The element barn - Geochemistry: An Introduction - page 292

Atomic weight: 55.847
Condensation temperature: 1336 K
Complexes in water: hydroxides, chlorides
Reactions limiting solubility in water:

Fe(OH) 2 OH − ( Fe(OH) 3 ⇔ + − log K = 16.5)

Fe(OH) 4 OH − ( Fe(OH) 3 + H 2 O ⇔ + − log K = 4.4)

Residence time in seawater: 55 years

Iron is the most abundant element in the Earth. It is refractory and, by definition,
siderophile. It is the most abundant element of the core, both in the solid inner core and the
liquid outer core in which convection generates the terrestrial magnetic field. In contrast
with the core where Fe occurs in its metallic and most reduced form Fe 0 , iron in the mantle
is essentially in its Fe 2 + form. Ferrous iron (Fe 2 + ) substitutes for Mg 2 + in most silicate
mineral phases. Iron is, after Si and Mg, the third most abundant cation in the mantle. In
the upper mantle, ferrous iron is found in olivine, pyroxene, garnet, and amphibole. In the
deep mantle, it enters with Mg into the perovskite structure of ringwoodite and also into
the oxide structure of magnesio-wüstite (Fe, Mg) O. In igneous rocks, as in the crust in
general, it is hosted in amphibole and biotite and also, together with Fe 3 + ,Al 3 + ,Cr 3 + , and
Ti 4 + , in oxide minerals (magnetite, ilmenite). Ferric iron easily substitutes into the tetra-
hedral site of alkali feldspars, which is why so many granites turn reddish upon incipient
weathering. When exposed to the atmosphere or seawater at low temperature, Fe is nor-
mally oxidized to Fe 3 + . It is found in different forms of iron hydroxide (such as goethite,
hematite, and limonite) that dominate soils, sediments, as well as ferromanganese nodules
and encrustations from the deep sea. Iron-rich clay minerals and carbonates are uncom-
mon. Organic compounds contain important Fe-rich proteins that have different functions,
notably oxygen transport in the cell (porphyrins). Iron concentration in seawater is very
low, again because of the very low solubility of hydroxides.
The sites occupied by Fe 2 + and Fe 3 + are normally octahedral, but Fe 3 + can be found in
tetrahedral sites, especially in feldspars. During mantle melting, Fe 2 + has a neutral behav-
ior (neither compatible nor incompatible) owing to its lack of octahedral-site preference
energy (see Chapter 1 ). In contrast, Fe 3 + is highly incompatible. Silicates are very poor
electrical conductors so magmas cannot exchange significant amounts of electrons with
country rocks. Upon removal of ferrous iron into cumulate minerals, the Fe 3 + /Fe 2 + ratio
therefore increases in residual melts. The apparently more oxidizing conditions of differ-
entiated rocks are therefore not the result of an externally imposed higher “fugacity” of
oxygen, but simply reflect the increasing electron deficit in smaller and smaller quantities
of melt. Most basalts would contain 10-12 wt % total iron as FeO and granites 3-4 wt %.
Typically, about 15 wt % of the iron present in a primary basalt is in the form of Fe 3 + .
Even when small quantities of ilmenite and magnetite are present at the liquidus, iron con-
centration in melts does not change when the fractionating assemblage is dominated by
olivine and pyroxene: this is the case for the high-pressure differentiation of ocean island
basalts and for the wet differentiation of orogenic (calc-alkaline) magmas. When plagio-
clase is present, as in mid-ocean ridge and continental flood basalts (see Section 11.1 ) ,