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REGULAR ARTICLE
Deep soil organic matter—a key but poorly understood
component of terrestrial C cycle
Cornelia Rumpel &Ingrid Kögel-Knabner
Received: 20 November 2009 / Accepted: 12 April 2010 /Published online: 15 May 2010
#Springer Science+Business Media B.V. 2010
Abstract Despite their low carbon (C) content, most
subsoil horizons contribute to more than half of the total
soil C stocks, and therefore need to be considered in the
global C cycle. Until recently, the properties and
dynamics of C in deep soils was largely ignored. The
aim of this review is to synthesize literature concerning
the sources, composition, mechanisms of stabilisation
and destabilization of soil organic matter (SOM) stored
in subsoil horizons. OrganicC input into subsoils occurs
in dissolved form (DOC) following preferential flow
pathways, as aboveground or root litter and exudates
along root channels and/or through bioturbation. The
relative importance of these inputs for subsoil C
distribution and dynamics still needs to be evaluated.
Generally, C in deep soil horizons is characterized by
high mean residence times of up to several thousand
years. With few exceptions, the carbon-to-nitrogen
(C/N) ratio is decreasing with soil depth, while the
stable C and N isotope ratios of SOM are increasing,
indicating that organic matter (OM) in deep soil
horizons is highly processed. Several studies suggest
that SOM in subsoils is enriched in microbial-derived C
compounds and depleted in energy-rich plant material
compared to topsoil SOM. However, the chemical
composition of SOM in subsoils is soil-type specific
and greatly influenced by pedological processes. Inter-
action with the mineral phase, in particular amorphous
iron (Fe) and aluminum (Al) oxides was reported to be
the main stabilization mechanism in acid and near
neutral soils. In addition, occlusion within soil aggre-
gates has been identified to account for a great
proportion of SOM preserved in subsoils. Laboratory
studies have shown that the decomposition of subsoil C
with high residence times could be stimulated by
addition of labile C. Other mechanisms leading to
destabilisation of SOM in subsoils include disruption of
the physical structure and nutrient supply to soil
microorganisms. One of the most important factors
leading to protection of SOM in subsoils may be the
spatial separation of SOM, microorganisms and
extracellular enzyme activity possibly related to the
heterogeneity of C input. As a result of the different
processes, stabilized SOM in subsoils is horizontally
stratified.InordertobetterunderstanddeepSOM
dynamics and to include them into soil C models,
quantitative information about C fluxes resulting from C
input, stabilization and destabilization processes at the
field scale are necessary.
Keywords Subsoil .Organic matter .
Chemical composition .Carbon stabilization
Plant Soil (2011) 338:143–158
DOI 10.1007/s11104-010-0391-5
Responsible Editor: M. Francesca Cotrufo.
C. Rumpel (*)
CNRS, BIOEMCO (UMR CNRS-Université Paris
VI, XII-IRD-AgroParisTech), Campus AgroParisTech,
Thiverval-Grignon, France
e-mail: cornelia.rumpel@grignon.inra.fr
I. Kögel-Knabner
Lehrstuhl für Bodenkunde,
Technische Universität München,
Freising-Weihenstephan, Germany
Introduction
The soil is the largest active terrestrial reservoir in the
global carbon cycle. The estimates of the organic C
stocks in 0–100 cm depth in the world’s soils range
from 1,220 Pg (1 Pg=10
15
g) (Sombroek et al. 1993)
to about 1,550 Pg (Eswaran et al. 1993; Batjes 1996;
Jobbagy and Jackson 2000). Recent studies suggest
that the soils C pool may be even greater and could
account for 2,000 Pg (Janzen 2005). These higher
values may be mainly due to additional recent
estimations of the C pool stored in boreal soils under
permafrost conditions (Zimov et al. 2006; Tarnocai et
al. 2009). Although routine soil surveys collect C
stock data down to a depth of 1 m, scientists studying
the composition and mechanisms of stabilisation of
SOM have mainly focused on the A horizon with the
highest SOM concentrations.
Organic matter stored in subsoil horizons below
the A horizon has received increasing interest in
recent years as high proportions of total C stored
within the soil profile may be found in subsoil
horizons despite low OM concentrations (Batjes
1996; Jobbagy and Jackson 2000). The proportion
of SOM stored in the first meter of the world soils
below 30 cm depth ranges between 63 and 46%,
except for Podzoluvisols, where 30% of OC is stored
below the first 30 cm (Batjes 1996). A recent study
suggests that in the northern circumpolar permafrost
region, at least 61% of the total soil C is stored below
30 cm depth (Tarnocai et al. 2009). Therefore, subsoil
C may be even more important in terms of source or
sink for CO
2
than topsoil C. Another property of
subsoil C is its high radiocarbon age, which suggests
that a high proportion of this C is stable at longer
timescales (e.g. Scharpenseel et al. 1989; Paul et al.
1997). Subsoil horizons with low C concentrations
may not yet be saturated in organic C. It has therefore
been suggested that they may have the potential to
sequester organic carbon for example through higher
C input into subsoil by roots and DOC (Lorenz and
Lal 2005). Recent work, however, suggested that
increase of OM storage in subsoils may not be as
straight forward, because subsoil C may become
available to microbial decomposition following C
input (e.g. Fontaine et al. 2007) and/or mechanical
disruption (Xiang et al. 2008). It also has been found
that subsoil C may respond to land-use and/or
management change (Guo and Gifford 2002;Wright
et al. 2007; Follett et al. 2009). In the context of
climate change and SOM as potential sink for
atmospheric CO
2
, it is important to understand what
leads to sequestration of large amounts of old C
below the A horizon. OM input into subsoil horizons
occurs as root litter and root exudates, dissolved OM
and/or bioturbation. The relative importance of these
four sources may depend on climate, soil and
vegetation types. In subsoil horizons, environmental
conditions may be different from those in topsoil
horizons, and OM storage may be driven by specific
processes (von Lützow et al. 2006). For example,
Fierer et al. (2003) observed that mineralisation of
subsoil C may be much more sensitive to tempera-
ture change than those of topsoil C.
In the last 10 years, our understanding of subsoil C
has made significant progress, despite the methodo-
logical problems encountered due to the low C
concentrations. Scientists became aware of the impor-
tance of pedological functioning with regards to OM
storage in subsoil. In this review, we would like to
discuss these findings with regard to the following
questions:
&Is there a stabilisation of specific compounds in
subsoil horizons ?
&What are the mechanisms that lead to stabilisation
of OM in subsoils ?
&To what extend does subsoil OM participate in C
and N cycles ?
Sources of organic matter in subsoils
Four main sources of OM input into subsoils have
been identified: plant roots and root exudates,
dissolved organic matter and bioturbation. In addition
there may be translocation of particulate organic
matter and transport of clay-bound organic matter in
certain soil types. The relative importance of these
sources is dependent on climatic parameters, soil
inherent processes as well as land-use. For example,
high input of dissolved organic matter can be
expected under humid climate conditions and when
podzolisation is the main soil forming process. DOC
transport was largely studied in temperate forest soils
(Michalzik et al. 2001) and has been thought to be the
main source of subsoil OM under such conditions
(Kaiser and Guggenberger 2000). The sharp decrease
144 Plant Soil (2011) 338:143–158
of DOC concentrations with depth of mineral soil
observed by many authors was explained by strong
retention in the mineral soil due to adsorption (Qualls
and Haines 1992; Kaiser and Zech 1997). DOC fluxes
in the mineral subsoil were reported to range between
10 and 200 kg ha
-1
yr
-1
. These fluxes were not
controlled by carbon storage, precipitation, C/N ratio
or pH (Michalzik et al. 2001). Recently it was
reported, that harvesting management (Strahm et al.
2009) as well as land-use could influence these fluxes
and thus subsoil C pools. The process of DOC
movement and retention within the mineral soil was
found to be responsible for 20% of the total mineral
soil C stock to 1 m depth in a forest soil and 9% in a
prarie soil (Sanderman and Amundson 2008).
Another important source of subsoil OM are plant
roots. These were found to affect the placement of C
in soil. In a global review of root distributions,
grasses had the shallowest root profiles, trees were
intermediate and shrubs had the deepest profiles
(Jackson et al. 1996). Specific allocation patterns
through vegetation types were also found to govern
vertical SOC distribution (Jobbagy and Jackson
2000). The importance of roots for soil C sequestra-
tion was underlined by the fact that they have a high
potential to be stabilized in soil (Rasse et al. 2006).
Despite their importance as a subsoil C source, root C
flux to soil are poorly understood mainly due to
uncertainties associated with the measurement of total
root C input, in particular from root exudation and
root cell sloughing. Root litter production can be
estimated from root turnover. Root turnover can be
measured directly using observation of roots from
birth to disappearance with microrhizotrons (Kleja
et al. 2008). However, minirhizotrons are only able
to estimate the most dynamic roots (<1 mm) and not
roots with larger diameter (>1 mm) for which isotope
techniques as
14
Cand
13
C may be more suitable
(Majdi and Andersson 2005). Depending on the
method the longevity of roots was found in the order
of 1–18 years (Kleja et al. 2008; Gaudinsky et al.
2001). Using the microrhizontron technique root C
input in the mineral soil was estimated 73 to
78 gC m
-2
yr
-1
for a northern hardwood forest (Kleja
et al. 2008). At this site DOC input ranged between
11 and 26% of the total carbon input. However,
considering root litter and DOC decomposition rates,
the authors estimated DOC and roots equally
important for SOM buildup in soil.
Physical carbon transport down the soil profile as
colloidal Fe/Al-humus complexes is an important
process increasing SOM of volcanic subsoils (Osher et
al. 2003). In Alisols, Luvisols, Acrisols and Lixisols
(FAO taxonomy), SOM input into subsoils may occur
as organomineral complexes. These complexes were
found to be enriched in labile compounds stabilized
against decomposition (Schmidt et al. 2000). Particulate
organic matter such like black carbon seems to migrate
easily into deeper soil horizons (Dai et al. 2005;
Rumpel et al. 2009) and could constitute an important
input of chemically recalcitrant C into subsoil horizons
of fire affected ecosystems. Migration of particles can
be enhanced by bioturbation.
Earthworms, termites, ants, arthropods and tree roots
are efficient in burying soil while forming voids in the
form of burrows, nests, chambers, galleries and root
channels (Paton et al. 1995; Lavelle et al. 1997). These
“biofabrics”indicate that biotic activity declines non-
linearly with depth with notable differences between
horizons (Humphreys 1994). Data on biological
activity with depth was summarized by Ekschmitt et
al. (2008) and reveal that the abundance of organisms
in subsoils does not generally follow vertical SOC
distribution. Bioturbation affects directly as well as
indirectly inputs of SOC in subsoils (Wilkinson et al.
2009). Direct inputs include litter sequestration into
nests, termitaria, borrows etc and bioturbator waste
disposal in form of dead tissues. Indirect inputs of
SOC into subsoils may occur by infilling of biogenic
pits with litter, redistribution of SOC and subsurface
mixing and burial. Biologically-mediated soil burial
rates range between 1-2 m My
-1
.
Several sources for C input into subsoils have been
identified. The relative importance of these inputs has
rarely been studied. Data from different climatic regions
are necessary to elucidate completely the controls on the
different inputs. The importance of biological SOM
burial and mixing in subsoil horizons is an important
factor that needs to be considered in future studies.
Nature of SOM in subsoils
Mean residence time and SOM turnover
The radiocarbon age of SOM is usually increasing
with soil depth, and carbon present in the subsoil
horizons is characterized by a low
14
C activity
Plant Soil (2011) 338:143–158 145
(Scharpenseel et al. 1989; Paul et al. 1997; Krull and
Skjemstad 2003; Eusterhues et al. 2003). Mean
radiocarbon ages of subsoils of different soil types
are presented in Table 1. In 1-m depth the radiocarbon
age of all studied soil types exceeds 1,000 years. The
reasons for the increase of the
14
C age of SOM with
depth are not completely understood (Trumbore
2009). High
14
C age of SOM may indicate that
stabilized carbon compounds with long residence
times are found in subsoil horizons at higher concen-
trations. However, recycling of old, stabilized SOM in
subsoils through microbial biomass may also lead to
old
14
C age of chemically labile, newly synthesized
carbon compounds (Rethemeyer et al. 2005). Thus,
old
14
C age of subsoil OM may also be observed as a
result of continuous microbial recycling of labile
material (Gleixner et al. 2002). For SOM in subsoils,
it must additionally be considered that the
14
C activity
may be influenced by the contribution of substrate
inherent geogenic carbon, which is usually carbon
dead (= older than 50,000 years), not showing
14
C
activity. This may be the case for soils developed
from sedimentary parent substrates, such as loess. The
small amounts of carbon associated with loess
deposits represent a small proportion of SOC in
surface horizons, but could be significant at depth
(Paul et al. 2001; Helfrich et al. 2007). Therefore the
very old radiocarbon age of some soils may simply be
due to dilution of geogenic (rock-derived) dead
carbon with younger SOM. However, even for SOM
in soils developed from parent substrate free of
geogenic C (e.g. granite), radiocarbon ages of several
thousand years have been reported (e.g. Eusterhues et
al. 2003). Further indication for low turnover of
subsoil carbon was derived from stable carbon isotope
analysis on sites, where a C3 vegetation was replaced
by a C4 vegetation. At a site, where a forest
dominated by C3 vegetation was replaced by corn
(C4 species) monoculture in the United States,
incorporationofC4carbonreached4–15% in 50–
100 cm depth after 30 years (Collins et al. 1999).
This corresponds to mean residence times of 100 to
700 years. In an agricultural soil in France, 10 years
of continuous corn after wheat monoculture resulted
in 10, 5 and 2% corn derived SOC at 15, 50 and
100 cm depth, respectively (Rasse et al. 2006).
Elemental and isotopic composition of subsoil OM
Generally, the SOM content and its C/N ratio are
decreasing rapidly below the A horizon. Low C/N
ratios have been attributed to highly processed SOM.
In most subsoils, C/N ratio is approaching that of
microbes (Wallander et al. 2003). In subsoils, which
are generally characterized by a very low organic
matter content, high nitrogen content may be related
to the presence of mineral nitrogen sorbed to clay
surfaces. Mineral nitrogen was found to contribute to
about 20% to the total nitrogen of deep soil horizons
and even when subtracted, C/N ratios of most soils
are decreasing with depth (Jenkinson et al. 2008;
Krull and Skjemstad 2003). An increase in C/N with
depth in some soils may be explained by the presence
of charred material (Dümig et al. 2009).
The stable isotope ratios δ
13
Candδ
15
Nare
generally increasing with depth and degree of
decomposition in soils without vegetation change
(Balesdent and Balabane 1996; Balesdent et al.
1993; Nadelhoffer and Frey 1988). Several factors
may be responsible for a
13
C enrichment of subsoil
SOM: (a) increasing atmospheric
13
CO
2
due to fossil
fuel burning (the so-called Suess effect) may account
for a 1.5% increase since 1800 (Leavitt and Long
1988); (b) it may be ascribed to a preferential
stabilization of
13
C enriched compounds, such as
polysaccharides and amino acids and the preferential
decomposition of
13
C depleted compounds, such as
lipids and lignin. In some cases, mainly under C4
grassland a decrease of δ
13
C with depth was recorded
(Volkhoff and Cerri 1987; Martin et al. 1990; Gill and
Burke 1999; Krull et al. 2005; Dümig et al. 2008).
This may be explained by the accumulation of
13
C
depleted charred material, as C4 grasslands are prone
Table 1 Mean radiocarbon ages (years B.P.) for different
depths and soil types sampled in Europe, Australia, Israel,
Tunisia, Sudan and Argentina (data from Scharpenseel and
Becker-Heidmann, 1998)
Soil type n 20 cm 50 cm 100 cm
Year B.P.
Alfisol 13 960 2,400 4,800
Inceptisol 16 920 1,000 1,160
Mollisol 47 1,240 2,700 5,150
Spodosol 9 1,430 1,680 2,100
Vertisol 44 410 1,620 3,650
146 Plant Soil (2011) 338:143–158
to disturbance by fire. Finally, isotopic fractionation
during microbial respiration was thought to be another
mechanism leading to
13
C enrichment. Van Dam et al.
(1997)reported
13
C enrichment of 3% due to microbial
respiration. Recently, it was suggested, that this
mechanism does not contribute to
13
C enrichment of
SOM with increasing depth (Boström et al. 2007).
Instead, Boström et al. (2007) hypothesized that the
increase of δ
13
C of organic matter as well as the
13
C
enrichment of respired CO
2
with soil depth is caused
by an increasing contribution of microbial derived
carbon. This hypothesis may be supported by the
observation that the δ
15
N may increase by up to 10%
with depth (Högberg 1997). While it has been
suggested that the
15
N gradient is mainly due to the
mineralization-plant-uptake pathway followed by the
deposition of
15
N depleted litter to the soil surface, it
was also shown that microbial derived products
contribute to this gradient (Högberg 1997).
Is subsoil carbon microbial derived carbon?
The elemental and isotopic evidence suggests that SOM
in subsoils is more microbially processed than topsoil
OM and most probably has a higher proportion of
microbial derived compounds. When compiling data
published by Jenkinson et al. (2008), we noticed a
negative correlation between the soils C/N ratio and
their
14
C activity (Fig. 1). This indicates that N
containing compounds may survive in subsoil horizons
for a long time, as already reported by Gleixner et al.
(1999) for topsoils. Enrichment of microbial derived
aminosugars in subsoil horizons was found by Liang
and Balser (2008) who stated that ‘microbial residues
are refractory and that they contribute to terrestrial
carbon sequestration’. Further evidence for the impor-
tance of microbial over plant derived carbon in subsoil
horizons was obtained from the analysis of non-
cellulosic neutral carbohydrates. Microbial-derived
sugars associated with the mineral phase were found
to be positively correlated to the
14
C activity of the
bulk sample, suggesting that these easily degradable
substances are effectively stabilized by mineral inter-
actions (Rumpel et al. 2010). Microbial sugars in the
clay fraction of subsoil horizons were found to be
associated with poorly crystalline Fe oxides but this
was not the case for plant-derived lignin (Spielvogel et
al. 2008). Plant-derived lignin may not be stabilised in
subsoil horizons (Rumpel et al. 2004; Mikutta et al.
2006). Together, this points to the importance of
microbial-derived organic matter in subsoil horizons.
Microbial derived organic matter is enriched in
subsoil horizons with regards to plant-derived organ-
ic matter and therefore most probably occurs in
higher concentrations compared to topsoil horizons.
Most studies to date used biomarkers and, therefore,
little is known about the quantity of microbial products
in subsoils.
Importance of pedological processes
The question arises, what is the difference between the
composition of stabilized SOM in subsoil horizons
compared to the topsoil with most likely higher
contributions of plant-derived organic matter? Although
few chemical structural information is available due to
analytical limitations, current evidence suggests, that
the chemical composition of SOM in subsoil horizons
under temperate climate is soil-type specific and
dependent on pedological processes operating within
the soil profile (Rumpel et al. 2002;Krulland
Skjemstad 2003; Spielvogel et al. 2008; Kögel-
Knabner et al. 2008), rather than dependent on the
input of chemically stable compounds. This is consis-
tent with results on the chemical composition of
mineral-associated SOM in top- and subsoil horizons
in a recent study by Mikutta et al. (2009). This study
indicates that mineral-associated SOM in topsoil
horizons predominantly reflects the uniform source
vegetation whereas subsoil mineral-associated SOM
shows different contribution of organic compounds
related to the composition of the mineral phase. Thus,
organic and mineral compounds in subsoil horizons
have to be studied simultaneously to understand SOM
dynamics (Basile-Doelsch et al. 2005).
Stabilisation processes of SOM in deep soil
horizons
Old radiocarbon age of subsoil SOM as well as
specific elemental and stable isotope composition
could result from diffusion of processed OM at depth,
where they are adsorbed on unsaturated mineral
surfaces. However, such an explanation does not take
into account considerable input of unprocessed litter
through bioturbation. The importance of such input is
evidenced by the relation between deep soil carbon
Plant Soil (2011) 338:143–158 147
stocks and plant functional types on a global scale
(Jobbagy and Jackson 2000).
Chemical recalcitrance
Litter input into deeper soil may mainly occur as root
C (Rasse et al. 2005), which could be chemically
more recalcitrant than topsoil litter due to higher
concentrations of lignin and aliphatic material (Nierop
1998; Lorenz and Lal 2005). Moreover, this material
may already be old, as fine roots of forest trees were
recently found to live several decades (Gaudinsky et
al. 2001; Trumbore et al. 2006). This is longer than
expected, but still a lot shorter than the mean
residence time of several hundred or thousand of
years recorded for the surrounding soil. Of the three
mechanisms of stabilisation reported in the literature
(physical protection, physico-chemical interaction and
stable chemical structure), chemical recalcitrance due
to a highly stable chemical structure of certain organic
matter compounds (Derenne and Largeau 2001)is
controversially discussed. Krull et al. (2003) high-
lighted chemical recalcitrance as the only mechanism
by which SOC can be protected over long periods of
time whereas Marschner et al. (2008) provided
evidence that chemical recalcitrance may no longer
be regarded as a long-term stabilisation mechanism of
SOM. Chemical recalcitrant organic matter may be
enriched by acid hydrolysis (Quénea et al. 2005). This
method was frequently used for the isolation of
chemical recalcitrant C from subsoil horizons of
temperate and tropical soils (Trumbore and Zheng
Fig. 2 Relationship between HF soluble C and OC in dense fractions > 2 g cm
3
(Figures from Eusterhues et al. 2007)
y = 9.2x – 2.2
r2 = 0.79
10
20
30
40
50
60
70
80
90
100
0 2 4 6 8 10 12 14 16
C/N ratio
14C activity (pMC)
Young C
Old C
Fig. 1 Relationship be-
tween the
14
C activity and
the C/N ratio of subsoil
layers from Rothamstad
(Data from Jenkinson et al.
2008)
148 Plant Soil (2011) 338:143–158
1996; Paul et al. 2001; Leavitt et al. 1996; Poirier et
al. 2002; Paul et al. 1997; Rumpel et al. 2008). Acid
hydrolysis was thought to remove easily decompos-
able protein and polysaccharide material leaving
behind chemical recalcitrant structures, such as lignin,
black carbon and aliphatic molecules. Carbon in
subsoils may or may not be lost by acid hydrolysis
at higher amounts than topsoil C (Paul et al. 2001;
Rumpel et al. 2008; Montané et al. 2007). Acid
hydrolysis, however, has been found to produce
artefacts, (Allard et al. 1998) through the Maillard or
browning reaction (condensation reaction between
proteins and polysacchrides, Maillard 1912). More-
over, this method converts crystalline minerals to
more amorphous ones (Zimmermann et al. 2007). The
C isolated by acid hydrolysis from subsoil horizons is
several hundred or thousand years older than bulk
SOM (Paul et al. 2001; Rumpel et al. 2008). These
results suggest that acid hydrolysis may be able to
isolate subsoil C with long-term stability. Black
carbon, produced by vegetation fire is chemically
recalcitrant, thus may contribute to the stabilized
SOM pool in subsoils. Charcoal with a
14
C age of
several thousand years was isolated by many authors
from subsoil horizons (exp. Ali et al. 2005). More-
over, black carbon seems to be prone to vertical
migration (Dai et al. 2005; Rumpel et al. 2009) and
may therefore accumulate in deep soil horizons. Black
carbon was thought to be preserved in soil due to its
stable chemical structure. Several authors evidenced
that it is intimately associated with the mineral phase
(Brodowski et al. 2005; Czimczik and Masiello
2007). Stabilization of other, plant-derived OM types
in subsoil horizons seems to be more related to
adsoption and chemical binding to mineral surfaces or
the physical encapsulation of SOM than to recalci-
trance of specific chemical compounds (Schöning and
Kögel-Knabner 2006).
Association with soil minerals
One of the most prominent properties of SOM in
subsoils is its high susceptibility to C-loss upon
demineralisation with hydrofluoric acid (HF, Table 2).
This is in sharp contrast to carbon loss by HF
treatment of topsoil horizons and pure organic matter
which is usually lower than 20% (Skjemstad et al.
1994; Schmidt et al. 1997; Rumpel et al. 2006). For
many soil types of the temperate and the tropical
regions it was observed, that subsoil horizons lose
more carbon upon HF treatment than topsoil horizons
(Table 2). This is especially evident, when we look at
relative carbon losses compared to initial carbon
content of the sample. The absolute loss is sometimes
similar or higher in topsoil and subsoil horizons. The
carbon lost upon demineralisation of soil was hypoth-
esised to be intimately associated with the mineral
phase (mineral-bound carbon). This hypothesis was
supported by the observation of a strong correlation
between HF-soluble C and C recovered in the heavy
fraction of subsoil horizons (Eusterhues et al. 2007;
Fig. 2). Evidence for the importance of mineral
interaction for the stabilisation of subsoil OM was
provided by the observation that the radiocarbon
activity of SOM in subsoil horizons was related to
the contribution of HF soluble C. This was observed
in temperate (Eusterhues et al. 2003; Spielvogel et al.
2008) and in tropical soils (Rumpel et al. 2008).
However, HF-resistant and presumably mineral-
bound carbon in subsoil horizons is not always older
than bulk organic C (Eusterhues et al. 2007). The
reasons may be that: (a) the mineral associated soil
organic matter fraction of deeper horizons may be
** significantly different from matrix (p < 0.05)
***si
g
nificantl
y
different from matrix (
p
< 0.01)
Tongues
Matrix
C = 5,7 ± 2,1 mg g-1**
14C activity= 101 ± 2 pMC***
C = 2,7 ± 0,7 mg g-1
14C activity= 55 ± 1 pMC
Fig. 3 A picture from a subsoil showing the presence of
tongues and a strong small scale spatial heterogenity of carbon
content and
14
C activity (data from Chabbi et al. 2009)
Plant Soil (2011) 338:143–158 149
more easily replenished by input of young adsorbing
materials compared to particulate organic matter, (b)
stabilization due to interactions with the mineral
matrix is not as effective as the other mechanisms,
which stabilized the particulate organic matter, and/or
(c) only the insoluble part of the mineral-associated
organic matter fraction is stable or experiences
significant stabilization through interaction with the
mineral matrix.
Chemical oxidation treatments have been increas-
ingly applied to subsoil OM to isolate old, organic
matter stabilised by the interaction with the mineral
phase (Eusterhues et al. 2005; Helfrich et al. 2007;
Lorenz et al. 2009). The hypothesis that OM
stabilised by interaction with the mineral phase may
be isolated by chemical oxidation is based on the
observation that minerals such as clay and poorly
crystalline minerals protect organic matter against
oxidative destruction (e.g. Hosking 1932; Singer and
Huang 1993). Since chemical oxidation leads to
alterations similar to those observed during biodegra-
dation, the residual organic C might also be resistant
to biodegradation (Cuypers et al. 2002). A number of
different chemicals are used and a presentation of
those and their effect on organic matter and mineral
phases is beyond the scope of this review. Excellent
overviews on these methods were provided by
Mikutta et al. (2005) and von Lützow et al. (2007).
The amount of non-oxidisable C present in soil
samples from different depths along the soil profile
was found to be correlated to the
14
C activity of the
sample, suggesting similar to HF soluble C that SOM
in subsoil horizons becomes older, when a greater
proportion of SOM is associated with the mineral
phase (Eusterhues et al. 2005). Material isolated by
oxidation treatments from subsoil horizons is gener-
ally older than bulk SOM (Helfrich et al. 2007;
Kleber et al. 2005) and may therefore represent a
significant proportion of the stabilised OM. This may
be different for topsoils, where this treatment was
unable to isolate old SOM (Bruun et al. 2008). The
chemical nature of the material associated intimately
with the mineral phase yet needs to be elucidated.
To study the stabilising mineral agents in subsoil
horizons, correlations between oxidation resistant OM
and mineral parameters were run. The results suggest
that in acid subsoil horizons under temperate and dry
tropical forests, the amount of oxidation resistant
organic matter is strongly related to the content of
poorly-crystalline mineral phases (Kleber et al. 2005;
Lorenz et al. 2009). Other authors found that Fe and
clay minerals are the most important stabilising agents
of OM in temperate and tropical subsoils (Eusterhues
et al. 2003; Rumpel et al. 2008; Dick et al. 2005). In
subsoils of Podzols in a tropical environment, SOM
seems to be stabilised in Al-organic matter complexes
(Skjemstad 1992; Bardy et al. 2007). Al-organic
matter complexes and poorly crystalline Al mineral
phases may also be responsible for C stabilisation in
subsoil horizons of temperate forest and grassland
Table 2 Carbon loss after HF-treatment from topsoil and subsoil horizons
Soil type Climate Conc HF C-loss C-loss Source
% % of total C mg g
-1
topsoil subsoil topsoil subsoil
Haplic Podzol Temperate 10 19 45–82 7 1–41 Rumpel et al. (2002)
Dystric Cambisol Temperate 10 10 15–92 8 1–5 Rumpel et al. (2002)
n.n. Temperate 10 18–32 31–45 3–4 1-2 Lorenz et al. (2006)
Typic Haplorthods Temperate 2 26–31 55–67 20–56 15–44 Dai and Johnson (1999)
Inceptisol and Entisol Tropical 10 24–42 38–63 15–20 5–13 Lorenz et al. (2009)
Entisol, Alfisol, Inceptisol Tropical 10 24–32 33–70 3–73–5 Rumpel et al. (2006)
Haplic and Stagnic Luvisol Temperate 10 12–25 23–61 6–17 1–7 Schöning and Kögel-Knabner (2006)
Ferralsols Tropical 10 44–50 52–84 * * Dick et al. (2005)
Leptic Cambisol Temperate 10 19–51 32–59 8–26 2–13 Mueller and Kögel-Knabner (2009)
Entic Podzol Alpine 10 0 26–85 0 0–20 Favilli et al. (2008)
150 Plant Soil (2011) 338:143–158
soils in Hawaii (Torn et al. 1997), the U.S. (Masiello
et al. 2004; Rasmussen et al. 2005) and in Germany
(Spielvogel et al. 2008).
Evidence for the high potential of pedogenic
oxides to stabilise organic matter in subsoil hori-
zons was obtained under strongly acidic conditions.
However, there are indications that the sorptive
capacity of the oxides is strongly reduced in near
neutral soil environments (Gu 1994). Under near
neutral conditions protection of OM via complexa-
tion with Ca
2+
ions was proposed (Muneer and
Oades 1989a,b). Recently, strong organo–mineral
associations in the presence of reactive, positively
charged mineral phases such as hydrotalcite were
reported for young calcareous subsoils (Grunewald
et al. 2006). Moreover, as aggregate formation depends
on biological activity, this process may be more
important in neutral or calcareous soils compared to
acid soils.
Aggregate protected SOM
In subsoil horizons, physical protection of occluded
particulate organic matter (oPOM) was identified as a
second important stabilisation mechanism. The oPOM
fraction isolated from subsoil horizons was character-
ised by the highest
14
C age (Rasmussen et al. 2005)
compared to all other density fractions. The importance
of physical protection for OM stabilisation in subsoil
horizons does not only concern the POM fraction, but
also clay associated OM in microaggregates. Clay-
associated OM in microaggregates was more abundant
throughout a soil profile and had a higher
14
C age than
OM associated with the free clay fraction (Moni et al.
2010). The importance of aggregation for stabilisation
of OM in subsoil horizons at different scales is a topic
in need for additional research, because the mecha-
nisms leading to aggregate formation in top- and
subsoils may be quite different. In subsoils, biological
aggregate formation is supposedly less important than
physical processes. However, to the best of our
knowledge no studies were carried out addressing the
incorporation of SOM into the two different types of
aggregates in subsoil horizons. The importance of
physical processes for the formation of aggregates in
subsoils should be soil type specific. Therefore, we
suggest that more studies are needed which address the
interaction between aggregate formation and OM
protection in subsoil horizons.
Destabilisation processes of SOM in deep soil
horizons
One of the questions, which need to be answered
under the scenario of global change and human
interventions, is: “how fast does subsoil C respond
to environmental changes”(Trumbore 2009). First
evidence about the extent to which subsoil OM
participates in C and N cycles can be obtained from
sites after land-use change. For example, SOM stock
changes of up to 48% may occur in B horizons after
land-use change from forest to grassland (Müeller and
Kögel-Knabner 2009). Similar results were obtained
by Don et al. (2009), who showed that 23 to 29 years
after conversion of cropland to grassland the subsoil
horizons contained fresh carbon. The authors concluded
that subsoil C has to be taken into account when land-
use change effects on SOC are assessed. The same
conclusion was obtained by Osher et al. (2003), who
estimated, that they would have greatly overestimated
C losses after land-use change if they had not taken
into account the subsoil horizons, where some of the C
lost from the topsoil had accumulated. As was
discussed previously, the degree of response and the
mechanisms responsible for carbon stock changes in
subsoil horizons are poorly understood.
Characteristics of microbial biomass and activity
in subsoil horizons
The abundance of microbial biomass is generally
declining with soil depth (Taylor et al. 2002;Andersen
and Domsche 1989; Ekklund et al. 2001)aswellasits
activity (Fang and Moncrieff 2005). The opposite was
noted for the Cmic/Corg quotient, suggesting that in
deeper horizons a higher proportion of organic C is
present as microbial biomass (Agnelli et al. 2004).
Despite the low biomass contents, subsoils are meta-
bolically active and contain substantial numbers of
microorganisms (Taylor et al. 2002). Studies on
pesticide degradation showed, that microbial activity
was still significant below the rooting zone down to a
depth of 2.5 m (Ajwa et al. 1988; Charnay et al. 2005).
Similar observations were made concerning amino
acid biodegradation (Kemmitt et al. 2008). Microbial
activity in subsurface horizons was found to be similar
to that measured in surface soil when normalized to
biomass size (Blume et al. 2002). The absence of a
decreasing trend is interesting as several authors
Plant Soil (2011) 338:143–158 151
observed differences in microbial community structure
between surface and subsurface horizons. In general,
the genetic and metabolic diversity was reduced in
subsoil horizons (Goberna et al. 2005; Agnelli et al.
2004). Carbon sources for soil bacteria with increasing
depth increasingly consisted of older SOM derived C
(Kramer and Gleixner 2008), suggesting, that those
can be used as C sources by the soil organisms present.
Factors influencing the availability of subsoil OM
The availability of SOM to soil microorganisms in
subsoil horizons may be affected by depth-specific
environmental conditions (Table 3). Oxygen limita-
tions in subsoil horizons could be responsible for long
C turnover times. However, no literature evidence has
been published so far supporting this assumption.
Low oxygen content of 5% had no effect on
decomposition of subsoil SOM (Salomé et al. 2010).
This result obtained during laboratory incubation was
supported by the observation that POM was not
accumulating in subsoil horizons (Fontaine et al.
2007). In contrast, the conditions for the decomposi-
tion of plant material in subsoils were reported to be
better than in topsoil in a Mediterranean environment
due to higher moisture (Rovira and Vallejo 2002).
It has been reported that the degradation of old
organic matter in subsoil of a Cambisol under prairie
is absent due to energy limitation of the microbial
biomass as a consequence of reduced input of fresh
organic matter (Fontaine et al. 2007). The authors
showed that the decomposition of old SOM may be
stimulated by the addition of cellulose (priming
effect). Temperature and available nutrients may be
additional control mechanisms for C mineralisation in
subsoil horizons. Q10 values of subsoil OM were
found to exceed those of topsoil SOM (Fierer et al.
2003). In northern boreal regions, permafrost may be
the main factor stabilising subsoil C (Trumbore
2009), as soils with similar parent material and
vegetation without permafrost were found to contain
lower C stocks than their frozen counterparts (Rodionov
et al. 2007). Unter temperate climate conditions,
nitrogen availability seems to have a strong influence
on C as well as N mineralization in subsoil horizons of
mountain grassland soils (Garcia-Pausas et al. 2008).
In contrast, in a well managed agricultural soil
receiving regular fertiliser treatment nutrient inputs
did not affect subsoil C mineralization (Salomé et al.
2010). In this soil, addition of energy rich material to
subsoil OM was not inducing a priming effect. It was
rather the location of subsoil C within the soil matrix
that appeared to determine its availability (Salomé et
al. 2010).SimilarobservationsweremadebyXianget
al. (2008) who noticed that the mineralisation of
subsoil C could be greatly increased by inducing
drying wetting stress. However, the large majority of
these studies was carried out under controlled labora-
tory conditions. The relevance of these works at the
field scale yet needs to be demonstrated.
The spatial separation of microbes and degradable
substrates may be one of the most important factors
influencing C dynamics in subsoils (Von Lützow et al.
2006;HoldenandFierer2005). It has been recognised
that small scale variability of SOM has an important
influence on microbial activity and growth in subsoil
horizons. In a soil profile, spatial heterogenity of SOM
distribution does not only occur in vertical but also
Table 3 Effects of different factors on microbial activity in subsoil horizons
Soil type Depth Treatment Microbial activity Source
Cambisol 80–100 cm Physical disturbance +75% Salomé et al. (2010)
Fructose amendment No effect Salomé et al. (2010)
Low oxygen No effect Salomé et al. (2010)
Nutrient amendment No effect Salomé et al. (2010)
Fluventic Eutrochrept 30–55 cm Thrawing-freezing No effect Lomander et al. (1998)
Cambisol 60–80 cm Cellulose amendment +72% Fontaine et al. (2007)
Pachic Argiustoll 90–100 cm Drying-rewetting cycles +20–426% Xiang et al. (2008)
Pachic Haploxeroll 50–100 cm Temperature increase Q10 values >than those
of surface soil
Fierer et al. (2003)
Nutrient amendment +50–400% Fierer et al. (2003)
152 Plant Soil (2011) 338:143–158
horizontal direction. SOM compounds are not randomly
distributed, but most likely associated with soil struc-
ture. Pedological processes operating within a soil
profile may also lead to heterogeneous distribution of
stabilised SOM compounds. In fact, the organic C
content as well as the
14
C activity of SOM were found
to be horizontally stratisfied, within spatially distinct
regions of a Cambisol profile (Fig. 3,Chabbietal.
2009). Three main processes of SOM input into deep
soil horizons were identified to be responsible for the
heterogenous distribution of stabilised C compounds:
(1) preferential flow of DOC; (2) plant rooting
behaviour and (3) SOM transport by bioturbation.
Preferential flow paths, which were found to be
stable for decades may act as transport pathways for
young dissolved SOM into deeper soil horizons
(Hagedorn and Bundt 2002). The preferential flow
paths were also found to have increased microbial
biomass due to enhanced C and nutrient supply
(Bundt et al. 2001a). As a consequence, SOM in
such paths was found to be younger and less humified
than SOM in the soil matrix (Bundt et al. 2001b). The
transport of C into deeper soil horizons is strongly
dependent on soil texture and homogeneity of plant
cover (Chevallier et al. 2000). Preferential SOC
transport may occur in vertical cracks of clayey soil
and cause an increasing variability of SOC distribution
in subsoil horizons (Don et al. 2007).
Soil structure may also lead to a specific root
distribution within the soil profile which enhances
young carbon input into the mineral soil. Roots move
downwards through the soil, release exudates and
sloughed-off organic material, ultimately die and
decompose. This source for young C in subsoil
horizons may persist even for annual plants, as
recolonisation of ancient root channel has been
reported (Rasse and Smucker 1998). It was found
that soil C stock changes were related to fine root
length and therefore root activity rather than fine root
mass (Guo et al. 2005).
One important but neglected process leading to
heterogenous distribution of SOM is bioturbation.
Anecic earthworms for example reach 1-2 m depth
but can go down to 5 m. They transport fresh organic
detritus from the soil surface into the burrows while
mixing with the mineral soil (Lee 1985). Earthworms
from different soil horizons which showed a radio-
carbon age of several thousand years were found to
feed on young carbon (Scharpenseel and Becker-
Heidmann 1989). Therefore, they may transport fresh
OM deep into soil. This assumption is reported by a
recent study showing that SOM contents as well as
stabilised carbon are heterogenously distributed
around the earthworm borrows (Don et al. 2008).
All these observations suggest that spatial distri-
bution of old and young carbon within the soil profile
is an important issue and that the study of stabilised
SOM in subsoils is probably best approached by
visual analysis of the soil profile and adaptation of a
specific sampling scheme.
Modelling of SOM dynamics in subsoils
Models of turnover of SOM in subsoils need to
address the downward movement of C and its
stability in deeper layers. The vertical distribution of
Cand
14
C was modelled by including transport
mechanisms in SOM turnover models (O’Brian and
Stout 1978; Elzein and Balesdent 1995). In these
models,
14
C provides the best tracer of SOM
dynamics and transport. Model results indicate that
diffusion rather than convection is responsible for
most C transport to depth. Further progress was made
by including process-based information from stable C
and N isotopes (
13
C and
15
N). This study indicates
that DOC transport represents an important process in
annual grassland ecosystems (Baisden et al. 2002). In
a recent modelling exercise, Jenkinson and Coleman
(2008) intended to develop a version of the dynamic
model RothC-26.3 that can be applied to subsoils.
Their results showed that this can be achieved by
adding two parameters, one that moves organic C
down the profile and another which slows decompo-
sition at depth to the original model. However, the
processes responsible for the transport of organic
matter into subsoil horizons and the mechanisms
leading to retarded decomposition of organic matter in
subsoil horizons are poorly understood. A major
knowledge gab concerns the C fluxes in deep soil
horizons. To develop a C model for subsoils, more
accurate measurements of deep C inputs and deep C
losses are needed. More studies should be carried out
to quantify SOM input into subsoils (roots growth,
DOC perculation and bioturbation). The factors
increasing or slowing down subsoil SOM decompo-
sition need to be related to pedological characteristics
of soil. Moreover, dynamic SOM models accounting
Plant Soil (2011) 338:143–158 153
for spatial heterogeneity of total and stabilised
organic carbon within subsoil horizons need to be
developed.
Conclusion
Main C sources of subsoil OM are dissolved organic
matter, root biomass and physically or biologically
transported particulate organic matter. Organic matter
in subsoil horizons is characterized by a high
radiocarbon age but the reasons for this are not
entirely clear. Decreasing C/N ratio, increasing
δ
13
C and analysis of single OM compounds suggest,
that microbial-derived OM may contribute with
higher proportions to OM in subsoil horizons than
plant-derived OM.
It is likely that OM in subsoil horizons contains a
high proportion of C compounds stabilized by mineral
interactions. Numerous studies on acid subsoil hori-
zons reported the importance of amorphous Fe and Al
oxides for OM stabilization. Less information is
available for calcareous subsoils. Another important
stabilization mechanism in subsoil horizons may be
occlusion of OM in soil aggregates.
While we did not find literature evidence for
oxygen limitations of SOM degradation in subsoils,
it has been suggested, that unfavourable conditions
with regards to temperature, nutrients and energy
could limit the degradation of OM stored in subsoil
horizons. Most studies concerning these controlling
factors were conducted at the laboratory scale. The
evaluation of their relevance at field scale is very
limited.
The most important factor determining the degree
of alteration of OM seems to be related to the
spatially heterogenous distribution of fresh C and
the soil microbial biomass. Specific pedological
processes operating in a given soil type, which
influence this distribution (i.e. preferential flow and
bioturbation) should be taken into account, when C
dynamics in subsoils is to be studied.
To completely understand and model the dynamics
of OM in subsoils, quantitative data are necessary on
(1) the relative importance and controlling factors of C
sources, (2) the contribution of microbial versus plant-
derived OM, (3) the C fluxes from deep horizons
related to C input and output and C stabilization and
destabilization processes at field scale.
Acknowledgements Two anonymous reviewers are acknowl-
edged for their constructive comments, which greatly helped to
improve the manuscript. Additionally, we thank the organizers
of the conference on “Soil organic matter dynamics”in
Colorado Springs for financial support.
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