thus, the content of elements

sensitive to productivity changes or inherently involved in bio-

logical and authigenic processes, such as cadmium, barium and

silicon, can be successfully normalized with respect to the det-

rital signal. One limitation of this technique is that it is

restricted to areas of predominantly clay deposition, and cannot

thus be applied on the continental margins where silt is com-

monly the dominant matrix size fraction. Nevertheless, several

such rain ratios are commonly used. For example, cadmium has

been found to be depleted in the surface ocean relative to dee-

per waters (Boyle et al., 1976), indicating uptake by organisms

at the surface and regeneration from deeper sinking biomass.

Profiles of cadmium generally resemble those of phosphate

(especially enriched in upwelling regions), rather than silicate,

confirming shallow cycling in association with labile nutrients.

Inorganic nutrient indicators

More detailed information for the reconstruction of upper

ocean nutrient concentrations and water mass formation and

circulation has been retrieved from trace metal/calcium ratios

in the calcitic tests of planktonic organisms such as foramini-

fera and coccolithophorids. The ratio of Cd:Ca in carbonate

plankton tests has been used to reconstruct the cadmium con-

tent of the water in which the organisms grew and to detect

changes in the position of the nutricline relative to the thermo-

cline and water column stratification. The ratio of Si:Al content

of biogenic opal has been similarly used (Martinez et al.,

1996). Cd resembles phosphate in terms of its behavior in sea-

water (Figure G5). Therefore, measurement of Cd/Ca ratios in

the tests of shallow water column foraminiferal species has

been employed as a paleonutrient (PO

3–

) proxy in thermocline

waters (Boyle, 1988).

The response of benthic foraminiferal cadmium to deep-water

Cd concentrations has also been established (Boyle, 1992).

Benthic shell Cd/Ca and Ba/Ca have been used to interpret

changes in deep-water circulation since the long residence

time of Cd limits the impact of changes in mean ocean Cd on

glacial/inter-glacial timescales (<5%, Rosenthal et al., 1995),

relative to the short residence time of Ba (10 kyr). Coupled

Cd/Ca and Ba/Ca records have proven useful because Cd/Ca is

a labile nutrient tracer and Ba/Ca is a refractory nutrient tracer,

Figure G4 Rock-Eval van Krevelen plot for differentiating sources of

sedimentary organic matter.

344 GEOCHEMICAL PROXIES (NON-ISOTOPIC)

each involved differently in biogeochemical cycling. However,

there is an apparent depth-related change in the partition coeffi-

cient of Cd in calcitic foraminifera. The effects of algal symbionts

(Mashiotta et al., 1997) and temperature (Rickaby and Elderfield,

1999) may also influence such estimates.

Trace metal ratios in foraminifera are subject to elevation in

instances of (sedimentary) secondary overgrowth or carbonate

precipitation on shell surfaces, which can trap contaminants

between the shell and the overgrowth layer. Such overgrowths

form where Mn

is mobilized (under oxic conditions), and can

therefore be detected in shells with relatively elevated Mn/Ca

values (Boyle et al., 1995).

Inorganic temperature indicators

Strontium and magnesium occur in seawater in approximately

constant proportions to calcium, and all three elements have long

oceanic residence times. Based on this constant source ratio, var-

iations in Sr/Ca and Mg/Ca in foraminifera shells are due to the

influence of environmental conditions that control their incor-

poration or preservation during the process of calcification.

Experimental evidence indicates that temperature exerts the

dominant control on shell Mg/Ca. Cultured planktonic foramini-

fera Globigerina bulloides and Orbulina universa exhibit an

increase in shell Mg/Ca of 8–10%/



C (Nürnberg et al., 1996;

Lea et al., 1999). This temperature regulation of co-precipitation

of shell Mg probably involves both a thermodynamic response of

the inorganic distribution coefficient, and a temperature effect

on physiological processes regulating cellular Mg uptake. Sec-

ondary influences are attributable to seawater pH and salinity.

A smaller but similar kinetic influence of calcification has been

recorded in Sr/Ca of coccoliths. However, global foraminiferal

Sr/Ca data exhibit a linear decrease with increasing water

depth, suggesting that the Sr partition coefficient is most likely

a pressure related phenomenon (Rosenthal et al., 1997). Inter-

pretations based on Sr/Ca are also complicated by the much

smaller range of Sr/Ca values than Mg values in shells

(Figure G6).

Records of Mg/Ca ratios in benthic foraminifera also have

been interpreted as an independent proxy for reconstructing the

history of Quaternary bottom water temperatures (e.g., Martin

et al., 2002). Mg-paleothermometry offers an advantageover other

proxies of sea-surface temperatures (SST) because Mg/Ca is mea-

sured on the samephase as d

O (a function of both the oxygeniso-

topic composition of sea water and temperature of calcification)

and therefore should record the same calcification temperature.

Consequently, paired measurements of

O and Mg/Ca can

permit an accurate reconstruction of d

water

and, by inference,

paleosalinity. Paired measurements on the same sample may also

resolve leads and lags between changes in SST and sea level

(Lea et al., 2000, 2002; Figure G7). Unfortunately, there remain

some uncertainties about Mg enrichment in species that have

undergone gametogenesis, suggesting a significant alteration

of Mg/Ca during gametogenic calcite precipitation. This effect

has the potential to confound Mg paleothermometry in sedi-

mentary shells containing a varying mixture of gametogenic

and ontogenetic calcite. However, recent studies indicate a strong

covariance between shell Mg/Ca and ambient temperatures and

suggest that, with careful calibration (of single species), it may

be a reliable proxy for seawater paleotemperature reconstruction.

Element/Ca ratio analyses of coral skeletal material have also

been used to reconstruct SSTs of the tropical oceans. Scleractinian

corals secrete skeletons composed of aragonite (CaCO

), which

incorporates both Sr and Ca into its structure. Similar to foramini-

fera, the incorporation ratio of Sr/Ca is determined by both the

Sr/Ca of ocean water and the Sr/Ca distribution coefficient

between aragonite and seawater. The former has remained

Figure G6 Ocean-wide Cibicidoides (a) Mg /Ca, and (b) Sr / Ca data from

Little Bahama Bank (), Sierra Leone Rise and Ceara Rise (), and Ontong

Java Plateau (○) (reprinted from Geochimica et Cosmochimica Acta, 61,

Rosenthal Y., Boyle E.A., and N. Slowey, Temperature control on the

incorporation of magnesium, strontium, fluorine, and cadmium into

benthic foraminiferal shells from Little Bahama Bank: Prospects for

permission from Elsevier).

Figure G5 Dissolved PO

3–

water-column profile from the sub-tropical

western North Atlantic compared with Cd /Ca concentrations in calcitic

foraminifera at similar depths (reprinted from Geochimica et

Cosmochimica Acta, 61, Rosenthal Y., Boyle E.A., and N. Slowey,

Temperature control on the incorporation of magnesium, strontium,

fluorine, and cadmium into benthic foraminiferal shells from Little

Bahama Bank: Prospects for thermocline paleoceanography,

GEOCHEMICAL PROXIES (NON-ISOTOPIC) 345

essentially constant over 10

yr timescales because of its long ele-

mental residence times, while the latter is determined principally

by the temperature of seawater in which the coral grew (Becket al.,

1992). Not surprisingly, measurements of skeletal Sr/Ca on

samples of coral typically show a very strong coherence with

SST over seasonal and interannual timescales (e.g., Linsley et al.,

2000; Figure G8). Such SST proxy variations have also been used

to reconstruct El Niño Southern Oscillation (ENSO) changes (e.g.,

Tudhope et al., 2001). By making paired measurements of d

derived tropical sea-surface salinity (SSS) reconstructions agree

with SSS observations at the interannual timescale (Quinn and

Sampson, 2002). There also appears to be a relationship between

Figure G7 Sea surface temperature (SST) (based on Mg / Ca), d

calcite

(V-PDB scale), and d

water

(V-SMOW scale) records from Cocos Ridge core

TR163-19, using G. ruber-white variety (Lea et al., 2000). The d

water

record is calculated from the SST and d

calcite

values. Dashed

lines are the measured and calculated data. Heavy lines indicate filtered data (reprinted from Quaternary Science Reviews, 21, Lea D.W., Martin P.A.,

Pak D.K., and H.J. Spero, Reconstructing a 350 ky history of sea level using planktonic Mg /Ca and oxygen isotope records from a Cocos Ridge

Figure G8 Comparison of instrumental monthly SST for Rarotonga and near-monthly Rarotonga coral Sr / Ca, spanning the interval from 1981

to 1997 (reprinted with permissions from Linsley B.K., G.M. Wellington, and D.P. Schrag, 2000, Decadal sea surface temperature variability in

the sub-tropical South Pacific from 1726 to 1997 A.D., Science, 290, pp. 1145–1148, Copyright (2000) AAAS. Permission from AAAS is

required for all other uses).

346 GEOCHEMICAL PROXIES (NON-ISOTOPIC)