113.08 The Great Oxidation Event Transition
David C. Catling
University of Washington, Seattle, WA 98195
13.08.1 INTRODUCTION
13.08.2 CONTROLS ON O2 LEVELS

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3

13.08.2.1 Hydrogen escape to space and oxidation
13.08.2.2 The net source flux of oxygen
13.08.2.3 The oxygen sink fluxes
13.08.2.4 General principles of oxygen change
13.08.2.5 The definition of ‘anoxic’ versus ‘oxic’ atmospheres

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6
8
10
10

13.08.3 ATMOSPHERIC CHEMISTRY THROUGH THE GREAT OXIDATIONEVENT (GOE)11
13.08.3.1 Constraints on atmospheric gases before the GOE
13.08.3.2 Constraints on atmospheric gases after the GOE

1215

13.08.4 EXPLAINING THE RISE OF O2

17

13.08.4.1 Hypotheses for an increasing source of O2
13.08.4.2 Hypotheses for an decreasing sink of O2

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13.08.5 CHANGES IN ATMOSPHERIC CHEMISTRY AND CLIMATEASSOCIATED WITH THE RISE OF O2

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13.08.5.1 The collapse of methane at the GOE
13.08.5.2 The formation of an ozone shield

23
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13.08.6 CONCLUSIONS

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Keywords: oxygenation, Great Oxidation Event, rise of oxygen, redox, methane

Synopsis
The Great Oxidation Event (GOE) was an increase in the concentration of
atmospheric oxygen from less than 1 part per million to 0.2-2% by volume some 2.4
billion years ago. In atmospheric chemistry, hydrogen-bearing reduced gases, such as
methane and hydrogen, are adversaries of O2. When the concentration of one goes up,
the other declines. The redox chemistry of the pre-GOE atmosphere should have been
dominated by methane and hydrogen. Before the GOE, oxygen was driven to tracelevels by reaction with volcanic and metamorphic reductants, including dissolved
cations (e.g., Fe2+) in surface waters and reducing gases (H2, CH4, CO, SO2 and H2S).
A tipping point”was reached when the flux of O2 associated with the burial of organiccarbon exceeded such O2 losses. Oxidative weathering then became significant.
Models suggest that a decline in the supply of reducing gases caused methane levels
to fall even before the GOE. Loss of greenhouse gases such as methane and ethaneplausibly caused global cooling. Multiple glaciations during 2.45-2.22 Ga suggest thatthe global climate and atmospheric composition oscillated until permanent
oxygenation was established. Subsequent levels of O2 were sufficient to protect the
Earth’s surface from harmful ultraviolet with a stratospheric ozone layer.

13.08.1 INTRODUCTION
In Chapter 13.05, Farquhar et al. described geologic evidence for a significant
increase in atmospheric O2 levels at 2.45-2.32 Ga, and in this chapter, I discuss why
such a transition occurred. This chapter revolves around a basic principle of2
atmospheric chemistry, namely, that the concentration of an atmospheric gas iscontrolled by a kinetic competition between production and loss. This competition is
described by fluxes, which, for a particular gas, are the rate of input to the atmosphereand the rate of chemical consumption. So, in its simplest formulation, there are only
two options for explaining the Great Oxidation Event (GOE): either an increase of theO2 source flux or a decrease of the O2 sink flux created an imbalance. Most
hypotheses in the literature argue that such changes in fluxes were probably gradual
but reached a “tipping point” when the O2 source flux exceeded the consumption flux
from reducing gases, allowing O2 to rise to a new equilibrium level. The higher, postGOE O2 concentration was clearly determined by a new and different balance offluxes. We can appreciate what controlled this new equilibrium with the naked eye.
The ~2.3 Ga appearance of red-beds, which are continental sediments stained by redcolored iron oxides, shows that atmospheric O2 was lost in a quantitatively significant
way to oxidative weathering after the GOE but not before. The oxidative weathering
flux must have been very small prior to the GOE given the absence of oxidation ofcontinental surfaces in the Archean. To understand the change at the GOE, we start bydescribing the source and sink fluxes of O2 as they are known today.
After describing general controls on O2, we discuss the other redox-sensitive
gases in the atmosphere before and after the GOE. In the redox chemistry of theatmosphere, hydrogen-bearing reduced gases, such CH4 and H2, are the chemical
adversaries of O2. Through a series of photochemical reactions, there is a rapid mutual
annihilation between O2 and the reducing gases, akin to combustion in its net effect.
In general, the redox chemistry of the pre-GOE atmosphere is dominated by methaneand hydrogen molecules that spend an average of ~104 years in the atmosphere before
being destroyed by reactions, whereas O2 behaves as a trace gas with a short lifetimein the Archean atmosphere of less than a day (Pavlov et al., 2001). This situation in
the Archean is the opposite of the modern atmosphere where methane and hydrogenare short-lived molecules with average lifetimes with respect to photochemical
destruction of ~10 and ~2 years (e.g., Warneck, 2000), while O2 has a lifetime ofabout 2 million years with respect to geologic sinks (Sec. 13.08.2.4, below). Possibly,
reduced biogenic sulfur gases were also important constituents of the late Archean
and pre-GOE Paleoproterozoic atmosphere. Models suggest that the GOE occurred
when the rate of supply of O2 to the atmosphere rose above the rate of supply of
reduced gases to the atmosphere (Claire et al., 2006; Goldblatt et al., 2006; Holland,2009). Because the atmosphere has a relatively small mass compared to the size oftypical fluxes, the oxic transition probably happened relatively quickly from ageological perspective. In photochemical simulations, the timescale is about 104 years(Claire, 2008), which is instantaneous in geologic time.
Finally, we discuss the implications of the GOE for episodes of climate
cooling and establishing an ozone layer. In detailed models, there is a decline in the
supply of reducing gases prior to the GOE, which means that the methane
concentration in the pre-GOE atmosphere should have fallen to a low level before O2
rose to a substantial concentration (Zahnle et al., 2006). The collapse in theconcentration of methane—a greenhouse gas, along with its photochemical product,ethane—is generally consistent with low latitude glaciation (“Snowball Earth”)
observed in the geologic record at ~2.4 Ga. In some geologic sections, such as the
Huronian of Canada, evidence for three glaciations appears between 2.45-2.22 Ga.
Some biogeochemical models are able to produce oscillations in climate andatmospheric chemistry before establishing a permanent rise of O2 (Claire et al., 2006).
In these models, oxygenation events fail and the atmosphere reverts back to being

3
weakly reducing before the Earth system attempts oxygenation again. Such behavior
in models is entirely dependent upon how the output of biogenic, redox-sensitive
gases is assumed to respond to the climate cooling associated the GOE.
Ozone (O3) is chemically derived from molecular oxygen. Once oxygenation
was permanent, even with the relatively low levels of 0.2-2% O2 by volume thatexisted in the middle Proterozoic, there was sufficient O2 to establish a stratosphericozone layer (Kasting and Donahue, 1980). The ozone layer would have protected theEarth’s surface from harmful ultraviolet in the range 200-300 nm. A counter-intuitiveeffect of the ozone layer was to decrease the net reaction rate in the tropospherebetween methane and oxygen by decreasing the input of ultraviolet photons. So, if
methane fluxes were still relatively large from poorly oxygenated and euxinic marine
waters in the Proterozoic, methane could have risen after the GOE to be a moderately
significant greenhouse gas, although not to such large concentrations as in theArchean.

13.08.2 CONTROLS ON O 2 LEVELS
13.08.2.1 Hydrogen escape to space and oxidation
There are both abiotic and biological net sources of O2 for the atmosphere, but
the latter source is dominant today by a factor of hundreds. Although the current
abiotic source of O2 is trivial, it is determined by the magnitude of the escape rate ofhydrogen to space, which is worth discussing in some detail because this flux was
almost certainly much bigger before the GOE. An anoxic atmosphere is relatively richin hydrogen-bearing reducing gases (such as methane, CH4, and hydrogen, H2), which
can drive relatively high rates of hydrogen escape. Hydrogen is the lightest gas andplanets the size of the Earth or smaller cannot hold on to hydrogen and undergo
cumulative oxidation. Today we observe the process of hydrogen escape on Saturn’s
moon, Titan, which has ~5% CH4 in its lower atmosphere. We can also observe a haloof hydrogen around the modern Earth directly, some of which are escaping. The
hydrogen atoms in the halo resonantly scatter Lyman-alpha (121.6 nm) ultraviolet
radiation from the Sun and have been imaged by spacecraft (Fig. 1). Today, thehydrogen escape flux is very small. But it helps to understand the hydrogen escapeprocess in order to appreciate its potential oxidizing effect on Earth’s earlyatmosphere when the flux of hydrogen escape was surely larger.
While hydrogen can escape to space, oxygen tends to remain behind because itis heavier. Indeed, it is universally accepted that the escape of hydrogen relative toheavier oxygen is responsible for the oxidized state of the atmospheres of Mars
(Hunten, 1979), Venus (Watson et al., 1984; Zahnle and Kasting, 1986), and theextremely tenuous O2-dominated atmospheres of several outer planet moons. The
latter includes Ganymede (Hall et al., 1998) and Europa (Smyth and Marconi, 2006),
which are satellites of Jupiter, and Rhea (Teolis et al., 2010), which is a moon of
Saturn. Some theories of oxygenation of the Earth give hydrogen escape a similarly
important role in oxidizing the Earth and enabling the build-up of photosynthetic
oxygen (Catling et al., 2001; Kasting et al., 1993). While hydrogen escape is theaccepted explanation for the oxidation state of atmospheres of other planets andsatellites, its role in the oxidation state of the Earth’s atmosphere remains an area of
debate, however.

4
The photolysis of gases in Earth’s upper atmosphere and the escape ofhydrogen to space provide an oxidation flux (Holland, 1978, p.296). It is easy to see
why. By the rules of redox chemistry, the hydrogen oxidation number decreases from
+1 in atmospheric compounds (such as H2O vapor, CH4 or H2S) to zero in H atoms
that escape to space, which means that hydrogen gets reduced. Reduction is always
accompanied by oxidation, so there must be an equivalent increase in oxidation
number, i.e., oxidation, for other species on Earth when hydrogen escapes,
irrespective of the details of the chemistry. While any particular redox reaction isredox-neutral, the oxidized and reduced components that are generated can beseparated into different reservoirs. For the Earth as a whole, reduction of hydrogen toelemental form followed by hydrogen escape represents partitioning of the reduced
component into space, leaving behind the oxidized component. The oxidation isirreversible because hydrogen that escapes into interplanetary space is lost forever
from the Earth’s gravitational field.
Oxygen-bearing gases can be photolyzed in the upper atmosphere to produce
atomic oxygen (O) that can undergo atmospheric reactions to make abiotic molecularoxygen, O2. For example, H2O and CO2 are photolyzed as follows,

H 2O + h! " H + OH
CO 2 + h! " CO + O(1)

Molecular oxygen, O2, can be produced from the O atoms and hydroxyl (OH) radicals
made from the reactions in eq. (1), as follows:
O + OH ! O 2 + H(2)

If hydrogen does not escape, O2 is lost in catalytic cycles and water vapor is reformed,
as follows:

H 2O + h! " H + OH
H + O 2 + M " HO 2 + M
3#…H 2 + OH " H 2O + H
H + HO 2 " 2OH(3)

H + H + M " H2 + M
2H 2 + O 2 " 2H 2OAlternatively, if hydrogen escapes to space, the products of photolysis cannot
recombine and O2 is leftover. Thus, it is the escape of hydrogen to space that isessential for making O2. So a net oxidation flux from photochemistry is always
determined by the escape flux of hydrogen.
Under a wide variety of circumstances, which apply to the Archean
atmosphere as well as the present one, the escape of hydrogen to space is constrainedby two “bottlenecks” in the atmosphere, shown schematically in Fig. 2 (following
Hunten, 1990). The first bottleneck is the tropopause cold-trap, which limits thetransport of hydrogen within water vapor to the upper atmosphere. The Earth’s
tropopause is the minimum temperature in the vertical temperature profile that occurs

5
at the top of the Earth’s troposphere. The troposphere extends vertically from the
Earth’s surface to ~17 km at the equator or to lower altitudes at other latitudes. The
tropopause acts as a brake on vertical transport in the atmosphere because the energybalance in the atmosphere changes from convection-dominated in the troposphere toradiation-dominated just below the tropopause. Convection is sufficiently strong in
the warm tropics to lift water vapor up and through the tropical tropopause. However,at the tropical tropopause, water vapor freezes into ice particles (hence the name coldtrap), so that very little water vapor enters the lower stratosphere. Consequently, thelower stratosphere is extremely arid with only ~3 ppmv H2O. Other hydrogen-bearing
constituents such as methane (CH4) and diatomic hydrogen (H2) have condensation
temperatures that are far too low to be cold-trapped in Earth’s atmosphere, so thesegases mix through the tropopause and their abundance in the lower stratosphere isabout the same as in the troposphere. Ultimately, stratospheric water vapor is broken
down at high altitudes through photochemical reactions. The resulting hydrogen,along with the hydrogen derived from methane and molecular hydrogen, istransported upwards and mixed by eddies through the stratosphere and mesosphere.
The second bottleneck for hydrogen escape occurs when the upwardly mixedhydrogen reaches the homopause, around 100 km altitude. The homopause is the level
above which gases separate out diffusively with altitude according to their mass.
Consequently, the air above the homopause becomes increasingly enriched inmolecules or atoms that have lower mass than the average. Hydrogen, being lighter
than all other gases, rises buoyantly through the background air by molecular
diffusion. Fortunately, the diffusion rate is readily calculable and sets a diffusionlimited escape rate, which is the maximum rate that hydrogen can be transported
upwards by diffusion and escape (Hunten, 1973). Essentially, diffusion acts asanother brake on the upward transport of hydrogen to space. Hydrogen is mostly
broken down into atomic form and escapes efficiently from the exobase, which is thelevel in the atmosphere, typically around 400-500 km altitude, above which collisions
between atoms and molecules are so infrequent that they can be neglected. This lackof collisions means that there is nothing to prevent hydrogen atoms with upwarddirected velocities exceeding the escape velocity being lost permanently tointerplanetary space from the exobase. So while the tropopause cold-trap modulates
the amount of hydrogen that can reach the upper atmosphere and escape, molecular
diffusion above the homopause sets the maximum rate that hydrogen can escape once
it gets to ~100 km altitude. Hence, the cold-trap and diffusion are said to be the two“bottlenecks” on hydrogen escape (Hunten, 1990).
The theory of diffusion-limited escape gives a hydrogen escape rate, ΦH, asfollows (Walker, 1977):
!H "

bH fT
= (constant) fT # (2.5 $ 1013 ) fT [atoms cm -2 sec-1 ]
H

(4)

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Here, bH is a binary diffusion parameter for hydrogen through air, where one uses anaverage of for H2 and H, weighted by their probable mixing ratios at the homopause
in order to obtain the constant in the second expression. The variable H is theatmospheric scale height for the background gas, given by H = kT/(mg) where k is
Boltzmann’s constant (1.38 × 10-23 J K-1), T is the mean atmospheric temperature at
the homopause, m is mean molecular mass, and g is gravitational acceleration. The
remaining term, fT, is the total hydrogen mixing ratio at the homopause, given by thesum of the mixing ratios of hydrogen in all of its chemical forms, weighted by the
number of hydrogen atoms each species contains, i.e.,fT = fH + 2 fH2 + 2 fH 2O + 4 fCH4 + ...

(5)

Here fH is the mixing ratio of H atoms, fH2 is the mixing ratio of H2 molecules, fH2O isthe mixing ratio of water vapor, and fCH4 is the mixing ratio of methane. Because thetotal hydrogen mixing ratio is approximately conserved with altitude from the lower
stratosphere to the homopause (as expected from theory and observation (Harries etal., 1996)), we can evaluate eq. (5) with values that are well known frommeasurements in the lower stratosphere.
The current hydrogen escape rate and the associated oxidation rate can easily
be calculated using the above theory. In today’s lower stratosphere, the concentrationsof hydrogen-bearing gases are 1.8 ppmv CH4, ~3 ppmv H2O and 0.55 ppmv H2.
Using eq. (5), the total mixing ratio of hydrogen, fT, is 14.3×10-6 (= [2(0.55)+4(1.8) +2(3)]×10-6). Consequently, the hydrogen escape rate from eq. (4) is 3.6 × 108 atoms
cm-2 sec-1. Given that the area of the Earth is 5.1×1018 cm2, the escape rate for thewhole Earth is (3.6×108 atoms cm-2 sec-1) × (5.1×1018 cm2) = 1.8×1027 atoms sec-1 =3 kg H sec-1 = 9.3×1010 mol H atoms yr-1 = 93,000 tonnes of H yr-1. By redox
balance, the escape of four moles of hydrogen atoms is equivalent to an oxidation fluxof one mole of oxygen (which we can deduce through the schematic redox balance of
4H + O2 = 2H2O). Consequently, the net oxidation flux for the whole Earth fromtoday’s escape of hydrogen to space is equivalent to 2.3 ×1010 mol O2 yr-1. This flux
is negligible compared to the biological flux of O2 (~1013 mole O2 yr-1), which we
discuss below. However, the hydrogen escape flux must inevitably have been larger
in the pre-GOE atmosphere, which was redox-dominated by hydrogen-bearing
species. A higher total mixing ratio of hydrogen in all its forms leads to a larger totalhydrogen mixing ratio, fT in eq. (5). Consequently, there is a proportionately greaterhydrogen escape rate, ΦH. For example, the anoxic Archean atmosphere probably had~103 ppmv CH4 (Kasting et al., 2001). In that case, the oxidation flux of hydrogenescape from eqs. (4) and (5) was significant and ~7 ×1012 mol O2 yr-1.

13.08.2.2 The net source flux of oxygen
Molecular oxygen in the modern atmosphere is produced almost entirely as
the waste product of oxygenic photosynthesis—the type of photosynthesis in whichgreen plants, algae and cyanobacteria use sunlight to split water molecules. Moregenerally, photosynthesis is the biological process by which carbon dioxide ischemically reduced with hydrogen with an overall scheme as follows:
2H 2 A + CO 2 + h! " CH 2O + H 2O + 2A (or A 2 )(6)

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Here, A designates an atom that varies depending on the type of photosynthesis and“CH2O” represents organic matter as carbohydrate. It is generally thought thatanoxygenic photosynthesis (which does not release O2) was ancestral to oxygenic
photosynthesis because it is chemically simpler and because the genetics of microbes
implies such a lineage (Xiong, 2006). The suspected evolutionary sequence is that thefirst photosynthetic microbes used H2 (eq. (6) with no ‘A’). These were probably
followed by microbes that used hydrogen sulfide, where “H2A” as H2S in eq. (6):2H 2S + CO 2 + h! " CH 2O + H 2O + 2S

(7)Photosynthesis that uses ferrous iron also evolved. Iron photosynthesis doesn’t quite
fit the generic scheme above but can be written:
4Fe 2+ + CO 2 + 11H 2O + h! " CH 2O + 4Fe(OH)3 + 8H +(8)In oxygenic photosynthesis, water is split. So, in eq. (6) “H2A” is H2O and thewaste product is O2 derived from the O atoms in H2O. The equation representing
oxygenic photosynthesis and the reverse processes of respiration or oxidative decay
can be simplified to:
!!!!!
!!!!!
" CH O + O
H 2O + CO 2 + energy #22
respiration, decay
photosynthesis(9)

In this equation, a mole of organic matter accompanies every mole of O2 generated byphotosynthesis. But within ~102 years, almost all the O2 produced in photosynthesis isused up in oxidizing organic matter back to carbon dioxide in respiration or decay.
Consequently, when we consider geologic timescales of millions of years, we canignore this fast, gross cycle of zero sum photosynthesis and focus on the net
production. In net production, about 0.1-0.2% of the organic carbon is buried, mostly
in deltaic sediments of continental shelves (Berner, 1982), where it escapes the fate ofbeing oxidized. The burial flux of organic carbon is estimated to be 10.0±3.3 Tmol C
yr-1 today (Holland, 2002), which, by the 1:1 stoichiometry of eq. (9), provides
10.0±3.3 Tmol O2 yr-1. (Here and elsewhere, 1 Tmol = 1012 mol). Berner (2004), p.
42, gives a somewhat lower estimate of 5.3 Tmol C yr-1 for the organic burial flux.
For consistency with other estimates, we will follow Holland’s numbers.
In the modern O2 balance, the burial of other reductants can also contribute toO2 production, of which sulfur and iron are the most significant (Fig. 3). At theseafloor, bacteria use organic carbon to make pyrite (FeS2) from the reduction ofseawater sulfate that diffuses into the sediments. When the pyrite is subsequently
buried, it effectively removes the reduced partner of O2 from being oxidized andprovides a source of O2. Pyrite generation and burial can be written as follows(Berner, 2004):

photosynthesis:

15H 2O + 15CO 2 = 15CH 2O + 15O 2

sulfate reduction: 15CH 2O +2Fe 2O3 + 16H + + 8SO 2!
= 4FeS2 + 23H 2O + 15CO 2
4
(10)
+
2!
net:
2Fe 2O3 + 16H + 8SO 4 = 4FeS2 + 8H 2O + 15O 2

8Holland (2002) estimates that the pyrite burial flux contributes 7.8±4.0 Tmol O2 yr-1.
Other minerals that are important for the redox balance include sulfates, which can beburied as evaporite minerals on continents, and ferrous iron. Their fluxes can be
estimated from examining the quantity of various redox-sensitive minerals in average
sedimentary rocks, as shown in Table 1.
Table 1. The source flux of O2 on modern geologic timescales (where thesedimentary sources follow Holland (2002)).
Process
Stoichiometry
2H2O → 4H(↑space) + O2

Hydrogen escape

Carbon burial
Pyrite burial
Pyrite burial
Sulfate burial
Fe2+ burial

O2 gain(Tmol/yr)
0.02

wt% in new
sediments
0.6 ± 0.1
0.6 ± 0.2

CO2 → C + O2
10 ± 3.3
Fe(OH)3 + 2H2SO4 → FeS2 +7.0 ± 3.6
15/2SO2 + 7/2H2O
0.2 ± 0.1
FeO + 2SO2 → FeS2 + 5/2O2
0.8 ± 0.4
0.3 ± 0.1
SO2 + 1/2O2 + H2O → H2SO4 -(0.3 ± 0.1)
1.5 ± 0.6
1/2Fe2O3 → FeO + 1/4O2
0.9 ± 0.4
Total rate of oxygen production = 18.4 ± 7.8

13.08.2.3 The oxygen sink fluxes
In recent geologic history, the flux of O2 from the sources shown in Table 1
must be approximately balanced by sink fluxes so that the O2 level remains relatively
stable. The persistence of advanced animals that rely on high O2 concentrations forthe past several hundred million years suggests that there were not wildly fluctuating
O2 levels during this period. Fig. 3 depicts the various sink fluxes that remove O2
from the atmosphere, which we now quantify.
The process of continental oxidative weathering occurs when O2 dissolved inrainwater or rivers reacts with reduced minerals on the continents, such as exposedorganic carbon, ferrous iron (Fe2+) minerals, and sulfides. Essentially, continental
weathering reverses the chemical equations that liberate O2 shown in Table 1.
Examining average continental rock allows estimates for the rate of continental
weathering of various minerals to be made. The estimates in Table 2 sum to a loss of
15.5 ± 6.7 Tmol O2 yr-1 during continental weathering.
Other sinks of O2 arise from reactions with various reductants that include
gases, such as H2, CO, CH4, H2S, and SO2, released from high temperature subaerial
volcanism (where rocks melt) and metamorphism (where the rocks are heated and but
do not melt), seafloor minerals encountered by O2 dissolved in percolating seawater,and gases and dissolved minerals released from hydrothermal vents into the oceans.
Table 2 lists various estimates from the literature for these sink fluxes.
Within the rather large uncertainty, the total source flux of 18.4 ± 7.8 Tmol O2
yr-1 from Table 1 agrees with the total sink flux of 22.1 ± 8.4 Tmol O2 yr-1 given in
Table 2. The sink fluxes in Table 2 imply that about 70% of the O2 in the atmosphereis eventually removed by oxidative weathering, while the remaining 30% is lost to
metamorphism, subaerial volcanism, seafloor weathering, and submarine volcanism,
going in order from the largest to smallest sink.

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Table 2: Modern long-term sink fluxes for O2. Sources are: 1. Holland (2002) 2.
Catling and Kasting (2012) and references therein. 3. Fiebig et al. (2009). 4. Using thelower estimate of 1.25-2.5 Tmol CH4/yr from Etiope et al. (2009) 5. Sleep (2005).
Loss Process
StoichiometryO2 loss
Ref.
(Tmol/yr)
Continental Weathering
Wt% inaveragerockundergoing
weathering
Carbon10.45 ± 0.1
C + O2 → CO2
7.5 ± 2.5
weathering
Sulfide1
0.6 ± 0.2
FeS2 + 15/2SO2 + 7/2H2O →7.0 ± 3.6
weathering
Fe(OH)3 + 2H2SO42+Fe1
1.5 ± 0.6
FeO + 1/4O2 → 1/2Fe2O3
1.0 ± 0.6
weathering
Subtotal for continental oxidative weathering flux = 15.5 ± 6.7
Surface volcanism
Gas flux(Tmol/yr)
H2
1±0.5
H2 + 1/2O2 → H2O
CO
0.1±0.05
CO + 1/2O2 → CO2
H2 S
0.03±0.015
H2S + 2O2 → H2SO4
SO2
1.8±0.6
H2O + SO2 + 1/2O2 → H2SO4
Subtotal for surface volcanism =Surface metamorphism
CH4 (abiotic)
0.3
CH4
1.25
(thermogenic)

CH4 + 2O2 → CO2 + H2O
CH4 + 2O2 → CO2 + H2O
Subtotal for metamorphic gases =Submarine volcanism
H2
0.1±0.05
H2 + 1/2O2 → H2O
H2 S
0.35±0.13
H2S + 2O2 → H2SO4
CH4 (ridge)
0.01±0.005
CH4 + 2O2 → CO2 + H2O
CH4 (off-axis)
0.25±0.12
CH4 + 2O2 → CO2 + H2O
Subtotal for submarine volcanic gases =Seafloor oxidation and subduction
Fe2+
12FeO + SO4= → 4Fe3O4 + S=
conversion tomagnetite
Total sink from subaerial and submarine reductants =0.5±0.3
0.05±0.03
0.06±0.03
0.9±0.3
1.5±0.7

2
2
2
2

0.6
2.5

3
4

3.1
0.05±0.03
0.7±0.4
0.02±0.01
0.5±0.3
0.8±0.7
1.2±0.3

22.1 ± 8.4

5

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13.08.2.4 General principles of oxygen change
For a steady concentration of O2 in the atmosphere-ocean system, the O2 source flux
FO2source, must be equal to the O2 sink flux, FO2sink. Oxygen levels will obviously
increase if the source flux exceeds the sink flux. For specificity, it is advantageous toconsider O2 changes mathematically, in terms of a flux balance equation. If we denote
the reservoir of O2 in the atmosphere-ocean system as RO2, it will change with time t,
as follows:d(RO2 )= FO2source ! FO2sinkdt= FO2source ! (Fvolcanic + Fmetamorphic + Fweathering )(11)Here, the O2 source flux arises from the relatively limited number of processes
shown in Table 1, principally, the burial of organic carbon. In contrast, the O2 sink
fluxes comprise a very large number of chemical reactions shown in Table 2. In thesecond line of eq. (11), we break up the sink fluxes into three types: Fvolcanic, which isthe flux of O2 that reacts with reductants from subaerial and submarine volcanism,
Fmetamorphic, which is the flux of O2 consumed by reductants from metamorphism, andFweathering, which is the flux of O2 consumed by reduced minerals in seafloor andcontinental weathering. Today, RO2 is 3.8 × 107 Tmol O2 and the various source andsink fluxes in Tmol yr-1 are shown in Tables 1 and 2. The average amount of time an
O2 molecule spends in the atmosphere-ocean system today is thus ~2 million years,
which is simply RO2 divided by the source flux from Table 1 of ~18 Tmol O2 yr-1.
This timescale is clearly geologically short. In the past any sustained difference
between FO2source and FO2sink would have changed RO2.

13.08.2.5 The definition of ‘anoxic’ versus ‘oxic’ atmospheres
Before considering the effect on the Earth’s atmospheric chemistry of theGOE, let us define the terms ‘anoxic’ and ‘oxic’ atmosphere. Oxygen is absent fromthe atmospheres of many other celestial bodies, such as the giant planets and Saturn’slargest moon, Titan. All of these atmospheres are ‘anoxic’. But they also shareanother characteristic, which is that their redox chemistry is dominated by H2 orhydrogen-bearing reducing gases, such as CH4.
Hydrogen-bearing reducing gases (including hydrogen itself) behave as if they
are the inverse of oxygen. Oxygen reacts rapidly with hydrogen-bearing reducing
gases, so a decrease in the concentration of one allows an increase of the other.
Anoxic atmospheres are all relatively rich in hydrogen-bearing reducing gases.
There is 82.2±2.6% H2, 96.3±2.4% H2, ~82.5% H2 and ~80% H2 in the atmospheresof Jupiter, Saturn, Uranus and Neptune, respectively, while Titan’s lower atmospherehas ~5% CH4. In contrast, Earth’s current atmosphere (21% O2, 1.8 ppmv CH4) andthat of Mars (0.13% O2, 15 ppmv H2) are both oxidizing. Given our observations ofplanetary atmospheres, it is clearly useful to define an ‘anoxic’ atmosphere as rich inhydrogen-bearing reducing gases and an oxic atmosphere as poor in hydrogen-bearing
reducing gases.

11
Our definition of anoxic versus oxic atmospheres helps us understand that theArchean atmosphere could have remained anoxic even if early oxygenic
photosynthesis pumped O2 into the atmosphere. Anoxia would have persisted if O2
had been efficiently removed and overwhelmed by a flux of reducing gases into theatmosphere (including, in this tally, any flux of reducing cations in the ocean incommunication with dissolved atmospheric O2). Specifically, we can define anoxygenation parameter as follows (Catling and Claire, 2005):
K oxy ==

O 2 source flux
non-weathering O 2 sink flux
Freductant_burial
Fmetamorphic + Fvolcanic;

K oxy > 1 gives an oxic atmosphere(12)

K oxy < 1 gives an anoxic atmosphere

If this parameter exceeds unity then O2 in the atmosphere is quickly removed byreactions with reducing species. The result is that hydrogen-bearing reducing gases
dominate the redox chemistry of the atmosphere even if there is a photosynthetic
source of O2. Detailed models that include photochemistry show that such a situation
can occur (Claire et al., 2006; Pavlov et al., 2001; Zahnle et al., 2006).

13.08.3 ATMOSPHERIC CHEMISTRY THROUGH
THE GREAT OXIDATION EVENT (GOE)
To understand the causes and consequences of the GOE, ideally, we would like tohave time series concentrations of the redox-sensitive gases in the atmosphere (O2,
CH4 and H2) before and after the GOE. Information about the levels of other gases,
such as CO2, N2O, and N2, would also help us understand associated climate change.
CO2 and N2O are greenhouse gases, while the concentration of N2 modulates the sizeof the greenhouse effect through the pressure-broadening of infrared absorption lines(Goldblatt et al., 2009). However, apart from some constraints on O2 levels, we havevery few reliable geologic determinations of the amounts of other gases. Most of ourinferences are very indirect. Table 3 contains estimates of either concentrations orlimits on concentrations of various atmospheric gases and marine sulfate. The basis
for these estimates is described in the next two subsections.
Table 3. A summary of our approximate knowledge of species concentrations before
and after the Great Oxidation Event. (See text for sources). PAL = Present
Atmospheric Level of a gas.
SpeciesAtmospheric O2
Atmospheric CO2
Atmospheric CH4
Atmospheric N2O
Atmospheric N2
Marine SO42-(aq)Anoxic late Archean< 1 ppmv≤4%~103 ppmv
unknown<2.1 PAL, probably lower
than 1 PAL<0.2 mM

Oxic Middle Proterozoic
0.2-2% by volume≤0.4%~102 ppmv
speculatively ~10-20 PAL
unknown, speculatively
similar to PAL~1-4 mM

12

13.08.3.1 Constraints on atmospheric gases before the GOE
The strictest constraint on pre-GOE O2 levels comes from mass-independent
fractionation (MIF) of sulfur isotopes, which suggests that the O2 level was less thanabout 1 ppmv in the Archean (Pavlov and Kasting, 2002; Zahnle et al., 2006). There
are four stable isotopes of sulfur, 32S, 33S, 34S and 36S. In mass-dependent
fractionation (MDF), isotopes are fractionated according to the difference between
their masses. Consequently, in MDF, 34S is fractionated twice as much with respect to
32
S as is 33S because the 34S-32S mass difference is twice that of 33S-32S. MDF applies
to most fractionation processes such as the diffusive separation of gases or the
kinetics of microbial metabolism. Expressed in terms of the delta notation (δxS =[(xS/32S)sample/[(xS/32S)standard)-1]), variations in δ33S are usually about half those in
δ34S for MDF, which defines a straight line relationship. In contrast, MIF is any
fractionation that deviates from the linear relationship. In simple terms:! 33S " 0.5(! 34 S) is mass dependent fractionation (precisely, ! 33S = 0.515! 34 S) (13)! 33S " 0.5(! 34 S) is mass independent fractionation (precisely, ! 33S " 0.515! 34 S) (14)The proportionality constant can be considered ~0.5 for pedagogical purposes but amore precise value is 0.515, as shown in parentheses above. Similarly, a MDF straightline defines the relationship of δ36S and δ34S as roughly δ36S ≈2δ34S or more precisely
δ36S = 1.90δ34S.
A notation called “cap-delta” is used to express the magnitude of the MIF, andrefers to a deviation from the MDF straight line, as follows:! 33S = " 33S # 0.515" 34 S
! 36S = " 36S # 1.9" 34 S

(15)
(16)

A graph of Δ33S for pyrites and sulfates over geologic time (see Chapter 13.05) showsan unambiguous step change around 2.4 Ga that has been interpreted as evidence forthe rise of O2 (Farquhar et al., 2000; Farquhar et al., 2010).
The relationship between MIF in sulfur isotopes and atmospheric redox
chemistry arises because photochemistry is the only process capable of producing
Δ33S up to ~12‰, which is observed in the pre-GOE record. In the lab, large MIF is
produced during photolysis of SO2 and SO, as follows (Farquhar et al., 2001):

SO 2 + h! = SO + O
SO + h! = S + O

(17)
(18)

The magnitude of the MIF depends on the isotopologue of the sulfur gas, i.e., whether
the gas is 32SO2, 33SO2, 32SO, etc. (Danielache et al., 2008; Lyons, 2009). Little MIF
is observed during the photolysis of H2S. Given the strong absorption of shortwavelength photons in the upper atmosphere by CO2 and H2O, for all practical
purposes, the photons in the early atmosphere that induce photolysis in eqs. (17) and(18) have wavelengths of 190-220 nm. Such ultraviolet photons can penetrate thetroposphere in an anoxic atmosphere that lacks a stratospheric ozone layer. The
elemental sulfur produced in eq. (18) polymerizes in anoxic atmospheres to make S8aerosols that rain out. However, sulfate aerosols are also produced from oxidizing

13
pathways because CO2 and H2O are present in any realistic early Earth atmosphereand are a source of photochemical oxidants (eq. (1)). Consequently, sulfate and S8aerosols can rain out with MIF signatures of opposite sign in Δ33S and can besegregated in the rock record because of different preservation pathways. Exactly how
this process works is still a matter of research.
The presence of MIF in sulfur isotopes has various implications for the ancient
atmospheric composition. In an oxic atmosphere, such as today’s, the sulfurcontaining gases are oxidized into sulfate, which is an isotopically homogenous single
species with no MIF that rains out and washes to the ocean. Photochemical models
show that the presence of MIF in sulfur isotopes requires three conditions, as follows(Zahnle et al., 2006):
1) Relatively abundant methane. Photochemical models show that enoughmethane or hydrogen is needed as a reducing agent to form and precipitate S8from the upper troposphere (Ono et al., 2003). The presence of sulfur MIF
doesn’t distinguish whether the reducing gas was hydrogen or methane.
Certainly, H2 is an expected biogenic gas as well as methane in the Archean(Hoehler et al., 2001). However, if the H2 concentration had been more than
~4×10-5 bar, it would be consumed by biology at the Earth’s surface andconverted into methane, according to experiments (Kral et al., 1998). Also,
biological thermodynamics predict atmospheric concentration ratios of
CH4/H2 in the range 10-30 (Kasting et al., 2001). Consequently, it isreasonable to conclude that sulfur MIF is indirect evidence that the Archean
atmosphere was relatively rich in methane. Models suggest 10s-1000s ppmv,
with the higher values giving larger rainout fluxes of S8, presumably linked tolarger Δ33S values.
2) Ground-level O2 < 1 ppmv. An anoxic atmosphere is required to allow theformation of both sulfate and S8 aerosols that carry an S-MIF signal (Pavlov
and Kasting, 2002). Also, lack of stratospheric ozone is required to allow
sufficiently shortwave photons to photolyze tropospheric sulfur-bearing gases.
3) Sufficient input of sulfur-bearing gases to the atmosphere. The production rateof polymers of elemental sulfur scales with the amount of sulfur gases. The
formation of S2 polysulfur depends approximately on the square of theconcentration of S produced from eq. (18) because it requires a reaction of theform S + S. By extrapolation, the production of S8 aerosols should depend on
the eighth power of sulfur concentrations and so S8 production is clearly going
to be very sensitive to fluxes of sulfur-bearing gases, as shown inphotochemical models (Zahnle et al., 2006). Such gases include SO2 fromvolcanism and possibly other S-bearing gases, perhaps biogenic ones.
The inference of trace levels of atmospheric O2 is consistent with conventional
geologic indicators of negligible O2 in the Archean. These include the absence ofcontinental redbeds, the lack of oxidized paleosols, the presence of detrital minerals inriverbeds that would otherwise dissolve in oxygenated waters, and ferruginous
conditions in the deep ocean that gave rise to banded iron formations on continental
shelves (see Chapter 13.5).
The record of Δ33S over geologic time may also tell us about concentrations ofother gases besides O2 and CH4. For example, if there had been very high CO2, which
is an efficient Rayleigh scatterer of UV, there would have been insufficient shortwave
photons reaching the lower atmosphere to produce sulfur-MIF. Farquhar et al. (2001)
suggest that levels of pCO2 less than 0.8 bar were required for the presence of sulfur-

14
MIF. This is consistent with a bound on the total atmospheric pressure of probably <1
bar deduced from fossil raindrop imprints at 2.7 Ga (Som et al., 2012).
In fact, there has been considerable debate about the amount of CO2 versus
methane in the Archean atmosphere. The general consensus is that Archean pCO2 washigher than today but insufficient for CO2-H2O greenhouse warming of the early
Earth without additional warming from other greenhouse gases such as methane ~103
ppmv and its photochemical product ethane (Haqq-Misra et al., 2008). Trace levels of
ethane derived from methane turn out to be important because ethane absorbs in theatmospheric window region 8-12 microns where most of the Earth’s infrared radiation
otherwise escapes to space. Of course, photochemical models show that higher levelsof methane and its photochemical products are expected in an anoxic atmosphere(Pavlov et al., 2000; Zahnle et al., 2006). The methane is assumed to be biogenic,
which may have been enhanced by the leaching of nickel from abundant nickel-richultramafic rocks on the early Earth (Konhauser et al., 2009). Nickel is a key cofactor
in several methanogen enzymes.
In an early study of paleosols (fossilized soils), Rye et al. (1995), suggested
that pCO2 was no more than ~10 times present levels based on the absence of siderite(FeCO3), which ought to form in anoxic porewaters in the presence of ferrous ironwith higher pCO2. However, Sheldon (2006) criticized this paper. He found that
calculated pCO2 would be comparable to modern levels (or even less) if up-to-datethermodynamic data were applied to the equilibrium between siderite and ironsilicates proposed by Rye et al. Because of uncertainty about the exact silicate
minerals in the equilibria, Sheldon suggested that the equilibria approach was flawed
and that a mass balance approach was better. With mass balance, he derived an
estimate of 23!3
PAL of CO2 for 2.2 Ga paleosols, where his PAL (Present
÷3
Atmospheric Level) was taken as 370 ppmv CO2. Driese et al (2011) applied the same
method to a 2.69 Ga paleosol to derive a pCO2 of 10-50 PAL at that time.
The Driese et al. (2011) and Sheldon (2006) pCO2 estimates come from apaleosol weathering equation of Holland and Zbinden (1988):$ K CO ! r
DCO ! # 'M
22
(mol cm -2 yr -1 ) = &+") pCO 2
3
t
L
10
&%
)(

(19)

On the right, the first term in the brackets accounts for CO2 dissolved in rainwater,while the second term corresponds to diffusion of CO2 through the soil. On the left isa column weathering rate where M is the number of moles of CO2 per cm2 required to
produce the observed paleosol weathering profile and t is the soil formation time. The
carbon dioxide partial pressure is pCO2 in atm, K CO is the Henry’s Law constant for2

CO2, r is the amount of rainfall absorbed by the soil (cm yr-1), κ is a conversion
constant that is the ratio of seconds in a year divided by the number of cm3 per mol ofgas at standard temperature and pressure (= 1.43×103 s⋅cm3/(mol yr)), DCO is the2

2 -1

diffusion coefficient for CO2 in air (0.162 cm s ), α is the ratio of the CO2 diffusion
coefficient in air to that in soil, and L is the total depth of the soil horizon. Holland
and Zbinden assign an uncertainty of a factor of 10 in either direction on α whereas
Sheldon suggests that the biggest uncertainty is a factor of 2 in either direction in t,
and he sets α = 0.1 with a variability of ±20 percent. Overall, the exact uncertainty on
pCO2 obviously depends on assumptions about r, α and t. But the general result is15that pCO2 was probably up to 100 PAL but not higher in the late Archean. This
constraint from paleosols is consistent with theoretical models that suggest greaterhydrothermal alteration of the early seafloor basalt that may have removed CO2 andlimited its abundance (Zahnle and Sleep, 2002).
In summary, the pre-GOE atmosphere likely had < 1 ppmv O2, high levels of
methane ~103 ppmv (but possibly falling to ~10 ppmv around 2.4 Ga), and levels of
CO2 up to 100 times the modern atmosphere’s. Constraints on the amount of N2
remain very limited at present. Evidence from Archean fossil raindrop imprints
suggests that the total pressure, pN2 was probably not much higher than today’s levelor perhaps lower (Som et al., 2012). A lower pN2 might be expected because thebiogenic release of N2 to the atmosphere in the modern nitrogen cycle comes from
nitrogen species that are oxidized prior to denitrification. This oxic part of the cycle
would have been very limited when O2 was only a trace gas.

13.08.3.2 Constraints on atmospheric gases after the GOE
The Canadian geologist Stuart Roscoe first noticed that large glaciation eventswere associated with atmospheric redox change around 2.4 Ga (Roscoe, 1969).
Roscoe’s fieldwork was in the 2.45-2.22 Ga Huronian Supergroup of Canada, which
is a section of mainly sedimentary rocks that contains detrital minerals of pyrite anduraninite stratigraphically below three diamictites. The diamictites are interpreted as
evidence for three glaciations. Reduced detrital minerals, which should dissolve inoxygenated waters, suggest anoxic conditions before the glaciations. In addition, pre2.45 Ga paleosols have upward depletion of total iron, which suggests weathering
under an anoxic atmosphere that was able to mobilize Fe2+ in solution (Prasad andRoscoe, 1996). Pre-2.45 Ga Huronian paleosols also have cerium profiles indicativeof anoxic weathering (Murakami et al., 2001). In contrast, the top of the Huronian
succession contains red beds and oxidized paleosols. To Roscoe, such an associationsuggested a causal link between the atmospheric oxygenation and glaciations.
Evidence for Paleoproterozoic glaciation is found in Australia, India, Russia,
Scandinavia, South Africa, and elsewhere in North America (Bekker et al., 2003;
Eyles and Young, 1994), which implies that it was global in extent. Moreover, theglobal climate is inferred to have been extremely cold because magnetic studies in theSouth African glaciations suggest low latitudes (Evans et al., 1997; Evans, 2003).
A possible explanation of the glaciations is that the rise of O2 could havediminished the greenhouse effect because an oxic atmosphere is chemically
incompatible with comparatively large amounts of methane that should have existed
in an anoxic atmosphere (Pavlov et al., 2000). A more nuanced view is that thegradual loss of reducing power in the late Archean atmosphere-ocean-crust system led
to an increasing pool of seawater sulfate, smaller biogenic methane fluxes, and thus adecline of atmospheric methane levels that initiated the cooling of the Earth (Zahnle
et al., 2006). Sulfur isotope evidence from the post-GOE glaciation supports this latterview (Papineau et al., 2007; Williford et al., 2011). Possibly, removal of CO2 throughsilicate weathering prior to the glacial times may have also contributed because of the
subaerial occurrence of large igneous provinces at 2.45 Ga (Kump and Barley, 2007;
Melezhik, 2006). In any case, the three Paleoproterozoic glacial episodes suggest thatthe atmosphere may have oscillated in composition and climate before finally settling
down to a permanently oxic state. Some biogeochemical models are able to reproduce
such oscillations in atmospheric composition and climate (Claire et al., 2006).
However, we note that the behavior of such models depends upon the16parameterization of the response of biogenic O2 and CH4 fluxes to global temperaturechange. There remains great uncertainty about the magnitude and scaling of such abiogenic response to global climate.
In general, the persistent absence of significant S-MIF throughout the latterpart of the Paleoproterozoic suggests that atmospheric O2 concentrations must haveincreased to > 1 ppmv. Although it is clear that the atmosphere was oxic, there are
relatively few constraints on the exact O2 levels. In the middle Proterozoic, theconcentration of marine sulfate is inferred to have been at a level of a few mM
compared to 28 mM today (Kah and Bartley, 2011; Kah et al., 2004; Shen et al.,2003) and it is likely that atmospheric O2 levels stayed somewhere between ~0.2-2%
in absolute terms throughout much of the Proterozoic. It is only much later, by ~551
Ma, that there are estimates of atmospheric O2 >3% in absolute concentration. At thattime, Canfield et al. (2007) calculate an O2 level >3% by assuming an O2 demand for
500-1500 m water depths of sediments at the Avalon Peninsula, Newfoundland. This
Neoproterozoic rise in O2 is also supported by a significant jump in Mo and V
enrichments in black shales (Scott et al., 2008) along with δ98Mo (Dahl et al., 2010).
A non-intuitive consequence of the rise of O2 in the Paleoproterozoic was aprobable re-establishment of atmospheric methane concentrations (Catling et al.,
2004; Claire et al., 2006; Goldblatt et al., 2006) that were enough to produce ~6-7°Cof greenhouse warming (Roberson et al., 2011). When O2 rises and the stratosphericozone layer is established, the net reaction between O2 and methane in the troposphereslows down because the troposphere is shielded from shortwave UV (Catling et al.,
2004; Claire et al., 2006; Goldblatt et al., 2006). Provided that there was a significant
methane flux, the methane levels could have been as high as ~102 ppmv until the
Neoproterozoic.
Today, there is an enormous flux of seafloor CH4 that never reaches theatmosphere, estimated as a few tens of Tmol CH4 yr-1 (Catling et al., 2007). Thismethane is consumed by microbial SO42- reduction at the CH4-SO42- transition zone in
sediments (D'Hondt et al., 2002), so that the ocean is an inconsequential source ofCH4 for the modern atmosphere. But in a Proterozoic ocean with extensive euxinia, a
big flux of methane could have vented from the seafloor to the atmosphere (Catling etal., 2002; Pavlov et al., 2003). Evidence for euxinia comes, in part, from molybdenum
isotopes. Under high oxygen conditions, uptake of Mo on marine solids preferentiallyremoves 95Mo relative to 98Mo, leaving an isotopically heavy ocean. However, under
low oxygen conditions, seawater is 95Mo-enriched. Sulfides in euxinic deep waters
sequester Mo efficiently and can capture the marine Mo isotope composition. Rocks
with 95Mo-enrichment from the 1.7-1.4 Ga MacArthur Basin, Australia, suggest that
the Mesoproterozoic deep ocean commonly had a euxinic area that was probably
several percent compared to today’s 0.3% (Arnold et al., 2004; Kendall et al., 2009;
Kendall et al., 2011).
Another redox-sensitive greenhouse gas that may have been moderately
important after the GOE is nitrous oxide, N2O (Buick, 2007). The source of N2O isdenitrification, when bacteria convert oxidized nitrogen compounds into N2 alongwith minor N2O release. The conversion of N2O to N2, which is the terminal step ofdenitrification, depends on an enzyme, nitric oxide synthase, which contains 12 Cu
atoms in its complex. Proterozoic oceans with widespread euxinia would tend toscavenge Cu into copper sulfides. If the oceans were copper-depleted, possibly more
nitrogen was released as N2O by a factor of 10-20. The concentration of N2O roughly
scales with its flux to the atmosphere. An N2O level 10-20 times that of today’s 0.3
ppmv could provide 3-5°C of greenhouse warming (Roberson et al., 2011). Of course,

17
there are no direct constraints on N2O levels, but the Mesoproterozoic is an era withno evidence of glaciation, which suggests an unusual duration of climate stability andpersistent greenhouse warming.
Carbon dioxide levels in the mid-Proterozoic were probably lower than thosein the Archean and Paleoproterozoic. Analysis of paleosols from around 1.1-1.0 Ga
suggests CO2 levels ≤ 10 PAL (Mitchell and Sheldon, 2010; Sheldon, 2006). Also,
microfossil evidence for in vivo calcification of cynanobacteria at 1.2 Ga suggests
pCO2 ≤ 10 PAL (Kah and Riding, 2007). In contrast, paleosol analysis around 1.8 Ga
implies CO2 levels up to ~50 PAL. Table 3 summarizes the various estimates of gas
concentrations before and after the GOE.

13.08.4 EXPLAINING THE RISE OF O 2
A consideration of what we mean by anoxic versus oxic atmospheres (Sec.
13.08.2.4) allows us to deduce what was required for a rise of O2. If we examine eq.
(11), it is clear that a positive value of d[O2]/dt, i.e., an O2 increase, only occurs when
the source of O2 exceeds the O2 sink. Alternatively, in terms of the parameter Koxy
introduced in eq. (12), to transition from an anoxic to oxic atmosphere, there are only
two choices: either the O2 source (Freductant_burial) increases or the non-weathering O2sink (Fmetamorphic + Fvolcanic) decreases. Consequently, various theories, described
below, which attempt to explain the GOE, are all variants of these two themes.

13.08.4.1 Hypotheses for an increasing source of O2
One group of hypotheses for the rise of O2 concern the idea that the source of
O2 increased because of enhanced rates of burial of organic carbon. These ideas
divide into two subgroups: A pulse of organic burial caused the rise of O2 or a longterm trend of increasing organic carbon burial led to a tipping point. Sometimes, thelatter concept is tied to models of continental growth, where it is argued that more
continental margin area allowed for more organic burial either directly or throughnutrient supply.
At one end of the spectrum is the idea that the rise of O2 happened almost
immediately (in geological time) after the origin of oxygenic photosynthesis(Kirschvink and Kopp, 2008). The counter-argument is that there are multiple lines of
independent evidence for the production of O2 several hundred million years before
the rise of O2 at 2.4 Ga. Geochemical evidence of small amounts of O2 before theGOE include iron speciation in marine sediments (Kendall et al., 2010; Poulton et al.,2010; Reinhard et al., 2009), trace amounts of molybdenum and rhenium (Anbar etal., 2007; Wille et al., 2007), sulfur isotope fractionation associated with the presenceof sulfate (Kaufman et al., 2007), and extremely 12C-enriched organic carbon
consistent with microbes that oxidize methane with oxidants derived from O2 or O2
itself (Hayes, 1994; Hayes and Waldbauer, 2006). Although controversial, organic
biomarkers for cyanobacteria and steranes derived from steroids biosynthesized with
O2 are reported up to ~0.3 Gyr before the rise of O2 (Eigenbrode et al., 2008;
Waldbauer et al., 2009). Conical stromatolites back to 2.7 Ga also indicate
cyanobacterial production of O2 through their disrupted, curled and contorted laminae
with enmeshed millimeter-scale bubbles (Bosak et al., 2009). Because one would
need to dismiss all the above lines of evidence, the hypothesis of Kirshvink and Kopp
(2008) has failed to garner much support in the Precambrian science community.

18
The idea that there was a pulse of burial of organic carbon arises fromexcursions in the carbon isotope record. Organic matter preferentially extracts thelighter isotope, 12C, from the ocean, so that seawater carbon, as recorded in marine
carbonates, should become more 12C-depleted (or higher in δ13C) if organic burial
rates increase. Two or three δ13C excursions in carbonates occur during the period
2.4-2.06 Ga, including the global 2.2-2.06 Ga Lomagundi excursion, named after a
district in Zimbabwe (now called Makonde) where 13C-rich dolomites were firststudied (Bekker et al., 2008; Bekker et al., 2006; Maheshwari et al., 2010; Melezhik
and Fallick, 2010; Melezhik et al., 2007; Schidlowski et al., 1976). Karhu andHolland (1996) suggested that the Lomagundi carbonates indicated a pulse in organiccarbon burial that may have caused the rise of O2 because of associated O2 release via
eq. (9). However, improved radiometric dating shows that the Lomagundi excursion
cannot have caused the GOE because it happened after the GOE at ~2.4 Ga. Instead,the excursion is probably an effect of the rise of O2.
The remaining possibility (on the O2 source side) is that there was a gradualincrease of organic carbon burial. Whether or not this happened depends upon aninterpretation of the somewhat noisy record of δ13C in carbonates and organic carbon,
for which there are differing views. One pioneering study suggested a secular increasein organic carbon burial fluxes based on boxcar averaging of the δ13C time series(DesMarais et al., 1992). However, the averaging started at 2.6 Ga when very
negative δ13C for organic carbon (δ13Corg) occur. These anomalously low δ13Corg
values were subsequently attributed to methanotrophy (Hayes, 1994; Hayes andWaldbauer, 2006; Hinrichs and Boetius, 2002). If so, the δ13Corg represents diagenetic
processing in sediments rather than an isotopic balance involving global marine
seawater. Another model concerns a gradual decrease of the uptake of 12C-depleted
carbon into hydrothermal carbonates from the early to late Archean, which would
allow the fraction of organic carbon burial to increase within the constraints of theδ13C record (Bjerrum and Canfield, 2004). This model assumes a gradient in the
13 12
C/ C ratio between the surface and deep seawater from a “biological pump”.
However, of δ13C measurements in 3.46 Ga seafloor do not show such a gradient(Nakamura and Kato, 2004). Nonetheless, analysis of the carbon isotope data can
produce a quite variable organic burial fraction over geologic time (see Chapter
13.09).
Continental shelves are places for burying organic carbon (Knoll, 1979), while
continental weathering is a source of phosphorus, so some hypotheses of increasingorganic burial tie organic burial to continental growth (Godderis and Veizer, 2000).
Indeed, bursts of continental growth have been inferred from zircon U-Pb ages, andspeculatively linked to O2 increases (Campbell and Allen, 2008). However, suchbursts are probably artefacts of crustal preservation from supercontinent formation(Hawkesworth et al., 2009). Indeed, the present continental crust volume wasprobably reached by ~2-3 Ga (Hawkesworth et al., 2010), which, at least on the older
end of these estimates, is 0.6 billion years before the GOE. For the younger end of
this age rage, the question is whether support is provided for a secular increase oforganic burial by the δ13C record or not.

13.08.4.2 Hypotheses for an decreasing sink of O2
An alternative and fairly common interpretation of carbon isotopes is that
there is no clear signal in the noisy record and that since 3.5 Ga the average δ13C

19
value of marine carbonates has remained roughly constant at ~0‰ while the average
δ13Corg has been about −25‰ (Schidlowski, 1988). In this view, the fraction of carbon
buried as organic relative to marine carbonate has remained roughly constant(Holland, 1984, 2002, 2009; Kump and Barley, 2007; Kump et al., 2001).
The fraction of carbon buried as organic carbon can be estimated from massbalance. On timescales exceeding ~105 years, which is the residence time of carbon
on the ocean, the total amount of carbon entering the atmosphere-ocean system must
balance the total amount leaving. Because the carbon is buried either as carbonate ororganic carbon, the burial fractions of each type of carbon, fcarb and forg, respectively,
will sum to unity:

fcarb + f org = 1(20)

Similarly, the input and output flux must balance for the isotopic composition ofcarbon, however one expresses it, whether a 13C/12C ratio or δ13C. So, in steady state,
the isotopic composition of carbon input from mantle and riverine sources, δ13Cin will
equal the output, δ13Cout. We can expand eq. (20) to include the isotopic compositionof organic matter δ13Corg, and carbonates δ13Ccarb, as follows:! 13Cin = ! 13Cout = f org! 13Corg + fcarb! 13Ccarb= f org! 13Corg + (1" f org )! 13Ccarb = ! 13Ccarb " f org # B(21)

Here, ΔB is the difference between carbonate and organic carbon isotopic
composition:! B = " 13Ccarb # " 13Corg(22)

Let us consider the δ13C values in the above equations. Today, volcanic gases
show that mantle carbon has δ13C around −5‰. This input value appears to havebeen constant over geologic time because ancient carbon in peridotitic xenoliths have
δ13C of about −5‰ (Pearson et al., 2003), as does carbon in mantle-derived basalts
and carbonatites (Mattey, 1987). Also, the riverine source from the weathering oforganic carbon and carbonates has a mean δ13C value of about −5‰ as well (Holser etal., 1988). So, δ13Cin ≈ −5‰. Marine sediments throughout geologic time have
δ13Ccarb ≈ 0‰. While δ13Corg has considerable scatter, it averages around −25‰
(Schidlowski, 1988). Thus, from eq. (22), we obtain ΔB ≈ 25‰. Substituting this andother values into eq. (21), we have −5‰ = 0 - forg(25‰). Solving for forg, we get forg
~0.2. Consequently, roughly 20% of incoming carbon is buried as organic carbon,
while the remaining 80% fluxes out as carbonate carbon.
Given the relative constancy of average 13Ccarb and δ13Corg since 3.5 Ga, somehave argued for relative constancy of the ratio of the burial rates of organic andcarbonate carbon (Holland, 1984, 2002, 2009; Kump and Barley, 2007; Kump et al.,2001). On the other hand, others have interpreted the data to suggest a larger
magnitude of ΔB in the Archean. For example, Hayes and Waldbauer (2006), suggestthat forg was ~0.15 throughout much of Earth’s history, and rose to ~0.24 since 0.9 Ga.
Canfield (2005) also favors lower forg in the Archean (see also Chapter 13.09).
If there had been relative constancy of the organic burial fraction within the20error bars of the data and oxygen photosynthesis arose before 2.7 Ga, then the only
way that Archean O2 could remained a trace atmospheric gas is if there had been alarge O2 sink from a volcanic and metamorphic reductants (Catling et al., 2001;
Holland, 2002; Kasting et al., 1993; Kump et al., 2001). Oxygen would have risen
once the sink declined.
Following this line of logic, one group of hypotheses for the GOE suggest that
there was a greater proportion of reducing gases in volcanic gases on the early Earthand that this fraction diminished in time. In other words, volcanic gases graduallybecame more oxidized (Holland, 2002; Kasting et al., 1993; Kump et al., 2001).
Volcanic gases have a redox state that is controlled by the redox state of the mantle.
Simply put, a more reducing mantle produces more reducing volcanic gases. So
volcanic gases could have been more reducing in the Archean if the Archean mantle
had been at a lower oxidation state than the Proterozoic mantle. However, analyses ofredox-sensitive chromium and vanadium abundances in basalts (Canil, 2002; Delano,
2001; Li and Lee, 2004) indicate that the mantle’s average oxidation state hasremained roughly constant since ~3.5 Ga. This is somewhat surprising because it
seems almost certain that the mantle has lost hydrogen to the atmosphere and thenspace (Catling et al., 2001; Kasting et al., 1993). However, the ability to lose reducing
power and stay at the same oxidation state can be explained by the mantle’s size andstrong redox buffering (Lee et al., 2005).
The mantle redox state is expressed in terms of oxygen fugacity F O2. The
term ‘fugacity’ generally represents the effective partial pressure of a gas inthermodynamic equilibrium with a particular mineral assemblage. But O2 itself is nota gas species in rock. Instead, the dissociation of H2O, CO2, and SO2 or the reductionof Fe3+ supplies oxidizing power, for which F O2 is a theoretical index of the redox
state (Frost, 1991). Petrologists define various synthetic mineral assemblages tocharacterize oxygen fugacities. These assemblages do not exist in nature, but are used
to express an equivalent F O2 for natural assemblages. A common synthetic
assemblage is the quartz-fayalite-magnetite (QFM or FMQ) buffer:
3Fe 2SiO 4 + O 2 ! 2Fe3O 4 + 3SiO 2(23)

The oxygen fugacity for this reaction is calculated as F O2 = 10−8.5, using atemperature and pressure-dependent parameterization (Mueller and Saxena, 1977)
with conditions appropriate for surface volcanism of 5 atm pressure and 1200oC, or
1473.15 K (Holland, 1984).
Volcanic gases can be assumed to attain a redox state defined by the mantle
oxygen fugacity. For example, the ratio of volcanic H2 to steam will be determined byan equilibrium reaction as follows:K

H2O
H 2O(g) !##" H 2 + 12 O 2(24)Here, KH2O is an equilibrium constant defined through the following:
K H2O =KF H 2 !F O 0.5
F2" rH 2 = H 2 = H2O
F H 2OF H 2O F O 0.5
2

(25)

21
In the second equality, rH2 is ratio of the hydrogen and water fugacities, F H2 / F H2O.
Gibbs free energy values (available online from the US National Institute ofStandards and Technology (NIST)) allow KH2O to be calculated as 1.31×10-6.
Consequently, substituting F O2 = 10−8.5 in eq. (25), we derive a modern value of rH2
as ~0.02, which is similar to typical measured volcanic gas compositions (Holland,1978). The geochemical studies mentioned above suggest that the average oxygen
fugacity F O2 in the Archean mantle differed from the modern mantle by no more than
0.3 log10 units. If F O2 were 0.3 log10 units lower than 10−8.5, rH2 would only increase
to 0.03. Volcanic gases with ~3% of the hydrogen as H2 compared to H2O instead of~2% today would make a negligible difference to the redox balance of the earlyatmosphere. Much larger changes in oxygen fugacity are necessary for significant
effect. For example, if the Archean mantle F O2 had been 10-11.5 atm, i.e., 3 log10 units
lower than today’s 10−8.5, about three-quarters of volcanic hydrogen would have been
emitted as H2 rather than water vapor, providing a huge sink for O2. But thegeochemical data do not allow for such drastic F O2 changes.
Although the mantle redox state has apparently remained roughly constantthrough time, it is still possible for the redox state of volcanic gases to have changed
through other mechanisms. For example, consider the H2/CO2 flux ratio. The flux of
H2 is a sink on O2 whereas the flux of CO2 is related to the O2 source flux, because O2
derives from carbon fixation and subsequent organic burial. In terms of fugacities, we
can write," F % (F+( H O flux +volcanic flux of H 2! fugacity ratio = $ H 2 ' * H 2O - ! rH 2 * 2- (26)
volcanic flux of CO 2
# F H 2O & ) F CO 2 ,) CO 2 flux ,From eq. (26), we can see that if the flux ratio of H2O/CO2 through volcanoeschanged in the past, the gas flux ratio of H2/CO2 would change even if rH2 (the ratioof fugacities F H2/ F H2O) had remained constant. Holland (2009) has put forward the
hypothesis that carbon accumulated in the surface reservoir over time from outgassing
so that the H2O/CO2 flux ratio through volcanoes decreased. In that case, theproportion of H2 relative to the carbon outgassing would have been greater in the past,
providing a larger sink on O2. Holland also suggests that the SO2/H2O flux ratio
increased with time because of greater geologic cycling of sulfur. Certainly, after
sulfate became abundant, subduction recycling of sulfur should have increased,because we know that some seawater sulfur cycles through subduction zone
volcanoes today. However, increased sulfur cycling is probably a positive feedback of
the GOE, rather than its cause, because marine sulfate derives from the oxidation of
sulfide. The history of carbon cycling is more complex. Some carbon cycle models
establish a maximum surface reservoir of carbon by about 3 Ga (Zahnle and Sleep,2002), which is much earlier than the GOE but there are many variables in suchmodels.
Kump and Barley (2007) provide an alternative way to effect a change in theredox properties of volcanic gases. They suggest that the predominant style ofvolcanism evolved from submarine to subaerial. Submarine volcanism has a large
area of relatively low temperature, off-axis emissions that are more reducing thangases from high temperature subaerial volcanism. Also, if there were shallower midocean ridges, thermodynamics predicts a larger Fe2+ flux from hydrothermal vents
compared to today, especially when seawater is depleted in sulfate (Kump and22
Seyfried, 2005). Today, the subaerial volcanic sink on O2 is twice that of thesubmarine volcanic gas. Together the subaerial and submarine volcanic sinks are afactor of ~4 smaller than the O2 flux released from organic burial according toinventories in Table 2 versus Table 1. So the proposal of Kump and Barley (2007)
requires large changes in the balance of volcanic emissions between subaerial andsubmarine. Another proposal on a similar theme is that the change in degassing
pressure between submarine and subaerial volcanism shifted the redox state of thegases, based on model calculations (Gaillard et al., 2011). However, gas compositionis sensitive to temperature, so the effect of reequilibration with lower temperaturesurroundings as gases emerge from magmas requires further investigation.
A final and similar idea for a decreasing sink on O2 is that the proportion of
reducing gases in the metamorphic gas flux decreased from the Archean to theProterozoic (Catling et al., 2001; Claire et al., 2006). Such gases, by definition, arise
from regions in the crust that do not melt, and so they are not constrained by the
mantle redox.
A couple of chemical examples show how the crust could have lost hydrogenand become progressively oxidized. Some Archean banded iron formations (BIFs)(see Chapter 13.07) contain siderite (FeCO3), which decomposes to magnetite during
thermal metamorphism, as follows (McCollom, 2003)
12FeCO3 + 2H 2O ! CH 4 + 11CO 2 + 4Fe3O 4(27)

In addition, serpentinization in hydrothermal systems generates hydrogen, which we
can write schematically as:(28)
3FeO + H 2O ! Fe3O 4 + H 2
Methane is often released from rocks where such serpentinization occurs because
microbes convert the hydrogen into methane (Chapelle et al., 2002; Sherwood Lollar
et al., 2006). Serpentinization certainly occurred in Archean greenstone belts, which
are elongate regions with an abundance of altered mafic to ultramafic igneous rocks.
In eqs. (27) and (28), ferrous iron is oxidized to ferric iron and hydrogen-bearing gas
is released. In the anoxic Archean atmosphere, high levels of methane or hydrogen
would have necessarily caused a significant flux of hydrogen to escape to space (seeSec. 13.08.2.1). Subsequent metamorphic processing of the more oxidized crust
would produce less reducing metamorphic gases as a matter of Le Chatelier’s
principle during each geologic cycle.
The proposal that the crust oxidized as a result of hydrogen loss is supported
by inventories of oxidized and reduced species in the continental crust (Fig. 4). If
photosynthesis were solely responsible for the oxidation state of the crust, the numberof moles of excess oxygen in ferric iron, sulfate and free oxygen should be balancedby an equal number of moles of organic carbon (see eq. (9)). The continental crust,ocean and atmosphere contain ~2.5×1021 mol of excess O2 equivalents (Catling et al.,2001), which is mostly in the form of ferric iron that started out as ferrous iron.
However, the quantity of reduced carbon in the crust is much smaller, <1.3×1021 mol(Wedepohl, 1995). Hydrogen escape to space is a plausible explanation for the
missing reductant because the expected quantity of oxidation is the correct size. If theconcentration of methane in the Archean atmosphere had been ~103 ppmv, ascalculated in photochemical models, the hydrogen escape rate would have been ~7
×1012 mol O2 yr-1 from the theory of Sec. 13.08.2.1. Over 0.4 b.y., the escape of23hydrogen at this rate would produce an irreversible quantity of 2.8 × 1021 mol O2equivalents—more than enough to account for the oxidation of the crust. Extra
oxidation could have been taken up by the mantle with little effect on its overall redox
state.
The overall result of a decline in the proportion of reducing gases in volcanic
or metamorphic gases is the same: a transition to an oxic atmosphere. Claire et al.
(2006) describe results of a box model that tracks the comings and goings of redox
fluxes in the Earth system. The model includes parameterized photochemistry thatallows calculation of the concentrations of atmospheric O2 and methane. In this box
model, crustal fugacity controls the oxidation state of metamorphic gases and, fornumerical simplicity, the model converts organic carbon to Fe2+ equivalents via aprocess of greenschist metamorphism, 6Fe2O3 + C → 4Fe3O4 + CO2. The model canbe run either assuming that the volcanic gas sink diminished or that the metamorphic
gas sink declined. Fig. 5 shows that either scenario can generate an oxic transition.
Common characteristics of such models are that there is negligible oxidative
weathering prior to the GOE and that a transition to an oxic atmosphere occurs when
Koxy (as defined in eq. (12)) reaches unity.

13.08.5 CHANGES IN ATMOSPHERICCHEMISTRY AND CLIMATE ASSOCIATED WITH
THE RISE OF O 2
13.08.5.1 The collapse of methane at the GOE
Earlier (Sec. 13.08.3.2), we discussed the possibility that a collapse of
atmospheric methane levels was responsible for Paleoproterozoic glaciation. This
would be consistent with increasing oxidation in the Earth system, particularly greater
concentrations of sulfate in the ocean and diminished methane fluxes to theatmosphere as a result of methanotrophy using sulfate (Catling et al., 2007).
We can consider the biogeochemistry and associated atmospheric change inmore detail. In the Archean microbial ecosystem models of Kasting (2001) and Claire
et al. (2006), the key redox-sensitive biogenic gases emitted to the atmosphere are O2
and CH4. Once oxygenic photosynthesis arose, it was the main source of sedimentaryorganic matter, which could be fermented by a consortium of microbes to makemethane. The net reaction can be deduced as follows:

2H 2O + 2CO 2 !photosynthesis!!!!
" 2CH 2O + 2O 2
2CH 2O !fermentation+methanogenesis!!!!!!!
" CH 4 + CO 2

(29)

Sum: CO 2 + 2H 2O " CH 4 + 2O 2Thus, oxygenic photosynthesis followed by fermentation and methanogenesis
produces fluxes of O2 (φO2) and CH4 (φCH4) in the ratio φO2/φCH4 = 2. Of course,
methane is a relatively inert molecule whereas O2 reacts with a plethora of gases andsoluble cations, particularly on a reducing early Earth. Consequently, we would not
expect a simple 2:1 mutual destruction of O2 and CH4 but rather an atmosphere withnegligible O2 and relatively abundant methane, as data and models suggest. However,24
once standing pools of oxidants, such as marine sulfate became more available, φCH4
should have dropped relative to φO2 because of methanotrophy (Valentine, 2002).
Photochemical models show that when the φO2/φCH4 flux ratio increases, anatmosphere can flip to a permanently oxic state. Fig. 6 shows the results of aphotochemical model, which was run as a thought experiment. In this model, theground-level methane abundance was held constant at 100 ppmv, while the groundlevel O2 mixing ratio was increased. The model calculated the fluxes of CH4 and O2
needed to maintain the prescribed atmosphere. This model used fixed outgassing
fluxes of ~1 Tmol S yr-1 (with SO2:H2S in 10:1 ratio), 2.7 Tmol H2 yr-1, and 0.3 Tmol
CO yr-1. This outgassing sink on oxygen adds up to only ~2.15 Tmol O2 yr-1 but one
can think of it as a net flux that ignores the zero sum of an organic burial flux thatwould be of order ~10 Tmol O2 yr-1 and a complementary ~10 Tmol O2 yr-1 geologic
gas sink. Fig. 6 also shows the rainout fluxes of SO42- and S8 aerosols generated by
the photochemistry. When the O2 mixing ratio increases above 10-6, i.e., >1 ppm, we
see that the S8 fallout flux drops off, which we interpret as related to the loss of massindependent fractionation (MIF) of sulfur isotopes (as described in Sec. 13.08.3.1).
The plus symbols (“+”) indicate the φO2/φCH4 flux ratio, which maps to the right handvertical axis. In Fig. 6, shaded regions require biogenic fluxes that Zahnle et al.
(2006) argue are unreasonable. On the anoxic part of the graph, φO2/φCH4 isapproximately 2 but not precisely because there are fluxes of redox-sensitive rainout
species, principally hydrogen peroxide and formaldehyde, which have to be tallied up
to describe the atmosphere’s redox budget precisely. Nonetheless, examination of Fig.
6 shows that oxic solutions on the right have higher φO2/φCH4 ratios than the anoxic
solutions on the left. But the change is not much. An increase in the flux ratio
φO2/φCH4 of less than 3% creates a flip between an anoxic atmosphere with O2 mixing
ratio ≤2×10-7 and an oxic atmosphere with O2 mixing ratio ≥6×10-4. Effectively, theincrease in φO2/φCH4 is equivalent to creating an imbalanced O2 source where O2
source exceeds the outgassing flux of reductants.

13.08.5.2 The formation of an ozone shield
After the GOE, the Earth was protected from shortwave ultraviolet by astratospheric ozone layer. Stratospheric ozone (O3) is derived from molecular oxygen.
The photolysis of O2 in the stratosphere produces oxygen atoms, which then combinewith other O2 molecules to make ozone, as follows:
O 2 + h! (" < 242 nm) # O + O(30)

O + O 2 + M ! O3 + M(31)

In eq. (31), ‘M’ is any air molecule—usually the most abundant one, N2. ‘M’ removes
energy liberated by the reaction, which is dissipated in molecular collisions. Ozone
atoms formed in eq. (31) are able to absorb ultraviolet photons when they photolyze.
Such absorption protects the Earth’s surface from the biologically harmful range ofultraviolet wavelengths of 200-300 nm.
The ozone column abundance, i.e., the number of molecules above a unit areaof the Earth’s surface, is a related to how much UV can be absorbed. This column
abundance has a non-linear dependence on O2 levels. Fig. 7 shows the results of25photochemical modeling (Kasting and Donahue, 1980). A relatively minor O2
concentration ≥1% of the modern atmospheric level is sufficient to generate an ozone
column that protects the surface from harmful ultraviolet. Thus, the small amount of
O2 expected after the GOE was enough to give rise to significant ozone shielding.
This presumably would have expanded ecological niches for land-based life.

13.08.6 CONCLUSIONSThe concentration of an atmospheric gas is controlled by a kinetic competition
between its production and loss. The Great Oxidation Event (GOE) occurred because
there was imbalance between the O2 source flux and the O2 sink flux, which led to an
accumulation of O2 in the atmosphere. Several lines of evidence suggest that O2 wasbeing produced several hundred million years before the GOE at 2.4 Ga (see Chapter
13.05). However, the Archean atmosphere would have remained anoxic if O2
consumption fluxes exceeded the long-term source of O2 from the flux of organiccarbon burial. In an anoxic atmosphere, the loss of O2 to continental oxidative
weathering is negligible, so the dominant sinks on O2 in the Archean must have beenreactions with volcanic and metamorphic reductants, including reducing gases (H2,
CH4, CO, SO2, H2S) and dissolved reducing cations (e.g., Fe2+) in directcommunication with the atmosphere. A “tipping point” would have been reachedwhen the flux of O2 from organic carbon burial exceeded the O2 loss to volcanic andmetamorphic sinks. After the GOE, oxidative weathering was a significant new sinkon O2 and helped set its new level.
In the redox chemistry of the atmosphere, hydrogen-bearing reduced gases,(such CH4 and H2) are the chemical adversaries of O2. When the concentration of one
goes up, the other declines, and vice-versa. Photochemical reactions cause rapid
mutual annihilation between O2 and the reducing gases. In general, we expect theredox chemistry of the pre-GOE atmosphere to have been dominated by methane andhydrogen (even though they were both probably at levels below a percent), whereas
O2 would have been a trace gas at less than 1 ppmv. It is opposite to the redox control
in the modern atmosphere. Models suggest that the GOE occurred when the rate ofsupply rate of supply of reduced gases to the atmosphere dropped below the O2 flux tothe atmosphere (Claire et al., 2006; Goldblatt et al., 2006; Holland, 2009).
The GOE appears in the rock record at the same time as global cooling. In
detailed models of the GOE, there is a decline in the supply of reducing gases prior to
the GOE, which means that the methane concentration in the pre-GOE atmospherefalls to low levels even before O2 rises to a substantial concentration (Zahnle et al.,2006). The loss of greenhouse gases such as methane and its photochemical product,ethane, are plausibly responsible for global cooling at ~2.4 Ga. It is possible that the
global climate and atmospheric composition oscillated until conditions of permanent
oxygenation were established, which would account for multiple glaciations during
2.45-2.22 Ga.
Moderate levels of O2 established in the Proterozoic (probably somewhere
between 0.2 and 2% O2 by volume) were still sufficient to establish a stratosphericozone (O3) layer (Kasting and Donahue, 1980) and protect the Earth’s surface fromharmful ultraviolet in the range 200-300 nm. A counter-intuitive effect of the ozone
layer was to decrease the net reaction rate in the troposphere between methane andoxygen. Provided that methane fluxes were still relatively large from a poorly
oxygenated and euxinic Proterozoic ocean, methane should have risen after the GOE
to be a moderately significant greenhouse gas during middle Proterozoic.

26

Acknowledgments
The author thanks Shawn Domagal-Goldman for improving the manuscript andNASA Astrobiology/Exobiology Program grant NNX10AQ90G for support.
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34

Figure Captions.

Figure 1 Hydrogen escape from the atmosphere is a directly observable phenomenonon the modern Earth. Above is an ultraviolet image of the Earth's dark hemispherewith the Sun behind it, taken from NASA’s Dynamic Explorer 1 spacecraft at 19,700
km altitude above 13°N, on February 16, 1982. The extended glow around the planet
comes from hydrogen atoms that scatter Lyman-α radiation from the Sun. A northern
auroral oval and equatorial glow are caused by emission from atomic oxygen andmolecular nitrogen. Isolated points are stars that are bright in the ultraviolet.
(Courtesy of NASA).
Figure 2 A conceptual diagram showing “bottlenecks” for hydrogen escape from the
Earth’s atmosphere. At the tropopause, most water vapor condenses and the dominant
form of hydrogen just above the tropopause is in the form of methane, even on themodern Earth. Hydrogen in all its forms is mixed up through the stratosphere andmesosphere but its transport is then rate-limited by diffusion through thethermosphere to the boundary with space, the exobase. By the time the hydrogen getsto the exobase, it is mostly in atomic form.
Figure 3 Schematic diagram showing the long-term sources and sinks of oxygen
fluxes.
Figure 4 The inventory of oxygen in the Earth’s crust. “Hard rocks” are high grade
metamorphic rocks and igneous rocks. “Soft rocks” are sedimentary rocks. The
imbalance between oxygen and reduced carbon shows that there is missing reductant.
(Based on tabulated values in Catling et al. (2001) and references therein).
Figure 5 Results from the box model of Claire et al. (2006). A,B: The upper row
shows the results when the model assumes that metamorphic reducing gases are theprimary sink on Archean O2. C,D: The lower row shows results where the model
assumes that volcanic gases are the main O2 sink. Qualitatively, there is no difference
to the results from either assumption. Fluxes are as follows: FW = continentaloxidative weathering, FE = hydrogen escape to space, FV = volcanic sink, and FM =metamorphic sink. KOXY is the oxygenation parameter defined in the text, whichreaches unity when the atmosphere goes through its “Great Oxidation Event” (GOE).
Figure 6 Atmospheric fluxes of chemical species as a function of the ground-level O2
mixing ratio. Mixing ratio is equivalent to a partial pressure in bars for a 1 bar
atmosphere. Results are shown from a photochemical “thought experiment” where theground-level CH4 mixing ratio is fixed to 100 ppmv and volcanic outgassing fluxes ofsulphur gases and reducing gases are also fixed (see text). Plus symbols (“+”) map to
the right-hand vertical axis and show the net ratio of the biogenic O2 flux (φO2) to CH4flux (φCH4) as a function of the ground-level O2 mixing ratio. The shaded regions
represent cases when the biogenic O2 fluxes are either implausibly high or low. (From
Catling et al. (2007)).

35
Figure 7 The ozone (O3) column abundance shown as the dashed line as a function of
the concentration of ground-level atmospheric oxygen. The O2 level is expressed as aratio to the Present Atmospheric Level (PAL). Today’s typical ozone column
abundance is shown as a dotted line. The ozone layer’s protective absorption of
biologically harmful ultraviolet radiation becomes significant at ~10-2 PAL.

36

Figure 1 Hydrogen escape from the atmosphere is a directly observable phenomenonon the modern Earth. Above is an ultraviolet image of the Earth's dark hemispherewith the Sun behind it, taken from NASA’s Dynamic Explorer 1 spacecraft at 19,700
km altitude above 13°N, on February 16, 1982. The extended glow around the planet
comes from hydrogen atoms that scatter Lyman-α radiation from the Sun. A northern
auroral oval and equatorial glow are caused by emission from atomic oxygen andmolecular nitrogen. Isolated points are stars that are bright in the ultraviolet.
(Courtesy of NASA).

37

Figure 2 A conceptual diagram showing “bottlenecks” for hydrogen escape from the
Earth’s atmosphere. At the tropopause, most water vapor condenses and the dominant
form of hydrogen just above the tropopause is in the form of methane, even on themodern Earth. Hydrogen in all its forms is mixed up through the stratosphere andmesosphere but its transport is then rate-limited by diffusion through thethermosphere to the boundary with space, the exobase. By the time the hydrogen getsto the exobase, it is mostly in atomic form.

38

Figure 3 Schematic diagram showing the long-term sources and sinks of oxygen
fluxes.

39

Figure 4 The inventory of oxygen in the Earth’s crust. “Hard rocks” are high-grade
metamorphic rocks and igneous rocks. “Soft rocks” are sedimentary rocks. The
imbalance between oxygen and reduced carbon shows that there is missing reductant.
(Based on tabulated values in Catling et al. (2001) and references therein).

40

Figure 5 Results from the box model of Claire et al. (2006). A,B: The upper row
shows the results when model assumes that metamorphic reducing gases are theprimary sink on Archean O2. C,D: The lower row shows results where the model
assumes that volcanic gases are the main O2 sink. Qualitatively, there is no difference
to the results from either assumption. Fluxes are as follows: FW = continentaloxidative weathering, FE = hydrogen escape to space, FV = volcanic sink, and FM =metamorphic sink. KOXY is the oxygenation parameter defined in the text, whichreaches unity when the atmosphere goes through its “Great Oxidation Event” (GOE).

41

Figure 6 Atmospheric fluxes of chemical species as a function of the ground-level O2
mixing ratio. Mixing ratio is equivalent to a partial pressure in bars for a 1 bar
atmosphere. Results are shown from a photochemical “thought experiment” where theground-level CH4 mixing ratio is fixed to 100 ppmv and volcanic outgassing fluxes ofsulphur gases and reducing gases are also fixed (see text). Plus symbols (“+”) map to
the right-hand vertical axis and show the net ratio of the biogenic O2 flux (φO2) to CH4flux (φCH4) as a function of the ground-level O2 mixing ratio. The shaded regions
represent cases when the biogenic O2 fluxes are either implausibly high or low. (From
Catling et al. (2007)).

42

Figure 7 The ozone (O3) column abundance shown as the dashed line as a function of
the concentration of ground-level atmospheric oxygen. The O2 level is expressed as aratio to the Present Atmospheric Level (PAL). Today’s typical ozone column
abundance is shown as a dotted line. The ozone layer’s protective absorption of
biologically harmful ultraviolet radiation becomes significant at ~10-2 PAL.