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Acknowledgements.
This work was supported by the National Science Foundation Materials Research
Science and Engineering Center at the University of Massachusetts.
Correspondence and requests for materials should be addressed to D.A.T. (e-mail: tirrell@polysci.umass.
edu).
Mineral control of soil organic
carbon storage and turnover
Margaret S. Torn
*
, Susan E. Trumbore
*
,
Oliver A. Chadwick
†
, Peter M. Vitousek
‡
& David M. Hendricks
§
*
Earth System Science Department, University of California, Irvine,
California 92697-3100, USA
†
Department of Geography, University of California, Santa Barbara,
California 93106, USA
‡
Department of Biological Sciences, Stanford University, Stanford,
California 94305-5020, USA
§
Department of Soil and Water Science, University of Arizona, Tucson,
Arizona 85721, USA
. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .
A large source of uncertainty in present understanding of the
global carbon cycle is the distribution and dynamics of the soil
organic carbon reservoir. Most of the organic carbon in soils is
degraded to inorganic forms slowly, on timescales from centuries
to millennia
1
. Soil minerals are known to play a stabilizing role,
but how spatial and temporal variation in soil mineralogy con-
trols the quantity and turnover of long-residence-time organic
carbon is not well known
2
. Here we use radiocarbon analyses to
explore interactions between soil mineralogy and soil organic
carbon along two natural gradients—of soil-age and of climate—
in volcanic soil environments. During the first
,
150,000 years of
soil development, the volcanic parent material weathered to
metastable, non-crystalline minerals. Thereafter, the amount of
non-crystalline minerals declined, and more stable crystalline
minerals accumulated. Soil organic carbon content followed a
similar trend, accumulating to a maximum after 150,000 years,
and then decreasing by 50% over the next four million years. A
positive relationship between non-crystalline minerals and
organic carbon was also observed in soils through the climate
gradient, indicating that the accumulation and subsequent loss of
organic matter were largely driven by changes in the millennial
scale cycling of mineral-stabilized carbon, rather than by changes
in the amount of fast-cycling organic matter or in net primary
productivity. Soil mineralogy is therefore important in determin-
ing the quantity of organic carbon stored in soil, its turnover time,
and atmosphere–ecosystem carbon fluxes during long-term soil
development; this conclusion should be generalizable at least to
other humid environments.
Our primary gradient was a chronosequence of six sites (300 yr to
4,100 kyr) located at the same elevation and climate in areas
dominated by native
Metrosideros polymorpha
trees and
Cibotium
tree ferns (Table 1)
3,4
. The soils formed from volcanic ash overlying
either lava or a lava-ash mixture. Soil organic carbon and radiocarbon
100
80
60
40
20
0
0
10
20
30
40
50
Soil carbon (%)
0.3 kyr
2.1 kyr
20 kyr
150 kyr
1,400 kyr
4,100 kyr
Substrate age
100
80
60
40
20
0
-1,000
-600
-200
200
No decay
1,000 yr
27,000 yr
170,000 yr
27,000 yr
37,000 yr
Turnover time
of deepest
horizon
Depth (cm)
∆
14
C of SOM (
‰
Depth (cm)
a
b
)
Figure 1
Chronosequences:
a
, carbon in soil organic matter (SOM) versus depth;
b
,
D
14
C of soil organic matter versus depth. Key for both panels in
a
. The depth
shown is the midpoint of each horizon. See Methods for details of analysis and
modelling. For the four youngest sites, modelled SOM turnover items are slow
relative to substrate age. They indicate that the soils are not at steady state and
are still accumulating C. At 300 yr, carbon and
14
C content were best explained by
assuming a constant C accumulation rate and infinitely slow turnover (the C
stabilized on minerals had effectively no turnover over 300 yr). The best fit for the
150-kyr site, assuming constant C inputs, was turnover of 175,000 yr. (The
precipitation gradient site with similar age, climate, and vegetation, Kohala-L, had
comparably depleted SOM; see Fig. 3.) If we model the same data assuming that
C ceased accumulating at the 150-kyr site tens of thousands of years ago, the
calculated turnover time is closer to 50,000 yr. The two oldest sites, with SOM
turnover much less than substrate age, can be assumed to be at steady state. The
third horizon (Bhs) at Kauai is a plinthite (stable iron oxide) layer with sharply
depleted radiocarbon content.
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171
content were measured by horizon down to the depth at which
relatively unweathered parent material was encountered (Table 1).
Soil depth remained nearly constant to 1 Myr, as a result of soil
collapse due to congruent dissolution of minerals. At the 4,100-kyr
site, however, the weathered lava extended
.
3 m deep. Between 1.2
and 1.3 m, organic content fell by nearly 100-fold to 0.01%, and we
stopped sampling at 1.5 m.
At all sites, carbon (C) content declined with depth, with the
steepest decline from the surface to 20–30 cm depth (Fig. 1a).
Nevertheless, a significant fraction of each site’s soil organic
matter (SOM) was below 20 cm, ranging from 33% at the two
youngest sites to 70% at the two intermediate-aged sites.
Radiocarbon content decreased with depth at all sites, with slope
depending on soil age (Fig. 2)
5–7
. To estimate average turnover time
of SOM (the reciprocal of the decomposition rate), we used
14
C/
12
C
ratios to constrain a non-equilibrium model of carbon stocks and
flows (see Methods). The surface horizons were dominated by fast-
cycling organic matter, shown by positive (‘bomb’)
D
14
C values
(Fig. 1b, Fig. 2b). In the subsurface horizons, soil C was depleted in
14
C, yielding modelled turnover times slower than 10,000 yr at the
20-kyr, 1,400-kyr and 4,100-kyr sites, and slower than 20,000 yr at
the 150-kyr site (Fig. 1b). The mineral horizons were clearly
dominated by long-residence-time C, often referred to as the passive
fraction of SOM
1
. At the youngest site, radiocarbon content in the
lowest soil horizons was consistent with C accumulating at a
constant rate with virtually no turnover.
The total stock of organic C in soil increased with substrate age up
to 150 kyr, to 60 kg m
−
2
, and then decreased to 31 kg m
−
2
at the
oldest site (Fig. 2a). Carbon inventory in the surface horizons (O
and A horizons) reached its maximum at 20 kyr, and varied less with
age than did deep soil C. A similar pattern was observed in the top
50 cm of these soils, based on 10 replicate soil pits per soil age
4
. Most
of the carbon sequestered after the first few thousand years of soil
development accumulated below 20 cm, in the passive fraction.
Similarly, most of the organic C lost in later soil development was
from the subsurface soil.
The decline in the quantity of soil C stored on older substrates
could result from a decline in productivity and/or an increase in
turnover. Net primary productivity does peak at 150 kyr, but is only
15% less at 4,100 kyr (ref. 8)—too small a decrease to explain the
50% loss of soil C. The profile-integrated turnover rate, however,
doubles between 150 and 4,100 kyr. Most of the decrease in soil
organic C late in soil development therefore can be attributed to
faster turnover of passive C, rather than to a decrease in plant
productivity.
What controls these changes in the storage and turnover of soil
organic C? Volcanic soils undergo dramatic changes in mineral
composition as they develop. Young volcanic soils are characterized
by high organic content and an abundance of non-crystalline
minerals (allophane, imogolite and ferrihydrite), which are the
primary weathering products. These relatively amorphous minerals
have a high degree of hydration, extensive surface area, and variable
charge
9,10
. As a result, they can form stable organic–mineral bonds
through anion and inner-sphere ligand-exchange reactions, and
their geometry may be well suited for physically protecting organic
matter
2,11
. Non-crystalline minerals are metastable. Given enough
0
20
40
60
0.1
1
10
100
1,000 5,000
-800
-400
0
400
0.1
1
10
100
1,000
5,000
0
250
500
0.1
1
10
100
1,000
5,000
0
250
500
0.1
1
10
100
1,000
5,000
Substrate age (kyr)
∆
14
C of SOM (
‰
Soil or
g
anic C (k
g m
–2
)
Non-crystalline
miner
als (k
g m
–2
)
Crystalline
miner
als (k
g m
–2
)
a
b
c
d
)
Figure 2
Soil inventory of carbon in soil organic matter (SOM;
a
),
D
14
C of SOM (
b
),
non-crystalline minerals (
c
), and crystalline minerals (
d
) versus age of soil
substrate. Filled circles, total profile; filled triangles, surface (O and A) horizons.
See Methods for details of soil inventory determinations.
0
5
10
15
20
21
0
15
30
45
60
Non-crystalline minerals (%)
R
2
= 0.45
-1,000
-800
-600
-400
-200
0
200
0
15
30
45
60
75
R
2
= 0.63
excluding 4,100 kyr site
Non-crystalline minerals (%)
∆
14
C of SOM (
‰
Soil carbon (%)
a
b
)
Figure 3
The quantity and turnover of soil C versus non-crystalline mineral
content for the six chronosequence sites.
a
, Soil C versus non-crystalline content
of mineral horizons (that is, excluding O and A horizons) from the chronose-
quence sites.
R
2
¼
0
:
44,
n
¼
35. Plotting symbols as in Fig.1a.
b
,
D
14
C versus non-
crystalline content of all horizons from the chronosequence sites. The solid line
shows the regression between mineral abundance and
D
14
C with chrono-
sequence
data,
excluding
4,100-kyr
site;
R
2
¼
0
:
62,
n
¼
39;
model:
D
14
C
¼
184
2
20 (noncrystalline %). The dashed line shows the regression for
the soil-age and precipitation gradients combined (precipitation gradient
n
¼
16,
described in Table 1); model:
D
14
C
¼
119
2
17 (noncrystalline %);
R
2
¼
0
:
63 for
data from all sites except 4,100 kyr,
n
¼
55. Plotting symbols for chronosequence
data as in Fig. 1a; symbols for precipitation gradient data are
L
, 150 kyr, and
W
,
400 kyr.
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time, they dehydrate to crystalline clays, including halloysite,
kaolinite, gibbsite, goethite and haematite
10,12
, that have a lower
surface area and charge density, and consequently a lower affinity
for SOM.
Mineral composition varied substantially with site age along the
chronosequence. The amount of non-crystalline minerals increased
up to 150 kyr and then declined with greater age (Fig. 2c). In
contrast, the amount of crystalline minerals remained low until
150 kyr, then increased steeply between 1,400 and 4,100 kyr (Fig.
2d), tracking the transition from Andisol to Oxisol.
We found that the abundance of non-crystalline minerals
accounted for
.
40% of the variation in organic C content across
all the mineral horizons, substrate ages, and soil orders (excluding
the O and A horizons dominated by fast-cycling plant litter; Fig. 3a).
Non-crystalline minerals also strongly influenced turnover of SOM.
Organic-matter
D
14
C was highly and negatively correlated with
abundance of non-crystalline mineral (
R
2
¼
0
:
62) except in the
oldest site, which had
,
10% non-crystalline content (Fig. 3b). In
contrast, there was no discernible correlation between the abun-
dance of crystalline minerals and C content or turnover across sites.
The amount of C stabilized per gram was much greater for non-
crystalline than for crystalline minerals.
Can this relationship between non-crystalline minerals and
organic C be used to predict turnover at other sites? To test this,
we analysed soil samples from a precipitation gradient of eight sites
formed on volcanic soils on the Island of Hawaii. Due to the effects
of climate on weathering processes, the six drier sites of this gradient
differed substantially in soil development and mineral composition
from the chronosequence soils. The amount of non-crystalline
minerals increased with precipitation across the precipitation gra-
dient (data not shown). We sampled 16 horizons (Table 1) and
found that the relationship between non-crystalline minerals and
D
14
C along the precipitation gradient was nearly identical to that
along the chronosequence (Fig. 3b). A linear regression from the
chronosequence accounted for 46% of the variance in
D
14
C on the
precipitation gradient; a regression with the two data sets combined
(
n
¼
55) yielded a coefficient of determination
R
2
¼
0
:
63 (Fig. 3b).
We conclude that the ability of these soils to retain C is due to the
capacity of non-crystalline minerals to stabilize large quantities of
organic C for thousands of years.
The mechanisms identified here may be generalizable to many
soils of humid environments, where weathering initially produces
metastable, reactive minerals that later transform to stable, less-
reactive products. In such environments common parent materials—
such as granite, siltstone, sandstone and basalt—that contain
feldspar, olivine, pyroxene or glass initially weather to non-crystalline
minerals. In addition, in cool, wet environments, pedogenic
processes create spodic horizons which have high concentrations
of non-crystalline minerals and organic matter, on a wide variety of
coarse-textured parent materials. In both cases, the non-crystalline
minerals transform over time to weakly reactive crystalline kaolin
and sesquioxide minerals
13
that have a much lower capacity to
stabilize C.
The weathering of other common minerals follows a similar
sequence of reactivity in humid environments. Biotite and musco-
vite micas initially weather to 2 : 1 clays such as vermiculite and
smectite; these in turn weather to kaolin (1 : 1 clay) and gibbsite with
large losses of mineral surface area and reactivity. These changes are
not as large as those observed in volcanic soils (Fig. 3)—globally, as
well as on our chronosequence, Andisols contain twice as much
organic soil C per m
2
as do Oxisols or any other soil order except
Histosols
14
—but the pattern of change in C storage and turnover
should be similar in kind to that shown here
15
. As mineral stability
increases over time, soil minerals lose capacity to stabilize SOM.
These results allow prediction of the direction and magnitude of
C fluxes caused by mineralogical changes associated with millennial
timescales of soil development. For example,
,
25% of the world’s
SOM is stored in soils that began developing after the last major
deglaciation
16,17
; the capacity of the mineral soils to store C will
change as they undergo further development, and they may even-
tually become long-term sources of atmospheric CO
2
.
Among the big uncertainties in our understanding of the global
carbon cycle are the current distribution of soil organic C, and how
much of this large reservoir of C exchanges with the atmosphere on
centennial and shorter timescales compared with millennial time-
scales. An understanding of mineral control over SOM storage and
turnover can improve our ability to model terrestrial carbon
cycling. In addition to the influence on long-term storage and
turnover shown here, C stabilization by non-crystalline minerals
also influences nutrient availability
9
and decomposition of labile
Table 1 Site descriptions
Site
Age of parent
material (yr)
Soil order
Elev.
(m)
Mean annual
temp. (
8
C)
Mean annual
precip. (mm)
Dominant vegetation
Horizons
Depth
(cm)
...................................................................................................................................................................................................................................................................................................................................................................
Chronosequence sites
...................................................................................................................................................................................................................................................................................................................................................................
Thurston
300
Inceptisol
1,176
16
2,500
Metrosideros
forest
Oe-4Bwb
0–38
Olaa
2,100
Inceptisol
1,200
16
2,500
Metrosideros
forest
Oe-C3
0–64
Laupahoehoe
20,000
Andisol
1,170
16
2,500
Metrosideros
forest
Oe-2CR
0–110
Kohala
150,000
Andisol
1,122
16
2,500
Metrosideros
forest
Oe-4CR1
0–98
Kolekole
1,400,000
Ultisol
1,210
16
2,500
Metrosideros
forest
Oa-CR2
0–100
Kokee
4,100,000
Oxisol
1,134
16
2,500
Metrosideros
forest
Oe-Crg
0–125
...................................................................................................................................................................................................................................................................................................................................................................
Precipitation gradient sites
...................................................................................................................................................................................................................................................................................................................................................................
Kohala-A
170,000
Andisol
77
23
160
Bufflegrass pasture
(
Cenchrus ciliaris
)
Bw3
69–104
Kohala-B
170,000
Mollisol
185
23
180
Bufflegrass pasture
Bw2
37–65
Kohala-D
170,000
Andisol
356
22
220
Bufflegrass pasture
Cr1
70–113
Kohala-E
170,000
Andisol
674
20
570
Bufflegrass pasture
Bw3
44–62
BC1
62–85
Kohala-H
170,000
Andisol
992
18
1,060
Kikuyu pasture
Bw2
36–52
(
Pennisetum clandestinum
)
Bw3
52–71
Kohala-J
170,000
Andisol
1,158
18
1,380
Kikuyu pasture
Bw2
47–67
Bw3
67–92
Kohala-L
170,000
Andisol
1,254
17
3,000
Metrosideros
forest
Bw1
25–40
Bw2
40–56
Cr1
56–73
Amalu
400,000
Andisol
1,494
14
2,000
Metrosideros
forest
2Bw1
25–31
2Bw2
31–46
2Bw3
46–62
2Bw4
62–80
...................................................................................................................................................................................................................................................................................................................................................................
All the sites were located on the Island of Hawaii, except the 1,400-kyr (Kolekole, Molokai) and 4,100-kyr (Kokee, Kauai) chronosequence sites, and were on undissected shield surfaces or
broad ridge tops.
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substrates
18
in surface soils. Ecosystem models at present represent
the mineral soil environment with a soil-texture parametrization
which is used to partition flows of organic C into fast- and slow-
cycling pools
19,20
. We suggest that these models should include soil
mineralogy as well as texture in determining storage and turnover
times, based on the varying influence of different minerals
18
. Soil
mineral composition can be estimated from soil maps tied to soil
classification schemes, which reflect climate, parent material and
developmental stage of soil.
Our results show that across landscapes and over long timescales,
the largest changes in the quantity and turnover of soil organic
carbon may be due to variation in passive (mineral-stabilized)
carbon deep in the soil. Passive carbon pools in turn are controlled
by soil mineralogy, which exerts an influence on soil carbon storage
of the same magnitude as that attributed to climate or vegetation in
other studies
21–24
. The predictability of soil mineralogy, as it varies
spatially as a function of climate and parent material and temporally
as a function of soil development
25,26
, means that an understanding
of how minerals influence soil carbon dynamics should yield
significant improvements to our understanding of the role of soils
in the global carbon cycle.
M
. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .
Methods
Carbon analysis and turnover calculation.
Soils for carbon analysis were
oven-dried at 110
8
C for 48 h, sieved (2 mm), ground, and analysed with a Carlo
Erba NA1500 autoanalyser. Radiocarbon content of SOM, measured by
Accelerator Mass Spectrometry at Lawrence Livermore National Laboratory on
targets prepared by sealed-tube, zinc reduction
27
, is expressed as
D
14
C, calcu-
lated as (
R
2
1)1,000.
R
, or fraction modern, is the per mil deviation from
14
C/
12
C ratio of oxalic acid in 1950
28
. Analytical error averaged 8‰ for values
close to modern. Carbonate was removed from Kohala B soils with dilute HCl
before radiocarbon analysis. Fast-cycling C (turnover time, years–century) has
positive
D
14
C because it has incorporated a significant proportion of the
14
C
that was released by atmospheric nuclear weapons testing from 1959 to 1963
5
.
Before weapons testing, atmospheric
D
14
C was approximately zero. Slow-
cycling C (turnover centuries–millennia) has negative
D
14
C owing to extensive
radioactive decay relative to
14
C inputs from more recently photosynthesized
organic material. We used a non-steady-state model to estimate turnover from
our measurements of C and radiocarbon content, and substrate age. Carbon
accumulation was modelled as
C
ð
t
Þ ¼
IT
ð
1
2
e
2
t
=
T
Þ
; radiocarbon content of
SOM was modelled as
R
SOM
ð
t
Þ ¼ ½
R
atm
ð
1
2
e
2
ð
l
þ
1
=
T
Þ
t
ÿ
=
½ð
1
þ
T
l
Þð
1
2
e
2
t
=
T
Þÿ
,
where
T
is turnover time,
C
0
¼
0,
I
is the annual C input to soil,
R
atm
is
fraction modern of the atmosphere, assumed equal to that of plant C input to
soil,
l
is half-life of
14
C (5,568 yr) and
t
is years of soil development. We assume
that C loss due to erosion is negligible relative to the rate of decomposition.
Soil inventory determination.
The C inventory is the sum of [(% C)(bulk
density)thickness] for each horizon. An analogous formula was used for
inventory of the mineral groups. The profile-total radiocarbon content is the
sum of [
R
(% C)(bulk density)thickness] for each horizon divided by total C
content of the profile, where
R
is fraction modern, defined above. Bulk density
of wax-coated clods was determined by displacement
4
. Soils for mineral
analysis were air-dried, sieved (
,
2 mm), ground and subject to sequential
extractions, with mass balance at each step, before analysis of Al, Fe, Si, and
mineral structure
29
. Organic matter was removed with hydrogen peroxide and
acetate. Non-crystalline mineral content was estimated by ammonium oxalate
dark extraction
30
. Crystalline sesquioxides and kaolin were extracted with
citrate dithionite
31
and NaOH
29
, respectively. These operationally defined
mineral fractions are expressed on a total-soil basis (including soil organic
matter). Mineral structures were identified by X-ray diffraction and Fourier-
transform infrared spectroscopy. Coarse fraction was negligible in the chrono-
sequence soils. To quantify collapse of the soil during weathering, zirconium
was used as a conservative element.
Received 30 December 1996; accepted 30 June 1997.
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Acknowledgements.
We thank S. Zheng and T. Crews for technical assistance; Lawrence Livermore
National Laboratory’s Center for Accelerator Mass Spectrometry for radiocarbon analysis; The Nature
Conservancy, US National Park Service, Parker Ranch, Kahua Ranch, Hawaii Division of Forestry and
Wildlife, Hawaii Division of State Parks, and the Joseph Souza Center for access to field sites; and
E. Davidson, J. Gaudinski and J. Harden for comments on the manuscript. This work was supported by
National Science Foundation, Mellon Foundation, Joan Irvine Trust, and Jet Propulsion Lab-California
Institute of Technology on contract to NASA.
Correspondence should be addressed to M.S.T. (e-mail: torn@ess1.ps.uci.edu).
Eclogite xenoliths in west
African kimberlites as
residues from Archaean
granitoid crust formation
Hugh Rollinson
Department of Geography and Geology, Cheltenham and Gloucester College of
Higher Education, Francis Close Hall, Swindon Road, Cheltenham GL53 0BL, UK
. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .
Eclogites are a comparatively rare but petrologically important
member of kimberlite xenolith suites. Their broadly basaltic
chemistry has led many authors to propose that they represent
ancient, subducted ocean crust
1–3
. Recent studies
4–6
, however, have
suggested an alternative origin and propose that kimberlitic
eclogites are residues from the process of Archaean granitoid
crust formation. Geochemical arguments in support of this new
Document Outline