Quantitative paleoenvironmental and paleoclimatic reconstruction ...

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N.D. Sheldon, N.J. Tabor / Earth-Science Reviews xxx (2009) xxx–xxx
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The fractionation of 16 O and 18 O among the reactants and products
is described by the alpha value (α):
18 α mineral =
18 O mineral
16 O mineral
18 O H2 O
16 O H2 O
ð45Þ
where 18 α is the oxygen isotope fractionation factor. The hydrogen and
carbon isotope fractionation factors are noted by D α and 18 α, respectively.
α values for isotope exchange reactions are very small,
which indicates that, in the absence of a more energetic reaction such
as dissolution and precipitation, isotope exchange will not occur. That
is, the isotope composition of a mineral is stable, and will remain
unchanged, after it has formed. Fractionation of 18 O and 16 O between a
mineral and water can vary, but changes in α are principally related to
the temperature of reaction such that:
10 3 lnα mineral−H2 O = A T 2 + B T + C ð46Þ
where A, B, and C are constants which describe the behavior of an
isotope for a specific mineral-water system, and T is temperature in K.
Note, however, that the quantity B/T does not appear in many isotope
fractionation equations. A list of fractionation equations for common
pedogenic minerals is given in Table 5.
Measurement of the isotope composition is usually reported with
respect to the difference between the measured 18 O/ 16 O value of the
sample with respect to the measured 18 O/ 16 O value of a standard.
These reported values, called δ (“delta”) values, are calculated as:
ð heavy X sample
δX = light X sample
× 1000 ð47Þ
1Þ
− heavy X standard
light X standard
where X is the element under consideration (C or O for example), heavy
and light refer to the concentration of heavy and light isotopes, respectively,
of element X in the sample and standard. δ values are
reported in per-mil units (parts per thousand), which is denoted by
the symbol, “‰”. Conventional presentation of δ values are δD for
hydrogen, δ 13 C for carbon and δ 18 O for oxygen. The standard that is
used for reporting δD and δ 18 O values is Standard Mean Ocean Water
(SMOW), whereas the standard used for reporting δ 13 C values is
PeeDee Belemnite. Therefore both SMOW and PDB have δ values that
are defined as 0‰ when referenced with respect to themselves.
7.1.2. Stable isotope fractionation factors of common pedogenic minerals
Asignificant amount of research has focused upon measurement and
determination of stable isotope fractionation factors (see reviews by
Hoefs, 1997; Valley et al., 1986; Savin and Lee, 1988; Chacko et al., 2001;
Sharp, 2007). Table 5 presents the stable isotope fractionation factors
(values of 10 3 lnα) of common soil-forming minerals, as determined
from laboratory syntheses, isotope exchange reactions and natural
experiments. At temperatures that reflect soil formation in the presence
of liquid water (~0–30 °C; Yapp, 2001a,b; Tabor and Montañez, 2005),
the majority of mineral-water oxygen isotope fractionation factors
(10 3 ln 18 α) have positive values, which indicates that 18 O is enriched in
the mineral product (Table 5; Fig.14A). Carbonates and silicates have the
largest (~25–35) and Fe(III)-oxides the smallest (~0–15) 10 3 ln 18 α
values, and Al- and Mn-oxides have intermediate values. In addition,
10 3 ln 18 α values decrease with higher temperatures for all minerals
considered herein. Minerals with little or no ionic substitution, such as
low-Mg calcite and kaolinite (Al 2 Si 2 O 5 (OH) 4 ) have relatively invariant
10 3 ln 18 α values under isothermal conditions. Mineral groups that
permit a great deal of metal-cation substitution, such as 2:1 phyllosilicate
minerals, result in variable 10 3 ln 18 α values under isothermal
conditions (Savin and Lee, 1988). For example, Tabor and Montañez
(2005) reported the chemistry of smectites from Permian–Pennsylvanian-age
paleosol profiles that correspond to differences in isothermal
10 3 ln 18 α values N1‰ among samples, whereas Koster and others (1999)
determined that nontronite (Na 0.3 Fe 3+ 2(Si,Al) 4 O 10 (OH) 2 ) 10 3 ln 18 α
values are ~4‰ lower than Fe-rich 10 3 ln 18 α values. This highlights the
importance of accurate mineralogical and chemical characterization of
pedogenic minerals in order to interpret appropriately the meaning of
their δ 18 Ovalues.
All mineral-H 2 O hydrogen isotope fractionation factors (10 3 ln D α;
Table 5; Fig. 14B) considered herein have negative values, indicating
that 1 H is enriched in the mineral product. Goethite has the most
negative (~−100), whereas gibbsite has the most positive (~− 5),
10 3 ln D α values, and phyllosilicates have intermediate values. Within
analytical uncertainty, goethite and gibbsite 10 3 ln D α values do not
Table 5
Light stable isotope fractionation equations for common soil-formed minerals.
Reaction 10 3 ln 18 α 10 3 ln D α 10 3 ln 13 α Reference
H 2 O(l)–H 2 O(v) 1.137⁎10 6 / T 2 −0.4156 ⁎ 28.844⁎10 6 /T 2 −76.248⁎
Majoube, 1971
10 3 /T−2.0667
10 3 /T+52.612
Calcite-H 2 O 2.78 ⁎10 6 /T 2 −2.89 –2.988⁎10 6 /T 2 +7.6663⁎10 3 /T−2.4612 (O'Neil et al., 1969; Kim and
O'Neil, 1997)
Calcite-CO 2 18.03 ⁎103 / T−32.42
Siderite-H 2 O 3.13⁎10 6 /T 2 − 3.50 Carothers et al. (1988)
Siderite-CO 2 See Fig. 17
Goethite-H 2 O 1.63⁎10 6 /T 2 −12.3 (Yapp, 1990; Yapp, 1987a,b; Bao and
Koch, 1999; Muller, 1995)
1.907⁎10 3 /T−8.004 −99.8
2.76⁎10 6 / T 2 −23.7
Hematite-H 2 O 1.63⁎10 6 /T 2 −12.3 (Yapp, 1990; Clayton and Epstein, 1961;
Bao and Koch, 1999)
0.413⁎10 6 /T 2 −2.56
0.542⁎10 6 /T 2 −5.221
0.733⁎10 6 /T 2 − 6.194
Gibbsite-H 2 O 2.04⁎10 6 /T 2 − (Vitali et al., 2000, 2001)
3.61⁎10 3 /T+3.65 − 5.1
Boehmite-H 2 O 2.11⁎10 6 /T 2 −4.4 XXX Bird et al. (1992)
Kaolinite-H 2 O 2.76⁎10 6 / T 2 −6.75 − 2.2⁎10 6 / T 2 −7.7 Sheppard and Gilg (1996)
Smectite-H 2 O 2.60 ⁎10 6 /T 2 −4.28
Sum of metal-oxygen
bonds
−2.2⁎10 6 /T 2 − 7.7 +
(2⁎Al− 4⁎Mg− 68 ⁎Fe)
(Delgado and Reyes, 1996; Savin and Lee, 1988;
Sheppard and Gilg, 1996)
Illite/smectite-H 2 O
−19.6⁎10 3 /T+25 (0–120 °C)
−45.3 ⁎10 3 /T+94.7 (0–150 °C)
(Yeh, 1980; Capuano, 1992; Hyeong and
Capuano, 2004)
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