Odekon M. Encyclopedia of paleoclimatology and ancient environments

Подождите немного. Документ загружается.

of events since that time, with some emphasis on the Younger

Dryas cold spell during the period of ice melting (Fairbanks,

1989). As noted by Olausson (1965) and many since, the

addition of large amounts of meltwater to the northern North

Atlantic interferes with the formation of deep water and thus

has many ramifications for the global thermohaline circulation

(Broecker and Denton, 1989).

The pioneer work in the reconstruction of the history of

the last million years was done on the cores obtained by the

Albatross during the Swedish Deep-Sea Expedition (1947–

1948, led by Hans Pettersson), by Gustaf Arrhenius, Cesare

Emiliani, Eric Olausson, Frances Parker, and Fred Phleger.

For the history of the last 100 million years, the pioneer studies

were done by a rapidly growing community of biostratigra-

phers and paleoceanographers on material recovered by the

drilling vessel Glomar Challenger, in the Deep Sea Drilling

Project (1968–1983). Among the contributors to the first of

the “Initial Reports or the Deep Sea Drilling Project” were

M. N. Bramlette, W. R. Riedel, W. A. Berggren, E. A. Pessagno,

J.D.Bukry,W.W.Hay,W.H.Blow,J.D.Hays,andA.G.Fischer.

Project Chief Scientist was M. N. A. Peterson. Just as the H.M.S.

Challenger symbolizes the beginning of modern oceanography ,

the Albatross and the Glomar Challenger areiconicforthebegin-

nings of paleoceanography. A third vessel, the JOIDES Resolution,

took over from 1985 and became the dominant tool for ocean

history exploration at that time.

The main result from the Albatross investigations was that

biogenic deep-sea deposits show cyclic sedimentation, a find-

ing that was linked to orbitally driven climate change and cor-

responding ice-age fluctuations. Much evidence has been found

since to support this notion, especially from the analysis of

oxygen isotopes of planktonic and benthic foraminifers, and

from time series analysis using Fourier methods (Hays et al.,

1976). The ice-age cycles are driven by changes in seasonal

contrast, much as initially suggested by the Serbian mathemati-

cian Milutin Milankovitch (1930). The dominant cycles reflect

the influence on seasonal contrast from changes in the obliquity

of Earth’s axis (41,000 years) and in the eccentricity of the

orbit (near 100,000 years) (Figure M6). Precession (the fact

that summer solstice migrates between positions close to and

far from the sun along the orbit) translates the eccentricity

signal into 23,000-year amplitude variations in climate, which

express themselves especially strongly in the productivity

cycles of upwelling regions. A striking change in response

of the climate system to the forcing occurs near 920,000 years

ago: a strong long-term cycle sets in as the amplitude of

the signal increases, first expressed as a doubling of the obli-

quity cycle, and then (near 700,000 year) moving into a

100,000-year cycle.

The main result from Glomar Challenger drilling concerned

the evolution of a cold ocean from a warm one, with all its

ramifications (Kennett, 1982). This general leitmotif governs

the evolution of marine organisms during the Cenozoic, from

nannoplankton to whales (which evolved blubber and baleen

to gather krill in cold upwelling regions).

Figure M5 Reconstruction of CCD variations for various regions of the

ocean. Dotted lines: According to Berger and Roth ( 1975). Solid and

dashed lines: according to van Andel et al. (1977), as drawn in Seibold

and Berger (1993). AFS: Auversian Facies Shift (Berger and Wefer, 1996).

The AFS denotes an overall change in sedimentation patterns, whereby

calcareous ooze becomes dominant in deep-sea deposits (carbonate

deposition being shifted from shelf to deep sea) and opaline sediments

are no longer widespread, but become concentrated below upwelling

regions, including the sea around Antarctica.

Figure M4 Deterioration of preservation of calcareous fossils (here:

planktonic foraminifers) as a function of water depth, and definition of

lysocline and CCD. The color of the calcareous ooze becomes darker

with depth, as the proportion of carbonate decreases.

MARINE BIOGENIC SEDIMENTS 529

The chief drawback of early drilling was the mixing of sedi-

ment during drilling. This was largely eliminated during the

Ocean Drilling Program (1985–2003), by sending a piston cor-

ing advice ahead of the drill bit. The cores allow for high-

resolution studies, on a thousand-year scale, and this opened

the door for “tuning” of the sedimentary record to Milankovitch

forcing, which can be reconstructed rather precisely (Shackleton

and Crowhurst, 1994). As a result, detailed records of climate

change emerged, including the details and differences of ice-

age cycles in various regions of the ocean. Other topics that

have been studied in detail are Antarctic ice buildup and the

response of the ocean in terms of changing productivity and

deep-sea circulation, the nature of the climate change at the

Eocene-Oligocene boundary (Prothero and Berggren, 1992),

and the sequence of a strange carbon isotope excursion at the

end of the Paleocene (thought to involve large-scale release

of methane from submarine clathrates; Dickens et al., 1995).

The series of events surrounding the grand extinction event

65 million years ago attracted much interest, as did many other

examples of natural experiments that allow an assessment of

how the climate machine works and how the ocean responds

to disturbance.

The great sustained stepwise cooling that characterizes glo-

bal climate history since the late Eocene (Auversian Facies

Shift, AFS, 40 million years ago) is well reflected in the oxy-

gen isotopes of benthic foraminifers from the deep-sea floor

of the Atlantic (Figure M7). The isotope index traces both ice

and temperature; thus, it illustrates both the cooling of deep

waters and the buildup of polar ice sheets, but at different

proportions depending on the time span considered. Ice on

Antarctica is thought to enter the picture around the time of

the AFS, followed by major changes in circulation and associ-

ated sedimentation patterns. However, great ice sheets have

presumably only covered the Antarctic continent since the

middle Miocene. At the Pacific equator, the first indications

for a narrowing of the zone of upwelling appear after the

AFS, in the shape of a band of opaline sediments (that is, sedi-

ments rich in radiolarians and diatoms). These sediments are

significant as a clue to the rise of zonal winds. An asymmetry

of zonal wind activity about the equator would have been

Figure M6 Oxygen isotope stratigraphy of Quaternary biogenic sediments from the western equatorial Pacific. Upper: Diagram showing how

ice-mass changes are translated into variations of oxygen isotope composition of planktonic foraminifers (concept introduced by C. Emiliani, 1955).

Lower: Stratigraphy of a sequence of cores from Ontong Java Plateau (modified after Berger and Wefer, 1992). Numbers: Emiliani isotope

stages. MPR: Mid-Pleistocene Revolution (climate shift to long cycles with large amplitudes). Top: Original isotope sequence. Below: 41-kyr cycle

contained in the series, and 100-kyr cycle, based on Fourier decomposition of the original record.

530 MARINE BIOGENIC SEDIMENTS

present, similar to that found today but more pronounced. The

southern trades helped to drive equatorial upwelling. A strong

asymmetry presumably persisted until well into the late Neo-

gene, when ice buildup in the Northern Hemisphere provided

a counterweight to southern polar ice fields (Flohn, 1985).

The low plankton diversity during the Oligocene is puz-

zling. The rapid increase of diversity in the early Miocene, with

the evolution of deep-living and shallow-living plankton, has

been linked to the development of a thermocline. Certainly,

the existence of a thermocline within the sunlit zone is a defin-

ing feature of the modern ocean. Photosynthesizing organisms

must stay in the sunlit zone sufficiently long, on average, to

produce more substance than they use up in respiration. When-

ever the mixed layer is substantially thicker than the sunlit

layer, this basic condition (Sverdrup’s criterion of net photo-

synthesis) is not likely to be met. A deep and/or ill-defined

thermocline will allow a thick mixed layer. This makes the

pelagic realm an unfavorable habitat for production. In addi-

tion, the lack of opportunity for migrating up and down through

the different environments defined by a thermocline precludes

diversification. We may conclude, therefore, that in the Oligo-

cene the thermocline was largely ineffective as an element

in structuring the pelagic environment; that is, the habitat of

plankton.

Strong coastal upwelling, based on silica and organic matter

content of hemipelagic sediments, is well documented for the

last 10 million years or so, especially in the Southern Hemi-

sphere (Summerhayes et al., 1992; Christensen and Giraudeau,

2002). Presumably, the resulting high abundance of plankton

and associated herring-like fish favored the rapid evolution of

warm-blooded predators, including sea birds, both with and

without flight, and marine mammals.

A direct comparison of warm oceans (pre-AFS) and cool or

cold oceans (post-AFS) is complicated by the fact that the

former are associated with large shelf seas and the latter are

not. Furthermore, the configuration of continents was quite

different, with a broad seaway once connecting the Pacific,

Indian and Atlantic Oceans in tropical latitudes – the Tethys

(Barron and Peterson, 1989; Barrera and Johnson, 1999). Any

efforts to relate differences or similarities in life habitat condi-

tions and biodiversity to the warm-cold contrast may well be

compromised by the presence or absence of large shelf seas

and the Tethys seaway (The Tethys persisted well into the Cen-

ozoic; Haq, 1981). There are some features of the warm ocean,

Figure M7 Stepwise cooling in the Cenozoic, since 40 million years ago (AFS, Auversian Facies Shift). The first step implies production of cold deep

water (seen in the extinction of benthic foraminifers); the second, buildup of Antarctic ice caps; the third, buildup of ice in the northern polar

regions (modified after a diagram in Miller et al., 1987).

MARINE BIOGENIC SEDIMENTS 531

however, that are quite different from those of a cold one on

a fundamental level. Foremost among these is a dearth of oxy-

gen in deep waters, which largely derives from the fact that

oxygen is less soluble in warm water than in cold. This, pre-

sumably, greatly contributed to the fact that anaerobic condi-

tions, reflected in the accumulation of black mud (now shale),

were widespread during Cretaceous periods, and especially so

during certain privileged intervals spanning the Albian-Aptian,

Cenomanian-Turonian and Santonian-Campanian boundaries.

The cause for extended anaerobic conditions, it is thought, is

a greatly increased effusion of volcanic magma to the surface

of the planet, bringing with it a corresponding increase of car-

bon dioxide in the atmosphere (Arthur et al., 1985). The system

responds by burying organic carbon, which counteracts the

geochemical imbalance. The burial of organic carbon preferen-

tially removes C-12 from the carbon cycle, driving up the ratio

between C-13 and C-12, which is then reflected in the stable

isotope composition of calcareous fossils deposited at the time

(Figure M8). The origin of much of the retrievable petroleum

on the planet is linked to anaerobic conditions in the middle

Cretaceous seas. The paleoceanographic implications of oxy-

gen stress are vast, including widespread denitrification, which

would have greatly impacted the level of productivity (For

background on productivity in the sea, and reading the record

of productivity in biogenic sediments, see Falkowski and

Woodhead, 1992).

The great concepts of orbitally driven climate cycles

during the Quaternary, of general cooling during the Ceno-

zoic, and of warm oxygen-stressed seas in the Cretaceous,

along with the revolutionary insights from plate tectonics,

have fundamentally changed the way geologists look at the

environmental history of the planet in the second half of the

twentieth century. However, perhaps the single most important

clue to the nature of Phanerozoic history, and especially

regarding evolution, has been the recognition of the impor-

tance of the end-of-Cretaceous (“K-T”) impact of a celestial

body with Earth, producing widespread extinction (for discus-

sions, see Berggren and van Couvering, 1984). It was within

biogenic deposits that the extinction event was first clearly

documented, in particular in calcareous shallow-water and

pelagic deposits (e.g., Berggren, 1962; Bramlette, 1965). The

evidence for impact was found in a famous pelagic carbonate

section near Gubbio in Italy (Alvarez et al., 1980), where the

exact horizon of extinction of planktonic foraminifers had

been previously established (Luterbacher and Premoli-Silva,

1964).

Studies of the K/T boundary have greatly contributed to the

reconstruction of the exact sequence of events, and especially

to elucidating the early response of the system to stress (invol-

ving strange plankton blooms: Thierstein and Okada, 1979)

and the course of recovery from the event in terms of biodiver-

sity (in a time-span counted in millions of years).

Wolfgang H. Berger and Gerold Wefer

Bibliography

Alvarez, L.W., Alvarez, W., Asaro, F., and Michel, H.V., 1980. Extraterres-

trial cause for the Cretaceous-Tertiary extinction. Science, 208,

1095–1108.

Arrhenius, G.O.S., 1952. Sediment cores from the east Pacific. Rep. Swed.

Deep-Sea Exped. 1947–1948, 5(1), 1–227.

Arthur, M.A., Dean, W.E., and Schlanger, S.O., 1985. Variations in the glo-

bal carbon cycle during the Cretaceous related to climate, volcanism,

and changes in atmospheric CO

. In Sundquist, E.T., and Broecker,

W.S. (eds.), The Carbon Cycle and Atmospheric CO

: Natural Varia-

tions Archean to Present. Washington, DC: American geophysical

union. AGU Geophysical Monograph, vol. 32, pp. 504–529.

Barrera, E., and Johnson, C.C. (eds.) 1999. Evolution of the Cretaceous

Ocean-Climate System. Boulder, CO: Geological Society of America.

GSA Special Paper 332, 445pp.

Barron, E.J., and Peterson, W.H., 1989. Model simulation of the Cretaceous

ocean circulation. Science, 244, 684–686.

Berger, W.H., and Roth, P.H., 1975. Oceanic micropaleontology: Progress

and prospects. Rev. Geophys. Space Phys., 13(3), 561–585 and 624–635.

Berger, W.H., and Wefer, G., 1992. Klimageschichteaus Tiefseesedimenten –

Neues vom Ontong-Java-Plateau (Westpazifik). Naturwissenschaften, 79,

541–550.

Berger, W.H., and Wefer, G., 1996. Expeditions into the past: Paleoceano-

graphic studies in the South Atlantic. In Wefer, G. et al., (eds), The

South Atlantic: Present and Past Circulation. Berlin: Springer,

pp. 363–410.

Berggren, W.A., 1962. Some planktonic foraminifera from the Maestrich-

tian and the Danian stages of southern Scandinavia. Stockholm Univ.

Contrib. Geol., 9,1–106.

Berggren, W.A., and van Couvering, J.A. (eds.) 1984. Catastrophes and

Earth History. Princeton, NJ: Princeton University Press, 464pp.

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

Cenozoic Geochronology. Geol. Soc. Amer. Bull., 96, 1407–1418.

Bolli, H.M., Saunders, J.B., and Perch-Nielsen, K. (eds.) 1985. Plankton

Stratigraphy. Cambridge, UK: Cambridge University Press, 1032pp.

Bramlette, M.N., 1965. Massive extinctions of biota at the end of Mesozoic

time. Science, 148, 1696–1699.

Broecker, W.S., and Denton, G.H., 1989. The role of ocean-atmosphere

reorganizations in glacial cycles. Geochim. Cosmochim. Acta, 53,

2465–2501.

Bukry, D., and Bramlette, M.N., 1969. Coccolith age determinations – Leg

1, Deep Sea Drilling Project. Initial Rep. Deep Sea Drilling Project, 1,

369–387.

Figure M8 General trends of carbon isotope stratigraphy (pelagic marine limestones) in the Cretaceous, suggesting the removal of organic carbon

during times of low oxygen availability (that is, during deposition of black mud) (modified from Arthur et al. 1985). OAE: Oceanic anoxic event.

532 MARINE BIOGENIC SEDIMENTS

Butcher, S.S., Charlson, R.J., Orians, G.H., and Wolfe, G.V. (eds.) 1992.

Global Biogeochemical Cycles. London/ San Diego,CA: Academic

Press, 379pp.

Christensen, B.A., and Giraudeau J. (eds.) 2002. Neogene and quaternary

evolution of the Benguela coastal upwelling system. Mar. Geol., 180

(Special Issue), 1–276.

CLIMAP Project Members, 1976. The surface of the ice-age Earth.

Science, 191, 1131–1137.

Daly, R.A., 1934. The Changing World of the Ice Age. New Haven: Yale

University Press, 271pp.

Dickens, G.R., O’Neil, R.R., Rea, D.K., and Owen, R.M., 1995.

Dissociation of oceanic methane hydrate as a cause of the carbon

isotope excursion at the end of the Paleocene. Paleoceanography, 10,

965–971.

Emiliani, C., 1955. Pleistocene temperatures. J. Geol., 63, 538–578.

Fairbanks, R.G., 1989. A 17,000-year glacio-eustatic sea level record:

Influence of glacial melting rates on the Younger Dryas event and

deep-ocean circulation. Nature, 342, 637–642.

Falkowski, P.G., and Woodhead, A.D. (eds.) 1992. Primary Productivity

and Biogeochemical Cycles in the Sea. New York: Plenum Press,

550pp.

Fischer, G., and Wefer, G. (eds.) 1999. Use of Proxies in Paleoceanogra-

phy. Berlin: Springer, 735pp.

Flohn, H., 1985. Das Problem der Klimaänderungen in Vergangenheit und

Zukunft. Darmstadt: Wissenschaftliche Buchgesellschaft, 228pp.

Funnell, B.M., and Riedel, W.R. (eds.) 1971. The Micropalaeontology of

Oceans. Cambridge, UK: Cambridge University Press, 828pp.

Haq, B.U., 1981. Paleogene paleoceanography: Early Cenozoic oceans

revisited. Oceanologica Acta, 4 (Supplement), 71–82.

Hay, W.W., 1988. Paleoceanography: A review for the GSA centennial.

Geol. Soc. Am. Bull., 100, 1934–1956.

Hays, J.D., Imbrie, J., and Shackleton, N.J., 1976. Variations in the Earth’s

orbit: Pacemaker of the ice ages. Science, 194, 1121–1132.

Kennett, J.P., 1982. Marine Geology. Englewood Cliffs, NJ: Prentice-Hall,

813pp.

Kennett, J.P. (ed.) 1985. The Miocene Ocean: Paleoceanography and bio-

geography. Geol. Soc. Am. Memoir, 163

,1–337.

Luterbacher, H.P., and Premoli-Silva, I., 1964. Biostratigrafia del limite

Cretaceo-Terziario. Riv. Ital. Paleontol., 70,67–128.

Milankovitch, M., 1930. Mathematische Klimalehre und astronomische

Theorie der Klimaschwankungen. Handbuch der Klimatologie, Bd 1,

Teil A, Berlin, Bornträger, 176pp.

Miller, K.G., Fairbanks, R.G., and Mountain, G.S., 1987. Tertiary oxygen

isotope synthesis, sealevel history, and continental margin erosion.

Paleoceanography, 2,1–19.

Olausson, E., 1965. Evidence of climatic changes in North Atlantic deep-

sea cores, with remarks on isotopic paleotemperature analysis. Prog.

Oceanogr., 3, 221–252.

Parker, F.L., 1962. Planktonic foraminiferal species in Pacific sediments.

Micropaleontology, 8(2), 219–254.

Prothero, D.R., and Berggren, W.A. (eds.) 1992. Eocene-Oligocene Climatic

and Biotic Evolution. Princeton, NJ: Princeton University Press, 566pp.

Purdy, E.G., 1974. Reef configurations: Cause and effect. In Laporte, L.F.

(ed.), Reefs in Time and Space. Tulsa, OK: Society of Economic

Paleontologists and Mineralogists. SEPM Special Publication 18,

pp. 9–76.

Reading, H.G. (ed.) 1986. Sedimentary Environments and Facies, 2nd edn.

Oxford: Blackwell Scientific, 615pp.

Seibold, E., and Berger, W.H., 1993. The Sea Floor. 3rd edn. Berlin:

Springer, 356pp.

Shackleton, N., and Crowhurst, S., 1994. Details that make the difference.

Oceanus, 36(4), 45–48.

Summerhayes, C.P., Prell, W.L., and Emeis, K.C. (eds.) 1992. Upwelling

Systems: Evolution since the Early Miocene. Piccadilly: Geological

Society of London. Geological Society of London Special Publication

64, 519pp.

Thierstein, H.R., and Okada, H., 1979. The Cretaceous/Tertiary boundary

event in the North Atlantic. Initial Rep. Deep Sea Drill. Proj., 43, 601–616.

van Andel, Tj.H., Thiede, J., Sclater, J.G., and Hay, W.W., 1977. Deposi-

tional history of the South Atlantic Ocean during the last 125 million

years. J. Geol., 85, 651–698.

Vincent, E., and Berger, W.H., 1981. Planktonic foraminifera and their use

in paleoceanography. In Emiliani, C. (ed.),“The Sea, vol. 7, the Oceanic

Lithosphere.” New York: Wiley-Interscience, pp. 1025–1119.

Wefer, G., and Berger, W.H., 1991. Isotope paleontology: Growth and com-

position of extant calcareous species. Mar. Geol.,

100, 207–248.

Winter, A., and Siesser, W.G., 1994. Coccolithophores. Cambridge, UK:

Cambridge University Press, 242pp.

Cross-references

Astronomical Theory of Climate Change

Carbonate Compensation Depth

CLIMAP

Coccoliths

Coral and Coral Reefs

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

Deep Sea Drilling Project (DSDP)

Diatoms

Foraminifera

Ocean Anoxic Events

Ocean Drilling Program (ODP)

Paleocene-Eocene Thermal Maximum

Quaternary Climate Transitions and Cycles

Radiolaria

MARINE CARBON GEOCHEMISTRY

Introduction

The ocean contains approximately 50 times as much dissolved

as is contained in the atmosphere, and the carbon cycle in

the ocean controls the pCO

of the atmosphere on time scales

of thousands of years. The atmospheric CO

concentration

today is rising because of human activities, such as fossil fuel

combustion and deforestation. The CO

rise is faster than the

century-scale ocean/atmosphere equilibration time, but the

ocean will ultimately absorb most of the CO

that is released.

In the past, changes in the carbon cycle of the ocean drove

changes in atmospheric pCO

synchronously with the glacial/

interglacial climate cycles. The orbital forcing which paces

the glaciations (called Milankovitch cycles, see Astronomical

Theory of Climate Change) is asynchronous across the equator;

extra warmth in the north is balanced by decreased heating in

the south. However, because of the rapid mixing of the atmo-

sphere, the glacial CO

cycles tend to warm or cool the entire

planet synchronously, and are largely responsible for the global

synchrony of the glaciations. The carbon cycle of the ocean is

therefore a primary determiner of the climate of the Earth, in

the past and in the future.

The carbon cycle in the ocean begins with gas fluxes across

the sea surface into the atmosphere. These depend on wind

speed, temperature, and the chemistry of the water at the sea

surface. Seawater chemistry of CO

is complicated by the pre-

sence of the carbonate pH buffer system pH. Only a tiny frac-

tion of the dissolved carbon is present as the uncharged

dissolved gas that can interact with the atmosphere. Because

of this chemistry, a slight change in the pH of seawater can

result in an enormous change in the concentration of dissolved

, and in the equilibrium pressure of CO

in the atmosphere.

The solubility of dissolved CO

depends on temperature and to

a lesser extent salinity, so that a warming parcel of seawater

will tend to push more CO

into the atmosphere. The CO

con-

centration of surface ocean waters is also affected by the

growth and sinking of phytoplankton in the surface ocean.

The carbon cycle of the ocean is therefore determined by a

MARINE CARBON GEOCHEMISTRY 533

combination of chemical, physical, and biological mechanisms

(see also Carbon cycle).

The physical chemistry of carbon in seawater

Carbon is a unique element in the biosphere, in part because of

its extremely varied chemistry. The oxidized form, CO

, is the

dominant form in the atmosphere. In seawater, oxidized carbon

undergoes hydration and pH reactions leading to generation of

bicarbonate (HCO



) and carbonate (CO

2

) ions. These species

are collectively called “dissolved inorganic carbon” or DIC.

The most reduced form of carbon is methane, CH

, while most

biological carbon exists at intermediate oxidation states analo-

gous to carbohydrates, which follow the approximate stoichio-

metry CH

O. Reduced forms of carbon are collectively called

“organic carbon.” Production of organic carbon from inorganic

requires an input of energy, such as the light energy used in

photosynthesis, or chemical energy used in chemosynthesis.

In the air, the CO

abundance can be defined as a concentra-

tion (molecules per volume), a proportion or “mixing ratio”

(molecules of CO

per molecule total), or as the proportion of

the total pressure that can be attributed to CO

. This last quan-

tity is called the “partial pressure,” and at sufficiently low total

pressures, the ideal gas law tells us that the partial pressure of a

gas like CO

is very close to the total pressure of the mixture

multiplied by the fraction of the molecules that are CO

. The

atmospheric CO

abundance is commonly expressed either as

a mixing ratio in units of parts per million (ppm), or as a partial

pressure in units of microatmospheres (10

6

atm = matm). The

numerical value of either of these measures in the atmosphere

is close to 370 in both cases (and is presently rising by approxi-

mately 1.5 ppm or matm per year). The abbreviation “pCO

”

stands for partial pressure (and is unfortunately easy to confuse

with the meaning of pH, defined as the negative logarithm

of the concentration of H

). For a detailed discussion of the

physical chemistry of inorganic carbon in seawater, see Stumm

and Morgan (1981).

A convenient way to think about CO

equilibrium between

air and water is to take the chemistry of a water sample,

and calculate how much CO

would exist in the air above it

in equilibrium. In this way, we can define the pCO

of a water

sample. This tells us nothing about the real physical pressure

that a diver, for example, would feel in the water, but tells us

what the composition of a bubble would be after it has had a

chance to equilibrate. In nature, water and gas phases are not

always in equilibrium; in fact, for CO

at the surface of the

ocean, disequilibrium is the rule. If the pCO

of a water sample

is higher than the pCO

of the air it is in contact with, then CO

will tend to evaporate from the water into the air, lowering

the pCO

of the water and raising the pCO

of the air as the

system approaches equilibrium (see Carbon dioxide, dissolved

(ocean)).

The solubility of any gas in water depends on tem-

perature, with increasing solubility as the water cools down.

gas is no exception. At a given concentration of dissolved

, the equilibrium partial pressure in the gas phase there-

fore decreases as the temperature drops (increasing solubility

means less remains in the gas phase). The surface ocean is

warm in the tropics and cold at the poles. Therefore, one com-

ponent of the carbon cycle in the ocean is a tendency for

to escape to the atmosphere in the warm tropics, where

it is less soluble, and to invade the ocean in the colder high

latitudes.

Partioning of carbonate species in seawater

gas proper is but a minor participant in a larger chemical

system called the carbonate buffer system. CO

gas combines

with water (is “hydrated”) to form carbonic acid (H

). Car-

bonic acid undergoes two proton dissociation reactions (loss of

) to form the charged (ionic) species bicarbonate (HCO



)

and carbonate ion (CO

2

). The hydration reaction is

þ H

O Ð H

ð1Þ

In practice, it is analytically difficult to distinguish hydrated

from unhydrated CO

, so these two species are lumped

together into a single category that for simplicity we will just

call dissolved CO

in the discussion that follows. The equili-

bration time of this reaction is roughly a minute, although it

can be accelerated in biological systems by the enzyme carbo-

nic acid anhydrase.

The dissociation reactions are governed by equilibrium expres-

sions, denoted K

for the loss of the first proton and K

for the sec-

ond. The dissociation constant K

can be written to determine the

ratio of the products to the reactants of the reaction

Ð HCO



þ H

ð2Þ

while the second, K

, applies to

HCO



Ð CO

2

þ H

ð3Þ

The ratios are taken between the activities of each of the

solutes, which can be approximated by the concentrations in

moles per liter, denoted by square brackets:

¼½HCO



½H

=½CO



and

¼½CO

2

½H

=½HCO





Values of the dissociation constants are given by Dickson

and Millero (1987).

Built upon these seemingly simple relations is a chemical

system of rather counter-intuitive behavior. One example is

the response of [CO

2

] to changes in the CO

concentration

of the solution (Equation 1). Carbonate ions determine whether

CaCO

from coral reefs or planktonic shells dissolves (see

Carbonate compensation depth). The response of an equili-

brium reaction to an addition of a species on one side of the

equilibrium is a tendency to shift the equilibrium toward the

opposite direction, to “mop up” the added species. The res-

ponse of the first dissociation (Equation 2) in isolation would

be an increase in HCO



and H

(an increase in H

is equiva-

lent to a decrease in pH). The response of the second dissocia-

tion reaction might seem to be another shift to the right,

increasing the concentration of CO

2

(and increasing the stabi-

lity of CaCO

). We could call this the “rising tide floats all

boats” response.

However, this simple expectation is wrong. The flaw in our

reasoning was to ignore the protons, which exist in much lower

concentration than any of the carbon species. The rightward

shift in Reaction 2 above generates a much more profound

change in H

, relative to its concentration, than the change in

HCO



therefore dominates the response of Reaction 3,

which shifts to the left rather than to the right as we would

assume. The behavior of the system can be seen more intuitively

by combining both reactions, eliminating the troublesome

534 MARINE CARBON GEOCHEMISTRY

protons (and thereby conceptually forbidding them from

changing their concentrations at all – a better approximation).

This new reaction is:

þ CO

2

Ð 2HCO



ð4Þ

When the reaction is written in this way, an addition of car-

bonic acid (or, equivalently, CO

) can be seen to consume

2

, provoking the dissolution of coral reefs, rather than their

growth. In spite of the fact that CaCO

contains carbon, it can-

not act as a sink for CO

unless a source of base is added.

The buffer chemistry of seawater has an immediate implica-

tion for the interaction of seawater with rising atmospheric

pCO

, even in the absence of solid CaCO

. For a simple unbuf-

fered gas like oxygen, for example, the solubility of the gas is

expressed by Henry’s law

h;O

¼ pO

=½O



which tells us that the concentration of the gas in solution goes

up proportionately to a rising partial pressure in the gas phase.

The relation also holds for CO

,as

h;CO

¼ pCO

=½CO



but the total amount of CO

that the seawater can absorb is

affected by the buffer chemistry as expressed in Reaction 4

above. Dissolved carbonate ion (CO

2

) reacts with added

, effectively hiding it from the atmosphere and from

Henry’s law. The buffer works by reacting most of the invading

with CO

2

to form HCO



, so the strength of the buffer is

determined by the ratio of the CO

and CO

2

concentrations.

Globally, the average ratio of CO

2

to CO

in surface waters

is about 10, so seawater can hold 10 times as much CO

as it

would without the buffer system. As atmospheric CO

rises,

however, sea surface CO

2

concentrations will drop, and the

buffer capacity of seawater will decline.

When the gas concentration in the surface ocean is out of

equilibrium with the atmosphere, it relaxes toward equilibrium

by a process known as gas exchange. The rate of gas exchange

depends on the wind speed, surface roughness, bubbles and

foam, the presence or absence of a surface microlayer, or skin

of oils, and the diffusivity of the gas. A typical gas exchange

rate is 3–5md

1

, meaning that a typical mixed layer of 100 m

thick at the ocean surface will approach equilibrium with

an e-folding time scale of 100 m/5 m d

1

= 20 days (see

Wanninkhof, 1992). For CO

, the factor of 10 buffering would

increase the time scale to about 200 days. For the ocean as a

whole, the rate at which gases invade the deep ocean is limited

by the circulation time scale of the ocean, which is about a

thousand years. For CO

, because the ocean is the larger reser-

voir, the atmosphere largely responds to ocean forcing, rather

than the other way around, on the time scale of ocean overturn-

ing circulation.

The balancing of the various dissolved carbon species across

the equilibrium reaction makes it somewhat more complicated

to think about the carbon cycle in the ocean. For example,

when seawater with one set of concentrations is mixed with a

different seawater sample with a different set of concentrations,

the mixed carbonate buffer system adjusts to re-establish pH

balance. That is to say, if we combine one solution with

10 mMCO

with equal parts of another that starts with 20 mM

, we do not necessarily obtain a solution that contains

15 mMCO

, because some of the CO

reacts with CO

2

the resulting mixture. This makes life complicated for oceano-

graphers, who would prefer that their chemical quantities be

“conservative to mixing.” The solution is a pair of derived

quantities for seawater that do have the required conservation

properties. One is the total concentration of all of the CO

spe-

cies; this is called the “total CO

,” or SCO

. If we mix a solu-

tion that contains 1,800 mM SCO

with equal parts of another

that contains 2,200 mM SCO

, we will obtain a solution that

contains 2,000 mM SCO

, with no tricky business.

The other derived quantity is called the alkalinity, which is

defined as the sum of all charges of the salts of weak acids in

solution. One way to think about alkalinity is that if we were

to add H

in an amount equal to the alkalinity, then all of

the carbon in solution would be in the form of CO

, free to

exchange with the atmosphere. Ignoring contributions from

borate and other minor species, the “carbonate alkalinity” of

seawater can be written as

alk ¼ HCO



þ 2CO

2

ð5Þ

where the factor of two in front of CO

2

reflects the two pro-

tons it would take to produce carbonic acid from carbonate.

The crucial point is that Reaction 4, above, has the same alka-

linity on both sides of the reaction. As pH equilibrium redistri-

butes carbon species toward one side or the other of Reaction

3, the alkalinity is unchanged. Therefore, if we combine two

water parcels with different alkalinities, the alkalinity of the

resulting mixture is a simple combination of the two. We say

that both alkalinity and total CO

are conservative to mixing,

and this simplifies our thinking about all kinds of processes

that occur in the ocean carbon cycle.

How can we relate the alkalinity and total CO

to the equi-

librium pCO

of surface waters, and exchange of CO

with the

atmosphere? As we have already seen, dissolved CO

gas itself

is only a minor part of the total carbon inventory of seawater,

and the fraction of carbon in the form of CO

depends on the

pH of the solution. Together, the alkalinity and the total CO

concentration can tell us everything there is to know about the

carbon chemistry of the seawater. At a constant alkalinity, if

we increase the total CO

concentration, the concentration of dis-

solved CO

and the equilibrium pCO

of the water sample will

go up. This is easy to imagine, as an increase in the number of

carbons without any change in the number of charges must lead

to an increase in the number of uncharged carbons, that is CO

In a similar way, if we maintain a constant total CO

and

increase the alkalinity, then the number of charges per carbon

atom goes up, and the proportion of uncharged CO

molecules

goes down.

A useful property of the alkalinity is that it does not care

about fluxes of CO

gas. An invasion of CO

from the atmo-

sphere will increase the total CO

of the seawater, and alter

the proportions of the various carbon species according to

Reaction 3, but since shifts along Reaction 3 have no effect

on the alkalinity, the alkalinity will be unchanged in response

to gas exchange. In an analogous way, if we were to add strong

acid to seawater somehow, this would change alkalinity while

leaving the total CO

unchanged. This latter is a hypothetical

example; natural fluxes of strong acid such as acid rain are gen-

erally negligible in the carbon chemistry of the surface ocean.

The biological pump

The other major mechanism governing the fluxes of CO

across the sea surface, nearly equal in magnitude to the effect of

MARINE CARBON GEOCHEMISTRY 535

temperature discussed above, is the growth and sinking of phyto-

plankton in the surface ocean. Geochemically, the effect of

phytoplankton growth is to convert dissolved carbon into particu-

late carbon. Particles can sink, generating a net flux of carbon

downwards in the ocean. Some of the phytoplankton (plants of

the sea) sink on their own, while another fraction of the sinking

carbon arises from the fecal pellets excreted by zooplankton gra-

zers. This process has been called the “biological pump” (see

Broecker and Peng, 1982). The atmosphere only sees the surface

ocean, so the effect of the biological pump is to change the ocean

chemistry so that the atmosphere sees, and therefore changes, the

chemistry of the atmosphere itself.

It is helpful to divide the biological pump into two components,

because these components have opposite effects on the CO

concentration and pCO

of the surface ocean. The first, and domi-

nant, component is the organic carbon pump, comprised of the

proteins, lipids, and other structural biochemical machinery of

the cell. This is called the “soft tissue” pump. Only a small fraction

of the plankton production is exported to depth; in very productive

regions, the export fraction may reach half, while in the open

ocean only 5–10% of the production is exported. The rest of

the biomass produced is respired in the surface ocean, returning

to dissolved carbon. Organic carbon production depletes the total

concentration of the seawater , but has only a small effect on

the alkalinity. Therefore, the effect of the soft tissue pump is to

decrease the pCO

of surface waters.

The second component of the biological pump is the pro-

duction and sinking of the calcium carbonate (CaCO

) shells

of some plankton species. This process is called the “hard

tissue” pump. The formation of CaCO

takes place by the

reaction

2þ

þ CO

2

Ð CaCO

ð6Þ

Because one mole of CO

2

accounts for one mole of total

but two equivalents of alkalinity (see Equation 5), the

effect of the CaCO

pump is a 2:1 change in alkalinity versus

total CO

. A decrease in alkalinity leaves a greater proportion

of uncharged carbon (CO

), so the net effect of the CaCO

pump on the pCO

of surface waters is an increase in pCO

When these two components of the biological pump are

combined in the surface ocean, the rate of the soft tissue pump

exceeds that of the hard tissues, so the overall effect of biologi-

cal activity in the ocean is to decrease the p CO

of the atmo-

sphere. CaCO

secreting plankton are favored in tropical

surface waters, but even here the rate of organic carbon sinking

from the surface ocean exceeds that of CaCO

by a factor of up

to 10. In polar waters colder than 10



C, CaCO

production

becomes even less important. If the biota of the ocean were

killed and the contents of the ocean allowed to redistribute

themselves in response to this, the pCO

of the atmosphere

would perhaps double.

Photosynthesis in the ocean is limited by a number of fac-

tors that determine the extent to which biology affects ocean

chemistry and atmospheric pCO

(Longhurst, 1998). The first

factor is light; production of biomass requires energy, which

is derived by sunlight via the photosynthetic apparatus.

Although open ocean waters are clearer than any freshwater

lakes on Earth, the incident sunlight is absorbed and scattered

by phytoplankton, other particulate and dissolved material,

and by the water itself. Although the absorption of light

depends on wavelength (red is absorbed most quickly, which

is why undersea scapes appear blue), in general the absorption

of intensity of “photosynthetically available radiation” (PAR)

decreases over a depth scale of about 100 m. Below this depth,

the metabolic costs of respiration exceed the available energy

from photosynthesis, and phytoplankton are unable to thrive.

The surface sunlit zone, capable of supporting phytoplankton,

is called the “euphotic zone.” The production of CaCO

does

not require energy in the same way as photosynthesis does,

and there is some secretion of CaCO

by zooplankton such as

foraminifera below the euphotic zone, but most CaCO

produc-

tion is associated with phytoplankton called coccolithophorids

in the euphotic zone.

A second factor influencing the availability of light to the

phytoplankton is mixing. Winds tend to mix the surface waters

of the ocean into a distinct mixed layer ranging from a few to a

thousand meters thick. The thickness of the mixed layer

depends on the contrast between the temperature (and to a les-

ser extent the salinity) of the surface ocean compared with the

deep ocean. In the tropics, warm surface waters are more buoy-

ant than the colder waters below. This density difference inhi-

bits mixing of waters across the stratification boundary, and

the mixed layer tends to be thin. Phytoplankton suspended in

the surface euphotic zone tends to remain in the euphotic zone.

However, at high latitudes, or in subpolar winter, the cooler

surface waters mix more readily with subsurface waters, allow-

ing the wind to create a deep mixed layer. The phytoplankton

spend at least some of their time at depths deeper than the

euphotic zone, limiting their growth. The “spring bloom” in

the North Atlantic is a clear demonstration of the effect of mix-

ing. As soon as spring warming stabilizes a shallow mixed

layer, the phytoplankton concentrations and photosynthesis

rates increase dramatically.

The end of the spring bloom in the North Atlantic (and in

other regions) is caused by another factor that limits phyto-

plankton growth: nutrient ions such as NO



and PO

3

, which

are essential to the manufacture of the biochemical machinery

of the cell. Nitrogen is used in the amino acid building blocks

of proteins, and phosphorus is used in DNA and in an energy

storage molecule called ATP. The elemental ratio of C:N:P in

phytoplankton is relatively constant, and has been named the

“Redfield ratio.” An astonishing fact about ocean chemistry is

that the ratio of N:P in seawater is close to that of phytoplank-

ton, so that to a first approximation, both elements get used up

at the same time. Carbon, in contrast, exists in much greater

abundance in seawater, so that the nutrients are always depleted

first. The mean concentration of PO

3

in the ocean is 2.2 mM;

when this value is multiplied by a revised C:P Redfield ratio

of 135 we estimate that approximately 200 mM of the roughly

2,000 mM SCO

concentration in the ocean (N10%) travels

with PO

3

as “metabolic” CO

. The metabolic CO

represents

the maximum impact of biology on the chemistry of the

surface ocean.

The effect of the soft-tissue biological pump in the ocean is

to deplete NO



and PO

3

in surface waters by exporting them

to the deep ocean. Primary production in the ocean is therefore

tied intimately to upwelling or mixing of nutrient-rich subsur-

face waters into the surface euphotic zone. One place where

this occurs is in near-shore waters, where winds drive offshore

currents, pulling subsurface waters up from below. Another

such place is along the equator, where the change of direction

of the Coriolis force pulls surface waters away from the equa-

tor, leaving a void that must be filled by upwelling subsurface

waters. A third mechanism for bringing nutrients to the surface

takes us back to the spring bloom example above; deep winter-

time mixing entrains nutrient-rich deep water into the euphotic

536 MARINE CARBON GEOCHEMISTRY

zone, recharging it with nutrients for the next spring bloom.

After the nutrients are used up by the spring bloom, phyto-

plankton concentrations and photosynthesis rates decrease

for the rest of the summer.

In contrast to the North Atlantic, some regions of the surface

oceans, such as the Equatorial Pacific, the North Atlantic, and

the Southern Ocean, maintain high concentrations of nutrients

in the surface euphotic zone all year round, with no spring

bloom. One of the more exciting discoveries of the past decade

is that the phytoplankton may be limited by the availability

of another nutrient, iron (Fe) (see Iron and climate change).

Phytoplankton require iron in the minute ratio of approximately

1:10

Fe:C. The oxic form of iron, Fe

, is relatively insoluble

in water, and so seawater tends to become depleted in iron rela-

tive to their stocks of NO



and PO

3

. When this water upwells,

it must await the delivery of iron by deposition as dust from the

atmosphere. The North Atlantic, according to this theory, is

rich in dust deposition, so supports a spring bloom, while the

equatorial Pacific is more remote, and is dust and iron limited.

There is a general correspondence between the temperature

of the surface ocean and its nutrient content, although this rela-

tionship is not absolute. In part, this is because colder surface

waters mix more readily with nutrient-rich subsurface. Deeper

mixing also creates a situation of light limitation for phyto-

plankton, as explained above. At any event, the net effect of

this correspondence is to set temperature and sea surface nutri-

ents in opposition to each other in their effects on atmospheric

pCO

. The cold temperatures in the Southern Ocean, for exam-

ple, tend to lower the pCO

of surface water, pulling CO

into

the ocean. However, the high concentration of nutrients in

Southern Ocean surface waters indicate that the SCO

concen-

tration is higher there than it would be if nutrients were

depleted, tending to increase atmospheric pCO

in the opposite

direction from the thermal forcing. On the other hand, the region

of the highest CO

flux from the ocean to the atmosphere is the

equatorial Pacific, where nutrient-rich waters are warmed by

the tropical sun, in other words where both temperature and

nutrient forcing are working in the same direction.

Most of the organic matter that sinks out of the euphotic

zone in the surface ocean never reaches the sea floor, typically

4 km below. It is subject to microbial degradation and to graz-

ing by zooplankton in the mid-water column. The sinking flux

of particles can be measured using a device called a sediment

trap, which consists of a large funnel or cylinder with a carou-

sel of sealable sample chambers that can sequentially store the

falling debris (Honjo, 1996). Several traps are typically tied to

a mooring, allowing an assessment of the sinking flux of car-

bon as a function of depth. These fluxes are generally found

to decrease drastically over the top 500 m of the water column.

Different chemical components of the debris are removed

selectively as they sink. The flux of organic carbon decreases

more quickly with depth than does the flux of CaCO

or SiO

(another type of shelly hard part, the mineral opal is secreted

by the phytoplankton diatoms) (see Diatoms). Nitrogen and

phosphorus are removed more quickly than carbon, as indicated

by increasing C:N and C:P ratios in sediment traps. The flux of

CaCO

appears rather uniform with depth in sediment traps, but

budgets for CaCO

production and deposition on the sea floor,

and for alkalinity in the water column, seem to indicate a sub-

stantial dissolution of CaCO

in the mid-water column. The

discrepancy with sediment trap results may be caused by biases

in trap efficiency in shallow waters (Yu et al., 2001)orby

CaCO

dissolution (Betzer et al., 1984) within the traps.

Organic carbon by itself is not sufficiently dense to sink

very quickly through seawater, if at all, and so attention has

focused on the role of CaCO

, SiO

, and wind-blown continen-

tal material as “ballast” (Armstrong et al., 2002; Klaas and

Archer, 2002). This theory is consistent with the sediment trap

data in that while the fluxes of the various constituents vary

widely in space and time, the ratio of organic carbon to ballast

in the deep traps, and in particular to CaCO

, seems much more

stable. Only a small fraction of the organic matter produced in

the surface ocean makes it to the deep sea, so the ballast model

predicts that the absolute flux of organic carbon to depth might

be determined by the fluxes of ballast. However, the ballast

model does not explain the trap data in the top kilometer of

the water column, where the organic carbon to ballast ratio is

higher and more variable than in the deep sea.

CaCO

compensation and the pH of the ocean

We have seen how the biological pump affects the pCO

of the

atmosphere, by pumping carbon and nutrients from the surface

ocean to the deep. In addition to this direct effect, the biological

pump has an indirect effect on pCO

, called carbonate compen-

sation (Broecker and Peng, 1987). CaCO

is produced biologi-

cally, mostly in the surface ocean. Some fraction of this CaCO

reaches the sea floor, and its chemical constituents (Ca

, one

SCO

, and two alkalinities) are permanently removed from

the ocean as it is buried. This output is balanced by input from

the dissolution of rocks, mostly sedimentary CaCO

, a process

called weathering on land. The idea is that the pH of the ocean

adjusts; acting as a stable negative feedback to insure that the

output of CaCO

by burial balances the input from weathering.

As we have seen, a change in seawater pH will have a large

impact on the speciation of carbon, and therefore on pCO

(see Paleo-ocean pH ).

The input of dissolved CaCO

to the ocean derives from the

chemical weathering of rocks on land (see Weathering and

climate). The saturation state for CaCO

dissolution is given

by the reaction

CaCO

! Ca

2þ

þ CO

2

; K

¼½Ca

2þ

½CO

2

ð7Þ

If the product of the reactants, Ca

and CO

2

, is less than

the value of K

then the solution is undersaturated and CaCO

will tend to dissolve. In practice, the CO

2

combines with CO

and water to yield the overall reaction

CaCO

þ CO

þ H

O ! Ca

2þ

þ 2HCO



ð8Þ

where the Ca

and HCO



reflect the dominant ion pair in most

river water. The reactant CO

is derived ultimately from the

atmosphere, but CO

partial pressures in soil gas are elevated

to an order of magnitude higher than atmospheric values by

the action of plants and the microbial respiration of soil carbon.

River water, reflecting its origin as soil water, is typically

strongly supersaturated in CO

, and close to saturation with

respect to CaCO

at that pCO

. In areas where the rate of fresh-

water runoff is extremely high, the dissolution reaction may not

have time to keep up with the flux of distilled rainwater, and

the solution concentrations will drop. The lithology, or rock

type, of a region also has an impact on the chemistry of rivers.

To a first approximation, however, we expect that the total flux

of dissolved CaCO

to the ocean depends primarily on the

total rate of fresh water runoff to the oceans, and secondarily

on temperature, soil biological activity ( pCO

), terrain type,

MARINE CARBON GEOCHEMISTRY 537

and other variables. Once the Ca

and alkalinity are delivered

to the ocean, they are used to produce CaCO

according to the

reaction

2þ

þ 2HCO



! CaCO

þ H

O þ CO

ð9Þ

returning the atmospheric CO

that weathering borrowed back

to the atmosphere.

CaCO

production occurs in shallow ocean waters and in

surface waters overlying the deep sea (Milliman, 1993). In

many cases, in coral reefs for example, the production rate of

CaCO

seems to mirror the rate of organic carbon production,

as though the organisms calcified in order to shift the pH bal-

ance back toward the acidic, ensuring a readily available supply

of dissolved CO

. Some fraction, roughly half, of the CaCO

produced in surface waters dissolves in the water column,

and the rest reaches the sea floor to be redissolved or buried.

The fate of CaCO

landing on the sea floor depends strongly

on depth in the ocean, because the solubility of CaCO

increases with pressure (Archer, 1996). This is because the dis-

solution reaction entails a small decrease in volume as CaCO

dissolves; an increase in pressure encourages the system to find

the volumetrically smaller, dissolved, form. In addition, there is

a surface/deep contrast in the concentration of CO

2

, resulting

from the biological pump. We have already seen that the effect

of the biological pump, dominated by the soft-tissue pump, on

the carbon chemistry of the water column is to deplete the sur-

face waters in SCO

. As a result of the pH equilibrium reaction

3, dissolved CO

2

is increased over the mean ocean value. The

concentration of CO

2

goes down with depth, but the amount

of CO

2

required in equilibrium goes up with depth. These

factors combine to create a “saturation horizon ” in the water

column, above which CaCO

is supersaturated, and below

which it is undersaturated. Some fraction of the CaCO

raining

to the sea floor lands on topographic highs, in sediments over-

lain by supersaturated water, and tends to be buried. CaCO

landing in abyssal sediments tends to dissolve. The distribution

of CaCO

on the sea floor resembles snow-capped mountains,

with preservation on the mountaintops and none in the valleys.

The depth of the “snow line,” where CaCO

is no longer

preserved in sediments, has been called the carbonate compen-

sation depth, or CCD. In practice, sedimentary CaCO

dissolu-

tion is complicated by the rate of diffusive contact of the pore

water with overlying water, and by other sediment reactions

such as organic carbon respiration. Not all of the CaCO

depos-

ited above the saturation horizon is buried, and some of the

CaCO

flux to undersaturated sediments may be preserved.

The CCD (a property of the solid sediment) is not the same

as the saturation horizon (a property of the water column). In

most parts of the ocean, the CCD is 0.5–1.5 km deeper than

the saturation horizon (see Carbonate compensation depth;

Carbon dioxide, dissolved (ocean)).

The pH-determining mechanism of the ocean, called carbonate

compensation (beware of the unfortunate overlap with the sepa-

rate concept of the carbonate compensation depth or CCD), arises

from the need for the ocean to balance the input and burial rates of

CaCO

. The way the ocean regulates burial to balance weathering

is through the pH of the ocean. The input of CaCO

from weather-

ing tends to drive the ocean towards the basic, as a source of excess

alkalinity. The neutralization of the primary CO

2

dissolution pro-

duct by atmospheric CO

to produce HCO



in Reaction 8 is a con-

sumption of the acid from the atmosphere and ocean by CaCO

base. If this source of base were unbalanced by CaCO

burial

in sediments, the ocean would get progressively more basic over

time. This would increase the proportion of the dissolved carbon

that is CO

2

, increasing the depth range over which CaCO

is stable and increasing the burial rate of CaCO

, until ultimately

the burial rate would balance the weathering rate. The time

scale over which this relaxation takes place is 5–10 kyr

(Archer et al., 1997).

Numerous factors can drive CaCO

compensation. An

increase in chemical weathering would drive the CCD deeper,

lowering atmospheric pCO

. Today, a significant fraction of

CaCO

deposition occurs on continental shelves and carbonate

platforms, which may have been exposed during times of low-

ered sea level such as the Last Glacial Maximum (LGM). A

decrease in shallow-water burial would have required an

increase in deep sea burial to compensate, also lowering

pCO

. A decrease in the production rate of CaCO

, driven per-

haps by ocean chemistry or climate change, would require that

a greater proportion of the CaCO

produced be buried, requir-

ing a deeper saturation horizon and CCD, resulting in lower

pCO

. Finally, a CaCO

compensation response in the future

will ultimately neutralize much of the fossil fuel CO

that

humankind is releasing to the atmosphere. The CO

invasion

will acidify the ocean, lowering [CO

2

] and decreasing CaCO

burial. Excess CaCO

source to the ocean will drive the ocean

toward the basic, neutralizing the excess CO

into the form of

dissolved HCO



in the ocean water column.

David Archer

Bibliography

Archer, D.E., 1996. An atlas of the distribution of calcium carbonate in

sediments of the deep sea. Glob. Biogeochem. Cycles, 10, 159–174.

Archer, D., Kheshgi, H., and Maier-Riemer, E., 1997. Multiple timescales

for neutralization of fossil fuel CO

. Geophys. Res. Lett., 24, 405–408.

Armstrong, R.A., Lee, C., Hedges, J.I., Honjo, S., and Wakeham, S.G.,

2002. A new, mechanistic model for organic carbon fluxes in the ocean

based on the quantitative association of POC with ballast minerals.

Deep Sea Res. II , 49, 219–236.

Betzer, P.R., Byrne, R.H., Acker, J.G., Lewis, C.S., and Jolley, R.R., 1984

The oceanic carbonate system: A reassessment of biogenic controls.

Science, 226, 1074–1077.

Broecker, W.S., and Peng, T.H., 1982. Tracers in the Sea. Palisades, NY:

Eldigio Press.

Broecker, W.S., and Peng, T.H., 1987. The role of CaCO

compensation in

the glacial to interglacial atmospheric CO2 change. Glob. Biogeochem.

Cycles, 1,15–29.

Dickson, A.G., and Millero, F.J., 1987. A comparison of the equilibrium

constants for the dissociation of carbonic acid in seawater media. Deep

Sea Res., 34, 1733–1743.

Honjo, S., 1996. Fluxes of particles to the interior of the open oceans. In

Ittekkot, V., Aschauffer, P., Honjo, S., and Depetris, P. (eds.), Particle

Flux in the Ocean. New York: Wiley.

Klaas, C., and Archer, D.E., 2002. Association of sinking organic matter

with various types of mineral ballast in the deep sea: Implications for

the rain ratio. Glob. Biogeochem. Cycles, 16, 10.1029/ 2001GB001765.

Longhurst, A., 1998. Ecological Geography of the Sea. San Diego, CA:

Academic Press, 398pp.

Milliman, J.D., 1993. Production and accumulation of calcium carbonate in

the ocean: Budget of a non-steady state. Glob. Biogeochem. Cycles, 7,

927–957.

Stumm, W., and Morgan, J.J., 1981. Aquatic Chemistry. New York: Wiley,

780pp.

Wanninkhof, R., 1992. Relationship between wind speed and gas exchange

over the ocean. J. Geophys. Res., 97(C5), 7373–7382.

Yu, E.-F., Francois, R., Bacon, M.P., Honjo, S., Fleer, A.P., Manganini, S.J.,

Loeff, M.M.R.v.d., and Ittekot, V., 2001. Trapping efficiency of

bottom-tethered sediment traps estimated from the intercepted fluxes

230

Th and

231

Pa. Deep Sea Res. I, 48, 865–889.

538 MARINE CARBON GEOCHEMISTRY