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
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The stable carbon (δ C) 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
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accomplish this, we examined changes in the δ C 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
proportion of C3/CAM/C4 biomass. For sites with low respiration, δ13C 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
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18
The stable carbon (δ C) and oxygen (δ O) composition of carbonate in soils and in
speleothems has been used to reconstruct late Quaternary paleoecological and paleoclimatic
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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. Carbonatecontaining 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,
13
speleothem δ C 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
(1) variation in δ18O composition (δ18Opc) as a function of altitude, because the δ18O value of
meteoric water (δ18Omw) decreases with increase in elevation (Rozanski et al., 1993), and (2)
variation in δ13C composition (δ13Cpc) 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 7301900 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 (1,0002,000 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 1,6001,900 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, shrubgrass 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 biologicallyderived 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+
–
Ca (aq) + 2HCO3 (aq)
CO2 (g) + H2O (aq) + CaCO3 (s)
(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
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global Pee Dee Belemnite (PDB) standard] as δ C and δ O, respectively, where:
(
)
δ (per mil) = Rsample/Rstandard – 1
18
x 1,000
(2)
16
and Rsample and Rstandard refer to the 13C/12C (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
from parent material, but is instead controlled by the δ C composition of soil CO2 (δ13Cconc). 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
proportion of atmospheric to soil-respired CO2 (δ13Cresp) 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
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13
CO2, δ Cconc should be 4.4‰ enriched relative to δ Cresp (Cerling et al., 1991).
We can illustrate the effects of soil respiration on δ13Cconc 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.
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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,
values do not vary much below about 40 cm. At ≥5 mmol/m2/hr, variation in δ13Cconc values at
100 cm soil depth is <1‰. Values are significantly higher at lower rates, reflecting a significant
atmospheric component.
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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 soilrespired CO2 will bear the isotopic signature of overall root biomass, which is the dominant
source of organic matter fueling microbial activity. Thus, δ13Cresp is determined by the relative
proportion of biomass from vegetation of the three major metabolic pathways. In turn, each of
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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‰,
•
C4: tropical grasses and shrubs of the genus Atriplex, δ C = -13‰, and
•
CAM: succulents of the family Cactaceae and genus Agave, intermediate δ C values.
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13
Temperature vs. δ13Cpc
13
We can predict δ13Cpc for a given temperature (T°C) by using the C isotopic
enrichment factor (ε) between CO2 and calcite from Romanek et al. (1992):
εCaCO3-CO2 =
11.98(±0.13) – 0.12(±0.1)*T(oC)
(3)
10
o
o
13
At 0 C, the enrichment factor is ~12‰, and at 25 C, δ Cpc is ~9‰.
Relationship Between δ13Cpc and δ13Cresp
The combined fractionation from diffusion and temperature-dependent fractionation results
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o
in a net enrichment of δ13Cpc with respect to δ Cresp, between 16.4 and 13.4‰ for 0 and 25 C,
respectively (Quade, 2007). As a result, in soils with high respiration rates with δ13Cresp 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
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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.
Controls on δ18Omw
Temperature
Rozanski et al. (1993) report a strong empirical relationship between the annual averages of
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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
In addition, summer monsoon (July-September) δ Omw values do increase 0.41‰/ C (R2 = 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 highelevation 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
δ18Omw and the monthly precipitation amount which became known as the ‘amount effect’. As a
result of rainout, δ18Omw 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
of raindrops, δ18Omw 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
the average δ Omw value for June is ≥3‰ lower compared to May δ18Omw, 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).
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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):
1000lnαCaCO -H O = (18030/T) – 32.42
3
(4)
2
where αCaCO -H O is the equilibrium fractionation factor between pedogenic carbonate and soil
3
2
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
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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 O18
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 Caveof-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 C3dominated 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 mesquitegrassland 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
δ13Cpc 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.
However, there are several problems with using δ13Csom to estimate δ13Cresp. 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
o
In the laboratory, carbonate samples were heated at 250 C 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
o
vacuuo combustion of samples at 900 C 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.
Calculation of Predicted δ18Opc 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
(calculated for winter and summer) to corresponding δ18Omw values. For this computation, we
derived equations for mean seasonal temperature (previously shown) and δ18Omw versus
elevation, where winter is defined as October-April and summer as May-September. To
calculate δ18Omw 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
In all, we measured δ18Opc and δ13Cpc 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).
For the elevation transect, we report ranges of δ18Opc and δ13Cpc values based on samples
obtained from >50 cm depth. Such values are the most likely to reflect δ13C 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
Rainfall δ18Omw values (y, in ‰) are related to elevation (x, in meters) by:
y = -10,638x – 103,295 (winter)
y = -12,664x – 104,865 (summer)
From this, we calculated δ18Opc values (y, in ‰) with elevation (x, in meters) by (Fig. 6):
y = -625x – 5068 (winter)
y = -625x - 2568 (summer)
20
Elevation Transect
The δ18Opc 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
δ13Cpc 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
from -6.7 to -9.9‰. At the Hirabayashi site, δ13Cpc 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
The δ13C 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
explained in the discussion section. δ13Cconc for Tumamoc is -14.8‰; for Hirabayashi, -21.1‰;
for COB West, -17.5‰; and for GC S-facing, -16.6‰. δ13Csom 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
Hirabayashi, respectively. For the COB Terrace and West sites, δ13Csom values are -18.9 and
22
-19.0‰, respectively, and -18.8‰ for the GC S-facing site. Finally, the abundance-weighted
δ13C value of modern plant biomass (δ13Cbulkplant) was estimated from the measured isotopic
composition of each plant, weighted by the fraction canopy cover for that plant. The δ13Cbulkplant
value for COB West is -23.2‰. δ13Cconc 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
It is our ultimate goal to relate δ13Cpc 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
first determine the respiration rate, for which knowledge of δ13Cresp is necessary. A simple
comparison between δ13Cpc (≥50 cm) and δ13Csom values suggests a significant atmospheric
influence on δ13Cpc 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
components of the carbon isotope system, including vegetation (δ13Cbulkplant), organic matter
13
(δ Csom), efflux (δ13Cresp), soil gas (δ13Cconc), and carbonate (δ13Cpc). Our current discussion is
23
limited to the set of measured values compiled for each site. Under ideal circumstances, the
δ13Cresp value is best estimated from soil CO2 fluxes collected throughout the year; however, we
do not currently have direct measurements of δ13Cresp. In lieu of this, we follow the suggestion of
13
Cerling et al. (1991) in using δ Csom as a proxy for δ13Cresp. Following the approach of Quade et
al. (2007), our respiration rate estimates assume:
•
δ13Cresp = δ13Csom, after applying a +0.36‰ Suess Effect adjustment to δ13Csom
•
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 ‰
decrease in the δ13C value of atmospheric CO2 and a corresponding change in plant tissue,
resulting in a ~0.36‰ decrease in δ13Csom in desert soils. Soil respiration rates reported in this
study include this correction.
24
In Table 1, we report δ13Cresp estimates calculated from δ13Csom and δ13Cbulkplant in the fashion
described above, as well as δ13Cconc and the average of δ13Cpc values from ≥50 cm depth. We
begin with the most complete profile sampled at Tumamoc, where δ13Cpc values decrease with
depth and resemble a CO2 diffusion profile (Fig. 7). An estimated δ13Cresp value of -22.5‰ and a
respiration rate of 0.22 mmol/m2/hr yield a modeled δ13Cpc profile that corresponds to observed
values from this profile. At 0.53 mmol/m2/hr, our modeled values intercept a single point. We
then modeled soil CO2 using a respiration rate of 0.41 mmol/m2/hr to correspond to modern soil
gas (adjusted for the Suess Effect). Measured δ13Cconc falls within the range of values
corresponding to observed δ13Cpc values when we use δ13Csom to estimate δ13Cresp. We conclude
that low soil respiration rates can explain δ13Cpc 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
estimated respiration rates correspond. Based on the model, observed δ13Cpc values from >60 cm
below the buried surface at Tumamoc reflect a ~30-55% atmospheric component.
At Hirabayashi, we estimate a δ13Cresp value of -22.8‰ and modeled a respiration rate of
2-4 mmol/m2/hr, concordant with observed δ13Cpc values (Fig. 9) and much higher than that
modeled for Tumamoc. A rate of 1.6 mmol/m2/hr results in a theoretical δ13Cpc profile that
intercepts the theoretical value of carbonate formed in equilibrium with modern soil gas, in
agreement with the estimate based on carbonate values. Results of modeling with the δ13Csom
value show that the deepest observed δ13Cpc values reflect a ~0-20% atmospheric component.
25
For COB West, we compared estimates of δ13Cresp obtained from both δ13Cbulkplant and
δ13Csom values. We did not model diffusion profiles for COB Terrace because δ13Csom and δ13Cpc
values there are similar to COB West, such that resultant respiration rates and % atmospheric
composition will also be similar. Using δ13Cbulkplant, we obtained an estimate of respiration rate
from comparison between δ13Csom and δ13Cpc values (Fig. 14), as well as from a comparison
between δ13Cbulkplant and δ13Cconc values. The mean δ13Cconc value (corrected for the Suess Effect)
is 2.2‰ lower than the mean value of soil CO2 in which carbonate ≥50 cm formed (calculated
from δ13Cpc values), and δ13Cbulkplant is 3.1‰ lower than δ13Csom (both corrected for the Suess
Effect). This finding suggests that either 1) the proportion of C4 and/or CAM biomass has
recently decreased, or 2) δ13Cbulkplant 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
proportions during the late Holocene, in which case such a deviation from δ13Csom (taken to
represent average late Holocene conditions) is not unique to the modern system. In this case,
δ13Csom, and not δ13Cbulkplant, will yield the most accurate estimate of the long-term δ13Cresp
relating to δ13Cpc 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
respiration rate is a plausible explanation for the higher-than-expected δ13Cpc values at COB, and
calculated that δ13Cpc 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
δ13Cpc diffusion profiles. The profile obtained for one of the Madera profiles has δ13Cpc values
confined to a 0.3‰ range, making it ideal for our modeling exercise. Until we have δ13Csom
measurements for Madera and GC Terrace profiles, however, we cannot estimate respiration
rates for these sites. By comparing δ13Csom with δ13Cpc values corresponding to the most
complete profile obtained at the GC S-facing site (Fig. 11), we estimated a δ13Cresp value of 18.44‰. With a respiration rate of 8 mmol/m2/hr, we obtained theoretical δ13Cpc values that
correspond to the upper limit of observed values at this profile as well as for most values
observed for the other profiles at this site (Fig. 11). We conclude that the true δ13Cresp is several
‰ more negative than that estimated from δ13Csom, and that the true respiration rate is
significantly lower than 8 mmol/m2/hr. Without a reliable δ13Cresp estimate, however, we cannot
determine the % atmospheric contribution to δ13Cpc values.
Oxygen
At, δ18Opc 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
between calculated δ18Omw values and real δ18Omw values due to local amount effects, or (2) 18Oenrichment of the soil water due to evaporative loss prior to carbonate formation. Because mean
seasonal precipitation δ18Omw 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
out as a source of variability (Wright, 2001). Still, with long-term δ18Omw data from rainfall
available for so few sites, it is not possible to directly examine δ18Omw 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
δ18Omw (Wahi, 2005). Groundwater values fall between predicted values for summer and winter
δ18Omw, suggesting the same for δ18Omw values at corresponding elevations. We conclude that
elevated δ18Opc 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
moisture were the main control on evaporation, we would expect δ18Opc values to follow a
negative linear trend with elevation (Fig. 6). At COB Terrace, however, observed δ13Cpc 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.
Observed δ18Opc 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
relative humidity in the soil, and lower respiration rates. δ13Cpc values at GC S-facing are more
positive than at GC Terrace, reflecting lower respiration rates and/or increased CAM and/or C4
proportions. We suspect that transpiration rates play a significant role in determining δ18Opc
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.
At COB West, δ13Cpc and δ13Csom values correspond well with values for COB Terrace,
suggesting that both sites have similar respiration rates. In contrast, δ18Opc 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
minimized during glacial periods, resulting in the observed concordance of δ13Cpc values.
Further, increased root transpiration activity that might result from increased annual precipitation
could moderate aspect-related differences in potential evaporation rates, resulting in δ18Opc
concordance.
29
SUMMARY AND CONCLUSIONS
It is clear from our results that the δ13Cpc values of soil carbonate formed in middle-to-late
Holocene soils reflect the combined influence of vegetation composition and soil respiration, and
that δ18Opc values also reflect changes in the ratio of transpiration to evaporation in addition to
δ18Omw. Although modeled δ13Cpc 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
and δ13Cpc values reflect a 35-45% and ≤55% atmospheric contribution, respectively.
At the same sites, δ18Opc 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
corresponding to our interpretation of low respiration rates from lower-than-expected δ13Cpc
values. In contrast, δ18Opc 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
sites with sufficiently high respiration rates; this interpretation is supported by δ13Cpc values from
Hirabayashi, from which we estimated a ≤20% atmospheric contribution.
At COB West, we observed a discrepancy between δ13Cresp values estimated from δ13Csom
versus δ13Cbulkplant, corresponding to observed δ13Cpc and δ13Cconc values, respectively. This
suggests that δ13Csom values reflect a δ13Cresp value consistent with observed δ13Cpc values, likely
reflecting average ecological conditions for the late Holocene. In contrast, plant biomass values,
which are concordant with modern soil gas values, appear to reflect modern δ13Cresp. Finally, we
detect no discernable difference between the δ13Cpc or δ18Opc 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
distribution of pedogenic carbonate formation and its influence on δ13Cpc at different depths, or
the influence of the depth distributions of C3 and C4 root systems in determining k. By
quantifying δ13Cconc, soil T, pCO2, δ13Cresp, 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
respiration will enable us to determine the degree of atmospheric contribution to δ13Cpc as well as
to assess the accuracy of using δ13Csom versus δ13Cbulkplant as a proxy for δ13Cresp. Further,
knowledge of seasonal respiration rates will enable us to quantify atmospheric contributions to
δ13Cconc, and seasonal measurement of the δ18O composition of soil CO2 (as a proxy for δ18Osw)
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
composition, one must first quantify the influence of soil respiration on δ13Cpc values. In turn,
this knowledge can be use to assess the influence of evaporation on δ8Opc values for
reconstructing relative, and possibly absolute changes in δ18Omw. Conversely, knowledge of the
δ18O composition of paleo-precipitation can enable the use of δ18Opc values to independently
assess the magnitude of atmospheric contributions to δ13Cpc 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
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37
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
vegetation, and soil gas, and estimated for δ13Cresp. δ13Cpc values are the average of
values from ≥50 cm soil depth; δ13Cconc values are corrected for the Suess Effect; and
δ13Cresp values are calculated from δ13Csom and δ13Cbulkplant 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).
Figure 3. δ13Cconc (VPDB) as a function of soil depth (cm) for soil respiration rates ranging
from 0.1 to 9.0 mmol/m2/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
pressure, δ13Catmosphere = -6.5‰ with pCO2 = 270 ppmV for the pre-industrial atmosphere,
o
free air porosity of 0.5, a tortuosity of 0.6, temperature = 21.7 C, δ13Cresp = -24‰, and
exponential production, where k = 22.5 cm.
Figure 4. Monthly weighted averages of δ18Omw 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).
Figure 5. δ13Cpc (VPDB) versus δ18Opc (VPDB) values from ≥50 cm in late Holocene soils
in southeastern Arizona.
Figure 6. δ18Opc (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.
Figure 7. δ18Opc (VPDB) and δ13Cpc (VPDB) versus depth for a late Holocene soil profile at
Tumamoc, for which δ13Cpc 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
T = 22.1 C, δ = 0.5, and δ13Cresp = -22.5‰. Open symbols represent values from other
profiles at Tumamoc, for which respiration rate was not modeled.
Figure 8. δ18Opc (VPDB) and δ13Cpc (VPDB) versus depth for 2 late Holocene soil profiles
at Madera, for which values are distinguished as solid versus open symbols.
Figure 9. δ18Opc (VPDB) and δ13Cpc (VPDB) versus depth for a late Holocene soil at
Hirabayashi. δ13Cpc is modeled (solid lines) for the respiration rates shown (mmol/m2/hr)
o
using the model described in Quade et al. (2007), where elevation is 1460 m, T = 17.0 C,
38
δ = 0.5, and δ13Cresp = -22.8‰..
Figure 10. δ18Opc (VPDB) and δ13Cpc (VPDB) versus depth for a late Holocene soil profile
at GC Terrace.
Figure 11. δ18Opc (VPDB) and δ13Cpc (VPDB) versus depth for a late Holocene
soil profile at GC S-facing, for which δ13Cpc is modeled (solid lines) for the respiration
rates shown (mmol/m2/hr) using the model described in Quade et al. (2007), where
o
elevation is 1900 m, T = 13.9 C, δ = 0.5, and δ13Cresp = -18.4‰. Open symbols represent
values from other profiles at this site for which respiration rate was not modeled.
Figure 12. δ13Cpc (VPDB) values from ≥50cm soil depth versus site elevation for late
Holocene soils in southeastern Arizona.
Figure 13. δ13Cpc (VPDB) versus δ13Csom (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
respiration rates (labeled) calculated for 25°C. Solid circles represent observed δ13Cpc
values. In general, these values plot above the field defined for high respiration-rate
soils, suggesting a significant atmospheric component.
Figure 14. δ13Cpc (VPDB) from 60 cm depth in a late Holocene soil at COB West, for which
δ13Cpc is modeled (solid lines) for the respiration rates shown (mmol/m2/hr) using the
o
model described in Quade et al. (2007) where elevation is 1650 m, T = 15.6 C, δ = 0.5,
and δ13Cresp = -18.6‰.
Figure 15. Predicted versus observed δ18Opc (VPDB) values, where predicted values are
calculated to form in equilibrium with mean annual δ18Omw.
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
ON
)
(‰
c
p
C pc
C res
C con
OZ
EC
G
)
)
(°C
)
)
(W
m
(m
G
(N
)
Y
(m
IN
TT
SE
P
MA
N
LO
T
LA
N
IT
AL
AT
MA
13
δ
13
13
δ
δ
IO
AT
EV
EL
C
LO
SOM PLANT
o
o
Tumamoc
730
31 43.423’
110 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
COB
West
Terrace
1660
32o 13.298’
111o 0.634’
15.6
539
Hillslope
limestone scrub
Terrace oak-juniper-grassland
-13.1
x
-18.6
-18.5
-21.8
x
-0.9
-0.7
GC
S-facing
Terrace
1900
31o 46.881’
110o 53.176’
13.9
605
Hillslope oak-juiper-grassland
Terrace Ponderosa-grassland
-15.1
x
-18.4
x
x
x
-5.0
-7.4
Figure 1
Cave-of-the-Bells
1660 m
Santa Rita Mountains
Madera
1105 m
Tucson
Basin
Tucson Mountains
Tumamoc
730 m
Hirabayashi
1450 m
Santa Catalina Mountains
Huachuca Mountains
Garden Canyon
1900 m
Temperature (oC)
Figure 2
Precipitation (cm/month)
Mt. Lemmon
(2425 m)
0.0
J
y
ar
nu
Ja
0
5.0
5
10.0
10
15.0
15
20.0
20
25
25.0
30.0
r
ua
br
Fe
y
ch
ar
ril
Ap
M
ay
ne
Ju
ly
Ju
st
gu
Au
pt
Se
em
r
be
O
er
ob
ct
Precipitation
N
em
ov
r
be
D
em
ec
r
be
Temperature
Precipitation
Temperature
Precipitation
Temperature
F M A M J J A S O N D
M
30 Tucson
(740 m)
35.0
35
0.0
0
5.0
5
10.0
10
15.0
15
20.0
20
25.0
25
30.0
30 Santa Ritas
(1310 m)
35.0
35
0.0
5.0
5
5
0
0
10.0
15.0
15
15
10
10
20.0
25.0
25
25
20
20
30.0
35.0
35
35
30
30
0
10
DDepth
e p t h (c(cm)
m)
(cm
20
6.0
6.0
7.0
7.0
30
8.0
8.0
4.0
5.0
5.0
4.0
40
50
60
0.1
90
0.05
0.01
0.01
0.5
0.05
1.0
0.1
2.0
0.5
3.0
1.0
3.0
80
2.0
70
100
-25
-20
-15
δ CCO2 (VPDB)
13
Figure 3
-10
-5
0
Tucson Precipitation wt. δ18Omw (VSMOW)
0
20
10
Figure 4
40
0
0
2
4
6
Month
8
10
12
Amount (mm)
30
Temperature (C)
5
-5
80
-10
0
δ18Opc (VPDB)
-12
-10
-8
-6
-4
-2
0
0
-4
-6
GC Terrace
GC S-facing
COB Terrace
COB West
Hirabayashi
Madera
Tumamoc
-8
-10
-12
Figure 5
δ13Cpc (VPDB)
-2
2000
1800
Elevation (m)
1600
GC Terrace
GC S-facing
winter
1400
COB Terrace
COB West
Hirabayashi
Madera
1200
Tumamoc
1000
summer
800
600
-12
-10
-8
-6
δ 18Opc (VPDB)
Figure 6
-4
-2
0
0A
pedogenic
structure
20-
stratified
gravels
carbon
oxygen
40-
stratified
alluvium
disconformity
60-
2Bw
2Bw
100-
2C
120-
3C1
140-
0.22
3C2
8.0
160-
0.53
Depth (cm)
80-
180200-
4
2
0
-2
-4
-6
-8
-1 0
-1 2
Figure 7
δ ‰ (VPDB)
0-
unbedded
gravels
20-
stratified
alluvium
A
carbon
oxygen
4060-
Depth (cm)
C
801001201401601802004
2
Figure 8
0
-2
-4
-6
-8
-10
-12
δ ‰ (VPDB)