Quantitative paleoenvironmental and paleoclimatic reconstruction ...

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26 N.D. Sheldon, N.J. Tabor / Earth-Science Reviews xxx (2009) xxx–xxx
tropospheric CO 2 is isotopically heavier; ~− 6.5 to − 8.0‰ during the
Holocene), but will be isotopically more negative than δ 13 Cvaluesofthe
same mineral that forms in a 2- or 3-CO 2 component soil. Minerals that
may result from one-component soil CO 2 mixing include calcite (Tabor
et al., 2007) and siderite (Mozley and Carothers, 1992; Mozley and
Wersin,1992; McMillan and Schwertmann,1988; Ludvigson et al.,1998;
White et al., 2005). It is less likely that carbon-bearing minerals that
require oxidizing conditions (e.g., Fe(CO 3 )OH in goethite; Yapp, 2001b)
or free drainage (e.g., occluded CO 2 in gibbsite; Bárdossy and White,
1979; Schroeder and Melear, 1999) will form in soils characterized by
one component of soil CO 2 .
7.2.2. Two-component soil CO 2
Two-component soil CO 2 mixing assumes that the only two sources
contributing to soil CO 2 are CO 2 from the open atmosphere (tropospheric
CO 2 ) and CO 2 from in situ oxidation of biological carbon in the
soil. Steady-state solutions to the one-dimensional Fickian diffusion
equation yield reasonable representations of δ 13 C values of soil CO 2
with depth in modern soils characterized by mixing of two sources of
CO 2 (Cerling, 1984). In these models, the pCO 2 and δ 13 C value of Earth's
atmosphere is the upper boundary condition, and there is a depthdependent
CO 2 production term that describes the oxidation of organic
carbon in the soil. Above a characteristic CO 2 production depth within
the soil, both the concentration and δ 13 C of soil CO 2 will progressively
approach values of the Earth's atmosphere. Thus, CO 2 in soil gas can be
generally regarded as a mixture of two isotopically distinct endmembers
in these soils: atmospheric CO 2 and CO 2 derived from in situ
oxidation of soil organic matter (Fig. 16). As mentioned earlier, soil
organic matter δ 13 C values are principally determined by (1) the δ 13 C
of CO 2 in the global troposphere and (2) the photosynthetic pathway
employed by primary producers within the soil profile. Therefore, for
times in the geological past when C 3 photosynthesis was the dominant
floral metabolism (pre-Miocene time; see 8.6) and soil organic matter
δ 13 C exhibited minimal temporal or spatial variation (e.g., Tabor et al.,
2004a,b; Peters-Kottig et al., 2006; Montañez et al., 2007), stratigraphic
changes in mineral δ 13 C values among paleosol profiles may
indicate either (1) variations in soil productivity (i.e., the concentration
of CO 2 derived from in situ oxidation of organic matter) or (2)
changes in atmospheric pCO 2 . In this regard, δ 13 C values of carbonbearing
pedogenic minerals which formed in the presence of twocomponent
soil CO 2 mixing have the potential to provide information
about (1) paleoecology and (2) the concentration of CO 2 in ancient
atmospheres. Pedogenic minerals which provide geochemical records
of two-component soil CO 2 mixing include calcite (Cerling,1991; Mora
et al., 1996; Yapp and Poths, 1996; Ekart et al., 1999; Tabor et al., 2004a,
b), goethite (Yapp, 1987a,b; Yapp and Poths, 1992, 1993, 1996; Yapp,
2004; Tabor et al., 2004a,b; Tabor and Yapp, 2005a), and gibbsite
(Schroeder and Melear, 1999; Tabor and Yapp, 2005a,b; Schroeder
et al., 2006). Of these minerals, only the geochemistry of calcite and
goethite are sufficiently understood to permit estimates of atmospheric
pCO 2 . Both pedogenic calcite and goethite are relatively
abundant in the terrestrial stratigraphic record, and thus have the
potential to provide a relatively continuous Phanerozoic record of
paleoatmospheric pCO 2 (see Section 7.4.3.1; Ekart et al., 1999) that
may be cross-checked by independent proxy methods (Yapp and
Poths, 1996; Tabor et al., 2004a,b).
Fig. 16. Steady-state model plot of soil depth in meters versus (A) soil CO 2 concentration and (B) δ 13 C values of soil CO 2 and pedogenic calcite. Solution to the soil CO 2 concentrations
assumes Fickian, one-dimensional diffusion, of the form
0 1
ð/ = D s Þ × 1− e − z = Z
pCO 2 = @
410 6 A + 300
22
where ϕ is the production rate of CO 2 in the soil in gC cm − 3 h − 1 ,D s is the diffusion coefficient for gas through the soil, z is depth in meters in the soil, and Z is a scaling depth in meters,
also called the characteristic production depth of CO 2 (Cerling, 1991) in the soil. This model soil is calculated with ϕ=0.00002 gC cm − 3 h − 1 , D s =0.0001, and Z=0.2, and 300 ppmV
as the upper boundary limit of the soil (i.e., a tropospheric pCO 2 of 300 ppmV). Soil CO 2 δ 13 C values were calculated as a mass balance of tropospheric CO 2 (300 ppmV through
profile) with δ 13 C value of − 6.5‰ (a reasonable pre-industrial age value; Schlesinger, 1997) and CO 2 from oxidation of organic matter (variable concentrations) with δ 13 C value of −
6.5‰ (i.e., organic matter δ 13 C=− 26.4‰). Soil calcite δ 13 C values were calculated using the CO 2 -calcite carbon isotope fractionation equation of Bottinga (1968) with assumed
calcite crystallization temperatures of 25 °C and mildly alkaline soil solution. See Text for discussion.
Please cite this article as: Sheldon, N.D., Tabor, N.J., Quantitative paleoenvironmental and paleoclimatic reconstruction using paleosols, Earth-
Science Reviews (2009), doi:10.1016/j.earscirev.2009.03.004