Geochemistry

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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. The realm of geochemistry extends beyond the Earth, encompassing the entire Solar System, 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.

History

Victor Goldschmidt (1909)

The term geochemistry was first used by the Swiss-German chemist Christian Friedrich Schönbein in 1838: "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] However, for the rest of the century the more common term was "chemical geology", and there was little contact between geologists and chemists.^[3]

Geochemistry emerged as a separate discipline after major laboratories were established, starting with the United States Geological Survey (USGS) in 1884, and began systematic surveys of the chemistry of rocks and minerals. The chief USGS chemist, Frank Wigglesworth Clarke, noted that the elements generally decrease in abundance as their atomic weights increase, and summarized the work on elemental abundance in The Data of Geochemistry.^[3][4]:2

The composition of meteorites was investigated and compared to terrestrial rocks as early as 1850. In 1901, Oliver C. Farrington hypothesised that, although there were differences, 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.^[5]

In the early 20th century, Max von Laue and William L. Bragg showed that X-ray scattering could be used to determine the structures of crystals. In the 1920s and 1930s, Victor Goldschmidt and associates at the University of Oslo applied these methods to many common minerals and formulated a set of rules for how elements are grouped. Goldschmidt published this work in the series Geochemische Verteilungsgesetze der Elemente [Geochemical Laws of the Distribution of Elements].

Subfields

Some subfields of geochemistry are:^[7]

• 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.^[8]
• Biogeochemistry is the field of study focusing on the effect of life on the chemistry of the earth.^[9]:3
• Cosmochemistry includes the analysis of the distribution of elements and their isotopes in the cosmos.^[2]: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.^[10]
• Organic geochemistry, the study of the role of processes and compounds that are derived from living or once-living organisms.^[11]
• Photogeochemistry is the study of light-induced chemical reactions that occur or may occur among natural components of the earth's surface.^[12]
• Regional geochemistry includes applications to environmental, hydrological and mineral exploration studies.^[13]

Chemical elements

The building blocks of materials are the chemical elements. These can be identified by their atomic number Z, which is the number of protons in the nucleus. An element can have more than one value for N, the number of neutrons in the nucleus. The sum of these is the mass number, which is roughly equal to the atomic mass. Atoms with the same atomic number but different neutron numbers are called isotopes. A given isotope is identified by a letter for the element preceded by a superscript for the mass number. For example, two common isotopes of chlorine are ³⁵Cl and ³⁷Cl. There are about 1700 known combinations of Z and N, of which only about 260 are stable. However, most of the unstable isotopes do not occur in nature. In geochemistry, stable isotopes are used to trace chemical pathways and reactions, while isotopes are primarily used to date samples.^[4]:13–17

The chemical behavior of an atom – its affinity for other elements and the type of bonds it forms – is determined by the arrangement of electrons in orbitals, particularly the outermost (valence) electrons. These arrangements are reflected in the position of elements in the periodic table.^[4]:13–17 Based on position, the elements fall into the broad groups of alkali metals, alkaline earth metals, transition metals, semi-metals (also known as metalloids), halogens, noble gases, lanthanides and actinides.


Another useful classification scheme for geochemistry is the Goldschmidt classification, which places the elements into four main groups. Lithophiles combine easily with oxygen. These elements, which include Na, K, Si, Al, Ti, Mg and Ca, dominate in the Earth's crust, forming silicates and other oxides. Siderophile elements (Fe, Co, Ni, Pt, Re, Os) have an affinity for iron and tend to concentrate in the core. Chalcophile elements (Cu, Ag, Zn, Pb, S) form sulfides; and atmophile elements (O, N, H and noble gases) dominate the atmosphere. Within each group, some elements are refractory, remaining stable at high temperatures, while others are volatile, evaporating more easily, so heating can separate them.

Differentiation and mixing

The chemical composition of the Earth and other bodies is determined by two opposing processes: differentiation and mixing. In the Earth's mantle, differentiation occurs at mid-ocean ridges through partial melting, with more refractory materials remaining at the base of the lithosphere while the remainder rises to form basalt. After an oceanic plate descends into the mantle, convection eventually mixes the two parts together. Erosion differentiates granite, separating it into clay on the ocean floor, sandstone on the edge of the continent, and dissolved minerals in ocean waters. Metamorphism and anatexis (partial melting of crustal rocks) can mix these elements together again. In the ocean, biological organisms can cause chemical differentiation, while dissolution of the organisms and their wastes can mix the materials again.^[1]:23–24

Fractionation

A major source of differentiation is fractionation, an unequal distribution of elements and isotopes. This can be the result of chemical reactions, phase changes, kinetic effects, or radioactivity.^[1]:2–3 On the largest scale, planetary differentiation is a physical and chemical separation of a planet into chemically distinct regions. For example, the terrestrial planets formed iron-rich cores and silicate-rich mantles and crusts.^[14]:218 In the Earth's mantle, the primary source of chemical differentiation is partial melting, particularly near mid-ocean ridges.^[15]:68,153 This can occur when the solid is heterogeneous or a solid solution, and part of the melt is separated from the solid. The process is known as equilibrium or batch melting if the solid and melt remain in equilibrium until the moment that the melt is removed, and fractional or Rayleigh melting if is removed continuously.^[16]

Isotopic fractionation can have mass-dependent and mass-independent forms. Molecules with heavier isotopes have lower ground state energies and are therefore more stable. As a result, chemical reactions show a small isotope dependence, with heavier isotopes preferring species or compounds with a higher oxidation state; and in phase changes, heavier isotopes tend to concentrate in the heavier phases.^[17] Mass-dependent fractionation is largest in light elements because the difference in masses is a larger fraction of the total mass.^[18]:47

Ratios between isotopes are generally compared to a standard. For example, sulfur has four stable isotopes, of which the two most common are ³²S and ³⁴S.^[18]:98 The ratio of their concentrations, R=³⁴S/³²S, is reported as

${\displaystyle \delta {}^{34}\mathrm {S} =1000\left({\frac {R}{R_{\mathrm {s} }}}-1\right),}$

where R_s is the same ratio for a standard. Because the differences are small, the ratio is multiplied by 1000 to make it parts per thousand (referred to as parts per mil). This is represented by the symbol ‰.^[17]:55

Equilibrium

Equilibrium fractionation occurs between chemicals or phases that are in equilibrium with each other. In equilibrium fractionation between phases, heavier phases prefer the heavier isotopes. For two phases A and B, the effect can be represented by the factor

${\displaystyle a_{\mathrm {A-B} }={\frac {R_{\mathrm {A} }}{R_{\mathrm {B} }}}.}$

In the liquid-vapor phase transition for water, a_l-v at 20 degrees Celsius is 1.0098 for ¹⁸O and 1.084 for ²H. In general, fractionation is greater at lower temperatures. At 0 °C, the factors are 1.0117 and 1.111.^[17]:59

Kinetic

When there is not equilibrium between phases or chemical compounds, kinetic fractionation can occur. For example, at interfaces between liquid water and air, the forward reaction is enhanced if the humidity of the air is less than 100% or the water vapor is moved by a wind. Kinetic fractionation generally is enhanced compared to equilibrium fractionation, and depends on factors such as reaction rate, reaction pathway and bond energy. Since lighter isotopes generally have weaker bonds, they tend to react faster and enrich the reaction products.^[17]:60

Biological fractionation is a form of kinetic fractionation, since reactions tend to be in one direction. Biological organisms prefer lighter isotopes because there is a lower energy cost in breaking energy bonds. In addition to the previously mentioned factors, the environment and species of the organism can have a large effect on the fractionation.^[17]:70

Cycles

Through a variety of physical and chemical processes, chemical elements change in concentration and move around in what are called geochemical cycles. An understanding of these changes requires both detailed observation and theoretical models. Each chemical compound, element or isotope has a concentration that is a function C(r,t) of position and time, but it is impractical to model the full variability. Instead, in an approach borrowed from chemical engineering,^[1]:81 geochemists average the concentration over regions of the earth called geochemical reservoirs. The choice of reservoir depends on the problem; for example, the ocean may be a single reservoir or be split into multiple reservoirs.^[19] In a type of model called a box model, a reservoir is represented by a box with inputs and outputs.^[1]:81[19]
Geochemical models generally involve feedback. In the simplest case of a linear cycle, either the input or the output from a reservoir is proportional to the concentration. For example, salt is removed from the ocean by formation of evaporites, and given a constant rate of evaporation in evaporite basins, the rate of removal of salt should be proportional to its concentration. For a given component C, if the input to a reservoir is a constant a and the output is kC for some constant k, then the mass balance equation is

${\displaystyle {\frac {dC}{dt}}=a-kC.}$

(1)

This expresses the fact that any change in mass must be balanced by changes in the input or output. On a time scale of t = 1/k, the system approaches a steady state in which C^steady = a/k. The residence time is defined as

${\displaystyle \tau _{\mathrm {res} }=C^{\text{steady}}/I=C^{\text{steady}}/O,}$

where I and O are the input and output rates. In the above example, the steady-state input and output rates are both equal to a, so τ_res = 1/k.^[19]

If the input and output rates are nonlinear functions of C, they may still be closely balanced over time scales much greater than the residence time; otherwise there will be large fluctuations in C. In that case, the system is always close to a steady state and a lowest order expansion of the mass balance equation will lead to a linear equation like Equation (1). In most systems, one or both of the input and output depend on C, resulting in a feedback that tends to maintain the steady state. If an external forcing perturbs the system, it will return to the steady state on a time scale of 1/k.^[19]

Abundance of elements

Solar System

Abundances of solar system elements.^[20]

The composition of the Solar System is similar to that of many other stars, and aside from small anomalies it can be assumed to have formed from a solar nebula that had a uniform composition, and the composition of the Sun's photosphere is similar to that of the rest of the Solar System. The composition of the photosphere is determined by fitting the absorption lines in its spectrum to models of the Sun's atmosphere.^[21] By far the largest two elements by fraction of total mass are hydrogen (74.9%) and helium (23.8%), with all the remaining elements contributing just 1.3%.^[22] There is a general trend of exponential decrease in abundance with increasing atomic number, although elements with even atomic number are more common than their odd-numbered neighbors (the Oddo–Harkins rule). Compared to the overall trend, lithium, boron and beryllium are depleted and iron is anomalously enriched.^{[23]:284–285}

The pattern of elemental abundance is mainly due to two factors. The hydrogen, helium, and some of the lithium were formed in about 20 minutes after the Big Bang, while the rest were created in the interiors of stars.^{[4]:316–317}

Meteorites

Meteorites come in a variety of compositions, but chemical analysis can determine whether they were once in planetesimals that melted or differentiated.^[21]:45 Chondrites are undifferentiated and have round mineral inclusions called chondrules. With ages of 4.56 billion years, they date to the early solar system. A particular kind, the CI chondrite, has a composition that closely matches that of the Sun's photosphere, except for depletion of some volatiles (H, He, C, N, O) and a group of elements (Li, B, Be) that are destroyed by nucleosynthesis in the Sun.^[4]:318[21] Because of the latter group, CI chondrites are considered a better match for the composition of the early Solar System. Moreover, the chemical analysis of CI chondrites is more accurate than for the photosphere, so it is generally used as the source for chemical abundance, despite their rareness (only five have been recovered on Earth).^[21]

Giant planets

Cutaways illustrating models of the interiors of the giant planets.

The planets of the Solar System are divided into two groups: the four inner planets are the terrestrial planets (Mercury, Venus, Earth and Mars), with relatively small sizes and rocky surfaces. The four outer planets are the giant planets, which are dominated by hydrogen and helium and have lower mean densities. These can be further subdivided into the gas giants (Jupiter and Saturn) and the ice giants (Uranus and Neptune) that have large icy cores.^{[24]:26–27,283–284}

Most of our direct information on the composition of the giant planets is from spectroscopy. Since the 1930s, Jupiter was known to contain hydrogen, methane and ammonium. In the 1960s, interferometry greatly increased the resolution and sensitivity of spectral analysis, allowing the identification of a much greater collection of molecules including ethane, acetylene, water and carbon monoxide.^{[25]:138–139} However, Earth-based spectroscopy becomes increasingly difficult with more remote planets, since the reflected light of the Sun is much dimmer; and spectroscopic analysis of light from the planets can only be used to detect vibrations of molecules, which are in the infrared frequency range. This constrains the abundances of the elements H, C and N.^[25]:130 Two other elements are detected: phosphorus in the gas phosphine (PH₃) and germanium in germane (GeH₄).^[25]:131

The helium atom has vibrations in the ultraviolet range, which is strongly absorbed by the atmospheres of the outer planets and Earth. Thus, despite its abundance, helium was only detected once spacecraft were sent to the outer planets, and then only indirectly through collision-induced absortion in hydrogen molecules.^[25]:209 Further information on Jupiter was obtained from the Galileo probe when it was sent into the atmosphere in 1995;^[26][27] and the final mission of the Cassini probe in 2017 was to enter the atmosphere of Saturn.^[28] In the atmosphere of Jupiter, He was found to be depleted by a factor of 2 compared to solar composition and Ne by a factor of 10, a surprising result since the other noble gases and the elements C, N and S were enhanced by factors of 2 to 4 (oxygen was also depleted but this was attributed to the unusually dry region that Galileo sampled).^[27]

Spectroscopic methods only penetrate the atmospheres of Jupiter and Saturn to depths where the pressure is about equal to 1 bar, approximately Earth's atmospheric pressure at sea level.^[25]:131 The Galileo probe penetrated to 22 bars.^[27] This is a small fraction of the planet, which is expected to reach pressures of over 40 Mbar. To constrain the composition in the interior, thermodynamic models are constructed using information on temperature from infrared emission spectra and equations of state for the likely compositions.^[25]:131 High pressure experiments predict that hydrogen will be a metallic liquid in the interior of Jupiter and Saturn, while in Uranus and Neptune it remains in the molecular state.^{[25]:135–136} Estimates also depend on models for the formation of the planets. Condensation of the presolar nebula would result in a gaseous planet with the same composition as the Sun, but the planets could also have formed when a solid core captured nebular gas.^[25]:136

In current models, the four giant planets have cores of rock and ice that are roughly the same size, but the proportion of hydrogen and helium decreases from about 300 Earth masses in Jupiter to 75 in Saturn and just a few in Uranus and Neptune.^[25]:220 Thus, while the gas giants are primarily composed of hydrogen and helium, the ice giants are primarily composed of heavier elements (O, C, N, S), primarily in the form of water, methane and ammonia. The surfaces are cold enough for molecular hydrogen to be liquid, so much of each planet is likely a hydrogen ocean overlaying one of heavier compounds.^[29] Outside the core, Jupiter has a mantle of liquid metallic hydrogen and an atmosphere of molecular hydrogen and helium. Metallic hydrogen does not mix well with helium, and in Saturn it may form a separate layer below the metallic hydrogen.^[25]:138

Terrestrial planets

Terrestrial planets are believed to have come from the same nebular material as the giant planets, but they have lost most of the lighter elements and have different histories. Planets closer to the Sun might be expected to have a higher fraction of refractory elements, but if their later stages of formation involved collisions of large objects with orbits that sampled different parts of the Solar System, there could be little systematic dependence on position.^[30]:3–4
Direct information on Mars, Venus and Mercury largely comes from spacecraft missions. Using gamma-ray spectrometers, the composition of the crust of Mars has been measured by the Mars Odyssey orbiter,^[31] the crust of Venus by some of the Venera missions to Venus,^[30] and the crust of Mercury by the MESSENGER spacecraft.^[32] Additional information on Mars comes from meteorites that have landed on Earth (the Shergottites, Nakhlites, and Chassignites, collectively known as SNC meteorites).^[33]:124 Abundances are also constrained by the masses of the planets, while the internal distribution of elements is constrained by their moments of inertia.^[4]:334

The planets condensed from the solar nebula, and much of the details of their composition are determined by fractionation as they cooled. The phases that condense fall into five groups. First to condense are materials rich in refractory elements such as Ca and Al. These are followed by nickel and iron, then magnesium silicates. Below about 700 kelvins (700 K), FeS and volatile-rich metals and silicates form a fourth group, and in the fifth group FeO enter the magnesium silicates.^[34] The compositions of the planets and the Moon are chondritic, meaning that within each group the ratios between elements are the same as in carbonaceous chondrites.^[4]:334

The estimates of planetary compositions depend on the model used. In the equilibrium condensation model, each planet was formed from a feeding zone in which the compositions of solids were determined by the temperature in that zone. Thus, Mercury formed at 1400 K, where iron remained in a pure metallic form and there was little magnesium or silicon in solid form; Venus at 900 K, so all the magnesium and silicon condensed; Earth at 600 K, so it contains FeS and silicates; and Mars at 450 K, so FeO was incorporated into magnesium silicates. The greatest problem with this theory is that volatiles would not condense, so the planets would have no atmospheres and Earth no atmosphere.^{[4]:335–336}

In chondritic mixing models, the compositions of chondrites are used to estimate planetary compositions. For example, one model mixes two components, one with the composition of C1 chondrites and one with just the refractory components of C1 chondrites.^[4]:337 In another model, the abundances of the five fractionation groups are estimated using an index element for each group. For the most refractory group, uranium is used; iron for the second; the ratios of potassium and thalium to uranium for the next two; and the molar ratio FeO/(FeO+MgO) for the last. Using thermal and seismic models along with heat flow and density, Fe can be constrained to within 10 percent on Earth, Venus and Mercury. U can be constrained within about 30% on Earth, but its abundance on other planets is based on "educated guesses". One difficulty with this model is that there may be significant errors in its prediction of volatile abundances because some volatiles are only partially condensed.

Earth's crust

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 Earth's 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

Except in acid or siliceous igneous rocks containing greater than 66% of silica, known as felsic rocks, quartz is not abundant in igneous rocks. 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.

Earth's 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%); the remaining silicate minerals make up another 3% of Earth's crust. Only 8% of the Earth is composed of non-silicate minerals such as carbonates, oxides, and sulfides.^[35]

The other determining factor, namely the physical conditions attending consolidation, plays on the whole a smaller part, yet is by no means negligible. 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	Felsic	Intermediate		Mafic	Ultramafic
	Quartz 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.^[36]

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.^[37] 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.^[38] 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.^[38]

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 MoS_2(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.^[39]

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.^[39]

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.^[39]

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).

Geobiology

From Wikipedia, the free encyclopedia

 

The colorful microbial mats of Grand Prismatic Spring in Yellowstone National Park, USA. The orange mats are composed of Chloroflexi, Cyanobacteria, and other organisms that thrive in the 70˚C water. Geobiologists often study extreme environments like this because they are home to extremophilic organisms. It has been hypothesized that these environments may be representative of early Earth.^[1]

Geobiology is a field of scientific research that explores the interactions between the physical Earth and the biosphere. It is a relatively young field, and its borders are fluid. There is considerable overlap with the fields of ecology, evolutionary biology, microbiology, paleontology, and particularly biogeochemistry. Geobiology applies the principles and methods of biology and geology to the study of the ancient history of the co-evolution of life and Earth as well as the role of life in the modern world. Geobiologic studies tend to be focused on microorganisms, and on the role that life plays in altering the chemical and physical environment of the lithosphere, atmosphere, hydrosphere and/or cryosphere. It differs from biogeochemistry in that the focus is on processes and organisms over space and time rather than on global chemical cycles.

Geobiological research synthesizes the geologic record with modern biologic studies. It deals with process - how organisms affect the Earth and vice versa - as well as history - how the Earth and life have changed together. Much research is grounded in the search for fundamental understanding, but geobiology can also be applied, as in the case of microbes that clean up oil spills.

Geobiology employs molecular biology, environmental microbiology, chemical analyses, and the geologic record to investigate the evolutionary interconnectedness of life and Earth. It attempts to understand how the Earth has changed since the origin of life and what it might have been like along the way. Some definitions of geobiology even push the boundaries of this time frame - to understanding the origin of life and to the role that man has played and will continue to play in shaping the Earth in the Anthropocene.

History

A microbial mat in White Creek, Yellowstone National Park, USA. Note the conical microstructure of the bacterial communities. These are hypothesized to be a living analogue of ancient fossil stromatolites. Each cone has an oxygen gas bubble on top, the product of oxygenic photosynthesis by cyanobacteria in the multi-species microbial mats.

The term geobiology was coined by Lourens Baas Becking in 1934. In his words, geobiology "is an attempt to describe the relationship between organisms and the Earth," for "the organism is part of the Earth and its lot is interwoven with that of the Earth." Baas Becking's definition of geobiology was born of a desire to unify environmental biology with laboratory biology. The way he practiced it aligns closely with modern environmental microbial ecology, though his definition remains applicable to all of geobiology. In his book, Geobiology, Bass Becking stated that he had no intention of inventing a new field of study.^[4]Baas Becking's understanding of geobiology was heavily influenced by his predecessors, including Martinus Beyerinck, his teacher from the Dutch School of Microbiology. Others included Vladimir Vernadsky, who argued that life changes the surface environment of Earth in The Biosphere, his 1926 book,^[5] and Sergei Vinogradsky, famous for discovering lithotrophic bacteria.^[6]

The first laboratory officially dedicated to the study of geobiology was the Baas Becking Geobiological Laboratory in Australia, which opened its doors in 1965.^[4] However, it took another 40 or so years for geobiology to become a firmly rooted scientific discipline, thanks in part to advances in geochemistry and genetics that enabled scientists to begin to synthesize the study of life and planet.

In the 1930s, Alfred Treibs discovered chlorophyll-like porphyrins in petroleum, confirming its biological origin,^[7] thereby founding organic geochemistry and establishing the notion of biomarkers,^[8] a critical aspect of geobiology. But several decades passed before the tools were available to begin to search in earnest for chemical marks of life in the rocks. In the 1970s and '80s, scientists like Geoffrey Eglington and Roger Summons began to find lipid biomarkers in the rock record using equipment like GCMS.

On the biology side of things, in 1977, Carl Woese and George Fox published a phylogeny of life on Earth, including a new domain - the Archaea.^[9] And in the 1990s, genetics and genomics studies became possible, broadening the scope of investigation of the interaction of life and planet.

Today, geobiology has its own journals, such as Geobiology, established in 2003,^[10] and Biogeosciences, established in 2004,^[11] as well as recognition at major scientific conferences. It got its own Gordon Research Conference in 2011,^[12] a number of geobiology textbooks have been published,^[3][13] and many universities around the world offer degree programs in geobiology.

Major geobiological events

The geologic timescale overlain with major geobiologic events and occurrences. The oxygenation of the atmosphere is shown in blue starting 2.4 Ga, although the exact dating of the Great Oxygenation Event is debated.^[14]

Perhaps the most profound geobiological event is the introduction of oxygen into the atmosphere by photosynthetic bacteria. This oxygenation of Earth's primoidial atmosphere (the so-called oxygen catastrophe or Great Oxygenation Event) and the oxygenation of the oceans altered surface biogeochemical cycles and the types of organisms that have been evolutionarily selected for.

A subsequent major change was the advent of multicellularity. The presence of oxygen allowed eukaryotes and, later, multicellular life to evolve.

More anthropocentric geobiologic events include the origin of animals and the establishment of terrestrial plant life, which affected continental erosion and nutrient cycling, and likely changed the types of rivers observed, allowing channelization of what were previously predominantly braided rivers.

More subtle geobiological events include the role of termites in overturning sediments, coral reefs in depositing calcium carbonate and breaking waves, sponges in absorbing dissolved marine silica, the role of dinosaurs in breaching river levees and promoting flooding, and the role of large mammal dung in distributing nutrients.

Important concepts

Geobiology is founded upon a few core concepts that unite the study of Earth and life. While there are many aspects of studying past and present interactions between life and Earth that are unclear, several important ideas and concepts provide a basis of knowledge in geobiology that serve as a platform for posing researchable questions, including the evolution of life and planet and the co-evolution of the two, genetics - from both a historical and functional standpoint, the metabolic diversity of all life, the sedimentological preservation of past life, and the origin of life.

Co-evolution of life and Earth

A core concept in geobiology is that life changes over time through evolution. The theory of evolution postulates that unique populations of organisms or species arose from genetic modifications in the ancestral population which were passed down by drift and natural selection.^[17]

Along with standard biological evolution, life and planet co-evolve. Since the best adaptations are those that suit the ecological niche that the organism lives in, the physical and chemical characteristics of the environment drive the evolution of life by natural selection, but the opposite can also be true: with every advent of evolution, the environment changes.

A classic example of co-evolution is the evolution of oxygen-producing photosynthetic cyanobacteria which oxygenated Earth's Archean atmosphere. The ancestors of cyanobacteria began using water as an electron source to harness the energy of the sun and expelling oxygen before or during the early Paleoproterozoic. During this time, around 2.4 to 2.1 billion years ago,^[18] geologic data suggests that atmospheric oxygen began to rise in what is termed the Great Oxygenation Event (GOE).^[19][20] It is unclear for how long cyanobacteria had been doing oxygenic photosynthesis before the GOE. Some evidence suggests there were geochemical "buffers" or sinks suppressing the rise of oxygen such as volcanism^[21] though cyanobacteria may have been around producing it before the GOE.^[22] Other evidence indicates that the rise of oxygenic photosynthesis was coincident with the GOE.^[23]

Banded iron formation (BIF), Hammersley Formation, Western Australia

The presence of oxygen on Earth from its first production by cyanobacteria to the GOE and through today has drastically impacted the course of evolution of life and planet.^[19] It may have triggered the formation of oxidized minerals^[24] and the disappearance of oxidizable minerals like pyrite from ancient stream beds.^[25] The presence of banded-iron formations (BIFs) have been interpreted as a clue for the rise of oxygen since small amounts of oxygen could have reacted with reduced ferrous iron (Fe(II)) in the oceans, resulting in the deposition of sediments containing Fe(III) oxide in places like Western Australia.^[26] However, any oxidizing environment, including that provided by microbes such as the iron-oxidizing photoautotroph Rhodopseudomonas palustris,^[27] can trigger iron oxide formation and thus BIF deposition.^[28][29][30] Other mechanisms include oxidation by UV light.^[31] Indeed, BIFs occur across large swaths of Earth’s history and may not correlate with only one event.^[30]

Other changes correlated with the rise of oxygen include the appearance of rust-red ancient paleosols,^[19] different isotope fractionation of elements such as sulfur,^[32] and global glaciations and Snowball Earth events,^[33] perhaps caused by the oxidation of methane by oxygen, not to mention an overhaul of the types of organisms and metabolisms on Earth. Whereas organisms prior to the rise of oxygen were likely poisoned by oxygen gas as many anaerobes are today,^[34] those that evolved ways to harness the electron-accepting and energy-giving power of oxygen were poised to thrive and colonize the aerobic environment.

Modern, living stromatolites in Shark Bay, Australia. Shark Bay is one of the few places in the world where stromatolites can be seen today, though they were likely common in ancient shallow seas before the rise of metazoan predators.

The Earth has changed

Earth has not remained the same since its planetary formation 4.5 billion years ago.^[35][36] Continents have formed, broken up, and collided, offering new opportunities for and barriers to the dispersal of life. The redox state of the atmosphere and the oceans has changed, as indicated by isotope data. Fluctuating quantities of inorganic compounds such as carbon dioxide, nitrogen, methane, and oxygen have been driven by life evolving new biological metabolisms to make these chemicals and have driven the evolution of new metabolisms to use those chemicals. Earth acquired a magnetic field about 3.4 Ga^[37] that has undergone a series of geomagnetic reversals on the order of millions of years.^[38] The surface temperature is in constant fluctuation, falling in glaciations and Snowball Earth events due to ice-albedo feedback,^[39] rising and melting due to volcanic outgassing, and stabilizing due to silicate weathering feedback.^[40]

And the Earth is not the only one that changed - the luminosity of the sun has increased over time. Because rocks record a history of relatively constant temperatures since Earth’s beginnings, there must have been more greenhouse gasses to keep the temperatures up in the Archean when the sun was younger and fainter.^[41] All these major differences in the environment of the Earth placed very different constraints on the evolution of life throughout our planet’s history. Moreover, more subtle changes in the habitat of life are always occurring, shaping the organisms and traces that we observe today and in the rock record.

Genes encode geobiological function and history

The genetic code is key to observing the history of evolution and understanding the capabilities of organisms. Genes are the basic unit of inheritance and function and, as such, they are the basic unit of evolution and the means behind metabolism.^[42]

Phylogeny predicts evolutionary history

A phylogenetic tree of living things, based on rRNA data and proposed by Carl Woese, showing the separation of bacteria, archaea, and eukaryotes and linking the three branches of living organisms to the LUCA (the black trunk at the bottom of the tree).

Phylogeny takes genetic sequences from living organisms and compares them to each other to reveal evolutionary relationships, much like a family tree reveals how individuals are connected to their distant cousins.^[43] It allows us to decipher modern relationships and infer how evolution happened in the past.

Phylogeny can give some sense of history when combined with a little bit more information. Each difference in the DNA indicates divergence between one species and another.^[43] This divergence, whether via drift or natural selection, is representative of some lapse of time.^[43] Comparing DNA sequences alone gives a record of the history of evolution with an arbitrary measure of phylogenetic distance “dating” that last common ancestor. However, if information about the rate of genetic mutation is available or geologic markers are present to calibrate evolutionary divergence (i.e. fossils), we have a timeline of evolution.^[44] From there, with an idea about other contemporaneous changes in life and environment, we can begin to speculate why certain evolutionary paths might have been selected for.^[45]

Genes encode metabolism

Molecular biology allows scientists to understand a gene’s function using microbial culturing and mutagenesis. Searching for similar genes in other organisms and in metagenomic and metatranscriptomic data allows us to understand what processes could be relevant and important in a given ecosystem, providing insight into the biogeochemical cycles in that environment.

For example, an intriguing problem in geobiology is the role of organisms in the global cycling of methane. Genetics has revealed that the methane monooxygenase gene (pmo) is used for oxidizing methane and is present in all aerobic methane-oxidizers, or methanotrophs.^[46] The presence of DNA sequences of the pmo gene in the environment can be used as a proxy for methanotrophy.^[47][48] A more generalizable tool is the 16S ribosomal RNA gene, which is found in bacteria and archaea. This gene evolves very slowly over time and is not usually horizontally transferred, and so it is often used to distinguish different taxonomic units of organisms in the environment.^[9][49] In this way, genes are clues to organismal metabolism and identity. Genetics enables us to ask 'who is there?' and 'what are they doing?' This approach is called metagenomics.^[49]

3.4 billion year-old stromatolites from the Warrawoona Group, Western Australia. While the origin of Precambrian stromatolites is a heavily debated topic in geobiology,^[50] stromatolites from Warrawoona are hypothesized to have been formed by ancient communities of microbes.^[51]

Metabolic diversity influences the environment

Life harnesses chemical reactions to generate energy, perform biosynthesis, and eliminate waste.^[52] Different organisms use very different metabolic approaches to meet these basic needs.^[53] While animals such as ourselves are limited to aerobic respiration, other organisms can "breathe" sulfate (SO42-), nitrate (NO3-), ferric iron (Fe(III)), and uranium (U(VI)), or live off energy from fermentation.^[53] Some organisms, like plants, are autotrophs, meaning that they can fix carbon dioxide for biosynthesis. Plants are photoautotrophs, in that they use the energy of light to fix carbon. Microorganisms employ oxygenic and anoxygenic photoautotrophy, as well as chemoautotrophy. Microbial communities can coordinate in syntrophic metabolisms to shift reaction kinetics in their favor. Many organisms can perform multiple metabolisms to achieve the same end goal; these are called mixotrophs.^[53]

Biotic metabolism is directly tied to the global cycling of elements and compounds on Earth. The geochemical environment fuels life, which then produces different molecules that go into the external environment. (This is directly relevant to biogeochemistry.) In addition, biochemical reactions are catalyzed by enzymes which sometimes prefer one isotope over others. For example, oxygenic photosynthesis is catalyzed by RuBisCO, which prefers carbon-12 over carbon-13, resulting in carbon isotope fractionation in the rock record.^[54]

"Giant" ooids of the Johnnie Formation in the Death Valley area, California, USA. Ooids are near-spheroidal calcium carbonate grains that accumulate around a central nucleus and can be sedimented to form oolite like this. Microbes can mediate the formation of ooids.^[50]

Sedimentary rocks tell a story

Sedimentary rocks preserve remnants of the history of life on Earth in the form of fossils, biomarkers, isotopes, and other traces. The rock record is far from perfect, and the preservation of biosignatures is a rare occurrence. Understanding what factors determine the extent of preservation and the meaning behind what is preserved are important components to detangling the ancient history of the co-evolution of life and Earth.^[8] The sedimentary record allows scientists to observe changes in life and Earth in composition over time and sometimes even date major transitions, like extinction events.

Some classic examples of geobiology in the sedimentary record include stromatolites and banded-iron formations. The role of life in the origin of both of these is a heavily debated topic.^[19]

Life is fundamentally chemistry

The first life arose from abiotic chemical reactions. When this happened, how it happened, and even what planet it happened on are uncertain. However, life follows the rules of and arose from lifeless chemistry and physics. It is constrained by principles such as thermodynamics. This is an important concept in the field because it is represents the epitome of the interconnectedness, if not sameness, of life and Earth.^[55]

While often delegated to the field of astrobiology, attempts to understand how and when life arose are relevant to geobiology as well.^[56] The first major strides towards understanding the “how” came with the Miller-Urey experiment, when amino acids formed out of a simulated “primordial soup”. Another theory is that life originated in a system much like the hydrothermal vents at mid-oceanic spreading centers. In the Fischer-Tropsch synthesis, a variety of hydrocarbons form under vent-like conditions. Other ideas include the “RNA World” hypothesis, which postulates that the first biologic molecule was RNA and the idea that life originated elsewhere in the solar system and was brought to Earth, perhaps via a meteorite.^[55]

Methodology

A microbial mat growing on acidic soil in Norris Geyser basin, Yellowstone National Park, USA. The black top serves as a sort of sunscreen, and when you look underneath you see the green cyanobacteria.

While geobiology is a diverse and varied field, encompassing ideas and techniques from a wide range of disciplines, there are a number of important methods that are key to the study of the interaction of life and Earth that are highlighted here.^[3]

1. Laboratory culturing of microbes is used to characterize the metabolism and lifestyle of organisms of interest.
2. Gene sequencing allows scientists to study the relationships between extant organisms using phylogenetics.
3. Experimental genetic manipulation or mutagenesis is used to determine the function of genes in living organisms.
4. Microscopy is used to visualize the microbial world. Microscope work ranges from environmental observation to quantitative studies with DNA probes to high-definition visualization of the microbe-mineral interface by electron microscope (EM).
5. Isotope tracers can be used to track biochemical reactions to understand microbial metabolism.
6. Isotope natural abundance in rocks can be measured to look for isotopic fractionation that is consistent with biologic origin.
7. Detailed environmental characterization is important to understanding what about a habitat might be driving life’s evolution and, in turn, how life might be changing that niche. It includes and is not limited to, temperature, light, pH, salinity, concentration of specific molecules like oxygen, and the biologic community.
8. Sedimentology and stratigraphy are used to read the rocks. The rock record stores a history of geobiologic processes in sediments which can be unearthed through an understanding of deposition, sedimentation, compaction, diagenesis, and deformation.
9. The search for and study of fossils, while often delegated to the separate field of paleontology, is important in geobiology, though the scale of fossils is typically smaller (micropaleontology).
10. The biochemical analysis of biomarkers, which are fossilized or modern molecules that are indicative of the presence of a certain group of organisms or metabolism, is used to answer the evidence for life and metabolic diversity questions.^[8]
11. Paleomagnetics is the study of the planet's ancient magnetic field. It is significant to understanding magnetofossils, biomineralization, and global ecosystem changes.

Sub-disciplines and related fields

As its name suggests, geobiology is closely related to many other fields of study, and does not have clearly defined boundaries or perfect agreement on what exactly they comprise. Some practitioners take a very broad view of its boundaries, encompassing many older, more established fields such as biogeochemistry, paleontology, and microbial ecology. Others take a more narrow view, assigning it to emerging research that falls between these existing fields, such as with geomicrobiology. The following list includes both those that are clearly a part of geobiology, e.g. geomicrobiology, as well as those that share scientific interests but have not historically been considered a sub-discipline of geobiology, e.g. paleontology.

Astrobiology

Astrobiology is an interdisciplinary field that uses a combination of geobiological and planetary science data to establish a context for the search for life on other planets. The origin of life from non-living chemistry and geology, or abiogenesis, is a major topic in astrobiology. Even though it is fundamentally an earth-bound concern, and therefore of great geobiological interest, getting at the origin of life necessitates considering what life requires, what, if anything, is special about Earth, what might have changed to allow life to blossom, what constitutes evidence for life, and even what constitutes life itself. These are the same questions that scientists might ask when searching for alien life. In addition, astrobiologists research the possibility of life based on other metabolisms and elements, the survivability of Earth’s organisms on other planets or spacecrafts, planetary and solar system evolution, and space geochemistry.^[57]

Biogeochemistry

Biogeochemistry is a systems science that synthesizes the study of biological, geological, and chemical processes to understand the reactions and composition of the natural environment. It is concerned primarily with global elemental cycles, such as that of nitrogen and carbon. The father of biogeochemistry was James Lovelock, whose “Gaia hypothesis” proposed that Earth’s biological, chemical, and geologic systems interact to stabilize the conditions on Earth that support life.^[58]

Geobiochemistry

Stromatolites in the Green River Shale, Wyoming, USA, dating to the Eocene

Geobiochemistry is similar to biogeochemistry, but differs by placing emphasis on the effects of geology on the development of life’s biochemical processes, as distinct from the role of life on Earth’s cycles. Its primary goal is to link biological changes, encompassing evolutionary modifications of genes and changes in the expression of genes and proteins, to changes in the temperature, pressure, and composition of geochemical processes to understand when and how metabolism evolved. Geobiochemistry is founded on the notion that life is a planetary response because metabolic catalysis enables the release of energy trapped by a cooling planet.^[59]

Environmental microbiology

Microbiology is a broad scientific discipline pertaining to the study of that life which is best viewed under a microscope. It encompasses several fields that are of direct relevance to geobiology, and the tools of microbiology all pertain to geobiology. Environmental microbiology is especially entangled in geobiology since it seeks an understanding of the actual organisms and processes that are relevant in nature, as opposed to the traditional lab-based approach to microbiology. Microbial ecology is similar, but tend to focus more on lab studies and the relationships between organisms within a community, as well as within the ecosystem of their chemical and geological physical environment. Both rely on techniques such as sample collection from diverse environments, metagenomics, DNA sequencing, and statistics.

Geomicrobiology and microbial geochemistry

A vertical cross section of a microbial mat containing different organisms that perform different metabolisms. The green are presumably cyanobacteria, and teepee-like microstructures are visible on the surface.

Geomicrobiology traditionally studies the interactions between microbes and minerals. While it is generally reliant on the tools of microbiology, microbial geochemistry uses geological and chemical methods to approach the same topic from the perspective of the rocks. Geomicrobiology and microbial geochemistry (GMG) is a relatively new interdisciplinary field that more broadly takes on the relationship between microbes, Earth, and environmental systems. Billed as a subset of both geobiology and geochemistry, GMG seeks to understand elemental biogeochemical cycles and the evolution of life on Earth. Specifically, it asks questions about where microbes live, their local and global abundance, their structural and functional biochemistry, how they have evolved, biomineralization, and their preservation potential and presence in the rock record. In many ways, GMG appears to be equivalent to geobiology, but differs in scope: geobiology focuses on the role of all life, while GMG is strictly microbial. Regardless, it is these tiniest creatures that dominated to history of life integrated over time and seem to have had the most far-reaching effects.^[60]

Molecular geomicrobiology

Molecular geomicrobiology takes a mechanistic approach to understanding biological processes that are geologically relevant. It can be at the level of DNA, protein, lipids, or any metabolite.

Organic geochemistry

Organic geochemistry is the study of organic molecules that appear in the fossil record in sedimentary rocks. Research in this field concerns molecular fossils that are often lipid biomarkers. Molecules like sterols and hopanoids, membrane lipids found in eukaryotes and bacteria, respectively, can be preserved in the rock record on billion-year timescales. Following the death of the organism they came from and sedimentation, they undergo a process called diagenesis whereby many of the specific functional groups from the lipids are lost, but the hydrocarbon skeleton remains intact. These fossilized lipids are called steranes and hopanes, respectively.^[61] There are also other types of molecular fossils, like porphyrins, the discovery of which in petroleum by Alfred E. Treibs actually led to the invention of the field.^[8] Other aspects of geochemistry that are also pertinent to geobiology include isotope geochemistry, in which scientists search for isotope fractionation in the rock record, and the chemical analysis of biominerals, such as magnetite or microbially-precipitated gold.

Ediacaran fossils from Mistaken Point, Newfoundland. Ediacaran biota originated during the Ediacaran Period and are unlike most animals around today.

Paleontology

Perhaps the oldest of the bunch, paleontology is the study of fossils. It involves the discovery, excavation, dating, and paleoecological understanding of any type of fossil, microbial or dinosaur, trace or body fossil. Micropaleontology is particularly relevant to geobiology. Putative bacterial microfossils and ancient stromatolites are used as evidence for the rise of metabolisms such as oxygenic photosynthesis.^[62] The search for molecular fossils, such as lipid biomarkers like steranes and hopanes, has also played an important role in geobiology and organic geochemistry.^[8] Relevant sub-disciples include paleoecology and paleobiogeoraphy.

Biogeography

Biogeography is the study of the geographic distribution of life through time. It can look at the present distribution of organisms across continents or between microniches, or the distribution of organisms through time, or in the past, which is called paleobiogeography.

Evolutionary biology

Evolutionary biology is the study of the evolutionary processes that have shaped the diversity of life on Earth. It incorporates genetics, ecology, biogeography, and paleontology to analyze topics including natural selection, variance, adaptation, divergence, genetic drift, and speciation.

Ecohydrology

Ecohydrology is an interdisciplinary field studying the interactions between water and ecosystems. Stable isotopes of water are sometimes used as tracers of water sources and flow paths between the physical environment and the biosphere.