Elsevier Encyclopedia of Geology

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This reaction was the major control on the amount

of carbon dioxide in the atmosphere before the onset

of photosynthesis, respiration, and organic-matter se-

questration. During subduction, carbonate minerals

carried in the descending oceanic crust are heated,

releasing carbon dioxide and water as volatile com-

ponents in arc magmas, completing the cycle back to

the atmosphere. In hydrothermal vents within basaltic

crust, water oxidizes Fe

2þ

in olivine, yielding hydro-

gen, magnetite, and serpentine. In gas-phase high-

temperature reactions, hydrogen can reduce carbon

dioxide to methane, but it is not likely that much

hydrogen was maintained in the atmosphere, owing

to its high rate of escape from the top of the atmos-

phere. The currently held view is that the early sec-

ondary atmosphere was, at most, weakly reducing,

with the major components – carbon dioxide, nitro-

gen, water vapour, and carbon monoxide – supple-

mented by minor mixing fractions of hydrogen and

methane. Estimates have been made of the rate of

decline of carbon dioxide over geological time that

provided the major greenhouse forcing to the atmos-

phere, while tracking secular changes in solar lumi-

nosity. Initially high partial pressures of carbon

dioxide, possibly supplemented by biogenic methane,

would have kept surface temperatures warm enough

to avoid freezing of the planet until the crisis condi-

tions of the snowball Earth in the Late Proterozoic.

Since the era of severe Proterozoic glaciations, there

has been a generally gentle decline in levels of carbon

dioxide, maintaining the stability of the biosphere in

the face of rising solar luminosity (of the order of

about 6% per billion years).

Atmospheric Evolution

Early Anoxic Atmospheres

Free oxygen was rare or absent on the early Earth, so

that the atmosphere was effectively anoxic. Con-

straints on ancient levels of atmospheric oxygen

have been sought for many years, and most models

of the early Earth consider oxygen levels to have been

very low (less than 10

2

of present atmospheric level

(PAL)) before about 2.3 Ga. Evidence to support this

comes from detrital gold, uraninite, and pyrite grains

in sediments older than 2.3 Ga, which it has been

argued could not have been preserved in a high-

oxygen atmosphere. Other geochemical indications,

such as Fe

2þ

/Fe

3þ

in palaeosols and discussions of the

possible origins of banded iron formations (see Sedi-

mentary Rocks: Banded Iron Formations) and of

carbon isotopic signatures in Archaean and Palaeo-

proterozoic organic matter and carbonates, have also

been used to support the contention that the early

atmosphere was essentially devoid of oxygen. In the

absence of a ready source of free oxygen and, there-

fore, no effective ozone (O

) screen from solar ultra-

violet, atmospheric water vapour would have

photodissociated into hydrogen and oxygen. Since

the oxygen formed by this reaction produces just

enough ozone to inhibit the reaction, accumulation

of oxygen by this means is limited. Oxygen combines

readily with hydrogen to form water vapour, so little

oxygen is lost from the atmosphere to space. The

bottom of the stratosphere acts as a cold trap and

maintains a constant temperature of 56



C, keeping

water from escaping from the top of the atmosphere

(Figure 5).

Sulphur isotopes: a unique window onto early atmos-

pheric chemistry Further evidence in support of an

early anoxic atmosphere comes from the unusual

chemistry of sulphur isotopes. Most physical processes

fractionate the isotopes of an element because of

the relative mass differences between the isotopic

species. For instance, because the mass difference bet-

ween

Sand

S is half that between

Sand

S, any

mass-dependent (defined by D

S ¼ 0%;where

S ¼ 0.515d

S – d

S) physical effects on the ratio

Sto

S are expected to be about half those on

the ratio of

Sto

S. However, certain unusual gas-

phase chemical reactions in the absence of free oxygen

are capable of producing anomalous sulphur isotopic

values that do not follow the mass-dependent relation-

ship described above. The recent discovery of the signal

of the mass-independent fractionation effect (where

S 6¼ 0%) in geological materials has important im-

plications for understanding the chemical evolution of

Figure 5 Temperature and pressure profile of the Earth’s

atmosphere.

ATMOSPHERE EVOLUTION 201

the atmosphere. It has been found that sedimentary

sulphide and sulphate minerals in Archaean rocks con-

tain sulphur where the levels of the minor isotope (

deviate from the mass-fractionation line with other

post-2.3 Ga terrestrial sulphur minerals. It appears

that the signal of this chemistry can be transferred

from the atmosphere to surface materials and there-

after to the geological record. The mass-independent

fractionation effect of sulphur in ancient sulphides

records chemical reactions in early anoxic terrestrial

atmospheres, which were destroyed in the Early Pro-

terozoic (2.4–2.3 Ga), and tracks the rise in levels of

atmospheric oxygen (Figure 6).

Metabolic Energy and the Rise of Oxygen

Changes in atmospheric composition have had a pro-

found effect on the evolution of life. Sequence analyses

used to compare genomes have been used to infer the

phylogenetic relatedness of living organisms and their

evolutionary history in terms of molecular biology and

metabolic style. It has been found that three great

domains exist within the ribosomal RNA tree of life:

Bacteria, Archaea, and Eukarya (Figure 7)(see Precam-

brian: Prokaryote Fossils). The organisms with the

deepest branches in the tree, corresponding to the

most ancient pedigrees, close to the root in Figure 7,

are hyperthermophilic (heat-loving) chemoautotrophs

(organisms that obtain their metabolic energy from

chemical disequilibria). It appears that at least some

of the earliest organisms lived off chemical energy, in

much the same way as contemporary hot-spring micro-

bial communities. Such environments are widespread

in volcanic centres on land and on the seafloor, and they

wouldhavebeenevenmorewidespreadonageologic-

ally restive early Earth. However, hyperthermophilic

organisms are restricted to zones around hydrothermal

vents where optimal chemical and temperature condi-

tions persist. Photosynthesizers, on the other hand, are

less restricted and have evolved to inhabit environ-

ments away from those of their chemosynthetic ances-

tors into the open ocean.

Light from the Sun provides a readily accessible,

stable, and inexhaustible energy source to drive meta-

bolic reactions. Photosynthesis uses light energy to cap-

ture electrons and move low-energy-state ions to higher

energy states. The energy released can then be used

to power metabolic reactions such as growth and

reproduction. Photosystem I was the earliest photosyn-

thetic pathway and used light-activated chlorophylls

and electron donors such as H

SandFe

2þ

to yield S

and Fe

3þ

as oxidative products. At a later stage, photo-

system II developed; this uses water as the electron

Figure 6 Mass-independent sulphur isotopic compositions track changes in atmospheric composition with time. (Reproduced from

Mojzsis

et al. (2003).)

202 ATMOSPHERE EVOLUTION

donor to reduce carbon dioxide to biologically useful

organic compounds such as carbohydrates. Liquid

water and carbon dioxide are available to life in prac-

tically unlimited quantities; sunlight drives the reac-

tion, and cyanobacteria that use it, although relative

late-comers in phylogenetic terms, are arguably the

most successful organisms in the history of life on

Earth. It is photosystem II that releases free oxygen as

a by-product in the reversible reaction:

6CO

þ 6H

O þ light energy $ C

þ 6O

The gradual build-up of oxygen in the atmosphere

was accomplished by the slow toilsome efforts of cy-

anobacteria over many hundreds of millions of years.

The immense importance of this metabolic reaction

to the history of all life beyond the microbial level

cannot be overstated (see Earth System Science).

Rise in atmospheric oxygen Initially, the vast

amounts of reduced iron (Fe

2þ

) and other chemical

species dissolved in the oceans, and perhaps

some reduced gases in the atmosphere (such as me-

thane), provided a sink for the oxygen produced by

photosynthesis. As these sinks were exhausted, local

concentrations of oxygen would have appeared, cre-

ating a dilemma for the early anaerobic microbial

biosphere. High levels of oxygen produce toxic rad-

icals in the environment, which cause extensive

damage to cell components. This would have been

a strong driving force for natural selection leading

to adaptation: microbes that did not immediately

Figure 7 The phylogenetic tree of life based on comparative 16s rRNA sequences. (Reproduced from Pace NR (2001) The universal

nature of biochemistry. Proceedings of the National Academy of Sciences USA 98: 805–808.)

ATMOSPHERE EVOLUTION 203

go extinct survived by retreating, for example, to

anaerobic environments, such as deep in crustal

rocks, organic-rich muds, and animal digestive tracts,

where they continue to thrive today. New forms of

life emerged with superoxide dysmutase and catalase

enzymes, which catalyze the reduction of oxic free

radicals to water. Finally, another group (the

Eukarya; Figure 7) developed aerobic metabolism,

providing energy yields approximately twenty times

greater than those of anaerobic metabolism (see Pre-

cambrian: Eukaryote Fossils). The mitochondria –

cell organelles contained in most Eukarya, which fa-

cilitate oxidative metabolism – arose from the endo-

symbiosis of proteobacteria. These organelles

actually contain their own subunits of DNA that

further implicate a eubacterial heritage.

The increased energy yield of aerobic metabolism

set the stage for the evolution of all life above the

unicellular level. The energy source that supports

the aerobic biosphere is the Sun – thus established

oxygenic photosynthesis and aerobic respiration

governed the flow of carbon and oxygen through

the atmosphere–hydrosphere system. During the Ar-

chaean the atmosphere and hydrosphere initially con-

tained small amounts of free oxygen from photolysis

of water vapour and then increasing amounts from

oxygenic photosynthesis. By the end of the Archaean–

Early Proterozoic, oxygen levels had begun to creep

upwards, balanced by oxygen-consuming reactions

such as weathering, hydrothermal activity, respir-

ation, oxidation of organic matter, and differential

rates of organic-matter burial. By about 2 Ga, levels

of oxygen were about 1% of PAL. A transition era

ensued, with oxygen levels fluctuating around the 1%

PAL value and the oceans remaining reducing and

sulphidic. After about 1.8 Ga, oxygen-consuming re-

actions were generally exhausted and free-oxygen

concentrations reached about 10% of PAL. Over

time, these levels stabilized at near ‘normal’ Phaner-

ozoic atmospheric concentrations. It remains unquan-

tified how rapidly oxygen levels increased during the

Proterozoic.

The rise in atmospheric oxygen had an acute effect

on the surface environment, not only because of its

toxicity to many microbial organisms (providing the

motive force to drive the evolution of more efficient

aerobic metabolisms) but also by establishing an ef-

fective ultraviolet screen. Ozone is far more effective

than diatomic oxygen at absorbing ultraviolet. The

ozone screen formed in the stratosphere from accu-

mulating O

that photodissociated to produce free

oxygen radicals (O



), which then recombined with

to make ozone. This ozone screen effectively

made dry land habitable for plants and animals

by the Palaeozoic. Industrial pollutants such as

chlorofluorocarbons are now severely damaging the

ozone layer.

A Neoproterozoic Snowball Earth

The increased oxidation of the surface zone in the

Middle Proterozoic was probably a consequence of

the sequestration of large quantities of organic carbon

in sediments. From the oxygenic photosynthesis reac-

tion, it can be seen how this scenario leads to a net loss

of carbon dioxide and a net increase of oxygen in the

atmospheric reservoir. The boosted greenhouse effect

of heightened levels of carbon dioxide before about

2.5 Ga was lost to the Earthat a most inopportune time.

Levels of solar luminosity were about 10%–15%

lower at 2.2 Ga than at present (Figure 1). Loss of

insolation from the carbon dioxide greenhouse

spelled disaster for the Proterozoic Earth. The planet

froze over, locking the oceans in ice and creating a

high-albedo feedback loop that kept the planet frozen

for extended periods, until levels of carbon dioxide

increased due to passive volcanism and outgassing.

Heightened levels of carbon dioxide warmed the at-

mosphere and dark dusty ice reduced the albedo,

causing catastrophic melting and massive planetary

warming. Subsequent weathering in a carbon diox-

ide-rich atmosphere, combined with massive algal

blooms, led to enhanced burial of organic carbon,

drawdown of carbon dioxide levels, and a renewed

snowball Earth. The cycle is thought to have been

broken by the secular increase in solar luminosity

and a steady redistribution of the continents via

plate motions. The survivors of these repeated ‘ice-

house’ and ‘hothouse’ Earths were multicellular or-

ganisms that inherited a more stable environment

high in oxygen (see Palaeoclimates).

The Phanerozoic Atmosphere

Phanerozoic Atmospheric Changes

Geochemical evidence from the study of palaeosols,

coupled with data from carbonate and organic

carbon in sediments as well as sedimentary pyrite

and the chemistry of sedimentary silicate minerals,

has been used to improve models of the carbon cycle

of palaeoenvironments. These models have been used

to document and explain fluctuations in levels of

oxygen and carbon dioxide over Phanerozoic time.

Long-term changes in the carbon dioxide and oxygen

concentrations in the Phanerozoic atmosphere are

summarized in Figure 8.

In the first part of the Phanerozoic, the Early

Palaeozoic (Cambrian–Ordovician), evidence indi-

cates that levels of carbon dioxide were about fifteen

times PAL. These declined to within a few percent of

204 ATMOSPHERE EVOLUTION

PAL at the end of the Carboniferous. Massive carbon

burial and increased solar radiation by the Carbon-

iferous meant that atmospheric oxygen concentra-

tions were greater than at any other time in Earth

history (see Palaeozoic: Carboniferous). This may

have allowed the existence of the huge Carboniferous

insects observed in the fossil record (see Fossil Inverte-

brates: Insects). Insects rely on diffusion-limited flow

Figure 8 Phanerozoic temperature history. (Reproduced from with permission from Berner RA. Geocarb II: A revised model of

atmospheric CO

over phanerozoic time, Am. Jour. Sci., 294: 56–91)

ATMOSPHERE EVOLUTION 205

of oxygen through their exoskeletons: to be big they

need to live in conditions of high atmospheric oxygen,

so that sufficient oxygen passively diffuses into their

blood to power muscles for flight.

Reduced greenhouse forcing and glaciation at the

end of the Palaeozoic with the assembly of Pangaea

reduced organic-matter burial, and there was a slow

rise in carbon dioxide levels to around six times PAL

in the Permian and Triassic. Levels of carbon dioxide

gradually decreased, and stabilized near present-day

levels in the Mesozoic. Oxygen tends to track carbon

dioxide inversely; geological evidence and palaeocli-

mate models suggest a maximum of near 35% oxygen

in the atmosphere at the beginning of the Permian.

During the Permian, the oceans were highly stratified,

with carbonate-rich water at depth that was depleted

in oxygen. This system was unstable: ocean hypoxia

could occur if ocean circulation intensified enough

to mix deep-water carbon dioxide and hydrogen

sulphide into surface waters. A protracted (20 Ma)

whole-ocean hypoxia event is considered to be a

major mechanism in the Permian–Triassic extinction

event, which wiped out 90%–95% of all marine

species (see Palaeozoic: End Permian Extinctions).

Carbon Dioxide and Climate Changes

High-resolution information about changes in atmos-

pheric chemistry over the past 160 000 years can be

obtained by studying the record of trapped gases in ice

cores from the Greenland and Antarctic ice-caps. Fur-

thermore, oxygen isotopic values from marine sedi-

ments, marine planktonic and benthonic fossils, cave

deposits, and other sources can be used to estimate

marine palaeotemperatures. Direct measurements of

carbon dioxide and methane concentrations in ice

cores permit assessment of past atmospheric levels of

these gases, providing factors to incorporate into

models of past air temperatures and sea-levels.

Data from deep ice cores taken in polar re-

gions, coupled with complex palaeoclimate models

(Figure 9), show large fluctuations in atmospheric

carbon dioxide, oxygen and methane levels, leading

to long-term temperature changes of the order of

6



C or more. There is a strong correlation between

levels of atmospheric greenhouse gases and palaeo-

temperature. The periodicities in these data provide

clear evidence of the role of Milankovitch forcing by

changes in the Earth’s orbital parameters. The two

strongest Milankovitch cycles observed correspond

to the 26 000 year precession of the equinoxes and

the 100 000 year period of rotation of the Earth’s

orbital axis (see Earth: Orbital Variation (Including

Milankovitch Cycles)). Changes in greenhouse-gas

concentrations appear to follow rather than guide

long-term climate, suggesting that Milankovitch

cycles are the prime mechanism bringing the Earth

into a greenhouse or icehouse condition. Greenhouse

gases do provide positive feedback at the beginning of

temperature changes by boosting insolation. Anthro-

pogenic emissions of greenhouse gases are not

governed by Milankovitch cycles and represent a sep-

arate and increasingly important climate-forcing

mechanism. Figure 9 shows that carbon dioxide con-

centrations changed by almost 100 ppm(vol.) towards

the end of the last glaciation. Modern levels of carbon

dioxide are near 370 ppm(vol.) and rising. Almost all

of this change has occurred since the Industrial Revo-

lution, and high-resolution monitoring at the Mauna

Loa observatory shows clear diurnal and annual

cycles in carbon dioxide levels (Figure 10), with a

mean annual increase of 1.16 ppm(vol.) year

1

. This

is over one hundred times the rate of increase of

carbon dioxide levels inferred from all available

Figure 9 Phanerozoic carbon dioxide and oxygen concentra-

tions.

Figure 10 Changes in the concentration of atmospheric carbon

dioxide over the last 60 years as measured at Mauna Loa, Hawaii.

Data provided by D. Keeling and T. Whorf.

206 ATMOSPHERE EVOLUTION

geological data, with no clear stabilization of these

levels in sight.

Other greenhouse gases Neglecting the small contri-

bution from internal heating driven by radioactive

decay, planetary surface temperatures are governed

by the amount of solar radiation received and its

interaction with the atmosphere. Re-radiation to

the surface of infrared radiation by greenhouse

gases, chiefly carbon dioxide but also water vapour,

methane, nitrous oxide species (NO

), and chloro-

fluorocarbons (CFCs), keeps the present average sur-

face temperature some 33 K above the black-body

temperature of 255 K. Although they are present in

small amounts relative to carbon dioxide and water

vapour, methane

nitrous oxide species, and CFCs

(which can form only from industrial processes) are

very effective greenhouse gases. Methane in the at-

mosphere (concentration of 1.7 ppm(vol.)) is domin-

antly formed by biological processes and absorbs

infrared radiation approximately 21 times more ef-

fectively than carbon dioxide; its levels have increased

rapidly in the last several centuries. Gaseous nitrous

oxide compounds are present at low concentrations

(ca. 0.3 ppm(vol.)), are produced by biological

nitrification, and have a long residence time in the

atmosphere. The concentration of CFCs is very low

(less than 0.003 ppm(vol.)), but they have a green-

house effect ten thousand times greater than that of

carbon dioxide and likewise have long residence

times in the atmosphere.

Conclusions

The splendour of contemporary life on Earth, as re-

vealed by the geochemical, palaeontological, and mo-

lecular phylogenetic records, took billions of years to

achieve, and its evolution occurred for the most part

within the envelope of the atmosphere and hydro-

sphere. In the context of astrophysical changes to

the Sun and geophysical changes to the solid Earth,

the atmosphere has evolved through a complex set of

interrelated cycles, within which biology has been of

central importance. Life has affected the planetary

atmosphere, and biological evolution was in turn

affected by it. The gaseous envelope of the planet

evolved in response to changing geophysical condi-

tions, such as mantle heat flow and solar luminosity.

Early in Solar System history, Venus and Mars appar-

ently had geochemical paths that paralleled that of

the Earth, probably including liquid water at their

surfaces. However, long ago they diverged from a

physicochemical regime that promoted habitability.

The nascent biosphere on these planets, if it ever

existed, was either burned to a crisp (on Venus) or

freeze-dried (on Mars). In the far future, as the lumi-

nosity of the Sun continues to increase, Earth will go

the way of Venus.

Further Reading

Hoffman PF and Schrag DP (2000) Snowball Earth.

Scientific American 282: 68–75.

Holland HD (1984) The Chemical Evolution of the Atmos-

phere and Oceans. Princeton: Princeton University Press.

Margulis L (1984) Early Life. Boston: Jones and Bartlett.

Mojzsis SJ and Harrison TM (2000) Vestiges of a

beginning: clues to the emergent biosphere recorded in

the oldest known sedimentary rocks. GSAToday 10: 1–6.

Pace NR (2001) The universal nature of biochemistry.

Proceedings of the National Academy of Sciences USA

98: 805–808.

Royer DL, Berner RA, Montan

ez IP, Tabor NJ, and Beerling

DJ (2004) CO

as a primary driver of Phanerozoic

climate. GSA Today 14: 4–10.

Sagan C and Mullen G (1972) Earth and Mars: evolution

of atmospheres and surface temperatures. Science 177:

52–56.

ATMOSPHERE EVOLUTION 207

AUSTRALIA

Contents

Proterozoic

Phanerozoic

Tasman Orogenic Belt

Proterozoic

I M Tyler, Geological Survey of Western Australia,

East Perth, WA, Australia

Introduction

The Proterozoic is the period of geological time

extending from the end of the Archaean, 2500 million

years ago (Ma), to the start of the Phanerozoic (the

base of the Cambrian System), 545 Ma. The Protero-

zoic is divided into the Palaeoproterozoic (2500–

1600 Ma), the Mesoproterozoic (1600–1000 Ma),

and the Neoproterozoic (1000–545 Ma). In Austra-

lia, Proterozoic rocks are present to the west of the

‘Tasman Line’ that separates ‘Proterozoic Australia’,

where geophysical datasets show that Precambrian

basement is continuous beneath Phanerozoic sedi-

mentary basins, from the Tasmanides, where predom-

inantly Palaeozoic basement is overlain by Mesozoic

and younger sedimentary basins. Extensive exposures

of Proterozoic rocks are present in western Australia,

in northern, central, and north-eastern Australia, and

in southern Australia and Tasmania. Proterozoic

Australia is made up of three distinct cratons, the

West Australian Craton, the North Australian

Craton, and the South Australian Craton (Figure 1).

These probably formed originally as parts of larger

cratons or continental blocks (the South Australian

Craton together with the previously adjacent part of

Antarctica formed the Mawson Craton) and are

dominated by Archaean and Palaeoproterozoic to

Mesoproterozoic crust. The three cratons are separ-

ated by two predominantly Mesoproterozoic to

Neoproterozoic orogenic belts, the Paterson Orogen

and the Albany–Fraser Orogen. The Palaeoprotero-

zoic to Neoproterozoic Pinjarra Orogen is present

along the western margin of Australia.

Plate tectonic models can be applied to Proterozoic

Australia. The increasing availability of high-quality

geochronological data has highlighted the presence of

distinct tectonostratigraphic terranes with differing

geological histories within orogenic belts, and geo-

physical datasets reveal the heterogeneous nature of

the crust throughout Proterozoic Australia. However,

Proterozoic plate-tectonic processes may differ from

modern processes, and the real lack of accretionary

complexes and ophiolites, and significant differences

in the geochemical and isotopic compositions of igne-

ous rocks, may reflect changes in the nature and

composition of the oceanic lithosphere through

time. Palaeomagnetic evidence is placing greater con-

trols on the movement and relative positions of the

constituent crustal components. Diverse cratons and

continental blocks aggregated during the Palaeopro-

terozoic and Early Mesoproterozoic to form Protero-

zoic Australia, which then played an integral part

in the formation and breakup of the Meso- to Neo-

proterozoic supercontinent, Rodinia (Figures 2–7;

summarized in Table 1). Proterozoic Australia is

host to a wide variety of minerals , including world-

class deposits of iron, uranium, gold, copper–gold,

lead–zinc–silver, and diamond orebodies (Figure 8).

Neoarchaean to Palaeoproterozoic

Assembling Proterozoic Australia:

(2770–1600 Ma)

West Australian Craton

Within the West Australian Craton, large areas of

Archaean rocks are exposed in the geologically dis-

tinct Pilbara and Yilgarn cratons (Figures 1 and 2).

The Hamersley Basin was initiated on the southern

part of the Pilbara Craton at the start of the Neoarch-

aean. West- to south-westerly directed rifting within

cratonized granite–greenstone basement represents a

distinct change to a Proterozoic and Phanerozoic tec-

tonic style. The flood basalts (2770–2690 Ma) of the

Fortescue Group dominate the rift-related lower part

of the basin. These were buried beneath a breakup

unconformity overlain by a passive margin sequence

characterized by cherts and banded iron formations

208 AUSTRALIA/Proterozoic

(BIFs) of the uppermost Fortescue Group and the

Hamersley Group (2600–2440 Ma), which were

deposited on a shelf or platform open to an ocean

(see Sedimentary Rocks: Banded Iron Formations).

A collisional, intracontinental back-arc setting has

been proposed for the BIFs and mafic and felsic mag-

matic rocks of the upper part (2470–2440 Ma) of the

Hamersley Group. The overlying Turee Creek Group,

which includes probable glacial deposits, and the

lower Wyloo Group were deposited in the McGrath

Trough, a foreland basin developed in front of a

northward-verging fold-and-thrust belt during the

Ophthalmian Orogeny (2200 Ma) (Figure 2). The

Glenburgh Terrane in the southern part of the Gas-

coyne Complex within the Capricorn Orogen to the

south contains basement from 2550–2450 Ma and

may represent a remnant of the colliding continent

that drove this orogeny.

In the Yilgarn Craton, tectonic processes typical of

the Archaean developed the Eastern Goldfields gran-

ite–greenstone terrane 2700–2600 Ma, at the same

time as the upper Fortescue Group on the Pilbara

Craton. In the southern Capricorn Orogen, rifting

and continental breakup took place in the gneiss ter-

ranes and granite–greenstones along the northern

margin of the Yilgarn Craton. The 2150-My-old

rocks of the basal Yerrida Basin (Figure 2) formed

initially in a sag basin followed by an abrupt change

to a rift setting. The development of the Bryah and

Padbury basins 2020–1900 Ma at the north-western

margin of the Yilgarn Craton may reflect the de-

velopment of a back-arc basin, which was then over-

lain by a foreland basin during the Glenburgh

Orogeny 2005–1960 Ma (Figure 2). Geochemical

and isotopic data indicate that suprasubduction

zone magmatism formed the Dalgaringa Supersuite

at an Andean-type margin during this event; this may

represent the coming together of the combined

Pilbara Craton and Glenburgh Terrane with the

Yilgarn Craton.

The Capricorn Orogeny from 1830–1780 Ma

(Figure 3) has previously been regarded as marking

the collision between the Pilbara and Yilgarn

Cratons. Neodymium isotopes show that melting of

pre-existing, Early Palaeoproterozoic crust without

the introduction of mantle-derived material formed

Figure 1 North, South, and West Australian cratons, showing the main outcrops of Precambrian rocks in Australia.

AUSTRALIA/Proterozoic 209

voluminous granitic magmatism, which, together

with associated deformational and metamorphic

events, extended across the entire orogen. The Capri-

corn Orogeny may represent an intracratonic event

resulting from reactivation, possibly during amal-

gamation of the West Australian Craton with the

North Australian Craton. The 1840- to 1805-

My-old upper part of the Ashburton Basin and the

1805-My-old Blair Basin developed as a foreland

basin along the northern margin of the orogen at this

time. Uplift of the Gascoyne Complex and the Yilgarn

Craton, together with a presently unexposed Early

Palaeoproterozoic terrane, provided the sediment de-

posited in the upper Ashburton Basin and the Blair

Basin, the 1840-My-old upper part of the Yerrida

Basin and the 1840- to 1800-My-old Earaheedy Basin

(Figure 3).

In the Rudall Complex along the eastern margin

of the Pilbara Craton (Figure 3), the Talbot Terrane

contains quartzite, amphibolite, and serpentinite as

inclusions within complex orthogneisses that contain

components that crystallized, respectively, 2015,

1970, and 1800 Ma. A younger sequence (from

<1790 Ma) of clastic metasedimentary rocks was

possibly deposited in a foreland basin. In the adjacent

Connaughton Terrane, a sequence of deformed and

metamorphosed mafic volcanic rocks, and chemical

and clastic sedimentary rocks, may have been de-

posited in a rift prior to 1780 Ma. Deformation,

metamorphism, and further granite intrusion took

place in the Talbot and Connaughton terranes

during the Yapungku Orogeny (1790–1760 Ma).

West-verging thrusting and high-P metamorphism

may have accompanied collision of the West Austra-

lian Craton with the North Australian Craton, which

palaeomagnetic evidence indicates were unified by

1700 Ma.

The Mount Barren Group and equivalent sediment-

ary rocks were deposited on the south-eastern part of

the Yilgarn Craton 1700 Ma (Figure 4). Reactiva-

tion of the Capricorn Orogen took place between

1670 and 1620 Ma with the intrusion of granitic

rocks and the occurrence of medium- to high-grade

metamorphism, which was synchronous with local-

ized shear zones and hydrothermal alteration. In the

Albany–Fraser Orogen, the Biranup Complex con-

sists of granitic gneisses that include both Archaean

(2630–2575 Ma) (Figure 2) and Palaeoproterozoic

(1700–1600 Ma) protoliths (Figure 3). In the Pin-

jarra Orogen, a monzogranite in fault contact with

Figure 2 Assembly of the West Australian Craton, 2800–1900 Ma.

210 AUSTRALIA/Proterozoic

Elsevier Encyclopedia of Geology - vol I A-E

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