organic matter because of the increased level of conservation). The issue of the oxygen balance during the Paleoproterozoic is still debated to this day, and in particular, the
possible role a decrease in the quantity of reduced gas
degassed by the mantle is questioned (see references in
Catling and Claire 2005). However, there is no doubt about
the accumulation of O 2 between 2.3 and 2.1 Ga (Catling and
Claire 2005). The GOE is also recorded in the variations in
the isotopic composition of iron in sedimentary pyrites
between 2.6 and 1.8 Ga, which is suspected to be related to
the isotopic composition of seawater (Fig. 26.6). The d
56 Fe
of these sediments shows widely dispersed values before
2.3 Ga, but this dispersion decreases considerably after
2.3 Ga. These changes are explained by the existence of an
iron rich ocean, fed by hydrothermal springs (containing
dissolved iron Fe
2 + ) before 2.3 Ga. This ocean periodically
purges itself of a certain amount of its iron during the precipitation of iron oxides during upwelling episodes. Such
purges brought water to more or less oxygenated zones. This
precipitation extracts the isotope 56 preferentially from the
iron; this process of fractional precipitation of iron is thus
capable of modifying the d
56 Fe isotopic composition of the
deep ocean so long as all the iron carried by the upwellings
has not precipitated and a part of it returns to the deep ocean.
After 2.3 Ga and the GOE, the Proterozoic ocean became
stratified, characterized by a permanently and thoroughly
oxygenated surface zone and an anoxic deep ocean. This
configuration allows the precipitation of all of the
hydrothermal Fe
2 + brought by the upwellings (see references in Catling and Claire 2005). In this way, the isotopic
composition of the deep ocean is no longer affected, since all
the iron carried by upwellings precipitated, and stabilized at
between 0 and 1‰. Finally, recent data using the independent fractionation of the mass of isotopes of sulfur D
33 S
confirm progressive oxidation of the surface layers from 2.4
to 2.3 Ga onwards (Papineau et al. 2007).
The consequences for the GOE climate were significant.
The residence time of methane in the atmosphere decreased
strongly and the concentration of CH 4 in the atmosphere
probably fell to around 300–100 ppmv, a decrease by a
factor of 5–16 compared to Archean values. This caused
significant cooling, which could not be immediately compensated for by the low level of CO 2 . This triggered the
Huronian glaciations during the time it took for the paleothermostat to allow the partial pressure of CO 2 to rise. This
major glacial phase may have included at least one episode
of total glaciation, but paleomagnetic data on the position of
the continental masses remain scanty and difficult to interpret. However, recent paleomagnetic studies show that the
Huron glaciation was probably similar to the ‘standard’
glaciations of the Quaternary. However, the GOE context is
completely different in terms of atmospheric composition,
paleogeography, solar insolation… and therefore the rhythm
of this glaciation is still an open question.
The Proterozoic
Following the Huronian glaciations, the Earth seems to have
been subjected to a warm climate persisting over most of the
Proterozoic, lasting for about 1 billion years. To date, no
trace of glaciation has been found in the timespan 2–0.8 Ga.
The measurement of d
13 C from acrytarches (microfossil
remains of cysts of photosynthetic eukaryotic organisms)
allows a rough estimation of the atmospheric CO 2 level.
Indeed, the difference between these ratios and the mean
ratio of oceanic carbonates (close to 0‰) defines, with some
assumptions, the isotopic fractionation involved in photosynthesis that occurred during the Calvin cycle. This fractionation depends, among other things, on the CO 2 pressure
of the water and can therefore be linked to atmospheric CO 2
pressure. At 1.4 Ga, the level of CO 2 was about 10–200
Fig. 26.6 The d
56
Fe of
Precambrian sediments
350
Y. Goddéris et al.
possible role a decrease in the quantity of reduced gas
degassed by the mantle is questioned (see references in
Catling and Claire 2005). However, there is no doubt about
the accumulation of O 2 between 2.3 and 2.1 Ga (Catling and
Claire 2005). The GOE is also recorded in the variations in
the isotopic composition of iron in sedimentary pyrites
between 2.6 and 1.8 Ga, which is suspected to be related to
the isotopic composition of seawater (Fig. 26.6). The d
56 Fe
of these sediments shows widely dispersed values before
2.3 Ga, but this dispersion decreases considerably after
2.3 Ga. These changes are explained by the existence of an
iron rich ocean, fed by hydrothermal springs (containing
dissolved iron Fe
2 + ) before 2.3 Ga. This ocean periodically
purges itself of a certain amount of its iron during the precipitation of iron oxides during upwelling episodes. Such
purges brought water to more or less oxygenated zones. This
precipitation extracts the isotope 56 preferentially from the
iron; this process of fractional precipitation of iron is thus
capable of modifying the d
56 Fe isotopic composition of the
deep ocean so long as all the iron carried by the upwellings
has not precipitated and a part of it returns to the deep ocean.
After 2.3 Ga and the GOE, the Proterozoic ocean became
stratified, characterized by a permanently and thoroughly
oxygenated surface zone and an anoxic deep ocean. This
configuration allows the precipitation of all of the
hydrothermal Fe
2 + brought by the upwellings (see references in Catling and Claire 2005). In this way, the isotopic
composition of the deep ocean is no longer affected, since all
the iron carried by upwellings precipitated, and stabilized at
between 0 and 1‰. Finally, recent data using the independent fractionation of the mass of isotopes of sulfur D
33 S
confirm progressive oxidation of the surface layers from 2.4
to 2.3 Ga onwards (Papineau et al. 2007).
The consequences for the GOE climate were significant.
The residence time of methane in the atmosphere decreased
strongly and the concentration of CH 4 in the atmosphere
probably fell to around 300–100 ppmv, a decrease by a
factor of 5–16 compared to Archean values. This caused
significant cooling, which could not be immediately compensated for by the low level of CO 2 . This triggered the
Huronian glaciations during the time it took for the paleothermostat to allow the partial pressure of CO 2 to rise. This
major glacial phase may have included at least one episode
of total glaciation, but paleomagnetic data on the position of
the continental masses remain scanty and difficult to interpret. However, recent paleomagnetic studies show that the
Huron glaciation was probably similar to the ‘standard’
glaciations of the Quaternary. However, the GOE context is
completely different in terms of atmospheric composition,
paleogeography, solar insolation… and therefore the rhythm
of this glaciation is still an open question.
The Proterozoic
Following the Huronian glaciations, the Earth seems to have
been subjected to a warm climate persisting over most of the
Proterozoic, lasting for about 1 billion years. To date, no
trace of glaciation has been found in the timespan 2–0.8 Ga.
The measurement of d
13 C from acrytarches (microfossil
remains of cysts of photosynthetic eukaryotic organisms)
allows a rough estimation of the atmospheric CO 2 level.
Indeed, the difference between these ratios and the mean
ratio of oceanic carbonates (close to 0‰) defines, with some
assumptions, the isotopic fractionation involved in photosynthesis that occurred during the Calvin cycle. This fractionation depends, among other things, on the CO 2 pressure
of the water and can therefore be linked to atmospheric CO 2
pressure. At 1.4 Ga, the level of CO 2 was about 10–200
Fig. 26.6 The d
56
Fe of
Precambrian sediments
350
Y. Goddéris et al.