Odekon M. Encyclopedia of paleoclimatology and ancient environments

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Marine environments

Initial isotopic analysis of foraminifera in Cenozoic marine

sediment cores (Emiliani, 1972) was focused on the oxygen

isotope composition as related to temperature or sea level (ice

volume). Interpretation of d

C measurements on the same

cores was delayed somewhat because of even more compli-

cated environmental influences (Hoefs, 1987), but Shackleton

and Kennett (1975) did find benthic forams consistently

depleted relative to planktonic forams by about 1%. In part,

this is a consequence of planktonic forams incorporating dis-

solved carbonate species enriched by preferential removal of

the isotopically light dissolved carbon during photosynthesis.

There is now a 65 million-year record of d

C variation of car-

bonates in deep-sea sediment cores (Zachos et al., 2001) that

shows overall variation of almost 4%, with the lowest values

occurring in the last 6 million years, and evidence of precipi-

tous but brief d

C excursions (2–3%), one of the largest

occurring 55 million years ago.

The influences on this variability have been studied.

Shackleton et al. (1983) used the d

C record of forams to infer

greatly reduced CO

atmospheric concentrations ( pCO

) dur-

ing the last full-glacial period, now known to be true in light

of subsequent paleoatmosphere measurements in ice cores.

Pagani et al. (1999) have since used compound-specific d

analysis of alkenones derived from algae organic matter

compared with d

C of planktonic forams to more precisely

estimate pCO

(Figure C22). This is based on foram d

representing the isotopic composition of aqueous CO

used in

photosynthesis, and the fractionation between organic d

and carbonate d

C being a function of aqueous CO

concen-

tration (itself a function of pCO

). The trend in calculated

pCO

thus derived suggests CO

concentrations were very

low 5–10 million years before a major continental C

plant

expansion occurred around 8–4 million years ago, according

to evidence from the d

C of teeth, soil carbonates, and terres-

trial organic matter (Cerling et al., 1998). This implies that fac-

tors other than pCO

must have driven that C

boom.

Corals have provided a wealth of paleoenvironmental infor-

mation from their chemical and isotopic composition. Coral

C variations have been attributed to a number of effects,

such as variations in food source or dissolved carbonate isoto-

pic composition (related to ecosystem productivity), but there

is a growing body of evidence that at least in some cases light

may be a dominating (and reconstructable) influence, whereby

high levels promote more photosynthesis and less discrimina-

tion against

C (Swart et al., 1996).

Quay et al. (1992) demonstrated the merit in repeated mea-

surement of d

C of dissolved inorganic carbon over two dec-

ades to help estimate ocean uptake of atmospheric CO

over

time and space. Recent d

C investigations of calcareous

sponges (Böhm et al., 2002), which deposit annual and subann-

ual banding like corals, have further revealed their potential to

reliably reconstruct d

air

over hundreds of years or more.

A current area of great scientific interest is the possibility

that methane (CH

) in the form of hydrates or clathrates has

periodically been released catastrophically from deep shelf

sediments as a result of decreasing sea level, increasing tempera-

ture of the ocean water, or even submarine slope failure. Because

methane is a powerful greenhouse gas, there is concern that any

such future events might exacerbate global warming. Isotopic

analysis of benthic foraminifera over the past 60,000 years from

the Santa Barbara Basin reveals periodic large releases of

methane (Kennett et al., 2000). Interstadials (warm intervals)

are characterized by foram d

C values depleted by several per-

mil. Detailed compound-specific d

C analysis of diplopterol, a

product of methanotrophic bacteria, in sediments of the Santa

Barbara Basin further confirms that the most negative values

(interpreted as highest marine methane) occur during intersta-

dials (Hinrichs et al., 2003).

Past and future

Stable carbon isotopes have been a highly valuable tool for the

past 50 years in helping to build our current understanding of

terrestrial and aquatic paleoenvironments, variously including

past moisture conditions, temperature, light, plant productivity,

the carbon cycle, and more. In some cases, d

C has provided a

direct proxy of absolute conditions in past systems, and in

others it has helped to constrain the range of likely paleoenvir-

onmental conditions or provided independent support for condi-

tions inferred from other evidence. With exciting improvements

in mass-spectrometry, d

C measurements will unquestionably

reveal even more about past changes in the Earth system.

Advances such as compound-specific analysis are beginning

to provide isotopic measurements of biomarkers exclusive to

Figure C22 Atmospheric CO

concentrations between 16 and 5

million years ago, reconstructed from the difference between d

of alkenones of organic matter and foram carbonates. (After Pagani

et al., 1999.)

Figure C21 Shift in isotopic composition of bone collagen in the

Fremont paleoculture of Utah from enriched values consistent with

maize in the diet before

AD 1100 and a C

-dominated diet afterward

(after Coltrain and Leavitt, 2002). The period in which maize could be

cultivated is attributed to farther northward penetration of the

southwestern summer monsoon.

CARBON ISOTOPES, STABLE 135

land or aquatic organisms, to species or genera, and even to

particular plant organs. Improvements in both the speed of ana-

lysis and ability to analyze extremely small quantities of com-

pounds extracted from bulk samples enable even inherently

microscopic materials such as pollen to be analyzed.

Steven W. Leavitt

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Coplen, T.B., 1996. New guidelines for reporting stable hydrogen, carbon

and oxygen isotope-ratio data. Geochimica et Cosmochimica Acta,

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500,000-year stable carbon isotopic record from Devils Hole, Nevada.

Science, 263, 361–365.

Ehleringer, J.R., Cerling, T.E., and Helliker, B.R., 1997. C

photosynthesis,

atmospheric CO

, and climate. Oecologia, 112, 285–299.

Emiliani, C., 1972. Quaternary paleotemperatures and the duration of the

high-temperature intervals. Science, 178, 398–401.

Farquhar, G.D., 1983. On the nature of carbon isotope discrimination in C

species. Aust. J. Plant Physiol., 10, 205–226.

Farquhar, G.D., O’Leary, M.H., and Berry, J.A., 1982. On the relationship

between carbon isotope discrimination and the intercellular carbon

dioxide concentration in leaves. Aust. J. Plant Physiol., 9, 121–137.

Frappier, A., Sahagian, D., Gonzalez, L.A., and Carpenter, S.J., 2002. El

Niño events recorded by stalagmite carbon isotopes. Science, 298, 565.

Hoefs, J., 1987. Stable Isotope Geochemistry. New York: Springer.

Hinrichs, K.-U., Hmelo, L.R., and Sylva, S.P., 2003. Molecular fossil

record of elevated methane levels in Late Pleistocene coastal waters.

Science, 299, 1214–1217.

Huang, Y., Street-Perrott, F.A., Metcalfe, S.E., Brenner, M., Moreland, M.,

and Freeman, K.H., 2001. Dominant control on glacial-interglacial var-

iations in C3 and C4 plant abundance. Science, 293, 1647–1651.

Kennett, J.P., Cannariato, K.G., Hendy, I.L., and Behl, R.J., 2000. Carbon

isotopic evidence for methane hydrate instability during Quaternary

interstadials. Science, 288, 128–133.

Leavitt, S.W., and Long, A., 1989. Drought indicated in carbon-13/ carbon-

12 ratios of southwestern tree rings. Water Resour. Bull., 25, 341–347.

Monger, H.C., Cole, D.R., Gish, J.W., and Giordano, T.H., 1998. Stable

carbon and oxygen isotopes in Quaternary soil carbonates as indicators

of ecogeomorphic changes in the northern Chihuahuan Desert, USA.

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Pagani, M., Freeman, K.H., and Arthur, M.A., 1999. Late Miocene atmo-

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Quade, J., and Cerling, T.E., 1995. Expansion of C

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Briffa, K.R., and Jones, P.D., 1997. Signal strength and climate relation-

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Cross-references

Alkenones

Carbon Cycle

Carbon Dioxide and Methane, Quaternary Variations

Carbon Isotope Variations over Geologic Time

Isotope Fractionation

Marine Biogenic Sediments

Marine Carbon Geochemistry

Methane Hydrates, Carbon Cycling, and Environmental Change

Radiocarbon Dating

Stable Isotope Analysis

CARBONATE COMPENSATION DEPTH

Principles

The carbonate compensation depth (CCD) is the particular

depth level in the oceans where the rate of supply of calcium

carbonate to the sea floor is balanced by the rate of dissolution.

In practice, the CCD is mapped as the level at which percent

carbonate values in marine sediments drop towards zero,

i.e., the depth transition from carbonate ooze to red clay or

siliceous ooze. Because the supply and dissolution rates of car-

bonate differ from place to place, the depth of the CCD is

variable. In the Indian Ocean and the Pacific, the CCD is typi-

cally found at depths between 3,500 and 4,500 m. In the North

Atlantic and the eastern South Atlantic, the CCD occurs deeper

than 5,000 m. Close to continental margins, the CCD tends

to shoal due to higher biological productivity and hence

respiration of organic matter in sediments, which releases meta-

bolic CO

, in turn causing the dissolution of CaCO

. Beneath

the equatorial upwelling zones, higher primary production

of planktonic carbonate shells depresses the CCD by up to

500 m compared with the oligotrophic realms towards the north

and south.

The depth of the CCD is determined by the kinetics of car-

bonate dissolution from sediments. If dissolution kinetics were

infinitely fast, the CCD would coincide exactly with the satura-

tion horizon of carbonate (the so-called chemical lysocline),

which in turn is a function of pressure, temperature, salinity,

and carbonate ion concentration of the overlying seawater

(Figure C23), with generally increasing solubility in deeper

waters. However, in real oceans the CCD is found at some

degree of undersaturation, i.e., a few hundred meters below the

lysocline (Archer, 1996). Because shells or skeletons made of

aragonite (mainly corals and pteropods) dissolve more readily

than those of calcite (coccoliths and foraminifers), the compen-

sation depth for aragonite (ACD) is positioned some 1,000–

2,000 m above the CCD. However, organic coatings of grains

136 CARBONATE COMPENSATION DEPTH

may inhibit the dissolution of sedimenting carbonate particles,

thereby depressing the positions of lysocline and CCD below

the predicted levels from carbonate chemistry. On the other

hand, oxidation of organic matter in the sediments releases

metabolic CO

and promotes CaCO

dissolution even above

the chemical lysocline (Milliman et al., 1999).

Over longer timescales, the depth of CCD in the oceans

is the expression of the mass balance of dissolved Ca

and

HCO



flowing into the oceans from weathering, and the

removal of CaCO

from the ocean by burial. After an equilibra-

tion time of the order of 10,000 years, CaCO

compensation

will have a profound impact on the buffering of the pH of the

ocean. An imbalance between sources and sinks of CaCO

will

drive the pH of the ocean toward restoring balance. For exam-

ple, excess input of dissolved CaCO

will tend to drive the pH

of the ocean toward higher, more basic values, deepening the

CCD and increasing CaCO

accumulation on the seafloor until

the rates return to balance.

Modern ocean distribution and Quaternary changes

The asymmetry of the modern ocean CCD pattern is largely a

product of the thermohaline circulation. The greatest depth

occurs in the eastern North Atlantic, where the production of

supersaturated deep water increases the CCD to >5,500 m.

The shallowest levels are in the northern North Pacific, where

deep waters are old and enriched in CO

. This pronounced

modern pattern has been termed basin-to-basin fractionation

and is illustrated by the fact that roughly 65% of the present

Atlantic seafloor is covered by carbonate ooze, while only

54% of the Indian Ocean sea floor and 36% of the Pacific

sea floor share that distinction (Archer, 1996).

Regarding Quaternary times, observations from deep-sea

sediment cores show that the CCD topography varied signifi-

cantly over the glacial to interglacial cycles characteristic of

the last 2.6 Myr of Earth history. The CCD fluctuations

of the Pacific (first described by G. Arrhenius) are com-

monly described as dissolution cycles, with high dissolution

in interglacials and vice versa. Carbonate productivity plays a

secondary role. On much of the Atlantic sea floor, the CCD

fluctuations tend to change to counter those of the Pacific, with

higher dissolution occurring during glacials. These changes in

carbonate preservation are ascribed to changes in the formation

of North Atlantic Deep Water, but also to changes in terrestrial

carbon storage and in the fractionation between shallow and

deep ocean carbonate deposition.

Pre-Quaternary history

Studies of longer cores from ocean floor drilling show that CCD

fluctuations were similar in the various oceans and therefore dri-

ven by global mechanisms (Figure C24). The mean ocean CCD

was about 1,000 m shallower than present during the Cretaceous

and early Paleogene, followed by deepening levels going into

Figure C23 Calcite saturation data of GEOSECS station 441 (5



S, 91



E) compared with carbonate contents (weight%) from surface sediment

samples of the eastern equatorial Indian Ocean. (After Peterson and Prell, 1985.)

Figure C24 Carbonate compensation depth variation in different ocean

basins through the late Cenozoic. (Redrawn from Van Andel, 1975.)

CARBONATE COMPENSATION DEPTH 137

the Oligocene. Another shallowing occurred during the mid

Miocene, before the CCD dropped to its present mean depth

near 4,300 m. The drastic drop at the end of the Eocene is global

in scale and reflects changes in the global carbon cycle, possi-

bly linked to an increase in alkalinity due to higher silicate

weathering caused by the Himalayan uplift (Francois and

Walker, 1992). It might also be linked to the lowering of sea level

at that time, reducing the area of shallow epicontinental seas

available for the accumulation of coral reef carbonate. The late

Neogene CCD drop is more pronounced in the North Atlantic

as compared with that of all other ocean basins. It reflects mostly

the increase in the production of North Atlantic Deep Water

related to the opening and closing of ocean gateways at that time

(Heinze and Crowley, 1997). However, there is still controversy

about the cause of the CCD fluctuations in Earth history.

Torsten Bickert

Bibliography

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

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

Francois, L.M., and Walker, J.C.G., 1992. Modelling the Phanerozoic car-

bon cycle and climate: Constraints from the

87 / 86

Sr isotopic ratio of sea-

water. Am. J. Sci., 292,81–135.

Heinze, C., and Crowley, T.J., 1997. Sedimentary response to ocean gate-

way circulation changes. Paleoceanography, 12, 742–754.

Milliman, J.D., Troy, P.J., Balch, W.M., Adams, A.K., Li, Y.H., and

Mackenzie, F.T., 1999. Biologically mediated dissolution of calcium car-

bonate above the chemical lysocline? Deep Sea Res. I, 46, 1653–1669.

Peterson, L.C., and Prell, W.L., 1985. Carbonate dissolution in recent sedi-

ments of eastern equatorial Indian Ocean: Preservation patterns and car-

bonate loss above the lysocline. Mar. Geol., 64, 259–290.

Van Andel, T.H., 1975. Mesozoic-Cenozoic calcite compensation depth and

the global distribution of calcareous sediments. Earth Planet. Sci. Lett.,

26, 187–194.

Cross-references

Carbon Cycle

Carbonates, Cool Water

Carbonates, Warm Water

Marine Carbon Geochemistry

Paleo-Ocean pH

Pre-Quaternary Milankovitch Cycles and Climate Variability

Quaternary Climate Transitions and Cycles

CARBONATES, COOL WATER

Introduction

The deposition of marine carbonate sediments in cool water

environments has been long recognized (James, 1997). Despite

this, the focus of study on carbonate sediments and their diagen-

esis has been carried out in warm, tropical waters in places

such as the Caribbean, the Persian Gulf, the Great Barrier

Reef, and Shark Bay. Not only do these places provide clear ana-

logs for many ancient deposits, they are also often located

in environments that are both pleasant to work in and easily

accessible. Cool-water carbonates are, by way of contrast, depos-

ited beneath some of the coldest and roughest waters of the

world such as the Southern Ocean and North Atlantic. Further-

more, many of the best-exposed examples occur in Australia

(Figures C25 and C26) and New Zealand, remote from European

and North American centers of carbonate research.

Such disadvantages aside, understanding cool-water carbo-

nates and their diagenesis is a vital component of carbonate

sedimentology. They are currently being deposited across

large areas of continental shelf, form extensive successions in

Cenozoic basins, and are also known from many ancient high

latitude environments. In addition, by virtue of their calcite-

dominated mineralogy, cool-water carbonates provide excellent

analogs for many Paleozoic limestones that were likewise

dominated by calcite, in contrast to modern tropical carbonates,

which are rich in aragonite.

History of investigations

Among the early studies, those of Chave (1967), Lees

and Buller (1972) and Lees (1975) stand out as having estab-

lished the conceptual framework for cool-water carbonate

deposition. Meanwhile, Connolly and von der Borch (1967)

and Wass et al. (1970) had discovered that cool-water carbo-

nates were being deposited across the entire length and width

of the southern Australian continental shelf. More recent

compilations include the 1988 special issue of Sedimentary

Geology (vol. 60), James and Clarke (1997), and Pedley and

Carannante (2006). The last compilation paid particular atten-

tion to the cool-water carbonates of the Mediterranean Basin.

Furthermore, ODP Leg 182 recovered a number of cores from

the southern Australian margin in the Great Australian Bight

(Feary et al., 2000).

Distinctive characteristics

James (1997) lists six characteristics for recognition of ancient

cool-water carbonates. In decreasing order of reliability, these

are:

1. Carbonates interbedded with glaciogene sediments

2. Carbonates situated in appropriate paleolatitudes

3. Glendonite-type minerals in carbonates

4. Carbonates composed wholly of the Heterozoan Association

5. Heterozoan Association outboard of mid to inner shelf/

ramp Photozoan Association

6. Trace element and isotope geochemistry of skeletons

Of these, 1–3 are self explanatory but not always present, while

6 can be equivocal because of diagenesis. The following sec-

tion therefore focuses on 4 and 5 as the characteristics that

are most common and most readily determined by geological

analysis.

Modern warm-water carbonates are dominated by photozo-

ans, organisms that are either photosynthetic or are zooanthel-

late, with symbiotic photosynthetic organisms. These form the

Photozoan Assemblage. In contrast, cool-water carbonates are

predominantly dominated by Heterozoans, organisms that are

non-photosynthetic and lack photosynthetic symbionts (James,

1997), and thus they are mostly dominated by bryozoans, with

varying proportions of mollusks, foraminifers, echinoderms,

brachiopods, barnacles, sponges, and coralline algae. As a

result of this dichotomy, warm-water benthic carbonates are

formed primarily in shallow water, especially in the first

30 meters where light is most intense. The biota that form

cool-water carbonates are not limited to the photic zones, apart

from coralline algae. Therefore, benthic carbonate production is

not depth-limited but continues across the entire continental

shelf and onto the upper continental slope. In addition to

138 CARBONATES, COOL WATER

the biota of the sediment surface, many of the organisms that

contribute carbonate to these sediments grow prolifically on

hard substrates, including rock reefs and sea floor hardgrounds.

These rocky areas are in some ways analogous to the coral

reefs of warm-water carbonates in that they provide a localized

factory for carbonate production (Nelson et al., 1988a). Simi-

larly, in shallow water, sea grass leaves provide a good sub-

strate for epibionts, and the sheltered environment of sea

grass banks (Figure C27) both encourages carbonate produc-

tion and traps sediment (Gostin et al., 1984). In deeper water,

bryozoan and sponge mounds are not uncommon (James et al.,

2000), appearing in the geologic record both as discrete bio-

herms and biostromes.

In addition to their taxonomic distinctiveness, the Heterozan

and Photozoan Assemblages differ in the dominant mineralogy

produced. The Photozoan biotas produce skeletons that are pre-

dominantly aragonite and high magnesium calcite. Heterozoan

biotas are predominantly calcite, ranging from low to high

magnesium contents. Aragonite is a comparatively minor com-

ponent. This has major implications for the diagenesis of cool-

water carbonates (see below).

Warm-water carbonates commonly contain ooids (formed in

agitated shallow marine environments), abundant micrite (carbo-

nate mud) formed either by disintegration of calcareous green

algae or direct precipitation from seawater super-saturated with

respect to carbonate, and intraclasts, formed by rapid sea floor

Figure C25 The “Twelve Apostles,” sea stacks formed by erosion of the Oligo-Miocene Gambier Limestone, a cool-water carbonate on the south

coast of Victoria, Australia (photo H. Apps).

Figure C26 Point Culver, Western Australia. Late Eocene cyclic cool-water chalky limestones rich in siliceous sponges, (photo by author).

CARBONATES, COOL WATER 139

cementation, again from super-saturated water. Ooids are absent

in cool-water environments where saturation states are lower.

True micrite is absent, and fine-grained carbonates are either car-

bonate silt formed by the erosion of other carbonates or cocco-

liths. Intraclasts are rare in cool-water successions, largely

because the lower saturation state of carbonates in cooler marine

waters reduces the rate of sea floor cementation.

Distribution

Modern marine carbonates occur wherever rates of biogenic

carbonate production exceed terrigenous sediment supply.

However, because biologic carbonate production is lower in

cool as opposed to warm water, cool-water carbonates are gen-

erally found in areas of reduced terrigenous sedimentation

(Chave, 1967). Gradations between warm and cool-water car-

bonates are known, such as on the west and southwest coasts

of Australia (Cann and Clarke, 1993; Collins et al., 1997).

However, a mean annual temperature of 22



C or less serves

as a useful dividing line between the warm and cool (or cold)

water depositional realms (James, 1997). Cooler waters typi-

cally have a lower carbonate saturation than warmer waters,

and this influences both the production of grains and the diag-

enesis of cool-water carbonates (see below).

Major settings for cool-water carbonate deposition include

shelves and ramps, offshore banks, and large embayments.

One of the largest and best-studied areas of cool-water carbo-

nate shelf deposition is the continental shelf of southern

Australia (James et al., 2001, and references therein). Rockhall

(Scoffin et al., 1980) and Porcupine (Scoffin and Bowes, 1988)

Banks in the North Atlantic are examples of modern offshore

banks characterized by cool-water carbonate sediments. Large

embayments floored by cool-water carbonates include Gulf

St. Vincent (Shepherd and Sprigg, 1976) and Spencer Gulf of

South Australia (Gostin et al., 1984).

In addition to temperature, several other processes can give

rise to a cool-water aspect in otherwise warm water carbonate

environments. One is turbidity, where high levels of suspended

sediment reduce the competitive advantage of the photozoan

biota so that, even in shallow water, the carbonate component

may be dominated by mollusks, bryozoan, and other hetero-

zoan biota. The sediments in Bonaparte Gulf in northern

Australia provide one such example (Clarke et al., 2001).

Another is high nutrient loadings, which, in shallow water, will

favor algal growth over that of zooanthellate organisms (James,

1997). In warmer waters, the algae will include calcareous

green forms, but in areas where water temperatures are mar-

ginal for calcareous green algae, red algae and other compo-

nents of the heterozoan biota will dominate. A third factor is

water depth. Below the photic zone, biota is typically hetero-

zoan in aspect. Deeper waters are cooler than shallow waters,

especially where there is a well-developed thermocline. Deep

shelf or upper slope carbonates of heterozoan aspect may there-

fore occur seaward either shallow cool-water shelf heterozoan

carbonates or shallow warm-water shelf photozoan carbonates

(James, 1997 ). Finally, cool or cold currents moving towards

the equator along continental margins can result in conditions

suitable for the Heterozoan Assemblage in low latitudes.

Diagenesis

Tropical carbonates often experience rapid seafloor cemen-

tation (Tucker and Wright, 1990). In contrast, cool-water carbo-

nates for the most part experience limited seafloor cementation

because of generally lower levels of carbonate saturation and

the slower reaction times at lower temperatures. Rao (1981a)

showed the presence of minor calcite seafloor cements in sedi-

ments on the Tasmania continental shelf. Ancient cool-water

carbonates similarly show only limited seafloor cementation

for the most part (e.g., Nicolaides, 1994 and Shubber et al.,

1997). During exposure to meteoritic waters, limestones con-

taining abundant aragonite have undergone extensive carbonate

dissolution, which is reprecipitated as calcite. Dissolution and

reprecipitation stabilizes the limestones and causes extensive

lithification even in young carbonates that have undergone

little burial, such as those of exposed Pleistocene reefs.

Because aragonite is much less abundant in most cool-water

carbonates, they do not undergo anywhere near as extensive

dissolution and reprecipitation. In consequence, many cool-

water carbonates that have undergone limited burial, such as

Figure C27 Seagrass beds and carbonate-rich sand, Esperance, Western Australia (photo by N. James).

140 CARBONATES, COOL WATER

the Oligocene carbonates of southern A ustralia, are only

weakly lithified, despite their age (James and Bone, 1989).

Most cementation in cool-water carbonates appears to be

related to burial. This contrasts with aragonite-rich carbo-

nates, where burial cements are more comm onl y associated

with infill of remnant porosity after extensive seafloor and

meteoric cementation. The Oligocene Te Kuiti Limestone of

New Zealand has been buried to a depth of up to 1,100 m

(Nelson et al., 19 88b) and is, as a result, fully lithified and

strongly stylolitized. It is likely that much of the carbonate

needed to fully cement calcitic, cool-water limestones comes

from pressure dissoluti on (Nicolaides and Wallace, 1997).

Pressure-dissolution appears to require at least 200 m of

burial to begin the process.

Examples of cool-water carbonates though

geologic time

Cenozoic cool-water carbonates are best known from southern

Australia, New Zealand,and southern Europe. The Australian

examples range in age from Eocene to Pliocene and are very

extensive, but have experienced little burial (rarely more than

a few tens of meters), although some (Nicolaides and Wallace,

1997) have had several hundred meters of burial (Figure C28).

Most formed on Cenozoic shelves and slopes (e.g., Nicolaides,

1994; Bernecker et al., 1997); however, some (Clarke et al.,

1996; Shubber et al., 1997) formed in sheltered embayments

similar to the modern gulfs of southern Australia. Because of

their very limited burial diagenesis, they provide a useful link

between the carbonates of modern cool-water environments

and those that have experienced deep burial and subsequent

uplift. Cool-water carbonates have also been documented

from the margins of the Mediterranean, including examples

from Spain (Franseen et al., 1997; Martin et al., 1996), Austria

(Nebelsick, 1992), Italy (Bassi et al., 2006), and Greece

(Spjeldnaes and Moisette, 1997). Cool-water carbonates of

New Zealand, such as those of the Oligocene Te Kuiti Group

(Nelson et al., 1988b), have experienced deep burial and are

fully lithified, unlike their Australian and Mediterranean coun-

terparts. New Zealand carbonates were deposited on both shelf

and platform environments (Nelson, 1978).

The well-known Cretaceous and Danian chalks of Europe

(Scholle et al., 1983) are now regarded as cool-water carbonates.

These were deposited in a range of environments directly analo-

gous to those of modern southern Australia, from high-energy

skeletal grainstones along the shores to offshore bryozoan

mounds and deep shelf coccolith lime mudstones (Surlyk,

1997). Other documented examples of Mesozoic carbonates with

cool-water affinities are the open shelf carbonates of Sardinia and

the southern Appenines (Carannante et al., 1997).

The first Paleozoic carbonates recognized as being of cool-

water origin were the Permian limestones of Tasmania. These

were shown by Rao (1981b) to have formed on a polar shelf mar-

ginal to the Gondwana glaciation on the basis of their dropstones

and close association with glaciomarine facies. In addition, they

were deposited at a high paleolatitude and had a primary calcite-

dominant mineralogy. Other cold-water carbonates of Tasmania

include the Darlington Limestone (Brill, 1982), deposited in

a near-shore facies. Equivalent sediments of the Northern

Hemisphere are found in the Permian of the Sverdrup Basin

(Beauchamp and Desrochers, 1997; Stemmerik, 1997).

Other known Paleozoic cool-water carbonates have been

described from the Early Carboniferous of western Canada

(Brandley and Krause, 1997; Martindale and Boreen, 1997),

to the Middle Ordovician of Canada by Brookfield (1988)

and Kentucky by Pope and Read (1997). Lavoie (1995) pro-

vided a Late Ordovician example from Quebec. These three

examples all occur at relatively low paleolatitudes. They are

attributed to a combination of factors including upwelling, ther-

mally stratified oceans, and cool or cold currents flowing

towards the paleoequator.

Jonathan Clarke

Figure C28 Oligo-Miocene Gambier Limestone exposed in cliffs near the Twelve Apostles, showing characteristic cyclic bedding (photo H. Apps).

CARBONATES, COOL WATER 141

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Cross-references

Carbonates, Warm Water

Glacial Sediments

Glaciomarine Sediments

Glendonite/ Ikaite

Marine Biogenic Sediments

CARBONATES, WARM WATER

Sediments composed of calcium carbonate form a large, but

not major, part of the world’s continental shelf seas and deep

ocean basins (Figure C29). The best-known examples of

these sediments are coral reefs and white beaches made up

of calcareous (CaCO

) sand grains, typically seen in tropical

environments.

What are warm-water carbonate sediments?

Most carbonate sediments are formed from the accumulation of

skeletons and shells constructed by marine organisms through

the precipitation of calcium carbonate (e.g., corals, mollusks

and foraminifera). These are called skeletal, or bioclastic, carbo-

nate sediments. Skeletal carbonates occur in both the warm and

the cold waters of the shelf areas in the world (Figure C29).

They are most common on continental shelves that are not

diluted by large amounts of eroded and transported siliciclastic

sediments, as is the case near major river mouths and adjacent

to mountainous coastal regions. The composition of skeletal

carbonates differs in warm and cool water areas because of

the different skeletal organisms that live in these areas. Today’s

warm or tropical waters are characterized by sea-floor, or

benthic, communities of the rapidly calcifying corals and calcar-

eous green algae, together with mollusks, foraminifera, echi-

noids, bryozoans, etc. Coral reefs today are restricted to

waters warmer than 18



C (Kleypas et al., 1999). Other carbo-

nate sediments are formed by the chemical precipitation of crys-

tals of calcium carbonate from seawater, with some authors

arguing for a microbial mediation to this process (Riding and

Awramik, 2000). Precipitation is favored in the warm, saline

waters of some tropical seas and results in spherical, sand-sized

grains known as ooids, aggregate grains, peloids and precipitated

lime mud (for details, see below). These are labeled as non-

skeletal carbonate sediments and are only found in warm-water

areas. Lees and Buller (1972) found that modern ooids and grain

aggregates form in waters with maximum annual temperatures

>20



C and in waters with salinities greater than 37% (parts

per thousand). This temperature together with the temperature

Figure C29 Global distribution of modern carbonate sediments. The two case studies discussed in the text are also located (after Bosence and

Wilson, 2003).

CARBONATES, WARM WATER 143

restriction for coral reefs (i.e., 18–20



C) may therefore be taken

as the boundary between what are regarded as warm (or tropical)

and cool (or temperate) water carbonate sediments.

Warm-water carbonates through geological time

The abundance of carbonate rocks formed in marine shelf

environments has varied considerably over geological time

(Figure C30). Not surprisingly, they were most abundant when

global sea levels were high and large areas of continents were

flooded by shallow, warm-water, shelf seas. These conditions

occurred several times in Earth history, during “greenhouse

periods” when there were prolonged periods with no major

ice caps covering the polar regions (Figure C30). The types

of warm-water organisms that produced skeletal carbonates

changed as life evolved, so that late Precambrian, early and late

Paleozoic, Mesozoic and Cenozoic bioclastic limestones are

quite different in their composition (Figure C30). An indepen-

dent measure of paleotemperature may be made by analyzing

the oxygen isotope ratios of original skeletal calcium carbonate

of fossils within bioclastic limestones. Similarly, the mineral-

ogy of ooids also changed through time in tune with green-

house (low magnesium calcite ooids) and icehouse periods

(aragonitic and high magnesium calcite ooids, Figure C30).

The abundance of ooids has also changed with time, and possi-

bly with sea-level changes, as peaks are found at times of major

long term sea-level change (Figure C30). Although the miner-

alogy of ooids has changed through time, and there is still

debate as to what controls this (see discussion in Tucker and

Wright, 1990), all indications are that ooids still represent

ancient environments of increased temperature and increased

salinity. The use of non-skeletal marine carbonate grains (such

as ooids and grain aggregates) has proven to be more reliable

than fossils as a warm-water indicator. This is because the for-

mation of non-skeletal grains is predominantly controlled by

the chemical conditions at or near the seafloor and, apart from

the changes illustrated in Figure C30, these do not appear to

have changed significantly through time. Studies on the nature

of warm-water carbonate rocks over the last 650 Ma indicate

that the present is not always the key to the past. A 400 Ma

warm-water limestone will differ in its preserved biota and ooid

mineralogy from its present day equivalent and care has to be

taken in undertaking any paleoenvironmental analysis based

on carbonate grain types.

Case studies of warm-water carbonate depositional

environments; humid versus arid settings

Two case studies of present-day environments accumula-

ting warm-water carbonate sediments are given below. These

examples are used extensively in the interpretation of ancient

environments as they illustrate the range of different sediment

Figure C30 Secular changes in warm-water carbonate grains with respect to sea-level changes and icehouse–greenhouse periods. For details, see

text (modified from many studies, for details see Tucker and Wright, 1990; Bosence and Wilson, 2003; Hallam, 1984; Vail et al., 1977).

144 CARBONATES, WARM WATER