The Stable Carbon and Oxygen Isotopic Composition of Pedogenic Carbonate and its Relationship to Climate and Ecology in Southeastern Arizona
By Andrew L. Kowler December 17, 2007
ABSTRACT
18 The stable carbon (δ13C) and oxygen (δ O) isotopic composition of terrestrial carbonate has been used to reconstruct late Quaternary paleoecological and paleoclimatic conditions, respectively, for many different regions of the world. Quantitative reconstructions of past variability in climate and the distribution of C3/CAM/C4 vegetation from carbonate in soils and speleothems depend upon a rigorous examination of the modern soil isotopic system. To
18 accomplish this, we examined changes in the δ13C and δ O in relation to modern climatic and ecological conditions along an elevation gradient in southeastern Arizona. Five sites were
18 selected for study, spanning 1,170 m of elevation. Along this gradient, δ13C and δ O values
from ≥50 cm soil depth range from -9.9 to -0.6‰ and from -9.4 to -1.3‰, respectively.
Modeling results suggest that δ13C values were determined by soil respiration rates and the
13 proportion of C3/CAM/C4 biomass. For sites with low respiration, δ C values from >50 cm
reflected an atmospheric contribution of up to 55% compared to <20% for sites with much
higher respiration rates. At the lowest respiration sites, maximum observed δ18O values from
>50 cm diverge from minimum (winter) predicted values by +4.3 to +7.1‰, reflecting the
influence of evaporation. In contrast, values for the highest respiration rate site fell entirely
between those predicted from winter and summer rainfall. The latter finding suggests that a
significant proportion of carbonate may form during winter, and that there is a positive
correlation between respiration rate and the ratio of transpiration to evaporation accounting for
soil drying.
Results suggest that the reconstruction of absolute changes in vegetation composition from
carbonate isotopic composition will require quantification of the influence of soil respiration on
δ13C values. In turn, this knowledge can be used to quantify the maximum extent of evaporation
1
on δ18O values, and to reconstruct minimum shifts in the δ18O composition of meteoric water.
Conversely, if the δ18O value of paleo-precipitation is known, carbonate δ18O values might serve as a proxy for past respiration rates and to infer the magnitude of atmospheric contributions to
δ13C values.
2
INTRODUCTION
Background
18 The stable carbon (δ13C) and oxygen (δ O) composition of carbonate in soils and in speleothems has been used to reconstruct late Quaternary paleoecological and paleoclimatic
13 18 conditions for many different regions of the world. The δ C and δ O composition of pedogenic
carbonate has been extensively modeled and analyzed (Cerling, 1984; Quade et al., 1989a;
Cerling and Quade, 1993; Quade et al., 2007) and is reasonably well understood. Carbonate-
containing paleosols suitable for paleoenvironmental reconstruction are widely distributed across
arid and semi-arid regions, with records extending back several million years (e.g. Cerling et al.,
1989; Cerling and Hay, 1986; Quade et al., 1989b; Gabunia et al., 2000; Levin et al., 2004;
Quade et al., 2004). Further, recent advances in uranium-series dating of carbonate rinds
promise to provide key age-control on soil isotopic archives (Ludwig and Paces, 2002; Sharp et
al., 2003).
In contrast to soils, speleothems are attractive targets for paleoclimate reconstructions because
they are datable and may behave as closed systems (Quade, 2004), providing near-continuous records
of climate change with much finer temporal resolution than soils (McDermott, 2004). Because of the
specific geologic setting required for speleothem formation, such records are far less abundant than
18 paleosol records. The δ O composition of carbonate in modern cave speleothems has been carefully documented (e.g. Ayalon et al., 1998; Bar-Mathews et al., 1995; Schwarcz et al., 1976; Harmon,
1979; Goede et al., 1982), while their δ13C composition is poorly understood. Nonetheless,
speleothem δ13C values have been widely interpreted under the assumption that carbonate in caves behaves as it does in soils (e.g. Bar-Mathews et al., 1997, 1999; Baskaran and Krishnamurthy, 1993;
Brook et al., 1990; Denniston et al., 1999, 2000, 2001; Dorale et al., 1992, 1998), although this is not
3
13 necessarily the case (Quade, 2004). While understanding the soil carbonate δ C system is a primary
condition for understanding that of caves, the scope of our study is limited to treatment of the soil
system.
Previous work on soil carbonate in the southwestern US has produced records of
paleoenvironmental change spanning the late Pleistocene to late Holocene, including a record
from the Ajo Mountains of south-central Arizona (Liu et al., 1996), and others from the Organ
Mountains of southern New Mexico (Buck and Monger, 1999; Cole and Monger, 1994; Monger
et al., 1998). However, soil records cannot be used to quantitatively reconstruct past ecological
conditions until the δ13C composition of pedogenic carbonate is understood in relation to modern environmental conditions. To this end, our study examines the carbon and oxygen isotopic system of modern soils with respect to modern ecological and climatic conditions in southeastern
Arizona.
More specifically, we seek to understand the isotopic composition of modern carbonate with
respect to changes in the distribution of C3, C4, and CAM vegetation along an elevation gradient in southeastern Arizona (Fig. 1). Using this approach, we can quantify for pedogenic carbonate
18 18 18 (1) variation in δ O composition (δ Opc) as a function of altitude, because the δ O value of
18 meteoric water (δ Omw) decreases with increase in elevation (Rozanski et al., 1993), and (2)
13 13 variation in δ C composition (δ Cpc) as a function of vegetation composition and soil
respiration rate. In turn, knowledge of these relationships will enable us to reconstruct past
climate and ecology from paleosol and related records.
4
Study Area
Field studies were conducted in the Basin and Range physiographic province in the
southwestern USA near Tucson, Arizona (Fig. 1). The sites studied range in elevation from 730-
1900 m above mean sea level, and are located across several mountain ranges, including the
Tucson and Catalina Mountains bounding the Tucson Basin, the Santa Rita Mountains to the
southeast, and the Huachuca Mountains further to the south.
Vegetation
The Basin and Range province of southeastern Arizona is complex ecologically as well as
climatically, straddling the boundary between the Sonoran and Chihuahuan Deserts and
containing elements of tropical, temperate, and arctic environments. Tropical grasses (C4) as
well as succulents and cacti (CAM) thrive in this strongly monsoonal environment, where
precipitation is a combination of summer and winter rains (Sheppard et al., 2002). Generally, the
proportion of C4 biomass diminishes as elevation increases, whereas C3 plants (trees, temperate
zone grasses, and most shrubs) are favored by cooler temperatures and greater winter precipitation.
Ecozones in the Tropical-Subtropical Desert biome of southeastern Arizona include:
Arizona Upland (650-1,100 m), semidesert grassland (1,000-2,000 m), interior chaparral (1,000-
2,000 m), oak (Quercus)–pine (Pinus) woodland (1,600-1,900 m), and Ponderosa pine (P. ponderosa) forest (>1,900 m) (Brown, 1982). The Arizona Upland subdivision is characterized
by C3 leguminous trees and shrubs, various cacti and succulents (CAM), and C4 grasses.
Dominant genera include the shrubs Larrea and Acacia, succulent Fouquieria splendens, and
grasses Aristida and Trichachne. South of Tucson, the upper limit of the Arizona Upland
5
ecozone merges with semidesert grassland. This ecozone extends from 1,000-2,000 m and can
be characterized as a C4 grassland expanse broken by shrubs, cacti, and succulents. Dominant genera include the C4 grasses Bouteloua, Aristida, and Trichachne in addition to CAM
succulents Opuntia, Yucca and Agave and the C3 shrub Prosopis. Also within this elevation range, interior chaparral vegetation is typical of the droughtier soils of the foothills and mountain slopes where C3 shrub cover can reach 60-70%, typically including Prosopis, Fouquieria splendens, Quercus, and Arctostaphylos. Thus, C3 plant abundance exceeds that of C4 plants in chaparral enclaves that occur within extensive tracts of semidesert grassland. From 1,600-
1,900 m, oak-pine-juniper savannas and woodland (Quercus arizonica, Pinus edulis, and
Juniperus deppeana) with well-developed C4 grassland under story occur on coarser soils in the
rugged mountainous terrain. From 1,900-2,000 m, the oak-pine woodland grades into Ponderosa
pine savanna commonly containing Quercus (several species) and an under story of C4 grasses, including Bouteloua and Schizachyrium. In the transitional zone from 1,600-1,900 m, shrub- grass associations grade to grassy woodlands. Differences in aspect strongly influence the dominant plant type. On south-facing slopes, the abundance of CAM and C4 vegetation may rival that of C3 vegetation, while on north-facing slopes C3 vegetation typically dominates.
Climate
The primary cause of climatic variability in the Southwest is due to shifts in mid-latitude and subtropical atmospheric circulation regimes as well as proximity to major moisture sources
(Adams and Comrie, 1997). Further, local climate variations occur as a result of the region’s
extreme topography. The most prominent feature of Southwestern climate is the North
American monsoon, which in our study area provides at least 50% of annual precipitation during
6
the summer months. The rest falls mainly from October through March as frontal storms
ushered in by Pacific westerlies, or as dissipating tropical cyclones in September and October
(Sheppard et al., 2002 and references therein). Temperature is also strongly seasonal. Although
these seasonal patterns in precipitation and temperature are independent of elevation, adiabatic
cooling results in reduced temperatures and increased precipitation with elevation increase (Fig.
2).
We calculated temperature dependence on elevation for the Tucson Basin using urban heat
island-adjusted values from the University of Arizona Coop station at 745 m (U.S. HCN, 2007),
and the Palisades Coop station on Mt. Lemmon in the Santa Catalina Mountains at 2422 m
(Desert Research Institute, 2007) for the period from 1965-1981. These stations yield a lapse
rate of 6.6°C/km for winter and 7.3°C/km for summer. Mean annual air temperature (MAAT)
varies from 20.4°C at 745 m, to 9.2°C at 2,422 m, with corresponding mean annual precipitation
(MAP) ranging from 290 mm to 790 mm (Table 1). From this, we calculated MAAT and MAP lapse rates of 6.6°C/km (R2 = 1, n = 2) and 29.6 cm/km (R2 = 1, n = 2), respectively.
Isotopic Composition of Soil Carbonate
Calcite in soils is either detrital limestone, or calcite formed authigenically. We refer to the latter as “pedogenic carbonate”, with which we are exclusively concerned in this study.
Pedogenic carbonate (CaCO3) is composed of (1) carbon, originating from both biologically-
derived and atmospheric CO2 in the vadose zone, and (2) oxygen, originating from meteoric
water (Cerling, 1984). Precipitation and dissolution of calcium carbonate in soils can be
summarized as follows:
7
2+ – CO2 (g) + H2O (aq) + CaCO3 (s) Ca (aq) + 2HCO3 (aq) (1)
The precipitation of calcite in soils results from supersaturation of the soil solution with
respect to bicarbonate, associated with the gradual dewatering of soil by transpiration and
evaporation, and offgassing of CO2. Provided that pCO2 remains relatively stable during carbonate formation, the soil solution will remain in equilibrium with the gas phase, such that the carbon system can be considered open with respect to the carbon isotopic composition of soil
CO2. In contrast, oxygen isotopes in soil water will evolve during evaporation according to
simple Rayleigh fractionation leading up to calcite precipitation, such that the oxygen system can
be considered closed with respect to the oxygen isotopic composition of meteoric water.
Stable carbon and oxygen composition is reported in the δ (per mil) notation [relative to the
13 18 global Pee Dee Belemnite (PDB) standard] as δ C and δ O, respectively, where:
δ (per mil) = (Rsample/Rstandard – 1) x 1,000 (2)
13 12 18 16 and Rsample and Rstandard refer to the C/ C (and O/ O) ratios in a sample or standard, respectively.
Carbon
Cerling (1984), Quade et al. (1989a), and Cerling and Quade (1993) provide evidence that,
-7 -6 2 because the rate of formation of new pedogenic carbonate (~10 - 10 mole/cm /yr) is negligible
-3 -5 2 13 in comparison to the efflux of soil-respired CO2 (~10 - 10 mole/cm /yr), δ Cpc is not inherited
13 13 from parent material, but is instead controlled by the δ C composition of soil CO2 (δ Cconc). In
13 turn, δ Cconc is controlled mainly by (1) the proportion of atmospheric contribution to soil CO2
8
as determined by soil respiration rate and resultant soil pCO2 levels, and (2) the proportionate contributions of C3, C4, and CAM vegetation to soil respiration.
13 Distribution of δ C(CO2) in Soils
Soil CO2 is comprised of soil-respired and atmospheric CO2, the former increasing in proportion along a down-profile diffusion gradient (Cerling and Quade, 1993; Quade et al.,
1989a). Soil respiration is a combination of root and microbial respiration, the latter resulting from the microbially-mediated oxidation of organic matter. It has been found that the
13 proportion of atmospheric to soil-respired CO2 (δ Cresp) at a given depth is a function of net soil
respiration, where higher respiration rates result in a lower proportion of atmospheric CO2 at shallower depths (Quade et al., 1989a).
In addition, it is necessary to make a distinction between “soil CO2”, which is the instantaneous concentration of CO2 at a given depth in a soil quantified in units of ppmV, and
“soil-respired CO2”, which refers to the flux of CO2 passing through a horizontal plane in a soil,
2 quantified in units of mmol/m /hr. In a natural system, the mass transfer of soil gas is
12 controlled by diffusion; owing to the difference between the diffusion coefficients for CO2 and
13 13 13 CO2, δ Cconc should be 4.4‰ enriched relative to δ Cresp (Cerling et al., 1991).
13 We can illustrate the effects of soil respiration on δ Cconc at different depths using the
production-diffusion model of Cerling (1984), as revised by Quade et al. (2007). Previous
modeling by Cerling (1984) and Quade et al. (1989a) assumed CO2 production to be uniformly distributed throughout the upper 1 m of a soil profile, with a characteristic production depth (k) of 50 cm. Instead, the revised model of Quade et al. (2007) assumes a k of 22.5 cm with an exponential decrease in production with increasing soil depth.
9
13 We modeled δ Cconc for a hypothetical soil under a nearly pure stand of C3 vegetation with
13 an abundance-weighted δ C value of -24‰ (Fig. 3). In soils with respiration rates ≤4
2 13 mmol/m /hr, δ Cconc remains relatively constant below about 60 cm, whereas above this rate,
2 13 values do not vary much below about 40 cm. At ≥5 mmol/m /hr, variation in δ Cconc values at
100 cm soil depth is <1‰. Values are significantly higher at lower rates, reflecting a significant
atmospheric component.
13 Photosynthetic Pathway vs. δ Cresp
It is very difficult to determine the proportionate contribution of root and microbial
respiration to net soil respiration. Cerling (1984) and Quade et al. (1989a) assume that soil-
respired CO2 will bear the isotopic signature of overall root biomass, which is the dominant
13 source of organic matter fueling microbial activity. Thus, δ Cresp is determined by the relative
proportion of biomass from vegetation of the three major metabolic pathways. In turn, each of
13 these can be distinguished on the basis of the δ C composition of plant tissue:
13 • C3: trees, most shrubs, and temperate grasses, δ C = -27‰,
13 • C4: tropical grasses and shrubs of the genus Atriplex, δ C = -13‰, and
13 • CAM: succulents of the family Cactaceae and genus Agave, intermediate δ C values.
13 Temperature vs. δ Cpc
13 13 We can predict δ Cpc for a given temperature (T°C) by using the C isotopic
enrichment factor (ε) between CO2 and calcite from Romanek et al. (1992):
o εCaCO3-CO2 = 11.98(±0.13) – 0.12(±0.1)*T( C) (3)
10
o o 13 At 0 C, the enrichment factor is ~12‰, and at 25 C, δ Cpc is ~9‰.
13 13 Relationship Between δ Cpc and δ Cresp
The combined fractionation from diffusion and temperature-dependent fractionation results
13 13 o in a net enrichment of δ Cpc with respect to δ Cresp, between 16.4 and 13.4‰ for 0 and 25 C,
13 respectively (Quade, 2007). As a result, in soils with high respiration rates with δ Cresp values
13 of -27 and -12‰, δ Cpc ≈ —12‰ for the pure C3 end member and +2‰ for the pure C4 end member, respectively.
Oxygen
The oxygen isotopic composition of carbonate in modern soils has received less systematic
18 18 study than carbon isotopes. δ Opc is determined by δ Omw as affected by soil temperature and the extent of evaporation of soil water prior to carbonate formation.
18 Controls on δ Omw
Temperature
Rozanski et al. (1993) report a strong empirical relationship between the annual averages of
18 o δ Omw and local surface air temperature in the mid-latitudes (+0.7‰/ C). Examination of the
Tucson data set by Wagner (2006) revealed that interannual and intraseasonal temperature for
18 the winter months (October-March) was poorly correlated with δ Omw values. However, the
monthly data exhibit a weak correlation with temperature, with values being higher for the
o 2 summer and lower for the winter (Fig. 4), resulting in a slope of 0.14‰/ C (R = 0.03, p < 0.025).
18 o 2 In addition, summer monsoon (July-September) δ Omw values do increase 0.41‰/ C (R = 0.15,
11
p < 0.1) with increasing temperature, and the monthly values exhibit a similar trend (Fig. 4)
(Wagner, 2006).
Elevation
In any region with topographic relief, orographic precipitation will occur as a vapor mass
18 rises and cools adiabatically, thus driving rainout and O depletion of the resultant precipitation
(Rozanski et al., 1993). For southeastern Arizona, this depletion can be quantified as -1.6‰/km
for both summer (May-September) and winter (October-April) (Wahi, 2005). This estimate,
which was based on stable isotope time-series data in precipitation obtained from several high-
elevation sites in the region, differs significantly from the global average of -2.8‰/km reported
by Rozanski et al. (1993).
Precipitation Amount
Dansgaard (1964) first quantified the strong inverse relationship between mean monthly
18 δ Omw and the monthly precipitation amount which became known as the ‘amount effect’. As a
18 result of rainout, δ Omw values decrease during months with greater precipitation. In contrast, as a result of the lower relative humidity below the cloud base and resultant secondary evaporation
18 of raindrops, δ Omw values decrease during months with low precipitation.
18 For the Tucson Basin, the anomalously high δ O value for May rainfall reflects the evaporative enrichment of 18O in light rains occurring in the very dry foresummer. In addition,
18 18 the average δ Omw value for June is ≥3‰ lower compared to May δ Omw, even though the average precipitation is similar for both months. This can also be explained by the amount effect, as much of the June precipitation occurs during the latter part of the month in years with
12
early monsoonal activity, when increased relative humidity minimizes secondary evaporation.
Statistical analysis of the Tucson precipitation data suggests that the amount effect has a larger influence than temperature on the inter-annual and inter-seasonal variations (Wagner, 2006).
Further, the trend in the monthly and seasonal data for the summer are -0.2‰/10 mm, while for the winter the rate of decrease in the monthly data is more than four times greater than that of the
2 2 seasonal averages, -0.35 (R = 0.11, p < 0.01) versus –0.08‰/10 mm (R = 0.27, p < 0.025)
(Wagner, 2006).
18 Controls on δ Opc
Soil Temperature
Pedogenic carbonate forms in equilibrium with soil water, whose oxygen isotopic
18 composition (δ Osw) is related to that of meteoric water (Cerling, 1984; Quade et al., 1989a).
The oxygen isotopic composition of pedogenic carbonate relates that of soil water and its temperature during carbonate precipitation (Cerling and Quade, 1993 and references therein), as defined by Kim and O’Neil (1997):
1000lnα = (18030/T) – 32.42 (4) CaCO3-H2O where α is the equilibrium fractionation factor between pedogenic carbonate and soil CaCO3-H2O water for a specified T, which is in K. Large changes in soil temperature correspond to small
18 o o o o changes in δ Opc as follows: about -0.20‰/ C at 0 C, and -0.24‰/ C at 30 C.
Evaporation
Evaporation and transpiration both serve as dewatering mechanisms following partial or complete saturation of the soil by meteoric water. While evaporative loss results in kinetic
13
18 fractionation and enriched δ Osw values, transpiration does not. Thus, the extent of evaporation
depends upon the influence of transpiration on soil drying (Hsieh et al., 1998). The isotopic
evolution of soil water can be modeled as a function of evaporation by applying simple Rayleigh
18 fractionation to soil water with an initial δ O value corresponding to that of meteoric water (e.g.
Quade et al., 2007).
Allison et al. (1984) observed that water in barren desert soils experiences the greatest
18 degree of O enrichment compared to nearby vegetated soils. Other studies (e.g. Hsieh et al.,
18 1998; Liu et al., 1996) have shown that the greatest variability in the δ O value of soil water
18 18 occurs within the upper 40 cm, while δ Osw correlates well with δ Omw at greater soil depths. In a detailed study of soils along a humid-to-arid climatic gradient in Hawaii, Hsieh et al. (1998)
18 observed for the sites with highest rainfall that δ Osw near the surface was up to 2‰ greater than
18 δ Osw in the deepest part of the profile, while at the driest sites the difference reached +8‰.
Similar to the findings of Allison et al. (1984), in arid soils they found a greater discrepancy
18 18 between δ Osw and δ Omw values below 40 cm compared to humid soils, correlating this to
lower vegetation density and attributing it to (1) a reduced influence of transpiration in soil
18 drying, and (2) low frequency of recharge events, and thus minimal dilution of partially O-
18 enriched antecedent soil water by relatively O-depleted meteoric water.
18 In sum, the potential influence of evaporation on δ Osw is regulated by the relative influence
of transpiration on soil drying, which itself varies as a function of vegetation density as well as
differences in the rate of soil water uptake related to seasonality and depth. Additionally, the
degree of evaporative loss depends upon the rate of evaporation, determined by soil temperature
and humidity, as well as the fraction of initial water remaining at the time of calcite precipitation.
14
Site Descriptions
Setting
A general description of the study sites is given here, and a more detailed summary of site
characteristics can be found in Table 1. Tumamoc is the lowest elevation site, situated on a late
Holocene stream terrace inset into an alluvial fan toeslope at the base of Tumamoc Hill in the
Tucson Mountains. South of Tumamoc, Madera is positioned on a late Holocene stream terrace
inset within the lower section of Madera Canyon, which originates in the Santa Rita Mountains
and is at this location incised into a late Pleistocene alluvial fan surface. Hirabayashi, on an inset
stream terrace of estimated late Holocene age, is the northern most site in our study. Both Cave-
of-the-Bells (COB) sites are in the Santa Rita Mountains; COB West is located on the West side
of a limestone hill mantled with recent colluvial deposits, and COB Terrace is positioned on a
stream terrace of estimated late Holocene age. Finally, two sites are located in Garden Canyon
(GC) of the Huachuca Mountains; the GC Terrace site is on a late Holocene alluvial deposit in
the canyon bottom, and the GC South-facing (GC S-facing) sits on a south-facing middle or late
Holocene debris flow.
Vegetation
Each site was assigned to an appropriate ecozone in Table 1 reflecting the general vegetation
characteristics, while we give a more detailed description of site ecology in this section. We
begin with COB West, the only site with survey data. Vegetation at COB West can be described
as “limestone scrub”, comprised of >15% C4, >25% C3, and >25% CAM plants, while vegetation at COB Terrace has a less significant CAM component and can be characterized as a C3-
dominated oak-juniper-grassland (Quade, 2007, pers. com.). The average δ13C values measured
15
for C3 and C4 plants at both COB sites are -26.7 and -13.6‰, respectively (Quade, 2007, pers.
com.). These values are consistent with global averages of about -27 and -12‰, respectively
(e.g. Ehleringer, 1988). Although C4 and CAM plants are present at all sites, they are not dominant.
Because there is no plant survey available for the Tumamoc, Hirabayashi, Madera, and both
GC sites, our estimates of C3/C4 biomass are based on coarse visual estimation, assuming that percent cover is representative of total plant biomass, and assuming all grasses are C4.
Vegetation at Tumamoc can be best described as creosote bush-grassland with significant mesquite and CAM components. Cover of trees and shrubs versus grasses, and thus C3/C4 proportions, appear similar among these sites. Vegetation at the Madera locality is mesquite- grassland with a significant ocotillo component. The Hirabayashi site is open oak-juniper
woodland with significant CAM and C4 grass cover (oak-juniper-grassland). Here, the distribution, height, and density of grasses, trees, and shrubs closely resemble conditions at COB
West, for which survey data reveals equivalent biomass for C3 and CAM (Quade, 2007, pers. com.). Vegetation at GC S-facing can also be characterized as oak-juniper-grassland, with the proportionate cover of trees to shrubs exceeding that at the Madera and Hirabayashi sites.
Vegetation at GC Terrace can be characterized as a closed-canopy Ponderosa-grassland, though it also contains some oak and juniper influence. Although C3/C4 composition appears similar for both GC sites, CAM vegetation is much more common on the south-facing slope, where it might account for as much as 25% of plant cover.
16
METHODS
Site Selection
In order to assure that we sampled carbonate formed in association with modern vegetation
and within the modern climate-meteoric water system, we confined our sampling to stream
terraces adjacent to, and therefore recently abandoned by, active washes. Consistent with
previous studies, we assume that soils which are characterized by incipient carbonate
development, and which have been recognized by Gile et al. (1966) and Machette (1985) as
Stage I, were formed within the last 1-5 ka. Further, by focusing on terrace settings we hoped to
minimize microclimatic and ecological variations associated with differences in aspect and
drainage.
Soil Sampling
We sought to obtain diffusion profiles in order to assess the significance of atmospheric
13 contributions to δ Cpc values, and whether these values are reflective of modern vegetation
composition. To accomplish this, we sampled at approximately 20 cm depth increments to ≥1 m
where possible, obtaining at least one full isotopic profile for each of five elevations along our
gradient. In addition to examining modern carbonate, we examined carbonate from paleosols
formed on basalt flows mantling the north and south slopes of Tumamoc Hill in the Tucson
Mountains, where the Tumamoc site is situated, to assess the long-term effects of aspect on
13 δ Cpc values. Lastly, all of the samples used for the elevation transect were collected from ≥50
13 cm to ensure that the signal we obtain from δ Cpc values is dominantly ecological, and not atmospheric.
17
In the field, we collected pedogenic nodules and clast coatings from freshly exposed trench faces and arroyo walls. We encountered two types of Stage I carbonate. In some soils, carbonate occurs as a soft, powdery coating on the underside of a clast, while in others it occurs in others as dense cement with a powder-coated exterior. We thoroughly scrubbed coatings prior to rinsing with deionized water, while abstaining from scrubbing the powdery samples in order to prevent inadvertent removal of the carbonate. Previous studies indicate that the cement coatings are more purely pedogenic in composition than the powder coatings, which can contain considerable amounts of detritus from the surrounding soil matrix (Amundson et al., 1988).
Also at each site, we collected soil organic matter (SOM) in order to measure its carbon
13 13 isotopic composition (δ Csom), from which to calculate a long-term average value for δ Cresp.
13 13 However, there are several problems with using δ Csom to estimate δ Cresp. For one, SOM may
not be well mixed across the site such that single plants may be overrepresented within a given
sample. To accommodate this, we amalgamated the <2mm fraction of 4-5 samples collected at
each site, then homogenized and ground the remaining fine fraction. Next, because the δ13C value of organic matter at depth can be altered by microbial activity (Nadelhoffer and Fry, 1988;
Biggs et al., 2002; Beidenbender et al., 2004), we confined our sampling to the uppermost 5 cm of mineral soil.
Vegetation Survey and Collection
Plant species and density were recorded along a line transect 80 m long at COB West, and accounted for canopy interception within 5 cm on either side of the measuring tape.
18
Laboratory
In the laboratory, carbonate samples were heated at 250oC for 3 hours in vacuuo, then
processed using an automated sample preparation device (Kiel III) fitted to a Finnigan MAT 252
18 13 mass spectrometer at the University of Arizona Stable Isotope Laboratory. δ O and δ C values
were normalized to NBS-19 based on internal lab standards. Precision of repeated standards are
18 13 ± 0.11‰ for δ O and ± 0.07‰ for δ C (1σ).
SOM samples were passed through a 250 µm mesh sieve, amalgamated, then pretreated with
2M HCl and repeatedly rinsed with deionized water. CO2 extraction was accomplished by in vacuuo combustion of samples at 900oC for 3 hours in the presence of silver foil and copper
oxide powder; we then determined % organic C manometrically as part of off-line purification.
Internal lab standards are calibrated relative to NBS-22 and USGS-24, and precision of repeated
internal standards was ± 0.09‰ for δ13C (1σ). Next, for measurement of δ13C values of plant tissue, three individuals from each plant—stem, root, and leaf—were ground, homogenized, and their δ13C composition measured with a Costech automated CHN analyzer connected to a
Finnigan Delta-plus XL continuous-flow mass spectrometer.
18 Calculation of Predicted δ Opc Values
18 We predicted the mean and range of δ Opc values that should result for carbonate formed in
equilibrium with local meteoric water along our elevation gradient by applying αCaCO3-H2O values
18 (calculated for winter and summer) to corresponding δ Omw values. For this computation, we
18 derived equations for mean seasonal temperature (previously shown) and δ Omw versus
elevation, where winter is defined as October-April and summer as May-September. To
18 calculate δ Omw values, we used long-term volume-weighted seasonal means compiled for the
19
Tucson Basin from Eastoe (1998) for the period from 1981-2004/5, in combination with a local
lapse rate of -1.6‰/km for both winter and summer derived by Wahi (2005), using the Tucson
data set in addition to data from higher elevations (Wahi, 2005 and references therein).
RESULTS
Modern Soils
18 13 In all, we measured δ Opc and δ Cpc of Holocene carbonate from one or more sites at each of 5 elevation stops, spanning a 1,170 m elevation gradient from 730 to 1900 m (Fig. 1; Table 1).
18 13 For the elevation transect, we report ranges of δ Opc and δ Cpc values based on samples
13 obtained from >50 cm depth. Such values are the most likely to reflect δ C composition of CO2 derived from soil respiration. We have also obtained depth profiles at all localities except at
COB, where we report the results of carbonate sampled from 60 cm soil depth (Quade, 2007, pers. com.).
Oxygen
18 Rainfall δ Omw values (y, in ‰) are related to elevation (x, in meters) by:
y = -10,638x – 103,295 (winter)
y = -12,664x – 104,865 (summer)
18 From this, we calculated δ Opc values (y, in ‰) with elevation (x, in meters) by (Fig. 6): y = -625x – 5068 (winter) y = -625x - 2568 (summer)
20
Elevation Transect
18 The δ Opc composition of Holocene carbonate below 50 cm ranges from -1.3‰ to -9.4‰, with both maximum and minimum values from 1,900 m, respectively (Fig. 6). At the Tumamoc locality, values from 4 profiles ranged from -2.7 to -5.8‰; two profiles at Madera ranged from
-6.5 to -8.6‰; 1 profile at Hirabayashi, from -8.7 to -9.3‰; 2 sites at COB with 1 profile each
(Terrace and West), from -5.6 to -6.3‰ and -8.1 to -8.5‰, respectively; finally, at GC one profile at the Terrace site yielded values from -6.6 to -9.4‰, while 4 profiles from the S-facing site yielded values from -5.4 to -8.5 ‰, with the exception of the highest value at -1.3‰.
Depth Profiles
At Tumamoc, we sampled carbonate from 10 to 155 cm, and values among 4
stratigraphically similar profiles range from -2.4 to -5.8‰ (Fig. 7). From 45 to 115 cm, the
central section is bounded by abrupt contacts, with moderate B horizon development in the upper
part. This section yielded values between -2.4 and -5.8‰; in the upper section, from -5.2 to
-4.0‰; and in the lower section, from -3.7 to -4.9‰. At Madera, values from 2 profiles range
from -6.5 to -9.8‰ between 60 and 170 cm (Fig. 8). At Hirabayashi, we sampled 1 profile (25
to75 cm) with values between -7.3 and -9.3‰ (Fig. 9). From the GC Terrace site, we obtained
one profile (15 to 95 cm) with values ranging from -6.6 to -9.4‰ (Fig. 10) and at the S-facing site, 4 profiles collectively yielded values of -5.4 to -9.1‰ from 25-180 cm, with the exception of the highest value at -1.3‰ (Fig. 11).
21
Carbon
Elevation Transect
13 δ Cpc values in Holocene carbonates from ≥50 cm soil depth range from -0.6‰ to -9.9‰
(Fig. 12; Table 1). At Tumamoc, values ranged from -1.2 to -4.5‰; at Madera, from -6.6 to
-9.8‰; at Hirabayashi, from -7.3 to -7.8‰; for the Terrace and West sites at COB, from -0.7 to
-0.8‰ and between -0.6 and -1.2‰, respectively (Fischer et al. in preparation, 2006); and lastly,
at the GC Terrace and S-facing sites, from -4.2 to -9.6‰, and from -0.5 to -6.6‰, respectively.
Depth Profiles
At Tumamoc, values range from -1.2 to -6.3‰ (Fig. 7). Values from the center section
range between -2.0 and -4.1‰; in the upper section, from -1.9 to -6.3‰; and in the lower
section, from -1.6 to -3.5‰. Between the 2 profiles sampled at Madera (Fig. 8), values range
13 from -6.7 to -9.9‰. At the Hirabayashi site, δ Cpc values range from -7.3 to -7.9‰ (Fig. 9).
For the GC Terrace site, values range from -5.2 to -9.6‰. Finally, carbonate from the S-facing
site yielded values from -0.6 to -7.1‰ (Fig. 11).
Organic Matter, CO2, and Plant Biomass
13 The δ C values for SOM, soil CO2, and plant biomass listed below are unadjusted measured
values, whereas corresponding values reported in Table 1 reflect adjustments which are
13 explained in the discussion section. δ Cconc for Tumamoc is -14.8‰; for Hirabayashi, -21.1‰;
13 for COB West, -17.5‰; and for GC S-facing, -16.6‰. δ Csom values range from -18.8 to
-22.9‰, and are lowest for the lowest elevation sites at -22.9 and -23.2‰, for Tumamoc and
13 Hirabayashi, respectively. For the COB Terrace and West sites, δ Csom values are -18.9 and
22
-19.0‰, respectively, and -18.8‰ for the GC S-facing site. Finally, the abundance-weighted
13 13 δ C value of modern plant biomass (δ Cbulkplant) was estimated from the measured isotopic
13 composition of each plant, weighted by the fraction canopy cover for that plant. The δ Cbulkplant
13 value for COB West is -23.2‰. δ Cconc values reported for the COB and Tumamoc localities are averages, calculated from successive monthly measurements made over the course of a year
(Quade, 2007, pers. com.), whereas values reported for the other sites reflect a single measurement made in early March of 2007.
DISCUSSION
Carbon
13 It is our ultimate goal to relate δ Cpc to C3/CAM/C4 abundance at each site, in effect
establishing a carbon and oxygen isotopic signature for the individual plant associations that
characterize distinct bioclimatic envelopes. As discussed and modeled earlier, the atmospheric
contribution to soil CO2 is significantly diminished below a certain soil depth. For a given soil, the depth at which this occurs and the significance of atmospheric influence below this depth both depends largely upon the respiration rate. To quantify the atmospheric component, we must
13 first determine the respiration rate, for which knowledge of δ Cresp is necessary. A simple
13 13 comparison between δ Cpc (≥50 cm) and δ Csom values suggests a significant atmospheric
13 influence on δ Cpc values (Fig. 13) for several of our study sites.
For some sites, we have sufficient data to quantitatively assess the relative importance of
13 respiration rate in determining δ Cpc values. We can accomplish this by examining the various
13 components of the carbon isotope system, including vegetation (δ Cbulkplant), organic matter
13 13 13 13 (δ Csom), efflux (δ Cresp), soil gas (δ Cconc), and carbonate (δ Cpc). Our current discussion is
23
limited to the set of measured values compiled for each site. Under ideal circumstances, the
13 δ Cresp value is best estimated from soil CO2 fluxes collected throughout the year; however, we
13 do not currently have direct measurements of δ Cresp. In lieu of this, we follow the suggestion of
13 13 Cerling et al. (1991) in using δ Csom as a proxy for δ Cresp. Following the approach of Quade et
al. (2007), our respiration rate estimates assume:
13 13 13 • δ Cresp = δ Csom, after applying a +0.36‰ Suess Effect adjustment to δ Csom
• k = 22.5 cm
• porosity = 0.5
• pressure as calculated from site-specific elevation and MAAT
• specified soil depths
• carbonate formed under similar-to-modern vegetation
To estimate k, Quade et al. (2007) relied on the depth distribution of root biomass, in lieu of
year-round measurements of soil CO2 at different soil depths, assuming that root-respired CO2 and most the dead tissue which fuels microbial respiration are both proportional to overall root biomass. Further, Richter (1987) provides evidence that CO2 production in soils is exponentially decreasing with increasing depth. By fitting a simple exponential function to root density in 19 desert soil profiles examined by Schenk and Jackson (2002), Quade et al. (2007) estimated k to be 22.5 cm. In addition, Quade et al. (2007) calculated that the Suess Effect causes a 1.5 ‰
13 decrease in the δ C value of atmospheric CO2 and a corresponding change in plant tissue,
13 resulting in a ~0.36‰ decrease in δ Csom in desert soils. Soil respiration rates reported in this
study include this correction.
24
13 13 13 In Table 1, we report δ Cresp estimates calculated from δ Csom and δ Cbulkplant in the fashion
13 13 described above, as well as δ Cconc and the average of δ Cpc values from ≥50 cm depth. We
13 begin with the most complete profile sampled at Tumamoc, where δ Cpc values decrease with
13 depth and resemble a CO2 diffusion profile (Fig. 7). An estimated δ Cresp value of -22.5‰ and a
2 13 respiration rate of 0.22 mmol/m /hr yield a modeled δ Cpc profile that corresponds to observed
values from this profile. At 0.53 mmol/m2/hr, our modeled values intercept a single point. We
2 then modeled soil CO2 using a respiration rate of 0.41 mmol/m /hr to correspond to modern soil
13 gas (adjusted for the Suess Effect). Measured δ Cconc falls within the range of values
13 13 13 corresponding to observed δ Cpc values when we use δ Csom to estimate δ Cresp. We conclude
13 that low soil respiration rates can explain δ Cpc values at Tumamoc, which initially seemed
higher than what we might have expected (-12 to -9‰) at a C3-dominated site. This conclusion is consistent with the findings of Quade et al. (1989a), who estimated a respiration rate of 0.18 mmol/m2/hr for a low elevation site in Death Valley (300 m). Further, we found that values from all Tumamoc sites fit within a range between the two modeled diffusion profiles to which the
13 estimated respiration rates correspond. Based on the model, observed δ Cpc values from >60 cm
below the buried surface at Tumamoc reflect a ~30-55% atmospheric component.
13 At Hirabayashi, we estimate a δ Cresp value of -22.8‰ and modeled a respiration rate of
2 13 2-4 mmol/m /hr, concordant with observed δ Cpc values (Fig. 9) and much higher than that
2 13 modeled for Tumamoc. A rate of 1.6 mmol/m /hr results in a theoretical δ Cpc profile that
intercepts the theoretical value of carbonate formed in equilibrium with modern soil gas, in
13 agreement with the estimate based on carbonate values. Results of modeling with the δ Csom
13 value show that the deepest observed δ Cpc values reflect a ~0-20% atmospheric component.
25
13 13 For COB West, we compared estimates of δ Cresp obtained from both δ Cbulkplant and
13 13 13 δ Csom values. We did not model diffusion profiles for COB Terrace because δ Csom and δ Cpc values there are similar to COB West, such that resultant respiration rates and % atmospheric
13 composition will also be similar. Using δ Cbulkplant, we obtained an estimate of respiration rate
13 13 from comparison between δ Csom and δ Cpc values (Fig. 14), as well as from a comparison
13 13 13 between δ Cbulkplant and δ Cconc values. The mean δ Cconc value (corrected for the Suess Effect) is 2.2‰ lower than the mean value of soil CO2 in which carbonate ≥50 cm formed (calculated
13 13 13 from δ Cpc values), and δ Cbulkplant is 3.1‰ lower than δ Csom (both corrected for the Suess
Effect). This finding suggests that either 1) the proportion of C4 and/or CAM biomass has
13 recently decreased, or 2) δ Cbulkplant values are biased toward the C3 component with respect to overall plant biomass.
Alternatively, the observed discrepancy might reflect regular variation in C3/CAM/C4
13 proportions during the late Holocene, in which case such a deviation from δ Csom (taken to represent average late Holocene conditions) is not unique to the modern system. In this case,
13 13 13 δ Csom, and not δ Cbulkplant, will yield the most accurate estimate of the long-term δ Cresp
13 relating to δ Cpc values. This holds true if organic matter turnover rates correspond to the length of time represented by the modern carbonate in our soils, most likely on the order of 102-103 years. In addition, the respiration rates we calculated for the modern and late Holocene analogues are similar: 0.58 and 0.37 mmol/m2/hr, respectively. We suggest that a low soil
13 respiration rate is a plausible explanation for the higher-than-expected δ Cpc values at COB, and
13 calculated that δ Cpc values (≥50 cm) are 35-45% atmospheric in composition, similar to
Tumamoc values.
26
At Madera (Fig. 8), GC S-facing (Fig. 11), and GC Terrace (Fig. 10), we were able to obtain
13 13 δ Cpc diffusion profiles. The profile obtained for one of the Madera profiles has δ Cpc values
13 confined to a 0.3‰ range, making it ideal for our modeling exercise. Until we have δ Csom measurements for Madera and GC Terrace profiles, however, we cannot estimate respiration
13 13 rates for these sites. By comparing δ Csom with δ Cpc values corresponding to the most
13 complete profile obtained at the GC S-facing site (Fig. 11), we estimated a δ Cresp value of -
2 13 18.44‰. With a respiration rate of 8 mmol/m /hr, we obtained theoretical δ Cpc values that
correspond to the upper limit of observed values at this profile as well as for most values
13 observed for the other profiles at this site (Fig. 11). We conclude that the true δ Cresp is several
13 ‰ more negative than that estimated from δ Csom, and that the true respiration rate is
2 13 significantly lower than 8 mmol/m /hr. Without a reliable δ Cresp estimate, however, we cannot
13 determine the % atmospheric contribution to δ Cpc values.
Oxygen
18 At, δ Opc values at both GC sites, COB Terrace, Hirabayashi, and Tumamoc correlate poorly with theoretical values predicted for carbonate formed in equilibrium with summer or winter precipitation (Figs. 6 & 15). We can attribute these differences to (1) a discrepancy
18 18 18 between calculated δ Omw values and real δ Omw values due to local amount effects, or (2) O- enrichment of the soil water due to evaporative loss prior to carbonate formation. Because mean
18 seasonal precipitation δ Omw values are not dominated by large, low frequency events which
average out over time, pronounced differences in storm frequency or intensity can likely be ruled
18 out as a source of variability (Wright, 2001). Still, with long-term δ Omw data from rainfall
18 available for so few sites, it is not possible to directly examine δ Omw in relation to elevation
27
and topography at all sites. Instead, δ18O values of surface groundwater—sampled near GC and
along a down slope elevation gradient into the Upper San Pedro Valley—can serve as a proxy for
18 δ Omw (Wahi, 2005). Groundwater values fall between predicted values for summer and winter
18 18 δ Omw, suggesting the same for δ Omw values at corresponding elevations. We conclude that
18 elevated δ Opc values, which are observed for most of our sites, can be explained only by
evaporation.
We recall that the degree of influence of evaporation on soil drying is partly a function of
local aridity, in addition to the relative influence of transpiration on soil drying. Further, if the
ratio of transpiration to evaporation is highest in soils with the highest respiration rates, we
would also expect to observe minimal evaporation and enrichment in these soils. If effective
18 moisture were the main control on evaporation, we would expect δ Opc values to follow a
13 negative linear trend with elevation (Fig. 6). At COB Terrace, however, observed δ Cpc values diverge from predicted values by a similar amount as Tumamoc (Figs. 6 & 15). While these sites are climatically dissimilar, they both have similarly low respiration rates, suggesting that the ratio of transpiration to evaporation, and not precipitation amount, is the dominant control on evaporation within the elevation range studied. Further evidence for this is provided by the remarkable concordance between observed and predicted values at Hirabayashi, for which we estimated high respiration rates.
Aspect
At the COB and GC localities, we sampled soils on hillslopes as well as stream terraces.
18 Observed δ Opc values at GC S-facing (Fig. 6) align with our expectation that carbonate formed in soil on a south-facing slope will be slightly more enriched in 18O due to better drainage, lower
28
13 relative humidity in the soil, and lower respiration rates. δ Cpc values at GC S-facing are more
positive than at GC Terrace, reflecting lower respiration rates and/or increased CAM and/or C4
18 proportions. We suspect that transpiration rates play a significant role in determining δ Opc values. However, this cannot be demonstrated conclusively in the absence of reliable respiration rates, which might be used as a proxy for the ratio of transpiration to evaporation.
13 13 At COB West, δ Cpc and δ Csom values correspond well with values for COB Terrace,
18 suggesting that both sites have similar respiration rates. In contrast, δ Opc values are more
positive at COB Terrace, but the reason for this is not clear. It is also not clear whether this
effect is due to aspect-related differences in vegetation composition, soil respiration rate, or both.
To determine this, we will need to analyze soils from the S-facing slope. In addition, local soil
moisture conditions may differ significantly from aboveground conditions, and could potentially
be of greater importance than transpiration in mediating evaporation.
Finally, we sampled pedogenic carbonate from the north and south slopes of Tumamoc Hill,
where cemented carbonates formed in Quaternary paleosols which developed on basalt. We
observe no significant difference between carbon or oxygen values (Fig. 16). One possible
explanation for this is if carbonate forms dominantly during glacial periods, when C3
proportions, and probably respiration rates, are higher than currently observed. If this is true, then aspect-related differences in vegetation composition and soil respiration rates would be
13 minimized during glacial periods, resulting in the observed concordance of δ Cpc values.
Further, increased root transpiration activity that might result from increased annual precipitation
18 could moderate aspect-related differences in potential evaporation rates, resulting in δ Opc concordance.
29
SUMMARY AND CONCLUSIONS
13 It is clear from our results that the δ Cpc values of soil carbonate formed in middle-to-late
Holocene soils reflect the combined influence of vegetation composition and soil respiration, and
18 that δ Opc values also reflect changes in the ratio of transpiration to evaporation in addition to
18 13 δ Omw. Although modeled δ Cpc values remain relatively constant below 50 cm, values might still reflect a significant atmospheric contribution in soils with low respiration rates, such as
COB and Tumamoc, where respiration rates are estimated to be between 0.2 and 0.6 mmol/m2/hr
13 and δ Cpc values reflect a 35-45% and ≤55% atmospheric contribution, respectively.
18 At the same sites, δ Opc values are >2‰ more positive than predicted values. We interpret this deviation as the result of evaporative enrichment of 18O in soil water leading up to carbonate
precipitation, owing to the minimal influence of transpiration on soil dewatering and
13 corresponding to our interpretation of low respiration rates from lower-than-expected δ Cpc
18 values. In contrast, δ Opc values at the other sites are concordant with, or much closer to predicted values, likely reflecting the significant influence of transpiration on soil dewatering at
13 sites with sufficiently high respiration rates; this interpretation is supported by δ Cpc values from
Hirabayashi, from which we estimated a ≤20% atmospheric contribution.
13 13 At COB West, we observed a discrepancy between δ Cresp values estimated from δ Csom
13 13 13 versus δ Cbulkplant, corresponding to observed δ Cpc and δ Cconc values, respectively. This
13 13 13 suggests that δ Csom values reflect a δ Cresp value consistent with observed δ Cpc values, likely reflecting average ecological conditions for the late Holocene. In contrast, plant biomass values,
13 which are concordant with modern soil gas values, appear to reflect modern δ Cresp. Finally, we
13 18 detect no discernable difference between the δ Cpc or δ Opc composition of pedogenic calcretes from the north and south aspects of Tumamoc Hill. This suggests to us that calcrete formation
30
occurs dominantly during cooler periods when differences in plant composition and soil
respiration rate, and thus the ratio of transpiration to evaporation, are less pronounced than
during arid intervals.
This study leaves many questions unanswered. We did not examine the seasonal
13 distribution of pedogenic carbonate formation and its influence on δ Cpc at different depths, or
the influence of the depth distributions of C3 and C4 root systems in determining k. By
13 13 quantifying δ Cconc, soil T, pCO2, δ Cresp, and flux rate at different depths and times throughout the year, we will be able to directly determine k. In addition, direct measurement of soil
13 respiration will enable us to determine the degree of atmospheric contribution to δ Cpc as well as
13 13 13 to assess the accuracy of using δ Csom versus δ Cbulkplant as a proxy for δ Cresp. Further,
knowledge of seasonal respiration rates will enable us to quantify atmospheric contributions to
13 18 18 δ Cconc, and seasonal measurement of the δ O composition of soil CO2 (as a proxy for δ Osw)
will enable us to quantify the influence of evaporation on soil water, as well as the relationship
between respiration rate and the ratio of transpiration to evaporation.
In conclusion, our continued examination of the carbon and oxygen systems of soils will
enable us to accurately reconstruct proportionate C3/CAM/C4 biomass from pedogenic
carbonate. Our results suggest that, in order to reconstruct absolute changes in vegetation
13 composition, one must first quantify the influence of soil respiration on δ Cpc values. In turn,
8 this knowledge can be use to assess the influence of evaporation on δ Opc values for
18 reconstructing relative, and possibly absolute changes in δ Omw. Conversely, knowledge of the
18 18 δ O composition of paleo-precipitation can enable the use of δ Opc values to independently
13 assess the magnitude of atmospheric contributions to δ Cpc values, especially important in the
absence of information about past soil respiration rates.
31
ACKNOWLEDGMENTS
I would like to thank my committee members for all their advice and assistance. In addition,
I appreciate the extensive assistance I have received from David Dettman, Jeff Pigati, and
Warren Beck while conducting lab work and data analysis.
32
REFERENCES
Adams, D.K. and Comrie, A.C., 1997. The North American Monsoon. Bulletin of the American Meteorological Society: 2197-2213.
Allison, G.B., Barnes, C.J., Hughes, M.W., and Leany, F.W.J., 1984. The effect of climate and vegetation on the oxygen-18 and deuterium profiles in soils. In Isotope Hydrology 1983: International Atomic Energy Agency 1983: 105-123.
Amundson, R.G., Doner, H.E., Sowers, J., and Chadwick, O.A., 1988. The relationship between modern climate and vegetation and the stable isotope chemistry of Mojave desert soils. Soil Science Society of America Journal 115: 213-223.
Ayalon, A., Bar Mathews, M., Sass, E., 1998. Rainfall–recharge relationships within a karstic terrain in the eastern Mediterranean semi-arid region, Israel: δ18O and δD characteristics. Journal of Hydrology 207: 18–31.
Bar-Matthews, M., Ayalon, A., Kaufman, A., 1997. Late Quaternary palaeoclimate in the eastern Mediterranean region from stable isotope analysis of speleothems at Soreq Cave, Israel. Quaternary Research 47: 155–168.
Bar-Matthews, M., Ayalon, A., Kaufman, A., Wasserburg, G.J., 1999. The eastern Mediterranean palaeoclimate as a reflection of regional events: Soreq Cave, Israel. Earth and Planetary Science Letters 166: 85–95.
Bar-Mathews, M., Ayalon, A., Mathews, A., Sass, E., and Halicz, L., 1995. Carbon and oxygen isotope study of the active water-carbonate system in a karstic Mediterranean cave: implications for paleoclimate research in semiarid regions. Geochimica et Cosmochimica Acta 60: 337-347.
Baskaran, M. and Krishnamurthy, R.V., 1993. Speleothems as a proxy for the carbon isotope composition of atmospheric CO2. Geophysical Research Letters 20(7): 603-606.
Biedenbender, S., McClaran, M., Quade, J., and Weltz, M. A., 2004. Landscape patterns of vegetation change indicated by soil isotope composition. Geoderma 119: 69-83.
Biggs, T., Quade, J., and Webb, R. H., 2002. The δ13C value of soil organic matter in a semi-arid grassland with mesquite (Prosopis) encroachment in southeast Arizona. Geoderma 110: 109-130.
Brook, G.A., Burney, D.,A., and Cowart, J.B., 1990. Desert palaeoenvironmental data from cave speleothems with examples from the Chihuahuan, Somali-Chalbi, and Kalahari Deserts. Palaeogeography, Palaeoclimatology, Palaeoecology 76: 311-329.
Brown, 1982 (ed.). Desert Plants: Biotic Communities of the American Southwest-United States and Mexico. Boyce Thompson Southwestern Arboretum special issue 4(1-4). Tucson, University of Arizona Press.
33
Buck, B.J. and Monger, H.C., 1999. Stable isotopes and soil-geomorphology as indicators of Holocene climate change, northern Chihuahuan Desert. Journal of Arid Environments 43: 357- 373.
Cerling, T. E., 1984. The stable isotopic composition of modern soil carbonate and its relationship to climate. Earth and Planetary Science Letters 71: 229-240.
Cerling, T. E. and Hay, R.L., 1986. An isotopic study of paleosol carbonates from Olduvai Gorge. Quaternary Research 25: 63-78.
Cerling, T. E. and Quade, J., 1993. Stable carbon and oxygen isotopes in soil carbonates. In: Swart, P., McKenzie, J.A., and Lohman, K.C. (eds.): American Geophysical Union Monograph 78: 217- 231.
Cerling, T.E., Quade, J., Wang, Y., Bowman, J.R., 1989. Carbon isotopes in soils and paleosols as ecology and paleoecology indicators. Nature 341: 138-139.
Cerling, T. E., Solomon, D.K., Quade, J., and Bowman, J.R., 1991. On the isotopic composition of carbon in soil carbon dioxide. Geochimica et Cosmochimica Acta 55: 3403-3406.
Cole, D.R. and Monger, H.C., 1994. Influence of atmospheric CO2 on the decline of C4 plants during the last deglaciation. Nature 368: 533-536.
Dansgaard, W., 1964. Stable isotopes in precipitation. Tellus 16(4): 436-468.
Denniston, R.F., Gonzalez, L.A., Asmerom, Y., Baker, R.G., Reagan, M.K., and Bettis, E.A., 1999. Evidence for increased cool season moisture during the middle Holocene. Geology 27: 815-818.
Denniston, R.F., Gonzalez, L.A., Asmerom, Y., Polyak, V.J., Reagan, M.K., and Saltzman, M.R., 2001. A high-resolution speleothem record of climatic variability at the Allerod-Younger Dryas transition in Missouri, central United States. Palaeogeography, Palaeoclimatology, Palaeoecology 176: 147-155.
Denniston, R.F., Gonzalez, L.A., Asmerom, Y., Reagan, M.K., and Recelli-Snyder, H., 2000. Speleothem carbon isotopic records of Holocene environments in the Ozarks Highlands, USA. Quaternary International 67: 21-27.
Desert Research Institute, 2007. Western Regional Climate Center Historical Climate Information. Accessible at http://www.wrcc.dri.edu/CLIMATEDATA.html.
Dorale, J.A., Edwards, R.L., Ito, E., and Gonzalez, L.A., 1998. Climate and vegetation history of the mid continent from 75 to 25 ka: a speleothem record from Crevice Cave, Missouri, USA. Science 282: 1871-1874.
Dorale, J.A., Gonzalez, L.A., Reagan, M.K., Pickett, D.A., Murrell & Baker, R.G., 1992. A high- resolution record of Holocene climate change in speleothem calcite from Cold Water Cave, northeast Iowa. Science 258: 1626-1630.
34
Eastoe, C.J., 1998. Isotopic data for Tucson Basin: summary of work to 1998 (unpublished).
Ehleringer, J.R., 1988. Carbon isotope ratios and physiological processes in aridland plants. In Rundel, P.W., Ehleringer, J.R., and Nagy, K.A. (eds.): Symposium proceedings: Applications of stable isotope ratios to ecological research, pp. 41-54 (Springer-Verlag, New York).
Gabunia, L. and thirteen others, 2000. Earliest pleistocene hominid cranial remains from Dmanisi, Republic of Georgia: taxonomy, geological setting, and age. Science 288: 1019-1025.
Gile, L.H., Peterson, F.F., and Grossman, R.B., 1966. Morphological and genetic sequences of carbonate accumulation in desert soils. Soil Science 101: 347-360.
Goede, A., Green, D.C., and Harmon, R.S., 1982. Isotopic composition of precipitation, cave drips, and actively forming speleothems at three Tasmanian cave sites. Halictite 20: 17-27.
Harmon, R.S., 1979. An isotopic study of groundwater seepage in the central Kentucky karst. Water Resources Research 15(2): 476-480.
Hsieh , J.C., Chadwick, O.A., Kelly, E.F., and Savin, S.M., 1998. Oxygen isotopic composition of soil water: Quantifying evaporation and transpiration. Geoderma 82: 269-293.
Kim, S.T. and O’Neil, J. R., 1997. Equilibrium and non-equilibrium oxygen isotope effects in synthetic carbonates. Geochimica et Cosmochimica Acta 61: 3461-3475.
Levin, N., Quade, J., Simpson, S., Semaw, S., and Rogers, M., 2004. Isotopic evidence for Plio- Pleistocene environmental change at Gona, Ethiopia. Earth and Planetary Science Letters 219: 93-100.
Liu, B., Phillips, F. M., and Campbell, A. R., 1996. Stable carbon and oxygen isotopes of pedogenic carbonates, Ajo Mountains, southern Arizona: implications for paleoenvironmental change. Palaeogeography, Palaeoclimatology, Palaeoecology 124: 233-246.
Ludwig, K.R. and Paces, J.B., 2002. Uranium-series dating of pedogenic silica and carbonate, Crater Flat, Nevada: Geochimica et Cosmochimica Acta 66(3), 487-506.
Machette, M. A.,1985. Calcic soils of the southwestern United States. In Weide, D.L. and Faber, M.L. (eds.):.Soils and Quaternary geology of the southwestern United States. Geological Society of America Special Paper 203: 1-21.
McDermott, F., 2004. Palaeo-climate reconstruction from stable isotope variations in speleothems: a review. Quaternary Science Reviews 23: 901–918.
Monger, H.C., Cole, D.R., Gish, J.W., and Giordano, T.H., 1998. Stable carbon and oxygen isotopes in Quaternary pedogenic carbonates as indicators of ecogeomorphic changes in the northern Chihuahuan Desert, USA. Geoderma 82: 137-172.
35
Nadelhoffer, K. J., and Fry, B., 1988. Controls on natural nitrogen-15 and carbon-13 abundances in forest soil organic matter. Soil Science Society of America Jounal 52: 1633-1640.
Quade, J., 2004. Isotopic records from ground-water and cave speleothem calcite in North America. In Gillespie, A., Porter, S.C., and Atwater, B.F. (eds.): The Quaternary period in the United States, pp. 205-220 (Elsevier, New York).
Quade, J., Cerling, T.E., and Bowman, J.R., 1989a. Systematic variations in the carbon and oxygen isotopic composition of pedogenic carbonate along elevation transects in the southern Great Basin, USA. Geological Society of America Bulletin 101: 464-475.
Quade, J., Cerling, T.E., and Bowman, J.R., 1989b. Development of Asian monsoon revealed by marked ecological shift during the latest Miocene in northern Pakistan. Nature 342: 163-166.
Quade, J., Levin, N., Semaw, S., Simpson, S., Rogers, M., Stout, D., 2004. Paleoenvironments of the earliest toolmakers. Geological Society of America Bulletin 116: 1529-1544.
Quade, J., Rech, J.A., Latorre, C., Betancourt, J.L., Gleeson, E., and Kalin, M.T.K., 2007. Soils at the hyperarid margin: the isotopic composition of pedogenic carbonate from the Atacama desert, northern Chile. Geochimica et Cosmichimica Acta 71: 3772-3795.
Richter, J., 1987. The Soil as a Reactor (Catena Verlag, Cremlingen).
Romanek, C. S., Grossman, E. T., and Morse, J. W., 1992. Carbon isotopic fractionation in synthetic aragonite and calcite: effects of temperature and precipitation rate. Geochimica et Cosmochimica Acta 56: 419-430.
Rozanski, K., Aruguas-Araguas, L., and Gonfiantini, R., 1993. Isotopic patterns in modern global precipiation. In Swart, P., McKenzie, J.A., and Lohman, K.C. (eds.): Continental Indicators of Climate. American Geophysical Union Monograph 78, pp. 1-36.
Schenk, H. J. and Jackson, R. B., 2002. The global biogeography of roots. Ecological Monographs 72(3): 311–328.
Schwarcz, H.P., Harmon, R.S., Thompson, P., and Ford, D.C., 1976. Stable isotope studies of fluid inclusions in speleothems and their paleoclimatic significance. Geochimica et Cosmochimica Acta 40: 657-665.
Sharp, W.D., Ludwig, K.R., Chadwick, O.A., Amundson, R., Glaser, L., 2003. Dating fluvial terraces by 230Th/U on pedogenic carbonate, Wind River Basin, Wyoming: Quaternary Research 59: 139-150.
Sheppard, P. R., Comrie, A. C., Packin, G. D., Angersbach, K., and Hughes, M. K., 2002. The climate of the US Southwest. Climate Research 21: 219-238.
U.S. Historical Climatology Network (HCN), 2007. Monthly temperature and precipitation data. Available at http://cdiac.ornl.gov/epubs/ndp019/ndp019.html. 36
Wagner, J.D.M., 2006. Speleothem record of southern Arizona paleoclimate, 54 to 2.5 ka. Unpublished Ph.D. thesis, University of Arizona.
Wahi, A. K., 2005. Quantifying mountain system recharge in the Upper San Pedro Basin, Arizona. Unpublished M.S. thesis, University of Arizona.
Wright, W.E., 2001. δD and δ18O in mixed conifer systems in the U.S. Southwest: the potential of δ18O in Pinus ponderosa Tree Rings as a Natural Environmental Recorder. Unpublished M.S. thesis, University of Arizona.
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TABLE AND FIGURE CAPTIONS
Table 1. Elevation, location, MAAT, MAP, vegetation, and δ13C (VPDB) values for transect sites. δ13C values are measured for pedogenic carbonate, abundance-weighted 13 13 vegetation, and soil gas, and estimated for δ Cresp. δ Cpc values are the average of 13 values from ≥50 cm soil depth; δ Cconc values are corrected for the Suess Effect; and 13 13 13 δ Cresp values are calculated from δ Csom and δ Cbulkplant values (both Suess-corrected).
Figure 1. Location of soil study sites in and around the Tucson Basin, in southeastern Arizona.
Figure 2. MAAT and MAP for 3 stations in southeastern Arizona: (1) Tucson (740 m), (2) Santa Rita Experimental Range (1310 m), and (3) Mt. Lemmon in the Santa Catalina Mountains (2425 m).
13 Figure 3. δ Cconc (VPDB) as a function of soil depth (cm) for soil respiration rates ranging 2 from 0.1 to 9.0 mmol/m /hr, using the diffusion model for soil CO2 (Cerling, 1984, 1991; Quade, et al., 1989a; Quade et al., 2007). Boundary conditions include 0.9 atm total 13 pressure, δ Catmosphere = -6.5‰ with pCO2 = 270 ppmV for the pre-industrial atmosphere, o 13 free air porosity of 0.5, a tortuosity of 0.6, temperature = 21.7 C, δ Cresp = -24‰, and exponential production, where k = 22.5 cm.
18 Figure 4. Monthly weighted averages of δ Omw values (red circles, error bars are one standard deviation), average temperature (green squares), and average precipitation amount (blue triangles) for Tucson from 1981-2005, adapted from Wagner (2005).
13 18 Figure 5. δ Cpc (VPDB) versus δ Opc (VPDB) values from ≥50 cm in late Holocene soils in southeastern Arizona.
18 Figure 6. δ Opc (VPDB) from ≥50 cm versus site elevation, showing values of carbonate predicted to form in equilibrium with summer and winter precipitation (solid lines), for sites studied in southeastern Arizona.
18 13 Figure 7. δ Opc (VPDB) and δ Cpc (VPDB) versus depth for a late Holocene soil profile at 13 Tumamoc, for which δ Cpc is modeled (solid lines) for the respiration rates shown (mmol/m2/hr) using the model described in Quade et al. (2007) where elevation is 730 m, o 13 T = 22.1 C, δ = 0.5, and δ Cresp = -22.5‰. Open symbols represent values from other profiles at Tumamoc, for which respiration rate was not modeled.
18 13 Figure 8. δ Opc (VPDB) and δ Cpc (VPDB) versus depth for 2 late Holocene soil profiles at Madera, for which values are distinguished as solid versus open symbols.
18 13 Figure 9. δ Opc (VPDB) and δ Cpc (VPDB) versus depth for a late Holocene soil at 13 2 Hirabayashi. δ Cpc is modeled (solid lines) for the respiration rates shown (mmol/m /hr) o using the model described in Quade et al. (2007), where elevation is 1460 m, T = 17.0 C,
38
13 δ = 0.5, and δ Cresp = -22.8‰..
18 13 Figure 10. δ Opc (VPDB) and δ Cpc (VPDB) versus depth for a late Holocene soil profile at GC Terrace.
18 13 Figure 11. δ Opc (VPDB) and δ Cpc (VPDB) versus depth for a late Holocene 13 soil profile at GC S-facing, for which δ Cpc is modeled (solid lines) for the respiration rates shown (mmol/m2/hr) using the model described in Quade et al. (2007), where o 13 elevation is 1900 m, T = 13.9 C, δ = 0.5, and δ Cresp = -18.4‰. Open symbols represent values from other profiles at this site for which respiration rate was not modeled.
13 Figure 12. δ Cpc (VPDB) values from ≥50cm soil depth versus site elevation for late Holocene soils in southeastern Arizona.
13 13 Figure 13. δ Cpc (VPDB) versus δ Csom (VPDB) for soils studied in southeastern Arizona. The lower two solid lines represent modeled isotopic values defining high respiration-rate soils (10 mmoles/m2/hr) between 0 and 25°C. The dashed lines represent a range of soil 13 respiration rates (labeled) calculated for 25°C. Solid circles represent observed δ Cpc values. In general, these values plot above the field defined for high respiration-rate soils, suggesting a significant atmospheric component.
13 Figure 14. δ Cpc (VPDB) from 60 cm depth in a late Holocene soil at COB West, for which 13 2 δ Cpc is modeled (solid lines) for the respiration rates shown (mmol/m /hr) using the o model described in Quade et al. (2007) where elevation is 1650 m, T = 15.6 C, δ = 0.5, 13 and δ Cresp = -18.6‰.
18 Figure 15. Predicted versus observed δ Opc (VPDB) values, where predicted values are 18 calculated to form in equilibrium with mean annual δ Omw.
Figure 16. δ18O (VPDB) and δ13C (VPDB) values of pedogenic calcrete sampled from the north and south aspects of Tumamoc Hill.
39
Table 1. Summary of Sites
E δ δ L 1 1 E 3 3 δ V 1 M M C C 3 L A L E O C T O A A S C c r C I L A P E O o e O N n s p A A T T p c N G Z c (m T ( ( ( L T O ‰ ( ( I ‰ ‰ IT ( ( ° N N m N W C m Y G E ) ) ) ) ) ) ) )
SOM PLANT Tumamoc 730 31o 43.423’ 110o 46.041’ 22.1 280 Terrace creosote-grassland -13.3 -22.5 x -2.1
Madera 1105 32o 20.406’ 110o 43.044’ 19.4 358 Terrace mesquite-grassland x x x -8.7
Hirabayashi 1450 31o 27.411’ 110o 22.324’ 17.0 480 Terrace oak-juniper-grassland -16.7 -22.8 x -7.7
West Hillslope limestone scrub -13.1 -18.6 -21.8 -0.9 COB 166032o 13.298’ 111o 0.634’ 15.6 539 Terrace Terrace oak-juniper-grassland x -18.5 x -0.7
S-facing Hillslope oak-juiper-grassland -15.1 -18.4 x -5.0 GC 1900 31o 46.881’ 110o 53.176’ 13.9 605 Terrace Terrace Ponderosa-grassland x x x -7.4 Santa Catalina Mountains
Hirabayashi 1450 m
Tumamoc 730 m
Tucson Mountains
Tucson Basin
Madera 1105 m Cave-of-the-Bells 1660 m
Santa Rita Mountains
Garden Canyon 1900 m
Huachuca Mountains
Figure 1 35.0 35 Precipitation Mt. Lemmon Temperature 3030.0 30 (2425 m) 25.0 25 2020.0 1515.0 1010.0 55.0 00.0
35.0 35 Precipitation Temperature C) 30.0 Santa Ritas o 30 (1310 m) 2525.0
2020.0
ature ( 15.0
r 15
1010.0
55.0
0.0 Tempe 0 Precipitation (cm/month) 35.0 35 Precipitation Temperature 3030.0 Tucson (740 m) 2525.0 2020.0
1515.0
1010.0
55.0
0.00 r r y ry il y ly st e er ar rch pr une u b be b ber ua a A Ma J Ju g m to m anu br M te em J e Au Oc ve F ep ec J F M A M J J AS S ONo ND D
Figure 2 0
10 5.05.0 0 4.0. 20 6.06.0 4 7.07.0 30 8.08.0
) 40 m c
50 h ( pt e
Depth (cm 60 Depth (cm) D
70
80 3.0 2.0 1.0 0.5 0.1 0.05 0.01 0.1 0.5 3.0 2.0 1.0 0.01 90 0.05
100
-25 -20 -15 -10 -5 0 δ13C (VPDB) CO2
Figure 3 5
30
0 (VSMOW) 20 mperature (C) e T mw O
18 -5 δ 10 80
-10 0
40 on Precipitation wt. on Precipitation Amount (mm) Tucs 0 0 2 4 6 8 10 12 Month
Figure 4 18 δ Opc (VPDB) -12 -10 -8 -6 -4 -2 0 0
-2
-4
-6 (VPDB) pc C 13 δ -8 GC Terrace GC S-facing COB Terrace COB West Hirabayashi -10 Madera T umamoc
-12
Figure 5 2000
1800
1600 GC Terrace
) GC S-facing winter 1400
(m COB Terrace
n COB West o
1200 Hirabayashi
evati Madera l
E Tumamoc 1000 summer
800
600 -12 -10 -8 -6 -4 -2 0
18 δ Opc (VPDB)
Figure 6 0- pedogenic A structure 20- stratified carbon gravels oxygen 40- stratified alluvium 60- disconformity 2Bw2Bw
80-
100- 2C
120- 3C1
140- Depth (cm) 3C2 160- 8.0 0.22 0.53 180-
200-
------0 2 4 1 1 8 6 4 2 2 0 δ ‰ (VPDB) Figure 7 4 n en 2 g oxy carbo 0
-2
-4 VPDB) (
-6 ‰ δ -8
-10
-12 C A 0-
40- 60- 80- 20-
200- 160- 100- 140- 180- 120- Depth (cm) Depth d de d tified unbe gravels stra alluvium Figure 8