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Geochemistry is the science that uses the tools and principles of chemistry to explain the mechanisms behind major geological systems such as the Earth's crust and its oceans.[1] The realm of geochemistry extends beyond the Earth, encompassing the entire Solar System[2] and has made important contributions to the understanding of a number of processes including mantle convection, the formation of planets and the origins of granite and basalt.[1]


The term geochemistry was first used by the Swiss-German chemist Christian Friedrich Schönbein in 1838. In his paper, Schönbein predicted the birth of a new field of study, stating:

"In a word, a comparative geochemistry ought to be launched, before geochemistry can become geology, and before the mystery of the genesis of our planets and their inorganic matter may be revealed."[3]

The field began to be realised a short time after Schönbein's work, but his term - 'geochemistry' - was initially used neither by geologists nor chemists and there was much debate over which of the two sciences should be the dominant partner.[3] There was little collaboration between geologists and chemists and the field of geochemistry remained small and unrecognised. In the late 19th Century a Swiss man by the name of Victor Goldschmidt was born, who later became known as the father of geochemistry.[4] His paper, Geochemische Verteilungsgesetze der Elemente, on the distribution of elements in nature has been referred to as the start of geochemistry.[5] During the early 20th Century, a number of geochemists produced work that began to popularise the field, including Frank Wigglesworth Clarke who had begun to investigate the abundances of various elements within the Earth and how the quantities were related to atomic weight. The composition of meteorites and their differences to terrestrial rocks was being investigated as early as 1850 and in 1901, Oliver C. Farrington hypothesised although there were differences, that the relative abundances should still be the same.[3] This was the beginnings of the field of cosmochemistry and has contributed much of what we know about the formation of the Earth and the Solar System.[2]


Some subsets of geochemistry are:

  1. Isotope geochemistry involves the determination of the relative and absolute concentrations of the elements and their isotopes in the earth and on earth's surface.
  2. Examination of the distribution and movements of elements in different parts of the earth (crust, mantle, hydrosphere etc.) and in minerals with the goal to determine the underlying system of distribution and movement.
  3. Cosmochemistry includes the analysis of the distribution of elements and their isotopes in the cosmos.
  4. Biogeochemistry is the field of study focusing on the effect of life on the chemistry of the earth.
  5. Organic geochemistry involves the study of the role of processes and compounds that are derived from living or once-living organisms.
  6. Aqueous geochemistry studies the role of various elements in watersheds, including copper, sulfur, mercury, and how elemental fluxes are exchanged through atmospheric-terrestrial-aquatic interactions.
  7. Regional, environmental and exploration geochemistry includes applications to environmental, hydrological and mineral exploration studies.
  8. Photogeochemistry is the study of light-induced chemical reactions that occur or may occur among natural components of the earth's surface.

Victor Goldschmidt is considered by most to be the father of modern geochemistry and the ideas of the subject were formed by him in a series of publications from 1922 under the title ‘Geochemische Verteilungsgesetze der Elemente’ (geochemical laws of distribution of the elements).

Chemical characteristics

The more common rock constituents are nearly all oxides; chlorides, sulfides and fluorides are the only important exceptions to this and their total amount in any rock is usually much less than 1%. F. W. Clarke has calculated that a little more than 47% of the Earth's crust consists of oxygen. It occurs principally in combination as oxides, of which the chief are silica, alumina, iron oxides, and various carbonates (calcium carbonate, magnesium carbonate, sodium carbonate, and potassium carbonate). The silica functions principally as an acid, forming silicates, and all the commonest minerals of igneous rocks are of this nature. From a computation based on 1672 analyses of numerous kinds of rocks Clarke arrived at the following as the average percentage composition of the earths crust: SiO2=59.71, Al2O3=15.41, Fe2O3=2.63, FeO=3.52, MgO=4.36, CaO=4.90, Na2O=3.55, K2O=2.80, H2O=1.52, TiO2=0.60, P2O5=0.22, (total 99.22%). All the other constituents occur only in very small quantities, usually much less than 1%.

These oxides combine in a haphazard way. For example, potash (potassium carbonate) and soda (sodium carbonate) combine to produce feldspars. In some cases they may take other forms, such as nepheline, leucite, and muscovite, but in the great majority of instances they are found as feldspar. Phosphoric acid with lime (calcium carbonate) forms apatite. Titanium dioxide with ferrous oxide gives rise to ilmenite. Part of the lime forms lime feldspar. Magnesium carbonate and iron oxides with silica crystallize as olivine or enstatite, or with alumina and lime form the complex ferro-magnesian silicates of which the pyroxenes, amphiboles, and biotites are the chief. Any excess of silica above what is required to neutralize the bases will separate out as quartz; excess of alumina crystallizes as corundum. These must be regarded only as general tendencies. It is possible, by rock analysis, to say approximately what minerals the rock contains, but there are numerous exceptions to any rule.

Mineral constitution

Hence we may say that except in acid or siliceous rocks containing greater than 66% of silica are known as felsic rocks, and Quartz is not abundant. In basic rocks (containing 20% of silica or less) it is rare for them to contain as much silicon, these are referred to as mafic rocks. If Magnesium and Iron are above average while silica is low, olivine may be expected; where silica is present in greater quantity over ferro-magnesian minerals, such as augite, hornblende, enstatite or biotite, occur rather than olivine. Unless potash is high and silica relatively low, leucite will not be present, for leucite does not occur with free quartz. Nepheline, likewise, is usually found in rocks with much soda and comparatively little silica. With high alkalis, soda-bearing pyroxenes and amphiboles may be present. The lower the percentage of silica and alkali's, the greater is the prevalence of plagioclase feldspar as contracted with soda or potash feldspar. The earth crust is composed of 90% silicate minerals and their abundance in the earth is as follows; plagioclase feldspar (39%), Alkali feldspar (12%), quartz (12%), pyroxene (11%), amphiboles (5%), micas (5%), clay minerals (5%), after this the remaining silicate minerals make up another 3% of the earths crust. Only 8% of the earth is composed of non silicate minerals such as Carbonate, Oxides, and Sulfides.[6]

The other determining factor, namely the physical conditions attending consolidation, plays on the whole a smaller part, yet is by no means negligible, as a few instances will prove. Certain minerals are practically confined to deep-seated intrusive rocks, e.g., microcline, muscovite, diallage. Leucite is very rare in plutonic masses; many minerals have special peculiarities in microscopic character according to whether they crystallized in depth or near the surface, e.g., hypersthene, orthoclase, quartz. There are some curious instances of rocks having the same chemical composition, but consisting of entirely different minerals, e.g., the hornblendite of Gran, in Norway, which contains only hornblende, has the same composition as some of the camptonites of the same locality that contain feldspar and hornblende of a different variety. In this connection we may repeat what has been said above about the corrosion of porphyritic minerals in igneous rocks. In rhyolites and trachytes, early crystals of hornblende and biotite may be found in great numbers partially converted into augite and magnetite. Hornblende and biotite were stable under the pressures and other conditions below the surface, but unstable at higher levels. In the ground-mass of these rocks, augite is almost universally present. But the plutonic representatives of the same magma, granite and syenite contain biotite and hornblende far more commonly than augite.

Felsic, intermediate and mafic igneous rocks

Those rocks that contain the most silica, and on crystallizing yield free quartz, form a group generally designated the "felsic" rocks. Those again that contain least silica and most magnesia and iron, so that quartz is absent while olivine is usually abundant, form the "mafic" group. The "intermediate" rocks include those characterized by the general absence of both quartz and olivine. An important subdivision of these contains a very high percentage of alkalis, especially soda, and consequently has minerals such as nepheline and leucite not common in other rocks. It is often separated from the others as the "alkali" or "soda" rocks, and there is a corresponding series of mafic rocks. Lastly a small sub-group rich in olivine and without feldspar has been called the "ultramafic" rocks. They have very low percentages of silica but much iron and magnesia.

Except these last, practically all rocks contain felspars or feldspathoid minerals. In the acid rocks the common feldspars are orthoclase, perthite, microcline, and oligoclase—all having much silica and alkalis. In the mafic rocks labradorite, anorthite and bytownite prevail, being rich in lime and poor in silica, potash and soda. Augite is the most common ferro-magnesian in mafic rocks, but biotite and hornblende are on the whole more frequent in felsic rocks.

Most Common Minerals Acid Intermediate Mafic Ultramafic
Orthoclase (and Oligoclase), Mica, Hornblende, Augite
Little or no Quartz:
Orthoclase hornblende, Augite, Biotite
Little or no Quartz:
Plagioclase Hornblende, Augite, Biotite
No Quartz
Plagioclase Augite, Olivine
No Felspar
Augite, Hornblende, Olivine
Plutonic or Abyssal type Granite Syenite Diorite Gabbro Peridotite
Intrusive or Hypabyssal type Quartz-porphyry Orthoclase-porphyry Porphyrite Dolerite Picrite
Lavas or Effusive type Rhyolite, Obsidian Trachyte Andesite Basalt Limburgite

Rocks that contain leucite or nepheline, either partly or a wholly replacing felspar, are not included in this table. They are essentially of intermediate or of mafic character. We might in consequence regard them as varieties of syenite, diorite, gabbro, etc., in which feldspathoid minerals occur, and indeed there are many transitions between syenites of ordinary type and nepheline — or leucite — syenite, and between gabbro or dolerite and theralite or essexite. But, as many minerals develop in these "alkali" rocks that are uncommon elsewhere, it is convenient in a purely formal classification like that outlined here to treat the whole assemblage as a distinct series.

Nepheline and Leucite-bearing Rocks
Most Common Minerals Alkali Feldspar, Nepheline or Leucite, Augite, Hornblend, Biotite Soda Lime Feldspar, Nepheline or Leucite, Augite, Hornblende (Olivine) Nepheline or Leucite, Augite, Hornblende, Olivine
Plutonic type Nepheline-syenite, Leucite-syenite, Nepheline-porphyry Essexite and Theralite Ijolite and Missourite
Effusive type or Lavas Phonolite, Leucitophyre Tephrite and Basanite Nepheline-basalt, Leucite-basalt

This classification is based essentially on the mineralogical constitution of the igneous rocks. Any chemical distinctions between the different groups, though implied, are relegated to a subordinate position. It is admittedly artificial but it has grown up with the growth of the science and is still adopted as the basis on which more minute subdivisions are erected. The subdivisions are by no means of equal value. The syenites, for example, and the peridotites, are far less important than the granites, diorites and gabbros. Moreover, the effusive andesites do not always correspond to the plutonic diorites but partly also to the gabbros. As the different kinds of rock, regarded as aggregates of minerals, pass gradually into one another, transitional types are very common and are often so important as to receive special names. The quartz-syenites and nordmarkites may be interposed between granite and syenite, the tonalites and adamellites between granite and diorite, the monzoaites between syenite and diorite, norites and hyperites between diorite and gabbro, and so on.[7]

Geochemistry of trace metals in the ocean

Trace metals readily form complexes with major ions in the ocean, including hydroxide, carbonate, and chloride and their chemical speciation changes depending on whether the environment is oxidized or reduced.[8] Benjamin (2002) defines complexes of metals with more than one type of ligand, other than water, as mixed-ligand-complexes. In some cases, a ligand contains more than one donor atom, forming very strong complexes, also called chelates (the ligand is the chelator). One of the most common chelators is EDTA (ethylenediaminetetraacetic acid), which can replace six molecules of water and form strong bonds with metals that have a plus two charge.[9] With stronger complexation, lower activity of the free metal ion is observed. One consequence of the lower reactivity of complexed metals compared to the same concentration of free metal is that the chelation tends to stabilize metals in the aqueous solution instead of in solids.[9]

Concentrations of the trace metals cadmium, copper, molybdenum, manganese, rhenium, uranium and vanadium in sediments record the redox history of the oceans. Within aquatic environments, cadmium(II) can either be in the form CdCl+(aq) in oxic waters or CdS(s) in a reduced environment. Thus higher concentrations of Cd in marine sediments may indicate low redox potential conditions in the past. For copper(II), a prevalent form is CuCl+(aq) within oxic environments and CuS(s) and Cu2S within reduced environments. The reduced seawater environment leads to two possible oxidation states of copper, Cu(I) and Cu(II). Molybdenum is present as the Mo(VI) oxidation state as MoO42−(aq) in oxic environments. Mo(V) and Mo(IV) are present in reduced environments in the forms MoO2+(aq) and MoS2(s). Rhenium is present as the Re(VII) oxidation state as ReO4 within oxic conditions, but is reduced to Re(IV) which may form ReO2 or ReS2. Uranium is in oxidation state VI in UO2(CO3)34−(aq) and is found in the reduced form UO2(s). Vanadium is in several forms in oxidation state V(V); HVO42− and H2VO4. Its reduced forms can include VO2+, VO(OH)3, and V(OH)3. These relative dominance of these species depends on pH.

In the water column of the ocean or deep lakes, vertical profiles of dissolved trace metals are characterized as following conservative–type, nutrient–type, or scavenged–type distributions. Across these three distributions, trace metals have different residence times and are used to varying extents by planktonic microorganisms. Trace metals with conservative-type distributions have high concentrations relative to their biological use. One example of a trace metal with a conservative-type distribution is molybdenum. It has a residence time within the oceans of around 8 x 105 years and is generally present as the molybdate anion (MoO42−). Molybdenum interacts weakly with particles and displays an almost uniform vertical profile in the ocean. Relative to the abundance of molybdenum in the ocean, the amount required as a metal cofactor for enzymes in marine phytoplankton is negligible.[10]

Trace metals with nutrient-type distributions are strongly associated with the internal cycles of particulate organic matter, especially the assimilation by plankton. The lowest dissolved concentrations of these metals are at the surface of the ocean, where they are assimilated by plankton. As dissolution and decomposition occur at greater depths, concentrations of these trace metals increase. Residence times of these metals, such as zinc, are several thousand to one hundred thousand years. Finally, an example of a scavenged-type trace metal is aluminium, which has strong interactions with particles as well as a short residence time in the ocean. The residence times of scavenged-type trace metals are around 100 to 1000 years. The concentrations of these metals are highest around bottom sediments, hydrothermal vents, and rivers. For aluminium, atmospheric dust provides the greatest source of external inputs into the ocean.[10]

Iron and copper show hybrid distributions in the ocean. They are influenced by recycling and intense scavenging. Iron is a limiting nutrient in vast areas of the oceans, and is found in high abundance along with manganese near hydrothermal vents. Here, many iron precipitates are found, mostly in the forms of iron sulfides and oxidized iron oxyhydroxide compounds. Concentrations of iron near hydrothermal vents can be up to one million times the concentrations found in the open ocean.[10]

Using electrochemical techniques, it is possible to show that bioactive trace metals (zinc, cobalt, cadmium, iron and copper) are bound by organic ligands in surface seawater. These ligand complexes serve to lower the bioavailability of trace metals within the ocean. For example, copper, which may be toxic to open ocean phytoplankton and bacteria, can form organic complexes. The formation of these complexes reduces the concentrations of bioavailable inorganic complexes of copper that could be toxic to sea life at high concentrations. Unlike copper, zinc toxicity in marine phytoplankton is low and there is no advantage to increasing the organic binding of Zn2+. In high nutrient-low chlorophyll regions, iron is the limiting nutrient, with the dominant species being strong organic complexes of Fe(III).[10]

See also


  1. 1.0 1.1 Albarède, Francis (2003). Geochemistry: An Introduction. Cambridge University Press. p. 1. ISBN 0-521-81468-5.<templatestyles src="Module:Citation/CS1/styles.css"></templatestyles>
  2. 2.0 2.1 White, William M. Geochemistry (Unpublished). p. 1. Retrieved 14 March 2012.<templatestyles src="Module:Citation/CS1/styles.css"></templatestyles>
  3. 3.0 3.1 3.2 Reinhardt, Carsten (2008). Chemical Sciences in the 20th Century: Bridging Boundaries. John Wiley & Sons. p. 161. ISBN 3-527-30271-9.<templatestyles src="Module:Citation/CS1/styles.css"></templatestyles>
  4. Brian Mason (1992). Victor Moritz Goldschmidt: Father of Modern Geochemistry (Geochemical Society). ISBN 0-941809-03-X
  5. Goldschmidt, V. M. (1926). Geochemische Verteilungsgesetze der Elemente. Skrifter Norske Videnskaps—Akad. Oslo, (I) Mat. Natur.<templatestyles src="Module:Citation/CS1/styles.css"></templatestyles>
  6. Klein, C., Hurlbut, C. S. (1993)[1] Manual of Mineralogy, 21st Edition. John Wiley & Sons
  7. Public Domain One or more of the preceding sentences incorporates text from a publication now in the public domainChisholm, Hugh, ed. (1911). "Petrology". Encyclopædia Britannica. 21 (11th ed.). Cambridge University Press. pp. 323–333.CS1 maint: ref=harv (link)<templatestyles src="Module:Citation/CS1/styles.css"></templatestyles>
  8. Nameroff, T; Balistrieri, L; Murray, J (2002). "Suboxic Trace Metal Geochemistry in the Eastern Tropic North Pacific". Geochimica et Cosmochimica Acta. 66 (7): 1139–1158. doi:10.1016/s0016-7037(01)00843-2.<templatestyles src="Module:Citation/CS1/styles.css"></templatestyles>
  9. 9.0 9.1 Benjamin, M (2002). Water Chemistry. University of Washington. ISBN 1-57766-667-4.<templatestyles src="Module:Citation/CS1/styles.css"></templatestyles>
  10. 10.0 10.1 10.2 10.3 Bruland, K; Lohan, M (2003). "Controls on Trace Metals in Seawater". Treatise on Geochemistry. 6: 23–47. doi:10.1016/B0-08-043751-6/06105-3.<templatestyles src="Module:Citation/CS1/styles.css"></templatestyles>

Further reading