Precambrian Research xxx (2006) xxx–xxx
Precambrian paleosols and atmospheric CO2 levels
Nathan D. Sheldon ∗
Department of Geological Sciences, Royal Holloway University of London, Egham, Surrey TW20 OEX, United Kingdom
Received 3 November 2005; received in revised form 13 February 2006; accepted 19 February 2006
Abstract
Precambrian atmospheric pCO2 levels were previously estimated using a simple thermodynamic model based on the mineral
assemblage in paleosols. In addition to theoretical flaws with that model, recalculation of pCO2 levels using more recent thermodynamic data gives significantly lower values. A new model based on the mass balance of weathering in paleosols gives consistent
results from three separate ∼2.2 Ga old paleosols. The calculated pCO2 value of 23×3
÷3 times present atmospheric levels is insufficient
to overcome the “faint young Sun” paradox, suggesting that another greenhouse gas such as CH4 was present at elevated levels
relative to the present atmosphere to account for the discrepancy. Applying this new approach to other Precambrian paleosols a
Proterozoic pCO2 curve is generated, which indicates consistently high pCO2 from 2.5 to 1.8 Ga ago, and a substantial drop in
atmospheric pCO2 at some time between 1.8 and 1.1 Ga ago.
© 2006 Elsevier B.V. All rights reserved.
Keywords: Precambrian; Paleosols; CO2 ; CH4 ; Proterozoic
1. Introduction
There has been a large amount of speculation about
the composition of the Earth’s Precambrian atmosphere
based on unusual sedimentary deposits such as banded
iron formations (BIFs) and possibly detrital uraninite and
pyrite (for contrasting views see Minter, 1999; Phillips
and Law, 1997), as well as on the basis of deposits such
as paleosols (e.g., Holland and Zbinden, 1988) that are
not restricted to the Precambrian. When they have not
been pervasively altered (e.g., by potassium metasomatism; Palmer et al., 1989), paleosols offer the most direct
evidence of past environmental conditions because they
formed at the Earth’s surface, in direct contact with the
atmosphere. Holland (1994) and his colleagues (Pinto
and Holland, 1988; Holland and Zbinden, 1988; Holland
et al., 1989) have suggested semi-quantitative methods
for estimating Precambrian oxygen and carbon dioxide
∗
Tel.: +44 1784 443615; fax: +44 1784 471780.
E-mail address: [email protected].
levels based on a number of simplifying assumptions.
Estimates of atmospheric carbon dioxide levels before
∼2.2 Ga ago were further constrained by a simple thermodynamic model, based on greenalite–siderite equilibrium, to have been less than 100 times present levels
(Rye et al., 1995), an amount insufficient to account for
the “faint young Sun” paradox (Kasting, 1993), which
has led recent investigators to propose that elevated
methane levels may have provided the additional greenhouse component necessary to maintain equable surface
temperatures throughout Earth’s history (Pavlov et al.,
2000, 2003). In this study, I examine the shortcomings
of the simple thermodynamic model, propose an alternative method of calculating atmospheric carbon dioxide
concentration using mass balance in paleosols, and generate a pCO2 curve for 2.5–1.0 Ga ago.
2. CO2 Paleobarometry
Rye et al. (1995) suggested that the equilibrium relationship between siderite and reduced iron silicates (SIS
0301-9268/$ – see front matter © 2006 Elsevier B.V. All rights reserved.
doi:10.1016/j.precamres.2006.02.004
PRECAM-2670;
No. of Pages 8
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N.D. Sheldon / Precambrian Research xxx (2006) xxx–xxx
herein) in paleosols could be used to constrain Precambrian atmospheric pCO2 . They argued that the lack
of siderite in Precambrian paleosols limits atmospheric
pCO2 levels since reduced iron silicates should precipitate instead of siderite if pCO2 is low. Greenalite was
selected by Rye et al. (1995) because it is a simple
low-temperature metamorphic iron silicate suitable to
Precambrian paleosols, many of which have been altered
by greenschist facies metamorphism at T ≥ 300 ◦ C (Rye
and Holland, 2000a,b). Siderite is not decarbonated
under common conditions of metamorphism up to, and
including, greenschist facies metamorphism (Rye et al.,
1995).
There are a number of theoretical and geological
issues with the simple model of Rye et al. (1995)
for determining Archean atmospheric pCO2 from paleosols. First, using a single equilibrium relationship
for all Archean and Proterozoic paleosols is misleading because the mineral assemblage present is not the
same in all Precambrian paleosols, and those differences
would give different equilibrium results. The presence or
absence of siderite is only relevant in cases where the rest
of the modeled mineral assemblage is present. Second,
greenalite is not known to form authigenically in soils,
and it would be better to apply an authigenic iron silicate
such as berthierine rather than a mineral in the post-burial
metamorphic assemblage. Third, Rye et al. (1995) measured in situ carbon dioxide in banded iron formations
(BIFs) as a check on their thermodynamic model results.
This provides only weak guidance because the chemistry
of a soil formed at the Earth’s surface bears little resemblance to the chemistry of a BIF formed in the ocean,
in part because decomposition of organic matter leaves
the effective pCO2 of the deep ocean about three times
that of the surface ocean (Kasting, 2004). Fourth, even
when a more likely mineral assemblage is used in the calculations including the reduced iron silicate berthierine,
which is present in the ∼2.2 Ga Waterval Onder paleosol, estimated carbon dioxide levels are inconsistent
with geologic evidence as described below. Fifth, the
oldest occurrence of siderite in a paleosol, which the SIS
paleobarometer would indicate formed under conditions
of atmospheric CO2 greater than 60–100 times present
levels, is ∼0.98 Ga, a time for which there is no independent evidence of CO2 levels that high (e.g., Retallack
and Mindszenty, 1994), suggesting siderite formation
may be tied to other factors such as productivity by
soil biota (e.g., Sawicki et al., 1995) rather than atmospheric composition. This idea is further supported by
anoxic weathering experiments where no siderite was
formed from the dissolution of Fe-biotite even at very
high pCO2 levels (1 atm; >2000 PAL) where Rye et al.’s
(1995) model would predict siderite stability (Murakami
et al., 2004).
3. Equilibrium model results
The original SIS CO2 paleobarometer gives an upper
limit of ∼100 times present atmospheric CO2 levels
based on the following equilibrium between siderite and
greenalite:
Fe3 Si2 O5 (OH)4(green) + 3CO2(g)
= 3FeCO3(sid) + 2SiO2(aq) + 2H2 O
(1)
Rye et al. (1995) used thermodynamic data from
Eugster and Chou (1973) and calculated the phase
boundary shown in Fig. 1. In their paper, Rye et al.
(1995) allowed that Eugster and Chou’s (1973) data for
greenalite were poorly constrained. Recalculating the
SIS paleobarometer using revised data for greenalite
from the computer program SUPCRT92 (Johnson et al.,
1992; database updated as SLOP98 by E. Shock) gives a
significantly different result (Fig. 1), with Precambrian
pCO2 levels estimated to be only significantly lower
than present atmospheric pCO2 levels (PAL) over a
reasonable range of aSiO2(aq) . Most modern surface
waters are near quartz saturation (aSiO2(aq) = 10−4.047 ),
which, if it was also the case in the Precambrian, would
indicate that atmospheric pCO2 was just 10−5.72 , or
0.006 PAL (Fig. 1). As discussed above greenalite is
unlikely to form authigenically in soils, so perhaps the
SIS paleobarometer approach is sound, but greenalite is
the wrong reduced iron silicate.
Berthierine (thermodynamic data from Fritz and Toth,
1997), a serpentine group ferrous iron silicate that forms
in chemically reducing conditions but rarely as a soil
mineral (e.g., Sheldon and Retallack, 2002), is part of
the mineral assemblage of the ∼2.2 Ga old Waterval
Fig. 1. SIS paleobarometer phase diagram for 25 ◦ C. Calculated pCO2
phase boundaries vary little over a plausible range for Earth surface
temperatures (10–30 ◦ C).
N.D. Sheldon / Precambrian Research xxx (2006) xxx–xxx
Onder paleosols (Retallack and Krinsley, 1993). A similar expression to Eq. (1) can be written for equilibrium
between berthierine and siderite and kaolinite as follows:
Fe2 Al2 SiO5 (OH)4(bert) + SiO2(aq) + 2CO2(g)
= 2FeCO3(sid) + Al2 Si2 O5 (OH)4(kaol)
(2)
where kaolinite is inferred to have been part of the
pre-metamorphic mineral assemblage (Retallack and
Krinsley, 1993; Rye and Holland, 2000a,b). Calculated
pCO2 levels are higher than for the recalculated SIS paleobarometer of Rye et al. (1995), and close to present day
values (Fig. 1). Furthermore, the berthierine at Waterval Onder is not a pure end member, thus it would have
an activity of less than one, further lowering the calculated pCO2 level. The choice of Al-mineral is relatively
unimportant as calculations based on gibbsite instead of
kaolinite (not shown) also result in pCO2 values below
present day conditions.
Equilibrium modeling using greenalite or berthierine SIS paleobarometers suggests one of two conclusions, either Precambrian pCO2 levels were comparable
to present atmospheric values, or the whole approach
is unreliable. Given the “faint young Sun” paradox
(Kasting, 1993), such low levels are difficult to reconcile
with the need for a much stronger than present Precambrian greenhouse. Furthermore, by simplifying atmospheric composition down to a single chemical reaction
there is a risk of over-interpreting the absence of evidence of a single mineral phase, in this case siderite.
4. Other equilibrium models
In addition to the widely cited approach of Rye et
al. (1995), two other recent studies have attempted to
estimate atmospheric pCO2 using equilibrium relationships and the Rye et al. (1995) study as a “jumping
off point”. Ohmoto et al. (2004) inverted the logic of
Rye et al. (1995) and proposed that massive siderite
beds in BIFs prior to ∼1.8 Ga imply that atmospheric
pCO2 must have greatly exceeded 100 PAL. They further suggest that atmospheric pCO2 may have been high
enough that no additional greenhouse gas was necessary
because the inhibition on siderite formation was elevated
pO2 values in the atmosphere and soil rather than insufficient pCO2 (Ohmoto et al., 2004). There are a few
problems with both the approach and the conclusions.
First, Ohmoto et al. (2004) assert that siderite stability
requires pO2 = 10−60±5 , a value well below any existing estimate for the Precambrian, but as Sleep (2004)
correctly points out, even at the naturally maintained
pO2 minima (pO2 = 10−13 ) reaction kinetics dictate that
3
oxidation processes are unlikely to inhibit siderite formation. Furthermore, even at the present day pO2 value
of 0.21, siderite formation in near-modern soils is not
restricted (e.g., Achyuthan, 2003) to poorly aerated settings as Ohmoto et al. (2004) asserted. Thus, even if
pO2 levels were higher than most researchers believe (as
Ohmoto et al., 2004 believe), siderite should form at the
Earth’s surface if pCO2 is significantly elevated. Second,
BIFs can only provide weak guidance for Earth surface
(and hence atmospheric) conditions because they form
in at least moderately deep water (i.e., below storm wave
base) and should have enhanced pCO2 relative to the surface as a result of the decomposition of organic matter
(Kasting, 2004).
A second thermodynamic approach set a lower limit
for atmospheric pCO2 by considering equilibrium relationships for rinds on sub-aerially deposited pebbles
from the 3.2 Ga Moody Conglomerate (Hessler et al.,
2004). Hessler et al. (2004) considered three possible
equilibrium relationships that might apply to the formation of the carbonate rinds, which, in contrast to the
other SIS paleobarometers, involve no iron silicate rather
than no siderite. Their calculated minimum pCO2 of
2.51 × 10−3 bar (∼6.8 PAL) requires an unusual geologic setting in which iron silicates are very unstable
at the Earth’s surface, whereas if iron silicates were stable the estimated pCO2 minimum is significantly higher,
though still insufficient to account for the “faint young
Sun”. Without more investigation into Archean weathering conditions it is not clear which of their two scenarios
is most likely. It is also unclear how common siderite
coated pebbles (sans iron silicates) are in the geologic
record, and thus how often they may provide a pCO2 constraint. However, where appropriate, the Hessler et al.
(2004) thermodynamic model appears to provide a useful constraint. Thermodynamic models in general oversimplify the problem of estimating Precambrian pCO2
levels, and with slightly different starting assumptions
(e.g., SIS paleobarometers of Rye et al. (1995), Ohmoto
et al. (2004), and this paper) or thermodynamic data
(e.g., this paper), vastly different conclusions may be
drawn. Extreme caution is thus advocated in their application, because even the mostly widely cited one (Rye
et al., 1995) has a number of fundamental flaws, and
other approaches (Hessler et al., 2004) may have only
limited use if they apply to relatively rare sedimentary
occurrences.
5. An alternative approach using mass balance
One common method of assessing gains and losses
of different elements in soils and paleosols is through
4
N.D. Sheldon / Precambrian Research xxx (2006) xxx–xxx
constitutive mass balance (e.g., Brimhall and Dietrich,
1987; Chadwick et al., 1990; Brimhall et al., 1991).
Mass balance can be reduced to two concepts, strain (ε)
of an “immobile” element and transport (τ) of a second element with respect to the immobile element. The
open system mass-transport function for element j in the
weathered sample (w) is defined as follows:
ρw Cj,w
τj,w =
[εi,w + 1] − 1.
(3)
ρp Cj,p
where ρw is the density of the weathered material, Cj,w
the chemical concentration (weight percentage) of element j in the weathered material, ρp the density of the
parent material, and Cj,p is the chemical concentration
(weight percentage) of element j in the parent material.
If τ j,w = 0 (i.e., element j was immobile), then εi,w can
be solved for as follows:
ρp Cj,p
εi,w =
− 1.
(4)
ρw Cj,w
where εi,w is the strain on immobile element i in the
weathered sample.
During silicate weathering, total CO2 consumption
can be approximated by the following reactions:
MgO + 2CO2 + H2 O → Mg2+ + 2HCO3 −
CaO + 2CO2 + H2 O → Ca
2+
+ 2HCO3
−
(5)
(6)
Na2 O + 2CO2 + H2 O → 2Na+ + 2HCO3 −
(7)
K2 O + 2CO2 + H2 O → 2K+ + 2HCO3 −
(8)
where for example, each mole of base cation (e.g., MgO)
liberated requires two moles of CO2 (e.g., Holland and
Zbinden, 1988). Although other elements are weathered,
the parent rock concentration of the bases in reactions
5–8 are at least 1–2 orders of magnitude larger than
any other mobile element. Because K may be remobilized metasomatically and accumulated in paleosols after
burial (e.g., Palmer et al., 1989), it has to be dealt with
differently than Ca, Mg, and Na (see below).
Mass transfer values (Eq. (3)) for individual cations
can be transformed into mass fluxes as follows:
Cj,p Z=Dj,w
mj,flux (g/cm2 ) = ρp
τj,w(z) dZ
(9)
100 Z=0
where Z is the depth in the soil profile and Dj,w is the
total depth of the profile (e.g., Chadwick et al., 1990).
Because many Precambrian paleosols have been deeply
buried and compacted (e.g., Retallack, 1986; Holland
and Zbinden, 1988), it is necessary to decompact the
profiles to their original thickness rather than using their
present thickness. Sheldon and Retallack (2001) derived
the following equation for calculating the compaction
(C) of soils and paleosols due to burial:
C= −Si
,
−1
(10)
F0
eDk
where Si is initial solidity, F0 initial porosity, D the depth
of burial in km, and k is an empirically derived constant.
Each of the variables has been determined for different soil orders, and using the constants for a Vertisol
(Retallack, 1986; Retallack and Krinsley, 1993) the calculated compaction of the ∼2.2 Ga old Waterval Onder
paleosol was 72% of the profile’s original thickness,
while direct measurement of folded veins gave 67–73%
compaction (Sheldon and Retallack, 2001), suggesting
that Eq. (10) gives suitable estimates for the compaction
of even deeply buried Precambrian paleosols.
After the compaction-corrected mass fluxes are calculated for each individual base cation and converted to
moles, the total flux of CO2 required for the observed
weathering (M) is given by
M(mol CO2 /cm2 ) = 2
mj,flux .
(11)
Soils do not form instantaneously, so to calculate a true
flux, M must be divided by time (T). The time-averaged
flux (M/T) is a product of two distinct sources of CO2 ,
CO2 in the atmosphere being added by rainfall to the
soils (Xrain ), and CO2 added by direct diffusion into the
soils (Xdiff ). Holland and Zbinden (1988) quantified Xrain
and Xdiff as follows:
M
(mol/cm2 year)
T
= Xrain + Xdiff ≈ pCO2
KCO2 r
DCO2 α
+κ
L
103
(12)
where pCO2 is the partial pressure of atmospheric
CO2 (atm), KCO2 the Henry’s Law constant for CO2
(discussed below), r rainfall rate (cm/year), DCO2 the
diffusion constant for CO2 in air (0.162 cm2 /s; CRC
Handbook, 1995), α the ratio of diffusion constant
for CO2 in soil divided by the diffusion constant
for CO2 in air (discussed below), L the depth to the
water table, and κ is a constant which is the ratio
of seconds in a year divided by the number of cm3
per mole of gas at standard temperature and pressure
(1.43 × 103 (s cm3 )/(mol year)). Implicit in Eq. (12) is
the assumption that the partial pressure of atmospheric
CO2 is much larger than the partial pressure of CO2 at
the water table (depth = L) and that the CO2 diffusion
constant and gradient are constant with depth. There
also is no explicit term including any CO2 processes
involving a terrestrial biosphere (i.e., it is assumed to
N.D. Sheldon / Precambrian Research xxx (2006) xxx–xxx
play a negligible role), a point that will be discussed
further below. Eq. (12) can be rearranged to solve for
atmospheric pCO2 as follows:
T
K
CO2 r
103
M
+κ
DCO2 α
L
= pCO2 .
(13)
Thus, by quantifying M using direct measurements of
paleosol mass balance and estimating T and L, it is
possible to calculate the partial pressure of atmospheric
CO2 at the time the paleosols formed. One substantial
advantage of this approach over the thermodynamic
approaches discussed above it that it is not redox sensitive (i.e., atmospheric pO2 does not matter), removing
one of the primary sources of uncertainty in all of the
SIS paleobarometers.
6. pCO2 levels 2.2 Ga ago
To test this approach, atmospheric pCO2 values have
been calculated for a variety of model conditions using
data for the Hekpoort paleosol from Rye and Holland
(2000a), who analyzed three separate cores through
the paleosol. Though these “profiles” are time equivalents to the widely studied Waterval Onder profile (e.g.,
Retallack, 1986; Retallack and Krinsley, 1993; Rye and
Holland, 2000a; Beukes et al., 2002), the completeness
of that profile has recently been questioned (Beukes et
al., 2002), so it has not been used here. Using the data
of Rye and Holland (2000a) also allows the robustness
of the method to be examined since results from three
separate profiles of the same paleosol can be compared.
Basalt-parented paleosols are ideal for this type of calculation because the depth of the water table is essentially
the thickness of the soil given the relative porosity of
basalt as compared to most soils.
A comprehensive discussion of the sources of uncertainty in this model can be found in the Data repository,
but a summary of the sensitivity of the various parameters can be given as follows (Table 1): (1) changing
the Henry’s Law constant over a range of values from
0.031 to 0.045 mol atm−1 results in a difference between
the highest and lowest pCO2 values of less than 5%,
so 0.034 mol atm−1 is used for all of the calculations;
(2) changing α by ±20% results in pCO2 variations of
∼15%; the estimate of 0.1 from Holland and Zbinden
(1988) is used as the “best guess”; (3) including the
weathering of K2 O by assuming that it would have the
same τ j,w values as Na2 O, which behaves chemically in a
fairly similar manner, the calculated pCO2 varies by less
than 5%, but this parameter has been included for completeness; (4) over the range of semi-arid to sub-tropical
5
Table 1
Parameters and sensitivities
Parameter
Range
Resulting pCO2
variation (%)
KCO2
α
Include K2 O
r
0.031–0.045 (0.034)
0.08–0.12 (0.1)
n.a.
50–150 (100)
5>
∼15
5>
15>
The parenthetical value is the one used calculations discussed in
the text. The other constant values used was DCO2 = 0.162 (CRC
Handbook, 1995); L was variable between paleosols and based on
the decompacted depth of the profile (BB3 = 1079 cm, BB8 = 1186 cm,
BB14 = 810 cm).
mean annual precipitations (50–150 cm/year), the resulting pCO2 variation is less then 15%, so the present
day global mean annual precipitation of ∼100 cm/year
(Holland and Zbinden, 1988) is used because it is difficult to assess how that value may have varied in the
Precambrian.
The largest difference in the calculated pCO2 values results from lengthening the formation time from
10,000 to 50,000 years, a range of typical values for
thick soils forming on basalt (Sheldon, 2003). Because
of the age of the Hekpoort paleosols, it is not possible to
distinguish the age of the overlying basalt flow from the
flow that underlies the paleosols, so the formation time
could be have been significantly longer than the range
of values examined here (which would decrease calculated pCO2 levels; Data repository figure DR-1). Based
on near-modern analogues and the thickness of the profiles (Sheldon, 2003), 50, 000×2
÷2 years is the best guess
for the formation time of the Hekpoort paleosols.
Using the “best guess” values for each of the parameters, the resulting pCO2 values vary between 18 and
31 times present levels (PAL) for the three Hekpoort
paleosols. As discussed above, the largest source of
uncertainty is formation time of the paleosols, which
was probably within a factor of two of the “best guess”.
Following Holland and Zbinden, and to be conservative, if the other lesser sources of uncertainty add half
again as much uncertainty, pCO2 2.2 Ga ago is estimated
to have been at least 23×3
÷3 PAL. This value is significantly lower than upper limit of 100 PAL calculated by
Rye et al. (1995), but higher than revised estimates from
SIS paleobarometers, and also would be insufficient to
account for the reduced insolation of the “faint young
Sun” (Kasting, 1993). The likely solution to this CO2
“greenhouse deficit” is that another greenhouse gas must
have been present at higher levels than at present. The
best candidate is CH4 (Pavlov et al., 2000, 2003), a trace
gas in the present atmosphere that is a >20 times more
6
N.D. Sheldon / Precambrian Research xxx (2006) xxx–xxx
effective greenhouse gas than CO2 , though this ratio
may have been different in a low-O2 Precambrian atmosphere. Nonetheless, mass balance calculations using
paleosols also indicate that atmospheric CO2 levels were
high and that calculations assuming CO2 levels for the
Precambrian just a few times pre-industrial levels (e.g.,
Pavlov et al., 2000) may overestimate CH4 levels somewhat.
A final consideration is what role, if any, an emergent biosphere may have played in modulating soil pCO2
levels. The calculations described herein assume that
there was no terrestrial biosphere or that it played a
negligible role. However, modern soils may have soil
pCO2 values up to 100 times the atmospheric pCO2
level (Brook et al., 1983) due to in situ production of
CO2 and slow diffusion. Theoretical calculations indicate that even a much smaller biosphere than at present,
given the correct soil substrate, could induce modern
levels of weathering (Keller and Wood, 1993). Thus, the
key question is whether there was a significant terrestrial
biosphere to speak of. The answer at present is equivocal.
There is some evidence for Precambrian soil methanogenesis/methanotrophy (Watanabe et al., 2000; Rye and
Holland, 2000b), but it is not clear yet if these phenomena were widespread or confined to local niches. Brady et
al. (1999) found that basalt weathering intensity is 2–18
times greater under lichens than in abiotic experiments,
although confined to a very shallow depth (mm-scale).
Using this somewhat inexact analogue for a Precambrian
biosphere as a guide suggests that while there may have
been some in situ enhancement of soil CO2 levels relative
to the atmosphere, it was likely only to a shallow depth,
and perhaps only a locally important process. Without
additional insight into the Precambrian terrestrial biosphere it is not possible to quantify what part of the
calculated atmospheric pCO2 levels were due to biotic
enhancement, but given the other sources of uncertainty
in the calculations it is unlikely that the results would be
fundamentally different.
7. Toward a Precambrian CO2 curve
A further test of this approach is to analyze other Precambrian paleosols both to see if the values obtained
are consistent with one another and also with independent geological evidence. Using data from the
Ville Marie (2.5 Ga; Panahi et al., 2000), Drakenstein (2.0 Ga; Wiggering and Beukes, 1990), Flin Flon
(1.8 Ga; Holland et al., 1989), Sturgeon Falls (1.1 Ga;
Zbinden et al., 1988), and Sheigra (0.98 Ga; Retallack
and Mindszenty, 1994) paleosols, a pCO2 curve for
2.5–1.0 Ga ago can be generated (Fig. 2). Two conclu-
Fig. 2. Atmospheric pCO2 2.5–1.0 Ga ago. The timing and rate of the
drop in pCO2 between 1.8 and 1.0 Ga ago are unconstrained by existing data. The shaded region approximates the range (upper and lower
limits) of CO2 -only solutions for a habitable Earth from Kasting (1987,
1993). While the lower part of Kasting’s (1987, 1993) error envelope
intersects the upper part of the error envelope in this study in places,
the “best guess” scenarios remain substantially different. For example,
at 2.5 Ga ago, the “best guess” of Kasting (1987, 1993) is 0.15 bar of
CO2 (>400 PAL), whereas the new paleosol “best guess” results predict
just 0.009 bar of CO2 (∼25 PAL). Thus, an additional greenhouse gas
such as CH4 is necessary to provide the addition greenhouse forcing
to maintain a habitable planet.
sions can be drawn: (1) atmospheric pCO2 was fairly
constant and elevated from 2.5 to 1.8 Ga ago; (2) at some
time between 1.8 and 1.1 Ga ago atmospheric pCO2
dropped significantly. Without additional data, it is not
possible to constrain when this occurred, but this drawdown was subsequently reversed at some time in the
late Precambrian given that early Paleozoic pCO2 levels
appear to have been comparable to the Paleoproterozoic
(Fig. 2; Berner and Kothavala, 2001). Some further guidance in answering the question of what happened in the
interim is available from estimates of pCO2 based on
microfossils. Kaufman and Xiao (2003) examined the
␦13 C composition of ∼1.4 Ga old microfossils and estimated atmospheric pCO2 of >10–200 PAL. The low end
of that range is compatible with the pCO2 curve presented here, with the high end lying just outside the error
envelope of the paleosol-based curve, suggesting a general correspondence between the two methods, though
further work is needed to confirm the estimate from the
microfossils and to find paleosols in this time interval
that can be compared directly.
8. Conclusions
Attempts to calculate the Earth’s atmospheric conditions before ∼2.2 Ga ago are fraught with difficulties
because the Precambrian rock record is sparse when
compared with that of the Phanerozoic. In spite of this,
many attempts have been made to estimate the com-
N.D. Sheldon / Precambrian Research xxx (2006) xxx–xxx
position of the atmosphere because reconstruction of
early atmospheric pCO2 is a crucial component of any
discussion of the Earth’s surface temperature history,
the “faint young Sun” paradox, and history of life on
Earth. Paleosols, which formed in direct contact with
atmosphere, have been used in concert with a simple siderite–greenalite thermodynamic model to suggest
that atmospheric pCO2 levels have always been below
60–100 times the present atmosphere (Rye et al., 1995).
In addition to theoretical problems with that model,
recalculation of pCO2 levels using revised thermodynamic data gives unrealistically low atmospheric pCO2 .
Furthermore, it is almost certainly an oversimplification
to calculate atmospheric pCO2 on the basis of a single
equilibrium expression. A new model based upon mass
balance gives pCO2 levels that can be constrained to be
at least 23×3
÷3 times present atmospheric levels for 2.2 Ga
ago. These pCO2 levels are still insufficient to make up
for a “faint young Sun” (Fig. 2), suggesting that some
other greenhouse gas such as methane was present and
provided enough additional greenhouse forcing to maintain equable conditions at the Earth’s surface throughout
its history. Atmospheric pCO2 levels appear to have varied little from 2.5 to 1.8 Ga ago, staying above 20 times
present levels. At some time between 1.8 and 1.0 Ga ago,
pCO2 dropped down to much lower levels, but the timing
and rate of this change cannot be constrained by existing
data.
Acknowledgments
Thanks are owed Mark Reed for numerous discussions of this manuscript, and Rob Rye for discussing the
Bank Break cores at an early stage of this project. An
earlier version of this manuscript benefited from comments by Dick Holland and an anonymous reviewer, and
this version benefited from three thoughtful anonymous
reviews.
Appendix A. Supplementary data
Supplementary data associated with this article can be
found, in the online version, at doi:10.1016/j.precamres.
2006.02.004.
References
Achyuthan, H., 2003. Petrologic analysis and geochemistry of the late
Neogene-early Quaternary hardpan calcretes of western Rajasthan,
India. Quaternary Int. 106/107, 3–10.
Berner, R.A., Kothavala, Z., 2001. GEOCARB III: a revised model of
atmospheric CO2 over Phanerozic time. Am. J. Sci. 301, 182–204.
7
Beukes, N.J., Dorman, H., Gutzmer, J., Nedachi, M., Ohmoto, H.,
2002. Tropical laterites, life on land, and the history of atmospheric
oxygen in the Paleoproterozoic. Geology 30, 491–494.
Brady, P.V., Dorn, R.I., Brazel, A.J., Clark, J., Moore, R.B., Glidewell,
T., 1999. Direct measurements of the combined effects of lichen,
rainfall, and temperature on silicate weathering. Geochim. Cosmochim. Acta 63, 3293–3300.
Brimhall, G.H., Lewis, C.J., Ford, C., Bratt, J., Taylor, G., Warin, O.,
1991. Quantitative geochemical approach to pedogenesis: importance of parent material reduction, volumetric expansion, and
eolian influx in laterization. Geoderma 51, 51–91.
Brimhall, G.H., Dietrich, W.E., 1987. Constitutive mass balance relations between chemical composition, volume, density, porosity,
and strain in metasomatic hydrochemical systems: results on
weathering and pedogenesis. Geochim. Cosmochim. Acta 51,
567–587.
Brook, G.A., Folkoff, M.E., Box, E.O., 1983. A world model of carbon
dioxide. Earth Surf. Processes Landforms 8, 79–88.
Chadwick, O.A., Brimhall, G.H., Hendricks, D.M., 1990. From a black
to a gray box — a mass balance interpretation of pedogenesis.
Geomorphology 3, 369–390.
CRC Handbook, 1995. Lide, D.R. (Ed.), Handbook of Chemistry and
Physics, 76th ed. CRC Press.
Eugster, H.P., Chou, I.M., 1973. The depositional environments of
Precambrian banded iron-formations. Econ. Geol. Bull. Soc. Econ.
Geol. 68, 1144–1168.
Fritz, S.J., Toth, T.A., 1997. An Fe-berthierine from a Cretaceous
laterite. Part II. Estimation of Eh, pH and pCO2 conditions of formation. Clays Clay Miner. 45 (4), 580–586.
Hessler, A.M., Lowe, D.R., Jones, R.L., Bird, D.K., 2004. A lower
limit for atmospheric carbon dioxide levels 3.2 billion years ago.
Nature 428, 736–738.
Holland, H.D., 1994. Early Proterozoic atmospheric change. In: Bengtson, S. (Ed.), Early Life on Earth: Nobel Symposium No. 84.
Columbia University Press, New York.
Holland, H.D., Feakes, C.R., Zbinden, E.A., 1989. The Flin Flon paleosol and the composition of the atmosphere 1.8 BYBP. Am. J. Sci.
289, 362–389.
Holland, H.D., Zbinden, E.A., 1988. Paleosols and the evolution of the
atmosphere: part I. In: Lerman, A., Meybeck, M. (Eds.), Physical
and Chemical Weathering in Geochemical Cycles. Kluwer Academic Publishers, New York.
Johnson, J.W., Oelkers, E.H., Helgeson, H.C., 1992. SUPCRT92: a
software package for calculating the standard molal thermodynamic properties of minerals, gases, aqueous species, and reactions
from 1 to 5000 bar and 0 to 1000 C. Comput. Geosci. 18, 899–947.
Kasting, J.F., 1987. Theoretical constraints on oxygen and carbon dioxide concentrations in the Precambrian atmosphere. Precambrian
Res. 24, 205–229.
Kasting, J.F., 1993. Earth’s early atmosphere. Science 259, 920–926.
Kasting, J.F., 2004. Archaean atmosphere and climate. Nature 432,
doi:10.1038/nature03166.
Kaufman, A.J., Xiao, S., 2003. High CO2 levels in the Proterozoic
atmosphere estimated from analyses of individual microfossils.
Nature 425, 279–282.
Keller, C.K., Wood, B.D., 1993. Possibility of chemical weathering before the advent of vascular land plants. Nature 364, 223–
225.
Murakami, T., Ito, J.-I., Utsunomiya, S., Kasama, T., Kozai, N.,
Ohnuki, T., 2004. Anoxic dissolution processes of biotite: implications for Fe behavior during Archean weathering. Earth Planet.
Sci. Lett. 224, 117–129.
8
N.D. Sheldon / Precambrian Research xxx (2006) xxx–xxx
Minter, W.E.L., 1999. Irrefutable detrital origin of the Witwatersrand
gold and evidence of eolian signatures. Econ. Geol. 94, 665–670.
Ohmoto, H., Watanabe, Y., Kumazawa, K., 2004. Evidence from massive siderite beds for a CO2 -rich atmosphere before ∼1.8 billion
years ago. Nature 429, 395–399.
Palmer, J.A., Phillips, G.N., McCarthy, T.S., 1989. Paleosols and their
relevance to Precambrian atmospheric composition. J. Geol. 97,
77–92.
Panahi, A., Young, G.M., Rainbird, R.H., 2000. Behavior of major and
trace elements (including REE) during Paleoproterozoic pedogenesis and diagenetic alteration of an Archean granite near Ville Marie,
Quebec, Canada. Geochim. Cosmochim. Acta 64, 2199–2220.
Pavlov, A.A., Hurtgen, M.T., Kasting, J.F., Arthur, M.A., 2003.
Methane-rich Proterozoic atmosphere? Geology 31, 87–90.
Pavlov, A.A., Kasting, J.F., Brown, L.L., Rages, K.A., Freedman, R.,
2000. Greenhouse warming by CH4 in the atmosphere of early
Earth. J. Geophys. Res. 105, 11981–119990.
Phillips, G.N., Law, D.M., 1997. Hydrothermal origin for Witwatersrand gold. Soc. Econ. Geol. Newslett. 31 (1), 26–33.
Pinto, J.P., Holland, H.D., 1988. Paleosols and the evolution of the
atmosphere. Part II. In: Reinhardt, J., Sigleo, W.R. (Eds.), Paleosols
and Weathering Through Geologic Time: Principles and Applications. Geological Society of America, Special Paper 216.
Retallack, G.J., 1986. Reappraisal of a 2200 Ma-old paleosol near
Waterval Onder, South Africa. Precambrian Res. 32, 195–232.
Retallack, G.J., Krinsley, D.H., 1993. Metamorphic alteration of a
Precambrian (2.2 Ga) paleosol from South Africa revealed by
backscattered electron imaging. Precambrian Res. 63, 27–41.
Retallack, G.J., Mindszenty, A., 1994. Well preserved Precambrian paleosols from northwest Scotland. J. Sediment. Res. A64,
264–281.
Rye, R., Holland, H.D., 2000a. Geology and geochemistry of paleosols
developed on the Hekpoort Basalt, Pretoria Group, South Africa.
Am. J. Sci. 300, 85–141.
Rye, R., Holland, H.D., 2000b. Life associated with a 2.76 Ga
ephemeral pond? Evidence from Mount Roe #2 paleosol. Geology 28, 483–486.
Rye, R., Kuo, P.H., Holland, H.D., 1995. Atmospheric carbon dioxide concentrations before 2.2 billion years ago. Nature 378, 603–
605.
Sawicki, J.A., Brown, D.A., Beveridge, T.J., 1995. Microbial precipitation of siderite and protoferrihydrite in a biofilm. Can. Mineral.
33, 1–6.
Sheldon, N.D., Retallack, G.J., 2001. Equation for compaction of paleosols due to burial. Geology 29, 247–250.
Sheldon, N.D., Retallack, G.J., 2002. Low oxygen levels in earliest
Triassic soils. Geology 30, 919–922.
Sheldon, N.D., 2003. Pedogenesis and geochemical alteration of the
picture Gorge Subgroup, Columbia River Basalt, Oregon. Geol.
Soc. Am. Bull. 115 (11), 1377–1387.
Sleep, N.H., 2004. Archaean palaeosols and Archaean air. Nature 432,
doi:10.1038/nature03167.
Watanabe, Y., Martini, J.E.J., Ohmoto, H., 2000. Geochemical evidence for terrestrial ecosystems 2.6 billion years ago. Nature 408,
574–577.
Wiggering, H., Beukes, N.J., 1990. Petrography and geochemistry of a
2000–2200-Ma-old hematitic paleo-alteration profile on Ongeluk
Basalt of the Transvaal Supergroup, Griqualand West, South
Africa. Precambrian Res. 46, 241–258.
Zbinden, E.A., Holland, H.D., Feakes, C.R., Dobos, S.K., 1988. The
Sturgeon Falls paleosol and the composition of the atmosphere
1.1 Ga BP. Precambrian Res. 42, 141–163.