Earth and Planetary Science Letters, 71 (1984) 229-240 229 Elsevier Science Publishers B.V., Amsterdam - Printed in The Netherlands

[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 CO 2 in soils can be described in terms of the proportion of biomass using the C 4 photosynthetic pathway and the CO 2 respiration rate of the soil; at low soil respiration rates significant atmospheric CO 2 mixing can occur. In general, the carbon isotopic composition of soil carbonate is related to the proportion of C 4 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 ca rbona te forms under ar id to sub-humid cl imat ic condi t ions [1,2]. In general , it is found in relat ively d ry soils where grasses or mixed grasses and shrubs are the dominan t vegetat ion. U n d e r these condi t ions soil p H is general ly 7 or above, in cont ras t to fores ted soils where p H is below 6. Auth igenic soil ca rbona te is c o m m o n in soils where mean annual rainfal l is less than 75 cm while it is rare ly found in soils receiving more than 100 cm prec ip i t a t ion per year.

Stable i sotopes in soil ca rbona tes are useful t racers of the inf luence of c l imate on soi l - forming processes. Oxygen isotopes in meteor ic waters are re la ted to cl imate, especial ly mean annual temper- a ture [3-5]. The ca rbon isotopic compos i t ion of soil CO 2 dur ing the growing season is related to the ca rbon isotopic compos i t ion of the b iomass [6-9] which is re la ted to the p ropor t ion of p lan ts that use the C 4 pho tosyn the t i c pa thway [10,11]. The flux of water in a soil receiving 50 cm of rain per year is 2.8 moles cm -2 y - l ; typical fluxes of

biogenic CO 2 f rom grass land soils are on the o rder of 5 × 10 -3 moles cm 2 y - a [12,13]. Because soil ca rbona te forms at much lower rates than this, typica l ly 1 × 10 -6 to 1 × 10 -5 mole cm -2 y -1 [14-16], it is l ikely that the oxygen and ca rbon isotopic compos i t ion of soil ca rbona te is con- t rol led by the oxygen isotopic compos i t ion of meteor ic waters and by the carbon isotopic com- pos i t ion of soil CO2, respectively.

Few previous studies have examined the iso- topic compos i t ion of soil ca rbona te (e.g., [17-22]). The re la t ionship be tween the isotopic compos i t ion of meteor ic water and soil ca rbona te and be tween vegeta t ion and soil ca rbona te has not been estab- l ished by these studies because they have ex- amined carbona tes formed on l imestone pa ren t m a t e r i a l [17,19,21,22] or have s tud ied soil ca rbona tes formed dur ing several c l imatic episodes or dur ing unknown cl imatic condi t ions [17-22]. This s tudy examines soil ca rbona tes formed dur ing abou t the pas t 10,000 years; it is only for this t ime interval that one can use the modern values for the i sotopic compos i t ion of meteor ic waters and the

0012-821X/84/$03.00 © 1984 Elsevier Science Publishers B.V.

230

modern vegetation as an estimate of the isotopic composition of soil water and soil CO 2 that pro- duced the soil carbonates examined. In addition, some results on paleosol carbonates are reported because paleosol carbonates may give an indica- tion 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 mas- sive, 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 propor- tion of C 4 biomass present when soil carbonate formation occurred. Samples studied and their lo- calities 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.

3. Results and discussion

The results of all isotopic analyses of modern soil carbonate are given in Table 1. In addition,

estimates of the fraction of C 4 flora and of the isotopic composition of meteoric water are in- cluded for each site.

3.1. Relationship between the oxygen isotopic com- position 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 mod- ern 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

0

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O

oo -15

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-F-

13

I" 16 +

i z

_1_.,o + o:+

[ I I I -20 -15 - I0 -5 0

~ 1 8 0 S M O W METEORIC WATER

I

Fig. 1. Oxygen isotopic composition of modern soil carbonates from Africa, North America and Europe plotted against esti- mated isotopic composition of meteoric water for each locality. Numbers refer to locality numbers in Table 1.

TABLE 1

Isotopic composition of modern soil carbonate, the estimated fraction of C 4 biomass, and the meteoric water

231

estimated isotopic composition of local

813 CcaCos(PDB) 81s Ocaco3(PDB) %C 4 biomass 81SOn2o(SMOW)

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)

+0.5±0.5 +0.3±0.5 95 a -1 .0 b -3 .2± 0 .3 -1 .7± 0 .2 60 c -3 .0 b -2 .9± 0 .5 -3 .3± 0 .2 75 c -4 .0 b

-10 .3±0.4 -4 .9± 0 .4 10 ~ -7 .9 r -8 .5± 0 .5 -5 .8± 0 .9 26 h -5 .0 f -4 .3± 2 .3 -4 .4± 0 .5 50 ~ -6 .0 j -5 .0± 0 .4 -9 .1± 0 .3 30 ~ -10.5 j -6 .5± 1 .3 -14.2±1.1 20 i -15.0 j +1.1±1.1 -13 .5±0.2 0 i -14.5 k -6 .5± 0 .3 -13 .7±0.2 25 i -14.2 I -6 .9±0.1 -11.3±0.5 10 m -12.0 j -2 .7± 0 .7 -9 .3± 0 .2 55 m -10.0 j

a Estimated from Tieszen et al. [31] and Livingstone and Clayton [51]. b Unpublished data. c This site contains a significant proportion of the C 3 plant Acacia drepanolobium in its biomass. See discussion in text. a From Salomons and Mook [19]. e By comparison to similar climates in western North America. See Teeri and Stowe [32]. r International Atomic Energy Agency [33]. g From Magaritz et al. [18]. Deviation quoted is the standard error of the mean for 5 nodule populations. h 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 the o x y g e n i s o t o p i c c o m p o s i t i o n of m e t e o r i c w a t e r a n d t h a t of m o d e r n soil c a r b o n a t e . I f t he i s o t o p i c c o m p o s i t i o n of 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 in t h e geo log ic r e co r d , it s h o u l d b e a g o o d i n d i c a t o r o f p a s t m e t e o r i c w a t e r c o m p o s i - t i ons .

T h e o x y g e n i s o t o p i c c o m p o s i t i o n of soil c a r b o n a t e f r o m s ing le h o r i z o n s or soil p ro 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 of a b o u t 0.5%0 o r less (see a l so 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 th i s m a y i m p l y c a r b o n a t e for- m a t i o n b y d i f f e r e n t m e c h a n i s m s w i t h i n a s ing le soil p rof i le , or i t 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 d i f f e r e n t c l i m a t i c c o n d i t i o n s . I t p r o b a b l y d o e s n o t 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 th i s i m p l i e s t h a t the soil b e h a v e s as a c l o s e d sys tem.

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 the isotopi ,c c o m - p o s i t i o n of soil c a r b o n d iox ide . Soi ls h a v e h i g h e r C O 2 p a r t i a l p r e s s u r e s t h a n d o e s the a t m o s p h e r e ; th i s r e su l t s f r o m C O 2 p r o d u c e d b y r o o t r e s p i r a t i o n a n d b y m i c r o b i a l o x i d a t i o n of o r g a n i c m a t e r i a l in the soil. T o g e t h e r , t he se a re r e f e r r e d to 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 fo r C O 2 in soi l 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 he ne t soil r e s p i r a t i o n ( Q ) b e d i s t r i b u t e d e q u a l l y o v e r s o m e d i s t a n c e ( L ) so t h a t t he r a t e of C O 2 p r o d u c t i o n is q}* = Q / L . T h u s :

OC~* a2C * ~ - 7 - 0 = D * - -~ +q}* (1) s Oz 2

for steady state, where Ds* is the diffusion coeffi- cient for CO 2 in soil [35], and C* represents the

232

C O 2 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:

at z = 0, C* = C~' (2)

=0 (3) at z = L, 0z

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 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 con- centration 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 res- piration rates for grassland soils during the grow- ing season are typically 6 × 10 -3 to 9 × 10 3 moles m-2 hr ~, those during the dry or cool non-grow- ing 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 respira- tion. 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 en- vironmental conditions [11,48]. Thus, values of

-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]:

D v (MY + M~) MBMa 11/2 D ~ - M rM~ ( 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 de- scribed by:

where Co z, 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> respec- tively. Using the notation:

8~ = ( R ~ 1 ) x 1000 for 8s, 8~, and ~ RpDB , ( 8 )

where 8 i is the permil value for soil air, the atmo- sphere, 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):

233

([ Z2 . "]t ¢* z 2 D,*.

where:

1 + RpoB( 10~0 + 1)

log Pcoz

-4.0 -3.0 -2.0

c,o~ \ ~" O~ e S "~ _ j ~, ^,e

O

A. I00

! 0 ~ 9 - Z ~ - 1 5 . ~

- 2 /

'~ 3cp

I o o I I I 0 I 2 3 4 5 6 S o i l R e s p i r a t i o n R a t e ( 1 0 - 3 m o l e s 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.

× 1000

Using 6, = -27%° and 6 a = -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 molec- ular diffusion. For soil CO 2 concentrations of 10 20, limits of -22.2%e and -8.5%~ are calcu- lated for the isotopic composition of soil CO 2 derived from soil respired CO 2 with isotopic com- positions 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 implica- tions concerning soil carbonate formation, which will be discussed below.

3.3. Relationship between the carbon &otopic com- position of soil carbonate and the proportion of C 4 biomass

The carbon isotopic composition of soil carbonate is best considered with respect to the carbon isotopic composition of the soil atmo- sphere. The isotopic composition of the soil atmo- sphere, as shown above, is related to the propor- tion 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 tem- peratures 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 sensi- tive 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 locali- ties where grasses make up the flora; four localities are included in this study where these distributions cannot be used: the two samples from the litera- ture (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 distri- bution of 613C in soil carbonate, an isotopic frac- tionation (103 in a) of -10 .36 [54] at 25°C for CO2-CaCO 3 will be used. Using this to calculate the isotopic composition of soil carbonate in equi- librium with the CO 2 of the pure component end- members, 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 composi- tion of soil carbonate and the fraction of C 4 bio- mass 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 I I I I

5 ioo

-"1 § o

~ - I 0 -

-151 0 I L I i 1 t I • . 5 0

F R A C T I O N C 4 F L O R A

Fig. 3. Carbon isotopic compos i t ion of soil carbonate com- pared to est imate of the fraction of C 4 plants in local flora (data from Table 1), M i x i n g l ines of calcite in equi l ibr ium with soil CO 2 and a tmospher ic CO 2 shown for reference. These are ca lculated us ing an isotopic fract ionat ion 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 carbon isotopic compos i t i on for soil CO 2 f rom a 100% C 3 flora, soil C O 2 from a 100% C4 flora, and a tmospher ic CO 2, respectively. N u m b e r s refer to local i ty num- bers in Table 1.

3.4. Implications for soil carbonate formation

Soil carbonate formation is generally consid- ered 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 evapotran- spiration is the dominant mechanism for soil water loss in areas covered by vegetation [2,55]. Few studies have examined the effect of evapotrans- piration on the isotopic composition of soil solu- tions 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 con- cerning 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 re- gions 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 in- crease in ion concentrations that results from ion exclusion during ice formation. Because soil respi- ration 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 pro- cess (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 atmo- spheric component of 60-80%.

235

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 com- position changes resulting from evapotranspira- tion, 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 forma- tion [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 composi- tion 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 de- pleted 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 grass- lands. "Normal continental" soil carbonates are represented in the shaded area, which is bounded by the 0-30% admixture of the atmospheric corn-

236

0

-5

~ -I0 o

-co

-15

-20

I I I I I I [ [

Station where P_<15Omm for all months "MONSOON EFFECT" ~ / Average for months where 117,1 o

PslSOmm / . ' , / / AVerwOhge;ef°;m°~ohmS~~'o / " !!

/ /

1 I ! ! --

Io I "} °1

I Northern Canadian Waters I .I

I

- 2 5 J I I I I [ F I -I0 -5 0 5 I0 15 20 25 30

T M (°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 months 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 HzO-CaCO 3 system at different tempera- tures.

Certain climatic conditions can result in devia- tions from this "normal" continental trend. Coastal stations in the Mediterranean and in Australia all

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- 2 0 -15 -IO - 5 O 5

or the rn Conctdion WeArs

~.oX\'%~ \ - ,% C'~ o .,. "0 - "% o ,,,> , /

u . I

g

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A. Continentel C. Monsoonel ]

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- 2 0 -15 - IO -5 O 5 ~18 ~CARBONATE

WPDB Fig. 5. A. Relat ionship between isotopic composition of water and carbonate using the temperature relationship of Craig [59]. Path of mean isotopic composition of water versus mean an- nual temperature (from Fig. 4) is plotted. Vectors of s ummer heating and isothermal evaporat ion are also shown. B. Fields of "normal" 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 "monsoonal" carbonates are also shown. See discussion in text.

have meteoric waters whose isotopic composition is several permil enriched in ~80 relative to con- tinental 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.

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 environ- ments; 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 nega- tive; 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 poten- tial 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 condi- tions (400,000 years) may persist. Sedimentation rates must take this into account: very low sedi- mentation rates (less than 1 cm/1000 years) may have an isotopic record that spans several different

237

climatic regimes; this is probably the case for many calcretes. Higher sedimentation rates (greater than 10 cm/1000 years) may result in the pre- servation 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 par- tial 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 excel- lent correspondence between 14C ages on calcretes and on associated organic material at Olduvai Gorge [25] is evidence that in some cases di- agenetic alteration does not appear to be im- portant.

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 paleocli- matic 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 ob- served differences in 813C contents of soil

238

c a r b o n a t e s is n o t d u e to c a r b o n a t e f o r m a t i o n in a c l o s e d sy s t em , 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 of soil C O 2 c a n r e su l t f r o m c h a n g e s in t he 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 in t h e p r o p o r t i o n o f C 4 b i o m a s s s e a s o n a l l y or o v e r a l o n g t i m e i n t e rva l . T h e o b s e r v e d d i f f e r e n c e be - t w e e n the 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 di- o x i d e a n d t he a s s o c i a t e d o r g a n i c m a t t e r in soil is e x p e c t e d b e c a u s e o f the d i f f e r e n c e in d i f f u s i o n c o e f f i c i e n t s for 13CO z and 12CO2. In a d d i t i o n , t he p r e s e n c e of 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 sugges t s t h a t f r eez ing of soil so lu- t i o n s m a y b e a n a d d i t i o n a l c a r b o n a t e f o r m a t i o n m e c h a n i s m .

Acknowledgements

G . W . Cox , S. G r e e n , R i . H a y , J .L. R i c h a r d - son , 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 he m o d e r n c a r b o n a t e s u s e d in th i s s t u d y ; T . M . B o w n p r o v i d e d a s s i s t a n c e in the f ield; 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 ance . I t h a n k F . H . B r o w n , M . C . M o n a g h a n , a n d W.T. P a r r y for m a n y h e l p f u l d i s c u s s i o n s . R e v i e w e r s i m p r o v e d t he c l a r i t y of th i s m a n u s c r i p t . T h i s w o r k ( i n c l u d i n g 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 F o u n d a t i o n , the F o u n d a t i o n fo r R e s e a r c h in t he O r i g i n of M a n , t he N a t i o n a l S c i e n c e F o u n d a t i o n ( 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 he 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 eleva- tion 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 drepano- lobiurn. 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 forma- tion 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 tempera- ture of about 7°C.

(11) Wyoming, U.S.A. I: Laminar calcrete coatings on peb- bles 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.

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