Permafrost-affected soils of the Arctic account for 70 %
or 727 Pg of the soil organic carbon (C) stored in the
northern circumpolar permafrost region and therefore play a major role
in the global C cycle. Most studies on the budgeting of C storage and
the quality of soil organic matter (OM; SOM) in the northern circumpolar
region focus on bulk soils. Thus, although there is a plethora of
assumptions regarding differences in terms of C turnover or stability, little knowledge is available on the mechanisms stabilizing
organic C in Arctic soils besides impaired decomposition due to low
temperatures. To gain such knowledge, we investigated soils from
Samoylov Island in the Lena River delta with respect to the
composition and distribution of organic C among differently stabilized
SOM fractions. The soils were fractionated according to density and
particle size to obtain differently stabilized SOM fractions differing
in chemical composition and thus bioavailability. To better understand
the chemical alterations from plant-derived organic particles in these
soils rich in fibrous plant residues to mineral-associated SOM, we
analyzed the elemental, isotopic and chemical composition of
particulate OM (POM) and clay-sized mineral-associated OM (MAOM). We
demonstrate that the SOM fractions that contribute with about
17 kgCm-3 for more than 60 % of the C stock are
highly bioavailable and that most of this labile C can be assumed to
be prone to mineralization under warming conditions. Thus, the amount
of relatively stable, small occluded POM and clay-sized MAOM that
currently accounts with about 10 kgCm-3 for about
40 % of the C stock will most probably be crucial for the quantity
of C protected from mineralization in these Arctic soils in a warmer
future. Using δ15N as a proxy for nitrogen (N) balances
indicated an important role of N inputs by biological N fixation,
while gaseous N losses appeared less important. However, this could
change, as with about 0.4 kgNm-3 one third of the N is
present in bioavailable SOM fractions, which could lead to increases
in mineral N cycling and associated N losses under global warming.
Our results highlight the vulnerability of SOM in Arctic
permafrost-affected soils under rising temperatures, potentially
leading to unparalleled greenhouse gas emissions from these soils.
Introduction
For several millennia, organic matter (OM) accrued in the remote soils
of the Arctic, and only recently have researchers started to increasingly
understand the importance of these cold soils for the global carbon
(C) cycle and, thus, global climate (Ping et al., 2015). Estimates on
the northern circumpolar soil organic carbon (SOC) stock within the
first meter vary between 445 and 496 Pg (Tarnocai et al.,
2009; Hugelius et al., 2014). These C-rich soils are changing from a C
sink to a source due to global warming (Oechel et al., 1993;
Parmentier et al., 2017). The Arctic is strongly affected by climate
change with an increase in surface temperatures during the last 2 decades that has been more than twice as high as the global average
(Meredith et al., 2019). In a warming Arctic, C is lost both via
carbon dioxide and methane emissions and by lateral transport with
water (Plaza et al., 2019). The C that is released from
permafrost-affected soils due to anthropogenically accelerated thawing
is assumed to further enhance global warming and thus trigger
additional C release from permafrost, a phenomenon known as permafrost
C feedback (Davidson and Janssens, 2006; Schuur et al., 2015).
An analysis of soils from 10 North American ecosystem types reaching
from tropical forests to Arctic tundra demonstrated a pronounced
longer turnover time for soil organic matter (SOM) in cold regions in
comparison to other climate regions, as the C stabilization mechanisms
clearly differ (Frank et al., 2012). In temperate soils, the main
drivers for SOC sequestration are spatial inaccessibility (occlusion
in soil aggregates), binding to mineral particles (organo-mineral
associated OM) and intrinsic chemical recalcitrance of the OM itself
(Six et al., 2002; von Lützow et al., 2006). Besides these
specific mechanisms, environmental factors like waterlogging and low
temperatures inhibit the turnover of OM in cold regions (Oades, 1988),
with cryoturbation additionally supporting the conservation of SOM at
greater soil depth and thus in the permafrost (Kaiser et al.,
2007). These abiotic mechanisms fail as soon as permafrost collapses,
which leads to an increased decomposition of OM (Turetsky, 2004; Plaza
et al., 2019).
Already in 1982, Post et al. recognized a considerable variability in
C stocks in tundra soils, which illustrates that gaining more knowledge about the biogeochemical cycling of C in permafrost soil needs
to involve more analytical approaches that enable the assessment of possible
mechanisms of C stabilization. Thus, besides the quantification of
organic C (OC), there is a growing number of studies aiming to
elucidate the chemical composition of SOM and the processes and
mechanisms involved in C cycling and stabilization in
permafrost-affected soils (inter alia Torn et al., 2013; Mueller et al.,
2015; Strauss et al., 2017; Jongejans et al., 2018; Kuhry et al.,
2020).
With ongoing warming, the active layers in cold regions deepen, and
thus, microbial activity changes and the accessibility and
bioavailability of OM in hitherto frozen soil layers increase
(Mackelprang et al., 2011; Hultman et al., 2015). The depolymerization and
ammonification as well as nitrification of the long-sequestered
organic nitrogen (N) might also enhance mineral N availability in
these permafrost-affected soils, leading to increased emissions of the
highly potent greenhouse gas nitrous oxide (Elberling et al., 2010;
Wilkerson et al., 2019). The importance of mechanisms restricting SOM
decomposition in permafrost soils will possibly shift from climatic
stabilization (Schmidt et al., 2011) to spatial inaccessibility and
association with minerals (Harden et al., 2012; Mueller et al., 2015)
with widely unknown consequences for the C stored in these soils.
Several studies estimated the vulnerability of C in permafrost soils
to microbial decay from the chemical composition of bulk SOM (inter alia Herndon et al., 2015; Strauss et al., 2017; Tesi et al., 2016; Weiss
and Kaal, 2018; Wild et al., 2016; Xue et al., 2016; Zimov et al.,
2006). Yet, as SOM represents a continuum of a range of materials of
different compositions, from fresh plant litter to highly altered
compounds (Lehmann and Kleber, 2015) ruled by different stabilization
regimes, the investigation of bulk SOM alone is insufficient. The use
of more sophisticated approaches, separating SOM into different
fractions, allows for a more detailed understanding of the
stabilization mechanisms in soil (Golchin et al., 1994). So far, only
few studies (inter alia Dao et al., 2018; Diochon et al., 2013; Dutta
et al., 2006; Gentsch et al., 2015; Höfle et al., 2013; Mueller
et al., 2015; Xu et al., 2009) used fractionation approaches to
investigate the distribution and composition of OM pools in
permafrost-affected soils (Ping et al., 2015), most of them focusing
on the composition of specific fractions or using incubation
experiments.
The objective of our study is to gain detailed insights into the
chemical composition and stabilization mechanisms of SOM in Cryosols
from the Siberian Lena River delta under present conditions;
therefore, we used a physical fractionation approach to separate light
organic particles and OM associated with minerals, i.e., particulate
OM (POM; dominated by bits and pieces of plants and to a lesser extent
microbial residues) and mineral-associated OM (MAOM). As it is known
from temperate soils that POM and MAOM have different ecological
functions and contribute differently to C and N storage and cycling,
we expect that also in permafrost-affected soils, these soil C and N
pools show marked differences in their chemical composition and thus
vulnerability to climate change. Our two major hypotheses for the
present work are that (1) SOM in permafrost-affected soils is mainly
stored as POM resulting from a restricted decomposition due to
climatic stabilization and that (2) larger POM is characterized by a high
content of rather labile OM that mirrors the plant litter input,
whereas smaller POM particles and MAOM resemble microbial transformed
OM, independent of the original plant litter.
MethodsSite characteristics and soil sampling
Samoylov Island (72∘22′ N, 126∘30′ E) is located
in one of the main channels of the Siberian Lena River delta, the
largest delta of the Arctic. The island developed during the Holocene
and belongs to one of three river terraces. While the western third of
the island consists of an active floodplain, the eastern part is
covered by ice-wedge polygonal tundra that is typical for this terrace
(Boike et al., 2013). Located at 10 to 16 m a.s.l., the
Holocene river terrace is rarely flooded, and its plant cover
represents the characteristic wet sedge tundra vegetation (Zubrzycki
et al., 2013). The dominant soils of the terrace are Cryosols
according to the World Reference Base for Soil Resources (Zubrzycki
et al., 2013; IUSS Working Group WRB, 2014) and Orthels and Turbels
according to the US Department of Agriculture soil taxonomy (Soil Survey Staff, 2014). This
terrace has recently been reported to be covered by about 40 %
non-degraded polygonal tundra; 40 % collapsed polygons, slopes
and water bodies; and 20 % of polygons that show different stages
of degradation (Kartoziia, 2019). On the island, active-layer
thickness varies around 50 cm, and the thawing period lasts
approximately 129 days (Boike et al., 2013). The climate is Arctic, and
the 30-year mean (1961–1990) of the closest meteorological station in
Tiksi, about 110 km southeast, shows a mean annual air
temperature of -13.5 ∘C with a large amplitude between the
warmest (around 8 ∘C in July and August) and coldest (around -32 ∘C in January) months (Roshydromet,
2019). Precipitation is low on Samoylov Island, and, due to the
different geographic setting within the river delta, it has a mean of
125 mma-1, markedly lower than the 323 mma-1
measured in Tiksi (Boike et al., 2013; Roshydromet, 2019).
In this aerial image of Samoylov Island, the separation between
the floodplain in the west (with the white unvegetated sandy sediment) and
the Holocene terrace in the eastern part (with blue–gray spots indicating
shallow water and larger water bodies) is shown. Red crosses indicate the sampling
sites, and the identification numbers of the cores are given (Boike et al.,
2012).
We drilled four intact soil cores from ice-wedge polygon centers on
the Holocene river terrace (Fig. 1; Boike et al.,
2012) in April 2011
and May 2013 using a Snow, Ice and Permafrost Research Establishment coring
auger (Jon's Machine Shop, Fairbanks, AK, USA) with a length of 1 m
and a diameter of 76 mm with a STIHL BT 121 engine power head
(Andreas Stihl AG and Co. KG, Waiblingen, Germany). A detailed
description of the study area and the sampling of the soil cores can
be found in Zubrzycki et al. (2013) and Zubrzycki (2013).
All bulk soil samples were slightly acidic with lowest pH values of 4.9
and highest of 6.6; electric conductivity (EC) ranged from 66 to
240 µScm-1 with a mean of
115 µScm-1, and bulk density varied from 0.2 to
0.9 gcm-3 around a mean of 0.5 gcm-3.
Geochemical properties of bulk soils, physical soil fractionation and
chemical analyses of fractions
We separated the drilled cores according to visible mineral soil
horizons in a frozen condition and subsequently thawed and dried them at
40 ∘C in an oven. Our analyses focused on 23 selected
layers only, as shown in Table S1 in the Supplement.
The bulk soils were fractionated according to density and
particle size, following the approach described by Mueller and
Koegel-Knabner (2009). Due to the high amount of fibrous material in
these Cryosols, some modifications of the procedure were necessary to
yield mechanistically different SOM fractions. We unclenched 15 to
20 g – depending on the available amount of sample material
– of each soil sample with forceps and gently saturated them with
a sodium polytungstate solution with a density of
1.8 gcm-3 by slowly adding the salt solution with
a pipette. After 12 h, to ensure a complete and gentle
saturation, the floating free POM (fPOM; not embedded in stable
aggregates; cf. Golchin et al., 1994) was collected using a vacuum
system. The removal of the floating fPOM was repeated twice to ensure
a high recovery, and the obtained fraction was subsequently washed over
a sieve of 20 µm mesh size to remove excessive salt. Due
to the highly fibrous nature of the fPOM, the washing step also
yielded fine mineral particles, which adhered to the fPOM fibers. As
the C and N contents and C/N ratios of this mineral material were in
the exact same range of the clay-sized MAOM fraction, we added it
mathematically to this fraction for the calculation of the C stock. To
separate occluded POM fractions (oPOM; incorporated in water-stable
aggregates; cf. Golchin et al., 1994) from MAOM, the residual
samples were subjected to ultrasonication (Bandelin Sonopuls HD 2200,
Berlin, Germany) using a calibrated (Graf-Rosenfellner et al., 2018)
energy input of 300 JmL-1 after the fPOM removal. On the
lines of the fPOM fractions, oPOM was withdrawn using a vacuum system
and washed salt-free over a sieve of 20 µm mesh size by
repeated washing until the EC dropped below
2 µScm-1. During the washing of the oPOM through
a 20 µm sieve, we obtained the small oPOM (oPOMs) fraction
representing a fine particulate light OM (Mueller et al., 2015,
2017). The remaining heavy residues, constituting the MAOM, were
separated by wet sieving and sedimentation to obtain coarse and medium
sand (>200µm), fine sand (63–200 µm),
coarse silt (20–63 µm), medium silt
(6.3–20 µm), and fine silt and clay (<6.3µm;
further referred to as the clay-sized MAOM fraction). All SOM
fractions were analyzed for total C and N contents in duplicate by dry
combustion (EuroVector EuroEA3000 Elemental Analyzer,
Pavia, Italy). After the analyses of each sample, for better clarity
for the reader, C and N contents were calculated for the combined
sand- and silt-sized fraction per each bulk soil sample. Due to the
absence of carbonates (see pH values in Table S1), total C represents
OC. Coarse fractions >20µm were ball-milled and
homogenized prior to C and N measurements. The bulk soil C and N
contents were calculated from the sum of the physical fractions; C and
N stocks for the SOM fractions were also calculated, and overall C and
N stocks were projected to 1 m soil depth. The mass recovery rate
after fractionation was >90 % in all samples. In addition, to
reveal the microscale structure and illustrate possible source
materials (microbial vs. plant origin) scanning electron microscope
(SEM) images (JSM-7200F, JEOL, Freising, Germany) were obtained for
representative POM fractions.
Stable isotope measurements
The abundance of 15N and 13C POM and clay-sized MAOM fractions were determined using an isotope ratio mass spectrometer
(Delta V Advantage, Thermo Fisher, Dreieich, Germany) coupled to an
elemental analyzer (EuroEA, Eurovector, Pavia, Italy). A lab standard
(acetanilide) was used as a standard for every sequence in intervals,
and different weights as well were used to quantify the isotope linearity of the
system. The standard itself was calibrated against several suitable
international isotope standards from the International Atomic Energy
Agency (IAEA, Vienna, Austria) for both isotopes. The final correction of
isotope values was achieved with several international isotope
standards and other suitable laboratory standards that cover the range
of δ15N and δ13C results. Results are given in
delta values relative to air N2 for 15N and relative to
Vienna Pee Dee Belemnite (V-PDB) for 13C (Werner and Brand,
2001).
13C nuclear magnetic resonance spectroscopy
We subjected all fPOM, oPOM, oPOMs and selected clay-sized MAOM fractions to 13C cross-polarization magic-angle spinning (CP-MAS)
nuclear magnetic resonance (NMR) spectroscopy (Bruker DSX 200 spectrometer, Billerica, MA, USA). The
13C NMR spectra were recorded at 6800 Hz with an
acquisition time of 0.01024 s. During a contact time of
1 ms, a ramped 1H pulse was applied to avoid Hartmann–Hahn
mismatches. We executed measurements in 7 mm zirconium dioxide
rotors with a delay time of 1.0 s for large POM fractions
(fPOM and oPOM) and a reduced delay time of 0.4 s for oPOMs
and clay-sized MAOM fractions. The acquired number of scans (NS)
varied according to the examined fractions and the available sample
material. For most of the large POM fractions, an NS between 3000 and
10 000 provided sufficient signal-to-noise ratios, while most of the
oPOMs and clay-sized MAOM fractions required an NS of at least 10 000.
Tetramethylsilane was equalized with 0 ppm as a reference for
the chemical shifts. The spectra were integrated in different chemical-shift regions according to Beudert et al. (1989) with slight
adjustments according to Mueller and Koegel-Knabner (2009): -10 to
45 ppm (alkyl C), 45 to 110 ppm (O/N alkyl C), 110
to 160 ppm (aromatic C) and 160 to 220 ppm (carboxyl
C); spinning sidebands were included. Based on these integrated-shift
regions, we calculated the ratio of alkyl C and O/N alkyl C
(a / o-a ratio) as a proxy for the degree of decomposition of plant
residues according to Baldock et al. (1997). Furthermore, we
calculated the ratio of the integrated chemical-shift regions 70 to
75 ppm (O alkyl C of carbohydrates) and 52 to 57 ppm
(methoxyl C of lignin) according to Bonanomi et al. (2013), which
provides another proxy for the decomposition stage of plant residues
in relation to fresh plant source material (further referred to as
70–75 / 52–57 ratio). To translate the NMR spectra into OM compound
classes (carbohydrate, protein, lignin, lipid and carbonyl), we fitted the
NMR data using the molecular mixing model (MMM) developed by Nelson
and Baldock (Baldock et al., 2004; Nelson and Baldock, 2005). For
the MMM fitting, we utilized the following chemical-shift regions: 0
to 45, 45 to 60, 60 to 95, 95 to
110, 110 to 145, 145 to 165 and
165 to 215 ppm. We applied the five-component MMM (without
char) with N / C constraint.
Statistics
We plotted C/N ratios and C and N concentrations against the N and C
stable isotope ratios of SOM fractions using R to identify
interrelations. The R software, RStudio and the packages “Rcmdr” (with
the plugin “FactoMineR”), “Hmisc”, “Factoshiny” and “corrplot” were used for
principal component analysis (PCA), correlation matrices and the
compilation of plots (Lê et al., 2008; RStudio Team, 2016; R
Development Core Team, 2017). We used PCA and correlation matrices to
find correlations between the properties of different SOM fractions
(fPOM, oPOM, oPOMs and clay-sized MAOM), namely C and N contents,
decomposition proxies (C/N ratio of bulk soils and of SOM fractions,
a / o-a ratio and 70–75 / 52–57 ratio), stable isotopes and the results
from the MMM.
ResultsBiogeochemical bulk soil properties and distribution of SOM fractions
The bulk soil C contents over all cores and depth layers varied
between 31.6 and 144.0 mgg-1. The content of total N
ranged from 1.3 to 6.8 mgg-1 for all cores and depth
layers. While the C/N ratios ranged between 23 and 38 in three of
the four cores, the values of the bulk soils of the fourth core were
markedly lower (Table S1). The soil C stocks (projected to 1 m
soil depth) ranged between 20.4 and 31.4 kgCm-3 with
a mean of 27.5±11.8kgCm-3; the N stocks varied
between 0.7 and 1.9 kgNm-3 with a mean of 1.2±0.6kgNm-3 (Table 1).
C and N stocks (projected to 1 m soil depth) and C/N ratios of the
SOM fractions. Given are mean values and the standard deviation of free
(fPOM), occluded (oPOM) and small occluded (oPOMs) particulate organic
matter and of different-sized mineral-associated organic matter (MAOM).
SOM fractionC stockN stockC/N ratiokgCm-3kgNm-3fPOM14.0±4.60.3±0.146±16oPOM2.6±1.10.1±0.151±22oPOMs5.4±3.50.3±0.217±3Clay-sized MAOM4.6±2.20.4±0.213±1Silt-sized MAOM0.7±0.40.1±0.010±1Sand-sized MAOM0.2±0.10.0±0.010±3Sum27.5±11.91.2±0.6
The mass distribution of POM fractions varied throughout all depth
layers with proportions between 10.6 and 295.0 mgg-1
(fPOM), between 3.0 and 71.7 mgg-1 (oPOM), and between 3.9
and 267.2 mgg-1 (oPOMs). In particular, core 3 and 4 showed
larger amounts of fPOM and oPOM material at greater depth in between
layers dominated by MAOM (Table S1). The MAOM fractions ranged
between 37.2 and 244.5 mgg-1 (clay-sized), between 182.4
and 479.3 mgg-1 (silt-sized), and between 79.0 and
591.5 mgg-1 (sand-sized).
Elemental composition of SOM fractions
The highest C contents were detected in the fPOM and oPOM fractions,
with values ranging from 196.3 to 425.5 mgg-1C for the
fPOM and from 368.4 to 449.1 mgg-1C for the oPOM
fractions. Due to the highly fibrous structure of these Cryosols rich
in plant residues, fractionation was challenging for some of the
samples, leading to one outlier within the fPOM fractions and four
outliers within the oPOM fractions. We defined outliers as the
measurements laying outside the boxplots' whiskers, thus values lower
than 1.5 times the interquartile range below the lower quartile and
values higher than 1.5 times the interquartile range above the higher
quartile. We excluded these fractions from further calculations, as we
assume that they point to mineral particles, which we were not able to
separate fully from the very fibrous POM structures. The C content of
the oPOMs fractions ranged between 61.4 and
344.8 mgg-1C, and the C contents of the clay-sized MAOM fractions ranged between 51.5 and 117.9 mgg-1C, while silt- and sand-sized MAOM fractions showed the lowest C contents with 2.5 to
11.1 mgg-1C and 0.7 to 3.1 mgg-1C,
respectively (Fig. 2a).
The content (in mg g-1) of C (a) and N (b) of bulk soils (I) and SOM fractions
(free particulate OM – fPOM, occluded particulate OM – oPOM, small occluded
particulate OM – oPOMs – and clay-sized mineral-associated OM – MAOM) (II).
Results for the N content were 5.0 to 19.5 mgg-1N for
fPOM fractions, 3.4 to 23.7 mgg-1N for oPOM fractions
and slightly higher for oPOMs fractions with 4.6 to
26.4 mgg-1N. The N contents of the clay-sized MAOM fractions ranged between 3.8 and 10.1 mgg-1N, while silt- and sand-sized MAOM fractions contained markedly less N
(Fig. 2b). Large POM fractions (fPOM and oPOM) showed a wide variation
of C/N ratios with values between 22 and 76 for fPOM and between 18
and 113 for oPOM. The values of the oPOMs fractions were clearly lower
and had less variability with 13 to 25, while clay-sized MAOM fractions ranged between 11 and 16. The lowest C/N ratios were present
in silt- and sand-sized MAOM fractions with 8 to 12 and 6 to 19,
respectively. Large POM fractions had not only the widest C/N ratios
compared to oPOMs and mineral-associated OM within each soil
layer but also showed the largest variation (Fig. 3; Table S2).
C/N ratios of bulk soils (I) and SOM fractions (free particulate
OM – fPOM, occluded particulate OM – oPOM, small occluded particulate OM
– oPOMs – and clay-sized mineral-associated OM – MAOM) (II).
The contribution of C and N weighted for the amount of each specific
SOM fraction per soil layer showed a great variance in the amount of C
and N stored either as POM or MAOM. For C, this ranged between 211.5
and 807.0 mgC per gram of bulk soil for the large POM fractions
(fPOM and oPOM), between 13.7 and 479.7 mgC per gram of bulk soil
for oPOMs, whereas the clay-sized MAOM ranged between 59.4 and
431.4 mgC per gram of bulk soil (Table S2).
Over all analyzed soil layers, POM fractions accounted for 80 % of
the C stock (22.0±9.2kgCm-3), while the MAOM
fractions accounted for about 20 % (5.5±2.7kgCm-3). Overall, the fPOM fractions dominated the
C stock, with 14.0±4.6kgCm-3 representing about
half of the total C stock of all analyzed cores and layers. The
occluded POM fractions contributed less with 2.6±1.1kgCm-3 (oPOM) and 5.4±3.5kgCm-3 (oPOMs). The share of the clay-sized MAOM fractions in the C stock was 4.6±2.2kgCm-3, while
silt- and sand-sized MAOM fractions played only a subordinate role
(Table 1).
For the N stock, the contribution of the POM fractions sums up to
about 60 % (0.7±0.4kgNm-3) and that of the
MAOM fractions to about 40 % (0.5±0.2kgNm-3). The fPOM and oPOM fractions contributed
differently to the stock with 0.3±0.1kgNm-3 and
0.1±0.1kgNm-3, respectively. The oPOMs and
clay-sized MAOM fractions added similarly to the N stock with 0.3±0.2kgNm-3 and 0.4±0.2kgNm-3
but also showed the largest variation. Similar to C stocks, silt- and
sand-sized MAOM fractions had a negligible share in the N stocks
(Table 1).
Although overall the soil C and N storage was dominated by POM, the
distribution of POM- vs. MAOM-related C and N varied greatly with
depth, with some soil layers showing a dominance of MAOM for C and N
storage (Table S2).
Isotopic composition of SOM fractions
For POM and clay-sized MAOM fractions, we analyzed the content of
stable carbon (13C) and nitrogen (15N) isotopes. With
respect to δ15N, the values differed little between all
examined fractions: fPOM (-0.3 ‰ to 1.4 ‰), oPOM (0.2 ‰ to
2.4 ‰), oPOMs (0.0 ‰ to 2.9 ‰) and clay-sized MAOM
(-0.4 ‰ to 3.4 ‰), with the latter showing the
highest values (Table S2). With decreasing C/N ratios, a clear trend
towards more negative δ13C and lower δ15N values
was demonstrated for all POM fractions (Fig. 4). As shown by PCA
(Fig. 5), δ15N and δ13C showed positive
dependencies with the C/N ratios. As the deeper soil layers of core
4 were clearly dominated by MAOM with a narrow C/N ratio, the
overall δ15N (0.7 ‰) was lower compared to the
other three cores.
Natural abundance of δ13C and δ15N
plotted against the C/N ratios and the δ15N values plotted
against the N content of SOM fractions (free particulate OM – fPOM, occluded
particulate OM – oPOM, small occluded particulate OM – oPOMs – and clay-sized
mineral-associated OM – MAOM): the values of δ13C
(‰ relative to V-PDB) (a) and the values of δ15N (‰ relative to air N2) (b) in relation to
the C/N ratio (log-converted) of SOM fractions and δ15N
(‰ relative to air N2) plotted against N content
(in mg g-1) of the SOM fractions (c).
The δ13C values were similar for all fractions and the range
of the values and their variability was similar for fPOM (-31.2 ‰ to
-25.6 ‰), oPOM (-30.6 ‰ to -25.3 ‰), oPOMs
(-31.5 ‰ to -25.0 ‰) and clay-sized MAOM (-31.8 ‰ to
-24.1 ‰; Table S2). As for δ15N, also the
δ13C values of the soil material of core 4 differed from
those in the other cores showing clearly lower values. Thus, overall
the differences between the cores were larger than the differences
between the fractions. Also for the δ13C values, a relation
to the C/N ratios of all fractions was demonstrated. The C/N
ratios of the clay-sized MAOM asymptotically approached a limit when
plotted over δ15N and δ13C, whereas the POM
fractions showed a linear increase in the isotope content at higher
C/N ratios (Fig. 4).
Principal component analysis (PCA) of δ13C
(‰ relative to V-PDB), δ15N
(‰ relative to air N2), C and N content of the SOM
fractions (free particulate OM – fPOM, occluded particulate OM – oPOM, small
occluded particulate OM – oPOMs – and clay-sized mineral-associated OM – MAOM), C/N ratio of fractions and of bulk soils, and 13C CP-MAS
NMR-derived decomposition proxies (a / o-a ratio and 70–75 / 52–57 ratio). Dim: dimension.
13C NMR – the molecular level
The 13C CP-MAS NMR spectra of all examined SOM fractions showed
dominant peaks in the region of O/N alkyl C. The spectra of both large
POM fractions were clearly dominated by the shouldered major peak
around 70 ppm and a minor peak around 105 ppm. The
integration of the spectra fortified the dominance of O/N alkyl C
with about 70 % in the fPOM (n=22) and oPOM (n=19) fractions
(Table 2). In the regions of carboxyl and alkyl C, small peaks were
present, with only a small hump being present in the region of aromatic C. The differences between the spectra of the fPOM and oPOM
fractions (see Fig. S1 in the Supplement) and in their relative composition were only
minor; even shoulders and minor side peaks were comparable in the
majority of the samples. In contrast, spectra of the oPOMs (n=23)
and clay-sized MAOM (n=10) fractions showed pronounced peaks around
30 ppm in the region of alkyl C and around 170 to 175 ppm
in the region of carboxyl C. Throughout all samples, there was a shift
from a high percentage of O/N alkyl C in the large POM fractions to
a higher percentage of aromatic and alkyl C in oPOMs and clay-sized
MAOM fractions (Table 2).
Relative chemical composition of SOM fractions obtained by 13C
CP-MAS NMR spectroscopy and decomposition proxies (a / o-a ratio and
70–75 / 52–57 ratio). Given are mean values and the standard deviation of free
(fPOM), occluded (oPOM) and small occluded (oPOMs) particulate organic
matter and of clay-sized mineral-associated organic matter (MAOM).
SOM fractionRelative chemical composition1Alkyl CO/N alkyl CAromatic CCarboxyl Ca / o-a ratio270–75 / 52–57 ratio3% fPOM13.3±5.070.2±7.611.6±2.54.9±1.90.2±0.15.6±2.1oPOM12.5±6.068.5±8.412.2±4.06.5±2.60.2±0.17.4±3.3oPOMs25.2±5.952.1±6.314.0±3.08.5±2.20.5±0.22.6±0.3Clay-sized MAOM24.0±2.649.6±3.415.1±1.811.2±3.70.5±0.12.1±0.3
1 Relative chemical composition determined by the integration of the
following chemical-shift regions: -10 to 45 ppm (alkyl C), 45 to 110 ppm
(O/N alkyl C), 110 to 160 ppm (aromatic C) and 160 to 220 ppm (carboxyl C).
2 Ratio of alkyl C and O/N alkyl C according to Baldock et al. (1997).
3 Ratio of the chemical-shift regions 70 to 75 ppm and 52 to 57 ppm
according to Bonanomi et al. (2013).
To get more differentiated information about the degree of
decomposition of the OM, we calculated the a / o-a ratio for the SOM
fractions (Baldock et al., 1997). While fPOM and oPOM fractions
revealed identically low values and relatively large standard
deviations with 0.2±0.1, oPOMs and clay-sized MAOM showed
clearly higher values with about 0.5. In addition to the a / o-a ratio,
we applied the 70–75 / 52–57 ratio (Bonanomi et al., 2013) to the
SOM fractions and received results consistent with the a / o-a ratios:
fPOM and oPOM showed high values, indicating a low degree of
decomposition, while oPOMs and clay-sized MAOM showed very low
values. With this ratio, the large POM fractions showed a considerable
variance, while the deviation within oPOMs and clay-sized MAOM was
marginal (Fig. 6). Figure 7 illustrates the close relation between the
C/N ratio of the SOM fractions and the NMR-derived decomposition
proxies.
Decomposition proxies obtained by 13C CP-MAS NMR spectroscopy
for specific SOM fractions: both the a / o-a ratio (a) and 70–75 / 52–57 ratio (b)
of SOM fractions demonstrate the similarity of large particulate OM fractions – free and occluded POM (fPOM and oPOM) – and the conjunctive
characteristics of the small occluded particulate OM (oPOMs) fractions that
links large POM fractions and the clay-sized MAOM fraction.
Relation between decomposition proxies and the C/N ratio of distinct
SOM fractions: 13C CP-MAS NMR spectroscopy derived decomposition proxies for the a / o-a ratio (a) and 70–75 / 52–57 ratio (b) vs. the C/N ratio for free
particulate OM (fPOM), occluded particulate OM (oPOM), small occluded
particulate OM (oPOMs) and clay-sized mineral-associated OM (MAOM).
By modeling the molecular composition of the SOM fractions using the
MMM (Baldock et al., 2004; Nelson and Baldock, 2005), we obtained
a clear differentiation between the large POM fractions (fPOM and oPOM)
and small oPOM and clay-sized OM separates. The composition of the
fPOM and oPOM fractions was rather similar: the percentage of
carbohydrates (about 60 %) was highest, and at the same time, the
contribution of lipids (about 8 %) was lowest in these fractions
(Table 3). Overall, the composition of both large POM fractions was
similar with slightly lower amounts of protein and slightly higher
amounts of carbonyl in oPOM compared to fPOM. The usage of the MMM
revealed once more clear differences between the large POM fractions
and oPOMs and clay-sized MAOM. The latter fractions had a lower
percentage of carbohydrates (about 40 %), whereas the percentage
of protein and lipids was markedly higher. These fractions differed
mainly in the proportion of protein and lipids, with clay-sized MAOM
containing a larger proportion of protein but a smaller proportion of
lipids (Table 3). The proportion of carbonyl was overall low with high
deviations, while the percentage of lignin was rather constant
throughout all four examined fractions.
The PCA executed on the examined fractions showed a slight correlation
between the abundance of stable isotopes and NMR-derived decomposition
proxies; yet, it confirmed the close relation between fPOM and oPOM
and the positioning of oPOMs between large POM and clay-sized MAOM fractions (Fig. 5). The separation of the large POM fractions and
oPOMs fractions provided correlation matrices with more details on the
correlations (Fig. 8). While the PCA (Fig. 5) already hinted at this, the
correlation matrices demonstrated that in the large POM fractions both
δ15N and δ13C were slightly positively correlated
with the 70–75 / 52–57 ratio and negatively correlated with the
a / o-a ratio. The positive correlation between δ13C and the
a / o-a ratio was strong in the oPOMs fractions, and the negative
correlation between δ13C and the 70–75 / 52–57 ratio was
more pronounced, whereas δ15N was not correlated with the
70–75 / 52–57 ratio but negatively correlated with the
a / o-a ratio in the oPOMs fractions.
DiscussionCryoturbation determines bulk soil organic matter distribution
The projected mean C stock of 27.5±11.9kgCm-3 corresponds to that reported in
other studies from the Siberian Arctic (cf. Zubrzycki et al., 2014,
where the authors demonstrated values between 6.6 and
48.0 kgCm-3 in their overview). Besides the large
amount of sequestered C, a noteworthy amount of N is stored in
permafrost-affected soils. Despite often being named as a decisive factor
for plant growth in usually N-deficient tundra ecosystems (Weintraub
and Schimel, 2005), soil N stocks strongly dominated by polymeric
organic N might not be related to N availability for plants in the
form of amino acids or mineral N. The values for N stocks of
permafrost-affected soils reported by other authors (cf. Fuchs
et al., 2018, and Zubrzycki et al., 2013, who demonstrated N stocks ranging
between 1.1 and 2.2 kgNm-3) are similar to our results
of 1.2±0.6kgNm-3.
Results from the molecular mixing model with data obtained by
13C CP-MAS NMR spectroscopy and calculated according to the molecular mixing model by Baldock et al. (2004) and Nelson and Baldock (2005): five-component model (without char) with N / C constraint. Given are mean values and the
standard deviation of free (fPOM), occluded (oPOM) and small occluded
(oPOMs) particulate organic matter and of clay-sized mineral-associated
organic matter (MAOM).
SOM fractionMolecular mixing model CarbohydrateProteinLigninLipidCarbonyl% fPOM61.4±8.06.4±2.821.7±3.28.2±3.92.3±1.9oPOM60.9±9.96.4±3.821.2±5.57.6±4.83.9±4.0oPOMs41.9±5.917.0±3.221.3±4.318.9±5.10.9±2.0Clay-sized MAOM41.1±2.922.6±2.021.2±3.713.5±2.81.6±2.3
Correlation matrices of POM fractions: the large POM (occluded
particulate OM and free particulate OM) fractions (a) show different
correlations compared to small occluded particulate OM fractions (b). The
more intense the color and the smaller the ellipses are, the stronger the
correlation is: blue indicates a positive correlation, and red indicates a negative correlation; the
direction of ellipses is color-related.
The ample range of the bulk soil C/N ratios points to a wide
variance in composition and the degree of decay of SOM. The C/N
ratios notably differed both between the single depth layers and the
overall soil cores. The variable bulk soil C/N ratios with depth
can be assigned to the translocation of fresh plant-derived OM from
top- to subsoils by cryoturbation, leading to specific soil layers
which also can contain so-called cryoturbated pockets rich in rather
less decomposed OM with higher C/N ratios (Kaiser et al., 2007;
Krüger et al., 2014). Such an incorporation of OM in the subsoil is
also confirmed by the high amounts of POM present in these depth
increments dominated by rather fibrous plant residues. Between the
analyzed cores, soils from three cores showed wider C/N ratios
indicative of the dominance of plant-derived OM, while the fourth
core had lower C/N ratios, pointing to a larger amount of
microbial-derived OM. Generally, C/N ratios decrease with ongoing
decomposition (Kramer et al., 2003), as the proportion of
microbial-derived OM with its characteristically low C/N ratio
increases after the depolymerization of plant-derived organic
macromolecules. This goes along with the increased binding of
microbial residues to mineral particle surfaces and thus OM becoming
less bioavailable (Connin et al., 2001; Vitousek et al., 2002).
POM fractions dominate the C stock more strongly than the N stock
The large POM fractions (fPOM and oPOM) clearly dominated the C stocks
(∼17kgm-3) in the analyzed Cryosols, whereas small
POM (oPOMs) and clay-sized MAOM represented slightly more than one
third of the stored C (∼10kgm-3). This nicely
illustrates that rather large plant-derived fragments (see Fig. 9)
dominate the C storage in these OM-rich Cryosols. Especially fPOM,
mainly consisting of less decomposed plant material, largely
contributes to both C and N stocks. However, in contrast to the C
stocks, the oPOMs and clay-sized MAOM fractions act beside fPOM as
major contributors to the N stock. A probably accelerated degradation
of the fPOM fractions under continued warming could clearly alter the
major contribution of the fPOM to the C stock. At the same time, the
increased mineralization of fPOM could release vast amounts of N,
which are assumed to further foster microbial OM mineralization. This
would increase the importance of mineral N cycling such as microbial
ammonification–immobilization turnover, compared to organic N
cycling. As permafrost-affected soils are often waterlogged during the
thawing season with changing oxygen availability and anoxic soil
microsites, it can be assumed that in these soils nitrification and
denitrification accelerate as well, thereby leading to associated
increases in nitrous oxide emissions (Marushchak et al., 2011; Voigt
et al., 2017). While a shift from aerobic to anaerobic conditions can
hamper the overall decomposition of organic compounds, a shift from
anaerobic to aerobic conditions, e.g., when a thawed Arctic soil is
exposed to drying conditions, can accelerate decomposition (Keiluweit
et al., 2017). With regard to consequences for the role of plants for
C and N budgets, some studies point to more plant-available N leading
to a changing flora and an increasing plant biomass that will possibly
be able to counteract the soil C loss caused by thawing (Sistla et
al., 2013; Keuper et al., 2017), while others question that gains in
biomass will lead to a sufficient compensation for the loss in soil C
(Salmon et al., 2016). No matter which of the predictions proves true,
as the rather labile fPOM fractions store almost one third of the N in
these soils, thawing will lead to a profound change in N budget and N
cycling with presumably increasing N bioavailability and increasing
importance of mineral N cycling (Voigt et al., 2017; Altshuler et al.,
2019).
The C and N content, C/N ratio, and decomposition proxies based on
NMR spectra clearly group the particulate OM fractions into large POM
(fPOM and oPOM) and small POM (oPOMs) (Table 1; Fig. 3). While the
large POM fractions showed rather high C/N ratios, the C/N ratios
of oPOMs were considerably lower. This demonstrates that oPOMs
represent a discrete type of SOM consisting of smaller, more degraded
organic fragments intimately connected with mineral particles,
a presumption already made by Wagai et al. (2009) for small
particulate OM. We assume that the distinct fibrous structure of the large POM fractions (see Fig. 9) drives the differentiation into
large plant-derived, less decomposed POM and mostly microbial-dominated small POM in the studied permafrost-affected soils. The less
decomposed fibrous fPOM and oPOM are hotspots for microbial activity
and, thus, for the decay of these larger plant structures (Kuzyakov
and Blagodatskaya, 2015). These hotspots for the formation of MAOM in
Cryosols, plant residues in direct contact with silt- to clay-sized
mineral particle surfaces, were already demonstrated on intact Cryosol
cross sections using spectromicroscopic imaging (Mueller et al.,
2017). The fibrous large POM provides a distinct network that entraps particles of
smaller POM and MAOM and thereby retains especially the
small POM fraction restricting its bioaccessibility. The particles of small OM (oPOMs) act as a linking element between the fresh less
decomposed plant residues (fPOM and oPOM) and the clay-sized MAOM, to our
knowledge a phenomenon not described before in permafrost-affected
soils.
Isotopic composition demonstrates the fate of labile and recalcitrant
organic compounds from POM to MAOM and the importance of biological N
fixation
The δ13C values of all samples were well within the range of
those obtained for SOM derived from plants with a C3
metabolism (Sharp, 2007). During the decomposition of plant-derived
material, the changes in δ13C values are usually subtle and
are determined by a variety of factors, especially by the composition
of the original plant material (Ågren et al., 1996).
Nevertheless, SOM compounds rich in presumably more recalcitrant
macromolecules, like lignin or aromatic hydrocarbons, have lower
δ13C values than labile compounds, like carbohydrates
(Schmidt and Gleixner, 1997). Besides the δ13C value of the
original plant litter input, soil δ13C values depend on
several factors like climate, soil texture and major soil processes
(Nel et al., 2018). We found clear positive correlations between the
decomposition stage (a / o-a ratio, 70–75 / 52–57 ratio and C/N ratio)
of the large POM fractions and δ13C (Fig. 8), which nicely
illustrates the initial decomposition stage of the large POM with
a relative dominance of labile OM rich in carbohydrates. This is
supported by the negative correlation between δ13C and the
rather recalcitrant lipids (aliphatic C based on NMR spectra) both in
the large POM and oPOMs fractions (Fig. 8). This correlation reflects
well the relative increase in aliphatic compounds with progressing
decomposition (Benner et al., 1987), which corresponds to the fact
that aliphatic compounds commonly show a lower 13C abundance
(Schmidt and Gleixner, 1997). Thus, although we demonstrate clear
mechanistic differences between large (fPOM and oPOM) and small POM
(oPOMs) with respect to C sequestration, the decomposition in both OM
pools follows the same fundamental principles. The positive
correlation of δ13C in large and small POM with the bulk
soil C/N ratios demonstrates the dominance of the POM-C pool for the
bulk soil C pool. Thus, the positive correlation between 13C and
bulk soil C/N reflects the larger amount of undecomposed plant
residues in the large POM in some soil horizons, while it demonstrates
an increased amount of aliphatic moieties in small POM for other
soils horizons. Thus, the overall elemental composition of the bulk
soils can directly be linked to the 13C isotopic composition of
the fresh and more decomposed POM fractions.
Scanning electron micrographs of particulate organic matter fractions: while free particulate OM fraction (a) and occluded particulate
OM fraction (b) consist mainly of larger particles of plant-derived litter
with clearly visible cell structures and only minor indications for initial
decomposition, the image of the small occluded particulate OM fraction (c)
of the same sample clearly reveals the intricate association of small
organic particles and silt- and clay-sized soil minerals.
In Arctic ecosystems, N2 fixation is known as the major N
input into ecosystems (Granhall and Selander, 1973; Rousk et al.,
2017, 2018) with N fixation rates between 1 and
29 kgNha-1a-1, depending on which N2-fixing species (e.g., free-living or moss-associated cyanobacteria) is
dominating (Rousk et al., 2017). Furthermore, Arctic soils are known
to be dominated by organic N cycling rather than mineral N cycling
(Hobbie and Hobbie, 2008), while atmospheric N deposition is low in
this region (Hole et al., 2009). The soil δ15N values we
found are consistent with δ15N values reported for bacterial
N2 fixation as N source (Casciotti, 2009; Hoefs, 2015) but
also similar to values reported for plant-litter-derived OM (Connin
et al., 2001). Other studies reported stable or increasing
δ15N values with advancing decomposition (e.g., Ågren
et al., 1996; Connin et al., 2001). Increases in δ15N occur with enhanced decomposition, and N turnover is largely
dependent on gaseous N loss processes, such as ammonia volatilization
and nitrous oxide (N2O) and dinitrogen (N2) losses
through nitrification and denitrification, as the highest isotope
fractionation factors are reported for these processes, enriching the
heavier 15N isotope in soil, while 14N is preferably lost to
the atmosphere (Bedard-Haughn et al., 2003; Nel et al., 2018). By
illustrating decreasing δ15N with increasing OM
decomposition, our results seem to contradict this
presumption. Therefore, we assume that biological N2
fixation is a decisive control of δ15N in the studied soils,
as also recently shown for permafrost-affected soils of Tibet (Chang
et al., 2017). Such a dominant role of biological N2
fixation in regulating δ15N requires that
nitrification or denitrification and associated gaseous N losses as well
as atmospheric inputs are not significant for the studied soils, which
is in general agreement with the N cycle paradigm for the High Arctic
(Schimel and Bennett, 2004).
Distinct differences in chemical composition from large POM to MAOM
By using NMR spectroscopy, we were able to differentiate further
between large POM fractions (fPOM and oPOM) and oPOMs and clay-sized
MAOM, which also allowed a nice clustering of these materials into
distinctly different OM pools with respect to assumed bioavailability
(see the representative example in Fig. S1). The NMR spectra of both
large POM fractions were clearly dominated by a major peak around
70 ppm and a minor peak around 105 ppm, both relating
to polysaccharides (Koelbl and Koegel-Knabner, 2004). This was well
reflected by the calculated high amounts of carbohydrates, the high
70–75 / 52–57 ratios and low a / o-a ratios, which all point to the
rather labile undecomposed nature of the larger OM particles.
Based on both decomposition proxies pointing in the same direction, we
assume a high potential bioavailability for both large POM fractions
(fPOM and oPOM). Interestingly, when comparing the decomposition proxies
between these two POM fractions per single soil layer (Table S3), they
indicate a less pronounced decomposition for oPOM in most of the
samples. These patterns deviate from what is commonly observed in
temperate soils, i.e., an increased degree of decomposition with
decreasing POM size and advancing aggregate occlusion (fPOM<oPOM<oPOMs; Mueller and Koegel-Knabner, 2009). We
assume that this demonstrates a reduced bioaccessibility
(accessibility of OM by microorganisms and enzymes) of oPOM, which is
encrusted by mineral particles, leading to a reduced degree of
decomposition of the occluded as compared to the free POM. Thus, the
initial microbial decomposition of the surfaces of fresh plant
residues (fPOM) driven by microbial decay leads, in part, to the
formation of oPOM due to the association with minerals glued to the
POM surfaces by microbial residues (see Mueller et al., 2017),
e.g., extracellular polymeric substances (Tisdall and Oades, 1982;
Schimel and Schaeffer, 2012; Costa et al., 2018). We thus demonstrate
soil structure formation in Cryosols as driven by microbial activity
via the excretion of extracellular polymeric substances at POM
surfaces leading to the stabilization of rather labile POM without
necessarily leading to OM with high degrees of decomposition.
In contrast to the large POM fractions, the NMR spectra of oPOMs and
clay-sized MAOM were dominated by peaks around 30 ppm
representing long-chain structured aliphatic C derived for example
from macromolecules like cutin or suberin (Koegel-Knabner,
2002). Irrespective of the high amount of alkyl C, the dominating
group of compounds as calculated by the MMM were carbohydrates for
both fractions, oPOMs and clay-sized MAOM (Table 3). Both fractions
also showed distinct peaks around 170 to 175 ppm, representing
partly esterified carboxyl groups and amide C that stems predominantly
from proteins (Koelbl and Koegel-Knabner, 2004). Especially the
clay-sized MAOM showed distinctly higher amounts of protein C
(Table 3) compared to all POM fractions, which corroborates the
preferential association of N-rich microbial residues at mineral
surfaces (Kleber et al., 2007; Kopittke et al., 2018, 2020). This
highlights the fact that the association of OM with mineral surfaces
follows the same mechanisms as previously described for temperate
soils (Kleber et al., 2007). In the specific context of the studied
permafrost-affected soils, the oPOMs represented a kind of passage
fraction. Although it clusters with the clay-sized MAOM in the PCA
(Fig. 5), the small POM links to the large POM fractions as
illustrated in Fig. 7. Thus, in contrast to the larger, relatively
undecomposed plant residues, lipids and proteins contribute noteworthy
to the oPOMs fractions and the fine MAOM of the clay-sized
fraction. This clearly points to the increased amount of
microbial-derived compounds in these fractions, as already stated
above with respect to the C/N ratio and δ15N. Thus, the
MAOM in these soils is dominated by microbial-derived SOM rich in
biologically fixed N. As demonstrated in the PCA and shown in Fig. 9,
oPOMs is represented by degraded plant residues, fungal hyphae and
amorphous material which can be assumed to mainly represent microbial
necromass (Miltner et al., 2012). The PCA demonstrated that oPOMs
represents a linking fraction between the initial plant residues of
the large POM fractions and the microbial OM-dominated clay-sized
MAOM (Fig. 5). However, it falls short to assume that all OM in POM
fractions is fast-cycling and all MAOM is slow-cycling (Torn et al.,
2013). Our results underline that the large and rather undecomposed
POM fractions rich in carbohydrates might act as a highly bioavailable
substrate in a warmer future. This means that when active layers deepen
and the larger POM fractions become accessible to microorganisms,
oPOMs and clay-sized MAOM may represent a C pool that is less
bioavailable and thus presumably more persistent. Besides the
demonstrated occlusion of particulate OM, we were able to show the
quantitative importance of MAOM for the C storage in these High Arctic
soils. Thus, the oPOMs and clay-sized MAOM representing altered and
microbially transformed OM pools could gain importance regarding C
storage under further thawing conditions in soils of the
Arctic. Besides the importance for C sequestration, the high amount of
biologically fixed N of MAOM may also be released and foster the
microbial decay of the high amounts of C stored in larger POM fractions (Jilling et al., 2018).
Conclusions
Employing physical fractionation and molecular-level analyses, we show
that the SOM fractions that contribute with about
17 kgCm-3 for more than 60 % of the C stocks in the
investigated Arctic soils are presumably highly labile and vulnerable
to environmental changes. In the face of global warming, most of this
labile C, currently protected from decomposition by low temperatures,
will be prone to mineralization, with severe consequences for the C
stocks in Arctic soils. Our results clearly support our hypotheses
that the major amount of C and N is stored as POM, with large POM
resembling the composition of the initial plant litter. With
increasing decomposition and, thus, decreasing size of the OM
particles, the material gets microbially transformed, which leads to
MAOM dominated by microbial residues, as indicated by C/N ratios,
15N abundance and chemical composition. Organic C stored in small
occluded POM and clay-sized MAOM that accounts with
10 kgCm-3 for about 40 % of the C stock currently
will likely dominate the C pools, as it is less vulnerable to increased
mineralization in Arctic soils in a warmer future. Small occluded POM
was found to be acting as a transitional C pool between the larger POM fractions and MAOM, demonstrating the importance of the interfaces
between particulate plant residues and the fine mineral fraction as
hotspots for microbial activity and thus MAOM formation. Using
δ15N as a proxy for N balances, we demonstrate the important
role of N inputs by biological N fixation with an increasing contribution
to organic matter N at a higher degree of decomposition, while gaseous N
losses appear to be of minor importance. The large soil organic N stocks however might be at risk in future, as with about
0.4 kgNm-3 one third of the N is present in presumably
bioavailable SOM fractions, which could lead to increases in mineral N
cycling and associated N losses under the auspices of global warming.
Data availability
The data that support the findings of this study are available from the
corresponding author upon request.
The supplement related to this article is available online at: https://doi.org/10.5194/bg-17-3367-2020-supplement.
Author contributions
IP conducted analyses in the laboratory (elemental analysis and NMR
measurements) and wrote the paper. SZ was responsible for the sampling
and the selection of the respective cores. LCZF conducted analyses in the
laboratory (fractionation, elemental analysis and NMR measurements). FB
conducted stable isotope measurements. CWM developed the design of the
study. IP, MD, GA and CWM were responsible for data evaluation and the
interpretation of results. All authors discussed the data and results and
contributed to the final form of the paper.
Competing interests
The authors declare that they have no conflict of interest.
Acknowledgements
We thank Maria Greiner for her assistance with the physical soil
fractionation and the elemental analysis, Theresa Hautzinger for her support
in the laboratory, and Stefanie Mayer for her support with statistical
analyses in R.
Financial support
This study was supported through the Cluster of Excellence “CliSAP”
(no. EXC177), Universität Hamburg, funded through the German Research
Foundation (DFG) and the BMBF project CARBOPERM (grant no.03G0836A). The analyses
were partly supported by the DFG in the framework of the priority programme
1158 “Antarctic Research with Comparative Investigations in Arctic Ice
Areas” (no. MU 3021/8). The work of Michael Dannenmann was supported through the DFG NIFROCLIM
project (no. DA1217/4-1).
Review statement
This paper was edited by Yakov Kuzyakov and reviewed by three anonymous referees.
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