Minerals, particularly clay-sized minerals, protect soil
organic matter (SOM) from decomposition by microorganisms. Here we report
the characterization of SOM and the associated minerals over decades of
biodegradation, in a French long-term bare fallow (LTBF) experiment started
in 1928. The amounts of carbon (C) and nitrogen (N) in the study area declined over time for
six fractions (sand, coarse silt, fine silt, coarse clays, intermediate
clays, and fine clays). The
Soils represent an important carbon reservoir on the global scale: they store 3 times more carbon than the atmosphere (Batjes, 1996) and are currently considered as one of the solutions for climate change mitigation and adaptation in addition to food security as highlighted by the “4 per 1000” initiative (Soussana et al., 2017). Soil organic matter (SOM) encompasses compounds with residence times ranging from days to millennia (Trumbore, 2000), and the mechanisms controlling SOM turnover are actively debated (Dungait et al., 2012; Lehmann and Kleber, 2015; Schmidt et al., 2011).
Except to some extent for pyrogenic C (Lutfalla et al., 2017), the current thinking considers organomineral interactions (by adsorption or coprecipitation) as a dominant factor, rather than intrinsic chemical recalcitrance, for the long-term persistence of otherwise labile organic compounds in soil (Baldock and Skjemstad, 2000; Balesdent et al., 2000; Lehmann and Kleber, 2015; von Lützow and Kögel-Knabner, 2010; Schmidt et al., 2011; Sollins et al., 2006). However, the exact mineralogical nature of soils remains barely documented, especially at the submicrometer scale, and in particular for soils where the clay fraction is dominated by phyllosilicates (Barré et al., 2014).
Soil phyllosilicates are very diverse (Wilson, 1999). Among phyllosilicates, smectites are considered to have higher protective capabilities than illite and kaolinite, because of their higher specific surface area and cation exchange capacities (Bruun et al., 2010). Nonetheless, in situ experimental demonstrations are still lacking (Barré et al., 2014). To date, the influence of phyllosilicate mineralogy on the chemical composition of persistent SOM has only been documented in model systems (Mikutta et al., 2007).
Here, we studied samples from long-term bare fallow (LTBF) experiments. These vegetation-free experimental plots offer the unique opportunity to naturally concentrate persistent SOM (Barré et al., 2010; Rühlmann, 1999): with time, biodegradation occurs in the LTBF plots and the carbon content gradually decreases as there is no input of fresh organic carbon.
We used quadruplicate LTBF soil samples collected in 1929, 1939, 1951, 1981,
and 2008 and fractionated into six particle-size classes” sand (> 50
The present contribution addresses the following fundamental questions:
How much pluri-decadal persistent SOM is in the different
fractions? What is the chemical nature of pluri-decadal persistent SOM
in the different fractions? What is the long-term protective
capabilities of the different phyllosilicates?
We used archived samples of the “
We subjected soil samples to physical dispersion and particle-size
fractionation following a published protocol
(Balesdent et al.,
1998, Fernández-Ugalde et al., 2016) to separate the
following fractions: sand fraction (> 50
Approximately 50 g of soil was shaken overnight with 20 glass beads in 180 mL
of deionized water to break aggregates bigger than 50
Clay subsamples (2 g) were suspended in water and sonicated at 320 J mL
The mineral compositions of the total clay fraction and of the clay
subfractions (CC, IC, and FC) were determined by X-ray diffraction (XRD)
analyses. Air-dried, oriented deposits prepared using the filter transfer
method (Moore and Reynolds, 1989) were analyzed with a Cu
K
The total organic carbon (C) and nitrogen (N) contents of all particle-size fractions were measured by dry combustion in a CHN autoanalyzer (Carlo Erba NA 1500). The total C content is equivalent to the total organic carbon content, as the investigated soil samples do not contain carbonates.
In the present study, synchrotron-based C-NEXAFS data were collected using beamlines located at the Canadian Light Source (CLS, Canada) where the storage ring is operated at 2.9 GeV and at a current between 250 and 150 mA.
The “bulk” carbon speciation of clay subfractions was investigated by NEXAFS spectroscopy using the CLS beamline 11-ID-1 Spherical Grating Monochromator (SGM; Regier et al., 2007). See Supplement for details regarding the method. Each spectrum reported in the present study corresponds to an average of about 50 measurements. Of note, only the first 250 nm of the sample surface are probed using the SGM setup. Spectra were averaged, background subtracted, and normalized using the Igor Pro software.
STXM-based NEXAFS data were collected using the CLS beamline 10ID-1 (SM
beamline, Kaznatcheev et al., 2007),
which operates in the soft X-ray energy range (130–2500 eV) using an
elliptically polarized undulator. See the Supplement for details
regarding the method. All clay subfractions were analyzed for each sampling
time, i.e., 15 samples were analyzed. We first analyzed each sample at the
millimeter scale to identify carbon bearing regions of a few square
micrometers. We then collected STXM-based NEXAFS data over the 250–450 eV
energy range covering the C K-edge (280–295 eV), the K L
To obtain a more “quantitative” insight on the evolution of the molecular signatures of the investigated experimental samples with increasing bare fallow duration and to be able to compare the spectra, we performed (i) background subtraction, (ii) normalization to the carbon amount, and (iii) fit using Gaussian functions placed at fixed positions (e.g., 284.4 eV, quinones; 285 and 285.4 eV, aromatic; 285.8 eV, imines; 286.2, 286.6, and 287.1 eV, carbonyls; 287.7 eV, aliphatics; 288.2 eV, amides; 288.6 eV, carboxylic; 289.1 eV, aldehydes; 289.4 eV, hydroxyls; 289.9 eV, aliphatics; 290.3 eV, carbonates) (Bernard et al., 2012; Le Guillou et al., 2014, 2018).
Statistical analyses were conducted using the R free software environment for
statistical computing (
Initial C and N concentrations were low and very low in the sand and coarse
silt fractions, respectively, but much higher in the fine silt and clay
fractions (Table 1). The IC and CC subfractions displayed an initial OC
content (
Evolution of carbon content
(mgC g
Carbon and nitrogen contents and corresponding losses of organic carbon and nitrogen in each of the fractions on the first and last sampling dates. Standard deviations are indicated in parentheses. Data for bulk soil and for clay fractions are in bold as they will be discussed in length throughout the paper.
Evolution of the carbon to nitrogen ratio (
The XRD patterns of the three clay subfractions of a bare fallow sample
(plot 21) collected at
X-ray diffraction patterns of particle-size fractions in
one of the replicated plots of the LTBF experiment sampled at the beginning of the BF treatment. The
black line corresponds to air-dried preparations, and the grey line corresponds
to glycolated preparations.
The “bulk” C-NEXAFS spectra of the clay subfractions were very similar, with spectral features attributed to four main chemical moeities (Fig. 4): aromatic or olefin carbons (peak between 285 and 285.5 eV), carbonyl groups (shoulder between 286.1 and 287.1 eV), aliphatic carbons (shoulder at 287.7 eV), and carboxylic groups (intense peak at 288.6 eV). The peak at 290.3 eV is attributed to carbonates (Bernard et al., 2015). The features at approximately 283–284 eV are artefact features created by the presence of minerals that significantly absorb the beam throughout the carbon absorption energy range. Deconvolution of the data allowed for the semiquantitative estimation of the relative concentration of the four functional groups described above (aromatics/olefinics, carbonyls, aliphatics, and carboxylics) as a function of time and revealed no major evolution except for a slight but statistically significant increase in carboxylic moeities contained in the CC subfraction and a slight but statistically significant decrease in aliphatics in the FC subfraction (Fig. S3 in the Supplement).
Normalized C-NEXAFS spectra of the fine clay (FC) subfraction
of samples collected at different times, from
Four different types of assemblages were observed in the clay subfractions
(Fig. 5): (1) SOM-poor K-rich minerals, for which the abundance increased with
bare fallow duration; (2) organomineral complexes rich in C, N, and K; (3) organomineral complexes rich in C,
N, K, and Ca; and (4) K-poor particulate
OM. Of note, pure OM without any signal from K minerals was not identified
in the present samples (a possible explanation could be that the signal
corresponds to the average of at least 10 pixels of 40 nm
Selected NEXAFS spectra of the four types of particles identified in the clay subfractions by STXM. Individual spectra are shifted on the intensity axis for better discrimination. Unlabeled peaks at 297.1 and 299.7 eV correspond to potassium (K) L-edge peaks. The spectra correspond to OM-rich particles with very little mineral (darkest color), organomineral particles with K and Ca mineral phases, organomineral particles with K-minerals, and K-rich phases (lightest color).
Over the course of the bare fallow treatment, CC subfractions displayed particles
ranging from isolated OM (particulate OM) to mineral-rich, OM-bearing
particles (Figs. 5 and 6). After 22 years of bare fallow conditions, mineral
particles exempt of OM appeared in the CC subfractions.Conversely,
mineral particles exempt of OM and OM-rich particles
were virtually undetected in the IC and FC subfractions. These
subfractions exhibited a homogeneous signal similar to the spectra of
OM
STXM-NEXAFS compositional maps of organomineral particles
contained in the fine clay
After 79 years of bare fallow conditions, SOM remained in all of the soil fractions. Apart from the coarse silt fraction which contained very low amounts of SOM, the amount of OC and N remaining after 79 years of bare fallow conditions were, as expected, higher in the clay fractions (Table 1). Of note, the relatively high amount of C remaining in the sand fractions could be explained by the presence of pyrogenic carbon in these fractions (Table 1, Fig. 2).
The percentage of C and N remaining increased with decreasing particle-size for fine silt, CC, and IC; however, the same was not observed for FC, where much lower percentages of C and N remained, suggesting a higher labile SOM content in the FC fraction. This is likely due to the fractionation procedure which can favor the accumulation of labile dissolved OM in the FC fraction (Laird et al., 2001).
Although it was not possible to determine the nitrogen speciation, our
results show that persistent SOM associated with clays is highly enriched in
N. Indeed, particulate OM with high
NEXAFS showed no major shift in the chemistry of SOM after several decades of biodegradation under bare fallow conditions besides a slight increase in carboxylic moeities in the CC subfraction and a slight decrease in aliphatics in the FC subfraction, supporting the fact that persistent forms of carbon are slightly more oxidized than the initial forms of carbon (von Lützow and Kögel-Knabner, 2010). The present study also confirmed that persistent SOM is mainly composed of microbial material: on average, all spectra displayed typical patterns of SOM strongly enriched in microbial material (Keiluweit et al., 2012; Kleber et al., 2011).
Particulate OM could still be observed in the CC subfraction even after 79 years of bare fallow conditions. The NEXAFS spectra of these particles highlighted their polyphenolic nature (Keiluweit et al., 2010), suggesting that these were pieces of lignin-rich plant debris, possibly physically protected in submicron aggregates as shown in similar temperate Luvisols (Chenu and Plante, 2006). Thus, the present results show that pluri-decadal persistent SOM is made of N-rich oxidized SOM adsorbed to mineral surfaces and, to a lesser extent, of particulate OM, in agreement with the recently proposed “Soil Continuum Model (SCM)” (Lehmann and Kleber, 2015).
A simple mixing model allows for a rough estimation of the amount of C associated
with minerals in the CC subfractions. Assuming a
Solving the equation leads to
After 79 years of bare fallow conditions, the FC and IC subfractions, mostly composed of swelling clays, had a higher N content than the CC subfraction which contained both illites and swelling clays (Fig. 3). The CC and IC subfractions had a similar apparent C content after 79 years of bare fallow conditions, which was lower than that of the FC subfraction (Table 1). However, as seen above, TOC (total organic carbon) contents might be misleading in this particular case as CC contains significant amounts of POM (particulate organic matter) (estimated to be around 40 % of TOC in this fraction). Therefore, if we focus on C associated with clay minerals by organomineral interactions (either through adsorption or coprecipitation), we find that at all times, there is more C associated with IC and FC than to CC. These results may be due to the predominance of swelling clays in IC and FC or to the fact that finer clay minerals have a higher specific surface area and can therefore interact with more SOM. These results also suggest that swelling clays may better protect N-rich SOM in particular.
Spatially resolved observations at the submicron scale with STXM-NEXAFS clearly showed that mineralogy influences SOM stabilization. Indeed, several illite particles (identified by the presence of K and the absence of Ca) were devoid of OM (Figs. 5 and 6), and the relative abundance of OM-depleted illites increased over time. Conversely, smectite layers (identified by the presence of Ca, as interstratified illite/smectite in the presence of K or pure smectite in the absence of K) were always found in association with OM over the chronosequence. This demonstrates that among phyllosilicates, smectites might have higher SOM protective capabilities than illites. Our results do not allow us to make conclusions regarding the protective capability of kaolinite; thus, additional experiments are needed to confirm the present findings and investigate clays with other mineralogies.
The suggested higher capacity of smectites and mixed-layer
illite/smectite to protect OM compared to illites could be due to the
presence of calcium which might facilitate the formation of persistent
bounds between clay surfaces and OM. Organomineral assemblages rely on
different physicochemical interactions, depending on the chemical nature of
the OM and of the mineral phase, which include covalent bonding, ligand
exchange, or weaker interactions such as Van der Waals for instance
(Barré et
al., 2014; Kögel-Knabner and Amelung, 2014). Similarly to previous
studies (Chen et al., 2014), we observed a
colocation of C and Ca (contained in smectites). Although the exact
mechanism responsible for this colocation is not clearly identified, one
hypothesis is that Ca could facilitate the binding of negatively charged or
polarized organic compounds to negatively charged mineral surfaces via
cation bridging
(Lützow
et al., 2006; Mikutta et al., 2007; Rowley et al.,
2018).
It has been shown that the cation bridging mechanism can promote
organomineral interaction and reduce the bioavailability of adsorbed organic
molecules for negatively charged
Underlying research data are available upon request from the corresponding author.
Materials and methods on C-NEXAFS spectroscopy at the “bulk” scale,
STXM-based NEXAFS spectroscopy, and the C-NEXAFS data deconvolution procedure.
Additional figures of normalized NEXAFS spectra of coarse and intermediate
clay subfractions from all sampling dates (Figs. S1 and S2) and
corresponding deconvolution (Figs. S3 and S4). The supplement related to this article is available online at:
SL, PB and CC designed the study. SL prepared the samples and carried out the analyses with technical help from all co-authors. All authors participated in the discussion of the results and the paper preparation.
The authors declare that they have no conflict of interest.
The INSU EC2CO program is acknowledged for financial support. NEXAFS data were acquired at the beamline 11ID-1 at the CLS, which is supported by the NSERC, the CIHR, the NRC, and the University of Saskatchewan. Special thanks go to Jian Wang and Jay Dynes for their expert support of the STXM at the CLS and to Tom Regier and Adam Gillespie for their expert support on the SGM-beamline at CLS.
This paper was edited by Yakov Kuzyakov and reviewed by two anonymous referees.