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The Earth and its Lithosphere

The earth has been in a state of continual change since its formation. The major part of this change, involving volcanism and tectonics, has been driven by heat produced from the decay of radioactive elements within the earth. The other source of change has been solar energy, which acts as the driving force of weathering and is the ultimate source of energy for living organisms.

The solar system was probably formed about 4.6 billion years ago, and the oldest known rocks have an age of 3.8 billion years. There is thus a gap of 0.8 billion years for which there is no direct evidence. It is known that the earth was subjected to extensive bombardment earlier in its history; recent computer simulations suggest that the moon could have resulted from an especially massive collision with another body. Although these major collisions have diminished in magnitude as the matter in the solar system has become more consolidated, they continue to occur, with the most recent one being responsible for the annihilation of the dinosaurs and much of the other life on Earth. The lack of many overt signs of these collisions (such as craters, for example) testifies to the dynamic processes at work on the Earth’s surface and beneath it.

Chemical composition of the Earth

The earth is composed of 90 chemical elements, of which 81 have at least one stable isotope. The unstable elements are 43Tc and 61Pm, and all elements heavier than 83Bi.

Note that the vertical axis is logarithmic, which has the effect of greatly reducing
the visual impression of the differences between the various elements.

The chart gives the abundances of the elements present in the solar system, in the earth as a whole, and in the various geospheres. Of particular interest are the differences between the terrestrial and cosmic abundances, which are especially notable in the cases of the lighter elements (H, C, N) and the noble gas elements (He, Ne, Ar, Xe, Kr).

Given the mix of elements that are present in the earth, how might they combine so as to produce the chemical composition we now observe? Thermodynamics allows us to predict the composition that any isolated system will eventually reach at a given temperature and pressure. Of course the earth is not an isolated system, although most parts of it can be considered approximately so in many respects, on time scales sufficient to make thermodynamic predictions reasonably meaningful. The equilibrium states predicted by thermodynamics differ markedly from the observed compositions. The atmosphere, for example, contains 0.03% CO2 , 78% N2 and 21% O2 ; in a world at equilibrium the air would be 99% CO2.

Similarly, the oceans, containing about 3.5% NaCl, would have a salt content of 35% if they were in equilibrium with the atmosphere and the lithosphere. Trying to understand the mechanisms that maintain these non-equilibrium states is an important part of contemporary environmental geochemistry.

Structure of the Earth

Studies based on the reflection and refraction of the acoustic waves resulting from earthquakes show that the interior of the earth consists of four distinct regions. A combination of physical and chemical processes led to the differentiation of the earth into these major parts. This is believed to have occurred approximately 4 billion years ago.

The Earth's Core

The Earth’s core is believed to consist of two regions. The inner core is solid, while the outer core is liquid. This phase difference probably reflects a difference in pressure and composition, rather than one of temperature. Density estimates obtained from seismological studies indicate that the core is metallic, and mainly iron, with 8-10 percent of lighter elements.

Hypotheses about the nature of the core must be consistent with the the core’s role as the source of the earth’s magnetic field. This field arises from convective motion of the electrically conductive liquid comprising the outer core. Whether this convection is driven by differences in temperature or composition is not certain. The estimated abundance of radioactive isotopes (mainly U238 and K40 in the core is sufficient to provide the thermal energy required to drive the convective dynamo. Laboratory experiments on the high-pressure behavior of iron oxides and sulfides indicate that these substances are probably metallic in nature, and hence conductive, at the temperatures (4000-5000K) and pressures (1.3-3.5 million atm) that are estimated for the core. Their presence in the core, alloyed with the iron, would be consistent with the observed density, and would also resolve the apparent lack of sulfur in the earth, compared to its primordial abundance.


The mantle

The region extending from the outer part of the core to the crust of the earth is known as the mantle. The mantle is composed of oxides and silicates, i.e., of rock. It was once believed that this rock was molten, and served as a source of volcanic magma. It is now known on the basis of seismological evidence that the mantle is not in the liquid state. Laboratory experiments have shown, however, that when rock is subjected to the high temperatures and pressures believed to exist in the mantle, it can be deformed and flows very much like a liquid.

The upper part of the mantle consists of a region of convective cells whose motion is driven by the heat due to decay of radioactive potassium, thorium, and uranium, which were selectively incorporated in the crystal lattices of the lower-density minerals that form the mantle. There are several independent sources of evidence of this motion. First, there are gravitational anomalies; the force of gravity, measured by changes in elevation in the sea surface, is different over upward and downward moving regions, and has permitted the mapping of some of the convective cells. Secondly, numerous isotopic ratio studies have traced the exchange of material between oceanic sediments, upper mantle rock, and back into the continental crust, which forms from melting of the upper mantle. Thirdly, the composition of the basalt formed by upper mantle melting is quite uniform everywhere, suggesting complete mixing of diverse materials incorporated into the mantle over periods of 100 million years.

High-pressure studies in the laboratory have revealed that olivine, a highly abundant substance in the mantle composed of Fe, Mg, Si, and O (and also the principal constituent of meteorites) can undergo a reversible phase change between two forms differing in density. Estimates of conditions within the upper mantle suggest that the this phase change could occur within this region in such as way as to contribute to convection. The most apparent effect of mantle convection is the motion it imparts to the earth’s crust, as evidenced by the the external topography of the earth.

The crust

The outermost part of the earth, known also as the lithosphere, is broken up into plates that are supported by the underlying mantle, and are moved by the convective cells within the mantle at a rate of a few centimetres per year. New crust is formed where plates move away from each other under the oceans, and old crust is recycled back into the mantle as where plates moving in opposite directions collide.

The oceanic crust

The parts of the crust that contain the world’s oceans are very different from the parts that form the continents. The continental crust is 10-70 km thick, while oceanic crust averages only 5-7 km in thickness. Oceanic crust is more dense (3.0-3.1 g cm–3) and therefore “floats” on the mantle at a greater depth than does continental crust (density 2.7-2.8 ). Finally, oceanic crust is much younger; the oldest oceanic crust is about 200 million years old, while the most ancient continental rocks were formed 3.8 billion years ago.

  New crust is formed from molten material in the upper mantle at the divergent boundaries that exist at undersea ridges. The melting is due to the rise in temperature associated with the nearly adiabatic decompression of the upper 50-70 km of mantle material as separation of the plates reduces the pressure below. The molten material collects in a magma pocket which is gradually exuded in undersea lava flows. The solidified lava is transformed into crust by the effects of heat and the action of seawater which selectively dissolves the more soluble components.    

An animated view of seafloor spreading.


Plate collisions


Where two plates collide, one generally plunges under the other and returns to the mantle in a process known as subduction. Since the continental plates have a lower density, they tend to float above the oceanic plates and resist subduction. At continental boundaries such as that of the North American west coast where an oceanic plate pushes under the continental crust, oceanic sediments may be sheared off, resulting in a low coastal mountain range (see here for a nice animation of this process.) Also, the injection of water into the subducting material lowers its melting point, resulting in the formation of shallow magma pockets and volcanic activity. Divergent plate boundaries can cross continents, however; temporary divergences create rift valleys such as the Rhine and Rio Grande, while permanent ones eventually lead to new oceanic basins.

Collision of two continental plates can also occur; the most notable example is the one resulting in the formation of the Himalayan mountain chain.

The Earth is composed of 90 chemical elements, of which 81 have at least one stable isotope. Most of these elements have also been detected in stars. Where did these elements come from? The accepted scenario is that the first major element to condense out of the primordial soup was helium , which still comprises about one-quarter of the mass of the known universe.


Hydrogen is the least thermodynamically stable of the elements, and at very high temperatures will combine with itself in a reaction known as nuclear fusion to form the next element, 2He4. "Heavier" nuclei (that is , those having high atomic numbers, indicated here by the subscript preceding the element symbol), are more stable than "lighter" ones, so this fusion process can continue up to 56Fe, which is the most energetically stable of all the nuclides. Beyond this point, heavier nuclei slowly become less stable, so fission becomes more likely. FIssion, however, is not considered an important mechanism of primordialnucleosynthesis, so other processes are invoked, as discussed farther below.



Primordial Chemistry

According to the “big bang” theory for which there is now overwhelming evidence, the universe as we know it (that is, all space, time, and matter) had its origin in a point source or singularity  that began an explosive expansion about 12-15 billion years ago, and which is still continuing.

Following a brief period of extremely rapid expansion called inflation, protons and neutrons condensed out of the initial quantum soup after about 10–32 s. Helium and hydrogen became stable during the first few minutes, along with some of the very lightest nuclides up to 7Li, which were formed through various fusion and neutron-absorption processes. Formation of most heavier elements was delayed for about 106 years until nucleosynthesis commenced in the first stars. Hydrogen still accounts for about 93% of the atoms in the universe.

The main lines of observational evidence that support this theory are the 2.7K background radiation that permeates the cosmos (the cooled-down remnants of the initial explosion), and the abundances of the lightest elements. Conventional physics is able to extrapolate back to about the first 10–33 second; what happened before then remains speculative.

Stellar nucleosynthesis

All elements beyond hydrogen were formed in regions where the concentration of matter was large, and the temperature was high; in other words, in stars. The formation of a star begins when the gravitational forces due to a large local concentration of hydrogen bring about a contraction and compression to densities of around 105 g cm–3. This is a highly exothermic process in which the gravitational potential energy is released as heat, about 1200 kJ per gram, raising the temperature to about 107 K. Under these conditions, hydrogen nuclei possess sufficient kinetic energy to overcome their electrostatic repulsion and undergo nuclear fusion:

1H1-> 2He4 + 2 b+ + 2 g + 2 n

Hydrogen burning

There will be a net mass loss in above process, which will therefore be highly exothermic and is known as “hydrogen burning”. As hydrogen burning proceeds, the helium collects in the core of the star, raising the density to 108 g cm–3 and the temperature to 108 K. This temperature is high enough to initiate helium burning, which proceeds in several steps:

2He4 -> 4Be8 + g

The first product, 4Be8 has a half life of only 10–16 sec, but a sufficient amount accumulates to drive the following two reactions:

4Be8 + 2He4 -> 6C12 + g

6C12 + 1H1 -> 7N13 -> 6C13 + b+ + g

The size of a star depends on the balance between the kinetic energy of its matter and the gravitational attraction of its mass. As the helium burning runs its course, the temperature drops and the star begins to contract. The course of further nucleosynthesis and the subsequent fate of the star itself depends on the star’s mass.

Small stars

If the mass of the star is no greater than 1.4 times the mass of our sun, the star collapses to a white dwarf, and eventually cools to a dark, dense dead star.

Big stars

In larger stars, the gravitational forces are sufficiently strong to overcome the normal repulsion between atoms, and so gravitational collapse continues. The gravitational energy released in this process produces temperatures of 6  10K, which are sufficient to initiate a complex series of nuclear reactions known as the carbon-nitrogen cycle. The net reaction of this cycle is the further fusion of hydrogen to helium, in which C12 acts as a catalyst, and various nuclides of nitrogen and oxygen are intermediates. The temperature is sufficiently high, however, to initiate fusion reactions of some of these intermediates:

6C12 + 6C12 -> 10Ne20 + 2He4

8O16 -> 14Si28 + 2He4

8O16 -> 16S31 + 0n1


The intense gamma radiation that is produced in some of these reactions breaks some of the product nuclei into smaller fragments, which can then fuse into a variety of heavier species, up to the limit of26Fe56, beyond which fusion is no longer exothermic. The greater relative abundance of elements such as 6C128O16, and 10Ne20 which differ by a 2He4 nucleus, reflects the participation of the latter species in these processes. These exothermic reactions eventually produce temperatures of 8  109 K, while contraction continues until the central core is essentially a ball of neutrons having a radius of about 10 km and a density of 1014 g cm–3. At the same time the outer shell of the star is blasted away in an explosion known as a supernova.

Since 26Fe56 has the highest binding energy per nucleon of any nuclide, there are no exothermic processes which can lead to the formation of heavier elements. Fusion into heavier species is also precluded by the electrostatic repulsion of the highly charged nuclei. However, the process of neutron capture can still take place (this is the same process that is used to make synthetic elements). The neutrons are by-products of a large variety of stellar processes, and are present in a wide range of energies. Two general types of neutron capture processes are recognized. In an “s” (slow) process, only a single neutron is absorbed and the product usually decomposes by b-decay into a more proton-rich species.

Elements heavier than iron

26Fe56 + 0n1 -> 26Fe56 -> 26Fe59 -> 27Co59 + –1e0

This process occurs at rates of about 105 yr–1 , and accounts for the lighter isotopes of many elements. The other process (the “r”, or rapid process) occurs in regions of high neutron density and involves multiple captures at rates of 0.1-10 sec–1:

26Fe56 + 13 0n1 -> 26Fe56 -> 27Co59 + –1e0

This mechanism favors the heavier, neutron-rich isotopes and the heaviest elements.

Other elements

A few nuclei are not accounted for by any of the processes mentioned. These are all low-abundance species, and they probably result from processes having low rates. Examples are Sn112 and Sn114which may be produced through proton-capture, and H2, Li6, Li7, Be, B10 and B11, which may come from spallation processes resulting from collisions of cosmic ray particles with heavier elements

Formation of the solar system

The solar system is believed to have formed about 5 billion years ago as a result of aggregation of cosmic dust and interstellar atoms in a region of space in which the density of such material happened to be greater than average. Over 99.8% of this mass, which consisted mostly of hydrogen, collapsed into a proto-sun; the gravitational energy released in this process raised the temperature sufficiently to initiate the hydrogen fusion reactions discussed above.

The planets

The remaining material probably formed a disk that rotated around the sun. As the temperature dropped to around 2000K, some of the most stable combinations of the elements began to condense out. These substances might have been calcium aluminum silicates, followed by the more volatile iron-nickel system, and then magnesium silicates. The further aggregation of these materials, together with the other constituents of the cooling disk, is now believed to be the origin of the planets. Density estimates indicate that the planets closest to the sun are predominantly rocky in nature, and probably condensed first. The outer planets (Uranus, Neptune and Pluto) appear to consist largely of water ice, methane, and ammonia, with a smaller rocky core.

Formation of the Earth

The Earth formed by accretion of solid and particulate material that remained after the much more massive amounts of hydrogen and helium present in the original protoplanets had been dispersed out of the solar system. Gradually, the heat produced by decay of radioactive elements brought about partial melting of the silicate rocks; these lower density molten materials migrated upward, leaving the more dense, iron-containing minerals below. This process, which took about 2 million years, was the first of the three stages into which the chemical evolution of the earth is usually divided:

  1. Primary differentiation of the elements between the core and mantle.
  2. Secondary differentiation of the elements, reflecting relative ionic sizes, bonding properties, and solubilities (influencing phase behavior such as fractional crystallization, etc.)
  3. Tertiary differentiation, still operative, involving the interaction of the crust with the hydrosphere and atmosphere.

The above listing should not be taken too literally; all three kinds of processes have probably proceeded simultaneously, and over a number of cycles. Since the earth is losing approximately four times as much heat as is generated by radioactive decay, the principal driving force of primary and secondary differentiation has gradually slowed down. Partial melting of the upper mantle brought about further fractionation as silicon-containing materials of low density migrated outward to form a crust. In its early stages the stronger granitic rocks had not yet appeared, and the crust was mechanically weak. Upwelling flows of lava would break the surface, and the weight of the solidified lava would cause the crust to subside. In some places, magma would solidify underground, forming low-density rock (batholiths) that would eventually rise by buoyancy and push up overlying crust. These mountain-building periods probably occurred in 6-8 major episodes, each lasting about 800 million years.


At the same time, outgassing of solids released large amounts of HCl, CO, CO2, H2S, CH4 , SO2 , and SO3 into the primitive atmosphere. Large amounts of water were present in the primeval rocks in the form of hydrates, which were broken down as the result of the heating. Eventually when the outer crust cooled enough to permit condensation of the water vapor as rain, a new stage of chemical evolution began. The rain was initially highly acidic, equivalent to about 1M HCl; this reacted readily with the basic rocks having high contents of K, Na, Mg, and Ca, leaching them away and forming what would eventually evolve into the oceans. The partial dissolution of the rocks also resulted in large amounts of sediments, which played their own role in the transformation of the earth’s surface.

The continents

Within the crust, the lighter materials, being in isostatic equilibrium with the upper mantle, floated higher, and gradually became the nuclei of continents, which grew by accumulating similar material around their boundaries. This picture of continental development is supported by isotopic ratio studies which indicate that the nucleus of the North American continent, the Canadian Shield, is over 2.5 billion years old, while the peripheral parts are less than 0.6 billion years of age.

Primary differentiation of the elements

The more traditional geochemical view of primary differentiation begins with the assumption that the core of the earth is in a chemically reduced state, while the metallic elements constituting the mantle are almost entirely oxidized to their lower free energy cationic forms. Oxygen and sulfur acted as the major electron acceptors in this process, but the abundance of these elements was insufficient to oxidize much of the nickel or iron.

Iron as a reductant

Iron itself is believed to have played a crucial role in the primary differentiation of other metals and of oxidized metallic elements that iron is able to reduce. As the dense molten iron migrated in toward the core, it dissolved (formed a liquid alloy with) any other metals with which it came in contact, and it reduced (donated electrons to) those metallic cations that are less “active” metals than iron under these conditions. The resulting metal would then mix with more of the migrating liquid iron, and be carried along with it into the core.

Redox power of the elements

Accordingly, elements whose reduction potentials are more positive than iron (i.e., are lower-free energy electron sinks) are called siderophiles; these elements have a low abundance in the crust and upper mantle. The other two important classes of solid-forming elements are lithophile and chalcophile (see below.) These generally have more negative reduction potentials than iron, and are distinguished mainly by their relative affinity for oxygen or sulfur. The chalcophiles, of which Cu, Cd, and Sb are examples, tend to form larger, more polarizable ions which can associate with the sulfide ion. The lithophiles comprise those elements such as K, Al, Mn, and Si, which have smaller ions and which combine preferentially with oxygen. This broad classification is reflected in the dominant forms in which many of these elements occur in nature.

Secondary differentiation of the elements

The differential distribution of the elements within one of the main regions of the earth has been studied in detail only in that portion that is accessible, namely the upper crust. It is clear that fractional crystallization from the cooling magma has played an important role. The relative temperatures at which minerals crystallize is determined in large part by their lattice energies, which are in turn related to ionic sizes and charges. Minerals with small, highly charged ions will have higher melting points and should crystallize first. Thus the sodium-containing feldspar albite (NaAlSi2O8) is found nearer the surface than is its calcium analog anorthite (CaAlSi2O8). The less abundant elements often do not form minerals of their own, but may replace the ion of a more abundant mineral in its crystal lattice. This is known as isomorphous replacement, and it naturally depends on the relative ionic radii. Some ion pairs that undergo isomorphous replacement in minerals are K+ and Ba2+, Si4+ and Ge4+.

Phase behavior

The Phase Rule can be invoked to explain in a very rough way the differentiation of the elements into distinct solid phases.

P = C + 2 – F

Taking the degrees of freedom as 2 (fixed temperature and pressure), the six major elemental components (O, Si, Al, Fe, Mg and Na) can form up to six phases. Actually, more than 99% of igneous rocks comprise seven principal mineral phases. These are: the silica minerals, feldspars, feldspathoids, olivine, pyroxenes, amphiboles and micas.

The differential deposition of minerals is also influenced by the temperature-composition phase relations as exemplified by the ordinary two-component phase diagram. If the mineral that is rich in one component and which first crystallizes out is also more dense, then the richer ore will occur near the bottom of the deposit, while a more mixed ore (approaching the eutectic) will remain near the top.

Geochemical classification of the elements

Whether an element is concentrated in the crust or elsewhere depends on its chemical behavior and on the physical properties of its stable compounds. Geochemists have found it convenient to establish the following general classifications:

  • lithophiles ("rock-loving") elements are those such as Fe, Al, and Si which tend to occur as oxides (and to a lesser extent as chlorides and carbonates.) Elements in this, the largest of all the groups, are concentrated in the crust.
  • chalcophiles also occur in the crust, but mainly in combination with sulfur and the other chalcogen elements of Groups 15-16.
  • siderophiles refer to the elements such as Ni which have concentrated in the core along with Fe.
  • atmosphiles consist of N, H and their volatile compounds
    and the noble gases which concentrate in the atmosphere.

The structure and composition of the outer part of the lithosphere has been profoundly affected by interactions with the atmosphere over one-quarter of the surface area of the earth, and with the hydrosphere over the remaining area. Further modification of the outermost parts of the crust has occurred as the result of the activities of living organisms. These changes have transformed much of the outermost parts of the crust into an unconsolidated surface region called the regolith. Further weathering and translocation of soluble substances often results in a sequence of horizons consisting of sediments, soils, or evaporites.

Chemically, the earth’s crust consists of about 80 elements distributed in approximately 2000 compounds or minerals, many of which are of variable composition. Over 99% of the mass of the crustal material is made up of only eight of these elements, however:

O 466,000
Si 277,200
Al 81,300
Fe 50,000
Ca 36300
Na 28,300
K 25,900
Mg 20,900
Ti 4,400
H 1,400
Table 1: Average amounts of elements in crustal rocks,mg/g.  


The crust has its origin in the upwelling convection currents that bring mantle material near to the surface at the mid-ocean ridges. The reduced pressure causes it to melt into magma. The magma may solidify before it reaches the surface, forming basalt, or it may energe from the surface in a volcanic eruption. The oceanic crust consists mostly of the simpler silicate minerals, which are said to be basicor mafic. The more evolved, silicon-rich rocks found in the continental crust are known as acidic or sialic.

Oceanic crust is continually being extruded from regions of the plastic mantle that intrude upward to just beneath the ocean’s floor at the mid-ocean ridges. A corresponding amount of this crust is being returned to the lithosphere at subduction zones off the West coasts of the Americas, and in the process pushing up the mountain ranges that lie along these coasts. The subducted oceanic crust is reheated and combined with sedimentary material to undergo partial remelting and reworking; this is believed by some to be the origin of granite. Subduction proceeds at a rate of a few cm per year, and the complete cycle time is on the order of a few hundred million years.

The continental and oceanic crusts. This schematic cross section runs in a west-to-east direction, from just of the West coast of South America on the left to the African continent on the right. Oceanic crust is shown in black, with with continental crust in red.

Both oceanic and contintental crusts float on the more dense upper lithosphere, and gradually shift their positions as they push against each other, and in response to the slow convective motions in the medium that supports them. The continental crust is thicker than the oceanic crust, but it is also less dense, which allows it to float higher (and thus to differentiate continents from oceans.) The lower density also prevents it from being subducted. Recycling can occur indirectly as continental material erodes and is deposited as sediments on the ocean floor, but this is a much slower process and one that takes billions instead of millions of years. Some of the very oldest rocks, found in Greenland and Labrador, have been dated at 3.9 billion years, and thus approach the age of the Earth itself.

Chemistry of the crust

When magma crystallizes it forms igneous rock, the major component of the Earth’s crust. The crystallization is a complex process which is not entirely understood, due largely to the lack of sufficient thermodynamic data on the various components at high temperatures and pressures. It is known that the different components of magma have differing melting points and densities, and that the phase behavior of multicomponent systems based on some of these substances is quite complex, involving binary and ternary eutectics, solid solutions, the presence of dissolved water (under pressure), and incongruent melting. One consequence of this complexity is that the composition of the magma will change as crystallization takes place; different substances will crystallize at various stages, and the resulting solids may migrate toward the top or bottom of the region if their densities differ greatly from that of the magma

It is well known that larger crystals form when a melt cools more slowly. This principle affords a simple distinction between the coarser-grained plutonic rocks, which are believed to have been formed by gradual cooling of magma pockets within the crust, and the fine-grained volcanic rocks such as basalt. Under the influence of heat and pressure, particularly at plate boundaries, solid crustal material may undergo partial or complete remelting, followed by cooling and transformation into metamorphic rocks such as gneiss, micas, quartzite, and possibly granites.

Granite was once thought to be an igneous rock, originating from the crystallization of a particular kind of magma. The association of granitic rocks with mountainous regions, and the similarity of their compositions in widely scattered regions, lends credence to the more recent hypothesis that granitic rocks are of metamorphic origin.

Another class of rock is sedimentary rock, formed from the consolidation of material produced by weathering and other chemical, and biological processes. Sedimentary rocks cover about three-quarters of the land area of the earth; 80% are shales, 15% sandstones and 5% limestones.

Composition of rock

The chemical composition of rocks tends to be complex and variable, and can only be specified in a precise way at the structural level. The traditional way of expressing rock compositions is in terms of the mass percent of the oxides of the elements present in the rock.



common name

fresh rock

weathered rock

SiO2 silica 71.54 70.30
Al2O3 alumina 14.62 18.34
Fe2O3 ferric oxide 0.69 1.55
FeO ferrous oxide 1.64 0.22
MgO magnesia 0.77 0.21
CaO lime 2.08 0.10
Na2O soda 3.84 0.09
K2O potash 0.32 5.88
H2O water 0.32 5.88
others   0.65 0.54
total   100.07 99.70

Table 2: Chemical composition of a typical rock
(quartz-feldspar-biotite gneiss)

The figures are in percent by weight.


This does not mean that these oxidess, or the structural units they represent, are actually present as such in a rock. In the chemical analysis of rocks, oxygen is generally not determined separately. When it is, however, it is found in an amount that would be expected to combine stoichiometrically with the other elements present. Thus the composition of albite can be written as either NaAlSi3O8 or Na2O·Al2O3·6SiO2 . Some rocks contain varying ratios of certain elements. For example olivine, which can be considered a solid solution of Mg2SiO4 and Fe2SiO4, can be represented by (Mg,Fe)2SiO4; this implies that the ratio of metal to silica is constant, and that magnesium is ordinarily present in greater amount than iron.

The major structural elements of rock (both in the crust and in the mantle) are the silicate minerals, built from silicon atoms surrounded tetrahedrally by four oxygens. The simplest of these consist just of SiO44– tetrahedra interspersed with positive ions to achieve electroneutrality; olivine, (Mg,Fe)2SiO4 is a well known example. More commonly, the silicate groups polymerize by sharing one or more oxygen atoms at adjacent tetrahedral corners. Depending on the number of joined corners per silicate unit, this can lead to the formation of a wide variety of chains (pyroxenes, amphiboles) and sheets (micas), culminating in the complete tetrahedral polymerization that produces quartz, SiO2.

Higher degrees of polymerization are associated with higher ratios of Si to O, smaller quantities of positive ions, and higher melting points. Thus when magma cools, the first silicates to crystallize are the olivines, followed by chain and sheet minerals having progressively higher degrees of polymerization and smaller fractions of cations of metals such as Fe and Mg.

Distribution of elements; ores

Although some elements are distributed fairly uniformly throughout the crust, others occur at greatly enhanced concentrations in localized areas. There are two general processes that result in these localized excesses, which are called ores when their extraction and refining is economically feasible

The first of these relates to how well a metallic ion can fit into the silicate lattice structure. Ions having the right charge and size can readily enter this structure, displacing the more common ions of Fe, Al and Mg. Such ions (of which Ga3+ is an example) are readily soluble in other minerals and thus are widely disributed and do not concentrate into ores. Other ions may be too large (Cs and La), too small (Li, Be, B) or too highly charged (Nb, Ta, W) to be accommodated in silicate mineral structures; these elements tend to remain in the magma as it solidifies, finally forming solid minerals only in the last stages of cooling.

The other major source of ores is hydrothermal formation. Magma contains some water itself, and additional water from the surface is able to reach the heated rock near magma chambers. At the very high temperatures and pressures that prevail in these regions, the water can dissolve many compounds such as sulfides which are normally considered highly insoluble. When these superheated solutions rise to the surface the solids are re-deposited, often in highly concentrated form. Ores of Cu, Sn, W, and possibly some iron ores, as well as some native metals such as gold, are believed to be formed in this way.

Hydrothermal vents known as “black smokers” have been observed at sites of sea-floor spreading; the “smoke” consists of metallic sulfides which precipitate in the cold seawater. The veins of pyrites (FeS2) and similar sulfide minerals that are often observed in rock formations are the result of hydrothermal solutions that once penetrated cracks and fissures in the rock.

The hydrothermal vent photographed here emits both black and white "smoke".



Chemical weathering

The weathering of rocks at the earth’s surface is a complex process involving both physical and chemical changes. The latter tend in principle to be rather simple kinds of reactions involving dissolution, reaction with carbon dioxide, hydrolysis, hydration, and oxidation. The difficulty in studying them and in arriving at a quantitative description is that these reactions occur very slowly and may never reach an equilibrium state. A comparison of the two rightmost columns in Table 2 on page 14 provides some illustration of the overall effect of these changes, although it must be emphasized that these are relative composition data, and thus cannot show how much of a given component has been lost. In general, sodium, calcium and magnesium seem to be lost more rapidly than potassium and silicon, while iron and aluminum decrease very slowly. Individual rates are of course dependent on the particular structural units containing the element, and also vary somewhat with grain size and condition of the surface.

Action of water

Water is undoubtedly the most important weathering agent. Not only does it act as a solvent for ionic dissolution products, but it also brings other active agents such as carbon dioxide and oxygen into intimate contact with the rock material. As water percolates into the outermost layers of the crust, it extends the zone of weathering beneath the surface; the effects of this are quite noticeable in a number of buried sedimentary materials such as Paleozoic sandstones, which tend to be depleted of all but the most resistant minerals. Dissolution, the simplest of all the weathering processes, usually results in ionic species, some of which may react with water to yield acidic or alkaline solutions. Dissolution of silica, however, results in the neutral species H4SiO4. Reactions involving hydration and dehydration are very common, and since the free energy changes tend to be small, these reactions can usually take place in either direction under slightly different conditions. Thus gypsum and anhydrite are interconvertable at observable rates under common environmental conditions:

CaSO4·2H2O -> CaSO4 + 2 H2O

In many cases, however, the reaction products are not very well characterized, thermodynamic data is lacking, and the reactions proceed so slowly that they are not entirely understood. For example, both hydrous and anhydrous iron oxides can be found in similar geologic environments, but the little is known about the interconversion process, represented approximately as

Fe2O3 + H2O -> 2 FeOOH

Solid carbonates tend to dissolve in acidic solutions, including those produced when atmospheric carbon dioxide dissolves in water. Thus the major surface limestone deposits (largely CaCO3, with some admixture of MgCO3) tend to be highly eroded in non-arid regions, and the local groundwater may have Ca2+ as high as 0.1-0.2M. Thermodynamics can unambiguously predict the most stable oxidation state of a metal ion under given conditions of pH and oxidant concentration. The mechanisms tend to be very uncertain, however. For one thing, both the reactant and product can often exist in various states of hydration, and the dissolved species (which probably undergo the actual oxidation) often consist of polycations and complexed species.

Oxidation of iron

Compounds of Fe(II), for example, will always tend to oxidize to Fe(III) in the presence of air; the various oxides of iron are responsible for the bright colors seen in many geological formations, and in certain soils. Some of the net reactions that probably occur are

Fe2SiO4 + 1/2 O2 + H2O -> Fe2O3 + H2SiO4

2 FeCO3 + 1/2 O2 -> Fe2O3 + 2 H2CO3

An environmental side effect of the first process is the release of hydrated silica. Also, where both starting materials are present, the carbonic acid produced in the second reaction is believed to promote the dissolution of ferrous silicate, creating a source of Fe(II) ions that can be rapidly oxidized:

Fe2SiO4 + 4 H2CO3 -> 2 Fe2+ + 4 HCO3 + H4SiO4

Fe2+ + 4 HCO3 + 1/2 O2 -> Fe2O3 + 4 H2CO3

The oxidation of sulfides can produce strongly acidic solutions:

2 FeS2 + 15/2 O2 + 4 H2O -> Fe2O3 + 2 SO42– + 8 H+

The effects of this can be seen in formations containing outcrops of pyrite veins, where the surrounding rocks are heavily stained with yellow and brown Fe(III) oxides, and the groundwater tends to be highly acidic. This process is mediated by microorganisms, and is an important source of acid pollution associated with mines and mine tailings.

Sequence of weathering

The various components of rocks weather at different rates. The more basic components such as CaO and MgO tend to disappear first, especially if in contact with groundwaters containing high CO2 concentrations. For rocks in general, the first reaction is usually hydration, followed by hydrolysis which can be summarized by

4 KAlSi3O8(s) + 22 H2O -> Al4Si4O10(OH)8(s) + 8 H4SiO4(aq) + 4 K+ + 4 OH

in which other Group 1 or 2 cations might replace potassium. The product Al4Si4O10(OH)8 is kaolinite, a form of clay (see below).

In general, the rocks which crystallized first from the magma (the Ca-feldspars and olivines) weather more rapidly than do the lower-melting rocks.




Clays are the solid end products of the weathering of rocks. They are basically composed of alternating sheets of “SiO2” and “AlO6” units in ratios of 1:1 (kaolinite), 2:1 (montmorillonite and vermiculite) and 2:2 (chlorite). In between the sheets, and holding them together by hydrogen bonding are water molecules. Also present are cations such as K+, Ca2+ and Mg2+ which act to neutralize the negative charges of the oxide ions.

Structure of a clay, showing two layers of the stacked sheets of kaolinite.

Physical weathering

The major agents of physical weathering of exposed rocks are rapid changes in temperature (promoting fracture by differential expansion), the abrasive action of windborne material and glacier movement, and especially by the penetration of water into cracks and its subsequent freezing. The expansion of water on freezing can exert a pressure of 150 kg cm–2, whereas the tensile strength of a typical rock is around 120 kg cm–2 . The roots of some plants are able to penetrate rock quite effectively, producing comparable expansive pressures in subsurface rocks.

Composition and structure of soils



Soils are a product of the interaction of water, air, and living organisms with exposed rocks or sediments at the earth’s surface. A typical soil contains about 45% inorganic solids and 5% organic solids by volume. Water and air each make up about 20-30%.

A simple way of classifying soils is based on the relative quantities of clay, silt and sand in the solid component. For ordinary agricultural purposes, loams are considered the most ideal soil type.




Mineral Components of soil

The primary inorganic components of soils consist of sand and silt particles that come directly from the parent rocks. This fraction is dominated by quartz and feldspars (aluminosilicates). Secondary components are formed by chemical changes within the soil itself, or in sediments from which the soil derived. These are most commonly clays, but may also include calcite, gypsum, and sulfide minerals such as pyrites; the latter are formed by bacterial action under reducing conditions in the presence of organic matter.

The clays have an especially important effect on both the physical properties of the soil, and on its ability to store plant nutrients, including trace nutrients such as Mo and Mn. These properties are due to the high ion-exchange capacity of clays. The more highly charged cations such as Al3+ and Fe3+ tend to be more strongly absorbed within the inter-sheet regions than do Mg2+ or K+. As plants withdraw these latter cations from the soil water, more are released by the clay components, which thus act as nutrient reservoirs.

The ion-exchange properties of clays also help to maintain the pH balance of soils, through the exchange of H+ and cations such as Ca2+. The soil pH, in turn, strongly affects the solubility of nutrient cations, and thus their availability to plants. For example, the uptake of phosphorus (in the form of H2PO42– , is only efficient within the rather narrow pH range between 6 and 7. Below 6, dihydrogen phosphates of Fe and Al are precipitated, while insoluble Ca3(PO4)2 forms at higher pH’s.

Organic Components

Part of the organic matter of soil consists of organisms (mainly bacteria and fungi) and roots and root hairs. The remainder is largely in the form of fulvic and humic acids. These substances of indefinite composition are classified on the basis of their solubility behavior; fulvic acids remain in solution at pH 2, but humic acids, having molecular weights of 20,000 to 100,000, are precipitated. Both are flexible polyelectrolytes that interact strongly with their own kind and with inorganic ions.


Associated with the fulvic and humic fractions are a wide variety smaller molecules such as alkanes, amino acids, amino sugars, sulfur and phosphorus derivatives of sugars, etc. Part of the organic carbon in a fertile soil is recycled in 1-2 years; plant residues, which are the major source of soil organic matter, have a half-life of days to months. Once carbon gets incorporated into humic substances, it is locked into a much slower recycling process; the turnover times of fulvic acids are a hundred years or more, while those for fulvic acids are around a thousand years. For this reason, humic substances are the major reservoir of organic carbon in soil. Organic matter, particularly polysaccharides, binds strongly to the cation components of clay colloids; the two together act as cementing agents and strongly influence the consistency and structure of the soil.


Soil water is held by capillary action and adsorption with varying degrees of tenacity. This water binding strength is traditionally expressed in terms of the pressure, or “tension” that would be required to force the water out of the soil. The tension of capillary water varies over a wide range of 0.1-32 atm; only in the lower half of this range will it be available to plants, which can exert an osmotic pressure of up to about 15 atm. Water in excess of the capillary capacity fills larger voids and is called gravitational water. Its presence in surface soils corresponds to a flooded condition that inhibits plant growth by reducing soil aeration.


The gas phase within soil pores generally has a CO2 content of 5-50 times that of the atmosphere due to the action of organisms. O2 tends to be depleted to roughly the extent that CO2 is present in excess. Under conditions of poor aeration (i.e., exchange with the atmosphere), considerable quantities of N2O, NO, H2, CH4, C2H4, and H2S may also be present.