Introduction
The radiocarbon content of soil C provides a measure of how long C can
persist in soils (Trumbore, 2009). A
working hypothesis is that the relative strengths of mineral–carbon
interactions will be reflected in the radiocarbon content of the associated
organic C. For example, 1 : 1 silicate clays with inherently low surface
area, such as kaolinite, have limited sorptive capacity and retain C over
relatively short timescales (Heckman et al., 2009; Sollins et al.,
2009). In contrast, 2 : 1 clays
with high charge density and high surface area, such as smectite, have higher
affinity for C and thus retain it for relatively longer. In soils where the
predominant minerals are smectites, organic C has older radiocarbon ages than
in soils dominated by kaolinite (Wattel-Koekkoek et al., 2003; Poch et al.,
2015). Soils in which much of the C is associated with high surface area
short-range-order (SRO) minerals like Fe and Al oxyhydroxides contain organic
C that has persisted for many millennia (Torn et al., 1997).
In soils of mixed mineralogy, several organic-mineral interaction mechanisms
operate simultaneously, requiring multiple timescales for organic carbon
persistence to explain radiocarbon measurements (e.g., Schrumpf and Kaiser,
2015; Schrumpf et al., 2013; Wattel-Koekkoek and Buurman, 2004).
The ability to quantitatively link specific mineral stabilization mechanisms
with radiocarbon-based timescales of turnover is hampered because the
operationally defined procedures used to quantify soil mineral content mostly
differ from those used to separate organic C into fractions that differ in
radiocarbon content. Two main approaches have been used to address this
issue. One approach is to select samples for analysis from distinctly
different global environments and use samples dominated by single mineral
compositions (e.g., as described above; Wattel-Koekkoek et al., 2003).
Another approach is to sample soils along environmental gradients, and to
correlate C age with the abundance of specific mineral stabilization
mechanisms (e.g., Torn et al., 1997; Masiello et al., 2004; Lawrence et al.,
2015), but often without full quantification of all the possible controls on
C storage. Relatively few studies combine measures of the amounts and age of
C in soil with quantitative measures of mineralogy. In particular, more
studies are needed that focus on the C stabilization behavior of mature soils
where long-term depletion of primary minerals and ripening of secondary
minerals provides an environment dominated by well crystallized compounds
that have relatively low chemical reactivity (see Wattel-Koekoek et al., 2003; Torn et al., 1997). In regions
with long-term tectonic and climatic stability, such as parts of the tropics
and subtropics (Paton et al., 1995), it is possible that the differences in C
sorption between 2 : 1 and 1 : 1 clays could be one of the most important
controls on C storage and turnover.
Here, we analyze a lithosequence of arid to subhumid savanna soils developed
on the Kaapvaal Craton and associated post-Gondwana breakup lavas in Kruger
National Park (KNP) South Africa (SA). Low rates of landscape erosion and
exceptionally long soil residence times (Chadwick et al., 2013) ensure that
nearly all soil minerals have evolved past the metastable SRO stage and that
there are few free trivalent metal ions available for direct sorption by
organic ligands (Khomo et al., 2011, 2012). We evaluate radiocarbon
(14C) in bulk soil, and fractions separated by density into free
particulate and mineral-associated components. In parallel, we used chemical
extractions of bulk soils to quantify Fe oxyhydroxides and SRO minerals, and
quantified allied properties such as cation exchange capacity. Because of our
interest in the role of silicate clay mineralogy, particularly smectite
(2 : 1) and kaolinite (1 : 1), we separated the clay-sized fraction for
X-ray diffraction (XRD) analysis of mineralogy, and measured 14C on the
same fraction.
Our specific research questions reflect the inherent limitations in combining
different methods to quantify minerals and organic matter as follows. (1) How
do the amount and radiocarbon content of bulk, low density
(< 1.7 g cm-3), and dense (> 1.7 g cm-3)
fractions vary among soils developed on different parent materials present in
the Kruger National Park? (2) Can we define relationships between minerals
and the amount and mean turnover time (TT) (derived from 14C) of carbon?
(3) Can such relationships be extrapolated from specific soil samples to
entire soil profiles and across soils with contrasting mineralogy? Our
overall goal is to find relationships that allow us to predict the amount and
TT of carbon across broader landscapes with similar soil forming factors.
Materials and methods
Field sites
To evaluate mineralogical controls on C storage and turnover we sampled soils
across gradients in geology, climate, and topography in KNP. Soil residence
times, estimated from average regolith depth and erosion rates determined
using cosmogenic isotopes are > 105 years (Chadwick et al.,
2013), providing ample time for crystalline mineral differentiation,
ripening, and depletion of metastable SRO minerals. In addition to strong
geological differences across KNP, variation in clay mineralogy is imposed by
a regional north–south gradient in rainfall that ranges from about 470 to
740 mm annually, and locally by differentiation of clay content along
hillslopes. Under this setting, we can focus on organic matter–mineral
interactions associated with differences in silicate clays and secondary Fe
and Al oxyhydroxides in an environmental regime expected to have few SRO
minerals.
We sampled soils underlain by five geological units: rhyolite, granite, an
olivine-rich picrite basalt (black basalt), an olivine-poor basalt (red
basalt), and nephelinite (Venter et al., 2003) (Table 1, Fig. 1). Each of the
lithologies were sampled in the northern arid zone, with mean annual
temperature of 23 ∘C and ∼ 470 mm annual precipitation. We
also sampled soils developed on granite, gabbro, and mixed granite/gabbro
parent materials in the south of the park where rainfall ranges from
∼ 550 to 740 mm per year (Table 1). Samples were collected along
watershed divides, i.e., hill crests in the gently rolling landscape,
although we include data for soils collected along one toposequence at
550 mm of rain to broaden the range of minerals formed on granitic soils
(Khomo et al., 2011, 2012; Bern et al., 2011).
Profile names used in the text, parent material lithology, mean
annual precipitation (MAP), year of collection, locations, and
classification of the soil profiles used for this study.
Profile name
Lithology
MAP
Slope
Latitude
Longitude
Classification
(mm)
position
Year
(easting)
(northing)
GR-450-C
Granite
450
Crest
2004
322 713
7 452 153
Haplocambid
GA-450-C
Gabbro
450
Crest
2010
321 956
7 449 291
Calciustoll
RH-450-C
Rhyolite
450
Crest
2010
351 375
7 421 676
Haplocambid
NE-450-C
Nephelinite
450
Crest
2010
336 567
7 398 988
Ustorthent
BB-450-C
Black basalt
450
Crest
2009
341 888
7 420 588
Haplustert
RB-450-C
Red basalt
450
Crest
2009
344 120
7 421 754
Duritorrand
GR-550-C
Granite
550
Crest
2006
348 678
7 231 971
Ustorthent
GR-550-S
Granite
550
Seepline
2006
348 755
7 231 990
Dystrustept
GR-550-T
Granite
550
Footslope
2006
348 831
7 231 986
Natrusalf
MG-550-C
Mixed granite
550
Crest
2010
341 298
7 232 342
Ustorthent
MG-550-C2
Mixed granite
550
Crest
2010
341 298
7 232 342
Ustorthent
GA-550-C
Gabbro
550
Crest
2005
333 525
7 230 774
–
GA-740-C1
Gabbro
740
Crest
2010
329 124
7 218 015
Haplotorrert
GA-740-C2
Gabbro
740
Crest
2010
329 124
7 218 015
Haplotorrert
GR-740-C
Granite
740
Crest
2004
326 823
7 211 630
Dystrustept
Locations and lithology of parent materials where soils were
sampled for this study. Rainfall decreases from ∼ 740 mm a-1 in
the southern end of the park to ∼ 450 mm a-1 in the northern end
of the park.
Bulk soil characterization
Soil profiles were sampled by horizon to bedrock where possible and described
and classified using standard techniques. Soil depth ranged from 30 cm to
about 2 m. Following air-drying, the samples were sieved to < 2 mm
to remove rocks and roots. Air-dried samples were homogenized and subsampled
for physical, chemical, isotopic, and mineralogical analyses. Bulk density
was measured as the mass of oven-dry soil in a core of known volume. The
amount of clay-sized material (< 2 µm size fraction) was
determined by the hydrometer method (Soil Survey Staff, 2014). The
concentration of exchangeable base cations was determined by atomic
absorption spectroscopy after extraction with 1 M ammonium acetate buffered
at pH 7. Cation exchange capacity (CEC) was determined by extracting the
ammonium saturated samples with a 1 M potassium chloride solution and
determining ammonium by Lachat autoanalyzer. We report CEC corrected for the
contribution of organic matter by assuming a contribution of 200 cmol(+)
per kilogram of organic C (as measured using an elemental analyzer; Soil Survey Staff,
2014).
SRO minerals (aluminosilicate or Fe oxyhydroxides that are minimally
polymerized) were extracted from bulk soils using acid ammonium oxalate in the dark
(Schwertmann, 1973). Iron (Fe(o)) and Al (Al(o)) from the extract were
measured by inductively coupled plasma-optical emission spectrometry. We also
applied a standard dithionite citrate bicarbonate (DCB) extraction and report
the Fe concentration in this solution as Fe(d) (Mehra and Jackson,
1960). Total crystalline Fe oxyhydroxides are defined as Fe(d)–Fe(o). Carbon
and nitrogen (N) content were determined by combustion on a vario MAX cube CN
elemental analyzer. To determine if soils contained pedogenic soil
carbonates, inorganic C was determined on the residue after dry combustion of
bulk samples at 450 ∘C for 16 h (Steinbeiss et al., 2008) and
organic C was calculated as the difference between total C and inorganic C.
Carbonates were present in the red basalt and two of the arid-zone gabbro
profiles; carbonates in upper horizons were mostly present as individual
particles (i.e., not coatings) presumably derived from more massive
carbonates in a Bk horizon below.
Clay-sized material for XRD analysis
To isolate and prepare clay-sized material for XRD measurement of mineralogy,
we started with bulk soil material. Sand-sized material was removed first by
wet-sieving and then the clay-sized (< 2 µm) fraction for
XRD was extracted following dispersion with 5 % sodium hexametaphosphate
and 2 % H2O2 with three rounds of sedimentation and decantation
in a 1 L cylinder (Soil Survey Staff, 2014). The decanted material
containing clay-sized material was evaporated and freeze-dried. The 2 %
H2O2 treatment, a standard pretreatment for isolation of clay-sized
material for XRD measurement, also removed organic C and we waited for bubble
formation (presumably from oxidation of organic matter) to cease before the
first decanting procedure. Organic matter oxidized by this treatment included
free particulate organic C, such as small plant fragments, that would also
float in a solution of 1.7 g cm-3 (i.e., our low-density fraction in
2.4 below). However, this treatment can also remove organic C that is weakly
associated to clay-sized mineral surfaces. Thus, we refer to material
isolated this way as the “clay-sized XRD fraction” and assume that any C
still in that fraction must be strongly associated with mineral surfaces. We
estimate the C and C isotope content of C removed during isolation of the
clay-sized XRD fraction using mass balance.
Splits of the clay-sized material were subjected to standard clay mineral
identification routines including saturation with KCl and MgCl2 before
qualitative and quantitative analysis by XRD. For mineral identification,
peel-mounts of oriented clay-sized material were made by transferring the
sample onto microprobe glass slides from 0.42 µm cellulose nitrate
membrane filters where they had been oriented by vacuum (Pollastro, 1982).
For mineral quantification, the clay-size fraction was micronized in methanol
with 10 % corundum by sample weight, dried, passed through a
50 µm sieve, and placed into side-packed powder mounts (Eberl,
2003). XRD spectra were generated with a Siemens D500 diffractometer using Cu
Kα radiation fitted
with a graphite monochrometer configured to 35 mA and 40 kV. Mineral
quantification was done using the RockJock software (Eberl, 2003) and results
were summed by mineral group. Quantitative mineral data for the clay-sized
XRD fraction from the granitic toposequence at 550 mm rainfall have been
previously published (Khomo et al., 2011). All mineralogy data are normalized
to sum to 100 %.
Density separation
For depth intervals identified as A horizons, where fresh plant inputs are
largest, we performed a density separation on a subsample of soil. We used a
heavy sodium polytungstate liquid (1.7 g cm-3) to separate the sample
into free light fraction (fLF) and heavy fraction (HF) (Schrumpf et al.,
2013, modified to a density of 1.7 g cm-3). A density of
1.7 g cm-3
is sufficient to separate minerals from fresh particulate organic
material especially in soils with minimal SRO minerals that can have low
densities (Castanha et al., 2008). Between 10 and 15 g of soil was added to
100 mL sodium polytungstate solution and gently shaken on a horizontal
shaker for 10 min, ultrasonicated at 60 J mL-1 for 2.5 min, then
centrifuged at 3500 rpm for 30 min. Because of the low energy used, we
consider this fraction to be the fLF, i.e., there may still be some
additional particulate material of low density trapped in aggregates that are
not dispersed. The floating fLF was
concentrated on filter paper (1.6 µm glass microfiber discs) using
a light vacuum. The sinking material was resuspended in the heavy solution
and the steps repeated without the ultrasonic disaggregation until no more
floating fLF was observed (usually three times). The fLF was rinsed with a
liter of water to remove the heavy liquid, then freeze-dried. Visible, very
fine roots were removed by hand (Castanha et al., 2008); coarser roots were
removed by previously sieving. However, the efficiency of these procedures
varies so we used the density separation to ensure all fine roots were
removed from the sample. The remaining root-free fLF was ground to homogenize
it for C isotope measurements. Somewhat lower C contents in the root-free fLF
fraction in this paper compared to other published studies that used
different procedures are likely due to removal of roots, combined with small
amounts of mineral inclusion that were unavoidable as small-sized material
was difficult to separate using the centrifuge (no flocculants were added).
We estimated the overall mass balance of the procedure by combining the mass
and C contents of the different fractions (roots, root-free fLF, and HF) with
the mass and C content of the original bulk sample (Supplement Table S1); in
A horizons we lost 2–15 % of the original C (Table S2), likely through
dissolution in the dense liquid (Castanha et al., 2008)
Mineralogy determined on clay-sized XRD fraction for selected soils
(mostly mafic soils that had higher clay contents). Abbreviations are Hor,
soil horizon, Q: quartz, F: feldspars, Cal: calcite, O: oxides, K: kaolins, S: smectites,
Ch: chlorites, and M: micas. Complete data can be found in Table S1.
Identifier
Hor
Clay
Q
F
Cal
O
K
S
Ch
M
(%)
Percentage of clay-sized fraction
NE-450-C
A
30
1
1
0
15
24
47
0
12
NE-450-C
Bw1
40
1
1
0
14
26
48
0
9
BB-450-C
A1
39
1
0
0
0
6
92
0
0
BB-450-C
Bw1
43
2
0
0
0
5
93
0
0
RB-450-C
A1
36
1
0
0
0
0
99
0
0
RB-450-C
Bk2
46
2
0
0
0
0
98
0
0
GA-450-C
A
15
2
2
0
7
0
67
0
22
GA-450-C
Bw1
25
1
9
6
10
0
43
3
28
GA-740-C1
A
20
0
1
0
14
10
60
6
9
GA-740-C2
Bw1
25
0
1
0
8
16
68
6
1
GA-740-C3
Bw2
10
0
0
0
14
3
77
3
3
GR-550-C*
A
14
0
0
0
0
79
0
21
0
GR-550-C*
Bw2
17
0
0
0
0
79
1
21
0
GR-550-S*
A
6
0
0
0
0
76
17
7
0
GR-550-S*
Bw2
7
0
0
0
0
65
25
11
0
GR-550-T*
A
25
0
0
0
0
57
26
17
0
GR-550-T*
2Btn2
47
0
0
0
0
53
23
15
10
* Data are from Khomo et al. (2011).
Thus, our analyses of C and radiocarbon in fractions contain overlapping
information. For example, we can assume that all of the C and 14C
strongly associated to the clay-sized XRD fraction (ClayXRD) is
also found in the HF. The C removed during isolation of ClayXRD
contains a mixture of C associated with minerals larger than clay-sized
(e.g., a component of HF) as well as root and root-free fLF C (see also
Results, Fig. 4). While it would have been preferable to do a sequential
extraction, the separation of sufficient clay-sized material for mineralogy
and C isotope analyses from the HF fraction would have been costly and
required large amounts of material (particularly for sandy soils with low
clay content).
Similarly, our mineralogical information has some overlapping components. For
example, in the ClayXRD, we report Fe oxyhydroxides as the sum of
minerals such as goethite, magnetite, maghemite, and ilmenite (Table S1). We
do not normally refer to this fraction, but rather to the bulk soil
measurement of Fe oxyhydroxides, which we define as the Fe compounds that are
dissolved by a standard dithionite citrate extraction but not by the standard
oxalate extraction (i.e., Fe(d)–Fe(o)). These compounds are assumed to be
pedogenic Fe (though some amount of geogenic Fe is possibly also dissolved),
and includes the ClayXRD Fe oxyhydroxides as well as coatings on
minerals with sizes > 2 µm but < 2 mm.
Carbon isotopes
Radiocarbon (14C) was determined by accelerator mass spectrometry (AMS).
For determination of 14C in organic samples, an amount of material (bulk
soil, HF, fLF, or clay) needed to yield ∼ 1 mg C was weighed into a
precombusted quartz tube with CuO wire. The tube was evacuated, sealed with a
torch, and placed in a 900 ∘C furnace for 3 h.
The resulting CO2 was purified on a vacuum line, and an aliquot
was removed for determination of 13CO2 using a gas bench coupled to
an isotope ratio mass spectrometer (Xu et al., 2007). The remaining CO2
was reduced to graphite using a sealed tube zinc reduction method (Xu et al.,
2007), and isotopic compositions were measured at the W.M. Keck Carbon Cycle
AMS facility at the University of California, Irvine. Samples containing
inorganic C were acidified with 1N HCl until the solution pH was below 6, and
then dried and analyzed as above. Carbonates were normally present as
distinct sand-sized or larger grains, and were present only in the black
basalt and the dry gabbro samples; we also used the 13C signature of the
combusted sample to indicate that carbonate did not contribute significantly
to the measured sample C. For the black basal soil, we analyzed 14C in
pedogenic carbonates by collecting and purifying the CO2 evolved during
acidification, then reducing it to graphite as for organic C samples.
Radiocarbon data are reported as Δ14C, the deviation from unity, in
parts per thousand, between the ratio of 14C / 12C in the
sample divided by that of preindustrial wood (the standard). The potential
influence of mass-dependent fractionation of isotopes is accounted for by
reporting the 14C / 12C ratio corrected to a common δ13C value (-25 ‰), and assuming that 14C is fractionated
twice as much as 13C by mass-dependent processes (Stuiver and Polach,
1977). Therefore, differences in Δ14C between samples reflect time
or mixing rather than isotope fractionation. In these units,
Δ14C = 0 ‰ is equivalent to the standard. Values
> 0 ‰ indicate the presence of 14C produced by
atmospheric thermonuclear weapons testing in the early 1960s. Values
< 0 ‰ indicate that radiocarbon has had time to
radioactively decay (half-life = 5730 years). Long-term accuracy for
samples measured at the W.M. Keck CCAMS facility is ±3 ‰ for
radiocarbon expressed as Δ14C and ±0.1 ‰ for
δ13C.
We also used the radiocarbon data to estimate the mean TT of soil C in the
profile using a one-pool model that includes incorporation of bomb 14C
in the last few decades and assumes steady state (see Torn et al., 1997;
Trumbore, 2009). Specifically, we
used the SoilR package (Sierra et al., 2014) to calculate the predicted
radiocarbon signature for such a one-pool, steady-state model in the year of
sampling (R code is included in the Supplement). For cases where two turnover
times yielded the same Δ14C in the year of sampling (i.e., in cases
where Δ14C is > 0 ‰), we report both TTs for
the root-free fLF, but only the longer turnover time as it is more consistent
with the fluxes of C into and out of the mineral associated and bulk
fractions (see Gaudinski et al., 2000). The one-pool model is clearly an
oversimplification, but is useful for translating radiocarbon data into
average timescales of stabilization. The use of a mean TT also provides a way
to compare data from samples collected in different years (2004–2011;
Table 1). We want to emphasize that these TTs only have meaning in the
context of the assumptions used to generate them – they refer to C in a
single, homogeneous pool at steady state.
We report C concentration and isotope data for individual horizons as well as
whole-profile averages (e.g., as in Masiello et al., 2004). Mean C isotope
ratios and mean estimated turnover times for whole profiles were calculated
as averages, the carbon mass weighted by horizon and calculated from measured
bulk soil 13C and 14C values.
Statistics
Graphs, including regression analyses, were produced with R (R Core Team,
2015). Correlation matrices were produced using the R package Hmisc (Harrell
Jr. et al., 2016). The TTs were calculated with the SoilR package (Sierra et
al., 2014).
Results
Mineralogy
With few exceptions, these ancient soils contained low amounts (< 0.3 wt. %)
of oxalate extractable Fe and Al presumed to be derived from SRO
minerals (complete data are given in Table S1). Only
nephelinite-derived (0.5–0.9 wt. %) and subhumid gabbro soils had greater
(2.9 wt. %) concentrations of oxalate extractable Al + Fe. Crystalline Fe
oxyhydroxides determined as Fe(d)–Fe(o) ranged from 0 % in a
periodically anoxic “seep” zone in the granitic toposequence to 5.2 wt. %
in the nephelinite soil.
The clay-sized XRD fraction made up ≤ 15 % of the < 2 mm mass
for soils developed at crest positions on granites and rhyolites, but up to
35–50 % in the red and black basalt, arid-zone gabbro and nephelinite
soils, and at the toeslope of the granitic toposequence (Tables 2, S1).
With a few exceptions, the amount of clay-sized material increased
with soil depth.
Within the clay-size XRD fraction, the sum of smectite, kaolinite, micas and
chlorite and crystalline Fe minerals generally made up over 90 % of the
quantified mineralogy (Table 2). Smectite was present in all of the isolated
clays except the granite crest under relatively high (740 mm) rainfall and
dominated the clay fraction in red (> 90 %) and black basalts
(> 99 %; Table 2). Kaolinite was common in most soils but
rare in the arid-zone gabbro and the two basalts. Crystalline Fe oxide
minerals identified by X-ray diffraction made up 3–26 % of the
clay-sized fraction for most soils but < 1 % in the
smectite-dominated red and black basalts (Table 2).
Comparison of Fe oxyhydroxide abundance estimated by scaling quantitative XRD
in the clay fraction to the whole soil (i.e., multiplying the weight percent
clay times the Fe oxide content measured by XRD) with that measured by
extraction with DCB and oxalate in the bulk soil showed overall
correspondence (see Fig. S1) but with much scatter.
Comparison of 13C (a) and 14C (b) in
root-free free light fraction (root-free fLF) and heavy fraction (HF) organic
C for individual samples from A horizons (see Table 3). Felsic lithologies
(white circles) include soils developed on granite and rhyolite; mafic
lithologies (black circles) include soils developed on gabbros, basalts, and
nephelinite. The 1 : 1 correlation line is plotted for reference.
C in density fractions of A and B1 horizons
The root-free fLF isolated from A horizons had C concentrations of
1.0–3.7 gC kg-1 (Table 3), with lower concentrations in the soils
derived from rhyolite and the two basalts (1.0–1.6 % gC kg-1). As
mentioned above, removal of roots and the inclusion of some clay-sized
material during filtration can result in lower values than expected if the
root-free fLF is only fresh plant material. In general, carbon in the
root-free fLF made up only 10–20 % of bulk C, even in surface soils
(Table S1). Mineral-associated HF had lower C concentrations but comprised
more of the bulk soil mass and represented 40–70 % of bulk soil C for
granites, nephelinite and dry gabbro soils, and > 80 % in
other soils (Table 3).
Carbon and carbon isotope signatures of heavy (> 1.7 g cc-1)
and root-free free light fraction carbon. Turnover
times (TTs) were estimated using the steady state, one-pool model described
in the text. When two TTs were possible, we show both options for the
root-free fLF but only the longer one for the heavy fraction.
Identifier
Root-free free light fraction
Heavy fraction
Depth
Total C
Corg
δ13C ‰
Δ14C ‰
TT
Corg
δ13C ‰
Δ14C ‰
TT
(cm)
in HF
(gC kg-1)
(year)
(gC kg-1)
(year)
GR-450-C
0–23
0.59
3.70
-23.7
99.3
5, 75
0.05
-19.5
30.1
195
23–45
1.06*
2.46
-18.9
125.6
8, 50
0.04
-17.5
-35.9
510
GA-450-C1
0–2
na
3.42
-16.8
44.1
155
na
-13.6
6.2
275
2–12
na
4.23
-15.0
34.4
180
na
-14.5
20.9
225
RH-450-C
0–3
0.81
1.66
-23.2
98.4
8, 75
0.07
-18.2
72.9
55
3–15
0.83
0.19
-20.0
76.7
5, 100
0.05
-16.2
40.9
165
NE-450-C
0–2
0.65
3.13
-20.6
88.3
8, 85
0.42
-16.7
74.4
105
2-18
0.76
3.08
-19.5
64.4
4, 120
0.24
-14.9
1.8
300
RB-450-C
0–4
0.82
1.16
-16.6
60.0
3, 125
0.18
-14.8
-23.1
425
4–15
0.69
1.00
-16.6
30.0
195
0.14
-13.3
-95.2
985
15–30
0.92
1.00
na
na
na
0.14
-13.3
-152
1560
30–49
0.84
1.12
na
-53.0
330
0.13
-13.3
-216
2300
BB-450-C
0–3
n.d
1.00
-15.4
33.5
185
0.15
-13.7
-25.4
440
3–11
0.78
0.05
-16.6
33.5
185
0.16
-12.9
-65.6
735
GR-550-C
0–15
0.70
1.60
-22.2
57.7
1, 130
0.05
-18.1
50.3
145
15–41
na
1.45
-21.9
82.3
2, 95
0.04
-17.1
59.0
130
GR-550-S
0–2
0.72
2.25
-20.8
62.4
1, 120
0.05
-18.8
62.9
120
2–10
1.1*
2.13
-20.2
82.1
4, 95
0.04
-19.1
96.1
80
GR-550-T
0–8
0.91
2.12
-18.0
54.7
1, 135
0.08
-16.7
58.0
130
8–15
0.69
1.52
-19.9
74.4
105
0.05
-17.4
79.0
100
MG-550- C
0–3
0.84
3.69
-22.3
71.7
5, 125
0.09
-16.6
60.2
125
3–10
0.92
1.55
-20.6
80.5
6, 110
0.07
-15.3
54.1
135
GA-550-C
0–9
na
2.99
-28.2
98.3
5, 85
0.28
-14.4
45.8
140
9–24
na
1.52
-13.2
88
4, 100
0.23
na
38.4
175
GA-740-C1
0–3
0.70
3.47
-16.9
72.6
1, 100
0.16
-13.4
78.2
100
3–9
0.73
2.32
-14.8
40.5
165
0.15
-11.8
6.1
275
GA-740-C2
0–4
0.85
3.57
-16.1
67.6
4, 135
0.15
-13.1
85.2
90
4–24
0.76
3.53
-16.9
64.4
4, 140
0.18
-12.8
60.6
125
GR-740-C
0–8
0.41
3.27
-21.3
145.8
10, 45
0.10
-16.8
143.9
45
8–17
0.66
1.86
-17.8
95.0
4, 80
0.03
-15.1
109.4
70
* Values > 1 indicate the magnitude of errors associated with
density separations; na means not analyzed.
Root-free fLF δ13C ranged from -24 to -14.5 ‰
(Table 3), reflecting a mixture of C3 and C4 vegetation sources. We found no
relationship between root-free fLF δ13C or HF δ13C with
rainfall, but mafic soils were consistently more enriched in δ13C
in both fractions compared to felsic soils (Fig. 2a). Radiocarbon signatures
of root-free fLF (that includes char as well as plant fragments) varied from
values close to those measured in annual grasses in 2010 (+35 ‰ in
Δ14C) up to +145 ‰ (Fig. 2b). For surface horizons, the
one-pool model yielded two possible turnover times for most of the root-free
fLF Δ14C. Assuming the shorter of the two for soil A horizons
(normally 0–2 cm), yielded TTs from < 1 to 8 years, while assuming
the longer TT yielded 45–185 years (Table 3). For the black basalt and
arid-zone gabbro soil A horizons, only longer root-free fLF TTs
(125–185 years) were consistent with observed Δ14C signatures. The
TT of both root-free fLF and HF increased with depth. Fine roots picked from
root-free fLF in the red basalt soil had radiocarbon signatures equivalent to
TT < 1 year regardless of depth (Table S1) and δ13C
signatures of -12 to -16 ‰.
Measurements of C and isotopes in the bulk soil and clay-sized XRD
fraction, the same fraction analyzed for quantitative mineralogy (Table 2).
The fraction of bulk C in the clay-sized XRD fraction (FclayXRD) and
the characteristics of the C making up the rest of the bulk C are calculated
using Eqs. (1)–(3) given in the text. Horizon bottom depths (Bot. depth) > 10 cm
indicate B1 horizons; we excluded data from sampled depths > 20 cm.
We cannot rule out the potential for carbonates making up a small
fraction of the ClayXRD fraction for the dry gabbro (GA-450-C1) Bw
horizon sample as these were not acidified prior to combustion and
carbonates were found using XRD in this soil. All other ClayXRD
samples did not contain measurable carbonates (Table 2).
Bot.
Clay-sized (< 2–µm) XRD fraction
Bulk soil
Remaining C (“non-clay”)
depth
Corg
δ13C
Δ14C
TT
Corg
δ13C
Δ14C
TT
FclayXRD
Corg
δ13C
Δ14C
TT
Identifier
(cm)
(gC kg-1)
‰
‰
(year)
(gC kg-1)
‰
‰
(year)
( %)
‰
‰
(year)
NE-450-C
2
0.252
-16.9
-1.5
310
0.604
-17.8
65.0
110
0.13
6.5
-17.9
69
110
NE-450-C
18
0.114
-16.8
-129.5
1330
0.304
-15.4
8.0
270
0.15
3.4
-15.3
16
240
GA-450-C1
2
0.469
-14.8
-64.6
730
0.328
-14.9
20.1
225
0.22
2.9
-14.9
58
130
GA-450-C1
12
0.235
-13.9
-145.0
1485
0.190
-13.9
-28.8
465
0.31
1.7
-13.9
43
160
RB-450-C
4
0.250
-14.1
-125.3
1275
0.194
-14.9
-16.0
385
0.47
1.5
-16.0
149
30
BB-450-C
3
0.248
-14.3
-91.7
955
0.241
-13.7
-25.4
440
0.40
2.4
-13.4
22
220
GA-450-C2
2
0.129
-16.3
-4.4
320
0.162
-16.6
22.0
220
0.12
1.7
-16.6
25
210
GA-740-C1
4
0.096
-17.6
-91.7
955
0.218
-11.8
-62.1
690
0.09
2.3
-11.6
-61
690
GA-740-C2
9
0.117
-19.1
-160.0
1650
0.176
-13.6
88.1
85
0.14
1.9
-12.9
122
55
GA-740-C2
24
0.110
-17.7
-116.1
1180
0.224
-13.4
70.0
110
0.10
2.4
-13.0
83
90
The δ13C of HF averaged ∼ 3–4 ‰ more enriched than δ13C of root-free
fLF from the same soil (Fig. 2a). Radiocarbon signatures in mafic soil HF
were generally much more depleted in 14C than root-free fLF from the
same horizon (Fig. 2b). Felsic soils tended to have higher 14C values
in HF than mafic soils, though this was less true for root-free fLF
fractions.
Changes in C, 13C and 14C with depth
Soil depths increased with increasing annual rainfall. In all soils, C and
14C concentration decreased with depth. However, differences in
lithology and hence mineralogy were more important controls on C and C
isotopes than differences in rainfall (Fig. 3 and Table S2).
Notably, soils developed on nephelinite had the highest C concentrations
while felsic soils had the lowest (Fig. 3). Patterns in δ13C
by depth followed two general patterns. Felsic, gabbro, and nephelinite soils
had large (2–6 ‰) increases in between the surface and
∼ 10–30 cm depth then became depleted below (Fig. 3). Red
and black basalt soils experienced a ∼ 2 ‰
enrichment in δ13C to ∼ 10 cm depth and then
stayed constant. Radiocarbon declined with depth in all soils but in wetter
sites (> 550 mm annual rainfall) shifted towards higher 14C
values at the very bottom of the profile (BC or C horizons; Table S1).
The 14C signatures of organic C in the red and black basalt soils were
lower (< 0 ‰ at all depths, even at the surface)
compared to the other soils and were the most enriched in δ13C
at all depths (Fig. 3).
Depth profiles of bulk C (top-left), δ13C
(top-middle), and Δ14C (top-right) measured in bulk C for soil
profiles developed on selected lithologies. Also shown in the lower panels are
other soil bulk properties, total Fe oxyhydroxides (Fe(d)–Fe(o))
(bottom-left), total nitrogen (bottom middle), and cation exchange capacity
(CEC) corrected for organic matter contributions (bottom right; see text).
The B horizons of the red basalt and the two arid-zone gabbro soils
contained pedogenic carbonates at concentrations of up to several percent
with radiocarbon ages ranging from ∼ 4500–25 000 14C years,
substantially lower in 14C compared to organic C at the same
depths (see Table S1). The carbonates in B1 horizons were generally
found as distinct small but visible fragments (with 14C ages up to
∼ 4500 years) and likely derived from fragmentation and upward
mixing of older and more massive carbonates observed deeper in the soil.
Carbonates do not influence radiocarbon or signatures reported for organic
matter in these soils as they were removed prior to combustion; the
efficiency of removal can be observed in the similarity between isotope
signatures of organic C in red basalts (containing carbonate) and black
basalts (that did not have carbonate).
Carbon inventory and C strongly associated with clay-sized XRD
fraction
The concentration of C strongly associated with the ClayXRD fraction
ranged between 0.10–0.47 gC kg-1 across all soils. Both the amount of
C and its TT (ranging from 310 to 1330 years) increased with the amount of
smectite measured in the same fraction (Table 4; Fig. 5a). For
ClayXRD samples where smectite made up > 95 % of the total
mineral content, the mean TTs were in the range of 970–1250 years. Carbon
strongly associated with ClayXRD samples where only ∼ 45 %
of the mineral was smectites had the lowest TTs (340 years). We were
unable to measure radiocarbon in samples from granites due to the very low
yield of clay-sized material in these soils.
Assuming bulk C has two components, a portion strongly associated with clay
(ClayXRD) and the rest of the soil C (“non-clay”), i.e., C that
was removed with > 2 µm material (including coatings
and organic matter associated with larger grains) or by the 2 %
H2O2 treatment (including much of the LF but also organic C weakly
associated with clay-sized material). We estimated the amount and radiocarbon
signature of this “non-clay” component using mass balance as follows:
%Cnon-clay=(%Cbulk-FClayXRD×%CClayXRD)/(1-FClayXRD)Δ14Cnon-clay=Δ14Cbulk-FClayXRD×Δ14CClayXRD)/(1-FClayXRD),
where FClayXRD is the fraction of bulk C that is found in the
clay-sized XRD fraction:
FclayXRD=(%CclayXRD×%Clay)/(%Cbulk×100%).
For the basalts, C in the ClayXRD fraction made up 40–47 % of
Cbulk in the top 2 cm, increasing to 80–86 % in B horizons
(see Table S1). For all other soils, the amount of C strongly associated with
ClayXRD accounted for < 30 % of bulk C. Other than the
basalt-derived soils (and deeper B horizons in the gabbros; Table S1), most
bulk C was thus removed by the fractionation processes (size and
H2O2 treatment). As estimated from mass balance, the C removed
(Δ14Cnon-clay-sized) in the top 18 cm was (with one
exception) dominated by C fixed in the last 50 years (Table 4). The estimated
TT for non-clayXRD C ranged from 30 to 690 years, averaging
190 ± 190 years; Table 4). The C strongly associated with
ClayXRD (a fraction that includes not only clay minerals but up to
26 % Fe oxides) for the same samples averaged 1020 ± 460 years.
As noted previously, the fractionation methods applied in this study, based
on density and particle size, overlapped in what they measured (Fig. 4). For
example, all of the C in ClayXRD is a subset of HF. There is also
overlap between the root-free fLF and the C removed when isolating ClayXRD. The
distribution of isotopes and C among the various fractions for one soil
(illustrated as an example in Fig. 4) demonstrate these variations and
relationships and show that the biggest differences in radiocarbon are
between ClayXRD and root-free fLF fractions.
Comparison of C and C isotopes for a single soil sample (0–3 cm
depth in GA-740-C (gabbro parent material) indicating the interrelationships
among the different process-defined organic C fractions. For each fraction we
indicate the percent of total C (CT) it contains; fractions for
root-free fLF and HF do not add to 100 % because of contributions from
roots picked from the free light fraction (fLF) and C that dissolves in the
polytungstate solution and is not recovered. Colors indicate overlaps between
C among different fractions that exist. For example, C strongly associated with
the clay-sized fraction measured with XRD (ClayXRD; orange) makes
up part of the HF C. Free low-density C (fLF; green; including both
roots picked from the fLF and the root-free fLF fractions) makes up part of
the non-clay-sized fraction. Most of the fLF is likely removed from the
ClayXRD when it is treated with 2 % H2O2.
Mineral–carbon relationships at the profile scale
Profile-averaged properties were calculated to highlight variation across the
landscape (Table 5). This calculation introduced errors associated with
highly uncertain estimates of gravel and bulk density (see values in
Table S1), but as these errors are identical for the elements being compared
(e.g., profile-averaged concentrations), the profiles being compared should
share systematic biases (i.e., similar operator error). The calculation of
total profile C inventories (Table 5) are included to demonstrate the
importance of these factors in understanding profile-scale C storage and
dynamics. For example, the nephelinite soil had the highest C concentrations
(averaging 3.8 %C for the whole profile) but was estimated to have
80–90 % gravel (Table S1), so the estimated C inventory
(1.1 kg C m-2) is not the highest when compared to other soils (which
ranged from a low of 0.6 kg C m-2 in the dry granite soils to
11.4 kg C m-2 in the black basalt soils). Thus, it should be
remembered that the relationships derived here are for the < 2 mm
component of soil.
Profile-averaged (excluding BC / C horizons) properties for the
soils sampled in this study. Averages for C and C isotopes are calculated
from bulk values. C inventory (Cinv) is the sum for the profile in kgC m-2.
All averages are mass-weighted, except C isotopes which are
weighted by the mass of C in each horizon. Smectite content (Smec.) is estimated from
multiplying the fraction of total mass in the clay-sized XRD by the
%Clay (denoted in the table as Clay and expressed as percent of total
mass) by the percent of the clay-sized XRD identified as Smectite (Smec.;
Table 2). Profile-averaged values for ditionite citrate extracted iron
(Fe(d)) and oxalate extractable iron Fe(o) and aluminum (Al(o)), as well as
their difference, a measure of Fe oxyhydroxides in bulk soil, are all
expressed as weight % as in mass Fe per 100 g of soil.
Identifier
Depth
Cinv
pH
CEC*
Clay
Smec.
Fe(d)
Fe(o)
Fe(d)–Fe(o)
Alo
Corg
C / N
δ13C
Δ14C
TT
(cm)
(wt. %)
(gC kg-1)
(‰)
(‰)
(year)
RH-450-C
30
1.3
6.8
5.5
1.0
10
2.7
0.1
2.6
0.1
0.066
9.9
-16.2
28.0
230
GR-450-C
23
0.6
6.1
7.7
6.3
46
0.6
0.0
0.6
0.0
0.078
11.9
-20.2
30.4
200
NE-450-C
18
1.1
6.8
61.8
38.7
48
5.4
0.3
4.9
0.2
0.385
11.8
-15.9
9.9
235
BB-450-C
49
11.4
7.7
44.3
42.0
98
1.7
0.2
1.5
0.2
0.156
14.2
-13.5
-140.4
1500
GA-450-C
34
8.6
8.3
25.7
9.9
50
1.9
0.2
1.7
0.2
0.166
na
-15.0
-37.0
550
RB-450-C
70
8.6
7.0
50.1
46.2
93
2.5
0.1
2.4
0.1
0.153
14.8
-12.3
-156.4
1720
GR-550-C
62
3.5
5.4
3.2
14.8
1
0.4
0.1
0.3
na
0.032
21.2
-15.3
12.2
430
GR-550-S
41
1.8
5.1
2.6
7.5
21
0.1
0.0
0.1
na
0.023
19.6
-18.6
51.8
150
GR-550-T
46
3.6
7.0
29.9
42.7
24
0.2
0.1
0.1
na
0.050
14.1
-14.7
24.8
225
MG-550-C
38
2.6
6.9
7.8
15.0
41
1.4
0.3
1.1
0.2
0.082
14.2
-13.8
-68.6
755
GA-740-C1
44
7.0
7.2
37.7
17.7
69
2.6
1.1
1.5
0.4
0.153
12.2
-13.9
-47.5
250
GA-740-C2
25
4.5
7.3
31.5
25.8
25
2.5
2.0
0.5
0.3
0.149
11.9
-13.3
26.9
150
GR-740-C
93
5.1
5.7
7.0
3.7
10
0.4
0.0
0.4
0.0
0.036
14.8
-18.5
88.8
240
Depth indicates the depth to which the in-profile averages were
calculated (we excluded BC and C horizons). Values in bold for smectite content (Smec.) were not measured but are
assumed based on similar lithology values. We assumed average values for the
horizons above and/or below to fill in data for smectite content for depths
in a profile where no measurements were available (see Supplement).
Mass-weighted mean profile %Corganic correlated significantly
with mineral CEC (i.e., CEC corrected for organic matter contribution), and
bulk Fe oxyhydroxides determined from Fe(d) to Fe(o) (Fig. 6 and Table S2).
Together, bulk Fe oxyhydroxides and nonorganic CEC explained most of the
variation in carbon inventory across all soils. We found a significant
relationship between the amount of smectite (determined as the % of mass
in the clay-sized XRD fraction times the fraction of that mass that was
quantified by XRD as smectite minerals) and the mean TT (Fig. 6); total
clay-sized XRD fraction and the fraction of smectite each also correlated
individually with the horizon averaged bulk 14C, but not as well as
their product (see complete correlation matrix in the Supplement,
Table S2). The only highly significant
correlation for bulk profile 13C was with 14C (Fig. 6), though less
significant correlations were found between 13C and average clay
content, pH, and CEC (Table S2).
(left) Mean turnover time (TT) of C in the clay-sized XRD fraction
increases with smectite concentration; the linear relationship for A horizon
points (n= 7) is R-squared = 0.48 and p= 0.08. (right) Mean TT of
bulk C averaged for each profile compared to the fraction of the total C that
is found in the clay-sized fraction (FclayXRD). The mean TT of bulk
organic C correlates significantly with the fraction of organic matter
strongly associated with the clay-sized XRD fraction. Here, the linear
relationship R-squared = 0.54 and p = 0.004.
Across all the studied soils, the mean organic C content (gC kg-1)
in the soil profile was best predicted by Fe(d)–Fe(o) (left;
R-squared = 0.60, p = 0.003), and cation exchange capacity corrected for
organic matter content (C-corr CEC) (right; R-squared = 0.63, p = 0.0003).
The best predictors for profile-averaged C turnover times was the amount of
smectite clay (Fig. 5 (right)). Correlation matrices for other variables
in Table 5 are given as Table S2.
Discussion
Geological, climatic, and topographic variation in Kruger National Park give
rise to soils of varying mineral compositions. Different strengths of
association of organic C with these minerals lead to observed patterns in C
inventory and TTs across the sampled landscape. None of the soil properties
we measured showed a significant relationship with mean annual precipitation
(Table S2), indicating that any influence of climate on C amount and TT was
indirect, through mineralogy and possibly vegetation. This was true even for
the fractions, like root-free fLF, that are expected to be controlled by vegetation
and climate. Underlying lithology significantly influenced the amount of
clay-sized material, the amount of smectite in this material, CEC, and the TT
of bulk C (Table S2).
The variation in mineral composition and amount in Kruger Park allowed us to
identify simple, scalable relationships between measures of soil mineralogy
with C amount and TT. Overall, we find that no single mechanism can explain
both C inventory and TT, partly because our operationally defined fractions
failed in most cases (all but the ClayXRD in basalts) to isolate pure
mineral end members.
We expected that in soils with low concentrations of SRO minerals, the ratio
of smectite to kaoline would exert the strongest influence on C inventory and
TT. We indeed observed that the C strongly associated with smectite minerals
that made up > 90 % of the clay-sized minerals in basalts had
TTs averaging ∼ 1000 years even in the top 2 cm, and overall the TT of
C in the ClayXRD fraction correlated (weakly) with the amount of
smectite (Fig. 5a).
At the scale of the whole profile, the amount of smectite correlated
significantly with mean TT of C (Fig. 5b). Hence, we conclude that the C
strongly associated with smectite clay surfaces is responsible for the long
TT of C in the clay-sized XRD fraction, and that the total amount of smectite
clay in a soil profile exerts control on the overall TTs estimated from
14C of bulk organic C. Soils at the toeslope of the granitic catena with
HF C radiocarbon signatures +58 ‰ (i.e., TT of 130 years) still
had 39–49 % of their mass in the ClayXRD fraction. In that
fraction, ∼ 23–26 % of the mass was identified as smectites and
53–57 % as kaolin clay minerals; Table 2). This example demonstrates
that it is not merely the amount of clay-sized material in the soil, but the
amount of it that is smectite (i.e., 2 : 1 clay) that is key to long-term C
storage in Kruger soils.
These results are in accord with findings by Wattel-Koekkoek and Buurman (2004)
that C stabilized on smectite in surface horizons has turnover times
of 600–1400 years in soils from Africa and South America. Wattel-Koekkoek et
al. (2003) also showed that the older C associated with smectite tends to be
more aromatic, suggesting that 2 : 1 clays provide a long-term store for
fire-derived C. The aging of LF C with depth in fire-prone soils was shown to
be related to the presence of char in soils from other fire-adapted
ecosystems (Koarashi et al., 2012; Heckman et al.,
2009); where we analyzed this in the
red basalt soils, we found increased TT of LF C with depth as well (Table S1).
Though we did not measure the chemistry of root-free fLF C, the presence of
charred materials provides one possible reason for its low TT, particularly
in the red and black basalts where grass biomass is high and fires frequent
(Govender et al., 2006).
As is clear from our results, phyllosilicates provide just one mechanism for
C storage in ancient soil. Given the very strong relationship between our
bulk measure of crystalline Fe oxyhydroxides (i.e., Fe(d)–Fe(o)) and C
concentration across our soils (Fig. 6), it is reasonable to propose that Fe
oxyhydroxides also provide important mechanisms for storing organic C,
especially in soils with low smectite content. We cannot directly measure the
TT of C removed by bulk extractions, as both DCB and oxalate contain
dissolved organic C extraneous to the soil. However, we can infer from the
relatively short TT of C in soils with fewer smectite minerals that the TT
associated with the more abundant mineral phases (kaolinite and Fe
oxyhydroxides) is hundreds of years or shorter in A horizons. We thus expect
millennial C associated with smectite to remain relatively insensitive to
future changes in climate and land use, while the decadal–centennial cycling
C associated with the fLF, Fe oxyhydroxides and non-smectite clays like
kaolinite should respond faster.
At the pedon scale, clay content was not the best predictor of the amount
(r2= 0.50, p= 0.02) or TT (r2= 0.59, p= 0.03) of soil C,
though this relationship improved when only smectite clay was considered
(Fig. 6; r2= 0.75, p= 0.001). Given the long TT associated with C
stabilized by smectite, we conclude that even a small addition of
millennially aged C strongly associated with smectite contributes
substantially to the mean TT estimated from the bulk soil C. For example,
mixing 75 % C with a TT of 25 years with 25 % C with a TT of
1200 years yields a bulk TT of ∼ 320 years. Increasing the millennial
pool to 35 % changes the mean TT of bulk C to ∼ 450 years. The same
is not true for C stocks, however, which are not significantly correlated
with either clay (r2= 0.50, p= 0.08) or smectite clay (r2= 0.45, p= 0.12).
Somewhat unexpectedly, the subhumid-zone gabbro and arid-zone nephelinite
soils with the highest concentration of SRO minerals as determined from the
Fe(o) and Al(o) extract concentrations had younger C than would be predicted
based on expected relationships between SRO minerals and C age found in other
soils (see Torn et al., 1997; Kramer et al., 2012). SRO minerals are
particularly strong sorbers of C because their hydrated nanocrystals create
intimate mixtures of mineral and organic material that – in the absence of
drying and rewetting or redox pulses – tend to remain very stable (Chorover
et al., 2004; Thompson et al., 2006b; Buettner et al., 2014). However, when
SRO minerals are subjected to drying and rewetting or oxidation–reduction
pulses, they reorganize into larger, more well-ordered crystalline compounds
by ejecting C and water from the interior of their lattice structure (Ziegler
et al., 2003; Thompson et al., 2006a). The subhumid-zone gabbro and
nephelinite soils had younger C, and only 9–17 % of the C was associated
with the clay-sized XRD fraction, even though they have > 5 %
SRO mineral concentrations. The young C suggests that the SRO mineral phase
is likely a relatively transitory phase that forms as primary minerals in the
gravel and cobble fraction of the soil weather and the resulting SRO minerals
rapidly ripen to kaolinite, with any C that was sorbed into the SRO mineral
phase made available for microbial decomposition. In the same way, redox
oscillations under seasonal wet–dry cycles promote crystallinity of Fe and we
suggest that the Fe-bearing SRO minerals in these environments are likely
short-lived giving way to crystalline Fe forms where C is sorbed to surfaces
rather than within the less accessible lattice (Ziegler et al., 2003;
Chorover et al., 2004). Thus, although the availability of a large surface
area may promote stabilization of large amounts of C in these soils (e.g.,
nephelinite in Fig. 3), the relatively rapid TT of that C may be a reflection
of the short residence time of the minerals themselves and the short TT of C
sorbed onto 1 : 1 clays and crystalline Fe oxyhydroxides. Studies of
mineral–carbon interactions must consider not only the strength of C
association with various mineral phases (e.g., strong for SRO and smectite,
weak for kaolinite and oxyhydroxides), but also the timescale of mineral
stability in the soil profile and its pedogenic setting. Where SRO minerals
and oxyhydroxides are stable, the associated C tends to be old, but in
climates such as in Kruger the combination of a relatively short but strong
rainy season and a long intervening dry season can lead to relatively rapid
mineral transformation and hence rapid C turnover.
Factors that vary with soil depth exert controls on both C inventory and TT
in KNP soils, as has been reported in many other areas. These affect the
14C in all measured fractions. However, the rates at which age increased
with depth differed between soils and C fractions. For soils with the largest
amount of smectite clays (e.g., basalts), offsets in the TT for clay-sized
XRD, and the “non-clay-size” fractions were largest at the surface and
smallest at depth. In contrast, in soils with little smectite clay, the
offset between fractions was relatively uniform with depth. More work is
required to understand the stability of the different mineral phases
themselves, especially organic C associated with Fe (and Al) oxyhydroxides
phases, and how they interact with transport mechanisms in soil (e.g.,
Schrumpf et al., 2013; Schrumpf and Kaiser, 2015).
There appears to be a mineral/lithologic control on 13C variation in
KNP soils. This control can operate in at least two ways. First, the
root-free fLF δ13C seem to indicate greater C3-derived
vegetation inputs to soils with more felsic parent materials, and a
predominance of C4 inputs in more mafic soils (Fig. 2a). This is
consistent with the vegetation patterns on the ground (Scholes et al., 2003),
with C4 grasses dominating the basalt-soil landscapes. These patterns are
largely preserved in the mineral-associated C (Fig. 2a), although HF δ13C is consistently enriched compared to the LF δ13C. At
the profile scale, the strongest predictor of δ13C is Δ14C (or TT; Fig. 6; r2= 0.75, p= 0.005), followed by pH
(r2= 0.65, p= 0.016) and clay content (r2= 0.65, p= 0.017).
For the isolated clay fraction, the relationship between δ13C
and smectite was weak. Soils with the greatest amount of smectite are also
those with the greatest C4 vegetation, so it is unclear whether lithologic
control on C3 vs. C4 plants or fractionation associated with different
mineral stabilization mechanisms is responsible for the overall trends
observed in δ13C. Nonetheless, interpretations of
paleo-vegetation from bulk soils must be undertaken with care, as variations
in the mechanism of C stabilization across the landscape may affect the
δ13C signature as well as vegetation changes. More work is
needed to disentangle these relationships at broader spatial scales
encompassing climate and topographic gradients that will also involve
changes in mineralogy.
Large parts of the land surface contain old soils with low concentrations of
SRO minerals (Paton et al., 1995). We found good agreement in the age of C in
the most smectite-rich (basalt soils) clay-sized XRD fraction and those
reported by Wattel-Koekkoek et al. (2003) from soils collected at other sites
in Africa and South America. Smectite clay content may thus provide a useful
indicator for the fraction of C stabilized on millennial timescales over
large areas. While quantitative clay mineralogy is not an easy measurement,
the amount of smectite in our soils was broadly predictable from lithology
and from more easily measured soil properties such as CEC (corrected for
organic contributions; r2= 0.79, p < 0.01) or pH (when below
CaCO3 saturation; r2= 0.83, p < 0.005; Table S2).
Thus, across a range of landscapes and parent materials, one could predict
how much of the C in soils is cycling on slower timescales based on these
parameters, while overall C inventory is more related to crystalline Fe and
Al oxyhydroxides.