The main source for iron and steel has for more than half a century been Precambrian rock characterised by intricate interlayering of silica- and iron oxide-rich sediments known as banded iron formations or BIFs. They always appear in what were shallow-water parts of Precambrian sedimentary basins. Although much the same kind of material turns up in sequences from 3.8 to 0.6 Ga, by far the largest accumulations date from 2.6 to 1.8 Ga, epitomised by the vast BIFs of the Palaeoproterozoic Hamersley Basin in Western Australia. This peak of iron-ore deposition brackets the time (~2.4 Ga) when world-wide evidence suggests that the Earth’s atmosphere first acquired tangible amounts of free oxygen: the so-called ‘Great Oxidation Event’. Yet the preservation of such enormous amounts of oxidised iron compounds in BIFs is paradoxical for two reasons: the amount of freely available atmospheric oxygen at their acme was far lower than today; had the oceans contained much oxygen, dissolved ions of reduced Fe-2 would not have been able to pervade seawater as they had to for BIFs to have accumulated in shallow water. Iron-rich ocean water demands that its chemical state was highly reducing.
Oblique view of an open pit mine in banded iron formation at Mount Tom Price, Hamersley region Western Australia (Credit Google earth)
The paradox of highly oxidised sediments being deposited when oceans were highly reduced was resolved, or seemed to have been, in the late 20th century. It involved a hypothesis that reduced, Fe-rich water entered shallow, restricted basins where photosynthetic organisms – probably cyanobacteria – produced localised enrichments in dissolved oxygen so that the iron precipitated to form BIFs. Later work revealed oddities that seemed to suggest some direct role for the organisms themselves, a contradictory role for the co-dominant silica-rich cherty layers and even that another kind of bacteria that does not produce oxygen directly may have deposited oxidised iron minerals. Much of the research focussed on the Hamersley BIF deposits, and it comes as no surprise that another twist in the BIF saga has recently emerged from the same, enormous repository of evidence (Rasmussen, B. et al. 2015. Precipitation of iron silicate nanoparticles in early Precambrian oceans marks Earth’s first iron age. Geology, v. 43, p. 303-306).
The cherty laminations have received a great deal less attention than the iron oxides. It turns out that they are heaving with minute particles of iron silicate. These are mainly the minerals stilpnomelane [K(Fe,Mg)8(Si, Al)12(O, OH)27] and greenalite [(Fe)2–3Si2O5(OH)4] that account for up to 10% of the chert. They suggest that ferruginous, silica-enriched seawater continually precipitated a mixture of iron silicate and silica, with cyclical increases in the amount of iron-silicate. Being such a tiny size the nanoparticles would have had a very high surface area relative to their mass and would therefore have been highly reactive. The authors suggest that the present mineralogy of BIFs, which includes iron carbonates and, in some cases, sulfides as well as oxides may have resulted from post-depositional mineral reactions. Much the same features occur in 3.46 Ga Archaean BIFs at Marble Bar in Western Australia that are almost a billion years older that the Hamersley deposits, suggesting that a direct biological role in BIF formation may not have been necessary.
More on BIFs and the Great Oxidation Event
Cyanobacteria: earliest producers of oxygen in the Precambrian. Image via Wikipedia
The entire eukaryote domain of life, from alga to trees and fungi to animals, would not exist had it not been for the emergence of free oxygen in the oceans and atmosphere about 2.4 billion years ago; thanks in large part to the very much simpler photosynthetic blue-green bacteria. The chemistry behind this boils down to organisms being able to transfer electrons from elements and compounds in the inorganic world to build organic molecules incorporated in living things. Having lost electrons the inorganic donors become oxidised, for instance ferrous iron (Fe2+ or Fe-2) becomes ferric iron (Fe3+ or Fe-3) and sulfide ions (S2-) become sulfate (SO42-) and the organic products that receive electrons principally involve reduction of carbon, on the OilRig principal – Oxidation involves loss of electrons, Reduction involves gain. Since the Great Oxygenation Event (GOE), ferric iron and sulfate ions now account for 75% of oxidation of the lithosphere and hydrosphere while free oxygen (O2) is a mere 2-3 % (Hayes, J.M. 2011. Earth’s redox history. Science. V. 334, p. 1654-1655; an excellent introduction to the geochemistry involved in the GOE and the carbon cycle). Free oxygen is around today only because more of it is produced than is consumed by its acting to oxidize ferrous iron and sulfide ions supplied mainly by volcanism, and carbon-rich material exposed to surface processes by erosion and sediment transport.
Eukaryote life has never been snuffed out for the last two billion years or so, but it has certainly had its ups and downs. To geochemists taking the long view oxygen might well seem to have steadily risen, but that is hardly likely in the hugely varied chemical factory that constitutes Earth’s surface environments, involving major geochemical cycles for carbon, iron, sulfur, nitrogen, phosphorus and so on, that all inveigle oxygen into reactions. Tabs can be kept on one of these cycles – that involving carbon – through the way in which the proportions of its stable isotopes vary in natural systems. If all geochemistry was in balance all the time, all materials that contain carbon would show the same proportions of 13C and 12C as the whole Earth, but that is never the case. Living processes that fix carbon in organic compounds favour the lighter isotope, so they show a deficit of 13C relative to 12C signified by negative values of δ13C. The source of the carbon, for instance CO2 dissolved in sea water, thereby becomes enriched in 13C to achieve a positive value of δ13C, which may then be preserved in the form of carbonates in, for instance, fossil shells that ended up in limestones formed at the same time as organic processes were favouring the lighter isotope of carbon. Any organic carbon compounds that ocean-floor mud buried before they decayed (became oxidised) conversely would add their negative δ13C to the sediment. Searching for δ13C anomalies in limestones and carbonaceous mudrocks has become a major means of charting life’s ups and downs, and also what has happened to buried organic carbon through geological time.
A most interesting time to examine C-isotopes and the carbon cycle is undoubtedly the period immediately following the GOE, in the Palaeoproterozoic Era (2500 to 1600 Ma). From around 2200 to 2060 Ma the general picture is roughly constant, high positive values of δ13C (~+10‰): more organic carbon was being buried than was being oxidised to CO2. However, in drill cores through the Palaeoproterozoic of NW Russia carbonate carbon undergoes a sharp decline in its heavy isotope to give a negative δ13C (~-14‰) while carbon in organic-rich sediments falls too (to~-40‰): definitely against the general trend (Kump, L.R. et al. 2011. Isotopic evidence for massive oxidation of organic matter following the Great Oxidation Event. Science. V. 334, p. 1694-1696). Oxygen isotopes in the carbonates affected by the depletion in ‘heavy’ carbon show barely a flicker of change: a clear sign that the 13C δ13C deficit is not due to later alteration by hydrothermal fluids, as can sometimes cause deviant δ13C in limestones. It is more likely that a vast amount of organic carbon, buried in sediments or dissolved in seawater was oxidised to CO2 faster than biological activity was supplying dead material to be buried or dissolved. In turn, the overproduction of carbon dioxide dissolved in seawater to affect C-isotopes in limestones. Such an event would have entailed a sharp increase in oxygen production to levels capable of causing the oxidation (~ 1% of present levels). Yet this was not the time of the GOE (2400 Ma) but 300-400 Ma later. A possible explanation is a burst in oxygen production by more photosynthetic activity, perhaps by the evolution of chloroplast-bearing eukaryotes much larger than cyanobacteria.