Odekon M. Encyclopedia of paleoclimatology and ancient environments

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Sun prior to 3 billion years was offset by higher levels of

greenhouse gases such as carbon dioxide or methane (see Faint

Young Sun Paradox). The lack of siderite in paleosols older

than ~2 billion years has been interpreted to indicate an insuf-

ficient amount of atmospheric CO

to compensate for the faint

young Sun, and that therefore the major greenhouse gas was

. However, Ohmoto et al. (2004) use thermodynamic argu-

ments to conclude that massive Archean siderite deposits were

laid down under a CO

-rich atmosphere. The formation of iron

(II)-rich carbonate (Fe

0.53

0.44

) weathering rinds on con-

glomerate pebbles from the 3.2 billion year old Barberton

Greenstone Belt, South Africa also hints at higher paleo-

atmospheric CO

levels than present (Hessler et al., 2004).

Sulfates

Gypsum and anhydrite

Gypsum (CaSO

·2H

O) and anhydrite (CaSO

) are among

the most abundant evaporite minerals. Gypsum also crys-

tallizes within muds in desert playas or coastal sabkhas

(Figure M33a–c). Gypsum nodules grow directly by displacing

the enclosing mud sediment, coalescing into nodular layers.

Gypsum dehydrates to anhydrite at the higher temperatures and

pressures caused by burial to depths of several hundred meters

(McLane, 1995). Anhydrite precipitates directly from highly

saturated brines at temperatures above 22



C (McLane, 1995).

Mirabilite

Mirabilite (Na

·10H

O) is a soluble evaporite mineral

that precipitates under conditions of high salinity and at low

temperatures.

Barite

Barite (BaSO

) forms in hydrothermal and in sedimentary envir-

onments. It also precipitates from seawater, accumulating on the

ocean floor, often in association with plankton and organic matter

(Derry and Murray, 2004; Kastner, 1999). Furthermore, its insolu-

bility and resistance to diagenetic alteration under oxic conditions

help preserve it in marine sediments. These features make

barite accumulation rates a useful marker for ocean paleoproduc-

tivity (Parrish, 1998; Paytan et al., 1996)(seeOcean paleopro-

ductivity). Oxygen isotope ratios in marine barite record changes

in exposure to sulfur oxidation on continental shelves during

glacial-interglacial cycles and thus indirectly record fluctuations

in sea level (Turchyn and Schrag, 2004).

Phosphates

Francolite, or carbonate-fluorapatite (Ca

(PO

,CO

)

(F,OH), is

the chief mineral in phosphorite, a sedimentary rock rich in

phosphates (see Phosphorite). The chemical composition

of francolite is variable, with differing amounts of Na, Sr,

or Mg replacing Ca, and SO

substituting for PO

. Francolite

occurs as pellets, nodules, crusts, or laminations. Collophane

is an amorphous or cryptocrystalline form of sedimentary

phosphate.

Phosphorites accumulate on continental margins, epeiric

seas, and to a lesser extent on seamounts or ocean islands. They

are generally formed in zones of marine upwelling and high

biologic productivity (Parrish, 1998). The main source of phos-

phorus comes from vertebrate skeletal remains, or breakdown

of organic matter. Fish scales or bones frequently act as nuclei

for the buildup of francolite pellets or nodules.

Figure M33 (a) Gypsum textures typical of evaporitic environments – growth aligned gypsum (left), lensoidal gypsum (right) (after Warren, 1999).

(b) Gypsum rose, Chihuahua, Mexico (4.4 cm  3.2 cm, specimen, V. Gornitz; photograph, John Betts). (c) Gypsum (prismatic, twinned crystals)

and halite (cube) on fine-grained halite, salinas de otomo, Pisco Province, Ica Pagion, Peru (6.4 cm  2.5 cm; specimen, V. Gornitz; photograph,

John Betts).

MINERAL INDICATORS OF PAST CLIMATES 579

Silica minerals

The most common varieties of silica found near the earth’s

surface include quartz, chalcedony, chert, and opal.

Quartz

Low-quartz (a-SiO

) is the stable form of crystalline silica at

the earth’ s surface and one of the most abundant minerals in

crustal rocks. Quartz is a hard (H = 7 Mohs scale), chemically

resistant mineral that survives the rock weathering cycle in

nearly all climates, except for hot and humid tropical climates

(e.g., see Bauxite, above). It is a major constituent of eolian

desert deposits, sandstones, and many littoral beach sands.

In deserts, winds concentrate quartz grains into dunes and

huge sand seas (ergs). Fossil sand dunes show characteris-

tic cross-bedding and wind-ripple forms. The orientation of

cross-bedding and other sedimentological features enable the

reconstruction of paleowind directions (Parrish, 1998). Examples

of quartz-rich eolian sandstones include the Jurassic Navajo

and Kayenta Sandstones of Utah and Colorado, the Permian

Coconino Sandstone of the Grand Canyon region, New Red

Sandstone (Permo-Triassic, England) and the Yellow Sands –

Permian, England (Parrish, 1998). Quartz is also a significant

component of eolian dust in marine sediments and in ice cores.

Variations in concentrations and grain sizes of eolian dust grains

provide information on changes in glacial to interglacial wind

directions and intensities (Eolian dust, marine sediments).

Opal-A, opal-CT, chert, chalcedony

Opal-A is an amorphous, structurally disordered form of silica

which contains several percent of water. Widespread in deep

sea sediments (e.g., in siliceous oozes), it is largely derived

from the remains of marine micro-organisms, such as diatoms,

radiolarians, silicoflagellates, and siliceous sponge spicules

(Parrish, 1998). (However, some marine opal may be of volca-

nic origin). Over time, opal-A gradually transforms to opal-CT,

which shows domains of short-range order similar to the

atomic arrangement in cristobalite and tridymite. Further diag-

enesis leads to opal-C and chert.

Chert is a sedimentary rock consisting of opal, crypto- to

microcrystalline quartz, and chalcedony (a fibrous variety of

quartz). Extensive bedded chert deposits are mostly products

of former marine biogenic sedimentation. Bedded chert is often

associated with phosphorites and black shales (see Phosphor-

ite; McLane, 1995). Such cherts, or siliceous oozes, have

deposited in zones of marine upwelling that are diagnostic of

high organic primary productivity (Kastner, 1999; Parrish,

1998). Tracking shifts in the distribution of upwelling zones

over time may also reveal information about changes in

paleo-ocean circulation.

Clay minerals

Clay minerals are phyllosilicates, or sheet silicates, consisting

of layers of silica tetrahedra attached to layers of aluminum

or magnesium atoms, each of which is surrounded by six oxy-

gen atoms or hydroxyl ions in octahedral layers. Other cations,

such as iron, potassium, sodium, or calcium, can also be pre-

sent. The clays differ in the nature of the cations within and

between sheets and in their stacking arrangements. Important

clay minerals and their ideal chemical formulas are listed

in Table M2.

Clay minerals are produced by the chemical weathering

of rocks near the earth’s surface (Figure M34). The clay

detritus is removed by water erosion and accumulates in lacus-

trine, estuarine, and marine sediments. Clays are also major

constituents of terrestrial soils and of airborne dust. The rela-

tive abundance of clays is closely related to climate or to

environmental setting, although source rock mineralogy also

influences their development. Kaolinite (Figure M34a), for

example, is created by intense chemical weathering of warm,

humid climates, in which silica is leached out, leaving soils

enriched in alumina (Chamley, 1989). Chlorite (Figure M34b)

and illite (Figure M34c) tend to form in soils dominated by

physical weathering – both in colder, often formerly glaciated

regions, and hot, dry climates (Chamley, 1989). Consequently,

these clays generally reflect the composition of their parent

rocks. Smectite (Figure M34d), illite, chlorite, and mixed-

layer clays are more prevalent in temperate-cool, moist cli-

mates (Chamley, 1989). Well-crystallized iron-rich smectites

(e.g., nontronite) occur in sub-arid climates (Chamley, 1989).

Montmorillonite – a smectite – also forms by weathering

of volcanic ash. Sepiolite, palygorskite, and attapulgite are

three rather rare clay minerals found in calcrete soils, and in

warm, semi-arid to arid climates (Chamley, 1989; Parrish,

1998; Watson, 1992).

Glauconite, (K,Na,Ca)(Fe

,Fe

,Mg)(Si,Al)

(OH)

), is a

dark-green to black mineral, related to mica and mixed-layer

clays, usually found as tiny rounded grains that most likely

formed as fecal pellets. It occurs in weakly reducing environ-

ments, near upwelling zones on continental shelves and mar-

gins between depths of 100–250 m (Chamley, 1989),

characterized by relatively slow sedimentation rates (Kastner,

1999; Parrish, 1998; Velde, 2003), and is often associated

with phosphorites (McLane, 1995; Velde, 2003). The presence

of potassium in glauconite allows dating of sea level cycles

in marine sediments, via the potassium-argon or rubidium-

strontium methods (Velde, 2003).

Berthierine, (Fe

,Al)

(Si,Al)

(OH)

, a dark-greenish phy-

llosilicate closely related to kaolinite and serpentine, occurs in

near-shore or deltaic sediments (Velde, 2003) from chemically

reducing environments, but more rarely in terrigenous soils. Its

presence in early Triassic continental paleosols may indicate low

atmospheric oxygen levels, possibly resulting from catastrophic

methane release, mass mortality, and rise in CO

at the Permian-

Triassic boundary (Sheldon and Retallack, 2002).

Inasmuch as non-biogenic marine sediments are largely

derived from continental sources, the distribution of clay minerals

in the oceans closely corresponds to that of adjacent landmasses.

Thus, marine clays display a latitudinal zonation broadly parallel-

ing the major climate zones (Chamley, 1989;Evans,1992;Parrish,

1998). Kaolinite is abundant in equatorial waters, whereas chlorite

Table M2 Common clay minerals

Kaolinite Al

(OH)

Smectite group (expanding clays)

Montmorillonite (Na,Ca)

0.3

(Al,Mg)

(OH)

·nH

Beidellite (Na,Ca)

0.3

(Si,Al)

(OH)

·nH

Nontronite Na

0.3

(Si,Al)

(OH)

·nH

Saponite (Ca,Na)

0.3

(Mg,Fe)

(Si,Al)

(OH)

·nH

Illite KAl

(Si

Al)O

(OH)

Glauconite (K,Na,Ca)(Fe

,Al,Fe

,Mg)(Si,Al)

(OH)

Vermiculite (Mg

2.7

0.3

)(Si

Al)O

(OH)

·nH

Chlorite group (Mg

2.6

0.4

)Si

2.5

(Al,Fe)

1.5

(OH)

Berthierine (Fe

,Al)

(Si,Al)

(OH)

Sepiolite Mg

(OH)

·6H

Palygorskite (Mg,Al)

(OH)

·4H

580 MINERAL INDICATORS OF PAST CLIMATES

and illite concentrations increase toward higher latitudes. The dis-

tribution of smectites, on the other hand, is not quite as strongly

zoned, indicating the influence of other, non-climate factors.

Temporal changes in clay mineralogy record important paleo-

climate events. For example, Kennedy et al., (2006)associateda

significant increase in clay minerals relative to quartz during the

late Neoproterozoic (800–550 million years ago) with a growing

degree of terrigenous chemical weathering, clay-organic matter

burial, and increasing oxygen levels. Kaolinite and expandable

clays (i.e., smectites) became more abundant at the expense of

illite, chlorite, and mica during this period, consistent with devel-

opment of soils, possibly promoted by biotic activity. Variations

in relative abundances of kaolinite, smectite and mixed-layer

illite/smectite have been used as a paleoclimate proxy for wet ver-

sus dry climates during the Mesozoic of northwestern Europe

(Ruffell et al., 2002), with high kaolinite indicating humid periods.

Elsewhere, increases in smectite and kaolinite relative to illite

in marine sediments off Antarctica at around 33.5 Ma have been

correlated with a marked positive dO

shift and onset of signifi-

cant Antarctic glaciation (Robert and Kennett, 1997). Decreasing

kaolinite/illite, smectite/illite ratios after 33.4 Ma signal enhanced

physical weathering under cooler, drier climates. A similar pattern

of higher kaolinite/illite and smectite/illite ratios at ~33.5 and

33.4 Ma, succeeded by increased illite content is observed in cores

from the Great Australian Bight (Mallinson et al., 2003). These

cores also show 22-Kyr cyclicity in kaolinite/smectite ratios,

linked to precessional changes in precipitation, wind patterns

and runoff that affected the mineralogy and transport of clays

to the Bight.

Enhanced chemical weathering during warmer interglacials

generated higher concentrations of interlayered illite-smectite

in Lake Baikal sediments, Russia, whereas greater physical

weathering during colder, glacial periods resulted in chlorite

and illite enrichment (Yuretich et al., 1999). Changes in illite/

chlorite ratios in Chinese loess sections have been used to

reconstruct monsoonal activity during the late Pleistocene

(Zhao et al., 2005). In the southern Loess Plateau, the illite/

chlorite ratio and magnetic susceptibility are higher in inter-

glacial paleosols and lower in glacial loess. More intense

weathering of chlorite, leading to higher illite/chlorite ratios,

indicates a stronger monsoonal regime operating during inter-

glacial periods.

The distinctive mineral assemblage of high illite, low

smectite, and low kaolinite/chlorite ratio in the dust from the

GISP2 (Greenland) ice core pinpoints a likely source of

the clays in deserts of eastern Asia. Although dust composi-

tion (and sources) remained fairly constant during the last gla-

ciation, changes in total concentrations and grain sizes reveal

Figure M34 (a) Kaolinite, Keokuk geode, USA. (b) Iron-rich chlorite, Spiro Sandstone, Arkoma Basin, Oklahoma, USA. (c) Illite, Rotliegend Formation,

Northern Germany. (d) Smectite, Ap horizon, Webster soil, University of Minnesota Southern Agricultural Station, Waseca, Minnesota.

Reproduced with permission from the “Images of Clay” gallery published by the Clay Minerals Group of the Mineralogical Society and the Clay

Minerals Society (www.minersoc.org/ pages/ gallery/ claypix).

MINERAL INDICATORS OF PAST CLIMATES 581

significant variations in wind speed and strength during stadials

and interstadials (Biscaye et al., 1997).

Zeolites

Zeolites (Table M3) are a group of hydrated aluminosilicate

minerals with an open framework structure, into which various

large ions and water molecules can fit and which possess a cer-

tain degree of mobility. They are widespread in cavities in

basaltic lavas, as authigenic or diagenetic minerals in marine

sediments, and in low-grade metamorphic rocks. Phillipsite is

an authigenic zeolite that forms on the seafloor or at shallow

depths (Kastner, 1999) Clinoptilolite forms in silica-rich marine

sediments, derived from opal-A. Zeolites also form from the

alteration of volcanic glass, tuff, or lavas in high pH, alkali-

rich, semi-arid to arid environments. Analcime is common in

sodium-rich alkaline desert soils (Watson, 1992). Zeolites from

saline, alkaline lakes include analcime, chabazite, clinoptilolite,

erionite, mordenite, and phillipsite. These zeolites are often

associated with evaporites (Parrish, 1998). Zeolites in lacus-

trine deposits have been used as markers for dry, alkaline

environments, as for example at Lake Naivasha, Kenya, where

low lake levels are represented by the presence of chabazite

and phillipsite, with even greater aridity indicated by clinoptilo-

lite and analcime (Trauth et al., 2001). The East African lake

level fluctuations have been linked to precessional cycles in

springtime insolation and monsoonal precipitation.

Summary

Minerals that form at or near the earth’s surface in contact with

the atmosphere and hydrosphere can provide important infor-

mation about former climates. They have been used to recon-

struct past climates and changes over time, deduce changes in

atmospheric composition, trace eolian and precipitation pat-

terns, and serve as “hosts” for climate-sensitive stable isotopes

and trace elements.

The most useful mineral indicators or proxies are those that

are generated under relatively narrow climatic ranges or within

restricted environmental settings. Examples include evaporites,

low temperature minerals (e.g., ikaite and hydrohalite), residual

soils (e.g., iron and aluminum oxyhydroxides, kaolinite), and

certain authigenic minerals (that form in situ, e.g., glauconite,

bertherine, phillipsite, barite). Minerals that are sensitive to oxi-

dation (e.g., pyrite and uraninite) reflect past atmospheric oxy-

gen levels. Distinctive detrital minerals in eolian deposits act as

markers of wind direction and intensity. Similarly, minerals in

deltaic sediments point not only to variations in the relative

contributions from source tributaries, but can also indicate pre-

cipitation variability in source areas. Trace elements and stable

isotope ratios contained within minerals such as calcite (both

biogenic and non-biogenic) have provided a wealth of paleocli-

matological information.

Vivien Gornitz

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Table M3 Common zeolite minerals

Analcime NaAlSi

·H

Chabazite Ca

(Al

)·12H

Clinoptilolite (Na,K,Ca

0.5

)

(Al

)·24H

Erionite (Na,K

,Mg,Ca

1.5

)(Al

)·28H

Mordenite Na

KCa

(Al

)·28H

Phillipsite (K,Ca

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(Al

)·12H

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Watson, A., 1992. Desert soils. In Martini, I.P., and Chesworth, W. (eds.),

Weathering, Soils & Paleosols. Amsterdam: Elsevier, Chap.10,

pp. 225–260.

Yuretich, R., Melles, M., Sarata, B., and Grobe, H., 1999. Clay minerals

in the sediments of Lake Baikal: A useful climate proxy. J. Sed. Res.,

69, 588–596.

Zhao, L., Ji, J., Chen, J., Liu, L., Chen, Y., and Balsam, W., 2005.

Variations of illite/ chlorite ratio in Chinese loess sections during

the last glacial and interglacial cycle: Implications for monsoon

reconstruction. Geophys. Res. Lett., 32, L20718. doi:10.1029/

2005GL024145.

Cross-references

Atmospheric Evolution, Earth

Banded Iron Formations and the Early Atmosphere

Beachrock

Carbonates, Cool Water

Carbonates, Warm Water

Cryosphere

Dating, Magnetostratigraphy

Desert Varnish as Climate Proxy

Duricrusts

Eolian Dust, Marine Sediments

Eolian Sediments and Processes

Eolianite

Evaporites

Faint Young Sun Paradox

Glacial Geomorphology

Glendonite/ Ikaite

Ice Cores, Antarctica and Greenland

Ice Cores, Mountain Glaciers

Ice-Rafted Debris (IRD)

Laterite

Marine Biogenic Sediments

Marine Clay Minerals

Ocean Paleoproductivity

Paleosols, Pre-Quaternary

Paleosols–Quaternary

Phosphorite

Red Beds

Sedimentary Indicators of Climate Change

Speleothems

Stable Isotope Analysis

MONSOONS: PRE-QUATERNARY

Monsoon circulation has existed throughout geological history

whenever the tropics were occupied by both land and sea. Pre-

Quaternary monsoons, however, are poorly understood because

paleo-monsoon studies have been heavily biased toward varia-

tions of the Quaternary monsoon on the orbital and suborbital

time scales. Pre-Quaternary monsoon systems are often consid-

ered over a much longer time span, so they show more signifi-

cant variations in responding to tectonically-induced geographic

and topographic changes. Pre-Quaternary monsoons have so far

only sporadically been mentioned in the literature for most of the

geological periods, except for two better studied intervals: the

Permian and Triassic where a “megamonsoon” developed on

the megacontinent Pangaea, and the late Cenozoic, which led to

the establishment of the modern monsoon system.

Megamonsoon of the megacontinent

As monsoons are caused by land-sea contrast in the heating

rate, the ideal conditions for a monsoon system to develop exist

when only one major continent and one complementary ocean

exist on Earth. This was the real case around 250–200 Myr

ago, during the late Permian and Triassic, when all continents

MONSOONS: PRE-QUATERNARY 583

assembled into two major landmasses, Laurasia and Gondwa-

naland, that joined near the equator to form the supercontinent

Pangaea and the super-ocean Panthalassa (Figure M35a). The

extensive distribution of evaporites and many biogeographic

features indicate a maximum of continentality, leading to spec-

ulation about the association of monsoon-type seasonal rains

with large landmasses at that time (Robinson, 1973). Parrish

et al. (1982) used the basic principles of atmospheric and ocea-

nic circulation to reconstruct global paleo-precipitation maps

for seven time intervals of the Mesozoic and Cenozoic, and

found patterns of a strong summer monsoon low over Laurasia

and a winter monsoon high over Gondwanaland in the Triassic

and reversed monsoon features for opposite seasons. According

to these authors, the Triassic was distinguished by maximal arid-

ity with precipitation on the supercontinent provided only by

monsoons (Figure M35b). Kutzbach and Gallimore (1989)were

the first to use numerical modeling to explore the “megamonsoon

of the megacontinent.” Applying a low-resolution general cir-

culation model to the idealized Pangaea (Figure M35a), they

found extreme continentality with a hot summer and cold win-

ter, and large-scale summer and winter monsoon circulations

(Figure M35c,d. At the same time, Crowley et al. (1989) used

Figure M35 Megamonsoon of the Pangaea. (a) The idealized Pangaean continent. Fine dashed lines indicate the approximate outlines of modern

landmasses (after Kutzbach and Gallimore, 1989). (b) Schematic diagram illustrating monsoonal circulation in northern summer. Arrows show

surface winds, stippling indicates heavy seasonal rains (modified from Parrish and Peterson, 1988). (c) Modeled precipitation rate (mm d

1

)on

Pangaea for summer. (d), (e) Modeled surface winds on Pangaea for winter (d) and summer (e), note the seasonal reversal of the wind direction.

The gray bar shows the poleward limit of summer monsoon over land and the Intertropical Convergence Zone over ocean (modified from

Kutzbach and Gallimore, 1989).

584 MONSOONS: PRE-QUATERNARY

energy balance modeling (EBM) to study the late Permian cli-

mate. Both modeling results show an extremely wide annual

range of temperature (50



C) for hinterland and strong mon-

soon circulation over Pangaea.

The geological record of the “megamonsoon” was found

from the upper Triassic of North America: the Newark and

related basins in the east and the Colorado Plateau in the west

(Figure M36a). Late Triassic lacustrine deposits from the New-

ark Basin, New Jersey, which were about 10



N in the mon-

soon-prevalent tropics at that time, display clear evidence of

monsoon-driven seasonality and lake level fluctuations (Olsen,

1986). Part of the sequence consists of micro-laminated mud-

stones (varves) with 0.2–03 mm thin couplets of alternating

light and dark layers, implying significant seasonal contrast

Figure M36 Late Triassic monsoon records from North America. (a) Locations of the Colorado Plateau with the Chinle Formation (CL)

and, to the east, a chain of rifted basins containing the Newark Supergroup (NW). (b) Photograph of microlaminated mudstone showing

organic-rich/ carbonate-rich couplets as annual varves. (c) Lake-level fluctuations revealed in a section of the Newark lake sediments, showing

20-kyr, 100-kyr, and 400-kyr cycles. (d) Average spectral estimates of sediment cycles in the Newark Basin against the modern precession spectrum.

(Modified from Olsen and Kent, 1996 and Ruddiman, 2001).

MONSOONS: PRE-QUATERNARY 585

(Figure M36b). Detailed studies including spectral analyses on

the nearly 7,000 m long section reveal a full range of

precession-related periods of lake level change: the 20-kyr pre-

cession cycles, and 100-kyr and 400-kyr eccentricity cycles

which modulate the amplitude of precessional cyclicity

(Figure M36c,d). All these cycles are characteristic of the tropi-

cal response to orbital forcing and support a monsoon-climate

origin of lake level change, whereas the absence of obliquity

cycles preclude the possibility of direct linkages to high-

latitude climate systems (Olsen and Kent, 1996). Similar

records of the Pangaean “megamonsoon” were also found in

the Chinle Formation, Colorado Plateau, on the tropical wes-

tern margin of the supercontinent (Figure M36a; Dubial et al.,

1991). During the Pangaean interval, eolian, playa deposits

and evaporates were widespread on the Colorado Plateau, but

the Chinle formation represents an unusually wet episode with

well-developed fluvial and lacustrine deposits and paleosol

sequences, indicating abundant moisture brought about by

enhanced monsoon circulation and strong seasonality.

Establishment of the modern monsoon system

The modern Asian-Australian and African monsoons cover

most of the Eastern Hemisphere, and fundamental questions

in Cenozoic paleoclimatology ask when the modern monsoon

system was established and how it has evolved since then.

Three tectonic factors have been proposed to exert a control

over the evolution of Asian monsoon circulation: plateau uplift,

sea-land distribution, and closing of oceanic gateways. Contin-

uous records of the monsoon history provided by Deep Sea

Drilling Project/Ocean Drilling Program (DSDP/ODP) cruises

to the Arabian, Mediterranean and South China Seas, as well

as from the Loess plateau in central China, have been used to

verify the various tectonic hypotheses (Sun and Wang, 2005).

Tectonic forcing and numerical modeling

1. Plateau uplift. GCM experiments on the modern land-sea dis-

tribution indicate that strong monsoons can be induced by solar

forcing only when the elevation of Tibet-Himalaya has reached

at least half that of today (Prell and Kutzbach, 1992). A number

of studies in the late 1980s investigated the climatic conse-

quences of uplift (e.g., Ruddiman et al., 1989) and found that

uplift may have been responsible for both the global cooling

and significant strengthening of the Asian monsoon system in

the late Cenozoic. According to the prevalent hypothesis, uplift

of the Tibetan Plateau intensified around 8 Ma and caused

enhanced aridity over the Asian interior and the onset of the

Indian and east Asian monsoons (Figure M37a, Prell and

Kutzbach, 1997;Anetal.,2001). However, this cognition is

challenged by the new discovery of Miocene loess (Guo et al.,

2002) and other evidence that indicates an older age (see

below).

2. Sea-land distribution. The results of Atmospheric General

Circulation Model (AGCM) simulation by Ramstein et al.

(1997) indicate that the Paratethys, an epicontinental sea

stretching over Eurasia 30 Ma ago, had progressively

receded during the Miocene, resulting in major continentali-

zation of the Asian interior and enhancement of monsoon

circulation. They consider the retreat of the Paratethys Sea

as important as the uplift of the Himalayan/Tibetan Plateau

Figure M37 Records of Asian monsoon and aridity evolution over the past 12 Myr. (a) Globigerina bulloides %(>150 mm), ODP Site 722,

Arabian Sea; (b) Magnetic susceptibility (thin line) and grain size fraction (thick line) from the Loess Plateau (An et al., 2001); (c) Neogloboquadrina

dutertrei %, ODP Site 1146, South China Sea (Wang et al., 2003); (d) Dust flux (mg cm

2

kyr

1

), ODP sites 885 /886, north Pacific (Rea et al., 1998).

586 MONSOONS: PRE-QUATERNARY

for development of the Asian monsoon. As the shrinkage of

the Paratethys Sea occurred during the Oligocene-late Mio-

cene, its effects on the monsoon may have lasted from 30 to

10 Ma, significantly earlier than the 8-Ma date implied by

the uplift hypothesis.

3. Oceanic gateways. An oceanic circulation model reveals

that the “closure” of the Indonesian seaway 3–4 Ma ago could

be responsible for east African aridification (Cane and Molnar,

2001). The northward drift of the Australian Plate may have

switched the source of the Indonesian Throughflow current

fromthe warm South Pacific to the relatively cold North Pacific

waters. This would have decreased Sea Surface Temperatures

(SSTs) in the Indian Ocean and would have subsequently

reduced precipitation over east Africa, as well as the

overall strength of the Indian summer monsoon as recorded

in marine deposits.

All of these three factors are believed to be significant in the

development of the modern monsoon, but their relative roles

remain unclear. To single out the role played by each of the fac-

tors, many more long-term sequences and better constraints on

the timing of the tectonic and climate events are needed.

Geological records

Studies on the long-term evolution of the Asian monsoon started

with ODP Leg 117 to the Arabian Sea in 1987. A sudden increase

in the cool-water planktonic foraminifer Globigerina bulloides in

sediment cores of Leg 117 around 8.5 Ma ago was considered

by Kroon et al. (1991) as indicating the onset of monsoon-related

upwelling (Figure M37a). This date is very close to the rapid eco-

logical transition from C3-dominated to C4-dominated vegetation

around 7.4–7.0 Ma, as revealed by the d

C data of pedogenic car-

bonates from northern Pakistan in the Himalayan foreland. These

were interpreted as evidence for the origin or intensification of

the Asian monsoon system (Quade et al., 1989). In addition, the

dating of the extensive faults on the Tibetan Plateau suggests a sig-

nificant uplift/extension period at about 8 Ma. On the basis

of these findings and GCM simulations, Prell and Kutzbach

(1992) hypothesized that uplift of the Tibetan Plateau to at least

half of its present height at ~8 Ma caused the intensification of

the Asian monsoon.

In the Chinese Loess Plateau, the base of the loess-paleosol

sequences dated to about 2.6 Ma was previously taken as indicat-

ing the initiation of the East Asian monsoon (Liu and Ding, 1993).

Later, Chinese scientists found that the Red Clay underlying the

loess sequence was also of wind-blown origin and indicative

of monsoon transport (Table M4), so the history of eolian deposits

of the Loess Plateau should be extended to 7–8Ma(Figure M37b;

An et al., 2001). Since the uplift of the Tibetan Plateau can lead

to enhanced aridity in the Asian interior and to intensification

of the Asian monsoon system (Kutzbach et al., 1993), a nearly

concurrent beginning of both the Indian Ocean upwelling and dust

accumulation in central China about 8 Ma ago has been interpreted

as marking the onset of the Indian and East Asian monsoons,

which in turn implies a significant increase in the altitude of the

Plateau (An et al., 2001). However, the recent discovery by Guo

et al. (2002) of a Miocene loess sequence from Qinan, western

Loess Plateau, has further extended the Chinese dust history. Like

the Pleistocene loess, the Miocene loess evinces enhanced aridity

in the dust source areas and energetic winter monsoon winds

required for dust transport, whereas paleosols point to increased

moisture supply by summer monsoon winds. A total of 231 inter-

bedded loess-paleosol layers, representing a nearly continuous his-

tory of eolian dust accumulation from 22 to 6.2 Ma, indicates that

large source areas of eolian dust and energetic winter monsoon

winds existed since early Miocene, at least 14 Ma earlier than pre-

viously thought (Table M4).

Noticeable climate change from arid to humid climatic condi-

tions in East China occurred as early as around the Oligocene/

Miocene boundary. The synthesized data from oil exploration

and stratigraphic studies indicate that a broad aridity belt stretched

across China from west to east during the Paleogene, particularly

in the Paleocene, before it contracted to Northwest China in the

Neogene (Figure M37), suggesting a transition from a planetary

to monsoonal system in atmospheric circulation over the region.

This climate transition, now confirmed by abundant paleobotani-

cal/palynological and lithostratigraphic data, may further imply

that monsoonal moisture brought westward from the ocean to East

Figure M38 Distribution of arid climate zones in China based on paleobotanical and palynological data: (a) Oligocene; (b) Miocene. Open dots

denote humid vegetation, filled dots denote arid vegetation. The Paleocene and Eocene patterns are similar to those of the Oligocene, while

Pliocene and Pleistocene patterns are close to those of the Miocene. The drastic change between the Oligocene and Miocene implies the

beginning of the modern East Asian monsoon system (Sun and Wang, 2005).

MONSOONS: PRE-QUATERNARY 587

China as a response to the reorganization of the climate system

about 24 Ma ago was probably caused by an enhancement, if

not the first establishment, of the East Asian summer monsoon

(Sun and Wang, 2005). The loess-paleosol sequence at Qinan sup-

ports the Oligocene/Miocene climate transition. The existence of

the Asian monsoon at about 16–14 Ma has been reported from

northern Thailand, where middle Miocene mammalian faunas

were found to have adopted to a monsoon-styled wet climate.

Sedimentological data from ODP Leg 116, Bay of Bengal

also imply the intensification of uplift-induced monsoon in

the early Miocene. Although the dramatic d

C increase of total

organic carbon in Bengal Fan sediments at ca. 7 Ma supports

the development of the monsoon in the Himalayan foreland at

this time, the sediment accumulation records are in conflict

with the 8 Ma uplift model. Accumulation rates at several

DSDP/ODP Sites were high for the 17–7 Ma old Bengal

Fan, but decreased from 7 to 1 Ma with the clay mineral assem-

blages indicating reduced physical erosion and strengthened

chemical weathering (Derry and France-Lanord, 1997). No sig-

nificant change in sediment accumulation around 8 Ma has

been found at any ODP Leg 184 sites in the South China Sea

(Wang et al., 2000). On the other hand, if using the planktonic

foraminifer Neogloboquadrina dutertrei as an indicator of the

East Asian monsoon and enhanced productivity in the South

China Sea, its abundance peaks at 7.6 Ma and 3.2 Ma at Site

1146 on the northern slope correspond well to the Indian mon-

soon records (Figure M37c; Wang et al., 2003).

The dust record also indicates that the Asian monsoon

system has a longer history and greater variability both in space

and in time than previously thought. In the Miocene sequence

of Qinan, two intervals are distinguished by higher dust accu-

mulation: 15–13 Ma and 8–7 Ma (Guo et al., 2002). These

might represent periods of enhanced aridity in the source areas,

an interpretation supported by pollen data from Yumen, north-

east of Tibet. Increased aridity over Asia around 8–7 Ma also

explains a peak in dust accumulation rate in the North Pacific

(Figure M37d; Rea et al., 1998).

Monsoon evolution in geological history

The number and geographic coverage of the monsoon records

decrease with increasing age, resulting in relatively deficient

knowledge of the pre-Quaternary monsoon history. Although

it may be premature to discuss the evolution of the monsoon

system through the entire geological history, the examples of

the Pangaean and late Cenozoic times provide convincing

evidence that tectonically-induced changes in sea-land distribu-

tion and in topography have played the primary role in control-

ling the monsoon system over the 10

–10

year and longer

timescales. Tracing back the entire Phanerozoic history, the

monsoon system strengthened with increased size of continents

and increased altitude of mountains. With the collapse of the

megacontinent, for example, the “megamonsoon” circulation

over Pangaea was subsequently replaced by a basically zonal

circulation (Parrish et al., 1982). The late paleozoic American

Appalachians (estimated average altitude of 4,500 m) and the

European Variscan (2,000–3,000 m) in Pangaea might have

played a role similar to the late Cenozoic Himalaya-Tibet in

the intensification of the concomitant monsoon circulation

(Fluteau et al., 2001).

The development of polar ice-sheets is another major con-

trol of the monsoon system. Clemens et al. (1996) found a non-

stationarity in the phase of the summer monsoon system

relative to a growing Northern Hemisphere ice volume over

the past 3.5 Myr. During the initiation and the growth of the

Northern Hemisphere ice sheets, the phase of strong monsoons

moved away from the phase of maximum ice volume and sys-

tematically drifted in a similar pattern over the past 2.6 Myr.

An intensified winter monsoon resulting from the growth of

the boreal ice sheet was also reported from the Chinese loess

records (Liu and Ding, 1993). Accordingly, the nature of the

monsoon system in the ice-house vs. hot-house Earth must

have been different, and the transition between hot- and ice-

house regimes can be of paramount importance for reconstruct-

ing the monsoon history.

A trustworthy reconstruction of the monsoon history depends

on the proxies adopted. Care is necessary to evaluate how the mon-

soon proxies that were developed for the Quaternary period can be

applied to the earlier geological past. It is equally crucial to distin-

guish monsoon proxies from those driven by other factors. An

example is the late Miocene development of C4 vegetation, exhib-

ited by a d

C shift in pedogenic carbonates, which was considered

by Quade et al., (1989) as a signal of the onset of the Indian mon-

soon when first found in Pakistan. Since many subsequent find-

ings have occurred in various continents, this d

C shift has been

re-interpreted as representing a large-scale vegetation change

caused by decreased CO

concentration on the globe (Cerling

et al., 1997).

Regardless of all the complexity, paleo-monsoons are

increasingly important in pre-Quaternary studies. Climate var-

iations throughout geological history have been driven not only

by higher latitudes where ice-sheets develop but also by lower

latitude factors, especially by the monsoons.

Pinxian Wang and Qianyu Li

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Table M4 Development of the dust history in the Loess Plateau, China

Deposits Age Reference

Loess-paleosol sequence 0–2.6 Ma Liu and Ding, 1993

Red clay 2.6–8 Ma An et al., 2001

Qinan loess sequence 6.2–22 Ma Guo et al., 2002

588 MONSOONS: PRE-QUATERNARY