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Title
Age of soil organic matter and soil respiration: radiocarbon constraints on belowground C
dynamics
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Author
Trumbore, Susan
Publication Date
2000-04-01
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April 2000 399
BELOWGROUND PROCESSES AND GLOBAL CHANGE
399
Ecological Applications,
10(2), 2000, pp. 399–411
q
2000 by the Ecological Society of America
AGE OF SOIL ORGANIC MATTER AND SOIL RESPIRATION:
RADIOCARBON CONSTRAINTS ON BELOWGROUND C DYNAMICS
S
USAN
T
RUMBORE
Department of Earth System Science, University of California, Irvine, California 92697-3100 USA
Abstract.
Radiocarbon data from soil organic matter and soil respiration provide pow-
erful constraints for determining carbon dynamics and thereby the magnitude and timing
of soil carbon response to global change. In this paper, data from three sites representing
well-drained soils in boreal, temperate, and tropical forests are used to illustrate the methods
for using radiocarbon to determine the turnover times of soil organic matter and to partition
soil respiration. For these sites, the average age of bulk carbon in detrital and Oh/A-horizon
organic carbon ranges from 200 to 1200 yr. In each case, this mass-weighted average
includes components such as relatively undecomposed leaf, root, and moss litter with much
shorter turnover times, and humified or mineral-associated organic matter with much longer
turnover times. The average age of carbon in organic matter is greater than the average
age predicted for CO
2
produced by its decomposition (30, 8, and 3 yr for boreal, temperate,
and tropical soil), or measured in total soil respiration (16, 3, and 1 yr). Most of the CO
2
produced during decomposition is derived from relatively short-lived soil organic matter
(SOM) components that do not represent a large component of the standing stock of soil
organic matter. Estimates of soil carbon turnover obtained by dividing C stocks by hetero-
trophic respiration fluxes, or from radiocarbon measurements of bulk SOM, are biased to
longer time scales of C cycling. Failure to account for the heterogeneity of soil organic
matter will result in underestimation of the short-term response and overestimation of the
long-term response of soil C storage to future changes in inputs or decomposition.
Comparison of the
14
C in soil respiration with soil organic matter in temperate and
boreal forest sites indicates a significant contribution from decomposition of organic matter
fixed
.
2 yr but
,
30 yr ago. Tropical soil respiration is dominated by C fixed
,
1 yr ago.
Monitoring the
14
C signature of CO
2
emitted from soils give clues as to the causes of
seasonal and interannual variability in soil respiration in these systems.
Key words: belowground processes and global change; carbon cycle; global climate change;
radiocarbon; sink; soil organic matter; soil respiration; source.
I
NTRODUCTION
Interest in the storage and cycling of organic matter
in soils has increased recently because of its importance
to the global carbon cycle. Organic matter in detritus
and mineral soil organic matter, collectively referred
to here as SOM, is the major reservoir of carbon in
terrestrial ecosystems, storing some 1500 Pg (1 Pg
5
10
15
g) of carbon in the upper meter of mineral soils
(Jobba´gy and Jackson 2000). This is slightly more than
twice the amount of carbon present in the atmosphere
as CO
2
. However, merely knowing the size of the res-
ervoir of carbon stored in soils is insufficient for pre-
dicting its potential to influence atmospheric CO
2
con-
centrations. We must also know something about soil
carbon dynamics.
Not all of the carbon in soils interacts with atmo-
spheric CO
2
on the same time scale. Observational con-
Manuscript received 18 September 1998; revised 5 May 1999;
accepted 24 May 1999; final version received 6 July 1999. For
reprints of this Invited Feature, see footnote 1, p. 397.
straints require SOM to consist of several pools, with
characteristic turnover times of
#
yr, years to decades,
and over several hundred years (e.g., Parton et al.
1987). Most of the roughly 80–160 Pg C in surface
detritus (Matthews 1997) and an estimated 200–300 Pg
C in SOM is in forms that accumulate and decompose
on time scales of a century or less (Schimel 1995, Potter
and Klooster 1997). The remainder, constituting the
majority of carbon stored in mineral soils, is stable on
time scales of centuries to millennia.
Several issues have emerged recently to focus ques-
tions on the role of soils in the global carbon cycle
over decade-to-century time scales. The first is the role
that soils have played historically as sources or sinks
of carbon associated with changes in land management
(e.g., clearing of forest or abandonment of agricultural
land to forest regrowth; Jackson et al. 2000). In general,
forest clearing is associated with soil C loss and re-
growth is associated with soil C gain in temperate eco-
systems. Tropical soils have been observed to both gain
and lose C on forest conversion, depending on the pro-
400
INVITED FEATURE
Ecological Applications
Vol. 10, No. 2
ductivity of the subsequent pasture (e.g., Camargo et
al. 1999). Changes in C storage occur not only in upper
soil layers, but also affect distribution of organic matter
with depth (Jackson et al. 2000).
Soil C stores are predicted to respond to climatic
change because organic matter decomposition rates are
linked to soil temperature and moisture regimes. In
particular, enhanced decomposition associated with
warming of 1
8
C has been predicted to release between
11 and 30 Pg C to the atmosphere from soils (Schimel
et al. 1994). Given the general warming of nearly 0.5
8
C
over the past century, deduced from observations and
paleoclimate data, we should therefore expect soils to
presently act as a global net source of CO
2
to the at-
mosphere. Of particular interest is the response of large
stores of C in northern peatlands to potential future
warming.
Another issue of interest is the observed interannual
variability in the partitioning of fossil fuel derived CO
2
among atmospheric, terrestrial, and oceanic sinks. The
annual transfer of CO
2
from soils to the atmosphere
has been estimated at 60–80 Pg C/yr (Raich and Potter
1995), 12–16 times the annual rate of addition of fossil
fuel CO
2
to the atmosphere. If a significant portion of
that flux is derived from carbon that resides in SOM
for several years to decades, then one cause of inter-
annual variability may be due to differences in the tim-
ing of C addition to and release from soils (Fung et al.
1997).
A timely question is the present and future potential
for storage of fossil fuel derived CO
2
in soil organic
matter. With the Kyoto accords suggesting that en-
hancement of C sinks may be an acceptable way of
offsetting fossil fuel emissions (Bolin 1998), there is
increased interest in managing soils to take up carbon.
The ability of a reservoir to act as a net CO
2
sink
depends on both the fluxes of C into the reservoir and
the residence time of carbon. Small reservoirs that ad-
just quickly to enhanced C inputs are not good long-
term solutions for storing carbon, whereas large pools
that turn over more slowly may remain net sinks of
carbon long enough to partially offset fossil fuel ad-
dition to the atmosphere for a decade or two. Critical
to the ability to predict how much and how long carbon
can be stored in soils is an understanding of the resi-
dence time of carbon in SOM, and what processes de-
termine it.
Various tools have been developed to determine the
dynamics of SOM, including direct observations of C
fluxes and inventory, manipulations such as litter de-
composition experiments and
14
C labeling of substrates,
and chronosequence studies. This paper will focus on
the use of
14
C produced by atmospheric weapons testing
to study belowground C dynamics. The comparison of
14
C in soil respiration with
14
C in soil organic matter
fractions demonstrates the heterogeneous nature of soil
C, and may be used to determine the fraction of soil
respiration derived from decomposition of organic car-
bon with turnover times
.
1 yr. Examples taken from
boreal, temperate, and tropical forest soils are used to
demonstrate several issues: (1) determining turnover
time of soil C fractions, including sources of uncer-
tainty in the interpretation of radiocarbon data; (2) par-
titioning soil respiration into contributions from C re-
cently fixed from the atmosphere and C that has resided
for years to decades in soil organic matter; (3) inves-
tigating causes of seasonal and annual variability in
soil respiration; and (4) using
14
C measurements to ver-
ify whether managed systems are accumulating or los-
ing carbon relative to unmanaged counterparts.
T
HE
R
ADIOCARBON
T
RACER
Radiocarbon is useful for studying soil organic mat-
ter dynamics on two different time scales. Natural ra-
diocarbon, which is produced at an approximately con-
stant rate in the upper atmosphere during cosmic ray
bombardment, reflects the dynamics of organic matter
that has been stabilized by interaction with mineral
surfaces and resides in soils long enough for significant
radioactive decay to occur (
14
C half-life
5
5730 yr).
Discussion here, however, will focus on radiocarbon
produced during atmospheric testing of thermonuclear
weapons during the early 1960s. Tracing this global
isotopic ‘‘spike’’ for the C cycle into ocean and ter-
restrial C reservoirs provides information on exchanges
that occur on time scales of years to decades (Broecker
and Peng 1982). Radioactive decay on this time scale
is negligible, and bomb
14
C is essentially a conservative
tracer.
Although the first radiocarbon measurements in
SOM were made in the 1960s (e.g., Campbell et al.
1967), the recent advent of accelerator mass spectrom-
etry (AMS), with its smaller sample size requirements
and faster throughput capabilities, has led to a rapid
expansion in the use of
14
C to study organic matter
dynamics in terrestrial ecosystems. The use of AMS
also makes it feasible to monitor the
14
C signature of
CO
2
emitted from soils, and to sample soil atmosphere
CO
2
for isotopic measurement.
Radiocarbon data are reported as the permil (‰) de-
viation from a standard of fixed isotopic composition:
14
D
C
5
[
F
2
1]
3
1000 (1)
where
14
C
sample
12
12
C
F
5
. (2)
14
C
standard
12
12
C
The standard is selected so that
D
14
C is zero (
F
5
1.000)
for atmospheric CO
2
in 1950. Both sample and standard
are corrected for mass-dependent fractionation of iso-
April 2000 401
BELOWGROUND PROCESSES AND GLOBAL CHANGE
F
IG
. 1. Changes in the
D
14
C of atmospheric
CO
2
in the northern and southern (solid circles)
hemisphere since 1900, as recorded in air sam-
ples and wines (Stuiver and Quay 1981, Bur-
choladze et al. 1989, Manning and Melhuish
1995, Nydal and Lo¨vseth 1995). A
D
14
C value
of
1
1000‰ represents a doubling of the amount
of
14
C in atmospheric CO
2
. The southern hemi-
sphere (SH) record of bomb
14
C lags that of the
northern hemisphere (NH) due to the prepon-
derance of atmospheric testing in the northern
hemisphere and the time required to mix air
across the equator. Also shown is the expected
evolution of two steady state, homogeneous car-
bon reservoirs (in thenorthernhemisphere)with
mean carbon residence times of 6.6 and 50 yr
(dashed lines). The two lines cross in 1996.
topes, with all samples corrected to a common
d
13
C
value of
2
25‰. Radiocarbon data reported in this way
thus reflect the time since C was fixed and/or the
14
C
signature of CO
2
sources, and not isotopic discrimi-
nation during fixation or respiration of carbon. Positive
D
14
C values indicate the presence of bomb-produced
radiocarbon (
14
C activities higher than the 1950 at-
mosphere). Negative
D
14
C values indicate that the bulk
of the carbon has resided in soils long enough for sig-
nificant radioactive decay to have occurred.
Fig. 1 illustrates changes in the amount of radiocar-
bon in atmospheric CO
2
during this century. Analysis
of the
14
C content of known-age tree rings shows that,
for several thousand years prior to 1900, atmospheric
14
CO
2
levels remained roughly constant. Radiocarbon
decreased in the early part of the century because of
dilution by CO
2
derived from the burning of
14
C-free
fossil fuels. The major feature seen in Fig. 1 is a near-
doubling of atmospheric
14
C levels between 1955 and
1964, due to
14
C production by atmospheric nuclear
weapons testing. With the cessation of most atmo-
spheric tests in 1964, radiocarbon levels decreased as
the excess
14
C in the atmosphere began to be transferred
into ocean and terrestrial C reservoirs. The radiocarbon
signature of atmospheric CO
2
in the 1980s and 1990s
continues to decline at a rate of about 8‰ per year
(Levin et al. 1992) due to continued uptake by the
oceans and dilution as CO
2
derived from
14
C-free fossil
fuel is added to the atmosphere. The
D
14
CofCO
2
during
the northern hemisphere growing season in 1996 was
1
97
6
5‰ (Gaudinski et al.,
in press
). Because the
annual rate of decline is greater than the precision of
the
14
C measurement (
6
4–8‰, depending on the lab-
oratory and method used), it is possible to identify the
year in which C was fixed from the atmosphere to with-
in 1–2 years over the past 30 years.
Few other methods exist to trace carbon dynamics
on decadal time scales. Most rely on the presence of
some form of disturbance; for example, observing the
accumulation or loss of carbon associated with a
change in land management, or the change in
13
C fol-
lowing a vegetation change from C
3
to C
4
plants. The
latter method has been used extensively in tropical eco-
systems where C
3
forests are often replaced with C
4
grasses. However, it only gives information on C dy-
namics in the disturbed system. Ehleringer et al. (2000)
discuss other assumptions and limitation of using stable
isotopes to study SOM dynamics. A major advantage
of using the bomb radiocarbon tracer is that it can be
used in both disturbed and undisturbed ecosystems.
Another method proposed to study more slowly cy-
cling soil organic matter pools is long-term (
.
1-yr)
incubations (e.g., Townsend et al. 1995). The rationale
for using incubations is that all rapidly cycling organic
matter will decompose within the first year, allowing
estimation of the fraction of heterotrophic respiration
derived from remaining C, which has turnover times
.
1 yr. However, soil incubations are limited in that
they are not representative of in situ conditions. In
addition, the C not respired after one year is likely to
have a range of decomposition rates. Because those
components with the fastest decomposition rates will
dominate the production of CO
2
, care must be exercised
when calculating a decomposition rate from incubation
experiments.
Distribution of radiocarbon in SOM
The degree to which bomb
14
C may be found in soil
organic matter provides a direct measure of the degree
to which soils have incorporated C fixed from the at-
mosphere over the past
;
35 yr. This information may
be used in conjunction with other measures of C in-
ventory and flux to deduce the residence time of carbon
in different components of SOM.
Fig. 2 shows the distribution of C and
14
C with depth
for moderately well to well-drained boreal, temperate
402
INVITED FEATURE
Ecological Applications
Vol. 10, No. 2
F
IG
. 2. Carbon density (in g C/m
2
per cm depth) and radiocarbon signatures (shading) with depth for boreal, temperate,
and tropical forest soils. Zero on the depth axis is the mineral–organic boundary, negative depths are surface litter, and
positive depths are within the mineral soil. Carbon is also identified as dense fraction (
.
2 g C/cm
3
) or low-density fraction
(
,
2 g/cm
3
). See
Distribution of radiocarbon in SOM
for descriptions of sites and original references.
deciduous, and tropical forest soils, sampled 31–33 yr
after the 1964 peak in bomb
14
C. Surface detrital layers
are shown in addition to the upper portions of the min-
eral soil. In mineral horizons, SOM has been separated
by density. The dense fraction (
,
2 g C/cm
3
) is carbon
associated with soil mineral surfaces, whereas the low-
density fraction (
.
2 g C/cm
3
) is nonmineral-associated
organic matter of varying degrees of degradation. The
shading shows the radiocarbon content of the SOM in
each depth interval. High values (
.
100‰) indicate that
the soil organic matter has more
14
C than the 1996
atmosphere, and is thus dominated by C fixed
.
1yr
but
,
35 yr ago. Low values (
,
0‰) indicate that the
organic matter, on average, has resided in the soil long
enough for significant radioactive decay of
14
C to have
occurred (
.
300 yr). Intermediate values (0–100‰)
represent either C that turns over on time scales of
several decades to centuries, or a mixture of more rap-
idly (
.
100‰) and more slowly (
,
0‰) cycling ma-
terial.
The boreal soil is in a mature black spruce–feather
moss stand near Thompson, Manitoba, Canada, that
burned
;
120 yr ago (Harden et al. 1997, Trumbore and
Harden 1997). Since then,
;
30 cm of moss, root, and
leaf detritus has accumulated above a residual layer of
charred material underlain by humified organic matter
in which plant fragments can no longer be identified.
We infer from chronosequence studies that the site is
continuing to accumulate C in surface moss in 1996
(Harden et al. 1997), although this gain is offset by
carbon losses from net decomposition of deeper,
14
C-
depleted carbon (Goulden et al. 1998). The temperatey
forest soil is from the Harvard Forest in central Mas-
sachusetts, USA (Wofsy et al. 1993, Goulden et al.
1996). This mixed deciduous hardwood and conifer
forest has regrown on abandoned farm land. All of the
carbon in surface litter and the A horizon at this site
has accumulated on the top of a plowed layer since
abandonment, nearly a century ago (Gaudinski et al.,
in press
). The tropical site is a deep-rooting, seasonally
dry, but still evergreen broadleaf forest in eastern Ama-
zonia, Brazil (Trumbore et al. 1995, Camargo et al.
1999).
Although Fig. 2 shows only the upper 40 cm, the C
storage of both fast-cycling and more persistent C at
depths
.
40 cm is significant for all three soils. This
is particularly true of the tropical forest soil, which
contains as much C with decadal turnover times below
1m as in the upper 1m of soil (Trumbore et al. 1995).
Clay contents are highest in the mineral portion of the
tropical and boreal soils and low in the temperate soil.
The boreal and temperate soils are both developed on
postglacial surfaces, whereas the tropical soil is prob-
ably
.
100 000 yr old.
At each site, the total amount of carbon in soil de-
rived from the atmosphere since 1963 is on the order
of 2.5–4 kg C/m
2
, some 40–50% of the total organic
matter in the soil to 40 cm depth. The
14
C signatures
of organic matter fractions clearly differ, with most of
the bomb C in low-density organic matter (
,
2gC/
cm
3
). The abundance of bomb carbon decreases with
depth in the soil profile at all sites, although its presence
at depth may be inferred from the
14
C measured in soil
CO
2
(see
Soil respiration
).
