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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δ 18 O and δD values of altered bentonites from Cortijo de Archidona
(Southeastern Spain) ranged from 21.0 to 23.0‰ and − 87 to −102‰,
respectively, which corresponds to temperatures of crystallization
ranging from 9±3 °C to 21±3 °C. In their assessment of clay minerals
in an Upper Triassic paleosol profile derived from basalt, Tabor and
others (2004a) measured smectite δ 18 O and δD values ranging from
14.5 to 15.6‰ and −78 to −71‰, respectively. This range of smectite
δ 18 O and δD values corresponds to temperatures of crystallization
ranging from 49 to 57 °C. Such temperatures are far too high for
soil formation, and were interpreted to result from hydrothermal
alteration.
Tabor and Montañez (2005) adopted a similar approach to that of
Delgado and Reyes (1996) in order to estimate temperature of smectite
crystallization for smectite and kaolinite mixtures in Upper Pennsylvanian
and Lower Permian paleosols from Texas, U.S.A. Notable
differences between the approach of Tabor and Montañez (2005)
and Delgado and Reyes (1996) are (1) analysis of only the b0.2 μm
fraction of phyllosilicate and (2) consideration of the variability of
oxygen and hydrogen isotope fractionation between smectite and
water that is related to differences in the chemical composition of
smectite. The range of δ 18 O and δD values corresponds to temperatures
of phyllosilicate crystallization ranging from 22±3 °C to 35±3 °C,
where phyllosilicate δ 18 O and δD values from Upper Pennsylvanian
paleosol profiles indicated significantly lower temperatures than δ 18 O
and δD values from Lower Permian paleosol profiles.
Vitali et al. (2002) analyzed the b0.2 μm of mixed-mineral
phyllosilicate fraction from a middle Cenomanian paleosol profile in the
Dunvegan fm., British Columbia, Canada. Measured δ 18 OandδD values
ranged from −12.9‰ to −11.6‰ and −93‰ to −81‰, respectively, and
were interpreted to correspond to paleotemperatures of phyllosilicate
crystallization ranging from 9 to 16 °C. To date, very few studies have
focused upon the δ 18 OandδD values of soil and paleosol smectites and
other common pedogenic minerals such as gibbsite and related
aluminum hydroxides (Bird and Chivas, 1989; Vitali et al., 2000, 2001)
and illite (Yeh, 1980; Capuano, 1992; Gilg, 2003; Hyeong and Capuano,
2004). Nevertheless, these hydroxylated pedogenic minerals are very
common mineral products of pedogenic weathering, and these minerals
have the potential to become important proxies of paleotemperature
throughout much of the terrestrial Phanerozoic stratigraphic record.
7.5.3.3. Kaolinite. Much work has been done with the oxygen and
hydrogen isotope compositions of soil-formed kaolinites (Sheppard,
1977; Hassanipak and Eslinger, 1985; Bird and Chivas, 1988a,b, 1989;
Lawrence and Rashkes Meaux, 1993; Savin and Hsieh, 1998; Girard
and Savin, 1996; Giral-Kacmarcík et al., 1989; Girard et al., 2000). The
oxygen and hydrogen isotope kaolinite-water fractionation equations
(Sheppard and Gilg, 1996) in conjunction with the equation for the
meteoric water line of Craig (1961) relates the measured δ 18 O and δD
of kaolinite to temperature of crystallization:
○
T
ð KÞ = 3:04×10 6 !1
2
δ 18 O k −0:125×δD k +7:04
ð62Þ
As discussed earlier, only kaolinite samples from paleosol profiles
with δ 18 O and δD values that plot within the WESD and MSD should
be considered viable paleotemperature proxies. For the following
discussion, we also limit the dataset of kaolinite δ 18 O and δD values to
only those samples that are known to have formed within the timeframe
of an individual Epoch or less in the Cenozoic. Cretaceous
samples are limited to Campanian or younger kaolinites. This culling
of the literature data leaves 78 kaolinite samples, which are plotted
with respect to their corresponding paleolatitude of soil formation in
Fig. 22.
Several observations may be made from this dataset (Fig. 22).
(1) The mid-latitudes, from ~35 to 50° N and ~30 to 62° S, are the
only locations with published, pre-Pliocene kaolinite δ 18 O and δD
values. However, pedogenic kaolinite formation surely formed at
lower latitude tropical sites (e.g., Wilson, 1999; Jacobs et al., 2005).
Therefore, more attention should be given to determination of kaolinite
δ 18 O and δD values, as well as other potential geochemical proxy
data of paleotemperature, from low-latitude sites in the future.
(2) Kaolinite δ 18 O and δD values indicate their warmest mean
temperatures (24 °C) during Cretaceous and Eocene time, followed by
Paleocene (22 °C), and their lowest mean temperatures (20 °C) during
Oligocene and Miocene time.
(3) Kaolinite δ 18 O and δD values correspond to crystallization
temperatures that are significantly (N3 °C) higher than mean annual
surface-air temperatures at modern isolatitudinal sites. Because of the
large thermal-mass of soil compared to atmosphere, subsurface
temperatures are generally assumed to be approximately equivalent to
that of mean annual temperature (+1 to 2 °C) with little or no seasonal
change in temperature (e.g., Cerling and Quade, 1993; Buol et al., 2003).
Therefore, isotope values of pedogenic isotope compositions are
assumed to approximately record mean annual surface air temperatures
in most studies (Yapp, 2000; Tabor, 2007; Tabor and Montañez, 2005;
Tabor and Yapp, 2005c). Yet, long-term monitoring of ground and air
temperatures from mid-latitude sites indicates that subsurface soil
temperatures are strongly correlated with surface air temperatures, and
vary seasonally, from ~0° to 22 °C, at 1.0 m depth (Bartlett et al., 2006).
The rate of chemical weathering, and of crystallization of pedogenic
minerals, is strongly temperature dependent. Furthermore, crystallization
of soil minerals likely will not occur at 0 °C, as liquid water is
necessary for chemical reaction. Therefore, it is possible that strikingly
warm temperatures indicated by kaolinite δ 18 OandδD values in Fig. 22
more accurately reflect surface-air temperatures during warm-months
and growing seasons, and thus should not be compared with modern
mean annual surface air-temperatures.
Note that mean annual surface air temperatures in low-latitude
tropical sites do not vary annually by more than 1 or 2 °C. Therefore
soil subsurface temperatures do not vary by more than 1 or 2 °C, and
pedogenic mineral δ 18 O and δD values from those sites likely provide
a reasonable approximation of surface air temperatures. However, it is
possible that perennially warm conditions, similar to those in the
modern tropics, existed at higher mid-latitude sites during early
Neogene, Paleogene and Late Cretaceous. This reflects both that
paleosol profiles (Van Houten, 1961, 1982; Dury, 1971; Tardy et al.,
1990a,b) and fossil floras (e.g., Dorf, 1964) that are indicative of
modern, low-latitude tropical, sites (~15° N and S) occurred at
paleolatitudes as high as 60°N. Furthermore, marine paleotemperature
indicators suggest that mid-latitude surface oceans were significantly
warmer, as high as 30 °C, during intervals of the Paleogene
and Late Cretaceous (e.g., Takashima et al., 2006). In this regard,
paleosol kaolinite δ 18 O and δD values, and paleotemperature estimates
shown in Fig. 22, may provide reasonable estimates of mean
annual surface air paleotemperatures.
Future work of this sort should seek to (1) collect kaolinite δ 18 O
and δD values for paleotemperature estimates from places and times
that are poorly represented in Fig. 22, (2) produce similar databases
of δ 18 O and δD value with other common hydroxylated pedogenic
minerals, and (3) if these paleotemperatures are correct, evaluate
what mechanisms (such as atmospheric pCO 2 , ocean and atmospheric
circulation, etc.) might contribute to such high temperatures in the
mid-latitudes.
A major limitation to the application of single-mineral paleothermometers
from soils and paleosols is the requisite assumption of
chemical equilibrium with water δ 18 O and δD values that lie upon the
meteoric water line of Craig (1961). Changes in the position of the
MWL (in hydrogen vs. oxygen isotope space) might have occurred at
times during the geological past. The slope of the meteoric water line
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