Odekon M. Encyclopedia of paleoclimatology and ancient environments

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Cross-references

Atmospheric Evolution, Earth

Deep Sea Drilling Project (DSDP)

Loess Deposits

Monsoons, Quaternary

Mountain Uplift and Climate Change

Ocean Drilling Program (ODP)

Plate Tectonics and Climate Change

Precession, Climatic

Pre-Quaternary Milankovitch Cycles and Climate Variability

Varved Sediments

MONSOONS, QUATERNARY

Definition

The word “monsoon” is from the Arabic word “mausim,”

meaning season of winds, which is often applied to the seasonal

reversals of the wind direction along the shores of the Indian

Ocean and surrounding regions, especially in the Arabian Sea.

The prevailing winds blow from the southwest during half of

the year and from the northeast during the other (Figure M39).

Monsoons are driven by the land-sea thermal contrast as the

continent heats up and cools down faster than the ocean. Thus,

in summer, land reaches a higher temperature than the ocean and

the hot air over the land tends to rise, creating an area of persis-

tent low pressure. This produces an extremely constant wind

blowing toward the land. Associated rainfall is generated as the

moist ocean air is diverted upward by mountains, which causes

cooling, and in turn, condensation. In winter, the land cools off

quickly, but the ocean retains heat longer. The warmer air over

the ocean rises, creating a low pressure area and a breeze from

land to ocean. Because the temperature difference between the

ocean and land is less than in summer, the winter monsoon winds

are weak and variable. Monsoons are similar to sea breezes, but

they are much larger in extent, stronger, and are more constant.

Monsoon systems

As monsoons have become better understood, the definition

has broadened, indicating climatic systems in which the preci-

pitation is largely confined to the warm season. Monsoons

Figure M39 Asian monsoon circulation marked by the southwest or

summer and northeast or winter monsoon winds. The Indian landmass

receives maximum solar insolation during the summer monsoon

months (June–September).

MONSOONS, QUATERNARY 589

now include almost all of the phenomena associated with the

annual weather cycle within the tropical and subtropical parts

of Asia, Africa, Australia, North America and the adjacent seas

and oceans.

The southwest or summer monsoon occurs from June to

September, marked by wet and strong winds, providing maxi-

mum rainfall to the Asian landmass. The northeast or winter

monsoon season ranges from December to early March, during

which time the winds are dry and weak, and precipitation is

reduced over the land. However, Australia and Southeast Asia

receive enough rainfall during the winter monsoon. The North

American monsoon occurs from mid-July to September, affect-

ing mainly southwestern parts of the United States. The Indian

or Asian monsoon, which affects South and Southeast Asia as

well as northern Africa, is the best known and most regular

of the world’s monsoons, and one of the most spectacular fea-

tures of the tropical climate system, affecting 60% of world

population between 30



N and S latitudes.

The Indian monsoon system

The Indian monsoon system is believed to have evolved in the

late Oligocene or early Miocene, and appears to have had a

major shift in its intensity and seasonal duration 8.5 Ma

BP.

(e.g., Gupta et al., 2004). Rapid changes in the monsoon in

the late Quaternary and the Holocene are documented in

numerous papers based on proxies from marine records of the

summer monsoon winds (Schulz et al., 1998; Gupta et al.,

2003), cave deposits that record precipitation in Oman (Neff

et al., 2001; Fleitmann et al., 2003), peat deposits that indicate

humidity and temperature (Hong et al., 2003), runoff in the Bay

of Bengal (Kudrass et al., 2001), and fluvial sediments indicat-

ing the intensity of rainfall (Sharma et al., 2004; also

Figure M40). Strong monsoon events had a potentially dra-

matic effect on the fluvial systems (Goodbred and Kuehl,

2000), the terrestrial and marine fauna and flora, and human

populations in Asia during the Holocene (past 10,000 years)

(Gupta et al., 2003; Gupta, 2004).

Across the Asian region, the effects of the monsoon are pre-

served in numerous proxies that include tree rings, soils, ice,

lake deposits, peat deposits, cave deposits and marine sedi-

ments. Paleoproxies provide evidence of slow changes in the

monsoon related to solar insolation, and evidence of abrupt,

multi-decadal to millennial scale events throughout the Qua-

ternary and the Holocene.

Quaternary changes

As the climate entered the glacial-interglacial mode during the

Quaternary, the Indian monsoon system also underwent parallel

changes (e.g., Schulz et al., 1998; Leuschner and Sirocko,

2003). For instance, the intervals of cold spells in the North

Atlantic have been found to be aligned with intervals of weak

summer monsoon (e.g., Schulz et al., 1998), during which time

the winter monsoon strengthened (Fontugne and Duplessy,

1986). The summer monsoon oscillated with millennial–scale

variability (the Dansgaard-Oeschger (D-O) and Heinrich

events), concentrated at periodicities of 1,100, 1,450, 1,750

and 2,300 yr during the last glacial (Naidu and Malmgren,

1995; Sirocko et al., 1996). This pattern is almost similar to

that of changes in the Greenland ice cores (Schulz et al.,

1998; Leuschner and Sirocko, 2003) (also Figure M40).

Despite numerous assertions, the origins of millennial-scale

(sub-Milankovich or sub-orbital scale) abrupt changes in the

monsoon remain elusive. Conceptual and numerical models

indicate that these climate changes could originate in the tro-

pics (e.g., Hoerling et al., 2001) or in the North Atlantic (Bond

et al., 2001). Changes in tropical sea surface temperatures over

the Indian and Pacific Oceans are capable of producing

changes in the North Atlantic Oscillation in coupled climate

model simulations, resulting in warmer temperatures associa-

ted with the positive phase of the North Atlantic Oscillation

(Hoerling et al., 2001). Other changes linked to the hydro-

logical cycle, such as changes in the quantity of water vapor

in the atmosphere, or vegetation-driven greenhouse gases (Stott

et al., 2002), could produce a cooling effect that is amplified at

high latitudes. Alternately, the North Atlantic could drive the

observed changes in the tropics (Bond et al., 2001).

Holocene variability

Repeated occurrences of millennial-scale weak and strong

phases of the Indian summer monsoon characterize the

Holocene record (Figure M41 ). The summer monsoon was

stronger in the early Holocene, which is evident in the enorm-

ous Ganges-Brahmaputra sediment discharge (Goodbred and

Kuehl, 2000), rapid speleothem growth reported in Oman

(Burns et al., 2001) and the dominance of conifers and broad-

leaved trees in the Central Higher Himalayas (Phadtare, 2000).

During the period between 10,400 and 5,500 cal yr

BP., North-

ern Hemisphere temperatures peaked and the Indian, Southeast

Asian and African monsoons reached their maximum – the so-

called “Holocene Climatic Optimum.”

The early Holocene monsoon maximum was interrupted by

an abrupt cooling peak 8,200 cal yr

BP. (Alley et al., 1997),

during which time the summer monsoon over the Indian sub-

continent and tropical Africa weakened (Gasse, 2000). Succes-

sive shifts towards drier conditions in northern Africa and Asia

intensified after 5,500 cal yr

BP. (Overpeck et al., 1996;

Gasse, 2000), which contributed to the termination of ancient

civilizations in the region. Sustained drought due to aridifica-

tion led to societal collapse of the Egyptian and Mesopotamian

civilizations at around 4,200 cal yr

BP. (Weiss et al., 1993;

Cullen et al., 2000). The evaporation dropped during the arid

phase in India (Sharma et al., 2004), and the Indus Valley civi-

lization transformed from an organized urban phase to a post-

urban phase of smaller settlements with southward migration

of the population (Allchin and Allchin, 1997; Staubwasser

et al., 2003; Gupta, 2004).

The summer monsoon was weakest from 2,500 to 1,500

cal yr

BP. in the late Holocene (Anderson et al., 2002; Gupta

et al., 2003). During the Medieval Warm Period (

AD. 900–

1400), the summer monsoon strengthened (Gupta et al., 2003)

whereas during the most recent climatic event, the Little Ice

Age (

AD. 1450–1850), there was a drastic reduction in the

intensity of the summer monsoon (Gupta et al., 2003).

Centennial to millennial-scale changes and forcing

mechanisms

Previous studies have demonstrated that the interannual varia-

bility of the Indian summer monsoon is controlled by the

Eurasian snow cover (Bamzai and Shukla, 1999) and the

amplitude and period of El Niño Southern Oscillation (ENSO)

(Krishna Kumar et al., 1999). Long term variability in the

Indian summer monsoon on millennial scales may be attributed

to solar forcing (Neff et al., 2001; Fleitmann et al., 2003) and

590 MONSOONS, QUATERNARY

glacial interglacial boundary conditions (Burns et al., 2001).

Major monsoon fluctuations on decadal to centennial scales

during the Holocene have been attributed to changes in surface

boundary conditions (Overpeck et al., 1996) and the North

Atlantic climate due to changes in the North Atlantic Deep

Water production (Gupta et al., 2003), and solar activity (Neff

et al., 2001; Gupta et al., 2005).

The extent/duration of the snow and ice cover has been sug-

gested as a cause of monsoon variability due to the albedo feed-

back of snow and ice-fields (Shukla, 1987). Increased albedo of

snow delays the seasonal heating cycle over Asia and thus delays

and weakens the summer monsoon. Previous workers have also

suggested that fluctuations in the monsoon intensity may be

forced by variations in the tropical land-surface boundary condi-

tions (Gasse and van Campo, 1994). Increased summer monsoon

strength may lead to denser vegetation cover in Central Asia that

reduces surface albedo of the region, increasing the land-sea ther-

mal contrast and thus the summer monsoon strength.

Stuiver and Braziunas (1993) suggested that solar variability

could be an important factor affecting climate variation during

Figure M40 Summer monsoon proxy record for the late quaternary from the Arabian Sea. (a) Marine total organic carbon (TOC) record of

SO90–111KL off Pakistan, northeastern Arabian Sea (Schulz et al., 1998); (b) G. bulloides percentages in ODP Hole 723A; (c) GISP2 d

O record

(Grootes et al., 1993). B1–12: Bond events; H1–14: Heinrich events; D1–10: D-O events.

MONSOONS, QUATERNARY 591

the Holocene on the basis of analysis of

C records from tree

rings. The production rates of cosmogenic nuclides (

C and

Be) that reflect changes in solar activity appear to be closely

correlated to Bond Cycles (Bond et al., 2001). Although the

exact process by which global climate and solar forcing may

be linked is not clearly understood, variations in ultraviolet

radiation and cosmic ray flux may trigger abrupt climate

changes by altering the heat budget of the stratosphere

Figure M41 Visual correlation between proxy records. (a) G. bulloides percent from ODP Hole 723A, Arabian Sea (Gupta et al., 2003); (b)

hematite % in core MC52 in the North Atlantic (Bond, et al., 2001); (c) oxygen isotope record from Dongge Cave, southern China (Wang et al.,

2005); (d) Hongyuan peat deposit (Hong et al., 2003), (e) Qunf Cave d

O (Fleitmann et al., 2003), and (f) Hoti Cave d

O (Neff et al., 2001). Grey bars

highlight the Bond events 0–7 (Bond, et al., 2001).

592 MONSOONS, QUATERNARY

(van Geel et al., 1999) and changing the atmosphere’s optical

parameters and radiation balance (Kodera, 2004). The summer

monsoon intensity has been positively related to changes in

solar insolation not only at the Milankovitch scale (Clemens

et al., 1991; Leuschner and Sirocko, 2003), but also at decadal

and centennial scales, as documented by various recent studies

(Wang et al., 1999; Neff et al., 2001; Gupta et al., 2005). For

example, Gupta et al. (2005) observed a relation between inter-

vals of low sunspot activity and low intensity of the summer

monsoon during the Holocene.

Solar variability may lead to changes in the circulation pat-

tern of the atmosphere and may amplify the variability of the

oceanic circulation. Model results have explained the mechan-

ism of the dynamic response of atmospheres to solar forcing

(Shindell et al., 2001). A recent study from the Arabian Sea

suggests that even small changes in solar output can cause

pronounced changes in the tropical climate, emphasizing the

importance of a sun-monsoon link (Gupta et al., 2005).

Conclusions

Summer monsoon variability has important implications on the

human population of the Asian-African region as the summer

monsoon winds bring rainfall to most parts of the region, thus

enhancing agricultural productivity. During the early Holocene,

when the summer monsoon intensity was high, several Asian

rivers were in their full vigor, and plants and animals were

domesticated in the Indian subcontinent along the banks of

the River Indus and its tributaries. However, prolonged

droughts after the beginning of the arid phase (5,000 cal yr

BP.) in the Indian subcontinent led to population dislocation,

widespread migration, and state collapse of the Harappan Civi-

lization. Some people resorted to modifying their dwelling

environments by developing new strategies to optimize the uti-

lity of available water rather than migrating to safer places. It

appears that the hydrological changes in South Asia had the

greatest effect on human settlements, as opposed to the theory

that temperature changes controlled the migration of the popu-

lation. Intervals of weak summer monsoon and cold spells in

the North Atlantic coincided with low sunspot activity during

the Holocene, indicating that the response of both the tropics

and high latitudes to solar variability was direct. The monsoon

could be sensitive to relatively small changes in solar forcing.

The sun-monsoon link can be explained by a direct solar influ-

ence on the Intertropical Convergence Zone that controls mon-

soonal precipitation. It can thus be concluded that solar

variability plays an important role in driving multidecadal to

millennial-scale changes in monsoon climate.

Anil K. Gupta

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Cross-references

Ancient Cultures and Climate Change

Dansgaard-Oeschger Cycles

Deltaic Sediments, Climate Records

Heinrich Events

Hypsithermal

Millennial Climate Variability

Monsoons and Monsoons Climate in Encyclopedia of World Climatology

Monsoons, Pre-Quaternary

North Atlantic Oscillation (NAO) Records

Paleo-El Niño-Southern Oscillation (ENSO) Records

Sun-Climate Connections

MORAINES

A moraine is a glacial landform created by the deposition or

deformation of sediment by glacier ice. Many different types

of moraine exist, reflecting the many different processes

by which glaciers deposit and deform sediment and the many

locations and environments within the glacier system where

deposition can occur. Moraines are composed of till, which is

also highly variable, as its characteristics depend on the charac-

teristics of the debris supplied by the glacier as well as on the

processes and environment of glacial deposition.

The morphology and sedimentology of moraines can be

used to reconstruct the characteristics of former glaciers. Their

distribution reflects the geography of former glaciers and gla-

cial process environments. Sediment characteristics reflect the

source location of the debris: supraglacial debris is characteris-

tically angular, while basally derived debris is typically

facetted, subrounded and striated. Structures within moraines

can reveal seasonal and long term variations in processes of

sedimentation.

Moraines are classified both genetically according to the

process by which they are created and geographically accord-

ing to their position within the glacier system. The principal

processes by which moraines are created are the release of deb-

ris from ice by meltout and the deformation of proglacial or

subglacial sediments by ice motion. Supraglacial moraines

include lateral moraines, medial moraines and inner moraines.

Subglacial moraines occur parallel or perpendicular to ice

flow or in irregular patterns. Ice marginal moraines occur

around the edges of glaciers and are defined by their position

as either lateral or frontal moraines. Moraines marking the

maximum extent of a glacial advance are referred to as terminal

moraines. Moraines deposited at successive positions of the

margin during a period of progressive retreat are recessional

moraines. Marginal moraines at existing glaciers are typically

ridges of sediment resting partially on the edge of the glacier.

Upon deglaciation, ice-cored moraines lose their ice support

and therefore tend to shrink in size and may become structu-

rally unstable. Marginal moraines may be several tens of

meters in height, tens or hundreds of meters across, and may

stretch for hundreds of kilometers around the margins of large

ice sheets.

Moraines provide long-term storage within the glacier sedi-

ment transfer system, and a supply of debris to the proglacial

zone. Sediment flux within glaciated basins is very sensitive

to the position of glaciers relative to their moraines. When gla-

ciers lie behind marginal moraines the bulk of sediment pro-

duced at the margin can go into storage in the moraine belt

and not reach the proglacial region. When glaciers have no

marginal moraines, sediment passes directly into the proglacial

system. Moraines can also focus meltwater discharge, localiz-

ing fluvial processes and causing moraine-dammed lakes.

These lakes are potentially unstable and pose a serious threat

of catastrophic flooding. Bennett and Glasser (1996) and Benn

and Evans (1998) provide a useful review of this topic.

Peter G. Knight

Bibliography

Benn, D.I., and Evans, D.J.A., 1998. Glaciers and Glaciation. London,

UK: Arnold, 734pp.

Bennett, M.R., and Glasser, N.F., 1996. Glacial Geology. Chichester, UK:

Wiley, 364pp.

Cross-references

Glacial Geomorphology

Glacial Sediments

Tills & Tillites

594 MORAINES

MOUNTAIN GLACIERS

Mountain glaciers are constrained by the topography of the

underlying bedrock. They can assume a wide variety of forms

such as cirque glaciers and valley glaciers, which generally

have unidirectional flow, and ice caps, which usually sit atop

peaks and flow radially from central summit locations. Often,

ice caps contain outlet glaciers at their bases. Mountain glaciers

occupy the upper altitudes of many ranges on all the continents

except Australia (although they also occur in New Zealand and

New Guinea), and are most extensive on the Tibetan Plateau, in

Southern Alaska, and in the South American Andes. In fact,

70% of all the world’s tropical mountain glaciers are located

in the Peruvian Andes. Mountain glaciers around the Gulf of

Alaska cover an area of 32,400 square miles, about 13% of

the world’s ice coverage, and include some of the largest gla-

ciers outside the polar regions. In general, the closer to the

equator a glacier is located, the higher the elevation, since the



C isotherm (temperature at which ice ablates) increases in

altitude with decreasing latitude. Thus, Alaskan and high Cana-

dian mountain glaciers currently exist at altitudes as low as

3,000 meters above sea level (m a.s.l.), while some mountain

glaciers near the equator, such as those on Huascarán in Peru



S) and Kilimanjaro in Tanzania (3



S), can only survive at

altitudes greater than 5.500 m a.s.l. In the tropics, these glaciers

are vital for the local populations since they are often important

sources of drinking water and hydroelectric power.

Tropical mountain glaciers often occur in regions that are

affected by strong seasonal variations in precipitation (for

example, monsoon environments), thus wet and dry seasons

are recorded as easily discernable layers. Because these glaciers

are relatively small (compared with the polar ice sheets) and lie

close to the 0



C isotherm, they can respond quickly to varia-

tions in temperature and effective moisture, and thus serve as

reliable “barometers” of climate change. For these reasons, in

the last 20 years mountain glaciers, particularly ice caps, have

become valuable sources of information about past climate

change and thus have placed the current global climate change

in perspective. The length of the climate record contained in a

mountain glacier is constrained primarily by two things: the net

accumulation of snow each year, and the temperature of the

glacier at its base. A glacier that contains a climate record that

covers several millennia often (but not always) occurs in

regions where the annual snowfall is relatively low (less than

a meter per year), but more importantly the temperatures are

cold enough that the base of the ice is frozen to the bedrock.

This prevents the removal of annual layers from the bottom,

which is very important since ice is compressed during burial,

and more annual layers are contained in smaller increments

with depth. Mountain glaciers in such environments can be

found in the highest altitudes of the Andes Mountains, the

Tibetan Plateau, the Alaskan and Canadian Arctic region, and

of course in the mountain ranges in Antarctica and Greenland.

Temperate glaciers, in which some of the ice (usually at the

base) is above the pressure melting point, usually contain

shorter records of climatic and environmental variation, and

these occur at lower altitudes and latitudes.

The clearest evidence for the current major climate warming

comes from the tropical glaciers, recorded in both the ice core

records and in the drastic and accelerating retreats of both totalarea

and total ice volume. The recent warming trend has contributed

greatly to the accelerating rates of mountain glacier retreat, espe-

cially in the climatically stable tropics where the average intrasea-

sonal temperature change is usually no greater than 1–2



C. For

example, the Qori Kalis Glacier in Peru now retreats on average

over the last 15 years more than 60 meters per year, more than

10 times faster than during the first 15 years of measurement

starting in 1963. Similarly, the ice on Mount Kilimanjaro, Africa,

which covered ~4.3 square miles in 1912, only covers 0.94 square

mile today. If the current rate of retreat continues, the perennial

ice on Kilimanjaro will disappear within the next 15–20 years,

the first time in the past 11,000 years that Kilimanjaro will be

devoid of ice (Thompson et al., 2002). The documentation of

this can be seen in the two photos in Figure M42 taken in 1912

and in 2000.

Dramatic mountain glacier retreat is also observed at high lati-

tudes. The glaciers in southern Alaska are experiencing very high

rates of precipitation and runoff, up to 400 cm yr

–1

. Recent work

suggests that melting and calving of Alaskan glaciers make up

about one-half of the total global glacial contribution to sea level

rise (Arendt et al., 2002). However, while the glaciers of Alaska

are retreating due to marked warming in this region, some glaciers

in Norway and Sweden which had been advancing because of the

Figure M42 The ice cap on Kilimanjaro in Tanzania (3



S) in 1912 (left) and in 2000 (right).

MOUNTAIN GLACIERS 595

more northerly trajectory of storms that bring higher snowfall have

since 1999 have also begun to retreat as the temperature rise has

started to over shadow the increase in winter snowfall.

The rapid retreat of the world’s mountains glaciers causes

grave concern for three reasons. First, they are the world’s

“water towers,” and their loss threatens water resources neces-

sary for hydroelectric production, crop irrigation, and munici-

pal water supplies for many nations of the world. The ice

fields constitute a “bank account” that was built up over thou-

sands of years, and is drawn upon during dry times to feed and

energize populations downstream. The current melting is cash-

ing in on that rapidly dwindling bank account. Second, these

ice fields contain paleoclimate histories that are unattainable

elsewhere and, as the ice fields melt, the records preserved

therein are lost forever. The records are needed to discern

how climate has changed in the past in these sensitive regions

and to assist in predicting future change. Finally, one of the

important aspects of glacier retreat on a global scale is the

potential contribution that it makes to sea-level rise.

Lonnie G. Thompson

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

contribution to rising sea level. Science, 297, 382–386.

Thompson, L.G., Mosley-Thompson, E., Davis, M.E., Henderson, K.A.,

Brecher, H.H., Zagorodnov, V.S., Mashiotta, T.A., Lin, P.N., Mikha-

lenko, V.N., Hardy, D.R., and Beer, J., 2002. Kilimanjaro ice core

records: Evidence of Holocene climate change in tropical Africa.

Science, 298, 589–593.

Cross-references

Cirques

Cryosphere

Greenhouse (Warm) Climates

Ice cores, Mountain Glaciers

Sea Level Change, Post-Glacial

MOUNTAIN UPLIFT AND CLIMATE CHANGE

Fossil shells in 400 million year old limestone near the arid, 8.8 km

high summit of Mount Everest provide clear evidence for large-

scale uplift of the Himalaya. South of the range crest, the southern

flank of the Himalaya descends to near sea level and receives

several meters of rain from the southeastern monsoon each sum-

mer. To the north, the Tibetan Plateau rises to an average elevation

of 5 km and experiences a semi-arid climate.

The spatial juxtaposition of these contrasting physiographic

regions and their strikingly different climates raise two fundamen-

tal questions: how does mountain uplift affect climate and climate

change; and to what extent does climate change affect the uplift

of mountain belts? Everyday experience indicates that mountains

affect weather and climate. Clouds gather around mountain sum-

mits, rainfall is commonly greater at higher elevations, and preci-

pitation is typically reduced downwind of mountain ranges.

The provocative contention that climate could affect the growth

of mountain ranges (Molnar and England, 1990) has motivated

recent research efforts to explore the potential interactionsbetween

climate, erosion, and tectonics. Numerical models that link geody-

namic deformation with climate and erosion predict that gradients

in climate can drive analogous spatial gradients in erosion, which,

in turn, can modulate where deformation occurs within an orogen

(Willett, 1999; Beaumont et al., 2001). Such reasoning suggests

that the direction from which prevailing winds bring moisture

into a mountain belt could ultimately determine the patterns and

rates of strain and rock uplift within the orogen. Testing of these

predictions has only recently begun, but synthesis of geodetic,

geophysical, geological, climatic, and geomorphic data are begin-

ning to provide insights on the nature, magnitude, and rates of

the interactions that affect the Earth’s surface and its deformation.

Initially, the impact of mountain uplift on climate and cli-

mate change is described. Subsequently the potential effects of

climate on mountain uplift and evidence for such interactions

are discussed.

Mountain uplift and climate

Orographic precipitation

Orographic precipitation refers to the impact that topography

exerts on the distribution of precipitation. In the simplest form,

mountain ranges cause enhanced rainfall on the upwind side of a

range and generate decreased rainfall on the downwind flank.

As air masses move across a landscape, mountains present topo-

graphic barriers that force diversion of the air mass through some

combination of flow over the crest of the mountains and flow

around its flanks. Two competing characteristics of air masses

determine precipitation where moisture-laden air encounters

mountainous topography (Roe et al., 2002). First, if relative

humidityis constant in a staticcolumn of air then the moisture con-

tent decreases as temperature decreases with height (at a rate of

6.5



Ckm

1

). This lowered moisture content promotes decreas-

ing precipitation at higher elevations (Figure M43). Second, as

wind pushes moist air masses toward the flanks of mountains,

the air is forced upslope, causing a convergence of moisture in

the air column. The cooling that occurs as the air is lifted upslope

increases the relative humidity. When combined with moisture

convergence, the cooling leads to saturation and to rainfall of the

moisture in excess of saturation (Roe et al., 2002). Hence, as air

masses rise up the flanks of a mountain range, rainfall commonly

increases (Figure M43). On the downwind side, as descending

air masses warm adiabatically, their relative humidity drops and

rainfall abruptly decreases, thereby creating a rain shadow.

Both the amount of moisture in an approaching air mass and the

height of a given mountain range have finite limits. Two compet-

ing effects serve to determine the rainfall distribution. As more

and more water is wrung out of the air due to orographic rainfall

on upwind flanks of ranges, the reduced amount of remaining

atmospheric moisture reduces the rainfall potential at higher eleva-

tions. In contrast, the steepening of slopes that characterizes the

upper reaches of mountains (Brozovic et al., 1997) causes more

rapid moisture convergence and tends to increase rainfall poten-

tial. The net result is that, for ranges of modest relief (1–2km),

rainfall tends to increase near the range crest, whereas for ranges

with high relief (>3 km) rainfall can decrease on the upper parts

of the windward flanks. The latter effect is well illustrated

in the Himalaya, where the peak monsoonal precipitation of

4–5myr

1

occurs several kilometers below and >10 km to the

south of the range crest (Figure M44). By 10 km north of the

Himalayan crest, monsoonal rainfall has dropped to <1myr

1

and at altitudes exceeding 6.5 km, precipitation is negligible

(Harper and Humphrey, 2003).

596 MOUNTAIN UPLIFT AND CLIMATE CHANGE

As climate changes, the presence of mountains induces non-

linear changes on orographic precipitation. For example, the

typical average atmospheric cooling of 5–6



C during glacial

maxima dictates that air masses will, on average, hold less

moisture. Such a reduction means that the total precipitation

will be reduced overall and the zone of maximum rainfall

may be moved farther from the range crest in ranges with high

relief. As even less moisture creeps over the range crest, rain

shadows can deepen. Although few reliable proxies exist for

rainfall during glacial times, an observed steepening of regional

glacial snowline gradients across high ranges like the Himalaya

(Williams, 1983; Pratt et al., 2002; Burbank et al., 2003)

implies an enhanced orographic effect on precipitation during

glacial times.

Air mass trajectories

Mountain ranges represent physical barriers to air masses blown

toward them. Two behaviors are commonly observed as air masses

encounter topography: some air flows up and over the crest of the

range, whereas some is diverted around the range margins. The

balance between these two effects is determined by the height

and lateral continuity of the range and by wind strength. Even

when flowing over the mountains, much air flow is funneled

throughthe lowest topography. Ranges with isolated high summits

present less effective barriers than do more plateau-like ranges.

Hence, volcanic arcs will cause less diversion than block uplifts.

For example, both the Cascades of Washington and the Sierra

Nevada of California are oriented nearly perpendicular to the

prevailing storm tracks off the Pacific Ocean. Whereas isolated

Figure M43 Orographic precipitation for a range experiencing rock-uplift rates of 2 mm yr

1

.(a) Top: Theoretical predictions of orographic

precipitation with the range crest on the left of each panel. With increasing altitude (moving from right to left), air masses cool as they expand.

If relative humidity (RH) is held constant, the air holds less moisture and produces less rainfall (“cooling only” curve light gray line). Wind

blowing toward a mountain range causes vertical convergence of air masses and increases water content. Convergence and cooling increase RH,

driving condensation and enhancing rainfall (the orographic effect heavy, black line). Bottom: Trends in rainfall with distance from the range

crest affect equilibrium river profiles, such that river incision balances rock-uplift rates. With less rain toward the divide (cooling-only curve), the

river channels are steeper and the height of the range crest is 50% greater than in the absence of orographic feedbacks, which tend to lower

the range crest and channel gradients (heavy black line). Note the nearly 3-fold difference in ridge height between the “cooling-only” case and the

“convergence and cooling” case. (Modified after Roe et al., 2002.)

Figure M44 Example of orographic precipitation in the central Nepalese Himalaya. Orographic lifting caused by monsoonal storms from the Indian

Ocean that impinge on the Himalaya enhances rainfall. The greatest rainfall occurs at ~2,500 m elevation, where the summertime maximum

precipitation reaches 4myr

1

. Loss of moisture as the air masses cross the Himalaya creates a striking rain shadow to the north with only 10%

as much rain. (Modified from Burbank et al., 2003.)

MOUNTAIN UPLIFT AND CLIMATE CHANGE 597

volcanic peaks in the Cascades rise to elevations >3–4km,many

mountain passes lie below 2 km and much moisture funnels

through these passes. The regional patterns of past glacial snow-

lines (a function primarily of precipitation and temperature)

show that, similar to today, moisture was funneled up the major

drainages during glacial climates (Figure M45). Although the

Sierra Nevada also has summits rising to 4km,thepassesalong

its crest are commonly >3 km in elevation, thereby creating a for-

midable barrier across the path of approaching storms and leading

to diversion around the range’s southern and northern ends. Simi-

lar to other ranges forming block-like barriers, the moisture

that does penetrate the Sierra Nevada tends to be funneled along

major river valleys (Burbank, 1991).

The impact of an individual range on regional climates can

alter during climatic changes. As ice sheets accumulate at high

latitudes during glacial times, the average position of the jet stream

and the intertropical convergence zone migrates latitudinally. At

the same time, storm trajectories also change withrespectto ranges

that could potentially block or divert them. As a consequence,

the impact of mountainous topography on precipitation patterns

will change coevally with major climatic variability, as storm

tracks, total atmospheric moisture, and wind strength are altered.

Continentality and fringing ranges

As an air mass flows over ocean to land, it leaves its major

source of evaporative moisture behind. Once over land, the

normal extraction of moisture via rainfall commonly causes a

progressive drying of the air mass, such that downwind regions

will become increasingly dry, unless new moisture sources are

added. Thus, farther inland on large land masses, continentality

increases as the distance from major moisture sources grows.

For example, central Asia’s isolation from major moisture

sources creates continental deserts in mid-latitude regions that

normally experience temperate climates. Aridification of this

region occurred at least 20 Ma ago (Guo et al., 2002) and

was undoubtedly enhanced by the growth of the Himalaya that

blocked northward moisture transport from the Indian Ocean.

The existence of fringing mountain ranges along the margins of

continents can augment the magnitude of continentality. If the

mountains are oriented across major storm trajectories, orography

can efficiently extract moisture in the near-coastal region and cre-

ate residual air masses that are even more depleted in moisture

than they would be in the absence of coastal topography. Thus,

north-south-oriented ranges like the Andes, Cascades, and Sierra

Nevada have a larger impact on downwind continentality than

do east-west-trending ranges like the Unitas, European Alps, or

Pyrenees that are oriented more parallel to prevalent storm trajec-

tories. Even these storm-track-parallelranges enhance continental-

ity, however, because air masses diverted around their flanks travel

a longer path and lose more moisture than they would in the

absence of the range.

Plateau growth and monsoons

Summer monsoons, which can drop several meters of rainfall

over 2 –4 months, are one of the world’s most dramatic climatic

phenomena. The contrast in the heat capacity of water versus

soils results in differential responses to seasonal changes in

insolation. During summer months, terrestrial areas warm more

rapidly than the oceans, causing lowered atmospheric pressure

over the land as air masses warm, expand, and vertically

ascend. The consequent lateral atmospheric flow from sur-

rounding higher pressure zones over the oceans draws water

vapor toward the continental interiors. Once over land, release

of latent heat during precipitative condensation warms the sur-

rounding air even more and creates a positive feedback by low-

ering atmospheric pressure over the land still further and

enhancing the regional pressure gradient and resultant land-

ward flow. Thus, differential heating and release of latent heat

during rain storms can create a monsoon, an intensified sum-

mer precipitation cycle over continents.

Summer monsoons exist over northern Australia, the Ama-

zon, and western Africa, but the Asian monsoon is the best

known and typically the strongest. Much of the strength of

the Asian monsoon derives from the presence of the Tibetan

Plateau. The largest plateau on Earth, the Tibetan Plateau

encompasses  2.5 million km

, an area about one-third the size

of the continental United States, and averages 5 km in elevation

(Fielding et al., 1994). Although summertime temperatures in

Tibet are cool by most human standards, Tibet is the warmest

place on Earth at 5 km elevation. Because soils absorb, retain,

and re-radiate heat much more effectively than does air, the

Tibetan Plateau represents a hot island surrounded by an other-

wise cool atmosphere at 5 km above sea level. This warmth

drives the summertime low-pressure system over southern Asia

and is the pump that primes and modulates the strength of the

Asian monsoon.

Numerical simulations reinforce this contention (Kutzbach

et al., 1993), although definition of a specific altitudinal thresh-

old or areal extent required to trigger a strong monsoon remains

uncertain. Nonetheless, the causative linkage of a high and

wide plateau to development of a strong Asian monsoon under-

pins interpretations that use geologic indicators of the onset of

a strengthened monsoon as predictors of the timing of major

Figure M45 Pattern of glacier snow lines during the past glaciation as

reconstructed from the height of cirques in the northern Cascade

Range, Washington. The re-entrants (dark arrows) in the snowline

elevations (heavy contours) correspond primarily with low topography

along river drainages, whereas the salients (white arrows) correspond

primarily with higher topography. River valleys funnel more moisture to

the interior of the range, thereby lowering snowlines. (Modified after

Porter, 1977.)

598 MOUNTAIN UPLIFT AND CLIMATE CHANGE