Odekon M. Encyclopedia of paleoclimatology and ancient environments
Подождите немного. Документ загружается.
period. The glaciation of Eurasia had an ice sheet centered over
the British Isles in the west and that extended beyond the present
day shoreline onto the then emerged shelf of the North Sea. The
Fennoscandian Ice Sheet spread into the same area from the east
but did not coalesce into the British Isles at the LGM. To the east,
the Fennoscandian Ice Sheet extended into the northern part of
the Russian plain and the Barents Sea, which in the northwest
was alsocovered by ice from Novaya Zemlya. The Kara Sea,
except at its western margin, was not glaciated at the LGM, but
likely at an earlier stage of the last glacial cycle. The thickness of
the ice sheets cannot be reconstructed from glacial-geological
observations but it has been inferred to have been up to 4 km
from indirect evidence such as uplift rates of the land, and
from modeling.
By comparison, the Antarctic and Greenland Ice Sheets under-
went only small changes in volume and extent during the last gla-
cial cycle. Volume changes of the Antarctic Ice Sheet were largely
concentrated in West Antarctica and the Antarctic Peninsula.
At the LGM, grounding lines advanced close to the continental
shelf break almost everywhere as global sea level dropped. It is
unclear whether extensive ice streams continued to exist in the
Ross basin and elsewhere. Ice over interior East Antarctica was
generally thinner than today because of lower accumulation rates.
The Greenland Ice Sheet also extended beyond the present coast-
line to cover at least the inner continental shelf. Summer melting
ceased and mass was lost by calving only. The Greenland Ice Sheet
extended northwest to join the Laurentide Ice Sheet flowing out of
Canada. Ice retreat of the Antarctic and Greenland Ice Sheets was
in full progress between 13,000 and 8,000 years ago, several thou-
sand years later than for the North American and European Ice
Sheets (Huybrechts, 2002).
To date, no criteria have been developed to recognize paleo-
ice shelves. It has been speculated that the Antarctic continent
was entirely encircled by a vast ice shelf that merged imper-
ceptibly with the permanent pack ice and possibly extended
as far as 800 km beyond the continental shelf (Denton and
Hughes, 1981). Russian scientists have hypothesized a 1 km
thick Arctic ice shelf from glacial scouring of ridges at the
sea floor, but its dating is uncertain (Polyak et al., 2001).
Small ice caps and glaciers merged to larger glacier complexes
in Patagonia, the Alps, and many presently glaciated areas else-
where (Pyrenees, Caucasus, eastern Siberia, New Zealand Alps,
and other mountain areas). There is no convincing evidence for
the occurrence of a Tibetan Ice Sheet, but glacier fronts in the
Himalayas descended by about 1,000 m elevation.
In the Southern Hemisphere, the most striking difference
was the winter extent of sea ice, which was significantly greater
than it is today (Cooke and Hays, 1982). Summer extent was
probably similar to today’s winter extent. Permanent sea-ice
cover also existed in the Norwegian, Greenland, and Labrador
Seas coincident with a large shift in the path of the Gulf
Stream, which flowed due east at 40
N during the glacial max-
imum. A striking and unexplained difference was the compar-
ably small amount of cooling in the North Pacific and the
associated southward extension of sea ice at LGM compared
with equivalent latitudes in the North Atlantic.
There are no proxy data to infer the seasonal variation of snow
cover and river and lake ice. Nevertheless, winter cover must
have been greatly expanded, commensurate with Northern
Hemisphere continental coolings of up to –20
C (CLIMAP Pro-
ject members, 1981; Frenzel et al., 1992). There are also many
fossil periglacial features at the surface that show that permafrost
at the LGM was more extensive than today. In North America,
discontinuous permafrost extended up to 1,000 km south of the
southern margin of the Laurentide Ice Sheet. Permafrost was also
present on most of Europe, the Atlas Mountains, Turkey, Iran, the
whole of Siberia, the Tibetan Plateau, northern China, and Hok-
kaido. In the Southern Hemisphere, widespread permafrost
extended at mid altitude in southern America, as far north as
40
S, and was present in Patagonia, the southern island of New
Zealand, and on all sub-Antarctic islands. It occurred isolated
in mountainous areas of Tasmania and southeast Australia, and
even in the highest mountains of East Africa and the high Andes.
It never existed in South Africa, with the exception of a small
spot of mountain permafrost in the Drakensberg (Van Vliet-
Lanoë and Lisitsyna, 2001).
The last glacial cycle was enclosed by two periods when the
cryosphere was smaller than today. During the Last Interglacial,
which culminated 125,000 years ago, Northern Hemisphere tem-
peratures at mid-latitudeswere2–3
C warmer than today, increas-
ing towards the pole to more than 4
C. The Greenland Ice Sheet
may have retracted to a small central ice dome, thus providing
most of the melt to explain the inferred sea-level stand of appro-
ximately 6 m higher than today (Letréguilly et al., 1991; Cuffey
and Marshall, 2000). The Holocene Climatic Optimum occurred
around 6,000 years ago. The warming was of a similar magnitude
as that of the Eemian warming, but it was shorter-lived. At
that time, the Greenland Ice Sheet margin receded inland from
its present position. The other components of the cryosphere,
except the Antarctic Ice Sheet, were also smaller.
The cryosphere before the last glacial cycle
In geological terms, the last ice age ended a short time ago. It
was the last episode of a period of climate oscillations between
glacial and relatively ice-free environments that began about
2.3 million years ago. Over the last 800,000 years, such ice
ages had a period of about 100,000 years; before that,
41,000-year cycles dominated. The penultimate glacial cycle,
which ended 130,000 years ago, was probably the most severe.
Periods of extensive glaciation also existed prior to those
of the Pleistocene epoch. Though the record of early climates
is dominated by evidence of global warmth and likely almost
total absence of land and sea ice, there were intervals of
widespread ice cover during Paleoproterozoic times, at around
2.3 billion years ago and perhaps also between 3.1 and
2.9 billion years ago.
The most extreme glacial climate in the Earth’s history
probably occurred during the Late Precambrian between 900
and 600 million years ago, and is known as the Proterozoic
ice house. In its extreme version, ice sheets extended to the
equator and the world’s ocean was virtually entirely frozen
over (Hoffman et al., 1998). At that time, the Earth experienced
anomalously low atmospheric carbon dioxide concentrations,
possibly as a result of mountain building and continental
weathering, in addition to the Sun being 6% fainter than it is
today. Recovery from such a snowball Earth could have
resulted from volcanic outgassing of greenhouse gases. Glacia-
tion of the Earth also occurred in the Late Ordovician at 450
million years ago, and during Carboniferous-Permian times
between 320 and 250 million years ago. At this time, a single
supercontinent, Gondwanaland, lay over the South Pole, which
suggests that the presence of a large land mass at high latitude
is a critical element for glaciation.
Glaciation on Earth started again during the Cenozoic Era,
from 65 million years ago to the present. It started first in the
CRYOSPHERE 225
Southern Hemisphere and was punctuated by two periods of
significant ice growth. The first ice increase occurred near the
Eocene-Oligocene boundary 36 million years ago and saw fre-
quent ice sheets growing and decaying on East Antarctica,
reaching the margin of the continent in a few places (Barker
et al., 1999). Global ice sheet volume increased again in the
middle Miocene around 15 million years ago with the forma-
tion of a quasi-permanent East Antarctic Ice Sheet that may
have reached its maximum extent. The first period of ice
increase was probably driven by declining levels of atmo-
spheric carbon dioxide, while the second period has been
linked to the thermal isolation of Antarctica, as newly formed
continents moved to the north and the Antarctic Circumpolar
Current developed. The West Antarctic Ice Sheet likely formed
later, during the Late Miocene. By 2 million years ago, the
Antarctic glacial regime was much as at present and ice ages
started to occur on the Northern Hemisphere, as did mountain
glaciers in midlatitude highlands.
Philippe Huybrechts
Bibliography
Arendt, A.A., Echelmeyer, K.A., Harrison, W.D., Lingle, C.S., and
Valentine, V.B., 2002. Rapid wastage of Alaska glaciers and their con-
tribution to rising sea levels. Science, 297, 382–386.
Barker, P.F., Barrett, P.J., Cooper, A.F.K., and Huybrechts, P., 1999.
Antarctic glacial history from numerical models and continental margin
sediments. Palaeogeogr., Palaeoclimatol., Palaeoecol., 150, 247–267.
Church, J.A., Gregory, J.M., Huybrechts, P., Kuhn, M., Lambeck, K.,
Nhuan, M.T., Qin, D., and Woodworth, P.L., 2001. Changes in sea
level. In Houghton, J.T., Ding, Y., Griggs, D.J., Noguer, M., Van der
Linden, P.J., Dai, X., Maskell, K., and Johnson, C.A. (eds.), Climate
Change 2001: The Scientific Basis: Contribution of Working Group I
to the Third Assessment Report of the Intergovernmental Panel on
Climate Change. Cambridge, UK/New York: Cambridge University
Press, pp. 639–694.
Clark, P.U., and Mix, A.C., 2002. Ice sheets and sea level of the Last
Glacial Maximum. Quaternary Sci. Rev., 21(1–3), 1–7.
CLIMAP, 1976. The surface of the ice-age earth. Science, 191, 1131–1137.
CLIMAP Project Members, 1981. Seasonal reconstructions of the Earth’s
surface at the Last Glacial Maximum (Geological Society of America
Map and Chart Series, MC-36). Boulder, CO: Geological Society of
America.
Comiso, J.C., 2002. A rapidly declining perennial sea ice cover in
the Arctic. Geophys. Res. Lett., 29(20), 1956, doi:10.1029/2002
GL015650.
Cooke, D.W., and Hays, J.D., 1982. Estimates of Antarctic ocean seasonal
sea-ice cover during glacial intervals. In Craddock, C. (ed.), Antarctic
Geoscience Madison: University of Wisconsin Press, pp. 1017–1025.
Cuffey, K.M., and Marshall, S.J., 2000. Substantial contribution to
sea-level rise during the last interglacial from the Greenland ice sheet.
Nature, 404, 591–594.
Denton, G.H., and Hughes, T.J. (eds.), 1981. The Last Great Ice Sheets.
New York: Wiley, 485pp.
Fitzharris, B.B., 1996. The cryosphere: Changes and their impacts. In
Watson, R.T., Zinyowera, M.C., and Moss, R.H. (eds.), Climate Change
1995: Impacts, Adaptations and Mitigations of Climate Change:
scientific-technical analysis. Cambridge, UK: Cambridge University
Press, pp. 241–266.
Frenzel, B., Pecsi, M., and Velichko, A.A., 1992. Atlas of palaeocli-
mates and palaeoenvironments of the Northern Hemisphere, Late
Pleistocene-Holocene. Stuttgart: Gustav Fisher Verlag, 153pp.
Hoffman, P.F., Kaufman, A.J., Halverson, G.P., and Schrag, D.P., 1998. A
Neoproterozoic Snowball Earth. Science, 281, 1342–1346.
Huybrechts, P., 2002. Sea-level changes at the LGM from ice-dynamic
reconstructions of the Greenland and Antarctic Ice Sheets during the
glacial cycles. Quaternary Sci. Rev., 21(1
–3), 203–231.
Huybrechts, P., Steinhage, D., Wilhelms, F., and Bamber, J.L., 2000.
Balance velocities and measured properties of the Antarctic Ice Sheet
from a new compilation of gridded datasets for modeling. Ann. Glaciol.,
30,52–60.
Letréguilly, A., Reeh, N., and Huybrechts, P., 1991. The Greenland Ice
Sheet through the last glacial-interglacial cycle. Global and Planetary
Change, 4(4), 385–394.
Lindsay, R.W., and Zhang, J., 2005. The thinning of arctic sea ice, 1988–
2003: Have we passed a tipping point? J. Climate, 18, 4879–4894.
Lythe, M., and Vaughan, D.G., 2001. BEDMAP: A new ice thickness
and subglacial topographic model of Antarctica. J. Geophys. Res., 106
(B6), 11335–11351.
Magnuson, J.J., and 13 co-authors. 2000. Historical trends in lake and river
ice cover in the Northern Hemisphere. Science, 289, 1743–1746.
Polyak, L., Edwards, M.H., Coakley, B.J., and Jakobsson, M., 2001. Ice
shelves in the Pleistocene Arctic Ocean inferred from glaciogenic
deep-sea bedforms. Nature, 410, 453–457.
Rignot, E.J., and Thomas, R.H., 2002. Mass balance of polar ice sheets.
Science, 297, 1502–1506.
Rignot, E.J., Rivera, A., and Casassa, G., 2003. Contributions of the Patago-
nia Icefields of South America to sea level rise. Science, 302,434–437.
Robinson, D.A., Dewey, K.F., and Heim, R.R., 1993. Global snow cover
monitoring: An update. Bull. Am. Meteorol. Soc., 81, 1281–1299.
Romanovsky, V.E., and Osterkamp, T.E., 2001. Permafrost: Changes
and impacts. In Paepe, R., and Melnikov, V. (eds.), Permafrost
Response on Economic Development, Environmental Security and
Natural Resources Dordrecht: Kluwer, pp. 297–315.
Serreze, M.C., Holland, M.M., and Stroeve, J., 2007. Perspectives on the
Arctic’s shrinking sea-ice cover. Science, 315, 1533–1536.
Untersteiner, N., 1984. The cryosphere. In Houghton, J.T. (ed.), The Global
Climate. Cambridge, UK: Cambridge University Press, pp. 121–
140.
Van Vliet-Lanoë, B., and Lisitsyna, O., 2001. Permafrost extent at the Last
Glacial Maximum and at the Holocene Optimum. The CLIMEX Map.
In Paepe, R., and Melnikov, V. (eds.), Permafrost Response on
Economic Development, Environmental Security and Natural Resources
(NATO Science Series, 76). Dordrecht: Kluwer, pp. 215–225.
Vaughan, D.G., Bamber, J.L., Giovinetto, M.B., Russell, J., and Cooper, A.P.R.,
1999. Reassessment of net surface mass balance in Antarctica. J. Climate,
12, 933–946.
Zhang, T., Barry, R.G., Knowles, K., Heginbottom, J.A., and Brown, J.,
1999. Statistics and characteristics of permafrost and ground-ice dis-
tribution in the Northern Hemisphere. Polar Geogr., 23(2), 132–154.
Cross-references
Antarctic Glaciation History
Borehole Climatology: Climate Change from Geothermal Data
CLIMAP
Cordilleran Ice Sheet
Glaciations, Pre-Quaternary
Glaciations, Quaternary
Icehouse (cold) Climates
Last Glacial Maximum
Laurentide Ice Sheet
Mountain Glaciers
Periglacial Geomorphology
Pleistocene Climates
Proterozoic Climates
Scandinavian Ice Sheet
CYCLIC SEDIMENTATION (CYCLOTHEM)
The term “cyclic sedimentation” is generic and can be applied
to any type or scale of repetitive sedimentation (Einsele et al.,
1991). Conversely, the term “cyclothem” has a much more
specific application. The concept has mostly been applied to
stratigraphic successions of middle to late Pennsylvanian
(upper Carboniferous) rocks that were deposited in cratonic
basins of the eastern and mid-continental USA. Well-developed
226 CYCLIC SEDIMENTATION (CYCLOTHEM)
cyclic sedimentation was first noted in Illinois Basin (Eastern
Interior Coal Basin), within late Paleozoic rocks in the
state of Illinois. These lithological cycles exhibited vertical
oscillations between coal bearing, siliciclastic-rich facies and
carbonate-rich facies (Udden, 1912;Weller,1930, 1931).
The term “cyclothem” was subsequently applied to this style
of cyclic sedime ntation (Wanless and Weller, 1932).
The siliciclastic-rich facies were traditionally viewed as terre-
strial or continental facies, however, tidal influences are now
widely recognized (Archer and Kvale, 1993;Archeretal.,
1994a,b). The carbonate-rich facies commonly contain marine
fossils. Thus, cyclothems were commonly interpreted as resulting
from changes in base level brought about by sea level fluctuations.
These changes could be related to basinal tectonics, or most com-
monly, to global glacioeustatic changes that are related to Gond-
wanaland paleoglaciations. More recently, climatic oscillations
and the resultant changes in rates of clastic flux into subsiding
basins have been advocated as a potential driving mechanism
(Cecil, 1990).
A recurring problem that haunts the application of cyclo-
them models is that there are a great number of different mod-
els at different scales. Reger (1931) employed an unusual and
extreme application of the cyclothem paradigm for cycles in
the Appalachian Basin. In sections that did not conform to
his idealized model, he insisted that any missing units would
be given stratigraphic names, but referred to as “phantoms”
since they did not occur at that specific locality.
Commonly, a cyclothem model, which was derived from a
specific basin and from a specific stratigraphic interval, has
subsequently been applied far from its original areas of suitability.
On the other hand, some general models, which do have a wide-
ranging applicability, are named such that they imply occurrence
only within a specific basin. For example, a common model is
the “Illinois-Basin cyclothem,” which was applied originally by
Wanless and Weller (1932) to coal-bearing sequences in middle
Pennsylvanian rocks of southern Illinois (see also Wanless
and Wright, 1978). Initially, the concept was thought to be a
mapable entity and was applied as a formation-level stratigraphic
entity. Another popular model is the “Kansas cyclothem,” which
was described from non-coal bearing, more marine cycles in
upper Pennsylvanian rocks of eastern Kansas (Moore, 1936,
1949, 1964). Similar to the application within the Illinois Basin,
the cyclothem was used commonly as a mapable, lithostrati-
graphic entity.
The use of geographic descriptors for the names of these dif-
ferent types of cyclothems has continued to create confusion.
For example, the “Illinois-Basin cyclothem” not only occurs
in the middle Pennsylvanian of Illinois, but is also the typical
sequence seen in age-equivalent rocks of Kansas. Similar, the
“Kansas cyclothem” is the typical sequence observed in upper
Pennsylvanian in eastern Kansas but also in age-equivalent
strata of the Illinois Basin. Thus, there is a far greater similarity
of the Pennsylvanian system in Illinois and Kansas than is
commonly supposed. Largely, these misunderstandings have
been perpetuated by the superficial application of the standard
cyclothem models.
Although the cyclothem concept was originally conceived
in the 1930s for sequences in Illinois, it was subsequently dis-
carded as a meaningful stratigraphic entity in the land of its ori-
gin. However, the concept was developed in far greater detail
for strata within the U.S. mid-continent (particularly within
Kansas and adjacent states). Over a period of decades, rang-
ing from the 1930s to the 1960s, increasingly more complex
models of cyclothems were discussed by Moore (1931, 1936,
1949, 1964). It is perhaps Moore
’s near obsession and his dom-
ineering role in mid-continental geology that did much to
entrench the concept of a “cyclothem” into the field of sedimen-
tary geology. At its most complex, Moore’s models included
cycles of cyclothems, which he termed a “megacyclothem.”
Later, Moore’s megacyclothem concept was modified and
simplified by Heckel (1977, 1986). This adapted model was
itself termed a “cyclothem”; thus, Moore’s earlier definition
and Heckel’s subsequent modification of the term referred to
different scales of stratigraphic sequences. This also increased
the confusion over the historical use of the term. The reinter-
pretation of Moore’s cyclothem concepts included a reinterpre-
tation of organic-rich facies as deep-water deposits and an
inferred and hypothetical depth zonation based upon conodont
distributions (Heckel and Baesemann, 1975). West et al. (1997)
presented a summary of the various cyclothem models and dis-
cussions on their potential eustatic and climatic controls.
Attempts have made to quantify cyclothems and to develop
a mathematical or statistical model. Such attempts have never
been successful (see Pearn, 1964; West et al., 1997). Argu-
ments have been proposed that a search for cyclicity in litholo-
gical successions is an inherent human bias (Zeller, 1964).
There are many drawbacks that exist with the traditional
usage of the cyclothem concept. The biggest limitation is that
most of the work has focused on the more regular and recurring
parts of the cycle, which consist of the more-marine, carbonate-
dominated portions of the sedimentary cycle. These units,
which also include organic-rich laminated black shales, are
the easiest to correlate, particularly when using subsurface,
geophysical logs. The more clastic components, which have
been referred to as clastic wedges (Wanless, 1964; Archer
and Kvale, 1993) of the cycle, are far more variable and are
much harder to adapt to simple cyclothem models. Subsequent
work on the clastic wedges suggests that these facies were
formed by deltaic and tidal-estuarine depositional systems
(Archer et al., 1994a). From a sequence stratigraphic perspec-
tive, these clastic wedges include incised valley-fill (IVF)
facies (Archer et al., 1994b).
Ongoing appreciation and application of sequence-strati-
graphic models (see Van Wagoner et al., 1990), particularly to
the clastic components of cyclothems, suggest that the cyclothem
concept is largely obsolete. However, because of the extensive
historical literature that has applied the various permutations
of cyclothem models, the concept still creates considerable confu-
sion in studies of cyclic sedimentation. Postdating the influence of
Moore, the subsequent compendia of formal stratigraphic units
of Kansas no longer attempted to apply the cyclothem concept
(Zeller, 1968).
Allen W. Archer
Bibliography
Archer, A.W., and Kvale, E.P., 1993. Origin of gray-shale lithofacies
(“clastic wedges”) in U.S. midcontinental coal measures (Pennsylva-
nian). In Cobb, J.C., and Cecil, B. (eds.), Modern and Ancient Coal-
Forming Environments. Boulder, CO: Geological Society of America.
286, pp. 181–192.
Archer, A.W., Feldman, H.R., Kvale, E.P., and Lanier, W.P., 1994a.
Pennsylvanian (Upper Carboniferous) fluvi o- to tidal-estuarine
coal-bearing systems: Delineation of facies transitions based upon
physical and biogenic structures. In Calder, J.H., and Gibling, M.R.
(eds.), Palaeogeography, Palaeoclimatology, and Palaeoec ology,
106,171–18 5.
CYCLIC SEDIMENTATION (CYCLOTHEM) 227
Archer, A.W., Lanier, W.P., and Feldman, H.R., 1994b. Stratigraphy
and depositional history within incised-paleovalley and related facies,
Douglas Group (Missourian/ Virgilian; Upper Carboniferous) of
Kansas, USA. In Dalrymple, R.W., Boyd, R., and Zaitlin, B. (eds.),
Incised-Valley Systems: Origin and Sedimentary Sequences. Tulsa,
OK: Society of Economic Paleontologists and Mineralogists. 51,
176–190.
Cecil, C.B., 1990. Paleoclimate controls on stratigraphic repetition of
chemical and siliciclastic rocks. Geology, 18, 533–536.
Einsele, G., Ricken, W., and Seilacher, A., 1991. Cycles and events in
stratigraphy—basic concepts and terms. In Einsele, G., Ricken, W.,
and Seilacher, A. (eds.), Cycles and Events in Stratigraphy. Berlin:
Springer-Verlag, pp. 1–19.
Heckel, P.H., 1977. Origin of phosphatic black shales in Pennsylvanian
cyclothems of mid-continent North America. Am. Assoc. Pet. Geol.
Bull., 61, 1045–1068.
Heckel, P.H., 1986. Sea-level curve for Pennsylvanian eustatic marine
transgressive-regressive depositional cycles along midcontinental
outcrop belt, North America. Geology, 14, 330–334.
Heckel, P.H., and Baesemann, J.F., 1975. Environmental interpretations
of conodont distribution in Upper Pennsylvanian (Missourian) megacy-
clothems in eastern Kansas. Am. Assoc. Pet. Geol. Bull., 59,
1058–1074.
Moore, R.C., 1931. Pennsylvanian Cycles in the Northern Mid-Continent
Regions. Illinois Geological Survey Bulletin 60, pp. 247–257.
Moore, R.C., 1936. Stratigraphic Classification of the Pennsylvanian
Rocks of Kansas. Kansas Geological Survey Bulletin 22, pp. 1–256.
Moore, R.C., 1949. Divisions of the Pennsylvanian System in Kansas.
Kansas Geological Survey Bulletin 83, pp. 1–203.
Moore, R.C., 1964. Paleoecological Aspects of Kansas Pennsylvanian
Rocks of Kansas. Kansas Geological Survey Bulletin 169, pp. 287–380.
Pearn, W.C., 1964. Finding the ideal cyclothem. In Merriam, D.F. (ed.),
Symposium on Cyclic Sedimentation. Kansas Geological Survey
Bulletin 169, pp. 399–413.
Reger, D.B., 1931. Pennsylvanian Cycles in West Virginia. Illinois State
Geological Survey Bulletin 60, pp. 217–
239.
Udden, J.A., 1912. Geology and Mineral Resources of the Peoria Quad-
rangle, Illinois. U.S. Geological Survey Bulletin 506, pp. 1–103.
Van Wagoner, J.C., Mitchum, R.M., Campion, K.M., and Rahmanian, V.D.,
1990. Siliciclastic Sequence Stratigraphy in Well Logs, Cores, and
Outcrops: Concepts for High-Resolution Correlation of Time and
Facies. Tulsa, OK: American Association of Petroleum Geologists. 7,
55pp.
Wanless, H.R., 1964. Local regional factors in Pennsylvanian cyclic
sedimentation. In Merriam, D.F. (ed.), Symposium on Cyclic Sedimen-
tation. Kansas Geological Survey Bulletin 169, pp. 593–606.
Wanless, H.R., and Weller, J.M., 1932. Correlation and extent of
Pennsylvanian cyclothems. Geol. Soc. Am. Bull., 169, 593 –606.
Wanless, H.R., and Wright, C.R., 1978. Paleoenvironmental Maps of
Pennsylvanian Rocks, Illinois Basin and Northern Mid-Continent
Region. Boulder, CO: Geological Society of America.
Weller, J.M., 1930. Cyclical sedimentation of the Pennsylvanian Period and
its significance. J. Geol., 38,97–135.
Weller, J.M., 1931. The Conception of Cyclical Sedimentation during
the Pennsylvanian Period. Illinois State Geological Survey Bulletin
60, pp. 163–177.
West, R.R., Archer, A.W., and Miller, K.B., 1997. The role of climate
in stratigraphic patterns exhibited by late Palaeozoic rocks exposed in
Kansas. Palaeogeogr. Palaeoclimatol., Palaeoecol., 128,1–16.
Zeller, E.J., 1964. Cycles and psychology. In Merriam, D.F. (ed.),
Symposium on Cyclic Sedimentation. Kansas Geological Survey
Bulletin 169, pp. 631–636.
Zeller, D.E., 1968. The Stratigraphic Succession in Kansas. Kansas
Geological Survey Bulletin 189, pp. 1–81.
Cross-references
Coal Beds, Origin and Climate
Continental Sediments
Dating, Biostratigraphic Methods
Deltaic Sediments, Climate Records
Glacial Eustasy
Late Paleozoic Paleoclimates (Carboniferous-Permian)
Varved Sediments
228 CYCLIC SEDIMENTATION (CYCLOTHEM)
D
DANSGAARD-OESCHGER CYCLES
Dansgaard-Oeschger (D-O) cycles are oscillations of the cli-
mate system during the Wisconsinan glacial, where the climate
switched between a cold glacial climate and a “warm” glacial
climate. The glacial climate, which ended around 11,500 years
ago, was colder than the present warm period (the Holocene)
and variations were of much greater amplitude than those
recorded during the Holocene.
D-O cycles are not regular sinusoidal cycles, but are charac-
terized by rapid warming events. Such events occurred 24 times
during the Wisconsinan glacial (115–10 ky
BP) (ky ¼10
3
years). The magnitude of the warming over Greenland is esti-
mated to be approximately 10–15
C based on data from
Greenland ice cores. However, these temperatures did not reach
the warm temperatures of the Holocene. A cycle is character-
ized by a rapid transition from the cold glacial climate to a war-
mer phase (which is also called a D-O event) within only
20–50 years. After this rapid warming, the warmer climate per-
sists for hundreds to thousands of years. A slow decline in tem-
peratures (recorded in the d
18
O values) occurs during D-O
events, until the climate finally “jumps” back to the colder gla-
cial climate (see Figures D1 and D2). The cold phase does not
show any characteristic behavior; it may be described as stable
with a noisy signal. This cyclical behavior is what gives the
D-O events their characteristic sawtooth shape.
Ocean sediment records show an oscillation of 1,500 years
(Bond and Lotti, 1995) during the Wisconsinan glacial period.
These cycles are considered to originate from the D-O cycles
(Bond and Lotti, 1995; Alley et al., 1999).
Occurrence and labeling of D-O cycles
The record of D-O cycles has been labeled by dividing each
cycle into a warm and a cold phase, where the warm phases
are named Greenland interstadials (GI) and the cold phases are
named Greenland stadials (GS) (Björk et al., 1998; Table D1
and Figure D1). Each GI/GS is associated with a number
(1, 2, 3...) and each numbered GI/GS is then subdivided by
alphabetic letters (a, b, c...), which can be divided further by
numbers associated with warm and cold phases. This numbering
system allows the incorporation of data from new sites if addi-
tional variations or oscillations are found within a specific
GI/GS. For example, GI-1 contains three warm periods, where
the two oldest are the well-known Allerød and Bølling. GS-1
is also called Younger Dryas. Older Dryas (GI-1d), however, is
the cold spell between the before-mentioned Allerød (GI-1c)
and Bølling (GI-1e) in GI-1. This shows that the naming has
been confusing in the past. The INTIMATE name convention
(see Björck et al., 1998) has been used for the D-O cycles pre-
sented in Table D1, which shows the dating of each GI and
GS and its duration for the GRIP ice core.
History
D-O cycles are named after the two scientists who led the
work on the phenomena, namely the Danish scientist Willy
Dansgaard and the Swiss scientist Hans Oeschger. They were
the first to relate the changes in the d
18
O record from the
Greenland ice cores to a possible climate change over the North
Atlantic region, and to changes in the thermohaline circulation
(THC) (Dansgaard et al., 1983).
D-O cycles were first seen in the d
18
O record from the Camp
Century ice core (Hansen and Langway, 1966) that was drilled
in northwestern Greenland (see Oxygen isotopes and Paleotem-
peratures, proxy reconstructions to understand how d
18
Ois
used as a proxy for temperature). However, the climatic signif-
icance was not realized until the same behavior was seen in the
d
18
O record of the Dye 3 ice core (from southern Greenland).
Finally, the climatic relevance of the D-O cycles in Greenland
ice cores was confirmed when the two Summit drillings, GRIP
(Dansgaard et al., 1993) and GISP2 (Grootes et al., 1993), were
finished. At Summit, where horizontal movement is slow, the
changes in the oxygen isotope record cannot originate from
internal folding of the ice layers due to ice movement over
rough bedrock. Therefore, they must be a signal of actual cli-
mate changes occurring over most of Greenland. At first, the
signal was circumspectly attributed to climatic changes in the
Greenland/North Atlantic region. Later, the area was enlarged
to include Europe when evidence of D-O cycles was found in
Europe (Thouveny et al., 1994; Allen and Huntley, 2000). The
area influenced by the D-O cycles was once again enlarged to
encompass most of the Northern Hemisphere and even across
the Equator. Evidence of D-O cycles has been found in
paleo-data from different regions of the world, for an up-to-date
overview see Voelker et al. (2002) and Figure D3.
Dansga ard-O eschger cycle s in ocean sedime nt cores
Ocean sediment cores, particularly from the North Atlantic and
Nordic seas, have also shown evidence of D-O cycles
(Rasmussen et al., 1996; Elliot et al., 1998; Dokken and
Jansen, 1999; van Kreveld et al., 2000). In ocean sediment
cores, D-O cycles can be identified in many of the different cli-
mate proxy records, a few of which will be mentioned below.
Layers of ice-rafted debris (IRD) are observed just before
D-O events. Hence, it is thought that the warming is preceded
by large amounts of icebergs carrying the IRD, which is then
deposited in the North Atlantic region. The IRD signal is
Figure D1 The GRIP d
18
O record is shown for the last 115,000 years. The Dansgaard-Oeschger events (GI / warm phase of the D-O cycle) are
shaded and numbered accordingly (Bjo
¨
rk et al., 1998 ). Note the sawtooth shape. During MIS 3, the two Bond cycles, consisting of GI:12-11-10-9
and GI:8-7-6-5, LGM denotes last glacial maximum. Only the last 4,000 years of the last warm period (Eemian) is shown to the right. The time scale
is the GRIP time scale SS09 (after Johnsen et al., 1995).
Figure D2 The figure shows schematically the name convention of the
INTIMATE group for the Greenland stadials (GS) and Greenland
interstadials (GI). It should be clear from the figure that each
Dansgaard-Oeschger cycle consists of a GS and a GI. For the whole
glacial record see Figure D1.
Table D1 The transitions from stadials to interstadials and vice versa
use the GRIP SS09 time scale; the names for the Dansgaard-Oeschger
cycles are based on Bjo
¨
rck et al. ( 1998). The ages correspond to the GIs
and GSs shown in Figure D1. Stadial numbers that have a “ H” and a
number are stadials in which Heinrich events took place, and the
number of the Heinrich event. This is not part of the INTIMATE naming
convention, but is added for illustrative reasons
Events GRIP age
(years bp)
Duration Events GRIP age
(years bp)
Duration
Holocene 11,551
GI-1 14,491 1,780 GS-1 12,711 1,160
GI-2 21,631 240 GS-2
H1
21,391 6,900
GI-3 25,571 300 GS-3
H2
25,271 3,640
GI-4 28,311 300 GS-4 28,111 2,540
GI-5 30,011 560 GS-5
H3
29,451 1,140
GI-6 31,191 360 GS-6 30,831 820
GI-7 32,911 790 GS-7 32,121 930
GI-8 35,731 1,810 GS-8 33,921 1,010
GI-9 37,671 240 GS-9
H4
37,431 1,700
GI-10 38,991 680 GS-10 38,311 640
GI-11 40,851 1,180 GS-11 39,671 680
GI-12 44,371 2,560 GS-12 41,811 960
GI-13 46,751 860 GS-13
H5
45,891 1,520
GI-14 51,991 4,980 GS-14 47,011 260
GI-15 53,571 400 GS-15 53,171 1,180
GI-16 56,011 1,780 GS-16 54,231 660
GI-17 56,851 560 GS-17 56,291 280
GI-18 61,771 280 GS-18
H6
61,491 4,640
GI-19 69,751 2,060 GS-19 67,691 5,920
GI-20 74,041 2,430 GS-20 71,611 1,860
GI-21 83,091 7,620 GS-21 75,471 1,430
GI-22 88,700 2,609 GS-22 86,091 3,000
GI-23 103,946 7,846 GS-23 96,100 7,400
GI-24 106,180 GS-24 105,601 1,655
230 DANSGAARD-OESCHGER CYCLES
mostly of local character; the IRD has been shown to originate
from the Fennoscandian (also known as the Scandinavian Ice
Sheet), Greenland and Icelandic Ice Sheets. A few of these
IRD events are also called Heinrich events; in these cases, the
IRD mostly originates from the Laurentide Ice Sheet. Heinrich
events are very large outbreaks of IRD / icebergs and only
occurred six times during the Wisconsinan glacial (see
Table D1 ). IRD events associated with outbreaks of icebergs
are considered to imply a freshwater input into the North
Atlantic, with a succeeding cooling of the region. After each
cooling, the region warms very rapidly, this is seen for example
in the abundance of the planktonic species (near surface) Neo-
globoquadrina pachyderma (sin). N. pachyderma is a cold
water species and therefore it almost disappears during the
warm phases. Another indicator of these climate changes is
the d
18
O record in benthic (bottom) species, where the fluctua-
tions in marine sediments mirror the d
18
O fluctuations in ice
cores. For the benthic d
18
O, low isotopic values indicate sta-
dials (cold phases) and high values indicate interstadials (warm
phases). A climate change that can be detected in the fauna at
1,000 meter depth indicates that the climate change must have
affected not only the atmosphere, but also the ocean circulation
(Rasmussen et al., 1996) and deep water convection in the
North Atlantic in particular, which is important for the THC.
The changes in the ocean sediment cores indicate that the
THC was weaker during stadials, and became enhanced to
almost Holocene strength during interstadials. There are also
indications of changes in the geographical occurrence of deep
water convection. It can be seen that convection only occurred
south of Iceland during the stadials, and both south and north
of Iceland (in the Nordic seas) during the interstadials
(Ganopolski and Rahmstorf, 2001).
Dansga ard-O eschger cycle s – a global phenom enon?
As mentioned above and shown in Figure D3, evidence of D-O
cycles has been found over large parts of the globe. Therefore, it
is particularly puzzling why there is no clear signal of D-O cycles
in Antarctic ice cores (e.g., compare Figure D4, profiles A and E).
Antarctic ice cores show a different signal in the d
18
O record,
looking completely different to the Greenland d
18
O record (see
Figure D4). In Antarctica, the sawtooth shape is not present at
all. The cycles look more like a triangle during the warm phase,
with a slow warming, which, after peaking, immediately starts
to slowly decline back to its cold period values. Due to uncer-
tainty in the time scales of the Greenland and Antarctic ice cores,
it was long discussed which came first – the Greenland or the
Antarctic warming, or whether they were synchronous. This con-
troversy was finally resolved when Blunier et al. ( 1998) mea-
sured atmospheric methane in microscopic air bubbles in the
ice cores. The methane is well mixed globally within a year,
and is therefore an inter-hemispheric marker that can be used to
link the records from Greenland and Antarctica, respectively. It
was shown that Antarctica was warming while Greenland was
in the cold phase, and when the Greenland d
18
O record jumped
into a warm phase of a D-O cycle, Antarctica started to cool. This
asynchronous see-saw mechanism between the two hemispheres
still lacks a full explanation. However, it is believed to arise from
changes in the oceanic heat transport from the Southern Hemi-
sphere to the Northern Hemisphere, which is believed to have
differed during the different phases of the D-O cycles.
The trigger of Dansgaard-Oeschger events/cycles
D-O cycles have become a large research field within paleo-
climatology. Evidence of D-O cycles in many different
Figure D3 Global spatial distribution of sites where impacts of Dansgaard-Oeschger cycles have been recorded in paleorecords. Black dots indicate
sites with a clear D-O type oscillation, open circles show sites where no or unclear D-O cycles have been found (after Voelker et al., 2002).
DANSGAARD-OESCHGER CYCLES 231
paleorecords from around the world have made the understand-
ing of these rapid and abrupt climate changes an intriguing
challenge.
Retrieval of ocean sediment cores has shed some light on the
magnitude of the impact of the D-O cycles. However, looking to
the ocean for a possible mechanism for the triggering of D-O
events is problematic, as the natural time scale of the THC is
much longer than the 20–50 years it takes to switch into a D-O
event. Furthermore, the ice sheets react on much longer time
scales, and there is no astronomical forcing with a period of
1,500 years. Sea ice has been mentioned as a possible trigger
mechanism that would influence the deep water convection,
and hence the THC. Another possibility that has been proposed
is surging glaciers from the surrounding ice sheets, but simulta-
neously surging glaciers also lack an explanation or a trigger.
Therefore, the triggering mechanism is still being debated.
Data analysis (Alley et al., 2001) and model studies
(Ganopolski and Rahmstorf, 2002) have shown that a possible
trigger of the D-O events could be stochastic resonance in the
climate system under glacial conditions. A stochastic resonance
implies that small changes, i.e., in the freshwater flux into the
northern North Atlantic, could cause large rapid changes of
the ocean circulation during the glacial period and hence large
climate changes such as the D-O cycles (Ganopolski and
Rahmstorf, 2002).
Irene A. Mogensen
Bibliography
Allen, J.R.M., and Huntley, M.B., 2000. Weichselian palynological records
from southern Europe; correlation and chronology. Quaternary Int.,
73 / 74, 111–125.
Alley, R.B., Clack, P.U., Keigwin, L.D., and Webb, R.S., 1999. Making
sense of millennial-scale climate change. In Clark, R.B., Webb, P.U.,
Keiwin, L.D. (eds.), Mechanisms of Global Climate Change at
Millennial Time Scales. Washington DC: American Geophysical Union,
pp. 385–394.
Alley, R.B., Anandakrishnan, S., and Jung, P., 2001. Stochastic resonance
in the North Atlantic. Paleoceanography, 16(2), 190–198.
Björck, S., Walker, M.J.C., Cwynar, L.C., Johnsen, S., Knudsen, K.-L.,
Lowe, J.J., Wohlfarth, B., and intimate members, 1998. An event strati-
graphy for the Last Termination in the North Atlantic region based on
the Greenland ice-core record: A proposal by the INTIMATE group.
J. Quaternary Sci., 13(4), 283–292.
Blunier, T., Chappellaz, J., Schwander, J., Dällenbach, A., Stauffer, B.,
Stocker, T.F., Raynaud, D., Jouzel, J., Clausen, H.B., Hammer, C.U.,
and Johnsen, S.J., 1998. Asynchrony of Antarctic and Greenland cli-
mate change during the last glacial period. Nature, 394, 739–743.
Bond, G.C., and Lotti, R., 1995. Iceberg discharges into the North Atlantic
on millennial time scales during the last glaciation. Science, 267,
1005–1010.
Dansgaard, W., Oeschger, H., Langway, C.C., Jr., 1983. Ice core indica-
tions of abrubt climate changes. In Palaeoclimatic Research and Mod-
els. Proceedings of Workshop, Brussels, Dec. 1982. Dordrecht/ Boston/
Lancaster: D. Reidel Publishing Company, pp. 72–73.
Dansgaard, W., Johnsen, S.J., Clausen, H.B., Dahl-Jensen, D., Gundestrup,
N.S., Hammer, C.U., Hvidberg, C.S., Steffensen, J.P., Sveinbjörnsdottir,
A.E., Jouzel, J., and Bond, G., 1993. Evidence for general instabi-
lity of past climate from a 250-kyr ice-core record. Nature, 364,
218–220.
Dokken, T.M., and Jansen, E., 1999. Rapid changes in the mechanism
of ocean convection during the last glacial period. Nature, 490,
458–461.
Elliot, M., Labeyrie, L., Bond, G., Cortijo, E., Turon, J.-L., Tisnerat, N.,
and Duplessy, J.-C., 1998. Millennial-scale iceberg discharges in the
Irminger Basin during the last glacial period: Relationship with the
Heinrich events and environmental settings. Paleoceanography, 13(5),
433–446.
Ganopolski, A., and Rahmstorf, S., 2001. Rapid changes of glacial climate
simulated in a coupled climate model. Nature, 409, 153–158.
Ganopolski, A., and Rahmstorf, S., 2002. Abrupt Glacial Climate change
due to Stochastic Resonance. Phys. Rev. Lett.
, 88(3), 1–4.
Grootes, P.M., Stuiver, M., White, J.W.C., Johnsen, S.J., and Jouzel, J.,
1993. Comparison of oxygen isotope records from the GISP2 and GRIP
Greenland ice cores. Nature, 366, 552–554.
Figure D4 The GRIP isotope profile is shown in (a). (b) shows a 8 ky high pass filtering of a after inversion, which is considered the
S-Atlantic temperature proxy. Folding b with e
-t / t
, where 1 /t = 1,000 years, results in the profile shown in c; this profile depends on the time
constant (t) used. The 0.5 to 8 ky band pass filtering of the Byrd isotope profile (e) is shown in (d). The profiles in c and d are quite similar in
shape and give a very high correlation coefficient of 0.87 for t = 1 ky in MIS 3 (22 to 58 ky
BP) (after Johnsen and Stocker, 2003).
232 DANSGAARD-OESCHGER CYCLES
Hansen, B.L., and Langway, C.C., Jr., 1966. Deep core drilling and core
analysis at Camp Century, Greenland 1961–1966. U.S. Antarct. J.,
1, 207–208.
Johnsen, S.J., and Stocker, T., 2003. Identifying the bipolar seesaw from
polar Ice Core Data using a simple Model. EGS-AGU-EUG Joint
Assembly, Nice, France, 6–11 April, 2003. Geophys. Res. Abstr. V.5,
13454, 2003.
Johnsen, S., Dahl-Jensen, D., Dansgaard, W., and Gundestrup, N., 1995.
Greenland palaeotemperatures derived from GRIP bore hole tempera-
ture and ice core isotope profiles. Tellus, 47B, 624–629.
Rasmussen, T.L., Thomsen, E., van Weering, T.C.E., and Labeyrie, L.,
1996. Rapid changes in surface and deep water conditions at the Faeroe
Margin during the last 58,000 years. Paleoceanography, 11(6),
757–771.
Thouveny, N., de Beaulieu, J.L., Bonifay, E., Creer, M.K., Guiot, J., Icole, M.,
Johnsen, S., Jouzel, J., Reille, M., Willams, T., and Williamson, D., 1994.
Climate variations in Europe over the past 140 kyr deduced from rock
magnetism. Nature, 371,503–506.
van Kreveld, S., Sarnhein, M., Erlenkeuser, H., Grootes, P., Jung, S.,
Nadeau, M.J., Pfaumann, U., and Voelker, A., 2000. Northeast Atlantic
sea surface circulation during the past 30–10 14C kyr. BP. Paleoceano-
graphy, 14(5), 616–625.
Voelker, A.H.L., and workshop participants, 2002. Global distribution of
centennial-scale records for Marine Isotope Stage (MIS) 3: A database.
Quaternary Sci. Rev., 21, 1185–1212.
Cross-references
Carbon Dioxide and Methane, Quaternary Variations
Foraminifera
Heinrich Events
Ice Cores, Antarctica and Greenland
Ice-Rafted Debris (IRD)
Interstadials
Isotopic Fractionation
Last Glacial Maximum
Laurentide Ice Sheet
Millennial Climate Variability
Oxygen Isotopes
Paleotemperatures, Proxy Reconstructions
Scandinavian Ice Sheet
Thermohaline Circulation
Wisconsinan (Weichselian, Würm) Glaciation
Younger Dryas
DATING, AMINO ACID GEOCHRONOLOGY
Dating Earth history using the racemization of amino acids pre-
served in biominerals belongs to the chemical family of dating
methods. Chemical methods differ from radioactive dating
techniques in that their reaction rates depend on one or more
environmental parameters, whereas radioactive decay remains
constant regardless of most environmental conditions. Amino
acids, derived from indigenous protein residues protected by
the skeletal hardparts of organisms, survive in most environ-
ments for thousands to millions of years. The extent of racemi-
zation (defined below) of these amino acids is dependent
primarily on the time elapsed since death of the organism and
the integrated thermal history experienced by the biominerals,
and to a lesser extent on vital effects unique to each taxon.
Amino acids are simple organic molecules that are the fun-
damental building blocks of proteins. They are exceptionally
stable, with lifetimes for most of the more stable forms exceed-
ing a million years at ambient temperatures. Although there are
a large number of possible amino acids, only about 20 are com-
mon constituents of proteins. Amino acid geochronology (often
referred to as simply Amino Acid Racemization, or AAR)
relies on the chiral nature of most amino acids. Chiral (derived
from the Greek word for “hand”) molecules are asymmetric, or
exhibit “handedness.” Chiral molecules are not superimposable
on their mirror image. All but the simplest protein amino acid
can exist in either a “left-” or “right-” handed configuration
(enantiomers). The designation of the handedness for chemical
compounds is determined by the direction they rotate plane-
polarized light when in solution. Molecules that rotate polar-
ized light are said to be optically active, whereas an equilibrium
mixture of left- and right-handed amino acids (a racemic mix-
ture) is “ achiral,” or optically inactive, and will not rotate polar-
ized light. The interaction of chiral molecules with other chiral
molecules can be different to interactions with achiral mole-
cules. A simple analog for a chiral compound is a glove, where
the left and right hand gloves are mirror images of each other
but they are not superimposable, whereas a ski pole is achiral.
Bare hands (also chiral) inserted into a box containing only
right-hand gloves, will only emerge with a gloved right hand,
never with a gloved left hand, whereas bare hands inserted
into a box of ski poles would always emerge with a ski pole,
regardless of which hand was inserted. These principles are
exploited by analytical procedures to separate chiral forms of
amino acids.
It is a quirk of nature that the proteins of almost all living
organisms utilize exclusively amino acids of the “left hand” con-
figuration. Left-handed forms are given the abbreviation “
L,”
for levorotatory (they rotate polarized light in the left-hand direc-
tion), whereas right-hand forms are dextrorotatory, or
D-amino
acids. All amino acids have a three-letter abbreviation. Thus the
common amino acid alanine, is abbreviated Ala, and its left-
and right-handed forms are then
D-Ala and L-Ala, respectively.
The extent of amino acid racemization is expressed as the propor-
tion of
D- relative to L-amino acids, or D /L.
Amino acid geochronology capitalizes on the original
uniquely left-handed character of amino acids preserved in ske-
letal hard parts or biominerals. Once the biological constraints
that created optically active assemblages of amino acids are
removed, usually at death of the organism, amino acids begin
to spontaneous invert to their right-handed, or
D-configuration.
The process by which amino acids of one configuration convert
into their opposite handedness is known as racemization. The
rate at which this inversion occurs is strongly dependent on
ambient temperature and, to a lesser extent, vital effects and
environmental pH. At room temperature (22
C), conversion
from an optically pure amino acid (protein from a living organ-
ism has a
D /L = 0) to a racemic mixture lacking optical activity
(
D/L = 1.0 for most amino acids) requires several thousands to
several tens of thousands of years for most amino acids. The
racemization reaction rate is exponentially related to tempera-
ture; at 160
C, a racemic mixture is reached in 10–20 h,
whereas in Arctic regions (10
C) equilibrium requires 1–2
million years. As a rule of thumb, at normal ambient tempera-
tures (0–25
C), the rate of racemization doubles for every 4
C
increase in temperature. Figure D5 illustrates the useful time
range for AAR dating under different mean annual tempera-
tures for isoleucine; aspartic acid racemizes about an order of
magnitude faster and is widely used for very young materials,
or extremely cold sites (e.g., Goodfriend, 1992; Goodfriend
et al., 1996; Figure D6). The effect of temperature on reac-
tion rate is readily apparent in a plot of racemization levels for
last interglacial “moderate” racemization-rate mollusks from
the European Arctic to the Mediterranean (Figure D7).
DATING, AMINO ACID GEOCHRONOLOGY 233
Racemization of amino acids in fossils
Fifty years ago, Abelson (1954) first showed that amino acids
could be preserved in biominerals for hundreds of millions of
years, and in a series of papers a few years later with P.E. Hare,
outlined the potential of amino acid racemization as a geologi-
cal dating method (Hare and Abelson, 1968). They showed that
mollusks of known age exhibited increasing levels of racemiza-
tion with increasing age, and the dating method of Amino Acid
Geochronology was launched. Recent summaries of the meth-
ods are given by Miller and Brigham-Grette (1989), Wehmiller
(1993), Wehmiller and Miller (2000).
Not all amino acids are suitable for geochronology. Glycine,
the simplest and one of the most common amino acids, is not a
chiral molecule. Other simple amino acids, such as alanine, are
created in fossils by the decomposition of other more complex
amino acids. Thermodynamically unstable amino acids, like
serine and threonine, have half-lives too short for most geologi-
cal applications. The most widely used slow-racemizing amino
acids are isoleucine, leucine, valine and glutamic acid, whereas
aspartic acid racemizes much more rapidly (Figure D6). All
of these amino acids except isoleucine have a single chiral
carbon atom, and can exist in either their
D-orL-forms, which
have identical physical and chemical properties except in their
interaction with other chiral substances. Isoleucine, in contrast,
has two chiral carbon atoms and can racemize about either
carbon, or both. In nature, isoleucine racemizes almost exclu-
sively about the alpha carbon, producing a new molecule,
D-alloisoleucine, which has different chemical and physical
characteristics, making it easier to separate in the laboratory.
This reaction is known as epimerization, and the extent of
isoleucine epimerization is abbreviated A/I or aIle/Ile.
The separation, detection and quantification of the
D- and
L-amino acids are accomplished by either gas or liquid chromato-
graphy.
L-isoleucine can be separated from its D-alloisoleucine
pair by conventional ion-exchange high-pressure liquid chroma-
tography (HPLC), whereas most other amino acids are separated
by gas chromatography (GC) or reverse-phase liquid chromato-
graphy (RPLC) using a chiral stationary or mobile phase to
separate the enantiomers.
Figure D6 Fast-racemizing amino acids provide superior resolution for
cold or very young samples. In this figure,
D/L aspartic acid and
D-alloisoleucine/ L-isoleucine (A /I) were measured in the marine bivalve
Hiatella arctica, collected from raised marine terraces along the north
coast of Alaska (mean annual temperature ca. 12
C). For the
younger shorelines,
D/L Asp provides a clearer separation of the sites,
but this difference decreases for the older shorelines. (mo = modern;
Pe = Pelukian shoreline (last interglacial); Wa = Wainwrightian shoreline;
Fi = Fishcreekian shoreline; Bi = Bigbendian shoreline; Co = Colvillian
shoreline (probably late Pliocene)). Modified from Goodfriend et al.
(1996).
Figure D7 D-alloisoleucine/ L-isoleucine (A /I) in moderate-
racemization-rate mollusks from last interglacial (ca. 125 ka) sites across
western Europe, ranging from Svalbard and Arctic Russia to the
Mediterranean Basin, plotted against current mean annual site
temperature. Last interglacial sites are dated by U /Th on corals or
correlated to the last interglacial on the basis of diagnostic marine
faunal elements. The data document the exponential dependency of
racemization rate on site temperature. Unpublished data from
G.H. Miller and P.J. Hearty.
Figure D5 Relation between the extent of isoleucine epimerization
(A / I) and sample age for effective diagenetic temperatures between
0 and +25
C showing the strong dependence of epimerization rate on
temperature. Lines derived from kinetics in eggshells of the Australian
emu (Dromaius) shown in Miller et al. (2000).
234 DATING, AMINO ACID GEOCHRONOLOGY