Introduction
Soil organic carbon (SOC) is one of the most important terrestrial C pools.
Carbon in soils can be protected against decomposition by three key
mechanisms: (i) inherent biochemical recalcitrance; (ii) organo-mineral
associations, by interaction of organic molecules with mineral surfaces; and
(iii) physical protection, making the SOC inaccessible to degraders/consumers
within soil aggregates (Sollins et al., 1996; Six et al., 2002). These
mechanisms are interactive (e.g., aggregation of organo-mineral associations
could already include biochemically recalcitrant SOC) and their contribution
to SOC stabilization is strongly influenced by soil environmental conditions
and landform (Salomé et al., 2010; Berhe et al., 2012; Dungait et
al., 2012; X. Wang et al., 2014). Distinguishing the role of a single mechanism
for stabilizing C is, hence, a difficult task and a matter of ongoing debate
and research (Berhe et al., 2012). For example, some authors argue that the
biochemical recalcitrance of organic molecules does not exist per se
and has always to be seen in an environmental context (Kleber, 2010),
while others argue that “recalcitrant SOM [soil organic matter] can be defined by intrinsic
molecular properties, but these properties may be fairly irrelevant under
specific environmental conditions” (von Luetzow and Koegel-Knabner, 2010).
Studying SOC dynamics in landscapes with lateral soil fluxes requires the
combined study of geomorphological, climatic, biogeochemical, and microbial
parameters (Park et al., 2014). In most studies on SOC dynamics, the
landscapes in which C exchange takes place are stable surfaces with no or
little lateral fluxes of soil or C. Hence, SOC fluxes in complex pool
models are limited to a vertical exchange between soil and atmosphere.
Connecting the different controls on SOC dynamics across soil depths and
topographic positions has only recently received attention in landscape
scale studies on SOC dynamics (Berhe et al., 2012; Doetterl et al., 2012).
Bringing the detailed knowledge that we possess about SOC dynamics at plot scales
and micro-scales into perspective at the landscape scale remains largely
neglected. Recently, soil redistribution in cropland and grassland systems
has been shown to lead to distinct qualitative and quantitative SOC
modifications along geomorphic gradients compared to soils in stable
landforms (Yoo et al., 2006; Dlugoss et al., 2011; Berhe et al., 2012).
Soils at eroding sites are usually C depleted while soils in depositional
settings can store more SOC due to burial of topsoil with eroded sediment,
potentially storing C for centuries (Van Oost et al., 2012; Hoffmann et al.,
2013; Johnson, 2014). The removal of weathered topsoil material from eroding
positions, the replacement of eroded SOC, and its burial at depositional
sites can potentially lead to a net sink for atmospheric C depending on the
fate of the eroded SOC (Harden et al., 1999; Doetterl et al., 2012; Wiaux et
al., 2014a). SOC at the depositional site is often regarded as more
stable with longer turnover times, depending on microbial activity,
environmental conditions (Wang et al., 2013), and biogeochemical
characteristics of the transported C fractions. However, areas (or
landscapes) with a fast burial can lead to the accumulation (storage) of
labile SOC, which is still vulnerable to decomposition if the conditions at
the site of burial change (Wiaux et al., 2014b). Thus, there is an ongoing
discussion about depositional sites of highly dynamic landscapes as C sink
or source.
Soil properties with relation to C stability, such as aggregate stability,
availability of reactive mineral surfaces or soil water saturation, differ
strongly between stable and dynamic landscapes, i.e., landscapes where high
rates of soil redistribution take place. This may have consequences for SOC
sequestration and stabilization in soils and requires the use of proxies to
investigate the stability of specific fractions. For example, the abundance
of amino sugars (AS) and the 14C derived age of a fraction can indicate
the potential of a specific mechanism to protect SOC against further
degradation. The large majority (> 99 %) of AS found in soils
are considered to be of microbial origin and constitute an important
building block in cell walls and extracellular polysaccharides (Glaser et al., 2004; Glaser and Gross, 2005; Simpson, 2004). In soils, only glucosamine (GluN), galactosamine
(GalN), and muramic acid are found in quantifiable concentrations. AS provide a generally readily available energy source for
microorganisms and are easy to decompose compared to more recalcitrant
organic matter such as lignin or lipids (Amelung et al., 2001, 2008; von
Luetzow et al., 2006; Roberts et al., 2007; Schmidt et al., 2011). AS turnover is, therefore, fast compared to other stabilized SOC
fractions. Turnover times of AS, however, are strongly affected by
interactions with the soil mineral matrix and the soil environment
(Bodé, 2013) and this can lead to an accumulation of AS after cell death
in soils (Guggenberger et al., 1999; Glaser et al., 2004). This fact can be
used to investigate the effectiveness of different stabilization mechanisms.
During the transport of sediment to and accumulation and burial at the site
of deposition, easily available SOC fractions have been decomposed. However,
at the depositional site, subsoil C is at least partly derived from buried
topsoil C. Hence, at the depositional site changes in the abundance of AS
and the 14C derived age of a fraction can indicate the potential of a
specific mechanism to protect SOC against further degradation (Z. Wang et al.,
2014).
Key soil and geomorphological properties and identified SOC
fractions as described by Doetterl et al. (2012) and Z. Wang et al. (2014) for
the investigated soils at different slope positions and soil depths.
Parameter
Units/Fraction
Stable
Eroding
Depositional
plateau
slope
footslope
Slope
[%]
3
10
4
Total erosion (-) & deposition (+)
[cm]
0
-200
350
Current erosion & deposition rate
[mm yr-1]
0
-1.32
1.16
Depth
[cm]
0–15
35–50
55–70
0–15
35–50
55–70
0–15
35–50
55–70
Bulk density
[g cm-3]
1.58
1.56
1.59
1.52
1.63
1.51
1.55
1.64
1.58
pH
7.0
7.1
7.0
7.2
7.3
7.1
6.8
6.8
6.9
Bulk soil
9.6
2.6
2.4
9.3
2.1
2.1
11.6
6.7
3.9
SOC
Macroaggregate
14.3
3.0
4.1
12.1
3.0
5.2
14.3
7.8
6.3
(g kg-1)
Microaggregate
7.3
2.6
2.2
7.3
2.8
2.4
9.3
5.9
3.5
Non-aggregated silt + clay
8.4
2.3
2.0
7.5
1.7
1.6
9.1
6.7
3.1
Relative C
Macroaggregate
48
19
24
46
19
27
58
37
35
(%bulk SOC)
Microaggregate
35
48
47
24
34
24
21
37
21
Non-aggregated silt + clay
17
33
29
30
47
49
21
26
44
Bulk soil
152
41
38
141
34
32
180
110
62
C stock
Macroaggregate
72
8
9
63
7
9
105
40
22
(gC m-2 cmdepth-1)
Microaggregate
54
19
18
33
12
8
37
40
13
Non-aggregated silt + clay
26
13
11
41
16
16
39
29
27
Our work is a conceptual approach to analyze and illustrate the mechanisms
of soil redistribution effects on C dynamics. For our study, we hypothesize
that the soil redistribution history of a soil profile influences the
present weathering status of the soil material, and is therefore a primary
control on the abundance and composition of the reactive soil mineral phase,
stabilizing SOC through geochemical (association of organic molecules with
minerals) and physical (aggregation) mechanisms. To complement our analysis
on the importance of different mechanisms to stabilize C along a geomorphic
transect, we measured the 14C and the abundance of microbial residues
in the form of AS in the different soil fractions. We further hypothesize
that with ongoing soil redistribution, the reactive soil mineral phase will
be highly variable, i.e., horizontally in space along the slope transect,
vertically with soil depth, and in time due to removal or burial of soil. We
use the total reserve in base cations (TRB in cmolc kg-1) to
characterize the weathering status of the soils along the geomorphic
transect and complement this with a qualitative and quantitative analysis of
clay minerals. If changes in the mineralogical composition of the
investigated soils due to soil redistribution occur at a faster pace than
weathering-related changes, distinct mineralogical differences between
profiles along the slope should be identifiable.
Material and methods
Study site and sampling
The resulting fractions of the applied fractionation scheme and
interpretation of the present carbon stabilization mechanisms in each
fraction.
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We performed our analysis on nine bulk soil samples and the dominating soil
C fractions that were collected in the study of Doetterl et al. (2012; Table 1) from a geomorphic transect on cropland situated in central Belgium
(50∘45′41.50′′ N; 4∘44′07.36′′ E) with a south/southeast
facing slope. The study area is characterized by a smooth rolling topography
with plateaus, slopes (up to 20 % slope) and dry valley bottoms. The
climate of the region is a temperate oceanic climate with mild winters and
cool summers (Köppen climate Cfb, 821 mm yr-1; 9.7 ∘C on
30-year average; IRM, 2011). The geological substrate is a several meters
thick Pleistocene aeolian deposit of calcareous loess, in which Luvisols
have developed and are overlaying Tertiary sands (Wouters and Vandenberghe,
1994). Soils are well drained and show no evidence of long-lasting
hydromorphic conditions. First traces of agriculture in the region date back
to the Late Bronze/Iron Age and both sites have been under continuous
agricultural land use since at least 1770 (Rommens et al., 2005; Lannoo,
2009). No detailed information about the long-term (i.e., several centuries)
crop rotation or SOC input is available. However, this study focuses on the
relative differences between geomorphic positions that were managed in the
same way. Soil samples were taken from a stable, non-eroding profile at the
hilltop plateau, an eroding profile from the hillslope shoulder (200 cm soil
loss), and a depositional profile in the colluvial valley bottom (350 cm soil
gain). In order to quantify the vertical and horizontal distribution of SOC
fractions, the soil cores were cut into the following depth intervals: the
topsoil (0–15 cm), the shallow subsoil (35–50 cm), and the deeper subsoil
(55–70 cm).
SOC fractionation
The SOC content was measured in duplicate on 1 g ground soil subsamples
using a VarioMax CN dry combustion analyzer (Elementar GmbH, Germany) with a
measuring range of 0.2–400 mg C g soil-1 (absolute C in sample) and a
reproducibility of < 0.5 % (relative deviation) on powdered soil
samples. To derive functional SOC fractions, Doetterl et al. (2012) used a
method based on the conceptual SOC fraction model proposed by Six et al. (1998, 2002; Fig. 1). For details on the gathered fractions see the
original study (Doetterl et al., 2012) and the concepts of Six et al. (1998,
2002). In short, the scheme consists of a series of chemical and physical
fractionation techniques applied to isolate functional SOC fractions,
differentiated by stabilization mechanisms (chemical, biochemical and
physical) which can also be associated with different (potential) turnover
times (see also von Luetzow et al., 2006). Background information to the key
soil and geomorphological properties as well as the abundance and
composition of identified SOC fractions as described by Doetterl et al. (2012) and Z. Wang et al. (2014) for the investigated soils can be found in
Table 1. For our analysis on the microbial composition and 14C age of
the fractions we analyzed the environment in which C is stabilized, hence
analyzing macroaggregates and microaggregates as a whole. For the mineralogical
analysis we are focusing on comparable fractions, hence comparing aggregated and non-aggregated silt and clay
fractions.
Soil physico-chemical characterization
Soil pH values were determined as the mean of two measurements per sample
after a response time of 30 min and 24 h respectively in a 1:2.5
soil / solution ratio in 25 mL 0.01 M CaCl2 potentiometrically with a
glass electrode using a portable multiparameter Meter HI9828 (Hanna
Instruments US Inc., USA). All samples have shown no reaction when treated
with 10 % HCl and are considered free of carbonates.
The soil weathering degree in the different soil depth layers was estimated
by measuring the total reserve in bases (TRB, the sum of total content in
Ca, Na, K, Mg, in cmolc kg-1) following Herbillon et al. (1988).
The TRB can be used to compare soil horizons relative to the parent material
to evaluate the weathering degree of soil material by assessing the relative
loss of Ca, Na, K, and Mg cations during weathering. Total elemental content
was determined by inductively coupled plasma–atomic emission spectrometry
(ICP-AES) on the bulk soil of the three profiles and three depths and the
parent loess material after borate fusion (Chao and Sanzolone, 1992).
A three-step sequential extraction scheme of pedogenic organo-mineral
associations and oxy-hydroxides (Stucki et al., 1988) was carried out in
duplicate in the following order: sodium-pyrophosphate at pH 10 (Bascomb,
1968), ammonium oxalate–oxalic acid at pH 3 (Dahlgren, 1994), and
dithionite–citrate–bicarbonate (DCB) at pH 8 (Mehra and Jackson, 1960). This
allows assessing the amount of Mn-, Mg-, Fe-, and Al-bearing phases in the
different fractions and their correlation with organic C in the different
SOC fractions. The specific extraction was performed on the bulk soil of
each sampling position, the (micro)aggregated (s + cm) and non-aggregated
(s + c) silt and clay associated SOC fractions separately. The last two of these
were identified as the key fractions to change in
their abundance along the hillslope between eroding and depositional
(sub-)soils (Doetterl et al., 2012) and are the building blocks for larger
aggregates. Each extract was analyzed for its Al, Fe, Mg, and Mn content by
ICP-AES.
In our sequential extraction, pyrophosphate (p) extractable components
are interpreted as predominantly organically complexed metals. Oxalate
(o) extractable components reflect the amorphous secondary Fe and Mn
oxides and poorly crystalline aluminosilicates (imogolite-type materials,
ITM). Dithionite (d) extractable components included
predominantly crystalline oxy-hydroxides of Mn, Mg, Fe, and Al. Several
authors could show that pyrophosphate extractable Al may not be attributable
only to Al bound to organo-metallic complexes since the alkaline extractant
could also extract Al from Al hydroxide phases and from poorly crystalline
aluminosilicates (i.e., Schuppli et al., 1983; Kaiser and Zech, 1996). The
results of the pyrophosphate extraction must, therefore, be treated with
caution due to uncertainty about the origin of the extracted minerals.
Hence, we limit our analysis to interpreting the abundance and spatial
patterns of the various extractable components in the bulk soil and the SOC
fractions and discuss the pedological implications of the observed
carbon/mineral correlations.
Soil mineralogy
The mineralogy of the clay-sized fraction was determined by X ray
diffraction (XRD, Cu Kα, D8, Brucker Advance) after K+ and
Mg2+ saturation, ethylene glycol solvation, and thermal treatments at
300 and 550 ∘C (Robert and Tessier, 1974). Clay minerals were
classified and peak-identified according to Brindley and Brown (1980). Quantitative
analyses were performed using the Rietveld method through the software
package Siroquant V4.0. Refinement of Rietveld parameters was carried out
following the instructions provided in the “Siroquant V4.0 Technical and
Clay Manuals” until a chi-square value < 3 was obtained. Key primary
minerals that can interfere with the spectra (quartz) and act as sources for
clay formation such as amphiboles (hornblende), feldspars (albite,
orthoclase, plagioclase), mica (muscovite, phengite, biotite) as well as
olivine and pyroxene have been quantified in order to analyze their relative
abundance in relation to the secondary aluminosilicates. For the latter, we
focused on the identification of expandable 2:1 layered minerals such as
smectite and vermiculite, the non-expandable 2:1 layered illite and chlorite
and the non-expandable 1:1 layered kaolinite. Note that no distinction has
been made between primary and secondary chlorite. In a final step, the
quantified primary and secondary minerals were then compared along the
geomorphic gradient, i.e., soil layers from the same depth across different
positions or soil layers from different depths but the same profile, to find
relevant trends.
Compound-specific analysis of AS
AS concentrations were determined using liquid chromatography for
the bulk soil, macroaggregates (> 250 µm), and microaggregates
(53–250 µm) as well as for the non-aggregated silt and clay fractions
(< 53 µm). Each sample was extracted in duplicate. AS extraction and analysis were based on the procedure described by
Bodé et al. (2009). Briefly, samples corresponding to 0.3 mg of N were
successively hydrolyzed with 6M HCl (20 mL g-1 of sample) at 105 ∘C for
8 h after which AS were purified on a cationic
exchange resin (AG50W-X8, 100-200 Mesh, Hydrogen form, Bio-Rad lab.) and
eluted by protons prior to liquid chromatography–isotope ratio mass spectrometry (LC-IRMS) analysis. The chromatographic separation
was performed using an LC pump (Surveyor MS-Pump Plus, Thermo Scientific,
Bremen, Germany) mounted with a PA20 CarboPac analytical anion-exchange
column (3 × 150 mm, 6.5 µm) and a PA20 guard column (Thermo Scientific,
Bremen, Germany). Basic AS (GluN and GalN) were eluted with 2 mM NaOH and a
column temperature of 15 ∘C. An analysis of the repeatability
between replicates shows deviations of ca. 11 % between the replicates for
GalN with contents of > 100 µg GalN g soil-1. For
smaller values, deviations were higher (ca. 31 %). For GluN with contents
of > 250 µg GluN g soil-1 deviations of c. 12 % were
observed, whereas smaller values differed about 40 % between replicates.
Hence, we limit our analysis of this data to observations that exceed the
abovementioned uncertainty and to a description of trends within the
data set.
Isotopic composition and <inline-formula><mml:math display="inline"><mml:msup><mml:mi/><mml:mn>14</mml:mn></mml:msup></mml:math></inline-formula>C age
Using 14C radiocarbon dating we estimate the age of the organic
material associated with the macroaggregates, microaggregates and the
non-aggregated silt and clay fractions at the depositional site in order to
see relative differences between fractions and burial depths. We
use these differences to evaluate the effectiveness of fractions in stabilizing C against decomposition after burial. In short,
roughly 1 mg C was sealed into an evacuated Pyrex tube and reduced to
graphite (Xu et al., 2007). Sample preparation backgrounds have been
subtracted, based on measurements of 14C-free coal. The radiocarbon
signature of the graphite was measured with accelerator mass spectrometry
(NEC 0.5MV 1.5SDH-2 AMS system) at the Keck-Carbon Cycle AMS facility at UC
Irvine (CA, USA). Results have been corrected for isotopic fractionation
according to the conventions of Stuiver and Polach (1977), with δ13C values measured on the prepared graphite using the AMS
spectrometer. Radiocarbon concentrations are given as fractions of a modern
oxalic acid standard and conventional radiocarbon age following the
conventions of Stuiver and Polach (1977).
Implementation and statistical analyses
All statistical analysis was realized using SAS 9.3 (SAS Institute Inc.,
Cary, NC) and R 2.11.1 (R Development Core Team, 2010). Differences between
the means of classes have been performed using multi-group ANOVA Bonferroni
corrections and Tamhane's T2. We also performed linear regressions to test
correlations between the abundance of the reactive soil phases and SOC in
the bulk soil and for the isolated fractions. All statistical tests were
evaluated using p < 0.1 as the level of significance.
Results
Total reserve in bases and degree of soil weathering along the
hillslope
The TRB for the loess parent material was 139 cmolc kg-1. For the
different soil layers, TRB ranged between 101 and 118 cmolc kg-1, which represents a TRB loss of 15–27 %. These values are
in the same range as reported for other Luvisol profiles in the region
(Brahy and Delvaux, 2001). Losses in TRB are mostly related to decalcification
(Ca losses compared to parent material: 80 ± 2 %) (Table 2). The
strongest losses of cations (lowest TRB in Table 2), and hence the most
advanced weathering stage, are observed for the deposited soils at the
foothill (relative difference of TRB at the depositional site compared to
the loess parent material: 25–27 %), while the layers at the eroding
hillslope have been facing the lowest losses of cations (highest TRB in
Table 2), and hence the weakest degree of weathering (relative difference of
TRB at the depositional site compared to the loess parent material 15–20 %). Nevertheless, these differences are small compared to differences
between soil types in the same region (Brahy et al., 2001) and indicate a
rather small effect of soil redistribution on TRB in comparison. At the
stable plateau and the eroding hillslope, a significant difference
(p < 0.1) with higher TRB at greater depth can be observed, but with
weaker increases of TRB with soil depth than reported from similar soils
(Brahy et al., 2001). In contrast, no effect of depth on TRB was observed in
the depositional profile.
Total Ca, K, Mg and Na content and TRB values for the soil profiles at several depths and for the parent material.
Parameter
Unit
Stable
Eroding
Depositional
Parent material
plateau
slope
footslope
(loess)
Depth
[cm]
0–15
35–50
55–70
0–15
35–50
55–70
0–15
35–50
55–70
> 100 cm
Ca
14
13
13
15
12
14
15
13
13
70
K
46
46
48
47
50
46
42
41
43
32
Mg
[cmolc kg-1]
15
20
19
17
19
19
14
14
16
18
Na
31
34
32
32
33
38
32
32
32
18
Total TRB
107
113
112
111
115
118
103
101
103
139
TRB relative
(%)
77
81
81
80
83
85
75
73
74
100
to parent material
The distribution of pyrophosphate (I), oxalate (II) and dithionite
(III) extractable Al, Mn and Fe between the bulk soil and its compartments
non-aggregated silt + clay (s + c) and aggregated silt + clay (s + cm) for
all depths and positions combined (bars). Dots represent the the C extracted
by pyrophosphate for the bulk soil and the fractions (Panel I only).
Different letters above bars indicate a significant difference (p < 0.1, tested for extracted element in the different SOC fractions and
treatments separately as indicated by differences in font types).
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Reactive soil phases
Pyrophosphate extractable components are generally abundant in
microaggregates, especially as Fe(p) (1118 ± 293 mg kg-1 soil) and
Al(p) (1647 ± 385 mg kg-1 soil) (Fig. 2). The ratio of
non-aggregated to aggregated pyrophosphate extracts in silt and clay is
generally between 1:3 and 1:7 and shows a clear association of these
components with soil microaggregates. In contrast, oxalate and DCB
extractable phases do not show such a pattern between aggregated and
non-aggregated soil samples. No significant differences in the amount of
oxalate and DCB extractable components in aggregated vs. non-aggregated soil
fractions were detected for Fe, Mg, and Mn. Al(o) showed about 40 % lower
concentrations in non-aggregated silt and clay (s + c) compared to
aggregated silt and clay (s + cm). Approximately the same amount of oxalate
extractable Al, Fe, Mg, and Mn is present in aggregated (1690 ± 378 mg kg-1 soil)
and non-aggregated (2421 ± 648 mg kg-1 soil) soil
samples, with 5 times more DCB extractable elements than oxalate in
both soil compartments.
Despite the small contribution of Mn(p) (40 ± 36 mg kg-1 soil) to
the total pyrophosphate extractable phase of the bulk soil compared to Al(p)
(122 ± 64 mg kg-1 soil) and Fe(p) (222 ± 115 mg kg-1 soil) and Mg(p) (101 ± 33 mg kg-1 soil), it is the only phase
with a significant correlation to SOC in the bulk soil across the whole
slope transect (Fig. 3). Table 3 shows that the positive correlation of
Mn(p) with SOC along the slope for the bulk soil is driven by the
non-aggregated silt and clay fractions. In contrast to Mn(p), a correlation
of Fe(p) and Al(p) with SOC could only be identified at the
depositional site. Poor or negative correlations were found for all
pyrophosphate extractable phases with SOC in the microaggregate silt and
clay (s + cm) fraction. Depth patterns of Al(p), Fe(p) and Mg(p) are not
consistent for eroding and stable profiles, but a significant decrease with
depth can be recognized at the depositional site. Mn(p), on the other hand,
decreases significantly (p < 0.1) with depth at all slope positions
except for s + c at the depositional site and for s + cm at the eroding and
depositional site.
Correlations (Pearson r) between pyrophosphate, oxalate and DCB
extractable solid soil phases and SOC or soil depth (*=p < 0.1).
SOC fraction
Position
Parameter
Pyrophosphate extractable
Oxalate extractable
DCB extractable
Al(p)
Fe(p)
Mg(p)
Mn(p)
Al(o)
Fe(o)
Mg(o)
Mn(o)
Al(d)
Fe(d)
Mg(d)
Mn(d)
Pearson r phase vs. depth
0.34
-0.45
-0.27
-0.95*
0.84*
0.09
-0.73*
0.37
0.74*
0.84*
0.60
0.61*
Non-eroding
Pearson r phase vs. SOC
-0.37
0.4
0.06
0.99*
-0.87*
0.02
0.84*
-0.09
-0.57
-0.68*
-0.33
-0.36
Range mg kg-1
70–223
118–410
62–119
6–73
647–1076
1512–1928
38–75
255–379
788–1280
6401–12161
81–132
43–60
Pearson r phase vs. depth
0.17
-0.12
0.1
-0.9*
0.09
-0.85*
-0.41
0.83
0.18
0.57
0.53
-0.53
Bulk soil
Eroding slope
Pearson r phase vs. SOC
0.04
0.34
0.13
0.99*
-0.38
0.75*
0.41
-0.95
-0.35
-0.72*
-0.54
0.55
Range mg kg-1
53–298
91–498
58–207
6–84
554–865
829–1531
42–66
191–282
783–1027
6277–9593
73–128
44–79
Pearson r phase vs. depth
-0.23
-0.67*
-0.68
-0.99*
0.83*
0.05
-0.43
0.99*
0.5
0.64*
0.29
0.25
Depositional
Pearson r phase vs. SOC
0.23
0.67*
0.69
0.99*
-0.83*
-0.05
0.43
-0.99*
-0.49
-0.64*
-0.29
-0.25
footslope
Range mg kg-1
58–180
106–320
69–133
18–112
432–636
930–1157
40–65
181–265
569–1234
5319–10324
67–146
41–60
Pearson r phase vs. depth
0.61*
0.05
0.24
-0.88*
-0.17
-0.33
0.19
-0.42
0.8
0.97*
0.12
0.56
Non-eroding
Pearson r phase vs. SOC
-0.47
0.2
-0.02
0.95*
-0.08
0.57
0.18
0.46
-0.92*
-0.99*
0.17
-0.39
Range mg kg-1
105–403
67–252
42–93
10–107
792–1842
1557–1858
38–75
152–301
1003–1626
7188–10760
100–137
53–57
Pearson r phase vs. depth
0.44
0.18
0.29
-0.74*
0.16
-0.92*
0.29
0.64*
0.21
0.59*
0.78*
0.54
Non-aggregated
Eroding slope
Pearson r phase vs. SOC
-0.27
-0.06
-0.12
0.72*
-0.49
0.81*
-0.36
-0.77*
-0.32
-0.81*
-0.82*
-0.46
s + c
Range mg kg-1
87–568
46–362
33–152
4–153
746–1035
686–1360
42–63
113–168
1016–1294
8288–10783
115–172
50–56
Pearson r phase vs. depth
-0.76*
-0.48
-0.5
-0.5
0.54
0.55
0.57
0.29
0.71*
0.76*
0.07
0.85*
Depositional
Pearson r phase vs. SOC
0.76*
0.48
0.5
0.5
-0.46
-0.51
-0.63*
-0.2
-0.71*
-0.76*
-0.07
-0.85*
footslope
Range mg kg-1
107–425
73–271
31–125
13–168
513–871
802–1085
44–62
68–175
855–1087
6612–8795
117–167
45–69
Pearson r phase vs. depth
0.80*
0.58
0.5
-0.69*
0.75*
0.62*
0.42
0.52
0.33
0.54
0.21
-0.4
Non-eroding
Pearson r phase vs. SOC
-0.96*
-0.76*
-0.64*
0.81*
-0.98*
-0.77*
0.07
-0.82*
-0.82*
-0.94*
-0.54
-0.21
Range mg kg-1
1290–1947
934–1189
165–219
62–169
460–760
1113–1476
36–44
85–166
647–1004
5797–9048
73–103
32–45
Pearson r phase vs. depth
0.68*
0.39
0.32
-0.31
0.72*
-0.41
0.52
0.93*
0.49
0.46
0.8*
0.65*
Aggregated
Eroding
Pearson r phase vs. SOC
0.57
0.71*
0.58
0.38
-0.81*
0.39
-0.55
-0.98*
-0.61
-0.62*
-0.8*
-0.72*
s + cm
Range mg kg-1
1117–1954
681–1599
142–291
89–216
463–688
812–1050
35–56
69–124
753–1064
7027–10764
83–136
34–40
Pearson r phase vs. depth
0.2
0.37
0.34
0.43
0.27
-0.38
0.22
0.52
0.2
0.52
0.61
-0.32
Depositional
Pearson r phase vs. SOC
-0.27
-0.25
-0.09
0
-0.22
0.4
0.16
-0.54
-0.24
-0.49
-0.21
0.37
footslope
Range mg kg-1
1137–2685
849–2021
132–356
58–221
286–529
640–945
32–50
39–66
557–789
5017–7972
78–114
29–33
Mean SOC (g kg soil-1) and pyrophosphate extractable Al (I), Mg
(II), Mn (III), and Fe (IV).
![]()
No strong correlation of oxalate extractable phases in the bulk soil, s + c
or s + cm with SOC could be identified. Oxalate and DCB and extractable
elements are generally not significant or negatively correlated to SOC,
except for a significantly positive correlation of Fe(o) to SOC in the bulk
soil and non-aggregated silt and clay fraction.
Clay-sized fraction mineralogy
The abundance of secondary and primary minerals in the clay fraction differs
largely along the transect and for different depths (Fig. 4), while the
concentration of quartz remains fairly constant across the different
transect profiles and soil depths (15 ± 1.6 %, data not shown).
Generally, the eroding slope profile has the highest number of investigated
primary minerals (26–45 %), while the lowest values are measured at the
stable plateau (31–36 %) and the deepest layer of the depositional site
(26 ± 0.6 %). The opposite trend can be observed for the abundance of
pedogenic secondary minerals with the highest contents measured at the
stable plateau (40–62 %) and the depositional site (38–52 %) while
the erosional profile shows the lowest values (35–42 %). A closer
analysis of the secondary mineral fraction of all investigated soils shows
the presence of both expandable (vermiculite, smectite) and non-expandable
2:1 layered clay minerals (illite; chlorite). Also larger amounts of 1:1
layered clays such as kaolinite were found, partly inherited from the parent
material and as to be expected in Luvisols of this region (Van Ranst et al.,
1982; Fig. 4). Within the clay fraction, the expandable 2:1 layered mineral
ranges between 6–26 % relative abundance with highest values at the
depositional topsoil (26.1 ± 0.5 %) and lowest values at the
eroding slope topsoil (5.9 ± 0.4 %). The abundance of illite is
highest at the erosional subsoils (32 ± 0.9 %), with no clear depth- or
slope-related pattern for the remaining samples (10–24 %), while the
highest relative abundance of chlorite has been measured at the depositional
deeper (55–70 cm) subsoil (49 ± 0.7 %). Kaolinite contents are
generally increasing with depth at the stable plateau (22–62 %) and
decreasing with depth at the eroding profile (46–22 %), while no pattern
was observed for the depositional profile (35–64 %).
Composition of the mineral phase and quantification of primary vs.
secondary minerals for the different slope positions and soil depths of the
clay sized fraction (no distinction between primary and secondary chlorite).
![]()
AS abundance
The contribution of AS mg g-1 SOC (Fig. 5) at the depositional site is several times
higher than at other positions, especially for glucosamine. Generally, AS
concentrations per unit C decrease with the size of the fraction and are
especially low in the non-aggregated silt and clay fraction. In addition,
the non-aggregated silt and clay fraction shows generally the highest
differences between topsoil and subsoil abundance of AS per unit C at all slope
positions; this is especially pronounced at the stable plateau and eroding
slope position. Moreover, the bulk soil AS values are similar for the
topsoil along the slope (GalN-C: 22.5 ± 1.9 mg g-1 SOC; GluN-C:
34.7 ± 22.2 mg g-1 SOC) with highest values at the depositional
site. In contrast, subsoils of the stable plateau (GalN-C: 7.2 ± 0.6 mg g-1 SOC;
GluN-C: 16.3 ± 4.6 mg g-1 SOC) and the eroding
slope (GalN-C: 6.3 ± 1.0 mg g-1 SOC; GluN-C:
15.7 ± 3.5 mg g-1 SOC) profile show significantly (p > 0.1) lower values
compared to subsoils at the depositional site (GalN-C: 25.1 ± 6.1 mg g-1 SOC; GluN-C: 54.2 ± 6.9 mg g-1 SOC). All fractions
contain a large portion of AS per unit C in the topsoil (GalN-C: 19.1 ± 3.5 mg g-1 C; GluN-C: 42.0 ± 9.7 mg g-1 C), while in
subsoils the contribution of AS to total C is higher in the macroaggregates
(GalN-C: 16.3 ± 2.5 mg g-1 C; GluN-C: 36.5 ± 7.3 mg g-1 C) than in the microaggregate
(GalN-C: 12.7 ± 6.0 mg g-1 C; GluN-C: 31.1 ± 13.0 mg g-1 C) and the
non-aggregated silt and clay fraction (GalN-C: 6.5 ± 5.7 mg g-1 C; GluN-C: 12.2 ± 12.2 mg g-1 C). Interestingly, at the
depositional site, the amount of AS per unit C in the microaggregate and
non-aggregated silt and clay fraction decreases with depth compared to the
macroaggregate fraction, where no such trend could be observed (Fig. 5).
The abundance of the two investigated AS is closely correlated to SOC in the
bulk soil and across all fractions (Table 4), whereas the correlation of
GalN to SOC is higher (r= 0.97*) than the correlation of GluN to SOC
(r= 0.71*). The smallest changes in the abundance of AS and SOC g soil-1 (Fig. 5) are
observed between topsoil and subsoil in the macroaggregate fractions (Ratio
topsoil vs. subsoil: 1.7–3.6), while changes in the microaggregate fractions
(Ratio topsoil vs. subsoil: 3.1–11.5) and especially in the non-aggregated
silt and clay fractions are much higher (Ratio top- vs. subsoil: 4.9–32.8).
Generally, the highest abundance of AS was found in topsoils, and the lowest
in deeper subsoil layers. But this trend is much weaker in the depositional
profile than in the non-eroding plateau and eroding slope profile.
AS per unit SOC in the bulk soil (I) and the fractions
(II–IV) along the slope and for different depths. Different letters above
bars indicate a significant difference (p < 0.1, for the two
different AS and the bulk soil and three fractions separately as indicated
by differences in font types).
![]()
Correlations (Pearson r) between SOC and the extracted AS
galactosamine (GalN) and glucosamine (GlcN) from different fractions (*=p < 0.1).
Bulk soil
Macroaggregate
Microaggregate
Non-aggregated s + c
Pearson r GalN-C
0.97*
0.87*
0.94*
0.92*
Pearson r GlcN-C
0.71*
0.76*
0.96*
0.87*
N Observations GalN-C
31
16
19
18
N Observations GlcN-C
31
14
16
20
Range GalN-C
12–325
33–360
4–248
2–206
µg g-1 soil
Range GlcN-C
26–860
60–887
23–546
4–403
µg g-1 soil
Mean GalN-C
296 ± 280
172 ± 93
100 ± 82
69 ± 65
µg g-1 soil
Mean GlcN-C
555 ± 679
396 ± 224
228 ± 183
138 ± 130
µg g-1 soil
Conventional radiocarbon ages for investigated fractions for
different depths at the depositional site.
Fraction
Depth
14C age
(cm)
(years BP)
Macroaggregates
0–15
975 ± 15
(> 250 µm)
35–50
2275 ± 20
55–70
3090 ± 20
Microaggregates
0–15
680 ± 15
(< 250–53 µm)
35–50
1680 ± 15
55–70
2670 ± 20
Non-aggregated silt
0–15
635 ± 15
and clay (< 53 µm)
35–50
1835 ± 15
55–70
2200 ± 15
Radiocarbon ages for fractions at the depositional site
The measured 14C ages (years BP) for the investigated fractions at the
depositional site differ widely for different soil depths (Table 5).
Generally, the 14C ages increase with depth and range between
635 ± 15 and 3655 ± 15 years. Consistently, the macroaggregate
fraction has the oldest 14C age at all depths of all fractions
(975 ± 15–3090 ± 20 years), representing distinctively older C
than associated with the respective microaggregate (680 ± 15–2670 ± 20 years)
or non-aggregated silt and clay fraction (635 ± 15–2200 ± 15 years). Furthermore, the data show a shift in the
relative age of microaggregate vs. non-aggregated C. In the topsoil and
shallow subsoil the measured 14C ages for these fractions are rather
similar (difference of approx. 50–150 yr). In the deepest subsoil a clear
distinction can be observed, with microaggregate associated C being 500 years older than C associated with non-aggregated silt and clay.
Hence, 14C ages for the non-aggregated silt and clay fraction in
subsoils indicate a relatively young age of this fraction compared to the
aggregated fractions.
Discussion
Organo-mineral interactions along the toposequence
The reactive mineral phases, predominantly manganese, that are related to
the abundance of SOC seem to be largely extracted with pyrophosphate (Fig. 3, Table 3). Hence, we suggest that manganese significantly influences the
dynamics of SOC in the investigated soil profiles by promoting the formation
of organo-mineral complexes and by the redistribution of these complexes
along the toposequence. The measured C concentrations in the pyrophosphate
extract (Panel I, Fig. 2) represented about 81 % of the total bulk C in
these soils. This points to a minor importance of well-crystallized minerals
to stabilize C, as extracted with DCB, and confirms earlier findings. Kodama
and Schnitzer (1980) could show that the presence of organic ligands in the
soil solution can prevent the formation of crystalline Al and Fe
oxy-hydroxides, e.g., due to strong complexation gradients of fulvic acids
with metals which prevent their complete hydroxylation. However, the absence
of correlations between SOC and poorly crystalline minerals extracted with
oxalate is contradictory to former studies (Kodama and Schnitzer, 1980;
Kaiser and Zech, 1996; Mikutta et al., 2009). Oxalate extractable soil phases
are generally assumed to be the most reactive group in many types of soils and
presumably strong candidates for stabilizing SOC with minerals by adsorption
on mineral surfaces (Huang et al., 1977; Kaiser and Guggenberger, 2003;
Kleber et al., 2005; Eusterhues et al., 2005). We interpret our result in
terms of (i) the applied extraction scheme, (ii) the specific soil
conditions in which interactions between the reactive mineral phase and SOC
take place, and (iii) the varying mobility of the extracted phases in
those soils.
First, in contrast to many other studies, our extraction is sequential, with
pyrophosphate, oxalate, and DCB applied in this order on the same sample.
Other studies that do not use pyrophosphate extract with oxalate a
potentially large part of organo-mineral associations. These associations
are in our design already partly extracted with pyrophosphate. Second, even
though concentrations of pyrophosphate extracted iron and aluminum were high
in microaggregates, no strong positive correlations with C were found for
these elements (Table 3, Fig. 2). The fact that pyrophosphate extractable
elements are generally abundant in aggregates raises questions about the
importance of the isolated elements for supporting the physical protection
of C against decomposition. One explanation could be that pyrophosphate
extractable Fe and Al are important for aggregate stability as ligands
between clay minerals, even without involving SOC, but not for SOC retention
itself. Fe(p) and Al(p) could act as proxies for the formation of aggregates
through the formation of organo-metallic complexes, but also through the
formation of Fe, Al oxy-hydroxides (Kaiser and Zech, 1996). Pyrophosphate in
that case might simply act as a strong reagent to disaggregate soil (Muneer
and Oades, 1989) and therefore the abundance of pyrophosphate extractable
elements in aggregates does not point at an identifiable group of
organo-mineral associations (Kaiser and Zech, 1996). Third, the pH-values in
our soils are near neutral (Table 1), which implies a pH buffer zone where Mn
is highly mobile as Mn2+, potentially forming organo-mineral complexes,
while the mobility of Fe and Al is strongly limited at pH > 6
(Lindsay, 1979). However, the data indicate a potential impact of oxalate
extractable Fe oxides to stabilize SOC where physical protection of SOC by
aggregation is weak, namely at the eroding site (Table 3). There, the
aggregation potential is small (Doetterl et al., 2012) due to the continuous
removal of aggregate-rich topsoil layers.
Importance of clay-sized mineral distribution
The presented data for the abundance of different clay-sized minerals (Fig. 4) clearly indicate a large variety of clay minerals along the slope
transect. In combination with the high TRB values at the eroding site, the
large abundance of primary minerals and illite (= non-expandable 2:1
layered clay) and the relatively low content of pedogenic clay minerals (1:1
and 2:1 layered clay minerals) indicate a lower level of weathering at the
eroding compared to the plateau and depositional site. These data are in
accordance with the general assumption that kaolinite is one of the end
products in the weathering sequence in soils and that the eroding profile is
less weathered than the stable and depositional profile (see Sect. 3.1).
The deepest layer of the eroding profile is the least weathered part of soil
along the sequence. Kaolinite there is mostly derived from the parent
material, hence lowest in comparison to other samples undergoing more
advanced weathering.
The applied quantitative approach allows us to document a relative depletion
of expandable 2:1 clay minerals (smectite and vermiculite) in the topsoil of
the eroding site (3.2 ± 0.4 %) and their enrichment in the topsoil of
the depositional site (15.4 ± 1.7 %). This suggests the preferential
mobilization of smectite and vermiculite at the eroding site and their
deposition at the foothill as soil gets mobilized and eroded mostly as
aggregates (Wang et al., 2010). Aggregates dominated by expandable clay
minerals have been identified to be less stable due to shrinking and
swelling than kaolinite or illite-rich aggregates (Fan et al., 2008) or
aggregates rich in amorphous and crystalline Al and Fe oxide (Goldberg,
1989). The lower stability of smectite or vermiculite-dominated aggregates
then leads to a preferential breakdown and mobilization of these aggregates,
resulting in smaller particles more susceptible to erosion, leading to
transport to and burial at depositional sites. Our data (Fig. 4 and Doetterl
et al., 2012) are consistent with the observation of the preferential export
of small particle sizes out of the catchment, which are often dominated by
smectite colloids (Gibbs, 1967). At the depositional site, weathering
continues and alters clay minerals which are involved in the formation of
aggregates and SOC stabilization, for example by transforming expandable
clay into non-expandable forms such as chloride (Fig. 4; depositional
footslope 55–70 cm). The two mechanisms together (weathering and transportation)
can hence explain the low amount of aggregates at eroding sites, and the
loss of aggregates with depth at the depositional site (Table 1) as
weathering continues. However, we cannot elaborate on whether the breakdown of
aggregates at the depositional site is induced first by decomposition of C
or mineral weathering. Nevertheless, no strong and significant correlations
between specific clay minerals and C content in different fractions or the
clay content and SOC concentration could be identified along the slope (data
not shown). Even though a direct quantification of the impact of clay
minerals on stabilizing SOC is not possible with our analysis, the data
indicate that in these soils with rather low bulk SOC concentrations of
1.2–0.2 %, clay surfaces are likely not the limiting factor for C
stabilization, even if the amount of highly reactive clay minerals such as
vermiculite or smectite is small. This is consistent with observations of
Duemig et al. (2012), where higher SOC loadings of clay minerals were
observed in consequence of a shortage of reactive surface area in clay-depleted soils compared to more clay-rich soils.
Conceptual figure showing the mineralogical changes along the
slope in relation to the abundance of AS in non-aggregated silt and clay
(s + c), microaggregates, and macroaggregates. Statement on
the level of soil weathering along the transect refers to both topsoil and
subsoil. Pie charts show the relative contribution of AS in each fraction to
total AS by multiplying AS g C-1 in a fraction with C stock estimates of
the respective fraction given in Table 1. The size of the pie chart
indicates the absolute AS content of these fractions combined in g m-2 and per cm depth of a profile. The amount of C extracted with
pyrophosphate in the bulk soil is given as a column next to each AS pie
chart. Panels at the top right show the distribution of primary and
secondary minerals along the slope in different depths (I) (see also Table 4) and the relative 14C derived ages of these fractions in topsoil and
subsoil for the depositional site (II), indicating differences between the
fractions (see also Table 5).
![]()
Radiocarbon ages in relation to microbial residues, respiration, and particulate organic
matter
The fact that the AS content per unit C in the macroaggregate and, to a
weaker extent, in the microaggregate and non-aggregated silt and clay
fractions at the depositional site does not decrease significantly after
burial, in comparison to very low subsoil AS concentrations at the eroding
site (Fig. 5), leads to the conclusion that burial of soil can lead to the
storage of an otherwise relatively easily decomposable part of SOC in soils
for decades and centuries. At the same time C in macroaggregates has shown
the oldest 14C ages of all isolated fractions (Table 5).
X. Wang et al. (2014) report, for a cropland slope from the same region with
similar topographic setting, a re-aggregation of deposited and buried C
within macroaggregates. Our data showing old 14C ages of macroaggregate
associated C are supportive of this, indicating indeed a preferential
accumulation of old carbon in macroaggregates at the depositional site
(Table 5). The measured old 14C ages and the high potential respiration
rates measured for these soils rich in macroaggregates (i.e., Doetterl et
al., 2012) indicate furthermore the presence of C sources with differing
availability to decomposers within the macroaggregate fraction. On one
hand, these aggregates represent a proper habitat for microorganisms to
thrive on C (Denef et al., 2001). Respiration data (presented in our former
work; Doetterl et al., 2012) clearly show that soils in this sequence with
large amounts of macroaggregate C also respire at a higher rate than soils
low on macroaggregate C. On the other hand, the presence of high amounts of
microbial residues (Fig. 5) can enhance the formation of stable aggregates
(Bossuyt et al., 2001; Cotrufo et al., 2012). Hence, macroaggregates
represent hotspots of microbial life and activity (Kolb et al., 2009) with
at the same time old 14C ages, indicating higher stability. A
conversion of SOC from non-microbial to microbial carbon within stable
aggregates would then not alter the 14C signature of the fraction, but
the potential stability of this fraction would still change. This is
supported by recent studies that highlight that the genetic variety of
microbial communities is maintained in buried surface soils, and that large
amounts of microbial biomass can be found in deeper layers of burial sites,
driven by the abundance of organic matter as a nutrient source (Helgason et
al., 2014). Hence, we argue that 14C ages and high AS contents in
macroaggregates indicate that microorganisms either feed on old carbon to
build up microbial biomass, if conditions are suitable, or these
environments contain considerable amounts of old C and are proper habitats
for microorganisms where their residues are preserved.
Soil redistribution and the depth-related changes of carbon at the depositional
sites
Our results show the ongoing rejuvenation of the soil at the eroding
hillslope by removal of soil material and exposure of less weathered material to
the surface. The lower loss of cations compared to the parent material (Table 2) indicates that the continuous removal of soil at the hillslope is acting
at a faster rate than the natural soil weathering and this keeps the eroding
profile TRB higher, and hence less weathered, than at the stable plateau
site. This is consistent with the absence of a depth-related trend in the TRB
at the depositional site (Table 2) as this material is buried former topsoil,
similar in TRB. Changes in C stocks with depth at the depositional sites can
shed light on the underlying mechanistic relationships as the buried topsoil
should have similar C content and composition to the present topsoil if no
changes in respect of the composition of C fractions have occurred, given the low C
contents in subsoils of soils not affected by burial processes (Table 1).
Hence, decreasing amounts of only certain fractions must be related to the
decomposition of C within these fractions. Furthermore, comparing this
decrease to C contents associated with other fractions allows assessment of
information on the effectiveness of protection through a specific set of
stabilization mechanisms after burial. A recently published study on SOC
stabilization along a 500 year chronosequence from wetland to cropland in
China illustrates the importance of understanding the interaction of
stabilization mechanisms for SOC by pointing out that “beyond providing
a physical barrier between microbes and carbon substrates, soil aggregation
could contribute to SOC stabilization by bringing organic matter and soil
particles together and promoting organo-mineral interaction” (Cui et al.,
2014).
Contradicting the findings of X. Wang et al. (2014), our data
indicate the strengthening effect of physical protection/aggregation on the
stabilization of C against decomposition. While the relative contribution of
C associated with the non-aggregated silt and clay fraction is increasing
with depth (Table 1), its age compared to C associated with aggregates is
becoming younger (Table 5). We conclude from this that non-aggregated C is
not as stable against decomposition as is often assumed (Yu et al., 2011).
Within aggregates, on the other hand, mineral surfaces are “locked away”
from new C input similarly to how aggregated C is not as easily
accessible for decomposers as non-aggregated C and can hence result in a
slower turnover of this C. Instead, mineral surfaces of the non-aggregated
silt and clay C can act very efficiently in stabilizing younger C
(Eusterhues et al., 2005, 2008; Duemig et al., 2011, 2012), resulting in
younger radiocarbon ages and at the same time similar C loading as for
microaggregates (Table 5). However, the amount of macroaggregates and microaggregates
is steadily decreasing with depth, indicating (i) the ongoing
weathering of soils and hence alterations of minerals (Fig. 6) that build up
the aggregates and (ii) that the slow but steady decomposition of C leads to a
deterioration of aggregates and, hence, this mechanism of protecting C
through physical isolation.