Timeline of World History: Year by Year from Prehistory to Present Day


    Year by Year: Big Bang
  BACK- NEXT-Proterozoic Eon    

Geologic  Timeline
of the Earth

Big Bang
Precambrian Time
(4567 to 542 mya)
Hadean Eon
(4567 to 3800 mya)
Archaean Eon
(3800 to 2500 mya)
Proterozoic Eon
(2500 to 542 mya)
Paleoproterozoic Era
(2500 to 1600 mya)
Mesoproterozoic Era
(1600 to 1000 mya)
Neoproterozoic Era
(1000 to 542 mya)
Paleozoic Era
(542 to 251 mya)
Mesozoic Era
(251 to 65.5 mya)
Cenozoic Era
(65.5 mya to today)
Quaternary Period
(2.58 mya to today)
3 million B.C.
30.000 B.C.
9000 B.C.

Bible Illustrationsby Gustave Dore
"Now the earth was formless
and empty, darkness was
over the surface of the deep..."

Genesis 1:2
Geologic and Biological Timeline of the Earth

Before Big Bang 

In a vacuum state with no space or time, physical laws would not seem appropriate. However, the law that states matter can neither be created nor destroyed implies another state here, i.e. a state of pure energy unbound by space and time. The chance fluctuation indicated below for the beginning of the Big Bang would have occurred in this energy field. This occurrence could have been like the breaking of a dam, or a puncture that explodes a filled tire, or a bomb that violently explodes upon detonation. The resulting tiny bubble of space-time provided an outlet for the enormous energy latent in the pre-space-time state. This of course gives no account of how or where or why the initial pure-energy state came about. We may never know, but we can always speculate.



Time, space, matter and energy came into being. Matter and energy began to define space and time. In 2002 scientists said experiments confirmed that only 5% of the universe was composed of ordinary matter. 65% was said to be "dark energy" and 30% was "dark matter." In 1998 Joseph authored “The Big Bang.” A 3rd ed. was published in 2004. In 2004 Simon Singh authored “Big Bang: The Most Important Scientific Discovery of All Time and Why You Need to Know About It.” In 2006 NASA released data backing the Big Bang theory that the universe sprang from marble size to infinity in less than a trillion-trillionth second.


Big-bang model, widely held theory of the evolution of the universe. Its essential feature is the emergence of the universe from a state of extremely high temperature and density—the so-called big bang that occurred 13.8 billion years ago. Although this type of universe was proposed by Russian mathematician Aleksandr Friedmann and Belgian astronomer Georges Lemaître in the 1920s, the modern version was developed by Russian-born American physicist George Gamow and colleagues in the 1940s.

The big-bang model is based on two assumptions. The first is that Albert Einstein’s general theory of relativity correctly describes the gravitational interaction of all matter. The second assumption, called the cosmological principle, states that an observer’s view of the universe depends neither on the direction in which he looks nor on his location. This principle applies only to the large-scale properties of the universe, but it does imply that the universe has no edge, so that the big-bang origin occurred not at a particular point in space but rather throughout space at the same time. These two assumptions make it possible to calculate the history of the cosmos after a certain epoch called the Planck time. Scientists have yet to determine what prevailed before Planck time.
  According to the big-bang model, the universe expanded rapidly from a highly compressed primordial state, which resulted in a significant decrease in density and temperature. Soon afterward, the dominance of matter over antimatter (as observed today) may have been established by processes that also predict proton decay. During this stage many types of elementary particles may have been present. After a few seconds, the universe cooled enough to allow the formation of certain nuclei.
The theory predicts that definite amounts of hydrogen, helium, and lithium were produced.

Their abundances agree with what is observed today. About one million years later the universe was sufficiently cool for atoms to form. The radiation that also filled the universe was then free to travel through space. This remnant of the early universe is the cosmic microwave background radiation—the “three degree” (actually 2.728 K) background radiation—discovered in 1965 by American physicists Arno A. Penzias and Robert W. Wilson.

In addition to accounting for the presence of ordinary matter and radiation, the model predicts that the present universe should also be filled with neutrinos, fundamental particles with no mass or electric charge. The possibility exists that other relics from the early universe may eventually be discovered.


Timeline of the Big Bang

The Hubble Ultra Deep Field showcases galaxies from an ancient era when the Universe was younger, denser, and warmer according to the Big Bang theory.

Extrapolation of the expansion of the Universe backwards in time using general relativity yields an infinite density and temperature at a finite time in the past. This singularity signals the breakdown of general relativity. How closely we can extrapolate towards the singularity is debated—certainly no closer than the end of the Planck epoch. This singularity is sometimes called "the Big Bang", but the term can also refer to the early hot, dense phase itself, which can be considered the "birth" of our Universe. Based on measurements of the expansion using Type Ia supernovae, measurements of temperature fluctuations in the cosmic microwave background, and measurements of the correlation function of galaxies, the Universe has a calculated age of 13.75 ± 0.11 billion years. The agreement of these three independent measurements strongly supports the ΛCDM model that describes in detail the contents of the Universe.


The earliest phases of the Big Bang are subject to much speculation. In the most common models, the Universe was filled homogeneously and isotropically with an incredibly high energy density and huge temperatures and pressures and was very rapidly expanding and cooling. Approximately 10−37 seconds into the expansion, a phase transition caused a cosmic inflation, during which the Universe grew exponentially. After inflation stopped, the Universe consisted of a quark–gluon plasma, as well as all other elementary particles. Temperatures were so high that the random motions of particles were at relativistic speeds, and particle–antiparticle pairs of all kinds were being continuously created and destroyed in collisions. At some point an unknown reaction called baryogenesis violated the conservation of baryon number, leading to a very small excess of quarks and leptons over antiquarks and antileptons—of the order of one part in 30 million. This resulted in the predominance of matter over antimatter in the present Universe.

The Universe continued to grow in size and fall in temperature, hence the typical energy of each particle was decreasing. Symmetry breaking phase transitions put the fundamental forces of physics and the parameters of elementary particles into their present form. After about 10−11 seconds, the picture becomes less speculative, since particle energies drop to values that can be attained in particle physics experiments. At about 10−6 seconds, quarks and gluons combined to form baryons such as protons and neutrons. The small excess of quarks over antiquarks led to a small excess of baryons over antibaryons. The temperature was now no longer high enough to create new proton–antiproton pairs (similarly for neutrons–antineutrons), so a mass annihilation immediately followed, leaving just one in 1010 of the original protons and neutrons, and none of their antiparticles. A similar process happened at about 1 second for electrons and positrons. After these annihilations, the remaining protons, neutrons and electrons were no longer moving relativistically and the energy density of the Universe was dominated by photons (with a minor contribution from neutrinos). A few minutes into the expansion, when the temperature was about a billion (one thousand million; 109; SI prefix giga-) kelvin and the density was about that of air, neutrons combined with protons to form the Universe's deuterium and helium nuclei in a process called Big Bang nucleosynthesis. Most protons remained uncombined as hydrogen nuclei. As the Universe cooled, the rest mass energy density of matter came to gravitationally dominate that of the photon radiation.

After about 379,000 years the electrons and nuclei combined into atoms (mostly hydrogen); hence the
  radiation decoupled from matter and continued through space largely unimpeded. This relic radiation is known as the cosmic microwave background radiation.

Over a long period of time, the slightly denser regions of the nearly uniformly distributed matter gravitationally attracted nearby matter and thus grew even denser, forming gas clouds, stars, galaxies, and the other astronomical structures observable today. The details of this process depend on the amount and type of matter in the Universe. The four possible types of matter are known as cold dark matter, warm dark matter, hot dark matter and baryonic matter. The best measurements available (from WMAP) show that the data is well-fit by a Lambda-CDM model in which dark matter is assumed to be cold (warm dark matter is ruled out by early reionization), and is estimated to make up about 23% of the matter/energy of the universe, while baryonic matter makes up about 4.6%.

In an "extended model" which includes hot dark matter in the form of neutrinos, then if the "physical baryon density" Ωbh2 is estimated at about 0.023 (this is different from the 'baryon density' Ωb expressed as a fraction of the total matter/energy density, which as noted above is about 0.046), and the corresponding cold dark matter density Ωch2 is about 0.11, the corresponding neutrino density Ωvh2 is estimated to be less than 0.0062.

Independent lines of evidence from Type Ia supernovae and the CMB imply that the Universe today is dominated by a mysterious form of energy known as dark energy, which apparently permeates all of space. The observations suggest 73% of the total energy density of today's Universe is in this form. When the Universe was very young, it was likely infused with dark energy, but with less space and everything closer together, gravity had the upper hand, and it was slowly braking the expansion. But eventually, after numerous billion years of expansion, the growing abundance of dark energy caused the expansion of the Universe to slowly begin to accelerate. Dark energy in its simplest formulation takes the form of the cosmological constant term in Einstein's field equations of general relativity, but its composition and mechanism are unknown and, more generally, the details of its equation of state and relationship with the Standard Model of particle physics continue to be investigated both observationally and theoretically.

All of this cosmic evolution after the inflationary epoch can be rigorously described and modeled by the ΛCDM model of cosmology, which uses the independent frameworks of quantum mechanics and Einstein's General Relativity. As noted above, there is no well-supported model describing the action prior to 10−15 seconds or so. Apparently a new unified theory of quantum gravitation is needed to break this barrier. Understanding this earliest of eras in the history of the Universe is currently one of the greatest unsolved problems in physics.



Earth as seen from Apollo 17

The age of the Earth is 4.54 billion years (4.54 × 109 years ± 1%). This age is based on evidence from radiometric age dating of meteorite material and is consistent with the ages of the oldest-known terrestrial and lunar samples. Following the scientific revolution and the development of radiometric age dating, measurements of lead in uranium-rich minerals showed that some were in excess of a billion years old.

The oldest such minerals analyzed to date – small crystals of zircon from the Jack Hills of Western Australia – are at least 4.404 billion years old. Comparing the mass and luminosity of the Sun to the multitudes of other stars, it appears that the solar system cannot be much older than those rocks. Ca-Al-rich inclusions (inclusions rich in calcium and aluminium) – the oldest known solid constituents within meteorites that are formed within the solar system – are 4.567 billion years old, giving an age for the solar system and an upper limit for the age of Earth.

It is hypothesised that the accretion of Earth began soon after the formation of the Ca-Al-rich inclusions and the meteorites. Because the exact accretion time of Earth is not yet known, and the predictions from different accretion models range from a few millions up to about 100 million years, the exact age of Earth is difficult to determine. It is also difficult to determine the exact age of the oldest rocks on Earth, exposed at the surface, as they are aggregates of minerals of possibly different ages.

Geologic time
Geologic time, the extensive interval of time occupied by the Earth’s geologic history. It extends from about 3.9 billion years ago (corresponding to the age of the oldest known rocks) to the present day. It is, in effect, that segment of Earth history that is represented by and recorded in rock strata.
The geologic time scale is the “calendar” for events in Earth history. It subdivides all time since the end of the Earth’s formative period as a planet (nearly 4 billion years ago) into named units of abstract time: the latter, in descending order of duration, are eons, eras, periods, and epochs. The enumeration of these geologic time units is based on stratigraphy, which is
  the correlation and classification of rock strata.

The fossil forms that occur in these rocks provide the chief means of establishing a geologic time scale. Because living things have undergone evolutionary changes over geologic time, particular kinds of organisms are characteristic of particular parts of the geologic record. By correlating the strata in which certain types of fossils are found, the geologic history of various regions (and of the Earth as a whole) can be reconstructed. The relative geologic time scale developed from the fossil record has been numerically quantified by means of absolute dates obtained with radiometric dating methods.

The geologic time scale provides a system of chronologic measurement relating stratigraphy to time that is used by geologists, paleontologists and other earth scientists to describe the timing and relationships between events that have occurred during the history of the Earth. The table of geologic time spans presented here agrees with the dates and nomenclature proposed by the International Commission on Stratigraphy, and uses the standard color codes of the United States Geological Survey.

Evidence from radiometric dating indicates that the Earth is about 4.570 billion years old. The geological or deep time of Earth's past has been organized into various units according to events which took place in each period. Different spans of time on the time scale are usually delimited by major geological or paleontological events, such as mass extinctions. For example, the boundary between the Cretaceous period and the Paleogene period is defined by the Cretaceous–Tertiary extinction event, which marked the demise of the dinosaurs and of many marine species. Older periods which predate the reliable fossil record are defined by absolute age.

Each era on the scale is separated from the next by a major event or change.

(mya = million years ago)

Precambrian Time (4567 to 542 mya)

Hadean Eon (4567 to 3800 mya)
Archaean Eon (3800 to 2500 mya)
Proterozoic Eon (2500 to 542 mya)
Paleoproterozoic Era
(2500 to 1600 mya)

   Siderian Period (2500 to 2300 mya)
   Rhyacian Period (2300 to 2050 mya)

   Orosirian Period (2050 to 1800 mya)

   Statherian Period (1800 to 1600 mya)
Mesoproterozoic Era
(1600 to 1000 mya)

   Calymmian Period (1600 to 1400 mya)
   Ectasian Period (1400 to 1200 mya)
   Stenian Period (1200 to 1000 mya)
Neoproterozoic Era
(1000 to 542 mya)

   Tonian Period (1000 to 850 mya)
   Cryogenian Period (850 to 630 mya)
   Ediacaran (Vendian) Period (630 to 542 mya)
Paleozoic Era
(542 to 251 mya)

Cambrian Period
(542 to 488.3 mya)

   Tommotian Stage (534 to 530 mya)
Ordovician Period (488.3 to 443.7 mya)
Silurian Period (443.7 to 416 mya)
Devonian Period (416 to 359.2 mya)
Carboniferous Period (359.2 to 299 mya)

   Mississippian Epoch (359.2 to 318.1 mya)
   Pennsylvanian Epoch (318.1 to 299 mya)
Permian Period (299 to 251 mya)
Mesozoic Era
(251 to 65.5 mya)

Triassic Period
(251 to 199.6 mya)
Jurassic Period (199.6 to 145.5 mya)
Cretaceous Period (145.5 to 65.5 mya)
Cenozoic Era
(65.5 mya to today)

Paleogene Period
(65.5 to 23.03 mya)      

Tertiary Period
(65.5 to 2.58 mya)
Paleocene Epoch (65.5 to 55.8 mya)
Eocene Epoch (55.8 to 33.9 mya)
Oligocene Epoch (33.9 to 23.03 mya)
    Neogene Period (23.03 mya to today)

Miocene Epoch (23.03 to 5.3 mya)

     Pliocene Epoch (5.3 to 2.58 mya)
    Quaternary Period (2.58 mya to today)

     Pleistocene Epoch (2.58 mya to 11,400 yrs ago)
      Beginning of the Stone Age
   - Neanderthal man spreads through Europe
   - Homo sapiens reduced to about 10,000 individuals
   Holocene Epoch (11,400 years ago to today)

   - 5,300 yrs ago: The Bronze Age
   - 3,300 yrs ago: The Iron Age
Geologic history of Earth

Geologic history of Earth, evolution of the continents, oceans, atmosphere, and biosphere. The layers of rock at the Earth’s surface contain evidence of the evolutionary processes undergone by these components of the terrestrial environment during the times at which each layer was formed. By studying this rock record from the very beginning, it is thus possible to trace their development and the resultant changes through time.
The pregeologic period

From the point at which the planet first began to form, the history of the Earth spans approximately 4.6 billion years. The oldest known rocks, however, have an isotopic age of only about 3.9 billion years. There is in effect a stretch of 700 million years for which no geologic record exists, and the evolution of this pregeologic period of time is, not surprisingly, the subject of much speculation. To understand this little-known period, the following factors have to be considered: the age of formation at 4.6 billion years ago, the processes in operation until 3.9 billion years ago, the bombardment of the Earth by meteorites, and the earliest zircon crystals.

It is widely accepted by both geologists and astronomers that Earth is roughly 4.6 billion years old. This age has been obtained from the isotopic analysis of many meteorites as well as of soil and rock samples from the Moon by such dating methods as rubidium–strontium and uranium–lead. It is taken to be the time when these bodies formed and, by inference, the time at which a significant part of the solar system developed. When the evolution of the isotopes of lead-207 and lead-206 is studied from several lead deposits of different age on Earth, including oceanic sediments that represent a homogenized sample of the Earth’s lead, the growth curve of terrestrial lead can be calculated, and, when this is extrapolated back in time, it is found to coincide with the age of about 4.6 billion years measured on lead isotopes in meteorites.

The Earth and meteorites thus have had similar lead isotope histories, and so it is concluded that over a period of about 30 million years they condensed or accreted as solid bodies from a primeval cloud of interstellar gas and dust—the so-called solar nebula from which the entire solar system is thought to have formed—at about the same time.

Models developed from the comparison of lead isotopes in meteorites and the decay of hafnium-182 to tungsten-182 in Earth’s mantle, however, suggest that approximately 100 million years elapsedbetween the beginning of the solar system and the conclusion of the accretion process that formed Earth. These models place Earth’s age at approximately 4.5 billion years old.

Particles in the solar nebula condensed to form solid grains, and with increasing electrostatic and gravitational influences they eventually clumped together into fragments or chunks of rock. One of these planetesimals developed into the Earth. The constituent metallic elements sank toward the centre of the mass, while lighter elements rose toward the top. The lightest ones (such as hydrogen and helium) that might have formed the first, or primordial, atmosphere probably escaped into outer space.

In these earliest stages of terrestrial accretion heat was generated by three possible phenomena:

(1) the decay of short-lived radioactive isotopes,
(2) the gravitational energy released from the sinking of metals, or

(3) the impact of small planetary bodies (or planetesimals).

  The increase in temperature became sufficient to heat the entire planet. Melting at depth produced liquids that were gravitationally light and thus rose toward the surface and crystallized to form the earliest crust.
Meanwhile, heavier liquids rich in iron, nickel, and perhaps sulfur separated out and sank under gravity, giving rise to the core at the centre of the growing planet; and the lightest volatile elements were able to rise and escape by outgassing, which may have been associated with surface volcanic activity, to form the secondary atmosphere and the oceans. This chemical process of melting, separation of material, and outgassing is referred to as the differentiation of the Earth.

The earliest thin crust was probably unstable and so foundered and collapsed to depth. This in turn generated more gravitational energy, which enabled a thicker, more stable, longer-lasting crust to form. Once the Earth’s interior (or its mantle) was hot and liquid, it would have been subjected to large-scale convection, which may have enabled oceanic crust to develop above upwelling regions. Rapid recycling of crust–mantle material occurred in convection cells, and in this way the earliest terrestrial continents may have evolved during the 700-million-year gap between the formation of the Earth and the beginning of the rock record. It is known from direct observation that the surface of the Moon is covered with a multitude of meteorite craters.
There are about 40 large basins attributable to meteorite impact. Known as maria, these depressions were filled in with basaltic lavas caused by the impact-induced melting of the lunar mantle. Many of these basalts have been analyzed isotopically and found to have crystallization ages of 3.9 to 4 billion years. It can be safely concluded that the Earth, with a greater attractive mass than the Moon, must have undergone more extensive meteorite bombardment. According to the English-born geologist Joseph V.

Smith, a minimum of 500 to 1,000 impact basins were formed on the Earth within a period of about 100 to 200 million years prior to 3.95 billion years ago. Moreover, plausible calculations suggest that this estimate represents merely the tail end of an interval of declining meteorite bombardment and that about 20 times as many basins were formed in the preceding 300 million years. Such intense bombardment would have covered most of the Earth’s surface, with the impacts causing considerable destruction of the terrestrial crust up to 3.9 billion years ago. There is, however, no direct evidence of this important phase of Earth history because rocks older than 3.9 billion years have not been preserved.
An exciting discovery was made in 1983 by William Compston and his research group at the Australian National University with the aid of an ion microprobe. Compston and his associates found that a water-laid clastic sedimentary quartzite from Mount Narryer in western Australia contained detrital zircon grains that were 4.18 billion years old. In 1986 they further discovered that one zircon in a conglomerate only 60 kilometres away was 4.276 billion years old; 16 other grains were determined to be the same age or slightly younger. This is the oldest dated material on Earth. The rocks from which the zircons in the quartzites and conglomerates were derived have either disappeared or have not yet been found. The ages of these single zircon grains are significantly older than those of the oldest known intact rocks, which are granites discovered near the Great Slave Lake in northwestern Canada. The latter contain zircons that are 3.96 billion years old.

Development of the atmosphere and oceans

Formation of the secondary atmosphere

The Earth’s secondary atmosphere began to develop at the time of planetary differentiation, probably in connection with volcanic activity. Its component gases, however, were most likely very different from those emitted by modern volcanoes. Accordingly, the composition of the early secondary atmosphere was quite distinct from that of today’s atmosphere. Carbon monoxide, carbon dioxide, water vapour, and methane predominated; however, free oxygen could not have been present, since even modern volcanic gases contain no oxygen. It is therefore assumed that the secondary atmosphere during the Archean—the time of the oldest known rocks—was anoxygenic.
The free oxygen that makes up the bulk of the present atmosphere evolved over geologic time by two possible processes. First, solar ultraviolet radiation (the short-wavelength component of sunlight) would have provided the energy needed to break up water vapour into hydrogen, which escaped into space, and free oxygen, which remained in the atmosphere. This process was in all likelihood important before the appearance of the oldest extant rocks, but after that time the second process, organic photosynthesis, became predominant.

Primitive organisms, such as blue-green algae (or cyanobacteria), cause carbon dioxide and water to react by photosynthesis to produce carbohydrates, which they need for growth, repair, and other vital functions, and this reaction releases free oxygen.
The discovery of stromatolites (layered or conical sedimentary structures formed by sediment-binding marine algae) in 3.5-billion-year-old limestones in several parts of the world indicates that blue-green algae existed by that time. The presence of such early carbonate sediments is evidence that carbon dioxide was present in the atmosphere, and it has been calculated that it was at least 100 times greater than the amount in the present-day atmosphere. It can be assumed that such abundant carbon dioxide would have caused retention of heat, resulting in a greenhouse effect and a hot atmosphere.

What happened to all the oxygen that was released? It might be surprising to learn that it took at least 1 billion years before there was sufficient oxygen in the atmosphere for oxidative diagenesis to give rise to red beds (sandstones that are predominantly red in colour due to fully oxidized iron coating individual grains) and that 2.2 billion years passed before a large number of life-forms could evolve. An idea formulated by the American paleontologist Preston Cloud has been widely accepted as an answer to this question. The earliest primitive organisms produced free oxygen as a by-product, and in the absence of oxygen-mediating enzymes it was harmful to their living cells and had to be removed. Fortunately for the development of life on the early Earth there was extensive volcanic activity, which resulted in the deposition of much lava, the erosion of which released enormous quantities of iron into the oceans.

This ferrous iron is water-soluble and therefore could be easily transported, but it had to be converted to ferric iron, which is highly insoluble, before it could be precipitated as iron formations. In short, the organisms produced the oxygen and the iron formations accepted it. Iron formations can be found in the earliest sediments (those deposited 3.8 billion years ago) at Isua in West Greenland, and thus this process must have been operative by this time. Early Precambrian iron formations are so thick and common that they provide the major source of the world’s iron. Large quantities of iron continued to be deposited until about 2 billion years ago, after which time the formations decreased and disappeared from the sedimentary record. Sulfides also accepted oxygen in the early oceans to be deposited as sulfates in evaporites, but such rocks are easily destroyed. One finds, nonetheless, 3.5-billion-year-old barite/gypsum-bearing evaporites up to 15 metres thick and at least 25 kilometres in extent in the Pilbara region of Western Australia. It seems likely that the excess iron in the early oceans was finally cleared out by about 1.7 billion years ago, and this decrease in the deposition of iron formations resulted in an appreciable rise in the oxygen content of the atmosphere, which in turn enabled more eolian red beds to form.

Further evidence of the lack of oxygen in the early atmosphere is provided by detrital uraninite and pyrite and by paleosols—i.e., fossil soils. Detrital uraninite and pyrite are readily oxidized in the presence of oxygen and thus do not survive weathering processes during erosion, transport, and deposition in an oxygenous atmosphere. Yet, these minerals are well preserved in their original unoxidized state in conglomerates that have been dated to be more than 2.2 billion years old on several continents. Paleosols also provide valuable clues, as they were in equilibrium with the prevailing atmosphere.
  From analyses of early Precambrian paleosols it has been determined that the oxygen content of the atmosphere 2.2 billion years ago was one hundredth of the present atmospheric level (PAL).

Fossils of eukaryotes, which are organisms that require an oxygen content of about 0.02 PAL, bear witness to the beginning of oxidative metabolism. The first microscopic eukaryotes appeared about 1.4 billion years ago. Life-forms with soft parts, such as jellyfish and worms, developed in profusion, albeit locally, toward the end of the Precambrian about 650 million years ago, and it is estimated that this corresponds to an oxygen level of 0.1 PAL. By the time land plants first appeared, roughly 400 million years ago, atmospheric oxygen levels had reached their present values.

Development of the oceans

Volcanic degassing of volatiles, including water vapour, occurred during the early stages of crustal formation and gave rise to the atmosphere. When the surface of the Earth had cooled to below 100° C (212° F), the hot water vapour in the atmosphere would have condensed to form the early oceans. The existence of 3.5-billion-year-old stromatolites is, as noted above, evidence of the activity of blue-green algae, and this fact indicates that the Earth’s surface must have cooled to below 100° C by this time. Also, the presence of pillow structures in basalts of this age attests to the fact that these lavas were extruded under water, and this probably occurred around volcanic islands in the early ocean. The abundance of volcanic rocks of Archean age (3.8 to 2.5 billion years ago) is indicative of the continuing role of intense volcanic degassing, but since the early Proterozoic (from 2.5 billion years ago), much less volcanic activity has occurred.
Until about 2 billion years ago there was substantial deposition of iron formations, cherts, and various other chemical sediments, but from roughly that time onward the relative proportions of different types of sedimentary rock and their mineralogy and trace element compositions have been very similar to their Phanerozoic equivalents; it can be inferred from this relationship that the oceans achieved their modern chemical characteristics and sedimentation patterns from approximately 2 billion years ago. By the late Precambrian, some 1 billion years ago, ferric oxides were chemically precipitated, indicating the availability of free oxygen. During Phanerozoic time (the last 542 million years), the oceans have been steady-state chemical systems, continuously reacting with the minerals added to them via drainage from the continents and with volcanic gases at the oceanic ridges.

Time scales

The geologic history of the Earth covers nearly four billion years of time. Different types of phenomena and events in widely separated parts of the world have been correlated using an internationally acceptable, standardized time scale. There are, in fact, two geologic time scales. One is relative, or chronostratigraphic, and the other is absolute, or chronometric. The chronostratigraphic scale has evolved since the mid-1800s and concerns the relative order of strata. Important events in its development were the realization by William Smith that in a horizontal sequence of sedimentary strata what is now an upper stratum was originally deposited on a lower one and the discovery by James Hutton that an unconformity (discontinuity) indicates a significant gap in time. Furthermore, the presence of fossils throughout Phanerozoic sediments has enabled paleontologists to construct a relative order of strata. As was explained earlier, at specific stratigraphic boundaries certain types of fossils either appear or disappear or both in some cases. Such biostratigraphic boundaries separate larger or smaller units of time that are defined as eons, eras, periods, epochs, and ages.

The chronometric scale is of more recent origin. It was made possible by the development of mass spectrometers during the 1920s and their use in geochronological laboratories for radiometric dating. The chronometric scale is based on specific units of duration and on the numerical ages that are assigned to the aforementioned chronostratigraphic boundaries. The methods used entail the isotopic analyses of whole rocks and minerals of element pairs, such as potassium–argon, rubidium–strontium, uranium–lead, and samarium–neodymium. Another radiometric time scale has been developed from the study of the magnetization of basaltic lavas of the ocean floor. As such lavas were extruded from the mid-oceanic ridges, they were alternately magnetized parallel and opposite to the present magnetic field of the Earth and are thus referred to as normal and reversed. A magnetic-polarity time scale for the stratigraphy of normal and reversed magnetic stripes can be constructed back as far as the middle of the Jurassic Period, about 170 million years ago, which is the age of the oldest extant segment of ocean floor.


Precambrian Time


Precambrian time


Precambrian time, period of time that extends from about 4.6 billion years ago (the point at which Earth began to form) to the beginning of the Cambrian Period, 542 million years ago. The Precambrian represents more than 80 percent of the total geologic record.

All life-forms were long assumed to have originated in the Cambrian, and therefore all earlier rocks were grouped together into the Precambrian. Although many varied forms of life evolved and were preserved extensively as fossil remains in Cambrian sedimentary rocks, detailed mapping and examination of Precambrian rocks on most continents have revealed that additional primitive life-forms existed more than 3.4 billion years ago. Nevertheless, the original terminology to distinguish Precambrian rocks from all younger rocks is still used for subdividing geologic time.

The earliest evidence for the advent of life includes Precambrian microfossils that resemble algae, cysts of flagellates, tubes interpreted to be the remains of filamentous organisms, and stromatolites (sheetlike mats precipitated by communities of microorganisms). In the late Precambrian, the first multicellular organisms evolved, and sexual division developed. By the end of the Precambrian, conditions were set for the explosion of life that took place at the start of the Phanerozoic Eon.

The Precambrian environment

Several rock types yield information on the range of environments that may have existed during Precambrian time. Evolution of the atmosphere is recorded by banded-iron formations (BIFs), paleosols (buried soil horizons), and red beds, whereas tillites (sedimentary rocks formed by the lithification of glacial till) provide clues to the climatic patterns that occurred during Precambrian glaciations.


One of the most important factors controlling the nature of sediments deposited today is continental drift. This follows from the fact that the continents are distributed at different latitudes, and latitudinal position affects the temperature of oceanic waters along continental margins (the combined area of the continental shelf and continental slope); in short, sedimentary deposition is climatically sensitive. At present, most carbonates and oxidized red soils are being deposited within 30 degrees of the Equator, phosphorites within 45 degrees, and evaporites within 50 degrees. Most fossil carbonates, evaporites, phosphorites, and red beds of Phanerozoic age dating back to the Cambrian have a similar bimodal distribution with respect to their paleoequators. If the uniformitarian principle that the present is the key to the past is valid (meaning the same geologic processes occurring today occurred in the past), then sediments laid down during the Precambrian would have likewise been controlled by the movement and geographic position of the continents. Thus, it can be inferred that the extensive evaporites dating to 3.5 billion years ago from the Pilbara region of Western Australia could not have been formed within or near the poles. It can also be inferred that stromatolite-bearing dolomites of Riphean rock, a sedimentary sequence spanning the period from 1.65 billion to 800 million years ago, were deposited in warm, tropical waters. Riphean rock is primarily located in the East European craton, which extends from Denmark to the Ural Mountains, and in the Siberian craton in Russia.

Today, phosphate sediments are deposited primarily along the western side of continents. This is the result of high biological productivity in nearby surface waters due to the upwelling of nutrient-rich currents that are moving toward the Equator. The major phosphorite deposits of the Aravalli mountain belt of Rajasthan in northwestern India, which date from the Proterozoic Eon, are associated with stromatolite-rich dolomites. They were most likely deposited on the western side of a continental landmass that resided in the tropics.



During the long course of Precambrian time, the climatic conditions of the Earth changed considerably. Evidence of this can be seen in the sedimentary record, which documents appreciable changes in the composition of the atmosphere and oceans over time.

Oxygenation of the atmosphere

Earth almost certainly possessed a reducing atmosphere before 2.5 billion years ago. The Sun’s radiation produced organic compounds from reducing gases—methane (CH4) and ammonia (NH3). The minerals uraninite (UO2) and pyrite (FeS2) are easily destroyed in an oxidizing atmosphere; confirmation of a reducing atmosphere is provided by unoxidized grains of these minerals in 3.0-billion-year-old sediments. However, the presence of many types of filamentous microfossils dated to 3.45 billion years ago in the cherts of the Pilbara region suggests that photosynthesis had begun to release oxygen into the atmosphere by that time. The presence of fossil molecules in the cell walls of 2.5-billion year-old blue-green algae (cyanobacteria) establishes the existence of rare oxygen-producing organisms by that period.

Oceans of the Archean Eon (4.0 to 2.5 billion years ago) contained much volcanic-derived ferrous iron (Fe2+), which was deposited as hematite (Fe2O3) in BIFs. The oxygen that combined the ferrous iron was provided as a waste product of cyanobacterial metabolism. A major burst in the deposition of BIFs from 3.1 billion to 2.5 billion years ago—peaking about 2.7 billion years ago—cleared the oceans of ferrous iron. This enabled the atmospheric oxygen level to increaseappreciably. By the time of the widespread appearance of eukaryotes at 1.8 billion years ago, oxygen concentration had risen to 10 percent of present atmospheric level (PAL). These relatively high concentrations were sufficient for oxidative weathering to take place, as evidenced by hematite-rich fossil soils (paleosols) and red beds (sandstones with hematite-coated quartz grains). A second major peak, which raised atmospheric oxygen levels to 50 percent PAL, was reached by 600 million years ago. It was denoted by the first appearance of animal life (metazoans) requiring sufficient oxygen for the production of collagen and the subsequent formation of skeletons. Furthermore, in the stratosphere during the Precambrian, free oxygen began to form a layer of ozone (O3), which currently acts as a protective shield against the Sun’s ultraviolet rays.

Development of the ocean

The origin of Earth’s oceans occurred earlier than that of the oldest sedimentary rocks. The 3.85-billion-year-old sediments at Isua in western Greenland contain BIFs that were deposited in water. These sediments, which include abraded detrital zircon grains that indicate water transport, are interbedded with basaltic lavas with pillow structures that form when lavas are extruded under water. The stability of liquid water (that is, its continuous presence on Earth) implies that surface seawater temperatures were similar to those of the present.
Differences in the chemical composition of Archean and Proterozoic sedimentary rocks point to two different mechanisms for controlling seawater composition between the two Precambrian eons.

  During the Archean, seawater composition was primarily influenced by the pumping of water through basaltic oceanic crust, such as occurs today at oceanic spreading centres. In contrast, during the Proterozoic, the controlling factor was river discharge off stable continental margins, which first developed after 2.5 billion years ago. The present-day oceans maintain their salinity levels by a balance between salts delivered by freshwater runoff from the continents and the deposition of minerals from seawater.


A major factor controlling the climate during the Precambrian was the tectonic arrangement of continents. At times of supercontinent formation (at 2.5 billion, 2.1 to 1.8 billion, and 1.0 billion to 900 million years ago), the total number of volcanoes was limited; there were few island arcs (long, curved island chains associated with intense volcanic and seismic activity), and the overall length of oceanic spreading ridges was relatively short. This relative shortage of volcanoes resulted in low emissions of the greenhouse gas carbon dioxide (CO2). This contributed to low surface temperatures and extensive glaciations. In contrast, at times of continental breakup, which led to maximum rates of seafloor spreading and subduction (at 2.3 to 1.8 billion, 1.7 to 1.2 billion, and 800 to 500 million years ago), there were high emissions of CO2 from numerous volcanoes in oceanic ridges and island arcs. The atmospheric greenhouse effect was enhanced, warming Earth’s surface, and glaciation was absent. These latter conditions also applied to the Archean Eon prior to the formation of continents.

Temperature and rainfall

The discovery of 3.85-billion-year-old marine sediments and pillow lavas in Greenland indicates the existence of liquid water and implies a surface temperature above 0 °C (32 °F) during the early part of Precambrian time. The presence of 3.5-billion-year-old stromatolites in Australia suggests a surface temperature of about 7 °C (45 °F). Extreme greenhouse conditions in the Archean caused by elevated atmospheric levels of carbon dioxide from intense volcanism (effusion of lava from submarine fissures) kept surface temperatures high enough for the evolution of life. They counteracted the reduced solar luminosity (rate of total energy output from the Sun), which ranged from 70 to 80 percent of the present value. Without these extreme greenhouse conditions, liquid water would not have occurred on the Earth’s surface.

In contrast, direct evidence of rainfall in the geological record is very difficult to find. Some limited evidence has been provided by well-preserved rain pits in 1.8-billion-year-old rocks in southwestern Greenland.

Worldwide glaciations

The presence of tillites (glacial sediments) indicates that extensive glaciations occurred several times during the Precambrian. Glacial deposits are not necessarily limited to high latitudes. In general, they are complementary to the carbonates, evaporites, and red beds that are climatically sensitive and restricted to low latitudes.

The oldest known glaciation took place 2.9 billion years ago in South Africa during the Late Archean; the evidence is provided by glacial deposits in sediments of the Pongola Rift in southern Africa. The most extensive early Precambrian Huronian glaciation occurred 2.3 billion years ago during the early Proterozoic. It can be recognized from the rocks and structures that the glaciers and ice sheets left behind in parts of Western Australia, Finland, southern Africa, and North America. The most extensive occurrences are found in North America in a belt nearly 3,000 km (1,800 miles) long extending from Chibougamau in Quebec through Ontario to Michigan and southwestward to the Medicine Bow Mountains of Wyoming. This probably represents the area of the original ice sheet.

Most details are known from the Gowganda Formation in Ontario, which contains glacial deposits that are up to 3,000 metres (9,850 feet) thick and that occupy an area of about 20,000 square km (7,700 square miles); the entire glacial event may have covered an area of more than 2.5 million square km. Paleomagnetic studies indicate that the Gowganda glaciation occurred near the paleoequator. Similar, roughly contemporaneous glacial deposits can be found in other parts of the world, suggesting that there was at least one extensive glaciation during the early Proterozoic.

The largest glaciation in the history of the Earth occurred during the late Proterozoic in the period between 1 billion and 600 million years ago. It left its mark almost everywhere. One of the best-described occurrences is in the Flinders Range of South Australia, where there is a sequence 4 km (2.5 miles) thick of tillites and varved sediments occupying an area of 400 by 500 km (250 by 300 miles). Detailed stratigraphy and isotopic dating show that three worldwide glaciations took place: the Sturtian glaciation (750 to 700 million years ago), the Varanger-Marinoan ice ages (625 to 580 million years ago), and the Sinian glaciation (600 to 550 million years ago).

What is the explanation for all these occurrences of glacial deposits? Some paleomagnetic studies have shown that the tillites in Scotland, Norway, Greenland, central Africa, North America, and South Australia were deposited in low or near-equatorial paleolatitudes. Such conclusions are, however, controversial, because it has also been suggested that the positions of the northern and southern magnetic poles may have migrated across the globe, leaving a record of glaciations in both high and low latitudes. There is the possibility that floating ice sheets could have traveled to low latitudes, depositing glacial sediments and dropstones below them.

Whatever the answer, the existence of such vast quantities of tillites and of such extensive glaciations is intriguing. It has been suggested that they record the existence of a frozen “snowball” Earth.
Precambrian life

Precambrian rocks were originally defined to predate the Cambrian Period and therefore all life, although the term Proterozoic was later coined from the Greek for “early life.” It is now known that Precambrian rocks contain evidence of the very beginnings of life on Earth (and thus the record of its evolution for more than 3.5 billion years), the explosion of life-forms without skeletons before the Cambrian, and even the development of sexual reproduction.


The earliest signs of life on Earth are in western Greenland where apatite (calcium phosphate) grains within a 3.85-billion-year-old meta-sedimentary rock have carbon isotope ratios that indicate an organic origin. The presence of organic hydrocarbon droplets in kerogenous sediments has been found in the 3.46-billion-year-old Warrawoona Group in the Pilbara craton of Western Australia. These are small amounts of Archean oil.
The first fossil evidence of terrestrial life is found in the early Archean sedimentary rocks of the greenstone-granite belts (metamorphosed oceanic crust and island arc complexes) of the Barberton craton in South Africa and in the Warrawoona Group, which are both about 3.5 billion years old. There are two types of these early, simple, biological structures: microfossils and stromatolites (sheetlike mats precipitated by communities of microorganisms).

Microfossils and stromatolites

The microfossils occur in cherts and shales and are of two varieties. One type consists of spherical carbonaceous aggregates, or spheroids, which may measure as much as 20 mm (0.8 inch) in diameter. These resemble algae and cysts of flagellates and are widely regarded as biogenic (produced by living organisms). The other variety of microfossils is made up of carbonaceous filamentous threads, which are curving, hollow tubes up to 150 micrometres (0.006 inch) long. Most likely, these tubes are the fossil remains of filamentous organisms. Hundreds of them have been found in some rock layers. The 2.8-billion-year-old gold reefs (conglomerate beds with rich gold deposits) of the Witwatersrand Basin in South Africa contain carbonaceous columnar microfossils up to 7 mm (slightly less than 0.3 inch) long that resemble modern algae, fungi, and lichens. They probably extracted gold from their environment in much the same way that modern fungi and lichens do.

Stromatolites are stratiform, domal, or columnar structures made from sheetlike mats precipitated by communities of microorganisms, particularly filamentous blue-green algae. The early Archean examples form domes as tall as about 10 cm (4 inches). Stromatolites occur in many of the world’s greenstone-granite belts. In the 2.7-billion-year-old Steep Rock Lake belt in Ontario, Can., they reach 3 metres (9 feet) in height and diameter.
Stromatolites continued to form all the way through the geologic record and today grow in warm intertidal waters, as exemplified by those of Shark Bay in Western Australia. They provide indisputable evidence that life had begun on Earth using algal photosynthesis in complex, integrated biological communities by 3.5 billion years ago.
These Archean organisms were prokaryotes that were incapable of cell division. They were relatively resistant to ultraviolet radiation and thus were able to survive during Earth’s early history when the atmosphere lacked an ozone layer. The prokaryotes were predominant until about 1.7 billion to 1.9 billion years ago, when they were overtaken by the eukaryotes (organisms possessing nucleated cells). The latter made use of oxygen in metabolism and for growth and thus developed profusely in the increasingly oxygen-rich atmosphere of the early Proterozoic. The eukaryotes were capable of cell division, which allowed DNA (deoxyribonucleic acid), the genetic coding material, to be passed on to succeeding generations.

By early Proterozoic time both microfossils and stromatolites had proliferated. The best-known occurrence of microorganisms is in the 2-billion-year-old, stromatolite-bearing Gunflint iron formation in the Huronian Basin of southern Ontario. These microbial fossils include some 30 different types with spheroidal, filamentous, and sporelike forms up to about 20 micrometres (0.0008 inch) across. Sixteen species in 14 genera have been classified so far. Microfossils of this kind are abundant, contain beautifully preserved organic matter, and are extremely similar to such present-day microorganisms as blue-green algae and microbacteria. There are comparable microfossils from the early Proterozoic in Minnesota and Michigan in the United States, the Belcher Islands in Hudson Bay in Canada, southern Greenland, Western Australia, and northern China. These microbiota lived at the time of the transition in the chemical composition of the atmosphere when oxygen began accumulating for the first time.

During the late Proterozoic, stromatolites reached their peak of development, became distributed worldwide, and diversified into complex, branching forms. From about 700 million years ago, however, they began to decline significantly in number. Possibly the newly arrived metazoans (multicelled organisms whose cells are differentiated into tissues and organs) ate the stromatolitic algae, and their profuse growth destroyed the habitats of the latter.
  There is the intriguing question as to when sexual division arose in life-forms. In the late 1960s, American paleobiologist J. William Schopf pointed out that the abundant microflora of the 900-million-year-old Bitter Springs Formation of central Australia includes some eukaryotic algae that have cells in various stages of division arranged into tetrahedral sporelike forms. These resemble the tetrad of spore cells of living plants known to develop by sexual division. In effect, by the end of the Precambrian the conditions were set for the explosion of life at the start of the Phanerozoic Eon.

Ediacaran fossils

Metazoans developed rapidly from the beginning of the Cambrian, when organisms acquired the ability to produce the protein collagen and, thus, skeletons and shells. However, more-primitive metazoans without skeletons—the Ediacara fauna—appeared earlier (more than 600 million years ago), after the end of the Varanger-Marinoan ice age at 580 million years ago and before the onset of the Cambrian Period at 542 million years ago. They are found as impressions of soft-bodied, multicellular animals in the rocks and have the form of tiny blobs, circular discs, or plantlike fronds ranging from less than 1 cm (less than 0.4 inch) to more than 1 metre (about 3 feet) long.

Precambrian geology

Major subdivisions of the Precambrian System

By international agreement, Precambrian time is divided into the Archean Eon (occurring between roughly 4.0 billion years ago and 2.5 billion years ago) and Proterozoic Eon (occurring between 2.5 billion and 542 million years ago). After the Precambrian, geologic time intervals are commonly subdivided on the basis of the fossil record. The paucity of Precambrian fossils, however, precludes the creation of small-scale subdivisions (epochs and ages) in this time period. Instead, relative chronologies of events have been produced for different regions based on such field relationships as unconformities (interruption in the accumulation of sedimentary rock due to erosion or nondeposition) and crosscutting dikes (intrusions of igneous rock that burrow through cracks in the original structures of surrounding rock).

These field relationships, combined with the isotopic age determinations of specific rocks, allow for some correlation between neighbouring regions.

The International Commission on Stratigraphy (ISC) and International Union of Geological Sciences (IUGS) divide the Archean Eon into the Eoarchean (approximately 4.0 billion to 3.6 billion years ago), Paleoarchean (3.6 billion to 3.2 billion years ago), Mesoarchean (3.2 billion to 2.8 billion years ago), and Neoarchean (2.8 billion to 2.5 billion years ago) eras. Likewise, they divide the Proterozoic Eon into the Paleoproterozoic (2.5 billion to 1.6 billion years ago), Mesoproterozoic (1.6 billion to 1 billion years ago), and Neoproterozoic (1 billion to 542 million years ago) eras. These definitions are based on isotopic age determinations.

Oldest minerals and rocks

The oldest minerals on Earth, detrital zircons from western Australia, crystallized about 4.4 billion years ago. They occur within sedimentary sandstones and conglomerates dated to about 3.3 billion years ago, but the environment in which they were formed is totally unknown. The rocks from which they came may have been destroyed by some kind of tectonic process or by a meteorite impact that spared individual zircon crystals. On the other hand, rocks containing these minerals may still exist on Earth’s surface but simply have not been found. Perhaps their very absence is indicative of something important about early terrestrial processes. Comparisons with the Moon indicate that the Earth must have been subjected to an enormous number of meteorite impacts about 4 billion years ago, but there is no geologic evidence of such events.

The oldest known rocks on Earth are the faux amphibolite volcanic deposits of the Nuvvuagittuq greenstone belt in Quebec, Canada; they are estimated to be 4.28 billion years old. The age of these rocks was estimated using a radiometric dating technique that measures the ratio of the rare-earth elements neodymium and samarium present in a sample.

The Acasta gneisses, found near Canada’s Great Slave Lake, are also among the world’s oldest rocks. Their age has been established radiometrically at 4.0 to 3.9 billion years. The Acasta gneisses are granitic and contain a single relict zircon crystal, which has been dated to 4.2 billion years ago and formed from granitic magma. They are thought to have evolved from older basaltic material in the crust that was melted and remelted by tectonic processes.


Significant geologic events


The Archean and Proterozoic eons within Precambrian time are very different and must be considered separately. The Archean-Proterozoic boundary constitutes a major turning point in Earth history. Before that time the crust of the Earth was in the process of growing, and so there were no large, stable continents. Afterward, when such continents had emerged, orogenic belts were able to form on the margins of and between continental blocks.
There are two types of Archean orogenic belts. The first occurs in upper crustal greenstone-granite belts rich in volcanic rocks that are probably primitive types of oceanic crust and island arcs (long, curved island chains associated with intense volcanic and seismic activity) that formed during the early rapid stage of crustal growth. The second occurs in granulite-gneiss belts that were recrystallized in the Archean mid-lower crust under metamorphic conditions associated with high-temperature granulite and amphibolite facies. Thus, granulites, which typically contain the high-temperature mineral hypersthene (a type of pyroxene), are a characteristic feature of many Precambrian orogenic belts that have been deeply eroded. In Phanerozoic orogenic belts, granulites are rare.

There are several other rock types that developed primarily during the Precambrian but rarely later. This restriction is a result of the unique conditions that prevailed during Precambrian time. For example, banded-iron formations are ferruginous sediments that were deposited on the margins of early, iron-rich oceans. Anorthosite, which consists largely of plagioclase, forms large bodies in several Proterozoic belts. Komatiite, a magnesium-rich, high-temperature volcanic rock derived from very hot mantle (part of the Earth between the crust and the core), was extruded in abundance during the early Precambrian when the heat flow of the Earth was higher than it is today. Blueschist, which contains the blue mineral glaucophane, forms in subduction zones under high pressures and low temperatures, and its rare occurrence in Precambrian rocks may indicate that temperatures in early subduction zones were too high for its formation.

The bulk of many of the world’s valuable mineral deposits (for example, those of gold, nickel, chromite, copper, and iron) also formed during the Precambrian. These concentrations are a reflection of distinctive Precambrian sedimentary and magmatic rocks and their environments of formation.


During the first third of geologic history (that is, until about 2.5 billion years ago), the Earth developed in a broadly similar manner. Greenstone-granite belts (metamorphosed oceanic crust and island arc complexes) formed in the upper Archean crust, and granulite-gneiss belts formed in the mid-lower crust. This was a time when the overall rate of heat production by the breakdown of radioactive isotopes was several times greater than it is today. This condition was manifested by very rapid tectonic processes, probably by some sort of primitive plate tectonics (more-modern plate-tectonic processes could not occur until the crust became cooler and more rigid). Most of the heat that escapes from Earth’s interior today does so at oceanic ridges. This manner of heat loss probably occurred during the Archean in much larger amounts. The oceanic ridges of the Archean were more abundant, longer, and opened faster than those in the modern oceans, and oceanic plateaus derived from hot mantle plumes (slowly rising currents of highly viscous mantle material) were more common. Although the amount of newly generated crust was probably enormous, a large part of this material was inevitably destroyed by equally rapid plate subduction processes. The main results of this early growth that still remain today are the many island arcs and oceanic plateaus in greenstone-granite belts and the voluminous Andean-type tonalites (a granitic-type rock rich in plagioclase feldspar) that were deformed to orthogneiss (gneiss derived from igneous rocks) in granulite-gneiss belts. Although most of the Archean oceanic crust was subducted, a few ophiolitic-type complexes have been preserved in greenstone-granite belts.

The late Archean (Neoarchean Era) was an important interval of time because it marks the beginning of the major changeover from Archean to Proterozoic types of crustal growth. The formation of the first major rifts characterized the significant events of this time. The first major rift valley known in the world, the Pongola Rift, emerged along the border of present-day Swaziland and South Africa; the intrusion of the first major basic dikes (such as the Great Dyke, which transects the entire Zimbabwe craton) and the first large stratiform layered igneous complexes (such as the Stillwater in Montana) formed; and the formation of the first large sedimentary basins (for example, the Witwatersrand in South Africa) also occurred. All of these structures indicate that the continental crust had reached a mature stage with considerable stability and rigidity for the first time during the late Archean. The Neoarchean represents the culmination that followed the rapid tectonic processes of the early Archean (Eoarchean and Paleoarchean) and middle Archean (Mesoarchean) eras. Because crustal growth took place at different times throughout the world, similar structures can be found in the early Proterozoic (Paleoproterozoic) Era.


During the early Proterozoic, large amounts of quartzite, carbonate, and shale were deposited on the shelves and margins of many continental blocks. This would be consistent with the breakup of a supercontinent into several smaller continents with long continental margins (combined areas of continental shelf and continental slope). Examples of shelf sequences of this kind are found along the margins of orogenic (mountain) belts, such as the Wopmay, bordering Canada’s Slave province, and also the Labrador Trough, bordering the Superior province.
The existence of stable continental blocks by the early Proterozoic allowed orogenic belts to develop at their margins by some form of collision tectonics. This was the first time that long, linear orogenic belts could form by “modern” tectonic processes that involved seafloor spreading, ophiolite obduction, subduction, and landmass collisions. Subduction lead to the creation of island arcs and Andean-type (formed by subduction at the continental margin) granitic batholiths. In addition, the collision of arcs and continents could now give rise to both sutures with ophiolites and to Himalayan-type (formed by continent-to-continent collision) thrust belts with abundant crustal-melt granites. These were key events in the evolution of the continents, and such processes have continued throughout Earth history.

During the late Proterozoic (Neoproterozoic Era), some orogenic belts, like the Pan-African belts of Saudi Arabia and East Africa, continued to develop. The intense crustal growth and the many orogenic belts that formed throughout the Proterozoic began to create large continental blocks, which amalgamated to produce a new supercontinent by the end of the Precambrian. Therefore, in the late Proterozoic many sedimentary basins were infilled with conglomerates and sandstones due to the deposition of material eroded from higher elevations. For example, the Riphean sequence in Russia and also the Sinian sequence in China were able to form on extensive cratons of continental crust.

Occurrence and distribution of Precambrian rocks

Precambrian rocks, as a whole, occur in a wide variety of shapes and sizes. There are extensive Archean regions, up to a few thousands of kilometres across, that may contain either greenstone-granite belts or granulite-gneiss belts or both. These regions are variously designated in different parts of the world as cratons, shields, provinces, or blocks. Some examples include: the North Atlantic craton that incorporates northwestern Scotland, central Greenland, and Labrador; the Kaapvaal and Zimbabwean cratons in southern Africa; the Dharwar craton in India; the Aldan and Anabar shields in Siberia in Russia; the Baltic Shield that includes much of Sweden, Finland, and the Kola Peninsula of far northern Russia; the Superior and Slave provinces in Canada; and the Yilgarn and Pilbara blocks in Western Australia. Linear belts, up to several thousand kilometres long, that are frequently though not exclusively of Proterozoic age include the Limpopo, Mozambique, and Damaran belts in Africa, the Labrador Trough in Canada, and the Eastern Ghats belt in India. Several small relict areas, spanning a few hundred kilometres across, exist within or against Phanerozoic orogenic belts and include the Lofoten islands of Norway, the Lewisian Complex in northwestern Scotland, and the Adirondack Mountains in the northeastern United States. Nevertheless, some extensive areas of Precambrian rocks, such as under the European and Russian platforms and under the central United States, remain overlain by a blanket of Phanerozoic sediments.

Archean rock types

Archean rocks occur in greenstone-granite belts that represent the upper crust, in granulite-gneiss belts that formed in the mid-lower crust, and in sedimentary basins, basic dikes, and layered complexes that were either deposited on or intruded into the first two types of belts.


These belts occur on most continents. The largest extend several hundred kilometres in length and measure several hundred metres in width. Today many greenstone-granite belts are regarded as tectonic “slices” of oceanic and island arc crust that have been thrust together to form tectonic collages similar to those in belts found in the present-day Pacific Ocean.

The greenstone sequence in many belts is divisible into a lower volcanic group and an upper sedimentary group. The volcanics are made up of lavas that are ultramafic (silica content less than 45 percent) and basaltic (silica content of 45 to 52 percent). The uppermost sediments are typically terrigenous (land-derived) shales, sandstones, quartzites, wackes, and conglomerates. All the greenstone sequences have undergone recrystallization during the metamorphism of greenschist facies at relatively low temperatures and pressures. In fact, the presence of the three green metamorphic minerals chlorite, hornblende, and epidote has given rise to the term greenstone for the recrystallized basaltic volcanics. Granitic rocks and gneisses occur within, adjacent to, and between many greenstone sequences.

Economic significance of Archean greenstone-granite deposits

Abundant mineralization has occurred in greenstone-granite belts. These belts constitute one of the world’s principal depositories of gold, silver, chromium, nickel, copper, and zinc. In the past they were termed gold belts because of the gold rushes of the 19th century that took place in areas such as Kalgoorlie in the Yilgarn belt of Western Australia, the Barberton belt of South Africa, and Val d’Or in the Abitibi belt of southern Canada. The mineral deposits occur in all the major rock groups: chromite, nickel, asbestos, magnesite, and talc in ultramafic lavas; gold, silver, copper, and zinc in basaltic to rhyolitic volcanics; iron ore, manganese, and barite in sediments; and lithium, tantalum, beryllium, tin, molybdenum, and bismuth in granites and associated pegmatites. Important occurrences are chromite at Selukwe in Zimbabwe, nickel at Kambalda in southwestern Australia, tantalum in Manitoba in Canada, and copper-zinc at Timmins and Noranda in the Canadian Abitibi belt.

Greenstone-granite rock types

The volcanics that comprise the lower portion of a greenstone sequence are made up of lavas noted for magnesian komatiites (ultramafic extrusive igneous rocks) that probably formed in the oceanic crust that are overlain by basalts, andesites, and rhyolites whose chemical composition is much like that of modern island arcs. Especially important is the presence in the Isua, Barberton, and Yellowknife belts of sheeted basic dike complexes cutting across gabbros and overlain by pillow-bearing basalts (basalts extruded underwater that form characteristic pillow-shaped hummocks). Volcanic sequences are capped by oceanic cherts and terrigenous sedimentary groups. The overall stratigraphy suggests an evolution from extensive submarine eruptions of komatiite and basalt (ocean floor) to more-localized stratovolcanoes (volcanoes constructed from alternating layers of ash and lava), which become increasingly emergent with intervening and overlying clastic sediments (clay-, silt-, and sand-sized sediments) that were deposited in trenches at the mouths of subduction zones. There are, however, regional differences in the volcanic and sedimentary makeup of some belts. The older belts in southern Africa and Australia have more komatiites, basalts, shallow-water banded-iron formations, cherts, and evaporites and fewer terrigenous (land-derived) sediments. On the other hand, the younger belts in North America have a higher proportion of andesites, rhyolites, and terrigenous and turbidite debris (sediments delivered to the deep ocean by density currents) but fewer shallow-water sediments. These differences reflect a change from the older oceanic-type volcanism (effusion of lava from submarine fissures) to the younger, more arc-type phenomena such as explosive eruption of pyroclastic materials (incandescent material ejected during violent eruptions) and lava flows from steep volcanic cones. Additional changes include an increase in the amount of trench (subduction zone) turbidites and graywackes and an increase in the availability of continental crust as a source for terrigenous debris.
  Ultramafic rocks (rocks with a very low silica content—less than 45 percent) are commonly altered to talc schists and tremolite-actinolite schists. There are some indications that several phases of metamorphism exist—namely, seafloor metamorphism associated with the action of hydrothermal brines that could occur at oceanic ridges, syntectonic metamorphism related to thrust-nappe tectonics, and local thermal contact metamorphism caused by intrusive granitic plutons pushing into cooler surrounding rock.

Structure and formation of greenstone-granite belts

The structure of many belts is complex. Their stratigraphic successions are upside-down and deformed by thrusts and major horizontal folds (nappes). They have been subsequently refolded by upright anticlines (convex folds of rock) and synclines. The result of this thrusting is the repetition of the same stratigraphic successions on top of one another, creating a massive deposit of material up to 10 to 20 km (6 to 12 miles) thick. Also, there may be thrusts along the base of the belts, as in the case of Barberton, showing that they have been transported from elsewhere. In other instances, the thrusts may occur along the borders of the belts, indicating that they have been forced against and over adjacent gneissic belts. The conclusion from structural studies is that many belts have undergone intense subhorizontal deformation during thrust transport.

Clearly, there are different types of greenstone-granite belts. To understand their origin and mode of evolution, it is necessary to correlate them with comparable modern analogues. Some, like the Barberton and Yellowknife belts, consist of oceanic-type crust and have sheeted dike swarms that occur in many ophiolites of Mesozoic-Cenozoic origin, such as in the Troodos Mountains in Cyprus. They are the hallmark of a modern oceanic crust that formed at an oceanic ridge. Also, like modern ophiolites, a few seem to have been covered by thrusting onto continental crust. Many belts, such as the Isua belt of Greenland and those in the Superior province of Canada, are very similar to modern island arcs. Geochemical data are revealing that some lavas were derived from depths of 1,000 to 2,700 km (620 to 1,680 miles) in the Earth’s mantle and not from shallower subduction zones, which are commonly 600 km (about 373 miles) deep. These rocks are comparable to oceanic plateaus in modern oceanic crust that were formed from plumes of hot magma from the very deep mantle. The Wawa belt, for example, has been shown to consist of an immature island arc built on oceanic plateau crust and overlain by a more mature arc. The Abitibi belt began as oceanic crust with island arcs and oceanic plateaus. Between the Wawa and Wabigoon island arcs lies the Quetico belt, consisting of metamorphosed turbidites and slices of volcanics that probably developed in a regularly overlapping accretionary prism in an arc-trench system, as seen today in the Japanese arcs. The Pilbara belts are similar to modern active continental margins, and they have been interthrust with older continental orthogneisses to form very thick crustal piles intruded by diapiric crustal-melt granites. This scenario is quite comparable to that of a Himalayan type of orogenic belt formed by collisional tectonics. In conclusion, most greenstone-granite belts are today regarded by geologists as different parts of interthrust oceanic crust–accretionary prism structures within island arcs of oceanic plateau systems that collided with continental gneissic blocks.

Age and occurrence of greenstone-granite belts

Greenstone-granite belts developed at many different times throughout the long Archean Eon. The Isua greenstone belt in West Greenland is about 3.85 billion years old. In the Zimbabwean craton, they formed over three successive periods: the Selukwe belt about 3.8 to 3.75 billion years ago, the Belingwean belts about 2.9 billion years ago, and the Bulawayan-Shamvaian belts about 2.7 to 2.6 billion years ago. The Barberton belt in the Kaapvaal craton and the Warrawoona belt in the Pilbara block are 3.5 billion years old. Globally, the most important period of formation was from 2.7 to 2.6 billion years ago, especially in the Slave and Superior provinces of North America, the Yilgarn block in Australia, and the Dharwar craton in India. Some of the better-documented belts seem to have formed within about 50 million years. It is important to note that while the Bulawayan-Shamvaian belts were forming in the Zimbabwean craton, flat-lying sediments and volcanics were laid down in the Pongola Rift and the Witwatersrand Basin not far to the north.

Greenstone-granite belts range from aggregates of several belts (as in the southern Superior province of Canada) to irregular, even triangular-shaped belts (as in the Barberton in South Africa) to synclinal basins (as in the Indian Dharwar craton). The irregular and synclinal shapes are commonly caused by the diapiric intrusion of younger granites.

Important occurrences are the Barberton belt in South Africa; the Sebakwian, Belingwean, and Bulawayan-Shamvaian belts of Zimbabwe; the Yellowknife belts in the Slave province of Canada; the Abitibi, Wawa, Wabigoon, and Quetico belts of the Superior province of Canada; the Dharwar belts in India; and the Warrawoona and Yilgarn belts in Australia.


The granulites, gneisses, and associated rocks in these belts were metamorphosed to a high grade in deep levels of the Archean crust; metamorphism occurred at a temperature of 750 to 980 °C (1,380 to 1,800 °F) and at a depth of about 15 to 30 km (9 to 19 miles). These belts, therefore, represent sections of the continents that have been highly uplifted, with the result that the upper crust made up of volcanics, sediments, and granites has been eroded. Accordingly, the granulite-gneiss belts are very different from the greenstone-granite belts. Granulite-gneiss belts may be regarded as variably preserved sections of continental cratons.

Economic significance of Archean granulite-gneiss deposits

The mid-lower crust is relatively barren of ore deposits as compared to the upper crust with its sizable concentrations of greenstones and granites, and therefore little mineralization is found in the granulite-gneiss belts. The few exceptions include a nickel–copper sulfide deposit at Selebi-Pikwe in the Limpopo belt in Botswana that is economic to mine, and banded-iron formations in gneisses in the eastern Hubei and Liaoning provinces of northwestern China that form the foundation of a major steel industry. There are subeconomic quantities of chromitite in the anorthosites of western Greenland, southern India, and the Limpopo belt; iron from a banded-iron formation at Isua in western Greenland; and tungsten in amphibolites of western Greenland.

Granulite-gneiss rock types

Orthogneisses of deformed and recrystallized tonalite (a granitic-type rock rich in plagioclase feldspar) and granite constitute the most common rock type. The geochemical signature of these rocks closely resembles that of modern equivalents that occur in granitic batholiths in the Andes. Where such rocks have been metamorphosed under conditions associated with amphibolite facies, they contain hornblende, biotite, or a combination of the two. However, where they have been subjected to conditions of higher temperature associated with the granulite facies, the rocks contain pyroxene and hypersthene and so can be called granulites.

The granulites and gneisses enclose a wide variety of other minor rock types in layers and lenses. These types include schists and paragneisses that were originally deposited on the Earth’s surface as shales and which now contain high-temperature metamorphic minerals such as biotite, garnet, cordierite, staurolite, sillimanite, or kyanite. There also are quartzites, which were once sandstones or cherts; marbles (either limestones or dolomites); and banded-iron formations. Commonly intercalated with these metasediments are amphibolites, which locally contain relict pillow structures, demonstrating that they are derived from basaltic lavas extruded underwater. These amphibolites have a trace element chemistry quite similar to that of modern seafloor basalts. The amphibolites are often accompanied by chromite-layered anorthosite, gabbro, and ultramafic rocks such as peridotite and dunite. All these rocks occur in layered igneous complexes, which in their well-preserved state may be up to 2 km (1.2 miles) thick and 100 km (60 miles) long. Such complexes occur at Fiskenaesset in western Greenland, in the Limpopo belt of southern Africa, and in southern India. These complexes may have formed at an oceanic ridge in a magma chamber that also fed the basaltic lavas, or they may be parts of oceanic plateaus. In many cases, the complexes, basaltic amphibolites, and sediments were extensively intruded by the tonalites and granites that were later deformed and recrystallized.
The result of this is that all of these rocks may now occur as metre-sized lenses in the orthogneisses and granulites.

Structure and occurrence of granulite-gneiss belts

The structure of the granulite-gneiss belts is extremely complex because the constituent rocks have been highly deformed several times. In all likelihood the basalts and layered complexes from the oceanic crust were interthrust with shallow-water limestones, sandstones, and shales; with tonalites and granites from Andean-type batholiths; and with older basement rocks from a continental margin. All these rocks, which are now mutually conformable (parallel to one another with uninterrupted deposition), were folded in horizontal nappes and then refolded. The picture that emerges is one of a very mobile Earth, where newly formed rocks were routinely compressed and thrust against other rocks.

Granulite-gneiss belts occur in a variety of environments. These may be extensive regions, such as the North Atlantic craton, which measures 1,000 by 2,000 km (about 620 by 1,240 miles) across and, before the opening of the Atlantic Ocean, was contiguous with the Scourian Complex of northwestern Scotland, the central part of Greenland, and the coast of Labrador; the Aldan and Ukrainian shields of continental Europe; the North China craton; large parts of the Superior province of Canada; the Yilgarn block in Australia; and the Limpopo belt in southern Africa. They may be confined to small areas such as the Ancient Gneiss Complex of Swaziland, the Minnesota River valley and the Beartooth Mountains of the United States, the Peninsular gneisses and Sargur supracrustals of southern India, the English River gneisses of Ontario in Canada that form a narrow strip between greenstone-granite belts, the Sand River gneisses that occupy a small area between greenstone-granite belts in Zimbabwe, and the Napier Complex in Enderby Land in Antarctica. Granulite-gneiss belts are commonly surrounded by younger, mostly Proterozoic belts that contain remobilized relicts of the Archean rocks, and the granulites and gneisses must underlie many Archean greenstone-granite belts and blankets of Phanerozoic sediment.

Age and correlation of granulite-gneiss belts

Isotopic age determinations from the granulite-gneiss belts record an evolution from about 4.0 to 2.5 billion years ago—more than a third of geologic time. Most important are the few but well-constrained age determinations of detrital zircons at Mount Narryer and Jack Hills in Western Australia that are more than 4 billion years old. Several regions have a history that began in the period dating from 3.9 to 3.6 billion years ago—western Greenland, Labrador, the Limpopo belt, Enderby Land, the North China craton, and the Aldan Shield. Most regions of the world experienced a major tectonic event that may have involved intrusion, metamorphism, and deformation during the period between 3.1 and 2.8 billion years ago; some of these regions, like the Scourian in northwestern Scotland, show no evidence of any older crustal growth. The best-documented region is in western Greenland, which has a long and complicated history from 3.85 to 2.5 billion years ago.

It is impossible to correlate the rocks in different granulite-gneiss belts. One granitic gneiss is essentially the same as another but may be of vastly different age. There is a marked similarity in the anorthosites in various belts throughout the world, and their similar relationship with the gneisses suggests that the belts have undergone comparable stages of evolution, although each has its own distinctive features. Little correlation can be made with rocks of Mesozoic-Cenozoic age because few modern orogenic belts have been eroded sufficiently to expose their mid-lower crust. The lack of modern analogues for comparison makes it particularly difficult to interpret the mode of origin and evolution of the Archean granulite-gneiss belts.


During middle and late Archean time (3 to 2.5 billion years ago), relatively stable, post-orogenic conditions developed locally in the upper crust, especially in southern Africa, where the development of greenstone-granite and granulite-gneiss belts was completed much earlier than in other parts of the world. The final chapters of Archean crustal evolution can be followed by considering specific key sedimentary basins, basic (basaltic) dikes, and layered complexes.

Along the border of Swaziland and South Africa is the Pongola Rift, which is the oldest such continental trough in the world; it is 2.95 billion years old, having formed only 50 million years after the thrusting of adjacent greenstone-granite belts. If there were earlier rifts, they have not survived, or, more likely, this was the first time in Earth history that the upper crust was sufficiently stable and rigid for a rift to form. It is 30 km (19 miles) wide, 130 km (81 miles) long, and within it is a sequence of lavas and sediments that is 11 km (7 miles) thick. It seems most likely that the rift developed as the result of the collapse of an overthickened crust following the long period of Archean crustal growth and thrusting in the Kaapvaal craton.

The 200-by-350-km (124-by-217-mile) Witwatersrand Basin contains an 11-km- (7-mile-) thick sequence of lavas and sediments that are 3 billion years old. The basin is famous for its very large deposits of gold and uranium that occur as detrital minerals in conglomerates. These minerals were derived by erosion of the surrounding greenstone-granite belts and transported by rivers into the shoreline of the basin. In all probability, the gold originally came from the komatiitic and basaltic lavas in the early Archean oceanic crust.
The Great Dyke, thought to be about 2.5 billion years old, transects the entire Zimbabwe craton. It is 480 km (about 300 miles) long, 8 km (5 miles) wide, and made up of layered ultrabasic rocks—gabbros and norites. The ultrabasic rocks have several layers of chromite and an extensive platinum-bearing layer that form economic deposits. The Great Dyke represents a rift that has been filled in with magma that was probably derived from a deep mantle plume.
The Stillwater Complex is a famous, 2.7-billion-year-old, layered ultrabasic-basic intrusion in the Beartooth Mountains of Montana in the United States. It is 48 km (30 miles) long and has a stratigraphic thickness of 6 km (3.7 miles). It was intruded as a subhorizontal body of magma that underwent crystal settling to form the layered structure. It is notable for a 3-metre- (9-foot-) thick layer enriched in platinum minerals that forms a major economic deposit. The basins, dikes, and complexes described above cannot be mutually correlated. They most resemble equivalent structures that formed at the end of plate-tectonic cycles in the Phanerozoic. They represent the culmination of Archean crustal growth.

Proterozoic rock types

What happened geologically at the time of the Archean-Proterozoic boundary 2.5 billion years ago is uncertain. It seems to have been a period of little tectonic activity, and so it is possible that the earlier intensive Archean crustal growth had caused the amalgamation of continental fragments into a supercontinent, perhaps similar to Pangea of Permian-Triassic times. The fragmentation of this supercontinent and the formation of new oceans gave rise to many continental margins upon which a variety of distinctive sediments were deposited. Much evidence suggests that in the period from 2.5 billion to 570 million years ago Proterozoic oceans were formed and destroyed by plate-tectonic processes and that most Proterozoic orogenic belts arose by collisional tectonics. Sedimentary, igneous, and metamorphic rocks that formed during this period are widespread throughout the world. There are many swarms of basic dikes, important sedimentary rifts, basins, and layered igneous complexes, as well as many orogenic belts. The rocks commonly occur in orogenic belts that wrap around the borders of Archean cratons. The characteristic types of Proterozoic rocks are considered below, as are classic examples of their occurrence in orogenic belts. The following types of rocks were formed during the early, middle, and late Proterozoic, indicating that similar conditions and environments existed throughout this long period of time.


The continents were sufficiently stable and rigid during the Proterozoic Eon for an extremely large number of basic dikes to be intruded into parallel, extensional fractures in major swarms. Individual dikes measure up to several hundred metres in width and length, and there may be hundreds or even thousands of dikes in a swarm, some having transcontinental dimensions. For example, the 1.2-billion-year-old Mackenzie swarm is more than 500 km (311 miles) wide and 3,000 km (1,864 miles) long and extends in a northwesterly direction across the whole of Canada from the Arctic to the Great Lakes. The 1.95-billion-year-old Kangamiut swarm in western Greenland is only about 250 km (155 miles) long but is one of the world’s densest continental dike swarms. Many of the major dike swarms were intruded on the continental margins of Proterozoic oceans in a manner similar to the dikes that border the present-day Atlantic Ocean and were similarly the result of the rise of mantle plumes into the crust.


There are several very important layered, mafic to ultramafic intrusions of Proterozoic age that were formed by the accumulation of crystals in large magma chambers. The well-known ones are several tens or even hundreds of kilometres across, have a dikelike or sheetlike (stratiform) shape, and contain major economic mineral deposits. The largest and most famous is the Bushveld Complex in South Africa, which is 9 km (5.6 miles) thick and covers an area of 66,000 square km (about 25,500 square miles). It was intruded nearly 2.1 billion years ago and is the largest repository of magmatic ore deposits in the world. The Bushveld Complex consists of stratiform layers of dunite, norite (a type of gabbro rich in orthopyroxene), anorthosite, and ferrodiorite (an iron-rich intrusive igneous rock that is basic to intermediate in composition) and contains deposits of chromite, iron, titanium, vanadium, nickel, and—most important of all—platinum. The Sudbury Complex in southern Canada, which is about 1.9 billion years old, is a basin-shaped body that extends up to 60 km (37 miles) across. It consists mostly of layered norite and has deposits of copper, nickel, cobalt, gold, and platinum. It is noted for its high-pressure structures and other manifestations of shock metamorphism, which suggest that the intrusion was produced by an enormous meteorite impact.


Quartzites, dolomites, shales, and banded-iron formations make up sequences that reach up to 10 km (6.2 miles) in thickness and that amount to more than 60 percent of Proterozoic sediments. Minor sediments include sandstones, conglomerates, red beds, evaporites, and cherts. The quartzites typically have cross-bedding and ripple marks, which are indicative of tidal action, and the dolomites often contain stromatolites similar to those that grow today in intertidal waters. Also present in the dolomites are phosphorites that are similar to those deposited on shallow continental margins against areas of oceanic upwelling during the Phanerozoic. Several early-middle Proterozoic examples of such dolomites have been found in Finland and northern Australia, as well as in the Marquette Range of Michigan in the United States, in the Aravalli Range of Rajasthan in northwestern India, and at Hamersley and Broken Hill in Australia. Other constituents of these dolomites include evaporites that contain casts and relicts of halite, gypsum, and anhydrite. Examples occur at Mount Isa in Australia (1.6 billion years old) and in the Belcher Group in Canada (1.8 billion years old). These evaporites were deposited by brines in very shallow pools such as those encountered today in the Persian Gulf.


Phanerozoic ophiolites are considered to be fragments of ocean floor that have been trapped between island arcs and continental plates that collided or that have been thrust onto the shelf sediments of continental margins. They consist of a downward sequence of oceanic sediments such as cherts, pillow-bearing basalts, sheeted basic dikes, gabbros, and certain ultramafic rocks (such as serpentinized harzburgite, which is primarily made of olivine and orthopyroxene; and lherzolite, which is mainly composed of olivine, clinopyroxene, and orthopyroxene). Comparable ophiolites occur in several Proterozoic orogenic belts and provide strong evidence of the existence of oceanic plates similar to those of today. The oldest is an ophiolite in the Cape Smith belt on the south side of Hudson Bay in Canada whose age has been firmly established at 1.999 billion years. There is a 1.96-billion-year-old ophiolite in the Svecofennian belt of southern Finland, but most Proterozoic ophiolites are 1 billion to 570 million years old and occur in the Pan-African belts of Saudi Arabia, Egypt, Yemen, and The Sudan, where they occur in sutures between a variety of island arcs.


Greenstone-granite belts such as those of the Archean continued to form in the Proterozoic, albeit in greatly reduced amounts. They are characterized by abundant volcanic rocks that include pillowed subaqueous basalt flows and subaerial and subaqueous volcaniclastic rocks. Magnesian komatiites are for the most part absent, however. Intrusive plutons are typically made of granodiorite. Examples occur at Flin Flon in central Canada, in the Birrimian Group in West Africa, and in the Pan-African belts of the Arabian-Nubian Shield. Generally, such rocks resemble those in modern island arcs and back-arc basins, and the presence of remnants of oceanic plateau is suspected.


These highly deformed and metamorphosed rocks are similar to those of the Archean Eon and occur in many Proterozoic orogenic belts such as the Grenville in Canada, the Pan-African Mozambique belt in eastern Africa and Madagascar, the Musgrave and Arunta ranges in Australia, and in Lapland in the northern Baltic Shield. They were brought up from the mid-lower crust on major thrusts as a result of continental collisions.

One of the world’s classic Proterozoic orogenic belts is the Wopmay Orogen, which is situated in the Arctic in the northwestern part of the Canadian Shield. This beautifully exposed belt formed within a relatively short time (between 1.97 and 1.84 billion years ago) and provides convincing evidence of tectonic activity of a modern form in the early Proterozoic. On the eastern continental margin here are red beds (sandstones) that pass oceanward and westward into stromatolite-rich dolomites deposited on the continental shelf to a thickness of 4 km (2.5 miles); these dolomites pass into submarine turbidite fans that were deposited on the continental rise. An island arc and a continental margin are located to the west. The history of the Wopmay Orogen can be best interpreted in terms of subduction of oceanic crust and collision tectonics.

The Svecofennian Orogen of the Baltic Shield extends in a southeasterly direction from northern Sweden through southern Finland to the adjoining part of western Russia. It formed in the period from 1.9 to 1.7 billion years ago. A major lineament across southern Finland consists of the suture zone on which occur ophiolite complexes representing the remains of oceanic crust. At Outokumpu there is copper mineralization in these oceanic crust rocks similar to that in the Cretaceous ophiolite at Troodos in Cyprus. On the northern side of the suture is a shelf-type sequence of sediments; on the southern side is a volcanic-plutonic arc. To the south of this arc lies a broad zone with thrusted gneisses intruded by tin-bearing crustal-melt granites, called rapakivi granites after their coarse, zoned feldspar megacrysts (that is, crystals that are significantly larger than the surrounding fine-grained matrix). The rocks in this zone probably formed as a result of mantle plume activity.

The Grenville Orogen is a deeply eroded and highly uplifted orogenic belt that extends from Labrador in northeastern Canada to the Adirondack Mountains and southwestward under the coastal plain of the eastern United States. It developed from about 1.5 to 1 billion years ago. Apart from an island arc situated today in Ontario, most of the Grenville Orogen consists of highly metamorphosed and deformed gneisses and granulites that have been brought to the present surface on major thrusts from the mid-lower crust. A result of the terminal continental collision that occurred at about 1.1 billion years ago was the formation of the Midcontinent (or Keweenawan) rift system that extends southward for more than 2,000 km (about 1,240 miles) from Lake Superior.
A type of crustal growth—one very different from that described above—took place in what are now Saudi Arabia, Egypt, Yemen, and The Sudan in the period from 1.1 billion to 500 million years ago. This entire shield, called the Arabian-Nubian Shield, is dominated by volcanic lavas, tuffs (consolidated rocks consisting of pyroclastic fragments and ash), and granitic plutons that formed in a variety of island arcs separated by several sutures along which many ophiolite complexes occur. Some of the ophiolites contain a complete stratigraphy that is widely accepted as a section through the oceanic upper mantle and crust. The final collision of the arcs was associated with widespread thrusting and followed by the intrusion of granitic plutons containing tungsten, tin, uranium, and niobium ore deposits. The island arcs grew from the subduction of oceanic crust in a manner quite comparable to that taking place today throughout Indonesia.

The Mozambique belt is one of the many Pan-African orogenic belts that formed in the period between 1 billion and 500 million years ago. It extends along the eastern border of Africa from Ethiopia to Kenya and Tanzania. It consists largely of highly metamorphosed, mid-crustal gneisses deformed by eastward-dipping thrusts very similar to the thrusts on the southern side of the Himalayas (formed as a result of the collision of India with Tibet during the early Cenozoic Era). To the east on the island of Madagascar, mid-crustal gneisses of similar age were brought to the surface by major late extensional collapse of the orogenic belt.

During the middle and late Proterozoic, thick sequences of sediment were deposited in many basins throughout Asia. The Riphean sequence spans the period from 1.6 billion to 800 million years ago and occurs primarily in Russia. The Sinian sequence in China extends from 800 to 570 million years ago, toward the end of the Precambrian time. The sediments are terrigenous debris characterized by conglomerates, sandstone, siltstone, and shale, some of which are oxidized red beds, along with stromatolite-rich dolomite. Total thicknesses reach over 10 km (6.2 miles). The terrigenous sediments were derived from the erosion of Proterozoic orogenic belts.


Evidence of the oldest known glaciation, which occurred 2.9 billion years ago, is preserved in the Pongola Rift in South Africa, though most Precambrian glaciations occurred during the Proterozoic. Evidence that ancient deposits are of glacial origin is obtained by comparing them with those left behind by the Quaternary ice sheets and with deposits associated with modern glaciers. The main sediments left behind by early Proterozoic glaciers are tillites containing rock fragments ranging in size from pebbles to boulders and distributed randomly in a fine-grained silty matrix. The surfaces of some pebbles have parallel scratches caused by having been rubbed against harder pebbles during ice transport. Locally, the basement rocks below the tillite also have been scratched, or striated, by the movement of the overlying boulder-strewn ice. Another type of glacial deposit is a varved (laminated) sediment composed of alternating millimetre-to-centimetre-thick layers of silt and clay, which closely resemble the layered varves that are laid down in modern glacial lakes at the front of retreating glaciers or ice sheets. Each of these layers defines an annual accumulation of sediment. Varved sediments may contain dropstones, which are fragments of rock that have dropped from an overlying floating ice sheet and that have sunk into and depressed the layers beneath them. When all these features are found together, they provide good evidence of ancient glaciations.

The most extensive early Proterozoic Huronian glaciation occurred 2.3 billion years ago in what is now northern North America. Glacial deposits, similar in age to those of the Huronian, are located in the Transvaal and Cape regions of South Africa, where they reach only 30 metres (100 feet) in thickness but extend over an area of 20,000 square km (7,700 square miles). Such deposits are also encountered in the Hamersley Basin of Western Australia, in east-central Finland and the adjoining part of northwestern Russia, near Lake Baikal in Siberia, and in central India, suggesting the occurrence of a wide-spread glaciation.

Evidence for the largest glaciation in Earth’s history, known as the Snowball Earth event, dates from the late Proterozoic between 1 billion and 600 million years ago. The principal occurrences of these global glacial deposits are in Europe (Scotland, Ireland, Sweden, Norway, France, the Czech Republic, and Slovakia), the Western Cordillera (Yukon, Can., to California, U.S.) of western North America and the Appalachians of the United States, eastern Greenland, Brazil, much of Africa (Congo [Brazzaville], Angola, Namibia, Zambia, Congo [Kinshasa], and South Africa), and much of Russia, China, and Australia. In addition to the Flinders Range deposits described above (see Worldwide glaciations), other notable deposits include the Port Askaig tillite on the island of Islay off northwestern Scotland, which is only 750 metres (2,460 feet) thick but records 17 ice advances and retreats and 27 periglacial periods (which are indicated by infilled polygons that formed under ice-free permafrost conditions). There are two major tillites in central Africa and Namibia (910 to 870 and 720 to 700 million years old, respectively) and two other such consolidated tills in eastern Greenland.

Correlation of Precambrian strata

The fact that Phanerozoic sediments have been so successfully subdivided and correlated is attributable to the presence of abundant fossil remains of life-forms that evolved and underwent changes over time. Precambrian sediments lack such fossils, thus preventing any comparable correlations. There are, however, stromatolites in Precambrian sediments ranging in age from about 3.5 billion to 540 million years that reached their peak of development in the Proterozoic. Stromatolites underwent evolutionary changes sufficient for Russian biostratigraphers to use to subdivide the Riphean sequence into four main zones throughout widely separated areas of former Soviet territory. Similar stromatolite-based stratigraphic divisions have been recognized in the Norwegian islands of Spitsbergen, China, and Australia. This stromatolite biostratigraphy still has relatively limited application, however. As a consequence, it is the chronometric time scale that is used to subdivide Precambrian time and to correlate rocks from region to region and from continent to continent.

The rocks within Proterozoic orogenic belts are invariably too deformed to allow correlation of units between different belts. Nonetheless, the techniques of geochronology—in particular, zircon dating—have improved considerably in recent years, with the result that rocks of approximately similar age on different continents can be mutually compared and regarded as equivalent. The isotopic dating of Archean rocks, especially with the use of zircons, has enabled similarities and differences in age to be determined, thereby aiding correlation.

Establishing Precambrian boundaries

There is no record of tectonic activity of any sort at the time corresponding to the Archean-Proterozoic boundary—about 2.5 billion years ago. This probably means that a supercontinent had been created by the amalgamation of innumerable smaller continental blocks and island arcs. Accordingly, this was a period of tectonic stability that may have been comparable to the Permian-Triassic when the supercontinent of Pangea existed. The main geologic events would have been the intrusion of basic dikes and the formation of sedimentary basins such as the Huronian on the U.S.-Canadian border, into which large volumes of clastic sediment (that is, sediment of predominantly clay, silt, and sand sizes) were deposited. Such sediments would have been derived by erosion of high plateaus and mountains that are characteristic of a large continental mass.

Brian Frederick Windley
Hadean Eon


The Hadean eon is often characterized by extreme volcanism as
Earth continued to cool

Hadean Eon

Hadean Eon, informal division of Precambrian time occurring between about 4 billion and about 4.0 billion years ago. The Hadean Eon is characterized by Earth’s initial formation—from the accretion of dust and gases and the frequent collisions of larger planetesimals—and by the stabilization of its core and crust and the development of its atmosphere and oceans. Throughout part of the eon, impacts from extraterrestrial bodies released enormous amounts of heat that likely prevented much of the rock from solidifying at the surface. As such, the name of the interval is a reference to Hades, a Greek translation of the Hebrew word for hell.

Earth’s surface was incredibly unstable during the early part of the Hadean Eon. Convection currents in the mantle brought molten rock to the surface and caused cooling rock to descend into magmatic seas. Heavier elements, such as iron, descended to become the core, whereas lighter elements, such as silicon, rose and became incorporated into the growing crust. Although no one knows when the first outer crust of the planet formed, some scientists believe that the existence of a few grains of zircon dated to about 4.4 billion years ago confirm the presence of stable continents, liquid water, and surface temperatures that were probably less than 100 °C (212 °F). Since Hadean times, nearly all of this original crust has subducted from the movements of tectonic plates, and thus few rocks and minerals remain from the interval. The oldest rocks known are the faux amphibolite volcanic deposits of the Nuvvuagittuq greenstone belt in Quebec, Canada; they are estimated to be 4.28 billion years old. The oldest minerals are the aforementioned grains of zircon, which were found in the Jack Hills of Australia.
Considerable debate surrounds the timing of the formation of the atmosphere as well as its initial composition. Although many scientists contend that the atmosphere and the oceans formed during the latter part of the eon, the discovery of the zircon grains in Australia provide compelling evidence that the atmosphere and ocean formed before 4.4 billion years ago. The early atmosphere likely began as a region of escaping hydrogen and helium. It is generally thought that ammonia, methane, and neon were present sometime after the crust cooled, and volcanic outgassing added water vapour, nitrogen, and additional hydrogen. Some scientists state that ice delivered by comet impacts could have supplied the planet with additional water vapour. Later, it is thought, much of the water vapour in the atmosphere condensed to form clouds and rain that left large deposits of liquid water on Earth’s surface.
The Moon is also thought to have formed during the Hadean Eon, and several theories of the Moon’s origin have been posited. The leading theory asserts that a collision between Earth and a celestial body the size of Mars ejected material that eventually coalesced into the Moon.

John P. Rafferty

Archaean Eon

Archean Eon

Archean Eon, also spelled Archaean Eon, the earlier of the two divisions of Precambrian time (about 4 billion to 542 million years ago).

The Archean Eon began about 4 billion years ago with the formation of the Earth’s crust and extended to the start of the Proterozoic Eon 2.5 billion years ago; the latter is the second division of Precambrian time. Records of Earth’s primitive atmosphere and oceans emerge in the earliest Archean (Eoarchean Era), and evidence of the earliest primitive life-forms—bacteria and blue-green algae—appears in rocks about 3.5 billion years old. Archean greenstone-granite belts contain many economic mineral deposits, including gold and silver.

The start of the Archean Eon is only defined by the isotopic age of the earliest rocks. Prior to the Archean Eon, the Earth was in the astronomical (Hadean) stage of planetary accretion that began about 4.6 billion years ago; no rocks are preserved from this stage. The earliest terrestrial materials are not rocks but minerals. In Western Australia some sedimentary conglomerates, dated to 3.3 billion years ago, contain relict detrital zircon grains that have isotopic ages between 4.2 and 4.4 billion years. These grains must have been transported by rivers from a source area, the location of which has never been found; it was possibly destroyed by meteorite impacts—quite frequent on both the Earth and the Moon before 4 billion years ago.

It is thought that the oxygen content in today’s atmosphere must have slowly accumulated through time starting with an atmosphere that was anoxic during Archean times. Although volcanoes exhale much water vapour (H2O) and carbon dioxide (CO2), the amount of free oxygen (O2) emitted is very small. The inorganic breakdown (photodissociation) of volcanic-derived water vapour and carbon dioxide in the atmosphere would have produced only a small amount of free oxygen. The bulk of the free oxygen in the Archean atmosphere was derived from organic photosynthesis of carbon dioxide (CO2) and water (H2O) by anaerobic cyanobacteria (blue-green algae), a process that releases oxygen as a by-product. These organisms were prokaryotes, a group of unicellular organisms with rudimentary internal organization.

Archean oceans were likely created by the condensation of water derived from the outgassing of abundant volcanoes. Iron was released then (as today) into the oceans from submarine volcanoes in oceanic ridges and during the creation of thick oceanic plateaus. This ferrous iron (Fe2+) combined with oxygen and was precipitated as ferric iron in hematite (Fe2O3), which produced banded-iron formations on the flanks of the volcanoes. The transfer of biologically produced oxygen from the atmosphere to the sediments was beneficial to the photosynthetic organisms, because at the time free oxygen was toxic to them. When banded-iron formations were being deposited, oxygen-mediating enzymes had not yet developed. Therefore, this removal of oxygen allowed early anaerobes (life-forms not requiring oxygen for respiration) to develop in the early oceans of the Earth.
Carbon dioxide emissions are abundant from modern volcanoes, and it is assumed that the intense volcanism during the Archean Eon caused this gas to be highly concentrated in the atmosphere. This high concentration most likely gave rise to an atmospheric greenhouse effect that warmed the Earth’s surface sufficiently to prevent the development of glaciations, for which there is no evidence in Archean rocks. The CO2 content in the atmosphere has decreased over geological time, because much of the oxygen formerly bound in CO2 has been released to provide increasing amounts of O2 to the atmosphere.
  In contrast, carbon has been removed from the atmosphere via the burial of organic sediments.

Throughout the Archean, oceanic and island arc crust was produced semi-continuously for 1.5 billion years; thus, most Archean rocks are igneous. The oldest known rocks on Earth, estimated at 4.28 billion years old, are the faux amphibolite volcanic deposits of the Nuvvuagittuq greenstone belt in Quebec, Canada. The second oldest rocks are the 4-billion-year-old Acasta granitic gneisses in northwestern Canada, and a single relict zircon grain dated to 4.2 billion years ago was found within these gneisses. Other ancient sediments and lavas occur in the 3.85-billion-year-old Isua belt of western Greenland (which is similar to an accretionary wedge in the trench of a modern subduction zone) and the 3.5-billion-year-old Barberton Complex in South Africa, which is probably a slice of oceanic crust. A huge pulse in the formation of island arcs and oceanic plateaus took place worldwide from 2.9 to 2.7 billion years ago.

Archean rocks mostly occur in large blocks hundreds to thousands of kilometres across, such as in the Superior and Slave provinces in Canada; the Pilbara and Yilgarn blocks in Australia; the Kaapvaal craton in southern Africa; the Dharwar craton in India; the Baltic, Anabar, and Aldan shields in Russia; and the North China craton. Smaller relicts of Archean rocks in various stages of obliteration occur in many younger Proterozoic and Phanerozoic orogenic (mountain) belts. Some Archean rocks that occur in greenstone-granite belts (zones rich in volcanic rocks that are primitive types of oceanic crust and island arcs) formed on or near the surface of the Earth and thus preserve evidence of the early atmosphere, oceans, and life-forms. Other rocks that occur in granulite-gneiss belts (zones of rocks that were metamorphosed in the Archean mid-lower crust) are exhumed remnants of the lower parts of the Archean continents and thus preserve evidence of deep crustal processes operating at the time.

In greenstone-granite belts there are many oceanic lavas, island arcs, and oceanic plateaus; therefore, they commonly contain rock types such as basalts, andesites, rhyolites, granitic plutons, oceanic cherts, and ultramafic komatiites (lavas enriched in magnesium, a special product of the melting of the hot Archean mantle). These igneous rocks are host to multitudes of economic mineral deposits of gold, silver, chromium, nickel, copper, and zinc, which are the mainstay of the economies of Canada, Australia, and Zimbabwe.

In granulite-gneiss belts the roots of many Andean-type active continental margins are exposed, the rocks being highly deformed and recrystallized during metamorphism in the deep crust. Common rocks are tonalites (a granitic-type rock rich in plagioclase feldspar) transformed into tonalitic gneisses, amphibolite dikes, and amphibolites derived from volcanic activity. Few mineral deposits occur in granulite-gneiss belts, in common with the deep crust of younger orogenic belts, which are relatively barren of ore concentrations.

Brian Frederick Windley

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