Odekon M. Encyclopedia of paleoclimatology and ancient environments

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diversity, but other types of coral accumulations found in deep

water are also of great size. These structures, sometimes called

“bioherms,” are commonly formed of skeletal debris accumu-

lated from one or a few species of azooxanthellate corals.

Zonation

Most accounts of coral reefs draw attention to the coral zones

that they invariably display. These zones have been likened to

zones of vegetation found on mountain slopes but on coral

reefs they are mostly created by decreasing light and turbulence

with increasing depth and, even in very clear water, this occurs

within a depth range of less than 50 meters. Contrary to popu-

lar belief, species diversity is not at its highest on the scenically

attractive reefs found in clear oceanic water. It reaches a peak

for any given region on fringing reefs protected from strong

wave action, where the water is slightly turbid. This is probably

because Acropora, by far the most common Indo-Pacific coral,

does not dominate these habitats, allowing for a higher diver-

sity of other groups of coral.

Distribution patterns

The global distribution of reefs and the global pattern of

coral diversity are linked, but not strongly (Figure C67). The

Indonesian/Philippines archipelago has the world’s greatest con-

centration of reefs and the greatest coral diversity. Other areas of

reef concentration are the Great Barrier Reef of Australia, the

Red Sea and the Caribbean, the latter having a much lower diver-

sity than all major Indo-Pacific regions. The reefs and coral faunas

of each geological period are a combination of inheritance

from the past and the outcome of more contemporary events.

The distribution of modern reefs are an inheritance of the

Tethys-dominated world of the Eocene (Figure C68), subse-

quently modified by the obliteration of both the Tethys Seaway

(Miocene) and the Central American Seaway (Pliocene).

Reef environments

Calcareous algae thrive, but corals are uncommon on parts of

reefs exposed to extreme wave action. Where there is some

protection and where the water is clear, Indo-Pacific reef flats

are usually dominated by Acropora. Unconsolidated rubble

accumulations occur primarily where conditions for coralline

algae growth are poor: where wave action is weak, the water

turbid, or the temperature frequently falls below 18



C. Condi-

tions for coral growth, however, may be good in such places.

Environmental factors. Substrate availability, sedimentary

regimes, bathymetry, tidal regimes, turbulence, water quality,

nutrients and biotic factors all affect reef development by con-

trolling coral growth. These operate on local scales in shallow

tropical seas primarily by affecting substrate conditions and

light availability. On larger scales of space and time, sea levels

and temperature are the most important limiting environmental

parameters.

Sea level. Because coral reefs built by zooxanthellate corals

can exist only in very shallow water, they are continually at the

mercy of sea-level changes due to the land or sea rising or fall-

ing. Darwin recognized this long before the phenomenon of

changing sea levels per se was actually discovered, yet even

now it is not generally appreciated that coral reefs have not

always looked like they do today. In times past, sea-level

changes have been so rapid and so frequent that what are reefs

today would have alternated between being exposed limestone

islands and submerged limestone pinnacles. The reefs of today

are artifacts of an unusually long period of relatively stable

sea level.

Temperature. Reef-building corals and the reefs that they

build have similar, tropical, distributions. This creates an inter-

esting dilemma: which comes first, the coral or the reef ? There

have been many theories, all assuming some effect of ocean

temperature. There is no correlation between coral diversity

and reef formation: only a small number of coral species are

needed to build the most massive reefs. There is, however, a

clear correlation between ocean temperature and zooxanthellate

coral distribution, and between the former and formation of

highly consolidated reefs. The latter only occur where sea tem-

perature does not fall below 18



C for extended periods of time,

a restriction that does not apply to those corals where approxi-

mately half of all species occur where temperatures regularly

fall to 14



C (Veron, 1995). The likely solution to the dilemma

is that temperature controls the formation of reefs through eco-

logical processes. Reef building requires carbonate production

and that requires coral-dominated ecosystems. In the tropics,

coral ecosystems exist because they are able to out-compete

algae-dominated ecosystems, but this is not the case in higher

Figure C68 The distribution of coral reefs of the Eocene. (After Veron, 2000.)

CORALS AND CORAL REEFS 205

latitudes. Thus, the highest latitude coral reefs in the world all

have mixtures of coral and algal communities, which compete,

and the corals struggle to get enough light and space to survive.

Symbiosis

Reefs of today are an ancient phenomenon without parallel in

nature – a specific example of a symbiotic relationship between

plants and animals. The sort of plants and animals is not critical,

but the relationship between them is. Symbiosis allows the limit-

less resources of sunlight and carbon dioxide to be harnessed and

used to construct a place to live. In effect, symbiosis allows reefs

to be built by animals because it gives them the energy-generat-

ing capacity of plants. What the animal gives the plant is a med-

ium in which to live that is stable yet exposed to sunlight. What

the plant gives the animal is an enhanced capacity to remove

metabolic waste, an enhanced capacity to calcify, a direct nutrient

source and a capacity to concentrate and re-cycle limited nutri-

ents, including nitrogen and phosphate. Symbiosis allows corals

to exist in an almost nutrient-free environment, including the

nutrient deserts of the emptiest oceans.

Clearly, the evolutionary advantage of symbiosis is very

great, but the evolutionary cost is great also. Zooxanthellate

corals are constrained to live near the ocean surface, the most

hostile of marine environments both physically and biologi-

cally. Most importantly, symbiosis constrains corals to live in

places where they must compete with macroalgae. Coral-algae

symbiosis is therefore ultimately responsible for the geographic

constraints of reefs as well as their construction.

When did symbiosis evolve? This question has been

debated almost since coral/algal symbiosis was first discov-

ered, and the debate continues. It is tempting to conclude that

the building and maintenance of wave-resistant carbonate plat-

forms requires sunlight as the primary energy source. If so,

such platforms, found as reefs worldwide, especially in the

Silurian, Devonian, Jurassic and Cenozoic, could only have

been built by calcifying plants or by plant-animal symbioses.

This is not true for the reefs of most other periods or those

which were not wave resistant or which were not exposed to

high-energy conditions. Nor can it be supposed that symbioses

or capacity for skeletonization are irreversible conditions.

Indeed, it is quite possible that corals are anemone-like organ-

isms of many unrelated groups that can, over geological time

intervals of environmental change, develop or lose the capacity

for symbiosis concurrently or separately from the capacity for

skeletonization.

John E. N. Veron

Bibliography

Beauvais, L., 1989. Jurassic corals from the circum-Pacific area. Mem.

Assoc. Australasian Paleontol. 8, 291–302.

Beauvais, L., 1992. Paleobiogeography of the Early Cretaceous corals.

Palaeogeogr. Palaeoclimatol. Palaeoecol., 92, 233–247.

Buddemeier, R.W., and Fautin, D.G., 1997. Saturation state and the evo-

lution and biogeography of symbiotic calcification. In International

Symposium on Biomineralization, Monaco, (Alleman, D., and Cuif,

J.-P. eds.), pp. 23–32.

Ezaki, Y., 1998. Paleozoic Scleractinia: Progenitors or extinct experiments?

Paleobiology, 24(2), 227–234.

Frost, S.H., 1981. Oligocene reef coral biofaces of the Vicentin, northeast

Italy. Soc. Econ. Paleontol. Mineral. Spec. Publ., 30, 483–539.

Jell, P.A., and Jell, J.S., 1976. Early middle Cambrian corals from western

New South Wales. Alcheringa, 1, 181–195.

Kauffman, E.G., and Fagerstrom, J.A., 1993. The Phanerozoic evolution of

reef biodiversity. In Ricklefs, R.E., and Schluter, D. (eds.),

Species Diversity in Ecological Communities. Chicago, IL: University

of Chicago Press, pp. 314–329.

Oliver, W.A., 1980. The relationship of the scleractinian corals to the

rugose corals. Paleobiology, 6(2), 146–160.

Perrin, C., 2002. Tertiary: The emergence of modern reef ecosystems.

Tulsa, OK: Society of Economic Paleontologists and Mineralogists.

SEPM Special Publication 72, pp. 587–621.

Romano, S.L., and Cairns, S.D., 2000. Molecular phylogenetic hypotheses

for the evolution of scleractinian corals. Bull. Mar. Sci., 67(3),

1043–1068.

Stanley, G.D., 2003. The evolution of modern corals and their evolutionary

history. Earth-Sci. Rev., 60, 194–225.

Stanley, G.D., and Fautin, D.G., 2001. The origins of modern corals.

Science, 291, 1913–1914.

Stolarski, J., and Roniewicz, E., 2001. Towards a new synthesis of evolu-

tionary relationships and classification of the Scleractinia. J. Paleontol.,

75(6), 1090–1108.

Veron, J.E.N., 1995. Corals in Space and Time: The Biogeography and

Evolution of the Scleractinia. Sydney, Australia: University of New

South Wales Press, 321pp.

Veron, J.E.N., 2000. Corals of the World (3 vols.). Townsville, Australia:

Australian Institute of Marine Science, 463, 429 and 489pp.

Veron, J.E.N., Odorico, D.M., Chen, C.A., and Miller, D.J., 1996.

Re-assessing evolutionary relationships of scleractinian corals. Coral

Reefs, 15,1–9.

Wood, R., 1999. Reef Evolution. Oxford: Oxford University Press, 414pp.

Cross-references

Cenozoic Climate Change

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

Early Paleozoic Climates (Cambrian-Devonian)

Late Paleozoic Paleoclimates (Carboniferous-Permian)

Mesozoic Climates

Neogene Climates

Ocean Anoxic Events

Plate Tectonics and Climate Change

CORDILLERAN ICE SHEET

The Cordillera of western Canada was repeatedly enveloped by

a continental ice sheet, known as the Cordilleran Ice Sheet, dur-

ing the Pleistocene and latest Pliocene (Flint, 1971; Clague,

1989; Jackson and Clague, 1991). At its maximum extent, the

Cordilleran Ice Sheet and its satellite glaciers covered almost

all of British Columbia, southern Yukon Territory, and southern

Alaska, and extended south into the northwestern conterminous

United States (Figure C69 and C70). The ice sheet, to a consid-

erable extent, was confined between the high mountain ranges

bordering the Canadian Cordillera on the west and east, but

large areas on the east flank of the Rocky Mountains and west

of the Coast Mountains were also covered by ice. Glaciers in

several mountain ranges, such as the Olympic, Cascade, and

Mackenzie Mountains, were more or less independent of the

ice sheet, even at times of maximum ice cover.

The Cordilleran Ice Sheet attained its maximum size in

British Columbia where it was up to 900 km wide and reached

to 2,000–3,000 m elevation over the plateaus of the interior

(Wilson et al., 1958). When fully formed, the ice sheet prob-

ably had the shape of an elongate dish, with gentle slopes in

the interior region and steeper slopes at the periphery. It closely

resembled the present-day Greenland Ice Sheet at such times.

More commonly, the interior of the ice sheet had an irregular,

206 CORDILLERAN ICE SHEET

undulating surface, with several ice divides that shifted through

time. These ice divides were subordinate to the main divide

along the axis of the Coast Mountains.

In western British Columbia, ice streamed down fjords and

valleys in the coastal mountains and covered large areas of

the Pacific continental shelf. Some lobes extended to the shelf

edge where they calved into deep water. Glaciers issuing from

the southern Coast Mountains and Vancouver Island Ranges

coalesced over the Strait of Georgia to produce a great outlet

glacier that flowed far into Puget Lowland in Washington State

(Waitt and Thorson, 1983). Glaciers streaming down valleys

farther to the east likewise terminated as large lobes in eastern

Washington, Idaho, and Montana (Waitt and Thorson, 1983).

Glaciers in southern Alaska were connected to the Cordilleran

Ice Sheet and are considered by most researchers to be part of that

ice sheet (Hamilton and Thorson, 1983). They flowed from

mountain source areas across the continental shelf to the south

and into the broad low country drained by the Yukon River to

the north. Most of interior Alaska and northern Yukon Territory

(Beringia) were extremely dry throughout the Pleistocene and

consequently remained unglaciated. The Cordilleran Ice Sheet

in southern Yukon was fed principally from the Selwyn and Cas-

siar Mountains, although some ice was supplied from the St.

Elias Mountains (Jackson et al., 1991).

To the east, ice flowed out of the Rocky Mountains and

locally coalesced with ice originating in the Keewatin sector

of the Laurentide Ice Sheet (Clague, 1989). Some valley gla-

ciers in the eastern Mackenzie Mountains also came into con-

tact with Laurentide ice (Duk-Rodkin and Hughes, 1991).

These ice masses coalesced only rarely, at times of maximum

glaciation, most recently about 20,000 years ago. At other

times, an ice-free zone (the “ice-free corridor”) existed between

the Cordilleran and Laurentide Ice Sheets.

Growth and decay of Cordilleran Ice Sheet

The Cordilleran Ice Sheet nucleated in the high mountains of

British Columbia, Yukon, and Alaska (Figure C71a and

C72a). Small mountain ice fields grew and valley glaciers

advanced when climate deteriorated early during each glacia-

tion (Figure C71b; Kerr, 1934; Davis and Mathews, 1944).

With continued cooling and an increase in precipitation, gla-

ciers expanded and coalesced to form a more extensive cover

of ice in mountains (Figure C72). The glaciers advanced

out of the m ountains and across plateaus and lowlands

(Figure C72c), eventually coalescing to form an ice sheet that

covered most of British Columbia and adjacent areas

(Figure C71c and C72d). During this period, which spanned

thousands of years, the major mountain ranges remained the

principal sources of ice, and ice flow was controlled by topo-

graphy. Ice thickened to such an extent during the final phase

of glaciation that one or more domes became established over

the interior of British Columbia, with surface flow radially

away from their centers. This full continental ice sheet phase

of glaciation was rarely achieved. The transition into this final

phase was accompanied by a local reversal of ice flow in the

Figure C69 The Cordilleran Ice Sheet about 17,000 years ago at

the maximum of the last glaciation. The upper surface of the ice

sheet reached up to 2,000–3,000 m a.s.l. Arrows indicate directions

of ice flow.

Figure C70 The Homathko Icefield in the southern Coast Mountains of

British Columbia. Most of the Cordillera of western Canada looked

something like this 16,000 years ago (photo by J. J. Clague).

CORDILLERAN ICE SHEET 207

Coast Mountains as the ice divide shifted eastward, from the

mountain crest to a position over the British Columbia interior.

A comparable westward shift and reversal of flow may also

have occurred locally in the Rocky Mountains. The flow rever-

sals resulted from the buildup of ice in the interior to

levels higher than the main accumulation areas in the flanking

mountains.

The model outlined in the preceding paragraph provides

a framework for conceptualizing the growth of the Cordilleran

Ice Sheet, but the actual history of the ice sheet is more

Figure C71 Schematic diagram showing growth and decay of the Cordilleran Ice Sheet. (a) Mountain area at the beginning of a glaciation. (b)

Development of a network of valley glaciers. (c) Coalescence of valley and piedmont lobes to form an ice sheet. (d) Decay of ice sheet by

downwasting; upland areas are deglaciated before adjacent valleys. (e) Residual dead ice masses confined to valleys. (Modified from Clague, 1989,

figure 1.13).

208 CORDILLERAN ICE SHEET

complicated (Clague, 1989). Ice did not build up in a uniform,

monotonic fashion; rather, periods of growth were interrupted

by intervals during which glaciers stabilized or receded.

Most glacial cycles terminated with rapid climate warming.

Deglaciation was characterized by complex frontal retreat in

peripheral glaciated areas and by downwasting accompanied

by widespread stagnation throughout the Cordilleran interior.

The western periphery of the ice sheet became unstable due,

in part, to the global rise in sea level that occurred at such

times. The British Columbia continental shelf was rapidly freed

of ice as glaciers calved back to fjord heads and valleys. Fron-

tal retreat also occurred elsewhere along the periphery of the

ice sheet, for example in northern Washington and southern

Yukon.

A different style of deglaciation has been documented for

areas of low and moderate relief nearer the center of the ice

Figure C72 Schematic maps of ice cover in British Columbia during the growth phase of the last glaciation, about (a) 35,000, (b) 30,000, (c) 25,000,

and (d) 18,000 years ago. Ice distributions are inferred from limited data and should only be considered approximations. (Modified from Clague

et al., 2004, figure 4).

CORDILLERAN ICE SHEET 209

sheet. Deglaciation in these areas occurred mainly by down-

wasting and stagnation and proceeded through four stages

(Fulton, 1967): (a) active ice phase – regional flow continued

but diminished as ice thinned (Figure C71d); (b) transitional

upland phase – the highest uplands became ice free, but regio-

nal flow continued in valleys; (c) stagnant ice phase – ice was

confined to valleys but was still thick enough to flow; and

(d) dead ice phase – ice tongues in valleys thinned to the point

that they no longer flowed ( Figure C71e).

The first areas to become ice-free were those near the per-

iphery of the ice sheet, for example British Columbia’s conti-

nental shelf and the plateaus of the northwestern United

States (Clague, 1989). Active glaciers probably persisted long-

est in high mountain valleys, but they may have coexisted with

large masses of dead ice on the plateaus of the Cordilleran

interior. In general, retreat in the interior proceeded from both

southern and northern peripheral areas towards the center of

the ice sheet. In detail, however, the pattern of retreat was com-

plex, with uplands in each region becoming ice-free before

adjacent valleys.

Decay of the Cordilleran Ice Sheet at the close of each glacia-

tion was interrupted repeatedly by glacier readvances (Clague,

1989). Most readvances affected relatively small areas and may

not have been synchronous from one region to another.

Glacial erosion and deposition

The ice sheet and the alpine glaciers from which it formed

modified the late Tertiary landscape of British Columbia and

Yukon Territory. Mountain areas are dominated by erosional

glacial landforms, whereas plateaus, coastal lowlands, and

intermontane valleys record both the erosional and depositional

effects of glaciers. In high mountains, classic alpine forms

were created, including cirques and overdeepened valley heads,

horns, and comb ridges. Most mountain valleys are typical gla-

cial troughs. Some valleys in the westernmost Cordillera

extend into fjords, which attain water depths of up to 750 m.

Much of the sediment produced by glacial erosion was tr ans-

ported beyond the periphery of the ice sheet. Considerable

sediment, however, was deposited in valleys, on coastal low-

lands, and on the plateaus of the Cordilleran interior as pro-

glacial and ice-contact sediments, mainly during the advance

and recessional phases of the last glaciation. Deposits of older

glaciations are less common, because they have been exten-

sively eroded by the last Cordilleran Ice Sheet. Even in areas

where these older deposits are present, they are covered by

younger sediments and, consequently, are poorly exposed.

Ice sheet glaciation and sea-level change

Growth and decay of the Cordilleran Ice Sheet triggered crustal

movements that were dominantly isostatic in origin (Clague,

1983; Clague and James, 2002). The crust was displaced down-

ward during periods of ice-sheet growth. Initially, the depres-

sion was localized beneath the mountain ranges that were loci

of glacier growth. The area of crustal subsidence grew larger

as glaciers advanced out of mountains and into lowlands. At

times of maximum ice cover, the entire area of the ice sheet

was displaced downward.

The amount of isostatic depression during times of maxi-

mum ice cover depended primarily on the thickness and extent

of the ice sheet, the length of time over which it formed, and

the structure and composition of the crust and mantle (James

et al., 2000). Isostatic depression was greatest beneath the cen-

ter of the ice sheet and decreased west of the Coast Mountains

and Strait of Georgia toward the continental margin, and south

into Washington State.

Elevated glaciomarine sediments and shoreline features

along the British Columbia and northern Washington coasts

provide evidence for isostatic depression at the end of the last

glaciation (Clague et al., 1982a; Thorson, 1989). The elevation

of the late-glacial marine limit differs in relation to distance

from former centers of ice accumulation and time of deglacia-

tion. In general, the marine limit is highest (ca. 200 m a.s.l.)

on the British Columbia mainland coast and drops towards

the west, southwest, and south. Many mainland fjords, how-

ever, were deglaciated after much of the local isostatic rebound

had occurred; consequently, the marine limit in those areas is

relatively low. Late-glacial shorelines on Haida Gwaii (Queen

Charlotte Islands) were lower than at present, indicating that

glacio-isostatic depression was less than the coeval global

(eustatic) lowering of sea level (Clague et al., 1982b; Clague,

1983; Josenhans et al., 1997; Barrie and Conway, 2002).

Rapid deglaciation at the end of each glaciation triggered

isostatic adjustments that were opposite in direction to those

that occurred during ice-sheet growth (Figure C73; Clague,

1983; Clague and James, 2002). Material moved laterally in

the mantle, from extraglacial regions towards the center of the

decaying ice sheet. Areas at the periphery of the ice sheet, which

Figure C73 Generalized patterns of sea-level change on the British Columbia coast since the end of the last glaciation. Deglaciation and isostatic

rebound occurred later in the Coast Mountains than on Vancouver Island. (Modified from Muhs et al., 1987, figure 10).

210 CORDILLERAN ICE SHEET

were deglaciated earliest, rebounded first. The total amount

of uplift in these areas, however, was less than at the center of

the ice sheet where ice thicknesses generally were greater. As

deglaciation progressed, the zone of rapid isostatic uplift

migrated in step with receding glacier margins (Figure C73;

Clague, 1983). The rate of uplift in each region decreased

exponentially with time, and rebound was largely complete

within several thousand years of deglaciation (Clague et al.,

1982a; James et al., 2000).

John J. Clague

Bibliography

Barrie, J.V., and Conway, K.W., 2002. Rapid sea-level change and coastal

evolution on the Pacific margin of Canada. Sediment. Geol., 150,

171–183.

Clague, J.J., 1983. Glacio-isostatic effects of the Cordilleran Ice Sheet,

British Columbia, Canada. In Smith, D.E., and Dawson, A.G. (eds.),

Shorelines and Isostasy. London: Academic Press, pp. 321–343.

Clague, J.J., 1989. Cordilleran Ice Sheet, chap. 1. In Fulton, R.J. (ed.),

Quaternary Geology of Canada and Greenland. Ottawa, ON: Geologi-

cal Survey of Canada, pp. 40–43.

Clague, J.J., and James, T.S., 2002. History and isostatic effects of the last

ice sheet in southern British Columbia. Quaternary Sci. Rev., 21,71–87.

Clague, J., Harper, J.R., Hebda, R.J., and Howes, D.E., 1982a. Late Qua-

ternary sea levels and crustal movements, coastal British Columbia.

Can. J. Earth Sci., 19, 597–618.

Clague, J.J., Mathewes, R.W., and Warner, B.G., 1982b. Late Quaternary

geology of eastern Graham Island, Queen Charlotte Islands, British

Columbia. Can. J. Earth Sci., 19, 1786–1795.

Clague, J.J., Mathewes, R.W., and Ager, T.A., 2004. Environments

of northwestern North America before the last glacial maximum. In

Madsen, D. (ed.), Entering America: Northeast Asia and Beringia

before the Last Glacial Maximum. Salt Lake City, UT: University of

Utah Press, pp. 63–94.

Davis, N.F.G., and Mathews, W.H., 1944. Four phases of glaciation with

illustrations from southwestern British Columbia. J. Geol., 52, 403–413.

Duk-Rodkin, A., and Hughes, O.L., 1991. Age relationships of Laurentide

and montane glaciations, Mackenzie Mountains, Northwest Territories.

Géogr. physique Quaternaire, 45,79–90.

Flint, R.F., 1971. Glacial and Quaternary Geology. New York, NY: Wiley,

892pp.

Fulton, R.J., 1967. Deglaciation Studies in Kamloops Region, An Area of

Moderate Relief, British Columbia. Ottawa, ON: Geological Survey of

Canada, 36pp.

Hamilton, T.D., and Thorson, R.M., 1983. The Cordilleran Ice Sheet

in Alaska. In Porter, S.C. (ed.), Late-Quaternary Environments of

the United States, Vol.1: The Late Pleistocene. Minneapolis, MN:

University of Minnesota Press, pp. 38–52.

Jackson, L.E., Jr., and Clague, J.J. (eds.), 1991. The Cordilleran Ice Sheet.

Géogr. physique et Quaternaire, 45, 261–377.

Jackson, L.E., Jr., Ward, B.C., Duk-Rodkin, A., and Hughes, O.L., 1991.

The last Cordilleran Ice Sheet in southern Yukon Territory. Géogr. phy-

sique et Quaternaire, 45, 341–

354.

James, T.S., Clague, J.J., Wang, K., and Hutchinson, I., 2000. Postglacial

rebound at the northern Cascadia subduction zone. Quaternary Sci.

Rev., 19, 1527–1541.

Josenhans, H., Fedje, D., Pienitz, R., and Southon, J., 1997. Early humans

and rapidly changing Holocene sea levels in the Queen Charlotte

Islands-Hecate Strait, British Columbia, Canada. Science, 277,71–74.

Kerr, F.A., 1934. Glaciation in northern British Columbia. Roy. Soc. Can.,

Trans, 28(4), ser. 3, 17–31.

Cross-references

Glacial geomorphology

Glacial isostasy

Glaciations, Quaternary

Laurentide Ice Sheet

Quaternary climate transitions and cycles

Sea level change, Quaternary

COSMOGENIC RADIONUCLIDES

In contrast to stable isotopes, cosmogenic isotopes or radio-

nuclides are continuously produced and destroyed by nuclear

reactions. The fact that their sources and sinks (the radioactive

decay) are well known makes them extremely useful tools in

environmental sciences.

Sources and sinks

The main source of cosmogenic radionuclides is nuclear inter-

actions caused by galactic cosmic rays. Solar cosmic rays are

too low in energy to contribute significantly to the total produc-

tion rate. Production induced by decay or fission of uranium or

thorium, for instance, is only important under special circum-

stances (for example, underground).

The galactic cosmic rays are composed of protons (91%),

helium nuclei (8%), and heavier nuclei (1%). Coming from

space, cosmic rays first interact with the atmosphere. When

they collide with nitrogen, oxygen or argon, a cascade of sec-

ondary particles is produced. Among them are mainly neutrons

and protons, which interact with atmospheric nuclei producing a

variety of cosmogenic nuclides with masses equal to or smaller

than the target nucleus. The majority of these nuclides is

unstable and they decay within a very short time into stable

nuclides. Only few cosmogenic nuclides live long enough to

be detected and useful for applications (Table C8).

Cosmogenic nuclides are detected using decay counting for

short-lived nuclides or accelerator mass spectrometry (AMS)

techniques for long-lived nuclides. The detection limits are on

the order of a million atoms.

Before reaching the atmosphere, the cosmic rays have to

cross the heliosphere, which expands to about 100 astronomi-

cal units (1.5  10

km) and is filled with solar plasma of a

very low density. This plasma streams out from the Sun (so-

called solar wind) and carries solar magnetic fields, which

deflect especially low-energy galactic cosmic ray particles

and therefore reduce the production rate of cosmogenic

nuclides depending on the solar activity. In addition, the geo-

magnetic dipole field prevents cosmic ray particles with too

low a magnetic rigidity (momentum per charge) from penetrat-

ing into the atmosphere. This geomagnetic shielding effect is

largest at low latitudes, where the field lines are parallel to

the Earth’s surface, and negligible at high latitudes, where the

charged particles are guided into the atmosphere by the field

lines. The dependence of the production rate on atmospheric depth

(expressed in g cm

–2

) and latitude is shown in Figure C74a

using

Be as an example. As a result of these interactions,

the intensity of the galactic cosmic rays is reduced within the

atmosphere by almost three orders of magnitude. Therefore,

the production of cosmogenic nuclides within the lithosphere

is much smaller. However, due to a large number of heavier tar-

get nuclei, a variety of additional cosmogenic nuclides is pro-

duced in the lithosphere that cannot be produced in the

atmosphere.

Some of the nuclides also have anthropogenic sources.

During the nuclear bomb tests from the 1940s to 1970s, the

atmospheric amount of several nuclides was considerably

increased:

C by a factor of 2 and

H and

Cl by almost three

orders of magnitude. Additional sources are nuclear power and

reprocessing plants.

COSMOGENIC RADIONUCLIDES 211

Figure C74 (a) Dependence of the

Be production rate on the atmospheric depth (0 g cm

–2

: top of the atmosphere; 1,033 g cm

–2

: sea level) and

the latitude (Masarik and Beer, 1999). (b) Dependence of the mean global

Be production rate on the solar activity (phi = 0, quiet sun; phi = 1000,

very active sun) and the geomagnetic field intensity (relative units, 1= present field) (Masarik and Beer, 1999). The dynamic range between high

solar activity/large magnetic field and quiet sun/no field is about an order of magnitude.

Table C8 List of some of the main cosmogenic nuclides with half-lives longer than 1 month that are produced by cosmic rays in the atmosphere

(after Masarik and Beer, 1999)

Nuclide Half-life Target Reaction Production rate [cm

–2

–1

] Decay (keV) Detection

H 12.34 yr N, O Spallation 0.28 b

–

(0.02) Decay

Be 53.4 d N, O Spallation 0.035 e (487) Decay

Be 1.5 706 yr N, O Spallation 0.018 b

–

(0.6) AMS

C 5,730 yr N

N(n, p)

C 2.0 b

–

(0.2) AMS/ Decay

Al 740,000 yr Ar Spallation 7  10

–5

* b

(5.1) AMS

Si 150* yr Ar Spallation b

–

(0.2) Decay/ AMS

Cl 301,000 yr Ar Spallation 8.8  10

–4

* b

–

(0.7) AMS

*Indicate values with a large uncertainty.

Figure C75 Deviations of the atmospheric

C ratio from the value of 1950 as a result of changes in the production rate caused by

geomagnetic (long-term trend) and solar (short-term fluctuations) modulation of cosmic rays. (Modified from Stuiver et al., 1998; Stuiver and

Braziunas, 1993.)

212 COSMOGENIC RADIONUCLIDES

Applications

The two main fields of application of cosmogenic radionuclides

are dating and tracing. In the case of dating, decay dating and

exposure dating can be distinguished.

Dating

Decay dating makes use of the law of radioactive decay:

AtðÞ¼A

 exp 

ln 2

1=2

 t



, with

A(t) = number of radioactive atoms, their concentration or

their isotopic ratio at time t

= A at the time zero

1/2

= half-life

If A

is constant or known and A(t) is measured, the time t

can be calculated. The crucial point is A

. By radiocarbon

dating of tree-rings with known ages, the deviation of A

(

C ratio) from A

of 1950 could be determined for

the past 12,000 years. These deviations in %, called ▵

are plotted in Figure C75. The curve shows a long-term trend

related to changes in geomagnetic field intensity and super-

imposed short-term fluctuations caused by changes in solar

activity Figure C74b, such as the well known Maunder,

Spoerer and Wolf minima (see insert in Figure C75). The curve

of Figure C75 provides the basis to convert radiocarbon ages

into calendar ages.

In contrast to decay dating, the exposure dating technique is

based on the increasing number of cosmogenic nuclides that are

produced in matter (e.g., rocks) previously shielded, for example,

by a glacier. If the production rate is known, the exposure time

can be calculated from the measured total number of produced

atoms (Lal, 1988).

Tracing

In the case of tracer applications, cosmogenic nuclides can

be used according to their specific properties to study a large

variety of different physical, chemical and biological processes.

They are particularly useful in investigating the behavior of

their stable counterparts, such as, for example, the carbon

cycle. Examples of

Be applications are given in the article

Be.

Jürg Beer

Bibliography

Lal, D., 1988. In situ-produced cosmogenic isotopes in terrestrial rocks.

Annu. Rev. Earth Planet Sci. Lett., 16, 355–388.

Masarik, J., and Beer, J., 1999. Simulation of particle fluxes and cosmo-

genic nuclide production in the Earth’s atmosphere. J. Geophys. Res.,

104(D10), 12099–12111.

Stuiver, M., and Braziunas, T.F., 1993. Sun, Ocean, Climate and

Atmospheric

, an evaluation of causal and spectral relationships.

The Holocene, 3, 289–305.

Stuiver, M., et al., 1998. INTCAL98 radiocarbon age calibration,

24,000–0 cal

BP. Radiocarbon, 40(3), 1041–1083.

Cross-references

Beryllium-10

Dating, Radiometric Methods

Maunder Minimum

Radiocarbon Dating

Sun-Climate Connections

CRETACEOUS WARM CLIMATES

Introduction

The Cretaceous period spans approximately 70 million years of

Earth’s history, from 135 to 65.4 million years ago. Warm

intervals of the Cretaceous account for almost half the duration

of the period. A cool temperate climate in the Early Cretaceous

was followed by a warm to extremely warm Middle and Late

Cretaceous, and the warmth persisted into the Paleogene.

The Cretaceous and Paleogene greenhouse interval was longer

in duration than the ensuing icehouse intervals of the Neogene.

The termination of the Cretaceous is coincident with a major

mass extinction event in Earth’s history and a meteorite or

bolide impact at the Cretaceous-Tertiary (K/T) boundary.

In describing the world during the Cretaceous period, geo-

chemical model results indicate that atmospheric CO

levels

were at least 2–10 times higher than present day values

(Berner, 1994). The lack of permanent ice in the polar regions

resulted in a low equator to pole thermal gradient. Processes

that distributed the warmth from the low to high latitudes were

the oceans and the atmosphere, but the amount of heat appor-

tioned to each process is debated for these warm intervals

(Barron et al., 1993). The driver of circulation lies in the tropi-

cal world, at least for atmospheric and surface ocean currents.

Subsurface circulation of the Late Cretaceous Campanian

oceans was probably initiated from high latitudes (DeConto

et al., 2000a).

The warm intervals of the Cretaceous are known for the

abundance of unique marine bivalve groups such as the ubiqui-

tous inoceramids and the reef-building rudists. Major terrestrial

floral groups such as the gymnosperms persisted during the

warm climates, and angiosperms flourished during the Late

Cretaceous. Dinosaurs of various groups, such as ceratopsians

hadrosaurs, tyrannosaurs, and ornithopods, roamed the land.

The techniques used to infer Cretaceous climates have evolved

over the decades. Traditional tools applied the distribution, abun-

dance, and ecological associations of biota toward climatic inter-

pretations. In particular, the geographic dispersion patterns of

bivalves, gastropods, and foraminifera allowed for the original

establishment of Cretaceous climate zones acrosstheglobe, as bio-

tas were designated the most sensitive indicators of paleoclimatic

conditions. Research directions then shifted to include stable iso-

topes to infer temperatures for the marine realm. Analyses focused

on isotopes of carbon and oxygen, first for whole rock, later for

ecological associations of foraminifers and inoceramid bivalves,

and eventually for individuals from a specific species group. Sin-

gle crystals of unaltered calcite, as preserved in the inoceramid

bivalves, or original shell material containing calcite or aragonite

provided the most reliable database from which thermal interpreta-

tions were developed. Currently, organic geochemical techniques

such as TEX

(see below) are applied toward the investigation

of Cretaceous warmth in the marine realm (see also: Geochemical

proxies (non-isotopic); Organic geochemical proxies). Research

on the terrestrial domain has not kept pace with that of the oceanic

world, but land-ocean correlations and interpretations of interac-

tions have accelerated in recent years.

Concurrent with the utilization of empirical and analytical

data toward climate sensors was the development of the numerical

climate model, three-dimensional General Circulation Models

(GCMs) simulating atmospheric and/or oceanic conditions. Mod-

els contributed to understanding mechanisms of the atmosphere

CRETACEOUS WARM CLIMATES 213

and oceans by providing sensitivity experiments encompassing

the entire globe. Advances in paleoclimate modeling, and limits

to the models as applied to paleoclimate studies, are summarized

in DeConto et al. (2000b). Valdes (2000) provides a comprehen-

sive review of factors included in models for warm climates, and

feedback mechanisms. One of the greatest challenges for model-

ing warm Cretaceous climates is in keeping the polar regions free

of ice while not overheating the tropics. Ocean heat transport, high

atmospheric CO

levels, and forest vegetation at high latitudes are

the proposed mechanisms for modifying polar climates during

Cretaceous warm intervals (Rind and Chandler, 1991;Barron

et al., 1993; Otto-Bliesner and Upchurch, 1997).

Atmosphere

Extra-terrestrial and terrestrial sources influenced Cretaceous

climate. Extra-terrestrial climatic forcing is expressed in Cre-

taceous rocks across the globe (Fischer, 1993; Park et al., 1993;

Herbert et al., 1999). Cycles of precession, obliquity, and eccen-

tricity, indicating a Milankovitch-style expression of climatic

conditions, modulated the distribution of energy and the surface

temperature of the Earth. Cool-wet and warm-dry climatic condi-

tions controlled deposition of sediments during these cycles. Sili-

ciclastic input was higher than carbonate production during the

cool-wet phase, resulting in deposition of calcareous shale and

marlstone, and carbonate deposition dominated during the

warm-dry phase of the cycle, resulting in the deposition of lime-

stone and chalk.

Terrestrial forcing of climate occurred through sea floor

spreading and volcanism that released CO

into the atmosphere.

Although the warmth of the Cretaceous is attributed primarily to

atmospheric CO

levels higher than those of today, the exact levels

of CO

contributing to the warmth are under investigation. Model

experiments center on 2–4 times present-day values to match the

distribution of temperatures inferred from macrofaunal and

floral paleobiogeographic ranges. Continued production of CO

over millions of years, and retention of this greenhouse gas in

the atmosphere, set up the conditions for warmth that extended

from the middle through the Late Cretaceous stages.

Across the Cretaceous globe, the effects of global warmth

resulted in the disappearance of the North and South Polar to Sub-

polar climatic zones, and their replacement by North and South

Cold-Temperate zones (Figure C76). No permanent ice existed

at the poles (Frakes, 1979), although slush caps and the periodic

development of ice on the South Pole have been suggested. Cold

Temperate, Mild T emperate, and Warm Temperate zones exten-

ded from the polar region through southern Canada. Episodically

during highstands, subtropical conditions extended from approxi-

mately the Canadian-US border southward to the tropics. The tro-

pical zone existed in the southern parts of the Gulf Coast states

and north-central Mexico. A supertropical or Supertethyan zone,

characterized by unusually high temperatures and salinities,

emerged episodically within the tropics and included the islands

of Puerto Rico, Cuba, and Hispaniola, as well as southern Mexico,

northern Venezuela, and northern Columbia. These global climate

zones were originally determined through the geographic map-

ping of bivalve genera and subgenera (Kauffman, 1973; Johnson

and Kauffman, 1996).

In North America, atmospheric conditions were influenced

by the presence of the Western Interior Basin, a fully marine

seaway that extended from Arctic Canada and Alaska to the

Gulf of Mexico during peak sea level highstand. This large sea-

way serves as a model for equivalent bodies of water across the

globe. A relatively flat landmass bordered the eastern coastal

plain, and a mountain belt developed to the west. A westerly

Figure C76 Cretaceous climate zones for paleolatitudes across the globe. Cold, Mild, and Warm Temperate Zones replaced the Polar Zones in the

highest latitudes. Subtropical, Tropical, and Supertropical Zones expanded during the warmest stages of the Cretaceous, as indicated in text

(base map for 80 Myr from C. R. Scotese).

214 CRETACEOUS WARM CLIMATES