Trace metal stable isotope biogeochemistry is the study of the distribution and relative abundances of trace metal isotopes in order to better understand the biological, geological, and chemical processes occurring in an environment. Trace metals are elements such as iron, magnesium, copper, and zinc that occur at low levels in the environment. Trace metals are critically important in biology and are involved in many processes that allow organisms to grow and generate energy. In addition, trace metals are constituents of numerous rocks and minerals, thus serving as an important component of the geosphere. Both stable and radioactive isotopes of trace metals exist, but this article focuses on those that are stable. Isotopic variations of trace metals in samples are used as isotopic fingerprints to elucidate the processes occurring in an environment and answer questions relating to biology, geochemistry, and medicine.
Isotope notation In order to study trace metal stable isotope biogeochemistry, it is necessary to compare the relative abundances of isotopes of trace metals in a given biological, geological, or chemical pool to a standard (discussed individually for each isotope system below) and monitor how those relative abundances change as a result of various biogeochemical processes. Conventional notations used to mathematically describe isotope abundances, as exemplified here for 56Fe, include the isotope ratio (56R), fractional abundance (56F) and delta notation (δ56Fe). Furthermore, as different biogeochemical processes vary the relative abundances of the isotopes of a given trace metal, different reaction pools or substances will become enriched or depleted in specific isotopes. This partial separation of isotopes between different pools is termed isotope fractionation, and is mathematically described by fractionation factors α or ε (which express the difference in isotope ratio between two pools), or by "cap delta" (Δ; the difference between two δ values). For a more complete description of these notations, see the isotope notation section in Hydrogen isotope biogeochemistry.
Naturally occurring trace metal isotope variations and fractionations In nature, variations in isotopic ratios of trace metals on the order of a few tenths to several ‰ are observed within and across diverse environments spanning the geosphere, hydrosphere and biosphere. A complete understanding of all processes that fractionate trace metal isotopes is presently lacking, but in general, isotopes of trace metals are fractionated during various chemical and biological processes due to kinetic and equilibrium isotope effects.
Geochemical fractionations Certain isotopes of trace metals are preferentially oxidized or reduced; thus, transitions between redox species of the metal ions (e.g., Fe2+ → Fe3+) are fractionating, resulting in different isotopic compositions between the different redox pools in the environment. Additionally, at high temperatures, metals ions can evaporate (and subsequently condense upon cooling), and the relative differences in isotope masses of a given heavy metal leads to fractionation during these evaporation and condensation processes. Diffusion of isotopes through a solution or material can also result in fractionations, as the lighter mass isotopes are able to diffuse at a faster rate. Additionally, isotopes can have slight variations in their solubility and other chemical and physical properties, which can also drive fractionation.
Biological fractionations In sediments, oceans, and rivers, distinct trace metal isotope ratios exist due to biological processes such as metal ion uptake and abiotic processes such as adsorption to particulate matter that preferentially remove certain isotopes. The trace metal isotopic composition of a given organism results from a combination of the isotopic compositions of source material (i.e., food and water) and any fractionations imparted during metal ion uptake, translocation and processing inside cells.
Applications of trace metal isotope ratios Stable isotope ratios of trace metals can be used to answer a variety of questions spanning diverse fields, including oceanography, geochemistry, biology, medicine, anthropology and astronomy. In addition to their modern applications, trace metal isotopic compositions can provide insight into ancient biogeochemical processes operated on Earth. These signatures arise because the processes that form and modify samples are recorded in the trace metal isotopic compositions of the samples. By analyzing and understanding trace metal isotopic compositions in biological, chemical or geological materials, one can answer questions such as the sources of nutrients for phytoplankton in the ocean, processes that drove the formation of geologic structures, the diets of modern or ancient organisms, and accretionary processes that took place in the early Solar System. Trace metal stable isotope biogeochemistry is still an emerging field, yet each trace metal isotope system has clear, powerful applications to diverse and important questions. Important heavy metal isotope systems are discussed (in order of increasing atomic mass) in the proceeding sections.
Iron
Stable isotopes and natural abundances Naturally occurring iron has four stable isotopes, 54Fe, 56Fe, 57Fe, and 58Fe. +Stable isotopes of ironIsotopeAbundance (%)54Fe5.84556Fe91.75457Fe2.119158Fe0.2819 Stable iron isotopes are described as the relative abundance of each of the stable isotopes with respect to 54Fe. The standard for iron is elemental iron, IRMM-014, and it is distributed by the Institute for Reference Materials and Measurement. The delta value is compared to this standard, and is defined as:
Delta values are often reported as per mil values (‰), or part-per-thousand differences from the standard. Iron isotopic fractionation is also commonly described in units of per mil per atomic mass unit.
In many cases, the δ56Fe value can be related to the δ57Fe and δ58Fe values through mass-dependent fractionation:
Chemistry One of the most prevalent features of iron chemistry is its redox chemistry. Iron has three oxidation states: metallic iron (Fe0), ferrous iron (Fe2+), and ferric iron (Fe3+). Ferrous iron is the reduced form of iron, and ferric iron is the oxidized form of iron. In the presence of oxygen, ferrous iron is oxidized to ferric iron, thus ferric iron is the dominant redox state of iron at Earth's surface conditions. However, ferrous iron is the dominant redox state below the surface at depth. Because of this redox chemistry, iron can act as either an electron donor or receptor, making it a metabolically useful species.
Each form of iron has a specific distribution of electrons (i.e., electron configuration), tabulated below: +Electron configurations of FeFe formElectron configurationFe0[Ar]3d64s2Fe2+[Ar]3d6Fe3+[Ar]3d5
Equilibrium Isotope Fractionation Variations in iron isotopes are caused by a number of chemical processes which result in the preferential incorporation of certain isotopes of iron into certain phases. Many of the chemical processes which fractionate iron are not well understood and are still being studied. The most well-documented chemical processes which fractionate iron isotopes relate to its redox chemistry, the evaporation and condensation of iron, and the diffusion of dissolved iron through systems. These processes are described in more detail below.
Fractionation as a result of redox chemistry To first order, reduced iron favors isotopically light iron and oxidized iron favors isotopically heavy iron. This effect has been studied in regards to the abiotic oxidation of Fe2+ to Fe3+, which results in fractionation. The mineral ferrihydrite, which forms in acidic aquatic conditions, is precipitated via the oxidation of aqueous Fe2+ to Fe3+. Precipitated ferrihydrite has been found to be enriched in the heavy isotopes by 0.45‰ per atomic mass unit with respect to the starting material. This indicates that heavier iron isotopes are preferentially precipitated as a result of oxidizing processes.
Theoretical calculations in combination with experimental data have also aimed to quantify the fractionation between Fe(III)aq and Fe(II)aq in HCl. Based on modeling, the fractionation factor between the two species is temperature dependent:
Fractionation as a result of evaporation and condensation Evaporation and condensation can give rise to both kinetic and equilibrium isotope effects. While equilibrium mass fractionation is present evaporation and condensation, it is negligible compared to kinetic effects. During condensation, the condensate is enriched in the light isotope, whereas in evaporation, the gas phase is enriched in the light isotope. Using the kinetic theory of gases, for 56Fe/54Fe, a fractionation factor of α = 1.01835 for the evaporation of a pool containing equimolar amounts of 56Fe and 54Fe. In evaporation experiments, the evaporation of FeO at 1,823 K gave a fractionation factor of α = 1.01877. Presently, there have been no experimental attempts to determine the 56Fe/54Fe fractionation factor of condensation.
Fractionation as a result of diffusion Kinetic fractionation of dissolved iron occurs as a result of diffusion. When isotopes diffuse, the lower mass isotopes diffuse more quickly than the heavier isotopes, resulting in fractionation. This difference in diffusion rates has been approximated as:
In this equation, D1 and D2 are the diffusivities of the isotopes, m1 and m2 are the masses of the isotopes, and β, which can vary between 0 and 0.5, depending on the system. More work is required to fully understand fractionation as a result of diffusion, studies of diffusion of iron on metal have consistently given β values of approximately 0.25. Iron diffusion between silicate melts and basaltic/rhyolitic melts have given lower β values (~0.030). In aqueous environments, a β value of 0.0025 has been obtained.
Fractionation as a result of phase partitioning There may be equilibrium fractionation between coexisting minerals. This would be particularly relevant when considering the formation of planetary bodies early in the Solar System. Experiments have aimed to simulate the formation of the Earth at high temperatures using a platinum-iron alloy and an analog for the silicate earth at 1,500 °C. However, the observed fractionation was very small, less than 0.2‰ per atomic mass unit. More experimental work is needed to fully understand this effect.
Biology In biology, iron plays a number of roles. Iron is widespread in most living organisms and is essential for their function. In microbes, iron redox chemistry is utilized as an electron donor or receptor in microbial metabolism, allowing microbes to generate energy. In the oceans, iron is essential for the growth and survival of phytoplankton, which use iron to fix nitrogen. Iron is also important in plants, given that they need iron to transfer electrons during photosynthesis. Finally, in animals, iron plays many roles, however, its most essential function is to transport oxygen in the bloodstream throughout the body. Thus, iron undergoes many biological processes, each of which have variations in which isotopes of iron they preferentially use. While iron isotopic fractionations are observed in many organisms, they are still not well understood. Improvements in the understanding the iron isotope fractionations observed in biology will enable the development of a more complete knowledge of the enzymatic, metabolic, and other biologic pathways in different organisms. Below, the known iron isotopic variations for different classes of organisms are described.
Iron reducing bacteria