Earth and Planetary Science Letters, 71 (1984) 229-240
Elsevier Science Publishers B.V., Amsterdam - Printed in The Netherlands
229
[3]
The stable isotopic composition of modern soil carbonate and its
relationship to climate
Thure E. Cerling
Department of Geology and Geophysics, University of Utah, Salt Lake City, UT 84112 (U.S.A.)
Received January 4, 1984
Revised version received August 29, 1984
The oxygen isotopic composition of modern soil carbonate is well correlated with the isotopic composition of local
meteoric water. The carbon isotopic cycle for CO2 in soils can be described in terms of the proportion of biomass using
the C4 photosynthetic pathway and the CO2 respiration rate of the soil; at low soil respiration rates significant
atmospheric CO2 mixing can occur. In general, the carbon isotopic composition of soil carbonate is related to the
proportion of C4 biomass present in soil, but soils that freeze to the depth of carbonate formation often have a
significant atmospheric component. This suggests that freezing of the soil solution should be considered as another
important mechanism for soil carbonate formation. Because of these relationships, the isotopic composition of soil
carbonate may be a paleoclimatic and paleoecologic indicator in cases in which diagenetic alteration has not occurred.
1. Introduction
Soil c a r b o n a t e forms u n d e r arid to s u b - h u m i d
climatic c o n d i t i o n s [1,2]. In general, it is found in
relatively d r y soils where grasses or m i x e d grasses
a n d shrubs are the d o m i n a n t vegetation. U n d e r
these c o n d i t i o n s soil p H is generally 7 or above, in
c o n t r a s t to forested soils where p H is below 6.
A u t h i g e n i c soil c a r b o n a t e is c o m m o n in soils where
m e a n a n n u a l rainfall is less than 75 cm while it is
rarely f o u n d in soils receiving m o r e than 100 c m
p r e c i p i t a t i o n per year.
Stable isotopes in soil c a r b o n a t e s are useful
tracers of the influence of climate on soil-forming
processes. O x y g e n isotopes in meteoric waters are
related to climate, especially m e a n a n n u a l tempera t u r e [3-5]. The c a r b o n isotopic c o m p o s i t i o n of
soil C O 2 d u r i n g the growing season is related to
the c a r b o n isotopic c o m p o s i t i o n of the b i o m a s s
[6-9] which is related to the p r o p o r t i o n of p l a n t s
that use the C 4 p h o t o s y n t h e t i c p a t h w a y [10,11].
T h e flux of water in a soil receiving 50 cm of rain
p e r year is 2.8 moles cm -2 y - l ; typical fluxes of
0012-821X/84/$03.00
© 1984 Elsevier Science Publishers B.V.
biogenic CO 2 from grassland soils are on the o r d e r
of 5 × 10 -3 moles cm 2 y - a [12,13]. Because soil
c a r b o n a t e forms at much lower rates than this,
typically 1 × 10 - 6 to 1 × 10 -5 mole cm -2 y - 1
[14-16], it is likely that the oxygen a n d c a r b o n
isotopic c o m p o s i t i o n of soil c a r b o n a t e is controlled b y the oxygen isotopic c o m p o s i t i o n of
m e t e o r i c waters a n d b y the c a r b o n isotopic comp o s i t i o n of soil CO2, respectively.
F e w previous studies have e x a m i n e d the isotopic c o m p o s i t i o n of soil c a r b o n a t e (e.g., [17-22]).
T h e relationship b e t w e e n the isotopic c o m p o s i t i o n
of meteoric water a n d soil c a r b o n a t e a n d b e t w e e n
vegetation a n d soil c a r b o n a t e has not been established b y these studies because they have exa m i n e d c a r b o n a t e s f o r m e d on limestone p a r e n t
m a t e r i a l [17,19,21,22] or h a v e s t u d i e d soil
c a r b o n a t e s f o r m e d d u r i n g several climatic episodes
or during u n k n o w n climatic c o n d i t i o n s [17-22].
This study examines soil c a r b o n a t e s f o r m e d d u r i n g
a b o u t the p a s t 10,000 years; it is only for this time
interval that one can use the m o d e r n values for the
isotopic c o m p o s i t i o n of meteoric waters a n d the
230
modern vegetation as an estimate of the isotopic
composition of soil water and soil CO 2 that produced the soil carbonates examined. In addition,
some results on paleosol carbonates are reported
because paleosol carbonates may give an indication of ancient climatic or ecologic conditions.
2. Terminology and methods
In order to use stable isotopes to study
carbonate formation and its relationship to climate,
it is necessary to eliminate the problem of
carbonate not formed in the soil. In this study
only those soils or paleosols were studied where
marine limestone makes up a negligible fraction of
the parent material. Throughout this paper, the
term soil carbonate is used to represent only the
carbonate in the soil or paleosol that is authigenic,
that is, formed in place. Calcretes represent a
special sort of soil carbonate deposit: that of massive, continuous horizons measuring up to several
meters in thickness. They are generally found in
regions that receive from 25 to 75 cm annual
rainfall.
Modern soil carbonates from Africa and North
America were analyzed. Samples were chosen from
areas where it was possible to estimate the proportion of C 4 biomass present when soil carbonate
formation occurred. Samples studied and their localities are briefly described in Appendix 1.
Carbonates were reacted with 100% phosphoric
acid to liberate CO 2. Modern soil carbonates were
roasted at 450°C under vacuum prior to reaction
with H3PO 4. Oxygen and carbon isotopic ratios
are reported in the standard notation relative to
the PDB standard. Water samples discussed in the
text were analyzed by equilibration with CO 2 gas
and subsequent mass spectrometer analysis; they
are reported relative to the isotopic standard
SMOW. Organic carbon from two soils was
analyzed by combustion at 800°C using CuO and
Ag foil [30] after CaCO 3 was removed with 10%
HC1.
estimates of the fraction of C 4 flora and of the
isotopic composition of meteoric water are included for each site.
3.1. Relationship between the oxygen isotopic composition of soil carbonate and meteoric water
It has previously been suggested that the oxygen
isotopic composition of soil carbonate may be
related to the oxygen isotopic composition of
meteoric water [20]. Salomons et al. [20] found
poor agreement between the expected 61~O values
of soil carbonate calculated from estimated
meteoric water compositions and the measured
6180 values of soil carbonate. They attributed this
observation to differences in mechanisms of soil
carbonate formation. However, because those
calcretes formed over much of the past million
years, it is possible that they formed under climatic
conditions different from the present so that modern isotopic values of meteoric water should not be
applied to those examples.
Samples in this study were chosen to maximize
the range of 6180 of meteoric water; Fig. 1 shows
I
I
The results of all isotopic analyses of modern
soil carbonate are given in Table 1. In addition,
I
0
iz
bJ
Z
0
m
n,.<
(~
13
I"
-5
..J
16
+
7
| |12
-FO
oo -15
-20
_1_.,o
[
-15
~180SMOW
3. Results and discussion
I
+ o:+
I
-I0
I
I
-5
0
METEORIC
WATER
Fig. 1. Oxygen isotopic composition of modern soil carbonates
from Africa, North America and Europe plotted against estimated isotopic composition of meteoric water for each locality.
Numbers refer to locality numbers in Table 1.
231
TABLE 1
Isotopic composition of modern soil carbonate, the estimated fraction of C 4 biomass, and the estimated isotopic composition of local
meteoric water
1. Olduvai Gorge, Tanzania (n = 5)
2. Laetoli, Tanzania (n = 3)
3. Nguu, Kenya (n = 3)
4. The Netherlands d (n = 10)
5. Israel g (n = 5)
6. Iowa, U.S.A. ( n = 6)
7. North Dakota, U.S.A. (n = 3)
8. Saskatchewan, Canada (n = 5)
9. Alaska, U.S.A. (n = 3)
10. Utah, U.S.A. (n = 3)
11. Wyoming, U.S.A. (n = 3)
12. Wyoming, U.S.A. (n = 3)
813CcaCos(PDB)
81s Ocaco3(PDB)
%C 4
+0.5±0.5
-3.2±0.3
-2.9±0.5
-10.3±0.4
-8.5±0.5
-4.3±2.3
-5.0±0.4
-6.5±1.3
+1.1±1.1
-6.5±0.3
-6.9± 0.1
-2.7±0.7
+0.3±0.5
-1.7±0.2
-3.3±0.2
-4.9±0.4
-5.8±0.9
-4.4±0.5
-9.1±0.3
-14.2± 1.1
-13.5±0.2
-13.7±0.2
-11.3± 0.5
-9.3±0.2
95 a
60 c
75 c
10 ~
26 h
50 ~
30 ~
20 i
0i
25 i
10 m
55 m
biomass
81SOn2o(SMOW)
-1.0
-3.0
-4.0
-7.9
-5.0
-6.0
-10.5
-15.0
-14.5
-14.2
-12.0
-10.0
b
b
b
r
f
j
j
j
k
I
j
j
a
b
c
a
e
r
g
h
Estimated from Tieszen et al. [31] and Livingstone and Clayton [51].
Unpublished data.
This site contains a significant proportion of the C 3 plant Acacia drepanolobium in its biomass. See discussion in text.
From Salomons and Mook [19].
By comparison to similar climates in western North America. See Teeri and Stowe [32].
International Atomic Energy Agency [33].
From Magaritz et al. [18]. Deviation quoted is the standard error of the mean for 5 nodule populations.
Proportion biomass estimated from percentage of grasses in region using the C 4 photosynthetic pathway using compilation of
Aaronsohn [50].
i Estimated from Teeri and Stowe [32].
J Estimated by comparison with other continental North American data compiled in [5] and [33] which was fit with a power function
relating the weighted mean isotopic composition of meteoric water and mean annual temperature.
k Estimated using data from Bethel and Barrow, Alaska [5].
i Average of four shallow groundwaters [34].
m Estimates from altitudinal distribution of C 4 biomass in Wyoming [57].
t h a t t h e r e is g o o d c o r r e l a t i o n b e t w e e n t h e o x y g e n
isotopic composition of meteoric water and that of
m o d e r n soil c a r b o n a t e . I f t h e i s o t o p i c c o m p o s i t i o n
o f soil c a r b o n a t e is p r e s e r v e d w i t h o u t d i a g e n e t i c
m o d i f i c a t i o n i n t h e g e o l o g i c r e c o r d , it s h o u l d b e a
good indicator of past meteoric water compositions.
The oxygen isotopic composition
o f soil
c a r b o n a t e f r o m s i n g l e h o r i z o n s o r soil p r o f i l e s
g e n e r a l l y s h o w a s t a n d a r d d e v i a t i o n o f a b o u t 0.5%0
o r less (see a l s o r e f e r e n c e s [18] a n d [20]). D e v i a t i o n s g r e a t e r t h a n t h i s m a y i m p l y c a r b o n a t e formation by different mechanisms within a single
soil p r o f i l e , o r it m a y i n d i c a t e f o r m a t i o n u n d e r
different climatic conditions. It probably does not
r e s u l t f r o m a R a y l e i g h f r a c t i o n a t i o n p r o c e s s as
p r e v i o u s l y s u g g e s t e d [20] b e c a u s e t h i s i m p l i e s t h a t
t h e soil b e h a v e s as a c l o s e d s y s t e m .
3.2. Model for the 6~SC(C02) distribution in soils
N o a d e q u a t e m o d e l d e s c r i b e s t h e isotopi,c c o m p o s i t i o n o f soil c a r b o n d i o x i d e . Soils h a v e h i g h e r
CO 2 partial pressures than does the atmosphere;
this results from CO 2 produced by root respiration
a n d b y m i c r o b i a l o x i d a t i o n o f o r g a n i c m a t e r i a l in
t h e soil. T o g e t h e r , t h e s e a r e r e f e r r e d t o as soil
respiration. T h e s t e a d y s t a t e c o n d i t i o n f o r C O 2 in
soil c a n b e c a l c u l a t e d a s s u m i n g t h a t t h e n e t soil
respiration (Q) be distributed equally over some
d i s t a n c e ( L ) so t h a t t h e r a t e o f C O 2 p r o d u c t i o n is
q}* = Q / L . T h u s :
OC~*
~-7-0=D*-
a2C *
s
-~ +q}*
Oz 2
(1)
for steady state, where Ds* is the diffusion coefficient for CO 2 in soil [35], and C* represents the
232
concentration in soil air. The superscript *
refers to bulk CO 2 without isotopic distinction.
The subscript s refers to soil. Assuming that the
soil is a one-dimensional box with an impermeable
base at depth L, and that the following boundary
conditions exist:
CO 2
at z = 0, C* = C~'
(2)
0z = 0
(3)
at z = L,
where C~' is the atmospheric concentration of CO2;
then [35]:
C*=-~-f Lz-~ +C~
(4)
Ds* can be related to the carbon dioxide diffusion
coefficient in air ( D ~ ) by [35]:
Os* = DairC p
-13%o and -27%0 are used for a pure C 4 biomass
and a pure C 3 biomass in the following discussion.
Because of their different masses, 12CO2 and
13CO2 have different diffusion coefficients; they
are related by [36,37]:
Dv
D~ -
(MY + M~)
M rM~
MBMa 11/2
( M ¢ + Ma ) J
(6)
where D r and D ~ are the diffusion coefficients of
the light and heavy isotopes of carbon in CO2, and
M Y, M B, and M~ are the masses 44 and 45 for
12CO2 and 13CO2, and the average atomic mass of
air, respectively. Equations of similar form can be
written describing the relationship of D* to D~ or
of D* to Dfi. In this analysis it is assumed that the
CO 2 concentrations are low enough so that D* can
be treated as constant throughout the soil profile.
The distribution of 13C(CO2) in soil can be described by:
where e is the free air porosity and p is a tortuosity
factor. Using Dair(CO2) of 0.14 cm 2 s -1, e of 0.24,
and O of 0.6 gives:
D* = 0.02 cm 2 s-1
which will be used in the following calculations.
D* will be treated as constant.
Using equation (4) and using values of soil
respiration rates [12,13] and an atmospheric concentration of CO 2 at 300 ppm, it is possible to
calculate CO 2 concentration profiles in soils (Fig.
2A). These are similar to observations of P(CO2)
with depth in many studies [9,35,38-41]. Soil respiration rates for grassland soils during the growing season are typically 6 × 10 -3 to 9 × 10 3 moles
m - 2 hr ~, those during the dry or cool non-growing season are typically about 1 × 10 -3 mole m -2
hr-~ [12,13,42]; during periods of soil freezing the
soil respiration rate may drop to almost zero [43].
The 8a3C value of the atmosphere is different
than that of the biomass and that of soil respiration. There are three major groups of plants with
different isotopic ratios. A survey of 650 species of
grasses [44-47] showed that C 4 grasses have an
average 8~3C value of -12.7%o, and C 3 grasses
have an average 813C value of -27.1%< CAM
plants have intermediate values depending on environmental conditions [11,48]. Thus, values of
where Coz, C~, g~¢, refer to the concentration of
13CO2 in the atmosphere, the concentration of
13CO2 in soil air, and the production rate of 13CO>
respectively. These are:
c? = ci*( R,
] for R a and R s
and:
R,
ll+Ri]
for R~,
where R S, R , , and R a represent the isotopic ratio
R i = (13CO2/12CO2), of carbon dioxide in soil air,
net respired CO 2, and atmospheric CO> respectively. Using the notation:
8~ = ( RRpDB
~
1 ) x 1000 for 8s,
, 8~, and ~
(8)
where 8i is the permil value for soil air, the atmosphere, or respired CO> respectively, and R pI)B is
the ratio (13C/12C) in the isotopic standard PDB.
Substituting into equations (4), (6), (7), and (8):
([
¢*
Z2 . "]t
z2
D,*.
233
× 1000
where:
1 + RpoB( 10~0 + 1)
log Pcoz
-4.0
"~
-3.0
-2.0
e S
jc,o~\ ~"~,O~^,e
_
O
A.
I00
0 ~
9
-
!Z
~ - 1 5 . ~
-2/
'~
3cp
Ioo
0
Soil
I
I
2
Respiration
I
3
Rate
I
4
Using 6, = -27%° and 6a = -6%~ (the estimated
pre-industrial value for atmospheric CO 2 [49]) it is
possible to calculate the isotopic composition of
the soil atmosphere. Fig. 2B shows the steady-state
distribution of 613C as a function of depth for
different soil respiration rates in the model soil
described above. These curves are compatible with
the following observations: soil CO 2 is often 3-7%~
heavier than the soil organic matter [7,9,38,39];
most soils sampled during periods of high soil
respiration do not show an isotopic gradient below
30 cm depth [9,39]; during periods of low soil
respiration a measurable isotopic gradient can be
observed [40]; and the net respired soil CO 2 is
about 4%~ lighter than the measured soil CO 2 [9].
This latter study [9] attributed this effect to molecular diffusion. For soil CO 2 concentrations of
10 20, limits of -22.2%e and -8.5%~ are calculated for the isotopic composition of soil CO 2
derived from soil respired CO 2 with isotopic compositions of - 27%~ and - 13%o, respectively, which
represent pure C 3 and C 4 biomasses, respectively.
This figure Shows that the atmospheric component
is likely to be important only when soil respiration
rates are quite low or at very shallow depths (less
than 10 cm), or both. This has important implications concerning soil carbonate formation, which
will be discussed below.
3.3. Relationship between the carbon &otopic composition of soil carbonate and the proportion of
C 4 biomass
5
( 10-3moles
6
m - 2 hr - i )
Fig. 2. A. Calculated steady state P(CO2) for different soil
respiration rates using the soil model described in the text
where L = 100 cm. B. Calculated steady state carbon isotopic
composition of soil carbon dioxide for different soil respiration
rates using the model described in the text. 3a3C values of
-27%o and -6%o were used for net soil respired CO 2 and
atmospheric COz, respectively.
The carbon isotopic composition of soil
carbonate is best considered with respect to the
carbon isotopic composition of the soil atmosphere. The isotopic composition of the soil atmosphere, as shown above, is related to the proportion of C 4 biomass, and the soil respiration rate.
C 4 plants are well adapted to conditions of high
water stress, particularly when that stress is related
234
to high temperatures. It has been observed that C 4
grasses are not present in floras when night temperatures fall below 8°C [32,47]. Vegetation in
regions of soil carbonate formation includes many
plant families whose members include both C 3 and
C 4 species (e.g., Poaceae, Chenopodiaceae,
Euphorbiaceae) although most trees common to
arid or semi-arid climates are C 3 species (e.g.,
Acacia). Grasses (Poaceae) are particularly sensitive to temperature and the distribution of C 4
grasses is known for many regions [31,32,46,47,
51,52]. These known distributions were used to
estimate the fraction of C 4 flora from those localities where grasses make up the flora; four localities
are included in this study where these distributions
cannot be used: the two samples from the literature (Israel and the Netherlands) and the samples
from Laetoli and Nguu which are in black cotton
soils and have a significant fraction of Acacia
drepanolobium present. Biomass estimates for
Laetoli and Nguu are based on the assumption
that the measured ~a3C values of soil organic
matter are due to the grass component because the
soil samples collected were situated away from
Acacia trees. Thus the values of - 1 5 and -13%o
represent grass populations of 85 and 100% C 4
components for Laetoli and Nguu, respectively. If
the Acacia biomass is about 25%, then the total C 4
biomass is about 60% and 75%, respectively, for
these two localities. Because of the uncertainty in
the Acacia biomass estimate, the overall estimate
in fraction C 4 biomass is less certain for these
localities than for other localities.
Newly-formed calcite in soils can probably be
considered to be in isotopic equilibrium with the
gas phase since rates of change in P(CO2) in soils
are much slower than those used in laboratories to
determine equilibrium isotopic fractionation (e.g.,
[53]). Some previous studies have considered
carbonate formation in soils to take place in a
closed system [18,20] so that the carbon isotopic
composition of the soil solution changes according
to a simple Rayleigh process. This is unlikely
because of the continuous CO2 flux due to soil
respiration. Thus, soil solutions are considered to
be in equilibrium with a CO 2 reservoir whose
isotopic composition is unchanged by carbonate
formation.
For the purpose of discussing the global distribution of 613C in soil carbonate, an isotopic fractionation (103 in a) of - 1 0 . 3 6 [54] at 25°C for
C O 2 - C a C O 3 will be used. Using this to calculate
the isotopic composition of soil carbonate in equilibrium with the CO 2 of the pure component endmembers, it is possible to show the mixing of these
reservoirs. Fig. 3 shows this mixing and gives the
613C values for modern soil carbonate analyzed in
this study; it shows that in general there is good
agreement between the carbon isotopic composition of soil carbonate and the fraction of C 4 biomass present. The high standard deviation for the
samples from Iowa result from the presence of two
isotopic populations with 613C values of about
- 2.9%0 and - 7.2%0, respectively.
I
I
I
I
I
[
I
5
I
I
I
ioo
-"1
§ o
~
I
-I0 -
-151 0
•
I
L
I
i
FRACTION
1
.5
t
C4
I
0
FLORA
Fig. 3. C a r b o n isotopic c o m p o s i t i o n of soil carbonate c o m pared to estimate of the fraction of C 4 plants in local flora
(data from Table 1), M i x i n g lines of calcite in equilibrium with
soil C O 2 and a t m o s p h e r i c CO 2 s h o w n for reference. These are
calculated using an isotopic fractionation factor (103 In a ) of
- 10.36%o at 25°C, and using ~13C values of - 22.2%0, - 8.5%~,
and -6%o for the c a r b o n isotopic c o m p o s i t i o n for soil C O 2
f r o m a 100% C 3 flora, soil C O 2 from a 100% C4 flora, and
a t m o s p h e r i c C O 2, respectively. N u m b e r s refer to locality n u m bers in Table 1.
235
3.4. Implications for soil carbonate formation
Soil carbonate formation is generally considered to result from carbonate supersaturation due
to evaporation, evapotranspiration, and lowering
of P(CO2) [1,2,17-21,28]. Of these, evaporation is
probably of little significance because evapotranspiration is the dominant mechanism for soil water
loss in areas covered by vegetation [2,55]. Few
studies have examined the effect of evapotranspiration on the isotopic composition of soil solutions but it is thought to be non-fractionating
[4,57]. Without detailed studies of the annual
oxygen and hydrogen isotopic variations in soil
waters, it is not possible to address this problem.
The carbon isotopic composition of these soil
carbonates does have important implications concerning soil carbonate formation. Of particular
interest are those samples that fall above the 10%
atmospheric component line. Samples below the
line (except Iowa which falls near the 10% line) are
from regions where the soil does not freeze to the
depth of carbonate formation; all samples with a
significant atmospheric component are from regions where the soil annually freezes to the depth
of carbonate formation. This implies that some
carbonate may be formed during periods of low
soil respiration rates. An important mechanism for
soil carbonate formation could be due to the increase in ion concentrations that results from ion
exclusion during ice formation. Because soil respiration rates are on the order of 0.25 × 10 3 moles
m 2 hr-1 or less during periods of soil freezing
(e.g., [43]) it is possible to get a high atmospheric
component in soil carbonate formed by this process (Fig. 2B). A single sample of cemented dune
material from Alaska [29] was examined with this
in mind; it yielded a 813C value of + 1.1 _+ 1.1%o
indicative of a high atmospheric component. To
calculate the atmospheric component for such a
sample one would of course have to compute the
atmospheric component based on a temperature of
0°C rather than 25°C as in Fig. 3. This results in
an upper limit of about + 8%o for the isotopic
composition of soil carbonate precipitated at 0°C
from a 100% atmospheric component. If applied
to this sample from Alaska, this implies an atmospheric component of 60-80%.
Studies of soil carbonate from single soils often
have a deviation of about _+0.5%° for 613C (e.g.,
[18,20]). Soil carbonate formed by solution composition changes resulting from evapotranspiration, from changes in P ( C O 2 ) , o r resulting from
soil solution freezing have isotopic signatures that
result from the three-component mixing described
above. This model suggests that variations in the
8~3C of soil carbonate may result from mixing of
these three reservoirs and from variations in the
proportion of C 4 floras seasonally or over a long
interval of time. Previously, these changes have
been attributed to closed system carbonate formation [19,20] or to contamination with detrital
material [18,22].
3.5. Relationship between 6180 and 613C in soil
carbonate
A coupling is expected between the 813C and
6~SO values for many soil carbonates. The fraction
of C 4 grasses is well correlated with night-time
temperatures [10,32]; hence, the 813C value for soil
CO 2 is expected to be higher in regions with high
night temperatures. The oxygen isotopic composition of meteoric waters from continental stations
receiving less than 1000 mm annual precipitation
is well correlated with mean annual temperature
[3 5] except for regions which have monsoonal
climatic conditions, which are often 4-6%o depleted in ~80 relative to continental stations of
similar temperatures (Fig. 4). In most regions, the
night-temperatures during the growing season and
mean annual temperature are well correlated; an
important exception are those climates buffered by
marine conditions which are expected to have a
disproportionately high percentage of C 3 plants.
Thus, San Francisco (U.S.A.) and St. Louis
(U.S.A.) both have mean annual temperatures of
13°C, but have July minimum temperatures of 12
and 19°C, and C4/(C 3 + C4) percentages of 8%
and 60%, respectively [32].
Fig. 5B shows a general model for the isotopic
composition of soil carbonates developed on grasslands. "Normal continental" soil carbonates are
represented in the shaded area, which is bounded
by the 0-30% admixture of the atmospheric corn-
236
I
I
I
I
I
I
[
~CARBONATE
•18 UPDB
[
Station where P_<15Omm for all months
0
"MONSOON EFFECT"
-20
~/
Average for months where
-IO
117,1 o
/
-5
O
5
Conctdion WeArs
orthern
/.',//
PslSOmm
AVerwOhge;ef°;m°~ohmS~~'o
-15
-20
-15
" !!
-5
~ -I0
~o~_5
/
o
/
1
I!
!
Io
-co
-15
~.oX\'%~ \
o
C'~
,,,> "%o, /
--
"}
-
,%
o .,."0-
I
°1
I
I
-20
5
Northern Canadian Waters
.II
u.I
-I0
J
-5
0
i
I
I
i
!
tuba
ecQ.
g
-25
I
-ioJ
I
0
I
5
I
I0
I
15
[
20
F
25
I
-5
30
TM (°C)
Fig. 4. Relationship of the weighted mean isotopic composition
of meteoric water from continental stations receiving less than
1000 mm of annual rainfall. Stations receiving more than 150
mm of rain in a single month are treated separately; the
weighted average for m o n t h s receiving less than 150 mm are
plotted separately from the weighted average of those months
receiving more than 150 mm. Data from IAEA compilations
[5,33].
ponent; areas affected by soil freezing are likely to
be closer to the 30% atmospheric component than
those not affected by soil freezing. A 100% C 4
flora was assumed to be represented by regions
with a mean annual temperature of 25°C; a 100%
C 3 flora was assumed to be represented by regions
with a mean annual temperature of 0°C; these are
similar to the percentages of grass floras for North
America under these temperature conditions [32].
All waters were assumed to fit the 61SO-tempera ture relationship of Fig. 4; that curve is plotted in
Fig. 5A, which also shows the equilibrium values
for the H z O - C a C O 3 system at different temperatures.
Certain climatic conditions can result in deviations from this "normal" continental trend. Coastal
stations in the Mediterranean and in Australia all
,30
• .. "
-r.o
-I0
-15
.."
~-Jo~
A. Continentel
B. Coestol
I
-20
I
-15
C. Monsoonel
D. Periglociel
I
-IO
I
-5
]
I
I
O
5
1
~18 ~CARBONATE
WPDB
Fig. 5. A. Relationship between isotopic composition of water
and carbonate using the temperature relationship of Craig [59].
Path of mean isotopic composition of water versus mean annual temperature (from Fig. 4) is plotted. Vectors of s u m m e r
heating and isothermal evaporation are also shown. B. Fields of
" n o r m a l " continental carbonates assuming 0-30% admixture
of atmospheric CO 2 with soil CO 2 and temperature relations
described in text. Fields of coastal carbonates, periglacial
carbonates, and " m o n s o o n a l " carbonates are also shown. See
discussion in text.
have meteoric waters whose isotopic composition
is several permil enriched in ~80 relative to continental stations [4,5]. This would result in a shift
of the oxygen isotopes to slightly heavier values
relative to the relationship shown in Fig. 5B;
calcretes from the Mediterranean region (Spain,
Libya, France, Italy, Crete [20]) plot in a field
similar to that of field B in Fig. 5B. Another
important effect is that of the "monsoon" effect.
237
Intense, heavy rainfalls often show a depletion in
180; monsoonal waters in India are depleted by
about 6%0 relative to normal continental waters
(Fig. 4). Calcretes from India [20] show this effect
and fall in field C in Fig. 5B. Another important
situation may be present in periglacial environments; if carbonate precipitation results primarily
from soil freezing when soil respiration rates are
low, very high 613C values are possible (field D in
Fig. 5B). One sample in our study from Alaska
may be representative of this condition; Dever et
al. [22] have also identified soil carbonate that they
attribute to periglacial conditions. The presence of
a significant non-grass component in the biomass
would also tend to make 313C values more negative; this was important in two of the samples
studied.
3.6. Use of the isotopic composition of soil carbonate
as a paleoclimatic indicator
Paleosol carbonates may be useful in some cases
in determining paleoclimates since, if they are
unaltered by diagenesis, they record information
concerning both the isotopic composition of
meteoric water and the proportion of C 4 biomass
present in the ecosystem. Two problems are of
immediate concern: overprinting and diagenesis.
Overprinting refers to the imposition of later
climatic information on previous information. This
could take place if the soil carbonate in a single
horizon formed during several different climatic
episodes. It is possible that some of the variation
observed in soil carbonates, especially calcretes, is
due to this problem. In evaluating this as a potential problem, one can consider the relationship
between the zone of soil carbonate formation and
the sedimentation rate. Soil carbonate horizons are
generally less than one meter thick; periods of
climatic change are on the order of 10,000 years or
more. Studies of terrestrial sequences in Africa [25]
and North America [58] show that intervals of
climatic change may be rapid ( < 50,000 years) but
long intervals of relatively constant climatic conditions (400,000 years) may persist. Sedimentation
rates must take this into account: very low sedimentation rates (less than 1 cm/1000 years) may
have an isotopic record that spans several different
climatic regimes; this is probably the case for
many calcretes. Higher sedimentation rates (greater
than 10 cm/1000 years) may result in the preservation of soil carbonate that retains climatic
information of a limited time span.
Diagenetic effects are also not easily evaluated
at this time. The carbonate precipitation and partial redissolution in soils is most sensitive in the A
and B horizons of soils; changes in P(CO2) and
water loss due to evapotranspiration are minimal
below the B horizon. The presence of micrite in
paleosols indicates that minimal recrystallization
has occurred. Co-existing micrite and sparite from
buried paleosol nodules in Eocene sediments in
Wyoming showed that 8~3C values are essentially
unchanged by recrystallization but 81SO values
were often 4-8%~ depleted in 180 in the sparite
(unpublished data). Pedogenic carbonates from
Olduvai Gorge show an excellent relationship with
time [25]. The persistence of 613C and 6XSO values
not compatible with today's climate, and the excellent correspondence between 14C ages on calcretes
and on associated organic material at Olduvai
Gorge [25] is evidence that in some cases diagenetic alteration does not appear to be important.
4. Conclusions
This study indicates that the oxygen and carbon
isotopic composition of soil carbonate is related to
the isotopic composition of meteoric water and to
the proportion of C 4 biomass present. Because of
this, soil carbonate can be an important paleoclimatic and paleoecologic indicator. Because the
proportion of C 4 biomass is related to temperature
and because the oxygen isotopic composition of
meteoric water is related to temperature, a positive
correlation between 613C and 6180 in soil
carbonates is sometimes found. Monsoonal,
coastal, and periglacial climates have different 613C
and 81SO relationships because of differences in
the isotopic composition of meteoric waters in
monsoons, influences of the oceans, and low soil
respiration rates, respectively.
In addition, this analysis indicates that the observed differences in 813C contents of soil
238
c a r b o n a t e s is n o t d u e t o c a r b o n a t e f o r m a t i o n i n a
c l o s e d s y s t e m , b u t r a t h e r t h a t d i f f e r e n c e s in t h e
i s o t o p i c c o m p o s i t i o n o f soil C O 2 c a n r e s u l t f r o m
c h a n g e s in t h e soil r e s p i r a t i o n r a t e a n d c h a n g e s i n
the proportion of C 4 biomass seasonally or over a
long time interval. The observed difference bet w e e n t h e i s o t o p i c c o m p o s i t i o n o f soil c a r b o n dio x i d e a n d t h e a s s o c i a t e d o r g a n i c m a t t e r i n soil is
expected because of the difference in diffusion
c o e f f i c i e n t s f o r 13COz and 12CO2. I n a d d i t i o n , t h e
p r e s e n c e o f a h i g h a t m o s p h e r i c c o m p o n e n t in s o m e
soil c a r b o n a t e s s u g g e s t s t h a t f r e e z i n g o f soil s o l u tions may be an additional carbonate formation
mechanism.
Acknowledgements
G . W . C o x , S. G r e e n , R i .
H a y , J.L. R i c h a r d s o n , a n d R.J. St. A r n a u d p r o v i d e d s o m e o f t h e
modern carbonates used in this study; T.M. Bown
p r o v i d e d a s s i s t a n c e i n t h e field; J.R. B o w m a n a n d
R. L a m b e r t p r o v i d e d l a b o r a t o r y a s s i s t a n c e . I t h a n k
F.H. Brown, M.C. Monaghan, and W.T. Parry for
many helpful discussions. Reviewers improved the
clarity of this manuscript. This work (including
p a l e o s o l s ) w a s s u p p o r t e d b y t h e L.S.B. L e a k e y
Foundation, the Foundation for Research in the
Origin of Man, the National Science Foundation
( B N S - 8 0 0 7 3 5 4 a n d B N S - 8 2 1 0 7 3 5 ) , a n d t h e U.S.
G e o l o g i c a l S u r v e y ( g r a n t to T . M . B o w n ) .
Appendix I
(1) Olduvai Gorge, Tanzania: 1.5 cm thick laminar calcrete
formed in aeolian tuff prior to deposition of the Namorod Ash
[23-25]. Calcretes at Olduvai Gorge have been shown to form
very quickly because of their highly reactive parent material
[24]. Dominant grass species include Digitaria macroblephara
and Sporabolus fimbriatus [26]. Olduvai Gorge has a mean
annual temperature of about 23°C.
(2) Laetoli, Tanzania: 1 cm nodules from black cotton soil
(vertisol). This site is on the edge of the Serengeti Plain but is
5-6°C cooler than Olduvai Gorge because of its higher elevation of about 1800 m. Vegetation is grassland with some Acacia
drepanolobium, the whistling thorn.
(3) Nguu, Kenya: 1 mm nodules from black cotton soil
(vertisol). Vegetation is grassland with some Acacia drepanolobiurn. Mean annual temperature is about 23°C.
(4) The Netherlands: Salomons and Mook [19] reported on
loess nodules from the Netherlands. The age of nodule formation is Late Pleistocene or Holocene.
(5) Israel: Magaritz et al. [18] have studied soil carbonates
from the coastal plain of Israel. Samples quoted in this study
include only those soil carbonates that had 14C ages less than
10,000 B.P.
(6) Iowa, U.S.A.: Loess nodules from the oxidized and
unleached zone of Wisconsin loess were collected from Logan,
Iowa. Carbonate content of the parent material is about 5%
CaCO 3 [27], but the loess nodules approach 100% calcite. Loess
deposition ceased about 14,000 B.P. in Iowa [27]; since these
nodules were from high in the loess profile, it is presumed that
they are of about this time or later. Arguments of Ruhe [27]
show that the oxidized and unleached zone must have formed
before 6800 B.P. Bouteloua, Andropogon, Agropyron, and Stipa
are important elements of this prairie flora [10,32]. This area
has a mean annual temperature of about 11°C.
(7) North Dakota, U.S.A. : The Williams soil series contains
1 mm soil carbonate nodules; this soil is formed on Wisconsin
glacial till that has about 5% detritat carbonate in the parent
material. This is in the mixed-grass prairie of central North
America; it has a mean annual temperature of about 5°C.
(8) Saskatchewan, Canada: Samples were taken from three
different soils in central Saskatchewan; details of these soils are
discussed in St. Arnaud [28]. The parent material contains a
few percent detrital carbonate; 14C dates on different soil size
fractions show decreasing age with decreasing particle size. The
less than 2/~m size fraction yields 14C ages of 1595-7200 B.P.,
whereas the coarse silt and sand size fraction yield ages of
> 30,000 B.P. Only carbonates in the less than 2 ~tm size
fraction are reported here. All soils were formed on glacial
deposits. This prairie site has a mean annual temperature of
about 2°C.
(9) Alaska, U.S.A.: Carbonate cemented sand dunes have
been reported at longitude 158°W and latitude 67°N [29]. Sand
from these dunes does not have calcite present, although a
stabilized dune field some 25 km distant is reported to have
some detrital carbonate present [29]. Carbonate formation is
thought to have taken place since or during late Wisconsin time
[29]. Estimated mean annual temperature is about - 7 ° C .
(10) Utah, U.S.A.: Laminar calcrete coating boulder near
Tintic, Utah. This desert grassland has a mean annual temperature of about 7°C.
(11) Wyoming, U.S.A. I: Laminar calcrete coatings on pebbles in post glacial fill or alluvium at about 2500 m elevation.
Pebbles were about 20 cm below the surface. This area has an
estimated mean annual temperature of about 4°C. This is an
alpine grassland.
(12) H@oming, U.S.A. 11: Laminar calcrete coatings on
orthoquartzite pebbles on postglacial alluvium at about 2000
meters elevation. Estimated mean annual temperature is about
6°C.
(13) Paleosol carbonates: Paleosol carbonates from buried
soils in Eocene sediments in Wyoming (U.S.A.) and from
Pliocene and Pleistocene sediments in Kenya, Spain, and
Tanzania also have been analyzed. Other paleosol and soil
carbonates from literature sources are discussed as well.
239
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