Global Change Biology (2007) 13, 1–12, doi: 10.1111/j.1365-2486.2007.01435.x
C4-derived soil organic carbon decomposes faster than its
C3 counterpart in mixed C3/C4 soils
J O N A T H A N G . W Y N N * and M I C H A E L I . B I R D w
*Department of Geology, University of South Florida, Tampa, FL, 33620, USA, wSchool of Geography and Geosciences, University of
St Andrews, St Andrews, Scotland, UK
Abstract
The large difference in the degree of discrimination of stable carbon isotopes between C3
and C4 plants is widely exploited in global change and carbon cycle research, often with
the assumption that carbon retains the carbon isotopic signature of its photosynthetic
pathway during later stages of decomposition in soil and sediments. We applied longterm incubation experiments and natural 13C-labelling of C3 and C4-derived soil organic
carbon (SOC) collected from across major environmental gradients in Australia to
elucidate a significant difference in the rate of decomposition of C3- and C4-derived
SOC. We find that the active pool of SOC (ASOC) derived from C4 plants decomposes at
over twice the rate of the total pool of ASOC. As a result, the proportion of C4
photosynthesis represented in the heterotrophic CO2 flux from soil must be over twice
the proportional representation of C4-derived biomass in SOC. This observation has
significant implications for much carbon cycle research that exploits the carbon isotopic
difference in these two photosynthetic pathways.
Keywords: C3-plants, C4, decomposition rate, soil organic carbon, stable carbon isotope
Received 13 February 2007 and accepted 7 May 2007
Introduction
Approximately 21–23% of current global primary productivity follows the C4 photosynthetic pathway (Lloyd
& Farquhar, 1994; Still et al., 2003; Suits et al., 2005).
Plants using this pathway occur predominantly in
tropical regions for ecophysiological reasons (Farquhar
et al., 1989), and models suggest that the fraction of total
productivity derived from C4 photosynthesis approaches unity throughout much of the arid/semi-arid
tropics (Still et al., 2003). Yet, stable carbon isotope data
from soil organic carbon (SOC) in many regions of
Australasia modelled as C4-dominated suggest a much
less significant contribution from C4 plants to the SOC
pool (Fig. 1; Bird & Pousai, 1997). Likewise, the particulate organic carbon fraction of suspended sediments
in rivers draining savanna regions of the Amazon Basin
and equatorial Africa show a relatively diminished
carbon isotopic signature of C4 plants (Bird et al.,
1991, 1998). Hence, there appears to be a significant
imbalance between the fraction of C4 carbon fixed
during photosynthesis and the fraction of C4-derived
Correspondence: Jonathan G. Wynn, tel. 11 813 974 9369,
fax 11 813 974 2654, e-mail: [email protected]
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Journal compilation r 2007 Blackwell Publishing Ltd
carbon represented in terrestrial carbon reservoirs, and
this may be partly attributable to a difference in the
behaviour of organic carbon produced by C3 and C4
plants during decomposition. This is significant because, a wide range of carbon cycle, global change
and paleoecology research makes use of the large
difference in 13C/12C ratios between C3 and C4 plants
(Cerling et al., 1989; Boutton et al., 1994; Ciais et al., 1995,
2005; Boutton, 1996; Fung et al., 1997; Koch, 1998; Archer
et al., 2000; Ehleringer et al., 2000; Guillet et al., 2001;
Randerson et al., 2002; Still et al., 2003; Krull et al., 2005;
Randerson, 2005; Suits et al., 2005) and much of
this research assumes that there is no inherent difference in behaviour between C3- and C4-derived organic
carbon once it enters the soil and sedimentary organic
carbon pools.
In this study, we applied natural abundance 13C-labelling during incubation (using C3 and C4 endmembers) to
test for a difference in the decomposition rates of C3- and
C4-derived SOC from soils collected from across the
Australian continent. We use direct observations of
13
C/12C ratios before and after from long-term soil organic
matter incubation experiments to elucidate any imbalance
between the decomposition rates of C3- and C4-derived
soil organic matter in mixed C3/C4 soils.
1
2 J. G. WYNN & M. I. BIRD
(a)
(b)
BAM
SARawak
Malaysia
CZEech
Republic
DAR
KAK
KAT
DER FIT
ANA
MUS
CHI
COE
TOP
TEN
(c)
LAS
BUN
BIR
INN
FRA
FRB
MOR
Fraction C4
1
MAC
0
CED
(d)
BOR
BIG
13CSOC
LIT
−18
400
N
Kilometers
0
400 800
GIP
GRI
STR
−32
Fig. 1 Maps of soil sampling regions with modeled stable carbon isotopic composition and C4 photosynthesis. (a) Sites in Australia
span the range of climates from tropical and temperate forests to savannas and deserts. Two sites outside Australia include tropical
rainforest in Sarawak, Malaysia and cool temperate forest in the Czech Republic. Color indicates d13CSOC of sandy soils, derived from an
empirical relationship of d13CSOC to W* (in Fig. 2) and continental W* data (from gridded climate data of the Australian Bureau of
Meteorology). (b) Fraction of C4 photosynthesis from global model of plant physiology, remote sensing, and global crop harvest spatial
data (from Still et al., 2003), compared with this study: (c) fraction of C4 photosynthesis modeled from the fraction of C4-derived soil
organic carbon. (d) Fraction of C4 photosynthesis modeled from the fraction of C4-derived ASOC.
Materials and methods
Sampling regions
We utilized a previously defined stratified sampling
approach that divides large scale ecosystem regions
(10 s of km2) into two types of soil sampling locations
locally dominated by endmember C3 or C4 vegetation
(‘tree’ and ‘grass’ samples; for details of methodology
see Wynn et al., 2006b). In this sampling regime, ‘tree’
sites are located at 1/2 canopy distance of trees, while
‘grass’ sites are located at 1/2 maximum distance
between trees. We used samples from a set of 29
ecosystem regions. Sampling regions were selected for
minimal anthropogenic disturbance, although all have
been grazed by either native species or livestock or
both. Sampling regions were selected throughout Australia (Fig. 1) so as to cover the wide range of environmental gradients across the continent, ranging from
tropical and temperate forests to deserts and savannas
(Table 1). Two sites outside Australia were also included
to extend the data set to wet tropical rainforests (Sarawak, Malaysia) and cool temperate forests (Czech
Republic).
Within each sampling region, 25 individual sampling
locations were collected to produce a ‘bulk’ sample
representative of the entire region. At each of these
sampling locations within the region, replicate samples
were taken from near trees (‘tree’ samples at 1/2 canopy
radius from trunks) and away from trees (‘grass’ samples at 1/2 maximum distance between trees). In each
soil sampling region, we took a total of 200 soil cores,
producing four ‘bulk’ samples representative of each
region (0–5 cm ‘tree,’ 0–5 cm ‘grass,’ 0–30 cm ‘tree,’ 0–
30 cm ‘grass’). Surface litter was removed when present,
and samples were taken at constant depths in steel or
PVC tubing. Samples were then air dried, and split
using a riffle splitter. The split samples were combined
such that each ‘bulk’ sample analyzed consists of
1/25th of each of the original 25 sample locations.
Laboratory methods
Where necessary (based on pH measurements on a 1 : 1
soil:water mixture), inorganic carbon was removed by
acidification with sulfurous acid. Organic carbon concentration and d13C of the CO2 produced by combustion
r 2007 The Authors
Journal compilation r 2007 Blackwell Publishing Ltd, Global Change Biology, doi: 10.1111/j.1365-2486.2007.01435.x
D E C O M P O S I T I O N R AT E S O F C 3 A N D C 4 B I O M A S S
Table 1
3
Environmental data for soil regions used in this study, and incubation period
Site
fw
MAT ( 1C)
MAP (mm yr1)
W* (mm yr1)
fo63 mm (05 cm)
t (year)
ANA
BAM
BIG
BIR
BOR
BUN
CED
CHI
COE
DAR
DER
FIT
FRA
FRB
GIP
GRI
INN
KAK
KAT
LAS
LIT
MAC
MOR
MUS
STR
TEN
TOP
CZE
SAR
0.09
0.60
0.43
0.05
0.73
0.49
0.54
0.77
0.48
0.51
0.32
0.20
0.56
0.90
0.48
0.40
0.16
0.44
0.36
0.40
0.52
0.60
0.58
0.40
0.64
0.15
0.38
0.80
0.90
27.2
26.8
15.8
23.1
16.2
21.2
17.6
26.0
26.1
27.5
27.7
27.5
21.3
21.2
14.0
12.5
21.3
26.9
27.1
21.3
14.5
18.7
20.5
25.9
11.8
24.4
26.5
7.1
26.8
443
1698
386
180
1015
1022
306
2159
1164
1475
644
570
1251
1294
657
1169
228
1249
904
245
467
1387
1308
1249
1886
380
609
588
3325
950
2616
1708
765
2835
2174
1474
2995
2400
2241
1196
1084
2519
2861
2187
3193
908
1986
1462
1115
1921
2688
2862
2305
3689
917
1051
3028
4663
0.06
0.04
0.02
0.04
0.05
0.14
0.10
0.00
0.12
0.12
0.07
0.07
0.00
0.00
0.01
0.01
0.06
0.07
0.11
0.06
0.03
0.02
0.00
0.08
0.01
0.15
0.11
0.05
0.05
3.87
3.77
4.56
4.60
4.60
4.54
4.57
3.77
4.57
4.60
3.87
3.88
4.50
4.50
4.50
4.56
4.56
4.60
4.56
4.57
4.60
4.57
4.60
4.60
4.56
4.56
4.60
4.56
3.89
fw is the fraction of woody vegetation for each region, MAT is mean annual temperature, MAP is mean annual precipitation,
W* is the index of mean annual availability of water. fo63 mm is the fraction of soil passing a 63 mm sieve, t is the incubation period.
of total SOC were measured by a combination of
dual-inlet mass spectrometry (Finnigan MAT 251,
Bremen, Germany) and elemental analysis-continuous
flow mass spectrometry (Micromass Prism III and
Finnigan Delta Plus XP; d13C 5 [{Rsample/Rstandard(VPDB)}
1] 1000%). Variance of each parameter is estimated
by a set of 20 samples from each region. These samples
were bulked from five sampling sites along five transects for each of the four sample types described above.
Each, thereby contains 1/5th of the original five sampling locations within each transect. Roots were not
removed from the soil, except those retained on a 2 mm
sieve before our analysis. All analyses were performed
on the fine fraction of air-dried soil (o2 mm).
We mathematically apportion the representation of
‘tree’ and ‘grass’ samples from 25 sampling locations
using the estimated fractional cover of woody vegetation at the 25 locations within each region. When
combined with measurements of bulk density, mass
concentration of carbon and stable isotope ratios, this
produces a robust estimate for the SOC inventory and
stable carbon isotopic composition of SOC for a given
ecosystem at this regional scale. This sampling approach also has the advantage of allowing differences
between these parameters at ‘tree’ and ‘grass’ sites to be
examined separately, while producing robust regional
estimates of SOC parameters.
In order to minimize the effects of secondary controls
on d13CASOC (within-photosynthetic-pathway variation
of d13CASOC, and carbon isotopic effects occurring during the transformation of living biomass to SOC), we
used the stratified sampling technique described above
and limited our data set to sandy soils in which the
fraction of material passing a 63 mm sieve is generally
o0.1 (Table 1). Such coarse-textured soils do not retain
the effects of 13C/12C fractionation during decomposition to the same extent as soils with significant proportions of finer particle sizes, as clay minerals stabilize the
13
C-enriched solid products of microbial decomposition
(Wynn et al., 2005).
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Journal compilation r 2007 Blackwell Publishing Ltd, Global Change Biology, doi: 10.1111/j.1365-2486.2007.01435.x
4 J. G. WYNN & M. I. BIRD
After initial analyses of the bulk SOC pool (sites and
analyses described in Wynn et al., 2006b), we selected
soil from 27 regions representing a wide range of C3/C4
plant fractions from across Australia, and two sites
collected outside the continent using similar methodology (Fig. 1). We incubated a 100 g split of four
composite samples from each region (a ‘tree’ and ‘grass’
sample from both 0–5 cm and 0–30 cm depth) in dark
conditions at laboratory temperature (not constant, but
uniform for all samples). Incubations were carried out
initially sterile polypropelene vessels (Bio-tites urine
specimen containers), ventilated with several punctures
through the cap. The soils were rewetted periodically to
field capacity, and allowed to dry uniformly at approximately 3-month intervals under laboratory temperature
(which was not held constant, but uniform for all
samples). Following approximately 4 years of this incubation cycling, all soil samples were rinsed once in
0.5 M K2(SO4) and twice in deionized water to remove
microbial biomass and the soluble products of decomposition which would have otherwise been removed by
leaching in the native soil (Vance et al., 1987). A split of
the remaining sample after incubation was then powdered for analysis of carbon concentration and d13C,
using a Costech elemental analyzer (Milan, Italy) and
Finnigan Delta Plus XP mass spectrometer. A second
split of each incubated sample was further subjected to
acid dichromate oxidation (Bird & Gröcke, 1997) in
50 mL centrifuge tubes filled with 0.1 M K2Cr2O7 in a
2 M solution of H2SO4. The samples were shaken for 72 h
in the solution. The solution was checked frequently,
with samples that had consumed the oxidizing agent
being replenished. The remaining soil material was
rinsed twice in deionized water and powdered for
carbon concentration and d13C analysis as above. The
carbon remaining after this digestion is defined as
oxidation resistant elemental carbon (OREC, Bird &
Gröcke, 1997), the relatively stable pool of SOC that
contains charcoal and other ‘black’ carbon that is unlikely to undergo significant change on short timescales.
Actual incubation time varied from 3.8–4.6 year and
these differences were taken into account in decomposition rate calculations described below.
We describe the ‘active’ pool of SOC using an analytical definition based on the mass difference between
total SOC and OREC. Assuming that OREC concentration and isotopic composition remained constant during incubation (only active SOC decomposed), we
calculated the ASOC concentration before and after
incubation by mass difference, and used these calculations to determine the amount of ASOC lost during
incubation. We then used mass and isotope balance to
calculate d13C of ASOC both before and after incubation
in addition to the d13C value of ASOC lost during
incubation, solving the following equations for four
unknowns (ASOCi, ASOCf, d13 CASOCi , d13 CASOCf ):
SOCi ¼ ASOCi þ OREC
SOCf ¼ ASOCf þ OREC
d13 CSOCi SOCi ¼ d13 CASOCi ASOCi þ d13 COREC OREC
d13 CSOCf SOCf ¼ d13 CASOCf ASOCf þ d13 COREC OREC;
where SOC, ASOC, and OREC are the carbon inventories (mg C cm2) for a given depth interval and subscripts i and f refer to before and after incubation.
Although this division of SOC pool structure differs
from the definition of active, intermediate and passive
pools of SOC in the CENTURY model (Parton et al., 1994),
our division follows that used in 14C models of SOC
turnover, which separate SOC into two simplified pools:
active and passive (Hahn & Buchmann, 2004).
We calculated relative contribution of C3 and C4
biomass to the SOC and ASOC pools before and after
incubation using a mixing model based on average
endmember values of C3 and C4 biomass
(25.4 2.7% and 12.5 1.1% for C3 and C4 vegetation). The SOC and ASOC pools were thus divided
proportionally into C3 and C4 components:
SOCC4;i þ SOCC3;i ¼ SOCi
ASOCC4;i þ ASOCC3;i ¼ ASOCi
d13 CSOCC4;i SOCC4;i þ d13 CSOCC3;i SOCC3;i
þ d13 CSOC SOC d13 CASOCC4;i ASOCC4;i
þ d13 CASOCC3;i ASOCC3;i þ d13 CASOC ASOC:
The subscripts C3 and C4 refer to the fraction of SOC
or ASOC derived from C3 and C4 photosynthesis.
The rates of decomposition of individual pools were
calculated from pre- and postincubation data (inventory and d13C before and after incubation). We used a
model of first-order decomposition to calculate the rate
constant for carbon (C) in the bulk SOC and ASOC
pools, as well as the C3 and C4 components of these two
pools, resulting in mean decomposition rate constant
(k), for a given pool (SOC, ASOC, C3, C4, etc.) over the
incubation period (t):
dC
¼ kC
dt
Cf ¼ Ci ekt :
We used calculations from annual climate data compiled for Australia to examine the climatic controls on
the initial d13C values of SOC pools using an index of
water availability (W*; Wynn et al., 2006b):
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D E C O M P O S I T I O N R AT E S O F C 3 A N D C 4 B I O M A S S
−16
(c)
OREC
>90% C4
−14
13C before incubation
(b)
Active SOC pool
−12
−20
0–5 cm
0–30 cm
mixed C3–C4
−12
>90% C4
>90% C4
0–5 cm
0–30 cm
−18
Total SOC pool
−14
0–5 cm
0–30 cm
mixed C3–C4
mixed C3–C4
−16
−18
−20
−22
−22
−24
−24
−26
−26
13C before incubation
(a)
5
−28
−28
−30
> 90% C3
> 90% C3
> 90% C3
−30
−32
−32
0
1000
2000
3000
4000
Index of annual available water (W*)
0
1000
2000
3000
4000
Index of annual available water (W*)
0
1000
2000
3000
4000
Index of annual available water (W*)
Fig. 2 Climatic control on the d13C of active soil organic carbon (SOC) (a), oxidation resistant elemental carbon (b) and total SOC (c).
Index of annual availability of water is equal to mean annual precipitation minus the amount of water that would evaporate given the
annual flux of global solar radiation at the surface, plus 4000 mm (the value of maximum evaporation with no precipitation). Error bars
are 1s of d13C from bulk SOC within five transects of five sampling locations in each region (total of 25 locations). d13C representative of
o1% and 480% C4-derived ASOC are shown in gray (based on mass balance equations and average d13C values of plants described in
the text).
Total SOC pool
Qs
þ Emax ;
W ¼ MAP rL
−12
where MAP is mean annual precipitation rate
(mm yr1), Qs is mean annual global solar radiation
(J m2 yr1), rw is density of liquid water (1000 kg m3
at 25 1C), and L is the latent heat of evaporation of water
(2.5 106 J kg1 H2O at 25 1C, and Emax is the maximum annual potential evaporation rate of water given
the maximum annual global solar radiation at earth’s
surface (1 1010 J m2 yr1) with no precipitation
(Emax 5 4000 mm yr1).
−16
13CSOC after incubation
−14
−18
−20
> 90% C3
mixed C3–C4
>90%
C4
0–5 cm
0–30 cm
−22
−24
−26
−28
Results
Stable carbon isotope values of SOC within this continental-scale dataset shows a strong relationship to the
index the annual availability of water (W*; Fig. 2). This
relationship also holds well for the active pool of SOC
for the ecosystem regions measured (d13CASOC; Fig. 2a).
In our entire dataset, d13CSOC does not reach values
representative of a 480% contribution from C4 biomass
(o17.5% for all ecosystem regions, Fig. 2c). Although
ASOC is generally more 13C-enriched than bulk SOC
(Fig. 2a, reaching values of nearly 14%), only one
ASOC region from the 11 regions modelled to have
480% C4 photosynthesis shows an initial d13CASOC
value that indicates such a high proportion of C4derived ASOC (Figs 1 and 2c). d13COREC shows a
relatively weak relationship to W*. Some d13C values
−30
−32
−32 −30 −28 −26 −24 −22 −20 −18 −16 −14 −12
13CSOC before incubation
Fig. 3 Comparison of d13C of the total soil organic carbon
(SOC) pool before vs. after incubation. 1 : 1 line, error bars and
d13C values representative of the fraction of C4-derived SOC as
in Fig. 2.
of OREC do reach a similar degree of 13C-enrichment to
those of the bulk SOC pool (18%, Fig. 2b).
Between 5% and 42% of the initial bulk SOC decomposed during the incubations. Using first order decomposition, we determined a range of 8–77 years for the
mean turnover time (1/k) of individual samples. For the
majority of C3-endmember soils (those with d13CSOC
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Journal compilation r 2007 Blackwell Publishing Ltd, Global Change Biology, doi: 10.1111/j.1365-2486.2007.01435.x
6 J. G. WYNN & M. I. BIRD
−12
−14
−16
13COREC
−18
−20
> 90% C3
mixed C3–C4
>90%
C4
0–5 cm
0–30 cm
−22
−24
−26
−28
−30
−32
−32 −30 −28 −26 −24 −22 −20 −18 −16 −14 −12
13CASOC before incubation
Fig. 4 Comparison of d13C of oxidation resistant elemental
carbon (OREC) vs. the d13C of the active pool of soil organic
carbon (SOC) before incubation. 1 : 1 line, error bars and d13C
values representative of the fraction of C4-derived SOC as in
Fig. 2.
o24.1% indicative of 490% C3 vegetation), there
was no significant change in d13CSOC over the course
of the incubations (Fig. 3). However, most soils with an
initial d13CSOC indicative of mixed C3–C4 vegetation
(424.1%) exhibited a significant negative shift in
d13CSOC of 1–3% for the bulk SOC pool (the sum of
ASOC and OREC).
OREC remaining in the soil after acid dichromate
oxidation retains a d13COREC value that broadly reflects
the d13C value of ASOC before incubation (Fig. 4).
However, there is generally a negative offset of several
per mil, notably in mixed C3–C4 environments, in
which OREC is up to 16% more 13C-depleted than
ASOC. d13COREC of C3-endmember soils generally
shows a negative offset of up to 4%, although in some
C3-endmember soils, OREC is slightly more 13C enriched than ASOC.
We calculated that between 5% and 99% of the initial
ASOC decomposed in individual samples incubated
(between 5% and 78% for regional averages), producing
a range of 1–67 years for the mean turnover time of
ASOC. Similar to the measurements for bulk SOC
described above, we found that for the majority of C3endmember soils, there was no significant change in
d13CASOC during incubation (Fig. 5a). However, most
soils with an initial d13CASOC indicative of mixed C3–C4
vegetation (424.1%) exhibited a significant negative
shift in d13CASOC of the ASOC pool that ranged from
1 to 12%. Using mass and isotope balance, we calcu-
lated the mass of ASOC lost, and the d13 CCO2 of the CO2
respired during incubation. Although there is more
scatter in the d13 CCO2 data for C3-endmember soils,
samples from mixed C3/C4 environments consistently
show 13C-enriched respiration as compared with the
d13C of initial ASOC (Fig. 5b). When separated between
‘tree’ and ‘grass’ sample sites that are used to determine
the spatial average values in Fig. 5a, our results show
that this shift occurs predominantly at the grass sites in
which C4-derived SOC is most abundant (Fig. 6). The
carbon isotopic shift that occurred during incubation
(Dd13CASOC, initial – final) shows is more pronounced in
soils with abundant C4-derived SOC (Fig. 7a). There is
also a strong relationship between the amount of carbon
lost during incubation, and the degree to which the
d13CASOC shifted during incubation (Fig. 7b). Many (but
not all) of the samples that showed the most pronounced d13CASOC shift also show the most significant
ASOC lost during incubation.
Using the model of first order decomposition of the
C4 and total pools of ASOC, we calculated that on
average, C4-derived ASOC decomposed at over twice
the rate of total ASOC in the mixed C3/C4 soils that we
analyzed (Table 2; kC4ASOC/kASOC ranges from 0.9
to 6.8, and the averages are 2.7 and 2.3 for the 0–5 and
0–30 cm pools, respectively). We also calculated a ratio
of the fraction of C4 biomass represented in the flux of
heterotrophic decomposition produced during incubation to the fraction of C4 biomass initially present in the
ASOC pool (fC4flux/fC4ASOC). This ratio ranges from
0.9 to 4.6 while the averages are 2.0 and 2.2 for the 0–5
and 0–30 cm pools, respectively.
Discussion
Stable carbon isotopic composition of SOC, ASOC, and
OREC
We interpret the relationships between W* and stable
carbon isotopic composition of SOC and ASOC in Fig. 2
predominantly as the result of the more efficient use of
available water by C4 plants in the water-limited environments typical of most of Australia (Farquhar et al.,
1989). C4 plants, and thus C4-derived SOC and ASOC
are increasingly abundant under water-stressed conditions of low W*. However, in our field-based data set,
the most 13C-enriched d13CSOC values are all more 13Cdepleted than 17.5%, even for ‘grass’ sampling locations, and never extend not as 13C-enriched as typical
pure C4 biomass (12.5%), or even the 13C-depleted
end member of C4 biomass (13.6%). Our field-based
data from the OREC pool of native soils confirm observations that biomass becomes progressively 13C-depleted during natural pyrolysis due to the selective loss
r 2007 The Authors
Journal compilation r 2007 Blackwell Publishing Ltd, Global Change Biology, doi: 10.1111/j.1365-2486.2007.01435.x
D E C O M P O S I T I O N R AT E S O F C 3 A N D C 4 B I O M A S S
(a)
7
(b)
Active SOC pool
−12
−12
−14
−14
13CASOC after incubation
−16
> 90% C3
mixed C3–C4
>90%
C4
> 90% C3
−16
−18
−18
−20
0–5 cm
0–5 cm
0–30 cm
0–30 cm
−20
−22
−22
−24
−24
−26
−26
>90%
C4 −28
−28
mixed C3–C4
−30
−30
−32
−32 −30 −28 −26 −24 −22 −20 −18 −16 −14
13CASOC before incubation
−32
−30 −28 −26 −24 −22 −20 −18 −16 −14 −12
13CASOC before incubation
Fig. 5 (a) Comparison of d13C of the active pool of soil organic carbon (SOC) (ASOC) before vs. after incubation. (b) Comparison of
d13C of CO2 respired during incubation to d13C of ASOC before. 1 : 1 line, error bars and d13C values representative of the fraction
of C4-derived SOC as in Fig. 2.
of thermally labile and 13C-enriched compounds such
as carbohydrates, and the selective carbonization of 13Cdepleted lignin-based compounds (Czimczik et al.,
2002; Krull et al., 2006; Turney et al., 2006). Figure 4
demonstrates a systematic difference between the d13C
values of ASOC and OREC, with OREC typically being
more 13C depleted by several per mil, and in some cases
as 13C-depleted as nearly 32%. In addition to the
potential of selective natural pyrolysis of 13C-depleted
compounds within the soil organic matter, this may also
be combined with the result of differential loss of C4derived OREC, which is typically finer grained than
woody C3 OREC, and hence more susceptible to loss by
aeolian transport (Bird & Cali, 1998).
A difference in the decomposition rates of C3- and
C4-derived SOC
After approximately 4 years of these incubation experiments, our d13C measurements of bulk SOC and ASOC
from C3-endmember soils show little to no evidence of
selective decomposition of individual organic compounds with variable 13C-content from within the total
C3-derived organic matter pool (Figs 3 and 5). This
effect was proposed by (Boutton, 1996) to explain many
changes in carbon isotopic composition during decom-
position, which would tend to selectively preserve 13Cdepleted stable compounds such as lignin and lipids at
the expense of more labile 13C-enriched components
such as carbohydrates and amino acids from the same
plants (Deines, 1980). The lack of a significant change in
the d13C values of these C3-endmember soils also likely
excludes the effect of selective preservation of 13Cenriched microbial biomass and/or solid products of
decomposition (Šantrůčková et al., 2000). This effect
might be expected for fine-textured soils in which the
solid products of microbial decomposition may be
bound by interaction with fine mineral particles (Wynn
et al., 2005, 2006a). In fact, controlling for this effect was
one aim of our sampling methodology. By limiting this
analysis to coarse-textured soils, we aimed to exclude
the effects of clay–humus complexation (Stevenson,
1978), and the selective preservation of 13C-enriched
compounds of humus in this state. The difference
between the trends of d13CSOC before and after incubation of C3-endmember soils compared to mixed C3/C4
soils indicates a role of the proportion of C4 biomass in
determining SOC decomposition dynamics.
While d13C values of C3-endmember soils do not
significantly change, our results of both SOC and ASOC
from mixed C3/C4 soils consistently indicate a pronounced negative shift in d13CSOC, which we interpret
r 2007 The Authors
Journal compilation r 2007 Blackwell Publishing Ltd, Global Change Biology, doi: 10.1111/j.1365-2486.2007.01435.x
8 J. G. WYNN & M. I. BIRD
Active SOC pool
−12
−14
> 90% C3
13CASOC after incubation
−16
mixed C3–C4
>90%
C4
−18
−20
0–5 cm tree
0–5 cm grass
−22
−24
−26
−28
−30
−32
−32 −30 −28 −26 −24 −22 −20 −18 −16 −14 −12
13 CASOC before incubation
Fig. 6 Comparison of d13C of the active pool of soil organic
carbon (SOC) pool before vs. after incubation (as in Fig. 5a), with
sampling locations separated by ‘tree’ and ‘grass’ sites (note that
‘grass’ sampling locations in C3 woodlands or forests are defined
at midpoints between trees – true grasses may not be present).
These data are mathematically apportioned according to the
fractional canopy cover to produce the regional estimates in
Fig. 5a. d13C values representative of the fraction of C4-derived
SOC as in Fig. 2. Error bars as in Figs 2–5 removed for clarity.
as a decrease in the proportion of C4-derived bulk SOC
at the end of the incubation (Figs 3 and 5). All ecosystem
regions with SOC within a range of initial d13CSOC
values typical of C3/C4 ecosystems (24.1% to
13.8%, 10–90% C4 plants) result in significantly more
13
C-depleted d13C values after incubation. Because this
shift in d13CSOC occurs in mixed C3/C4 soils, and we
find the d13CSOC of C3-endmember soils did not significantly change, we interpret this shift to indicate a
relative decrease in C4-derived SOC compared with C3derived SOC than was the case before incubation.
One possible explanation for the shift in d13CSOC
during incubation could be a relative enrichment of
13
C-depleted OREC in the soil, as the active pool of
ASOC decomposes. However, our measurements also
provide a means of testing for inherent differences
between degradability of C3- and C4-derived carbon
within the active SOC pool (ASOC), because we have
separately analyzed and removed the influence of the
selective preservation of 13C-depleted OREC. Comparison of d13C values of the ASOC pool before and after
incubation shows that the active pool of SOC also
shows the effect of selective decomposition of C4derived materials (Fig. 5a). The pronounced negative
shifts in d13CASOC from mixed C3/C4 soils in the
absence of a shift from C3-endmember soils indicate
that a greater proportion of C4-derived ASOC decomposed in the mixed C3/C4 soils. Several of the mixed
C3/C4 soils yielded final d13CASOC values that reflect
(b)
Active SOC pool
(a)
11
11
mixed C3–C4
9
(initial–final)
ASOC
∆13C
7
7
0–5 cm
0–5 cm
0–30 cm
0–30 cm
5
5
3
3
> 90% C3
1
1
−1
∆13CASOC (initial–final)
9
>90%
C4
−30 −28 −26 −24 −22 −20 −18 −16 −14
13CASOC before incubation
−1
0.20
0.40
0.60
0.80
Fraction ASOC remaining (F)
Fig. 7 Change in d13C of the active pool of SOC (ASOC) pool (initial–final), with respect to the initial d13C of ASOC (a) and the fraction
of ASOC remaining after incubation (b). d13C values representative of the fraction of C4-derived soil organic carbon as in Fig. 2. Error
bars removed for clarity.
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Journal compilation r 2007 Blackwell Publishing Ltd, Global Change Biology, doi: 10.1111/j.1365-2486.2007.01435.x
D E C O M P O S I T I O N R AT E S O F C 3 A N D C 4 B I O M A S S
Table 2 Decomposition rate constants and ratios for mixed
C3/C4 ecosystem regions
1/kASOC
(year)
1/kC4–ASOC
(year)
kC4–ASOC/
kASOC
fC4–flux/
fC4–ASOC
Site
0–
0–
0–
0–
0–
0–
0–
0–
5 cm 30 cm 5 cm 30 cm 5 cm 30 cm 5 cm 30 cm
ANA
BIR
CED
DER
FIT
INN
KAT
LAS
MUS
TEN
TOP
Ave.
1 s
5
30
16
9
12
2
1
8
12
14
11
11
8
1
10
31
16
19
4
4
15
14
24
24
15
9
*
1
5
8
*
2.2
2.2
4.0
*
*
2.8
*
2.4
1.2
3
12
5
11
13
7
4
*
1.5
3.1
2.0
0.9
2.7
1.9
1.2
1.2
3.0
2.1
1.9
2.3
0.9
*
5
2
*
6
4
7
12
6
3
*
6.8
3.4
5
5
7
2.0
4.6
2.5
1.9
3.0
1.3
1.0
1.3
2.2
1.7
0.9
2.0
1.0
1.1
1.7
3.2
5.1
2.4
2.5
1.1
1.2
2.2
1.9
1.8
2.2
1.2
k is the first-order decomposition rate constant for the laboratory incubations (shown here for total ASOC, and for C4derived ASOC of 0–5 and 0–30 cm samples). ASOC indicates
the entire active soil organic carbon pool, ‘C4–ASOC’ indicates
only the C4 fraction of this pool. Asterisks indicate samples
with d13CASOC values after incubation indicating a complete
loss of C4-derived ASOC. fC4 is the fraction of C4-derived
biomass represented, with subscripts for the heterotrophic flux
from the incubation (flux) and for the total ASOC pool
(ASOC). The ratio shown in the final two columns is thus
the ratio of C4-derived CO2 flux to C4-derived ASOC. Only
sites with differences of d13CASOC before and after incubation
40.3% are shown, i.e. those which fall off the 1 : 1 line
( 0.3%) in Fig. 5a.
a nearly complete loss of C4-derived ASOC, with final
d13CASOC values being more 13C-depleted than 24.1%.
This interpretation is supported by the relationship between the shift in d13CASOC values (initial – final) and the
amount of carbon lost during incubation (Fig. 7b) and the
proportion of C4-derived ASOC originally present (Fig.
7a). Samples with the most pronounced d13CASOC shifts
display the combined features of a substantial loss total
ASOC during incubation and a high proportion of the
initial ASOC being derived from C4 plants. Similarly, this
interpretation is supported by the fact that the sites with
significant d13CASOC shifts (mixed C3/C4 soils) show
more significant shifts for samples collected at the ‘grass’
sampling locations, as opposed to the ‘tree’ sampling
locations. This occurs for similar reasons – the relative
proportion of C4-derived ASOC is initially higher at
‘grass’ sampling locations and in general, more ASOC
was lost from the samples at these locations (open
squares in mixed C3–C4 field in Fig. 6 are more displaced
from the 1 : 1 line than closed squares).
9
We have previously noted that bulk measurements of
d13CSOC and d13CASOC from this data set are consistently
more 13C-depleted than is otherwise modelled for the
productivity of C4 plants at the sampling regions (Fig. 1),
or from observations of the C4 biomass present in the
regions sampled. We hypothesized several potential mechanisms for the selective preservation of C3-derived
biomass in the SOC pool: (1) greater input to SOC from
larger root systems of C3 plants (which are mostly woody
shrubs and trees in the tropics), (2) recent widespread
ecological change resulting in less C4 biomass (woody
weed thickening) caused either by burning and/or grazing (Archer et al., 2000) or by the CO2 fertilization effect,
which favours C3 over C4 photosynthesis (Berry &
Roderick, 2002), (3) greater preservation of C3-derived
biomass in the relatively stable pool of OREC or (4) a
difference in rates of decomposition of C3- and C4derived SOC due to differences in the quality of organic
matter (attributable to differences in lignin content and/
or mean particle size). Our results here identify (4) as the
primary mechanism that can account for our observations. Mechanisms (1) and (2) are discounted as they
derive from changes in the rates of input from C3 and C4
biomass and there were no inputs to ASOC during our
incubation experiments. This depends on the assumption
of homogeneity of the ASOC pool, which in our analysis
includes fine roots (o2 mm), and that differences in
decomposition rates are not due to differences in the
proportions of C3 and C4 fine roots. Because we have
isolated and removed the potential effects of selective
preservation of C3-derived OREC, we can eliminate
mechanism (3) in our observations of the ASOC pool.
From these observations, we conclude that there exists a
fundamental difference in the rates of decomposition of
C3- and C4-derived ASOC, although factor (3) may
account for a portion of the observed imbalance observed
in the bulk SOC pool.
We exclude the possibility of several other factors that
might be interpreted to have caused the observed
negative shifts in d13CASOC during the incubation period, such as selective decomposition of 13C-enriched
labile organic components of SOC [such as carbohydrates and amino acids, which can be enriched by
several per mil with respect to entire plants (Deines,
1980)]. In order to interpret this effect as a dominant
factor controlling our observations, we would need a
mechanism to account for the negative d13C shift occurring only in mixed C3/C4 soils and not C3-endmember
soils (Figs 3 and 5), and for the shift occurring primarily
at ‘grass’ sampling locations (Fig. 6). Furthermore,
recent studies of 13C-labelled lignin monomers have
shown that some lignin may decompose faster than
bulk SOC (Dignac et al., 2005). It is likely that variations
in the proportions of organic compounds present in C3
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Journal compilation r 2007 Blackwell Publishing Ltd, Global Change Biology, doi: 10.1111/j.1365-2486.2007.01435.x
10 J . G . W Y N N & M . I . B I R D
and C4-derived litter are a primary control on their
inherently different decomposition rates (cf. Meentemeyer,
1978; Melillo et al., 1982). However, within-pathway selective decomposition of 13C-enriched/depleted organic compounds is excluded. It is also possible that changes in
microbial communities and nitrogen availability during
the laboratory incubation may have affected the patterns
of d13CASOC during these incubation experiments, as observed by (Waldrop & Firestone, 2004). However, we do
not identify a mechanism by which these effects would
exclusively cause a shift in of mixed C3/C4 soils, and
predominantly at ‘grass’ sampling localities, except for the
proposed mechanism relating to an inherent difference in
the organic chemistry of C3 and C4 biomass. We also
intentionally excluded the effects of stabilization of the 13Cenriched products of microbial decomposition (Šantrůčková et al., 2000), by reducing our analysis to coarsetextured soils, which do not effectively retain this signature
which would have the opposite sense of our observed shift
in d13C of mixed C3/C4 SOC.
In order to quantify the interpreted difference between
the average decomposition rates of C3- and C4-derived
ASOC, we used a simple first-order decomposition model
and calculated the decomposition rate constant (k) of
ASOC using measurements of the fraction of each pool
remaining and the period of incubation (Table 2, values
shown as 1/k, ranging from 1–31 year). We also mathematically separated the ASOC pool into proportions of
C3- and C4-derived components, using a mixing model of
C3- and C4-endmember biomass, and similarly calculated
decomposition rate constants for each of these components
(1/k ranging from 1 to 13 year for C4 biomass). Although
these laboratory experiments show a difference in the
intrinsic decomposition rates of C3- and C4-derived
SOC, it is important to note that this may differ from real
field conditions. Recognizing that for a variety of reasons,
the decomposition dynamics in our controlled laboratory
experiments do not represent field decomposition rates,
we examined the ratio of the decomposition rate constant
for C4-derived ASOC to the rate constant for total ASOC.
Under these assumptions, we interpret that over the
course of the incubations, on average, C4-derived ASOC
decomposes approximately twice as fast as total ASOC
(Table 2). Similarly, there is over twice as much C4-derived
biomass represented in the CO2 flux as was represented in
the total ASOC pool before incubation (Table 2; Fig. 5b).
We conclude that C4 biomass is overrepresented in the
heterotrophic respiration during laboratory incubation in
the absence of fresh input of organic matter.
Implications for carbon cycle science
It is well-known that SOC consists of a spectrum of
organic compounds and pools, with an equally broad
spectrum of turnover times (Trumbore et al., 1996;
Davidson et al., 2000; Trumbore, 2000). Our results here
show that C4-derived materials are more abundant in
the labile components of this spectrum, while the more
stable components of ASOC are biased toward C3derived materials. This result has significant implications for carbon cycle science and paleoecology studies
that depend on measurements or estimates of the stable
carbon isotopic composition of the bulk SOC pool or
that of CO2 produced by heterotrophic decomposition.
Double deconvolution models of atmospheric carbon
budgets used to identify sources and sinks of CO2
depend heavily on estimates of carbon isotopic disequilibrium between terrestrial reservoirs and of fluxes from
those reservoirs (Ciais et al., 1995, 2005; Still et al., 2003;
Randerson, 2005). Inversion studies use atmospheric
measurements of CO2 concentration and d13C values
combined with models of land–air and sea–air gas
exchange to partition the surface exchange of atmospheric CO2 into component uptake by the terrestrial
biosphere and oceans (Ciais et al., 1995, 2005; Fung et al.,
1997; Randerson, 2005). Models of land–air CO2 exchange depend heavily on estimates of the contribution
from C4 ecosystems (Lloyd & Farquhar, 1994; Still et al.,
2003; Suits et al., 2005), as C4 plants differ significantly
from C3 plants in their discrimination against 13CO2
during photosynthesis (Farquhar et al., 1989). Thus,
accurate representation of the land–air flux requires
accurate estimates of the spatial and temporal patterns
of net carbon assimilation and respiration by C4 vegetation. One of the most poorly constrained components of
these models is the pattern of spatio-temporal shifts
between C3 and C4 vegetation (Randerson, 2005) due to
changes in land use, climate and atmospheric CO2
concentration. There is recognition that ecosystem respiration should be separated into autotrophic and
heterotrophic components in carbon cycle models (Tang
& Baldocchi, 2005), and we suggest these will differ
substantially in their d13C values, not only due to the
terrestrial Suess effect but also due to differences in the
rates of decomposition of C3 and C4 derived biomass.
Our results suggest that there will be a variable lag in
the heterotrophic respiration response of mixed C3–C4
ecosystems, such that the heterotrophic component of
ecosystem respiration will be biased towards the fastcycling C4 pool following a shift from C3 to C4 biomass.
Our calculations indicate that the fractional representation of C4-biomass in the heterotrophic CO2 flux from
soil to atmosphere is more than twice the C4 representation in the ASOC pool from mixed C3/C4 soils
collected across the Australian continent. We suggest
that if these laboratory-based results can be extended to
field conditions, and the observations from mixed
C3/C4 environments from Australia can be extended
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Journal compilation r 2007 Blackwell Publishing Ltd, Global Change Biology, doi: 10.1111/j.1365-2486.2007.01435.x
D E C O M P O S I T I O N R AT E S O F C 3 A N D C 4 B I O M A S S
to global models, the reduced rate of decomposition of
C3-derived SOC in mixed C3/C4 soils may produce a
more significant sink of 13C-depleted C3-derived carbon
in tropical soils than may otherwise be accounted for in
current models. Although we do not quantitatively
extend our interpretations to global mass and isotope
balance models of the atmosphere, these observations
would suggest an increase the magnitude of the deforestation disquilibrium term (Ddef) in mass balance
equations (Ciais et al., 2005), and thereby diminish the
magnitude of the ‘extra’ sink (Fb) required to balance
atmospheric CO2 observations.
A second area of research that will be impacted by
these results has used the use of variation in the d13CSOC
from depth profiles of soils to interpret temporal
changes in the proportion of C4 photosynthesis (Boutton et al., 1994; Boutton, 1996; Guillet et al., 2001; Jackson
et al., 2002; Krull et al., 2005). These studies assume
equivalence between the fraction of C4 photosynthesis
and its proportional representation in the SOC pool,
while we demonstrate that C4-derived ASOC decomposes at least twice as fast as the total ASOC pool. The
difference in the average turnover time for C3- vs. C4derived carbon in bulk SOC must be even greater than
we have observed for the ASOC pool, given the 13Cdepleted composition of OREC (Fig. 4) and the known
recalcitrance of the OREC pool (Bird et al., 1999). The
dominant interpretation of paleoecological studies of
C3–C4 plant distribution has been the documentation
of woody weed thickening, which is well established
from other lines of evidence (Archer et al., 2000) and
has a theoretically sound causal mechanism (Berry &
Roderick, 2002). However, our results suggest that the
magnitude and timing of thickening will be overstated
by simple linear d13C mixing models.
Third, our observations may also account for the field
observations of increases in SOC storage in arid and
semi-arid climates invaded by woody (C3) weeds (Jackson et al., 2002). As C3-derived organic carbon cycles
more slowly through the soil pool, we would expect to
be able to extend this observation to known increases in
SOC inventory in sandy soils invaded by C3 vegetation.
In other soils with significant fine particles, other processes such as overprinting by stabilization of some
components of the SOC pool would need to be considered in addition to a raw difference in turnover times.
Fourth, paleoenvironmental interpretations of the balance between C3 and C4 vegetation based on d13C from
organic matter preserved in paleosols (Cerling et al., 1989)
and in riverine and oceanic sediments (Bird et al., 1998)
must also take into account the relative rates of decomposition demonstrated here in order to avoid overestimation of the fraction of C3 productivity in the past. As
organic matter matures in the absence of fresh input of
11
biomass, we would expect an increase in the representation of C3-derived SOC at the expense of more labile C4derived SOC, which may not be well represented in
paleosol organic carbon or sedimentary organic carbon.
Finally, we demonstrate that SOC sequestration potential per unit productivity must be inherently lower
for C4 ecosystems. This has global implications for
carbon cycle science and policy regarding replacement
of tropical forests with pasture. Not only do ecological
changes from C3 to C4 systems lower primary productivity, we show here that biomass derived from tropical
C4 grasses will cycle through the soil faster than that
derived from trees, leading to lower stocks of SOC in
C4 ecosystems per unit of primary productivity.
Acknowledgements
We thank the Australian Cooperative Research Centre for Greenhouse Accounting for funding field work, and the labs of the
Earth Environment Group at the Research School of Earth
Sciences, Australian National University and the FEEA Stable
Isotope Laboratory at the University of St Andrews for analytical
support. The assistance of Lins Vellen, Youping Zhou, Delphine
Derrien, Joe Cali, Emilie Grand-Clement, Elliott Dale and Angus
Calder greatly facilitated the collection and analysis of the nearly
1.2 km of soil core collected. Lins Vellen caringly supervised the
incubating microorganisms.
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