Odekon M. Encyclopedia of paleoclimatology and ancient environments

Подождите немного. Документ загружается.

Winkler, M.G., and Van Helden, A., 1992. Representing the heavens:

Galileo and visual astronomy. ISIS, 83, 195–217.

Woodward, H.B., 1913. Jamieson–Obit. Geol. Mag., 50, p.332–333.

Zeuner, F.E., 1945/1959. The Pleistocene Period. London: Royal Society,

2447p.

Zeuner, F.E., 1946. Dating the Past. London: Methuen, 444pp.

Cross-references

Astronomical Theory of Climate Change

CLIMAP

Dendroclimatology

Glacial Geomorphology

Glaciations, Quaternary

History of Paleoclimatology

Oxygen Isotopes

Plate Tectonics and Climate Change

Sea Level Change, Post-Glacial

SPECMAP

Stable Isotope Analysis

Sun-Climate Connections

Tills & Tillites

HOLOCENE CLIMATES

Introduction and definition

The Holocene – or “wholly recent”–epoch is the youngest

phase of Earth history, which began when the last glaciation

ended, and for this reason it is sometimes also known as the

post-glacial period. In reality, the Holocene is one of many

interglacials that have punctuated the late Cenozoic ice age.

The term was introduced by Gervais in 1869 and was accepted

as part of valid geological nomenclature by the International

Geological Congress in 1885. A summary account of Holocene

environmental history is provided by Roberts (1998). A com-

mission of the International Union for Quaternary Research

(INQUA) has been devoted to the study of the Holocene, and

it is also a major focus of the PAGES (Past Global Changes)

program. Since 1991, the journal The Holocene has been dedi-

cated exclusively to Holocene research.

There are three principal schools of thought about how the

Holocene should be formally defined. Because the Holocene

still “lives,” debate has focused on defining its onset, namely

the Pleistocene-Holocene boundary. The first school of thought

(Watson and Wright, 1980) believes that the beginning of

Holocene is easily recognized in individual records of degla-

ciation or biotic change (e.g., in pollen diagrams), but that

because these changes are not exactly the same age every-

where, the Pleistocene-Holocene boundary should therefore

be time-transgressive. A second school prefers the “type-

section” approach commonly used in geology to define strati-

graphic boundaries. Mörner (1976), for example, proposed

that a sequence in southern Sweden might be the standard

reference point for the Pleistocene-Holocene boundary. The

third view, and one which is supported by INQUA, is that the

Holocene is simply defined as beginning 10,000 radiocarbon

(

C) years ago.

C chronologies count AD 1950 as being the “present-day”

and also underestimate true – or calendar – ages by several

centuries for most of the Holocene. Cross-dating of pines from

the western United States and northwest Europe has revealed a

systematic divergence between

C and tree-ring ages in sam-

ples more than 2,500 years old, with a discrepancy amounting

to 600 years (or 12%) at 5,000 y

BP. Moreover, the calibration

curve is not a smooth line but contains a number of deviations

that represent short-term variations in atmospheric

C produc-

tion. This is significant in that some

C dates have more than

one calendar age. Current estimates for the beginning of the

Holocene put it at around 11,500 y

BP in calendar years, based

on counts of tree rings and annual ice layers (Alley et al.,

1993; Gulliksen et al., 1998). It is now possible to calibrate

C dates for the whole of the Holocene (Stuiver and Reimer,

1993), and because it is better to use calendar rather than radio-

carbon years as a time frame for past global changes, the

former (i.e., cal. y

BP) will be used in this entry.

Various attempts have been made to subdivide the Holocene,

usually on the basis of inferred climatic changes. Blytt and

Sernander, for instance, proposed a scheme of alternating cool-

wet and warm-dry phases based on shifts in peat stratigraphy in

northern Europe. However, with the systematic application of

C dating it has become clear that Holocene vegetation and

climatic changes were more often than not time-transgressive,

and as a result most researchers prefer to use age (e.g., 8,000

cal. y

BP) rather than stage (e.g., Boreal-Atlantic transition) for

defining past events.

Data sources

For the last few centuries of the Holocene, instrumental data

are available and can be used to build up high-quality time-

series for individual climate variables. However, because the

climate system varies over timescales longer than can be

observed directly, it is also necessary to turn to longer-lived

human and natural archives. For intermediate time-scales,

documentary sources can be used (Bradley and Jones, 1992),

although these data were not always collected in a form ideal

for the analysis to which we wish to subject them. The length

of the “historical” time period also varies greatly in different

parts of the world, extending back around 5,000 years in regions

such as Egypt, but little more than a few decades in places like

Papua-New Guinea. A third group of techniques operates over

longer Holocene timescales. These are proxy-climate methods

such as dendroclimatology and pollen analysis, which can some-

times be combined within a single sequence to create a multi-

proxy approach.

In environments with little human disturbance, it is widely

believed that organisms – terrestrial, aquatic or marine – will

eventually achieve a state of equilibrium with the prevailing

climate. Modern data on climate-species relationships can

therefore be applied to paleoecological data from sediment

cores to produce quantitative estimates of past temperatures

or rainfall levels, using transfer functions or other multivariate

statistical methods. Climatic calibration of modern and fossil

pollen data has allowed maps to be produced showing tempera-

ture, precipitation or air mass distribution for different times

during the Holocene over eastern North America (Webb et al.,

1993). An alternative, complementary approach involves cali-

brating proxy-climate indicators, such as the stable isotope sig-

nature in cave speleothems, against local meteorological

records for the last 50–100 years. If a statistically significant

correlation can be found, the proxy-climate records can then

be used to extend the time series back to before the beginning

of recorded weather observations.

438 HOLOCENE CLIMATES

The last deglacial climatic transition

The Last Glacial Maximum occurred around 25,000–18,000 years

ago, according to calibrated radiocarbon dating. At this time, tem-

peratures on land were lowered by as much as 20



C although

ocean temperatures changed much less. Under the full glacial cli-

mate, ice sheets more than 4 km thick lay across northern Europe

and North America, and smaller ice caps existed in the Alps, the

southern Andes, and parts of East Asia. Climatic warming, and

the consequent deglaciation of these continental regions, took

place in a series of steps that appear to have been broadly syn-

chronous over large parts of the globe, although there may have

been some asynchrony between the Northern and Southern

Hemispheres.

Deglaciation, which took about 9,000 years to complete in

Europe and North America, was an unsteady process, and at var-

ious times towards the tail end of the Pleistocene ice sheets tem-

porarily re-advanced. In at least some cases, this was the result of

a temporary but sharp oscillation in climate, which represented a

false start to the present interglacial. In the circum-North Atlantic

region, this late Pleistocene thermal oscillation comprised a

warming event often known as the Bølling-Allerød interstadial

(~15,000–12,700 cal. y

BP) followed by a cooling event called

the Younger Dryas stadial (12,700–11,500 cal. y

BP). Evidence

from coleopteran (beetle) and chironomid (midge) fossils shows

that summer temperatures warmed rapidly to reach – or even

exceed – those of today soon after 15,000 cal. y

Ice cores provide some of our most important paleoclimatic

data for ice age climate, with Greenland possessing the most

important record in the Northern Hemisphere. Ice has accumu-

lated faster here than in Antarctica and it therefore offers a more

detailed sequence of past climatic change, with an annual time

resolution for the part covering the late Pleistocene and

Holocene. The GRIP and GISP2 cores drilled at the summit of

the Greenland ice sheet between 1988 and 1993 provide espe-

cially important climate records. They show the same late-glacial

climatic oscillation as revealed in biotic records, and confirm that

the warming steps were extremely rapid. The last of these, mark-

ing the transition from the Younger Dryas stadial to the Holocene

interglacial proper, involved a temperature rise of 7



C over no

more than 50 years (Alley et al., 1993). These sudden jumps in

climate may be linked to periodic on and off switching of the

thermohaline “conveyor belt ” circulation in the Atlantic Ocean

(Bard and Broecker, 1992). The Younger Dryas climatic reversal,

for example, was most probably the result of a partial shut-down

of the Atlantic conveyor caused by large pulses of meltwater

from adjacent Northern Hemisphere ice sheets.

The early Holocene climatic optimum

Because climatic change was especially rapid during the millennia

either side of the Pleistocene-Holocene boundary, many Earth sur-

face systems experienced lagged responses. Global (eustatic) sea

levels, for example, lay at about 40 m at 11,000 cal. y

BP and only

reached modern elevations around 6,000–5,000 years ago

(Edwards et al., 1993). Similarly, because soil formation and vege-

tation cover often lagged behind the climate, many landscapes

experienced a phase of temporary geomorphic instability during

the Pleistocene-Holocene transition, with high rates of erosion

and sediment flux.

In many mid-latitude Northern Hemisphere regions, a thermal

optimum can be identified between ~10,000 and ~6,000 cal. y

from the extension of species such as the water chestnut and

the pond tortoise north of their present European climatic limits.

The cause of these long-term climatic changes lies in astronomi-

cally-controlled variations in the Earth’s receipt of solar radiation

(Figure H6). At the start of the Holocene, one of the three Croll-

Milankovitch cycles – the precession of the equinoxes – lay

at the opposite point on the cycle from where it lies today. There-

fore, the Northern Hemisphere received nearly 8% more solar

radiation during summer months then than at present, and north-

ern summers were considerably warmer in consequence.

Because climatic conditions exert a powerful influence over

geomorphic processes, relict landforms can furnish indications

about past climates. In the tropics and sub-tropics, two of the

most climatically-sensitive geomorphic regimes are desert

dunes and lakes. As a result of the close relationship between

eolian processes and climatic aridity, fossil dunes have helped

to provide an indication of the changing extent and location

of the World’s arid zones (Goudie, 1983). Eolian landforms

not only indicate times of climatic aridity but also reveal former

wind strengths and directions. While dunes indicate periods of

climatic aridity, high-level lakes reflect periods of positive

water balance, and in the Holocene at least, these are associated

with phases of wetter climate. Direct geomorphic evidence of

former high levels derives from shoreline terraces and related

landforms left above present-day lakes.

It was recognized in the early twentieth century that deserts

were not always as arid as they are today. In addition to the

evidence of high-level lakes, ancient rock paintings from the

Sahara depict herds of savanna mammals, while pollen dia-

grams and rodent nests indicate much richer, often wooded,

vegetation cover. Within the tropics,

C dating has shown that

the last “pluvial” (or wet) period started during the terminal

Pleistocene and continued through the first half of the Holocene

(14,000–5,500 cal. y

BP). Water-balance calculations for these

enlarged early Holocene lakes indicate that rainfall in East

Africa, the Sahara, and northwest India increased by between

150 and 400 mm per year (Street-Perrott et al., 1991). This

increase was caused by an enhanced monsoon circulation that

brought summer rains further north than they occur today. The

Figure H6 Changing receipt of solar radiation and other climatic

controls in the Northern Hemisphere since the Last Glacial Maximum

(modified after Kutzbach and Street-Perrott, 1985)

HOLOCENE CLIMATES 439

stronger monsoon was itself generated by the “thermal opti-

mum,” which led to greater heating over the land – especially

over south Asia – and caused low pressure, which sucked in

the humid, monsoonal air masses from the south.

In mid-latitude regions such as the North American prairie,

a different pattern of climate change occurred around this time.

Pollen data show that the mid-western prairie-forest boundary

moved eastwards until 8,000 cal. y

BP, but subsequently shifted

back in the opposite direction. As this boundary is today con-

trolled by moisture availability, it is reasonable to use it as a

proxy climatic indicator, recording climatic conditions between

8,000 and 4,500 cal. y

BP that were drier than at present. Sedi-

ment cores from some of the many small lakes of the North

American Great Plains also record this climatic phase. Diatom

and geochemical analyses, for example, show high salinities

during the first half of the Holocene, while before and after this

phase the lake water was relatively fresh (Haskell et al., 1996).

This orbital forcing of the early Holocene climatic optimum

appears to have been amplified by the biosphere itself. Model-

ing experiments have shown that vegetation had a significant

effect on climatic conditions in key boundary areas, such as

the edges of forest and desert biomes, by changing the surface

albedo and moisture conditions. Along the boreal-tundra eco-

tone, for example, calculations suggest that the expanded forest

caused an additional warming of about 4



C in spring tempera-

tures 6,000–7,000 years ago, over and above that caused

directly by solar heating (Foley et al., 1994).

These longer-term climatic trends were interrupted on several

occasions during the Holocene by abrupt events of shorter, centen-

nial duration. In the tropics, they are recorded in abrupt lake-level

falls, dated in Africa to ~12,000, 8,200 and 5,200 cal. y

BP (Gasse,

2000). As with the rapid rises in temperature at the Pleistocene-

Holocene transition, these hydrological changes occurred so

rapidly that their true rate cannot be determined by

Cdating.

The abrupt arid events are also recorded by the Greenland ice

core record as sharp troughs in atmospheric methane, implying

a reduction in tropical wetlands, which are a major source of this

greenhouse gas (Figure H7).

Late Holocene climatic variability

Although less marked than during the early Holocene, the later

Holocene nonetheless experienced some significant shifts in

climate. Among these were the progressive cooling following

the Holocene thermal optimum and the climatic desiccation

of the northern sub-tropics after 6,000 cal. y

BP that created the

modern Saharan, Arabian and Thar deserts. The water levels of

some low-latitude lakes fell abruptly at this time (Figure H7),

suggesting rapid suppression of the monsoonal circulation sys-

tem, which controls tropical rainfall (Street-Perrott et al., 1991;

Gasse, 2000).

Superimposed on longer-term climatic trends linked to orbital

insolation forcing were shorter-term secular variations from wet

to dry, or cold to warm, recorded in the stratigraphy of peat bogs

(Charman et al., 1999) or the advance and retreat of mountain

glaciers. Whether these oscillations represent true, regular cycles

is uncertain, as is the question of the causes of these climatic

changes. Among the most likely controlling mechanisms for

sub-Milankovitch climatic variations are changes in the radiation

output of the Sun, linked to sunspot cycles, and the heat-

shielding effect of volcanic eruptions with large dust emissions.

Volcanic aerosols can play an important role in disturbing

the Earth’s climate, for example by creating a dust veil that

reduces the transparency of the atmosphere. After major

historically-recorded eruptions like that at Laki in Iceland in

1783, contemporary writers such as Benjamin Franklin recorded

severe late frosts and peculiar hazes in regions far removed

from the volcanoes themselves (Grattan and Brayshay, 1995).

Some earlier Holocene eruptions were substantially larger than

the Laki fissure eruption, and it has been suggested that they

may have had a proportionately bigger climatic impact (Zielinski

et al., 1994). For example, in about 1628

BC (3,578 cal. yBP), the

Aegean island of Santorini (Thera) exploded catastrophically,

forcing its Bronze Age population to flee or perish. Dendroclima-

tological studies have shown that the tree rings of Irish bog oaks

were unusually narrow at this time, while bristlecone pine growth

rings of the same age in California show signs of frost damage.

These have been interpreted as the result of a cooler and/or

wetter climate. However, in Turkey, which is warm and dry in

summer, the same climatic change appears to have given

rise to more, not less, favorable growing conditions and to

broader tree rings (Kuniholm et al., 1996). While the evidence

linking tree-rings, climate and volcanic eruptions is impressive,

it also suggests that in most cases the climate perturbation was

short-lived, typically a decade or less.

Increasing evidence is also being found to support a role for

variations in solar radiation upon Holocene climates, typically

based on the observation that cyclic fluctuations in climate proxies

have periodicities close to those found in the record of solar activ-

ity (Bond et al., 2001). In some cases, well-dated climatic anoma-

lies have been to correlate with a known forcing function, such as

European temperature minima coinciding with the Maunder mini-

mum in solar activity of

AD 1650–1715 (Mauquoy et al., 2002).

In addition to these external mechanisms, variations in

Holocene climate will also have been forced by cyclic

Figure H7 Tropical African lake-level record compared with

atmospheric methane concentration recorded in a Greenland ice core

(modified after Roberts, 1998; data from Gillespie et al., 1983; Blunier

et al., 1995). Arrows show abrupt arid events interrupting longer-term

climatic trends. Ages are in 1000 y

BP (cal).

440 HOLOCENE CLIMATES

processes originating within the Earth system, such as the

quasi-periodic ~1500-year long Bond cycles during which ice-

berg “armadas” drifted southward into the North Atlantic,

slowing down the oceanic thermohaline conveyor belt. A com-

parable but much higher-frequency auto-cyclic phenomenon is

the well-known El Niño-Southern Oscillation (ENSO).

Although ENSO-type events appear to have operated through-

out the Holocene, there is evidence, especially from high-

resolution stable isotope analysis of corals, that the current

ENSO frequency spectrum was not established until mid-

Holocene times (Gagan et al., 2000).

The period during the Holocene with the best-resolved and

most reliably calibrated climate data is the last millennium, includ-

ing both the Mediaeval Climate Anomaly (~

AD 1000–1250) and

Little Ice Age (~

AD 1250–1850, particularly AD 1590–1850) dur-

ing which major rivers like the Thames and the Tiber regularly

froze over. Data come from proxy-climate records, notably tree

rings, from historical observations and – for the last two

centuries – from instrumental records. Following the initial work

of Mann et al. (1999) a number of synthetic curves have been gen-

erated from available data sets for Northern Hemisphere tempera-

ture changes. Source of these have a “hockey-stick” shape, in

which temperatures fell progressively to reach a minimum in the

seventeenth century, beforerisingsharplyduringthelast150years.

It remains an unresolved question how much of the observed rise

of ~1



C since the mid-nineteenth century is due to anthropogenic

increases in atmospheric greenhouse gases, and how much is due

to natural variability. However, global temperatures are probably

now higher than at any time during the last 1,000 years, and it is

likely that most of the post-1980 warming is of human origin.

To judge by recent experience, drought-prone regions can be

expected to have been at least as severely affected by any cli-

matic fluctuations as mid- and high-latitude areas, not only

because of temperature change, but also due to rainfall variabil-

ity. During the last millennium, greater aridity appears to have

marked the North American Great Plains at the time of the

Little Ice Age to judge from lake salinity variations, while

increased dust flux onto the ice caps of the Andes suggests a

similar tendency towards drought during this time interval.

By contrast, East African lakes were at low levels at the time

of the Mediaeval Climate Anomaly and were often higher

during the Little Ice Age (Verschuren et al., 2004).

Conclusion

It is clear that since the end of the last glacial stage, climatic sta-

bility over decadal to millennial timescales has been far from

the norm. A Holocene perspective can help to disentangle the dif-

ferent factors that determined these variations in global climate.

Secular climatic cycles are recorded by many paleoclimatic

indicators, including peat bog profiles, salinity fluctuations in

non-outlet lakes, and in cores from ice caps. Some of these

archives, such as variations in tree ring widths, can be resolved

to individual years, and this permits investigation of the potential

causes of short-term changes to the climate.

A Holocene timescale is also valuable if we want to estab-

lish how typical current rates of climatic change are, compared

to those of the recent geological past. During the last major global

warming around 11,500 years ago, a gradual orbital forcing of cli-

matewas amplified by a series of positive feedback mechanismsto

cause an abrupt switch from glacial to interglacial conditions

within the span of a human lifetime. The current temperature rise

is therefore not unprecedented, although the climate is heading

towards conditions warmer than at any time during the Holocene.

Past climatic changes offer further warnings about the potential

instability of climate and the consequent hazards of climate predic-

tion. On at least three occasions within the last 10,000 years, the

climate deteriorated abruptly, only to return to its previous state

within a few centuries.

If it is climatically calibrated and homogenized, proxy data

for the Holocene can also be compared against numerical (glo-

bal circulation models) GCM experiments run with different

boundary conditions (COHMAP, 1988). Such data versus

model comparisons allow the sensitivity and reliability of

GCMs to be tested, hence improving their ability to predict

future climates. The climate during the early Holocene “ther-

mal optimum,” in particular, has been used to test and calibrate

numerical models of the atmosphere. On the other hand, the

Holocene does not offer a direct analogue for a future green-

house gas-warmed climate, because it was dominated by seaso-

nal, rather than mean annual, changes in the Earth’s net receipt

of solar radiation.

Neil Roberts

Bibliography

Alley, R.B., Meese, D.A., Shuman, C.A., Gow, A.J., Taylor, K.C., Grootes,

P.M., White, J.W.C., Ram, M., Waddington, E.D., Mayewski, P.A., and

Zielinski, G.A., 1993. Abrupt increase in Greenland snow accumulation

at the end of the Younger Dryas event. Nature, 362, 527–529.

Bard, E., and Broecker, W.S. (eds.), 1992. The Last Deglaciation: Absolute

and Radiocarbon Chronologies. NATO ASI Series. Berlin: Springer.

Blunier, T., Chappellaz, J., Schwander, J., Stauffer, B., and Raynaud, D.,

1995. Variations in atmospheric methane concentration during the

Holocene epoch. Nature, 374,46–49.

Bond, G., Kromer, B., Beer, J., Muscheler, R., Evans, M.N., Showers, W.,

Hoffmann, S., Lotti-Bond, R., Hajdas, I., and Bonani, G., 2001. Persis-

tent Solar Influence on North Atlantic climate during the Holocene.

Science, 294, 2130–2136.

Bradley, R.S. and Jones, P.D. (eds.), 1992. Climate since A.D. 1500.

London: Routledge.

Charman, D.J., Hendon, D., and Packman, S., 1999. Multi-proxy surface

wetness records from replicate cores on an ombrotrophic mire: Implica-

tions for Holocene palaeoclimate records. J. Quaternary Sci., 14,

451–463

COHMAP members, 1988. Climatic changes of the last 18,000 years:

Observations and model simulations. Science, 241, 1043–1052.

Edwards, R.L., Beck, J.W., Burr, G.S., Donahue, D.J., Chappell, J.M.A.,

Bloom, A.L., Druffel, E.R.M., and Taylor, F.W., 1993. A large drop

in atmospheric

C and reduced melting in the Younger Dryas,

documented with

238

Th ages of corals. Science, 260, 962–968.

Foley, J.A., Kutzbach, J.E., Coe, M.T., and Levis, S., 1994. Feedbacks

between climate and boreal forests during the Holocene epoch. Nature,

371,52–54.

Gagan, M.K., Ayliffe, L.K., Beck, J.W., Cole, J.E., Druffel, E.R.M.,

Dunbar, R.B., and Schrag, D.P., 2000. New views of tropical paleocli-

mates from corals. Quaternary Sci. Rev., 19,45–64.

Gasse, F., 2000. Hydrological change in the African tropics since the Last

Glacial Maximum. Quaternary Sci. Rev., 19, 189–212.

Gillespie, R., Street-Perrot, F.A., and Switsur, R., 1983. Post-glacial arid

episodes in Ethiopia have implications for climate prediction. Nature,

306, 681–683.

Goudie, A.S., 1983. The arid Earth. In Gardner, R., and Scoging, H. (eds.),

Megageomorphology. Oxford: Clarendon Press, pp. 152–171.

Grattan, J., and Brayshay, M., 1995. An amazing and portentious summer:

Environmental and social responses in Britain to the 1783 eruption of

an Iceland volcano. Geogr. J., 161, 125–134.

Gulliksen, S., Birks, H.H., Possnert, G., and Mangerud, J., 1998. The

calendar age of the Younger Dryas – Holocene transition at Kråkenes,

western Norway. The Holocene, 8, 249–260.

Haskell, B.J., Engstrom, D.R., and Fritz, S.C., 1996. Late Quaternary

paleohydrology in the North American Great Plains inferred from the

HOLOCENE CLIMATES 441

geochemistry of endogenic carbonate and fossil ostracods from Devils

Lake, North Dakota, USA. Palaeogeogr, Palaeoclimatol, Palaeoecol.,

124, 179–193.

Kuniholm, P.I., Kromer, B., Manning, S.W., Newton, M., Latini, C.E., and

Bruce, M.J., 1996. Anatolian tree rings and the absolute chronology of

the eastern Mediterranean, 2220–718 BC. Nature, 381, 780–783.

Kutzbach, J., and Street-Perrott, F.A., 1985. Milankovitch forcing of fluc-

tuations in the level of tropical lakes from 18 to 0 kyr BP. Nature,

317, 130–134.

Mann, M.E., Bradley, R.S., and Hughes, M.K., 1999. Northern Hemisphere

temperatures during the past millennium: Inferences, uncertainties and

limitations. Geophys. Res. Lett., 26, 759–762.

Mauquoy, D., van Geel, B., Blaauw, M., and Plicht, van der J., 2002. Evi-

dence from northwest European bogs shows “Little Ice Age” climatic

changes driven by variations in solar activity. The Holocene, 12,1–6.

Mörner, N.-A., (ed.), 1976. The Pleistocene/ Holocene boundary: A

proposed boundary-stratotype in Gothenberg, Sweden. Boreas, 5,

193–275.

Roberts, N., 1998. The Holocene. An environmental history. Oxford /

Boston: Blackwell.

Street-Perrott, F.A., Mitchell, J.F.B., Marchand, D.S., and Brunner, J.S.,

1991. Milankovitch and albedo forcing of the tropical monsoons:

A comparison of geological evidence and numerical simulations for

9,000 yr BP. Trans. R. Soc. Edinb., Earth Sci., 81, 407–427.

Stuiver, M., and Reimer, P.J., 1993. Extended

C database and revised

CALIB radiocarbon calibration program. Radiocarbon, 35, 215–230.

Verschuren, D., Briffa, K., Hoelzmann, P., Barker, P., Scott, L., Barber. K.,

Snowball, I., and Roberts, N., 2004. Holocene climate variability

in Europe and Africa: A PAGES PEP III time-stream 1 synthesis.

In Battarbee R.W., Gasse F., and Stickley C.E., (eds.), Past

Climate Variability through Europe and Africa. Dordrecht: Kluwer,

pp. 567–582.

Watson, R.A., and Wright, H.E. Jr., 1980. The end of the Pleistocene: A

general critique of chronostratigraphic classification. Boreas, 9,

153–163.

Webb, T., Bartlein, P.J., Harrison, S.P., and Anderson, K.H., 1993. Vegeta-

tion, lake levels and climate in eastern North America for the past

13,000 years. In Wright, H.E. Jr., Kutzbach, J.E., Webb, T.

Ruddiman, W.F., Street-Perrott, F.A., and Bartlein, P.J. (eds.), Global

Climates since the Last Glacial Maximum. Minneapolis: University of

Minnesota Press, pp. 415–467.

Zielinski, G.A., Mayewski, P.A., Meeker, L.D., Whitlow, S., Twickler, M.S.,

Morrison, M., Meese, D.A., Gow, A.J., and Alley, R.B., 1994. Record

of volcanism since 7000 B.C. from the GISP2 Greenland ice core and

implications for the volcano-climate systems. Science, 262, 948–952.

Cross-references

Arid Climates and Indicators

Astronomical Theory of Climate Change

Bølling-Allerød Interstadial

Climate Variability and Change, Last 1000 Years

COHMAP

Dating, Dendrochronology

Dendroclimatology

Eolian Sediments and Processes

Hypsithermal

Ice Cores, Antarctica and Greenland

Lake Level Fluctuations

Last Glacial Termination

Late Quaternary-Holocene Vegetation Modeling

Little Ice Age

Medieval Warm Period

Millennial Climate Variability

PAGES

Paleoclimate Modeling, Quaternary

Paleo-El Niño-Southern Oscillation (ENSO) Records

Pollen Analysis

Sea Level Change, Post-Glacial

Sun-Climate Connections

Thermohaline Circulation

Volcanic Eruptions and Climate Change

Younger Dryas

HOLOCENE TREELINE FLUCTUATIONS

Timberline and treeline

Treelines are climatically sensitive transitional zones between

closed forests and alpine, polar, or dry grassland communities.

In most cases, temperature and precipitation thresholds deter-

mine whether a site is treeless or not (Arno and Hammerly,

1993; Körner, 1999). However, many other factors, such as vege-

tation type, soil type, snow cover, topography, and wind can

locally co-determine the boundary of tree growth. Independently

from the biotic and abiotic causes that form this conspicuous

vegetation limit, the spatial position of treeline depends on the

definition of “tree.” The most commonly used criterion to define

“tree” at treeline is a minimum height between 8 and 2 m (Arno

and Hammerly, 1993; Körner, 1999). Some studies prefer to dis-

tinguish between timberline and treeline (e.g., using the height of

trees as a diagnostic criterion: 8 m for timberline and 2 m for tree-

line). The width of the transitional zone separating closed forests

from treeless plant communities is not uniform: polar treelines

and drought-caused treelines can form very broad transition

zones such as parklands with widely spaced trees (Arno and

Hammerly, 1993). Conversely, mountain treelines have rather

narrow transitional zones (i.e., 100–200 m of vertical extent).

It has been suggested that this range overlaps with the altitudinal

span between timberline and treeline as defined above (8 vs. 2 m

of tree height, Tinner and Theurillat, 2003).

The climatic sensitivity of treelines has led to much effort on

the reconstruction of past treeline positions in order to quantify

Holocene climatic changes. The significance of the method for

paleoclimatology lies mainly in two factors: (a) Treeline studies

have local to regional spatial resolutions and therefore provide

valuable information about climatic changes of mountainous or

polar sites, and (b) If combined with independent environmen-

tal proxies (e.g., oxygen isotopes as climatic proxy), treeline

studies can give insights into biosphere responses to rapid

environmental changes.

Reconstructing treeline fluctuations

The reconstruction of treeline fluctuations can be performed

using different approaches. Analysis of pollen, stomata, macro-

fossils, and megafossils (including tree rings) are the most

important methods. In addition, charcoal, soil, and phytolith

analysis may be applied to determine maximum treeline posi-

tions. The methods have different spatial resolutions. Since pol-

len is easily lifted, transported, and deposited by winds over

long distances, pollen records have rather low spatial resolu-

tions. Thus, it is difficult to reconstruct treeline fluctuations

by using the pollen approach alone. Macrofossil and megafossil

analyses are considered to be more reliable tools for recon-

struction of timberline fluctuations (Birks, 2001; Tinner and

Theurillat, 2003) since they have higher taxonomic and spatial

resolutions (meters to decameters) than pollen studies (deca-

meter to kilometers). A disadvantage of macrofossil and mega-

fossils is that, in contrast to pollen, they are not ubiquitous in

Holocene deposits. Producing treeline reconstructions requires

records from today’s treelines that reach back to the late gla-

cial. Two preconditions are essential: (a) the chronology of

the treeline records must be fixed to an absolute scale with suf-

ficient

C-dates of terrestrial fossils, and (b) the study sites

should be situated near to the present-day treeline position.

442 HOLOCENE TREELINE FLUCTUATIONS

To assess the magnitude of past climate changes, reconstruc-

tions of Holocene treeline fluctuations must be related to

modern treeline positions. For estimations of past temperature

changes it is assumed that today’s occurrence of trees is in

equilibrium with the climate of the past few decades. In regions

affected by strong human impact (e.g., the Alps), single trees in

remote areas may still indicate the potential altitudinal limit of

forest and tree growth. Using today’s lapse rates (under the

assumption of Holocene stability), it is possible to convert past

treeline fluctuations (m) into (summer) temperature changes

(



C) (Haas et al., 1998), although it must be considered that

higher treelines in the past reflected warmer periods that were

stable enough to allow subalpine forest to migrate up-slope

(100–200 years of potential response lags). In an attempt to

synthesize common patterns across two continents, we focus

on a few well-documented but representative case studies

from temperature-controlled mountain treelines in Europe and

Northern America.

Holocene treelin e dynamics

Since reconstructions of Holocene treeline fluctuations reflect

past local dynamics, comparisons among the sites are needed

to assess larger-scale patterns that could be related to climatic

changes. Comparisons among pollen and macrofossil stratigra-

phies in the Swiss and Italian Alps allowed the detection of

phases of warm (dry) and cool (wet) phases (Wick and Tinner,

1997) that are in agreement with lake-level fluctuations in the

alpine forelands (CE-1 to CE-8, Haas et al., 1998) and glacier

oscillations (Figure H8). The analysis of pollen and macrofos-

sils from lake and mire deposits along altitudinal transects per-

mits the reconstruction of treeline fluctuations during the

Holocene (Tinner and Theurillat, 2003). These reconstructions

show that the positions of timberline and treeline fluctuated

within a band of 100–180 m during the Holocene and that

treeline was at its uppermost limit (ca. 180 m higher than

today) during the period 10,000–6,000 cal. y

BP (Figure H9).

Severe diebacks of treeline vegetation occurred during cold

(humid) periods (CE-1 to CE-8 as well as during the Little

Ice Age; Figures H8 and H9). A severe treeline decline lasted

from 8,200 to 7,500 cal. y

BP (CE-3), with two minima at 8,000

and 7,600 cal. yr

B.P. (Figure H8, sites Gouillé Rion and Lengi

Egga). After 6,000–5,000 cal. y

BP, timberline and treeline pro-

gressively declined by about 180 m and 300–400 m, respec-

tively (Figure H9). The regression of densely forested areas

Figure H8 Comparison between pollen and macrofossil records of high-elevation sites and glacier fluctuations during the past 12,000 years;

(a) pollen percentages and macrofossil concentrations at Gouille

Rion (Swiss Alps); (b) pollen percentages at Lengi Egga (Swiss Alps);

(c) radiocarbon dates of wood and organic debris collected in front of Alpine glaciers (Hormes et al., 2001); (d) estimated length variation of

Swiss glaciers (Maisch et al., 1999); (e) central European cold-humid phases (Haas et al., 1998). +in a and b show the chronological position of

radiocarbon dates used for depth-age models.

HOLOCENE TREELINE FLUCTUATIONS 443

(timberline) in the Alps during the past 5,000 years was primarily

caused by human impact, whereas the course of treeline gives a

more realistic estimation of the climatic influence (Figure H9).

Assuming constant lapse rates of 0.7



C/100 m, it is possible to

estimate the range of Holocene temperature oscillations in the

Alps to 0.8–1.2



C between 10,500 and 4,000 cal. yBP, when

average (summer) temperatures were about 0.8–1.2



C higher

than today. After 4,000 cal. y

BP, summer temperatures succes-

sively declined to values comparable to those of the twentieth

century (Figure H9).

The general course of Holocene treeline fluctuations in the

Alps is similar to that reconstructed for Scandinavian sites.

Megafossil, macrofossil, and pollen studies suggest that treeline

was at its uppermost limit between 9,500(–9,000) and 6,500

(–6,000) cal. y

BP in Norway and Sweden, about 200–400 m

above today’s treeline position (Kullman, 1995; Dahl and Nesje,

1996; Barnekow, 1999; Barnett et al., 2001). At some study sites,

treelines declined temporarily between 8,300 and 8,000 cal. y

(Dahl and Nesje, 1996; Barnekow, 1999; Barnett et al., 2001).

Using today’s lapse rates, it has been estimated that temperatures

were 1.3–2.0



C higher than today during the period of highest

treeline positions.

A number of pollen profiles from western North America

provide multi-millennial year records of alpine timberline fluc-

tuations. These records were derived primarily from sediments

deposited in high-elevation lakes and bogs, and, in most cases,

limited dating control of the sedimentary records restricts their

application to the reconstruction of only the general trends in

treeline position throughout the Holocene. A more detailed pic-

ture of timberline fluctuations during the last 1,000 years has

emerged from the study of tree rings taken from living trees

or dead snags situated near the alpine timberline in the North

American Cordillera. Most pollen records indicate that the

alpine timberline was higher than present during the early to

mid-Holocene and a general and widespread decline was

underway by ca. 6,000–5,000 cal. y

BP (Hebda, 1995). Pollen-

based estimates of Holocene timberline position have been

corroborated by several complementary proxies for estimating

past environmental and climatic conditions. Taken together,

these data suggest that mean growing season temperature in

western North America was on the order of 1–4



C warmer

than present during the early-mid Holocene.

Perhaps the most direct and unequivocal evidence for past

high stands of alpine timberline in western North America is

Figure H9 The approximate Holocene timberline and treeline elevation (m) in the Swiss central Alps based on radiocarbon-dated macrofossil and

pollen sequences (modified from Tinner and Theurillat, 2003).

444 HOLOCENE TREELINE FLUCTUATIONS

radiocarbon-dated macrofossils that have been preserved in

alpine environments. In some cases these data are derived from

abundant conifer needles preserved in lacustrine sediments at

sites that are currently above natural treeline (Reasoner and

Hickman, 1989) and in others they are obtained from logs

(megafossils) preserved in alpine bogs or soil profiles (Kearney

and Luckman, 1983; Carrara, et al., 1991). Similarly, fossil

wood dating to the early Holocene has been recovered from

retreating termini of a number of glaciers in western Canada

(Luckman, 1988), which indicates that early Holocene forests

occupied areas now covered by ice. The recovery of wood ran-

ging in age from ca. 10,000–5,000 cal. y

BP from sites situated

100–200 m above the present timberline clearly indicates that

subalpine forests reached at least this elevation during the early

Holocene. Reasoner et al. (2001) compiled early-mid Holocene

macrofossil information from across southwestern Canada and

plotted this information relative to modern timberline. The

results indicate two clear periods of higher-than-present timber-

line during the early Holocene: an early phase from the onset of

the Holocene to approximately 8,200 cal. y

BP (7,500

CyBP)

and a later phase between 7,300 and 5,700 cal. y

BP (6,400

and 5,000

CyBP, Figure H10). These data are in general

agreement with European records (e.g., Figure H9) that indi-

cate the timberline was higher than present during the early-

mid Holocene and suggest the 8.2 event may have also

impacted alpine treelines in western North America. Additional

high-resolution records for the North American Cordillera are

necessary to resolve this issue.

Several small-scale fluctuations of the alpine timberline

have been documented during the last millennium, primarily

from dendrochronological analyses of living trees near alpine

timberline and in situ dead snags. For example, the abrupt syn-

chronous termination of 92 tree-ring records related to the

1690’s “cold snap” (Luckman, 1994) clearly indicates that sig-

nificant changes to the position of the alpine timberline have

occurred on decadal timescales and that these changes have been

driven by natural variability of the regional climate (Figure H11).

Linking the past to the future

It is clear that the position and composition of the timberline

ecotone has been sensitive to Holocene climate change. The

millennial-scale trends generally reflect gradual decline in

Northern Hemisphere insolation from approximately 11%

higher-than-present in the early Holocene. Superimposed on

this long-term trend are higher-frequency fluctuations related

to changes in oceanic circulation, volcanic activity, and solar

Figure H10 The approximate minimum early-mid Holocene treeline elevation in the southern Canadian Cordillera, based on radiocarbon-dated

wood macrofossils recovered from high-elevation sites (modified from Reasoner et al., 2001).

Figure H11 Length of record preserved in calendar-dated snags

recovered from the Athabasca Glacier Site, Central Canadian Rockies.

The standing and fallen snags occur at and above the present

treeline or where these snags are considerably larger than trees

presently growing at the site. One of the main Little Ice Age

advances of the Athabasca Glacier has been dated at 1714 (modified

from Luckman, 1994).

HOLOCENE TREELINE FLUCTUATIONS 445

irradiance or a combination of these factors. It is also clear that

anthropogenic climate forcing over the next 100 years is likely

to rival or exceed the warmest conditions of the Holocene.

Concerns about global warming involve adjustments, collapses,

migrations, or extinctions of boreal and alpine life. Surpris-

ingly, relatively few studies have addressed past responses of

ecosystems such as treeline communities to climatic change.

One of the reasons for avoiding this topic is that accurate

studies require independent climatic proxies and very high tem-

poral resolution (<10–20 years/sample). To assess how tree-

lines could respond to global change, high resolution studies

including macrofossil analysis are urgently needed.

Mel A. Reasoner and Willy Tinner

Bibliography

Arno, S.F., and Hammerly, R.P., 1993. Timberline - Mountain and Arctic

Forest Frontiers. Seattle: The Mountaineers, 304pp.

Barnekow, L., 1999. Holocene tree-line dynamics and inferred climatic

changes in the Abisko area, northern Sweden, based on macrofossil

and pollen records. Holocene, 9, 253–265.

Barnett, C., Dumayne-Peaty, L., and Matthews, J.A., 2001. Holocene cli-

matic change and tree-line response in Leirdalen, central Jotunheimen,

south central Norway. Rev. Palaeobot. Palynol., 117,119–137.

Birks, H.H., 2001. Plant macrofossils. In Smol, J.P., Birks, H.J.B., and

Last, W.M. (eds.), Tracking Environmental Change Using Lake Sedi-

ments. Dordrecht: Kluwer, 49–74.

Carrara, P.E., Trimble, D.A., and Rubin, M., 1991. Holocene treeline

fluctuations in the Northern San-Juan Mountains, Colorado, USA, as indi-

cated by radiocarbon-dated conifer wood. Arct.Alp.Res., 23, 233–246.

Dahl, S.O., and Nesje, A., 1996. A new approach to calculating Holocene

winter precipitation by combining glacier equilibrium-line altitudes and

pine-tree limits: A case study from Hardangerjokulen, central southern

Norway. Holocene, 6, 381–398.

Haas, J.N., Richoz, I., Tinner, W., and Wick, L., 1998. Synchronous Holo-

cene climatic oscillations recorded on the Swiss Plateau and at timber-

line in the Alps. Holocene, 8, 301–309.

Hebda, R.J., 1995. British Columbia vegetation and climate history

with a focus on 6 ka BP. Géographie physique et Quaternaire, 49,

55–79.

Hormes, A., Müller, B.U., and Schlüchter, C., 2001. The Alps with little

ice: Evidence for eight Holocene phases of reduced glacier extent in

the Central Swiss Alps. Holocene, 11, 255–265.

Kearney, M.S., and Luckman, B.H. 1983. Holocene timberline fluctuations

in Jasper National Park, Alberta. Science, 221, 261–263.

Körner, C., 1999. Alpine Plant Life. Berlin: Springer, 338pp.

Kullman, L., 1995. Holocene Tree-Limit and Climate History from the

Scandes-Mountains, Sweden. Ecology, 76, 2490–2502.

Luckman, B.H., 1988. 8000 year old wood from the Athabasca Glacier,

Alberta. Can. J. Earth Sci., 25, 148–

151.

Luckman, B.H., 1994. Evidence for climaticconditions between ca. 900–1300

A.D. in the Southern Canadian Rockies. Clim. Change, 26, 171–182.

Maisch, M., Wipf, A., Denneler, B., Battaglia, J., and Benz, C., 1999. Die

Gletscher der Schweizer Alpen. Gletscherhochstand 1850, Aktuelle

Vergletscherung, Gletscherschwund-Szenarien. Zürich: vdf, 373pp.

Reasoner, M.A., and Hickman, M., 1989. Late Quaternary Environmental-

Change in the Lake Ohara Region, Yoho-National-Park, British-

Columbia. Palaeogeogr. Palaeoclimatol. Palaeoecol., 72, 291–316.

Reasoner, M.A., Davis, P.T., and Osborn, G., 2001. Evaluation of proposed

early-Holocene advances of alpine glaciers in the North Cascade Range,

Washington State, USA: Constraints provided by palaeoenvironmental

reconstructions. Holocene, 11, 607–611.

Tinner, W., and Kaltenrieder, P., 2005. Rapid responses of high-mountain

vegetation to early Holcene environmental changes in the Swiss Alps.

Journal of Ecology, 93(5): 936–947.

Tinner, W., and Theurillat, J.-P., 2003. Uppermost limit, extent, and fluc-

tuations of the timberline and treeline ecocline in the Swiss Central

Alps during the past 11,500 years. Arct., Antarct., Alp. Res., 35,

158–169.

Wick, L., and Tinner, W., 1997. Vegetation changes and timberline fluctua-

tions in the Central Alps as indicator of Holocene climatic oscillations.

Arct. Alp. Res., 29, 445–458.

Cross-references

Dendroclimatology

Holocene Climates

Hypsithermal

Macrofossils

Palynology

Pollen Analysis

The 8,200-Year b.p. Event

HUMAN EVOLUTION AND CLIMATE CHANGE

Introduction

How, where, and why humans evolved are fundamental ques-

tions that capture our attention and spark curiosity. A number

of recent hominin fossil discoveries as old as 6 Ma and

advances in DNA research have now documented that humans

evolved in Africa and then migrated to other parts of the

world starting as early as 2.0 Ma. Hominins are living or fossil

members of the family Hominidae. Evidence of their early exo-

dus from Africa has raised other questions. Why did humans

leave the safety of trees? What were the drivers that may have

nudged hominins toward bipedalism and developed species

that failed, while only one ultimately succeeded? Were the

development of tools, the exodus from Africa, and the brain

development that led to language, art and imagination a passive

or direct response to some paleoenvironmental stresses?

During the last few decades, the need to understand the

Earth’s climate system and evaluate the potential threat of glo-

bal warming has yielded a wealth of knowledge on climate

and paleoclimate. This research has led to the exciting potential

of GCMs (global climate models) – computer models that

facilitate the analysis of present, past, and future climates.

However, understanding the paleoenvironmental and paleocli-

matic context of human evolution can only be accomplished

using an interdisciplinary approach integrating geologic, bio-

logic, and anthropologic data and the knowledge of global

climate change.

The hominin fossil record reaches back in time to 6 mil-

lion years ago and stone tools to 2.5 million years ago. The

general framework of global climate during the Plio-Pleisto-

cene is well documented from the oxygen isotope record from

deep sea cores, but the details of paleoclimate on the continents

is just now starting to emerge. There have been a few attempts

to synthesize within regions and to place key sites within a

broader temporal and spatial context, e.g., Turkana, northern

Kenya (Brown and Feibel, 1991; Feibel et al., 1991), Ologesai-

lie, southern Kenya (Behrensmeyer et al., 2002) and Olduvai

Gorge, northern Tanzania (Hay, 1976; Ashley and Hay,

2002). However, most geological studies of hominin-bearing

deposits have typically focused on correlation and chronology

of sites, with usually only local site-based descriptions accom-

panying each find. With the growing realization of the impor-

tance of tectonism, volcanism, and astronomic climate forcing

during the Plio-Pleistocene, there is now an opportunity to inte-

grate what is known about the paleoenvironment and paleocli-

446 HUMAN EVOLUTION AND CLIMATE CHANGE

mate from the geologic record and the evolutionary history

from archaeological and paleo-anthropological records. A num-

ber of scientists have attempted to relate climate to hominin

evolution, with some success (Brain, 1981; Behrensmeyer,

1982; Boaz and Burckle, 1983; Foley, 1994; deMenocal,

1995; Potts, 1996b; Feibel, 1997), and a series of international

workshops and conferences have been held to address the pro-

blem (Vrba et al., 1995; Bromage and Schrenk, 1999).

In order to examine the relationship between climate and

human evolution, the following will:

1. Review the phylogeny (family evolutionary record) of the

hominins

2. Describe the evolution of aspects of hominins that are dis-

tinctively “human”

3. Summarize briefly what is known about global climate dur-

ing the period of hominin evolution

4. Examine possible links between climate change and human

evolution

Hominin family tree

Despite the meager fossil record of hominins prior to appear-

ance of modern humans (probably less than a few thousand

individuals total) the main pattern of the hominin family tree

is evident (Figure H12). Prior to about 4 million years ago, the

record is so poor and the connections between fossils so uncer-

tain that each hominin fossil has been given its own genus

name: Sahelanthropus tchadensis, 6–7 million years from

Chad (Brunet et al., 2002); Orrorin tugenensis, 6 million years

from Kenya (Pickford and Senut, 2001; Senut et al., 2001);

Ardipithecus kadabba, 5.77–5.54 million years from Ethiopia

(Haile-Selassie et al., 2004); and Ardipithecus ramidus,

4.4 million years from Ethiopia (White et al., 1994). The fact

Figure H12 The graph traces the trend of climate during the Cenozoic era, showing a general cooling and culminating with the ice ages of the

Pleistocene epoch. The relative temperatures are interpreted from the oxygen isotope record of foraminifers found in marine sediment records.

Major events in primate evolution are indicated on the right (Tattersall, 1993) (modified from Partridge et al., 1995).

HUMAN EVOLUTION AND CLIMATE CHANGE 447