Carbon in Boreal Peatlands

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9 Carbon in Boreal Peatlands
Harri Vasander and Anu Kettunen
9.1 Introduction
Although peatlands have been developing on Earth since wetland plants
first existed, the great majority of present-day peatlands have originated
during the last 11,000 years (Gajewski et al. 2001; Chap. 3). Owing to dif-
ferent definitions of wetland ecosystems the global coverage estimates of
peatlands differ considerably. The estimates of the total area of peatlands
in the world have changed from 1 ×10
6
to 5 ×10
6
km
2
during the last
100 years as more inventory results have become available (Kivinen and
Pakarinen 1981). One reason for the differences is the required minimum
peat thickness (see Joosten and Clarke 2002 for a thorough discussion).
While not going deeper into that discussion,it is interesting to note that if
a geological definition with a minimum peat thickness of 30 cm is com-
pared with the biological definition (Laine and Vasander 1996) that
requires only potentially peat-forming vegetation, the ratio between geo-
logical and biological peatlands is approximately 0.7 for Finland and Swe-
den (Lappalainen 1996; Vasander et al. 2003) as well as for the European
part of the former USSR (calculated from the data of Tjuremnov 1949).
Joosten and Clarke (2002) and Charman (2002) estimate that the area of
boreal and subarctic peatlands with more than 30 cm of peat is approxi-
mately 3.5 ×10
6
km
2
.
Also, estimates of the C pool in peatlands have varied by over an order
of magnitude from 41.5 Pg (Buringh 1984) to 489 Pg (Schlesinger 1977)
owing to differences in estimates of average peat depth and bulk density.
The moderate estimate of 300 Pg (Sjörs 1981) represents about 13% of the
terrestrial C in the biosphere and is in line with later studies that estimate
270–370 Pg for northern mires (Turunen et al. 2002),249 Pg for boreal and
subarctic mires (Armentago and Menges 1986),and 234–252 Pg for boreal
mires (Lappalainen 1996). The often cited figure by Gorham (1991; 455
Pg) was calculated by using a total area of pristine boreal and subarctic
Ecological Studies,Vol. 188
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Boreal Peatland Ecosystems
© Springer-Verlag Berlin Heidelberg 2006

peatlands of 3.42 ×10
6
km
2
, a mean depth of the peat layers of 2.3 m, a bulk
density of 0.122 g cm
–3
, and a C content of 51.7 %.With a lower mean bulk
density of 0.091 g cm
–3
(Mäkilä 1994; Turunen 2003) and a mean depth of
1.7 m for all boreal and subarctic mires (Turunen 2003), the C pool of
northern mires would be 274 Pg, which is in accordance with most of the
aforementioned values. The large range in the C storage estimate mainly
reflects uncertainty in the depth of global peat deposits (Gorham 1991,
Botch et al. 1995, Clymo et al. 1998). A more accurate characterization of
the age–depth distributions of mires, especially from North America and
Russia,is needed (Bauer et al.2003,Yu et al. 2003,Borren et al 2004, Sheng
et al. 2004).
9.2 Carbon Cycle in Peatlands
Atmospheric CO
2
is fixed by plants via photosynthesis during the growing
season and subsequently is deposited as litter both on and in the soil. Net
primary production (NPP) in boreal peatlands is lower than in many
other ecosystems (Ruimy et al.1996; Frolking et al. 1998; Bubier et al.1999;
Chap. 8). Decay rates are also low as the water table lies near the soil sur-
face, leading to anoxic conditions. Peat accumulates whenever the rate of
organic matter production exceeds the rate of decay. While the NPP and
peat accumulation values in different kinds of peatlands differ widely,the
“efficiency” of peatlands, i.e., the ratio between peat accumulation and
NPP, varies between 1 and 20 % (Tolonen 1979; Tolonen et al. 1992; Warner
et al. 1993; Francez and Vasander 1995; Moore et al.2002; Feng 2002). The
peat accumulation rate has been related to peatland geographical location
(south greater than north), age (young greater than old), and type
(Korhola et al.1995).
Part of the photosynthesized C is returned to the atmosphere as CO
2
in
the maintenance and growth respiration of aboveground and below-
ground plant parts, and in the respiration of consumers such as soil ani-
mals and heterotrophic microbial communities. By measuring the total
flux of CO
2
it is impossible to distinguish these from each other, but by
other means it has been found that autotrophic and heterotrophic respira-
tion comprise about one third of the CO
2
uptake via photosynthesis dur-
ing the intensive growth period (Bubier et al.1998; Heikkinen et al.2002).
The rate of autotrophic respiration is regulated by photosynthesis, tem-
perature, and water and nutrient availability,while heterotrophic respira-
tion is controlled largely by soil temperature, oxic peat layer volume,
nutrients and soil pH, and the quality and quantity of decomposable
material (Chapman and Thurlow 1998; Chapin et al. 2002). Root-associ-
ated respiration follows the vegetation phenology and may account for
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10–45 % of the total soil CO
2
release,originating mainly from the turnover
of fine roots and from root exudates (Silvola et al. 1996a).
The remaining C is transformed into plant structures, especially into
the belowground parts of plants where the majority of the plant biomass
is located (Metsävainio 1931;Vasander 1982;Wallén 1986,1992; Sjörs 1991;
Saarinen 1996). Finally it is deposited as dead plant matter (litter) on and
in the soil (Fig. 9.1). Mosses grow upward, die gradually, and regulate the
vertical growth rate of peatlands. Besides adding stem, rhizome, and root
matter to the peat, vascular plants provide physical support for the
upward growth of mosses (Malmer et al. 1994).In the oxic surface parts of
the peat (acrotelm), litter initially is decomposed primarily by aerobic
bacteria, leading to the release of CO
2
, but eventually litter becomes cov-
ered by the gradually rising water table (Reader and Stewart 1972; Clymo
1984, 1992; Bartsch and Moore 1985; Laine et al. 1996; Scanlon and Moore
2000). In the water-saturated anaerobic part of the peat (catotelm),
decomposition is slow and a large portion of the total mineralized C is
released to the atmosphere as CH
4
.
C also flows in and out of the peatland in dissolved form (dissolved
organic C, DOC) (Urban et al. 1989; Sallantaus 1992; Schiff et al. 1998;
Moore 2003). As peatlands have very high C densities, the DOC output
from them usually exceeds the DOC input to them with water inflow. The
C leaching rate and loss as DOC from peatlands depend especially on
hydrologic throughflow rates and on peatland NPP (DeVito and LaZerte
1989; Sallantaus 1992; Moore 2003).C leaching and DOC losses from peat-
lands may be increasing owing to warming climate (Freeman et al.2001).
Also, the significance of episodic factors like fires causing C loss from
peatland ecosystems (Pitkänen et al. 1999; Turetsky et al. 2002,2004) may
increase in the future (Robinson and Moore 2000).
Ca leaches downward in the peat profile (Charman et al. 1994; Domisch
et al. 1998) and can reach the underlying mineral soil (Turunen et al. 1999)
especially during the early developmental phase of a peatland. Long-term
average C accumulation rates beneath Lakkasuo mire in central Finland
were 19 g m
–2
year
–1
at sites younger than 500 years and 1 g m
–2
year
–1
at
sites older than 500 years (Turunen and Moore 2003), while the average
estimates throughout the Holocene were 17–19, 20, and 24 g C m
–2
year
–1
in Finland, Russia, and Canada, respectively (Turunen et al. 2001, 2002;
Vitt et al.2000; Turetsky et al. 2002).
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Fig. 9.1. a Sphagnum-dominated bogs and poor fens may be considered to have a
diplotelmic acrotelm–catotelm structure. Mosses add new organic matter to the sur-
face. Most of the decay takes place in the upper oxic surface peat layer.Below this is
the anoxic catotelm layer with stagnant water where the decomposition of organic
matter is much slower. The lowest water level is the main regulator of the oxic
acrotelm layer and only a small amount of new photosynthetic organic matter is
transported to the anoxic catotelm layer for the methanogens. bIn sedge fens the
diplotelmic structure with a horizontal limit between acrotelm and catotelm does

9.3 Carbon Dioxide Uptake and Release
Photosynthesis is a light-controlled process in which CO
2
is the C source
and light is used as energy (Fig. 9.2). Other controlling factors of photo-
synthesis are CO
2
concentration, temperature, and water and nutrient
availability, as well as the leaf area (Mooney 1986). To tackle the spatial
and temporal variation in gas fluxes from boreal peatlands that partly
results from the considerable spatial variability in microtopography
(Figs. 9.1, 9.2), several approaches such as the chamber technique (Crill
1991; Carroll and Crill 1997) and eddy covariance (EC; Baldocchi 2003)
have been used.While the chamber technique is applicable for microscale
flux measurements, the EC technique can be applied to landscape-scale
measurements of net CO
2
flux and energy balance (Lafleur et al. 1997,
2001; Aurela et al. 1998; Soegaard and Nordstroem 1999; Vourlitis and
Oechel 1999; Hargreaves et al.2001;Frolking et al.2002). Chamber and EC
techniques have yielded quite similar C fluxes (Norman et al. 1997).
The difference between the gross uptake of CO
2
in photosynthesis (P
GP
)
and both autotrophic and heterotrophic respiration (R
a
and R
h
, respec-
tively; the sum of which is the total ecosystem respiration,or R
tot
) is called
net ecosystem production (P
NE
) (Curtis et al. 2002). P
NE
can be directly
measured with chamber and EC techniques, while the same is not true for
P
GP
. However, as P
NE
is small compared with P
GP
and R
tot
, accurate mea-
surements are needed to determine its rate. Ecologists and other soil-
related researchers consider P
NE
positive when the system acquires more
CO
2
from the atmosphere than is released back to it.When P
NE
is negative,
there is a net flow of CO
2
to the atmosphere. Meteorologists and other
atmosphere-related researchers consider flows to and from the atmos-
phere in the opposite way in terms of the positive or negative signs to the
fluxes.
As photosynthesis is highly dependent on the amount of light, P
NE
is
usually determined under different light conditions with artificial shad-
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not exist. Roots of sedges grow deep into anoxic peat, releasing root exudates and
producing easily degradable root litter directly in the anoxic layer where
methanogenic processes occur. These new photosynthates enable high methane
production.Also oxygen is transported deep into the peat by the aerenchymous tis-
sues of sedges forming oxic “pockets”to the anoxic peat layer.Also new organic mat-
ter is produced on the surface as well as in the deeper peat layers. Vascular plants
possessing aerenchymous tissues facilitate the transport of methane from anoxic
layers to the atmosphere.Increased root exudation in combination with transport of
methane in the intercellular space of sedges explains the higher methane emissions
in oligotrophic sedge fens in comparison with ombrotrophic bogs and mesotrophic
fens. (Photographs by Harri Vasander from Russian Karelia)

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Fig. 9.2. The key interactions in carbon cycling dynamics in different microsites of a
boreal peatland. Plants bind atmospheric carbon dioxide in photosynthesis ( func-
tion of plant community, temperature, water level, photosynthetically active radia-
tion, PAR) and supply part of the carbon to the peat matrix (function of litter qual-
ity,temperature,water level). In a hollow/flark with scarce field layer vegetation the
carbon supply is lower. In high hummocks, the carbon is mainly supplied to the
unsaturated layers where it is aerobically degraded. In lawn surfaces, low sedge
abundance results in low methane flux owing to the lack of substrate and,further on,
methane flux increases with an increase in sedge cover until at some point the
increased oxygen supply to the peat profile starts to decrease the methane flux.
Plants also supply the peat matrix with oxygen, which enhances methane oxidation
and inhibits methane production.Methane is transported upward by plants and dif-
fusion. Also ebullition occurs, especially in undefeated surfaces. Oxygen concentra-
tions determine the proportion of substrate to aerobic decomposition and methane
production. Aerobic decomposition ( function of plant community, temperature,
water level,peat chemistry) consumes oxygen. The amount of substrate available for
methanogenesis also depends on the substrate production rate (function of photo-
synthesis, root profile in peat) and substrate consumption rate (function of
methanogens, substrate concentrations).The methane oxidation rate increases with
increasing population of methane oxidizers and increasing methane and oxygen
concentrations.The population dynamics, i.e., biomass gain and dying of methane
producers and methane oxidizers, plays an important role when the water table
(WT) shows short-tem and long-term fluctuations. Carbon is also transported as
dissolved organic carbon (DOC; function of plant community, temperature, water
level, hydrology). The possible flow of CaCO
3
in the case of rheotrophic mires and
from soil layers beneath the peat in ombrotrophic mires (Lamers et al. 1999) is not
considered here. (Modified from Kettunen 2002 with the permission of Helsinki
University of Technology)

ing. Photosynthetically active radiation (PAR) is measured as photosyn-
thetic photon flux density (PPFD, micromoles per square meter per sec-
ond). R
tot
can be measured in total darkness by blocking the light entering
the chamber.P
GP
can then be calculated by adding P
NE
and R
tot
, assuming
that plant photorespiration equals plant respiration under dark condi-
tions. Over the past decade,many measurements have been made to deter-
mine the relationships between PPFD and P
NE
in different kinds of boreal
peatland ecosystems in Canada, Russia, Scandinavia, and the USA (see
review tables in Blodau 2002; Heikkinen 2003).Most of the measurements
have been made during the growing season.; however, more measure-
ments are now being made during other seasons when R
tot
continues even
under snow, but photosynthesis has ceased. Dark measurements during
the summer show that R
tot
varies from –1 to –7 µmol CO
2
m
–2
s
–1
(–1.0 to
–7.3 g CO
2
-C m
–2
day
–1
) (Martikainen et al. 1995;Alm et al. 1997; Frolking
et al. 1998; Ikkonen et al. 2001; Moore 2001; Moore 2002; Saarnio et al.
2003; Tatarinov et al. 2003). Respiration is quite similar in different kinds
of peatlands (poor – rich) and depends mostly on temperature and water
table variation (Updegraff et al. 2001; Moore et al. 2002; Chimner and
Cooper 2003).
When PAR increases to about 200 µmol m
–2
s
–1
P
GP
equals R
tot
(P
NE
=0);
with increasing PAR, peatlands become net CO
2
sinks. Average maximum
summer P
GP
values are 5 µmol CO
2
m
–2
s
–1
with 1,000–1,500 µmol m
–2
s
–1
PAR on minerotrophic peatlands and 2 µmol CO
2
m
–2
s
–1
on ombrotrophic
peatlands (Frolking et al.1998; Moore 2001). Based on the hourly values of
measured P
NE
, seasonal net exchange may be modeled (Alm et al. 1997;
Carroll and Crill 1997; Bellisario et al. 1998; Saarnio et al. 2003; Tuittila et
al. 2004).Growing season P
NE
estimates vary considerably, but may reach
values as high as 200 g CO
2
-C m
–2
(Moore 2001).
PAR has a clear diurnal cycle (Oechel et al. 1995). Diurnal rhythm was
also measured for CO
2
exchange on all microsites of a boreal bog (Ket-
tunen 2000), acting as net C sinks during the day when the photosynthe-
sis rate exceeded respiration and as net C sources during the night when
photosynthesis ceased,but respiration continued.Dry sites were found to
fix more C during the daytime than wet sites, which parallels reported
differences in CO
2
exchange across hydrological and vegetation gradients
(Bubier et al. 1998, 1999; Frolking et al. 1998; Christensen et al. 2000).
Respiration rates remained relatively low as the water table stayed close
to the peat surface (Silvola et al. 1996b) and, consequently, daily CO
2
exchange was clearly positive for all microsites for most of the season.
The results of Kettunen (2000) emphasized that CO
2
exchange is
extremely sensitive to variation in environmental factors on short-term
time scales and, consequently, annual C exchange estimates are also
affected by short-term variation (Bubier et al. 1998, 1999; Griffis et al.
2000; Soegaard et al. 2000).
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In boreal peatlands, seasons other than the growing season are impor-
tant contributors to annual CO
2
exchange. In spring, peatlands usually
function as a small source of CO
2
to the atmosphere, at least until plants
are released from under the snow cover and photosynthesis begins. Hum-
mocks are released first and they become the first microsites to turn into
sinks of CO
2
with the bright sunny days in spring (Bubier et al. 1998).
Because of the decreases in vascular plant leaf area and in the capacity of
vascular plant leaves to photosynthesize, during late autumn only
bryophytes photosynthesize (Silvola and Hanski 1979).
Significant proportions (10–40%) of CO
2
fixed during the growing sea-
son are released from arctic, subarctic, alpine tundra, and boreal peats
during the winter (Sommerfeld et al. 1993,1996; Zimov et al.1993; Brooks
et al. 1997;Oechel et al. 1997;Mast et al.1998; Alm et al.1999a; Fahnestock
et al. 1999;Panikov and Dedysh 2000; Aurela et al.2001,2002; Lafleur et al.
2001; Heikkinen et al.2002; Roehm and Roulet 2003).Soil biological activ-
ity partly controls the soil CO
2
production rate during the winter, while
vertical gas transport is controlled by snowpack characteristics such as
porosity, tortuosity, and depth (Sommerfeld et al. 1993; Mast et al. 1998).
Brooks et al.(1997) suggested that for an alpine tundra site the majority of
CO
2
flux during the winter originated from the thin organic layer at the
soil surface. The total as well as proportional winter CO
2
emissions
depend on the duration and timing of snow accumulation (Brooks et al.
1997; Aurela et al.2004,).
Lafleur et al. (2001) obtained an annual CO
2
uptake of 248 g
CO
2
m
–2
year
–1
(equivalent to 68 g C m
–2
year
–1
) at a boreal bog in central
Canada (growing season 200 days), while the corresponding uptake in a
subarctic fen in northern Finland (growing season 70 days) was
68 g CO
2
m
–2
year
–1
(equivalent to 19 g C m
–2
year
–1
; Aurela et al.2002).The
maximum daily net uptake at the subarctic site with no night was slightly
higher than that of the Canadian bog with clear diurnal light rhythm (9.3
and 8.3 g CO
2
m
–2
day
-1
; Aurela et al. 2002 and Lafleur et al. 2001, respec-
tively). Typical winter CO
2
efflux for the subarctic fen was 0.5 g m
–2
day
–1
,
leading to a seasonal winter efflux of 110 g CO
2
m
–2
, which was greater
than the total annual sink term (Aurela et al. 2002). Fluxes at this fen in
northern Finland have been measured continuously for 6 years
(1997–2002) and the annual CO
2
balance has consistently been positive
(15–192 g CO
2
m
–2
year
–1
equivalent to 4–53 g C m
–2
year
–1
) (Aurela et al.
2004).
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9.4 The Methane Cycle in Wet Ecosystems
9.4.1 Substrate Supply
Even though peat itself represents a large reservoir of C, the C in the peat
matrix is very resistant to decomposition in the anaerobic conditions that
prevail in the peat profile and, therefore,peat C can provide only a limited
substrate for methanogenesis (Kuder and Kruge 2001). Even though it has
long been known that recent C bound by vegetation can promote
methanogenesis by providing root exudates and easily decomposable lit-
ter (Rovira 1969),the full importance of plants to CH
4
fluxes has been real-
ized only recently. Supported by experimental evidence, CH
4
flux was sug-
gested to increase with the photosynthetic activity of plants in a variety of
wetland types (Whiting et al. 1991;Whiting and Chanton 1992,1993).The
significantly lower CH
4
fluxes from unvegetated compared with vegetated
surfaces and studies where clipping of vascular plants decreased CH
4
flux
provided indirect evidence for the importance of plants (Chanton et al.
1992a; Torn and Chapin 1993; Kelker and Chanton 1997; King et al. 1998;
Verville et al. 1998; Frenzel and Karofeld 2000; Christensen et al. 2003).
Furthermore, CH
4
fluxes were found to be correlated to sedge cover both
across microsites within a single mire (Bubier et al.1995a, b; Schimel 1995;
Bellisario et al. 1999; Tuittila et al.2000; Nykänen et al. 2002) and across
different mires (Nilsson and Bohlin 1993; Bubier 1995; Granberg et al.
2001; Nilsson et al. 2001). Positive correlations between CO
2
fixation and
CH
4
flux at sites covered by sedge vegetation have been reported
(Waddington et al. 1996; Friborg et al. 2000; Strack et al. 2004).Also pore
water CH
4
concentrations were found to increase from unvegetated to veg-
etated surfaces (Whiting and Chanton 1992).
Use of stable and radioisotopes of C have supported the link between
vascular plant root exudates and the methanogenic food chain. Having
found that
14
C-dated CH
4
collected directly from peat was 2,000 years
younger than adjacent peat, Charman et al. (1994) suggested that at least
part of the C in CH
4
originates from DOC in mire water and that root exu-
dates are one potential source of DOC. In some studies using C isotopes,
recent photosynthetic C was found to be the predominant substrate for
methanogenesis (van den Pol-van Dasselaar and Oenema 1999; Chasar et
al.2000), while in others this contribution was considered to be much
lower (King and Reeburgh 2002; King et al.2002; Christensen et al. 2003).
The variation is understandable as allocation of C to aboveground and
belowground parts of the plants and to exudation is known to be affected
by plant species,plant age,tillering stage, root damage, light intensity,soil
temperature, soil water stress, nutrient availability/deficiency, and soil
microorganisms (Rovira 1969; Shaver and Cutler 1979; Kummerow and
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Ellis 1984). Root exudation has been shown to increase the availability of
the uptake of many nutrients, for example phosphate (Hoffland 1992),
iron (Römheld 1991), manganese, and copper (Mench and Martin 1991),
through formation of complexes with nutrient ions. Exudation can also
promote the mineralization of nutrients by enhancing microbial activity
in the rhizosphere (Darrah 1993). Root exudation of Eriophorum angusti-
folium grown in ombrotrophic and oligotrophic mire conditions was
found to increase with decreasing availability of nutrients, suggesting that
high root exudation might be a compensative mechanism to deal with low
concentration of cations and low pH in mire ecosystems (Saarinen 1999).
On average,about 15% of the photosynthetically fixed C is estimated to be
released from the roots, mainly in microbial and plant respiration
(Saarnio et al. 1998; Saarnio and Silvola 1999).
Even if the proportion of root exudates is usually very small in compar-
ison to the net photosynthesis, this fraction of C consisting of, for exam-
ple, soluble sugars (glucose, fructose, sucrose, and other mono- and
oligosaccharides) and organic acids is readily available for microbes
(Rovira 1969; Russell 1977; Ström et al 2003). Methanogenic Archaea, in
general, rely on other anaerobic bacteria for the initial breakdown of com-
plex organic structures into simpler molecules (Svensson and Sundh
1992). Besides soluble sugar exudates, plant roots release secretions,
sloughed cells, and material from root turnover into the soil, and these
compounds can be exploited by microorganisms for biosynthesis and
energy production (van Veen et al. 1989).
9.4.2 Acetate and Hydrogen Pathways
In freshwater systems, CH
4
is formed either from acetate dissimilation
(acetate pathway) or bicarbonate reduction (hydrogen pathway) (Kelley et
al.1992; Westermann 1993) (Fig. 9.3), which differ in relation to tempera-
ture dependence and substrate availability (Ferguson and Mah 1983;
Svensson 1984; Westermann 1993). At low temperatures (between 10 and
15 °C), the acetate pathway was found to contribute 85–90% of the CH
4
produced; the contribution of the hydrogen pathway increased with
increasing temperature (Avery et al. 1999; Fey and Conrad 2000). In addi-
tion to temperature control, vegetation affects the pathways, so in vege-
tated sites where fresh organic matter is available owing to high plant pro-
ductivity, the acetate pathway dominates, while in unvegetated sites the
hydrogen pathway becomes important (Bellisario et al. 1999; Popp et al.
1999; Chasar et al.2000; Ström et al.2003). Recent studies using C isotope
methods (Chanton et al.1995;Avery et al.1999; Bellisario et al.1999; Popp
et al.1999; Chasar et al.2000) have shown that the acetate pathway clearly
dominates in northern peatlands during summer. Also, the acetate path-
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way is favored in the shallow subsurface peat, while the hydrogen pathway
becomes more dominant in older, less reactive deeper peat (Hornibrook et
al. 1997). The analysis of methanogenic communities from boreal mires
supported the idea that upper peat layers that receive fresh organic matter
harbor acetoclastic methanogens, while hydrogen-utilizing methanogens
prevail in deeper layers (Galand et al. 2002,2003).
9.4.3 Methane Production
CH
4
is formed as a terminal step of a very complicated anaerobic degrada-
tion chain (Cicerone and Oremland 1988) by methanogenic Archaea (Gar-
cia et al. 2000). In wetlands, changes in substrate availability and redox
conditions are suggested to control the CH
4
production rate and the
growth and death of methanogens (Conrad 1989, 1996;Morrissey and Liv-
ingston 1992; Valentine et al.1994). In addition, in principle, CH
4
produc-
tion is enhanced by an increase in temperature, but under in situ condi-
tions, substrate availability strongly affects the temperature response
(Dunfield et al.1993; Valentine et al.1994; Bergman et al.1998). Deeper in
peat, oxygen concentrations are lower,but the fresh organic C is supplied
mainly to the uppermost layers where plant roots survive (Svensson and
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Fig. 9.3. The internal cycles of methane and
carbon dioxide in peatlands showing the
major microbial acetogenic and
methanogenic processes.High molecular
weight organic matter (OM) is decomposed
by two different fermentation reactions (F1,
F2) into acetate (2CH
2
OÆCH
3
COO
–
+H
+
),
hydrogen and carbon dioxide
(CH
2
+H
2
OÆCO
2
+2H
2
): Aceticlastic reac-
tions (AC) allow further splitting of acetate
to produce methane and carbon dioxide
(CH
3
COO
–
+H
+
ÆCH
4
+CO
2
). Hydrogen and
carbon dioxide can recombine to form
methane by CO
2
reduction (CR:
CO
2
+4H
2
ÆCH
4
+2H
2
O) or again form
acetate by acetogenesis (AG:
2CO
2
+4H
2
ÆCH
3
COOH+2H
2
O). Not explic-
itly shown are processes such as the produc-
tion of bacterial biomass, the initial decay of
plant material,and the generation of high
molecular weight DOC. (Adapted from Eil-
rich and Steinmann 2003 with the permis-
sion of E. Schweizerbart’sche Verlagsbuch-
handlung, Stuttgart)

Sundh 1992; Schimel 1995). Maximal CH
4
production has been observed
at about 20 cm below the water table (Sundh et al.1994). Addition of
monovalent and divalent cations as well as acetate has been found to
increase CH
4
production in peat cores (Thomas and Pearce 2004).
9.4.4 Methane Oxidation
In wetlands, CH
4
oxidation is carried out by low-affinity CH
4
oxidizers
and oxidation rates depend on CH
4
and oxygen availability,which are con-
nected to peat moisture conditions,temperature, and the activity of CH
4
-
oxidizing bacteria in the peat matrix. Populations of CH
4
oxidizers
develop where CH
4
and oxygen profiles overlap in the peat profile (Con-
rad 1989, 1996; Sundh et al.1995; Segers and Kengen 1998). Changes in
CH
4
and oxygen concentrations during the growing season affect the pop-
ulation dynamics of methanotrophic bacteria (Svensson and Rosswall
1984; Whiting and Chanton,1993) that are reflected in the net flux of CH
4
.
On the basis of the microbial 16S ribosomal DNA similarities of near-sur-
face (20 cm) and deeper (6 m) peat layers, Steinmann et al. (P. Steinmann,
S. Huon, P. Rossi,B.Eilrich,S. Casati, unpublished results) have speculated
about the possibility of the presence of methanotrophs in the deepest
anoxic peat layers. A possible oxidizing agent could be solid or colloidal
trivalent iron. Temperature control has been suggested to be less impor-
tant for CH
4
oxidation than for CH
4
production (Dunfield et al.1993).
Estimates of the CH
4
fraction that becomes reoxidized before reaching the
atmosphere vary from 0 to 100% (Yavitt et al.1988, 1990; Moosavi and
Crill 1998; Frenzel and Karofeld 2000; Popp et al. 2000; Pearce and Clymo
2001).
9.4.5 Methane Transport
CH
4
is liberated from peat via three routes: diffusion, ebullition, and pas-
sage through plants (Conrad 1989; Chanton et al.1992b; Joabsson et al.
1999). In unvegetated surfaces, ebullition mainly dominates (van der Nat
and Middelburg 1998; van der Nat et al.1998). In vegetated surfaces,bub-
ble flux may become important during wintertime when plant biomass is
low (van der Nat and Middelburg 1998).Also, individual CH
4
bubbles can
be formed in normal pore water concentrations by “scavenging”CH
4
from
surrounding pore water (Baird et al. 2004). However, whenever vascular
plants are present, bubbling is rare and flux via plants tends to dominate
the diffusive flux (Sebacher et al. 1985; Morrissey and Livingston 1992;
Whiting and Chanton 1992; Schimel 1995;van der Nat et al.1998; Joabsson
et al. 1999; Frenzel and Karofeld 2000). Nevertheless, abrupt and high
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ebullition fluxes of “old” CH
4
from the catotelm may represent a large
source for CH
4
emissions from northern peatlands (Christensen et al.
2003; Glaser et al. 2004).
Some plants, like Phragmites and Typ ha, show active gas transport
based on pressure differences, while others,like Carex spp.,have only pas-
sive diffusion (Koncalová et al.1988; Chanton et al.1992a, 1993; van der
Nat and Middelburg 1998; van der Nat et al.1998; Popp et al.1999).Active
gas transport leads to a strong diurnal pattern in CH
4
flux and, conse-
quently, if no diurnal pattern is observed, plants that use active gas trans-
port are not present (Morrissey et al. 1993; van der Nat and Middelburg
1998; van der Nat et al.1998). The within-plant diffusion rate has been
found to be higher for Eriophorum angustifolium than for Carex aquatilis
(Schimel 1995).Also,temperature has been shown to have an effect on the
within-plant diffusion rate (Thomas et al. 1996). In spite of many exam-
ples for a plant-associated CH
4
oxidation (Frenzel 2000; Heilman and
Carlton 2001),there is growing evidence that a few plants,including Erio-
phorum sp.,do not support CH
4
oxidation (King et al.1990; Chanton et al.
1992b; Kelker and Chanton 1997; Frenzel and Rudolph 1998), possibly
owing to differences in the quality and type of root exudates between
those plant species that support plant-associated CH
4
oxidation and those
that do not (Frenzel and Rudolph 1998).
Kettunen et al. (1996) noticed that diurnal fluctuations in CH
4
emis-
sions tended to occur when the difference between the air temperature
and the peat temperature was large, i.e.,during the warm days in the early
season when deep peat layers had not warmed up. The large diurnal vari-
ations in peat temperatures apparently were related to the diffusion rate of
CH
4
in peat (Jähne et al. 1987) and possibly also to CH
4
production (Dun-
field et al. 1993; Westerman 1993) causing diurnal variations in CH
4
flux.
The result that diurnal fluctuation in the microsites,where they occurred,
could be correlated to peat surface temperatures only for short time peri-
ods indicates that the control mechanisms for CH
4
fluxes may change over
the growing season.
9.4.6 Relations Between Environmental Factors and Methane Flux
The CH
4
fluxes from wetlands, which are controlled by the dynamic bal-
ance between CH
4
production and oxidation rates in peat profiles and the
transport rate from peat to the atmosphere (Conrad 1989, 1996; Bubier
and Moore 1994), show high spatial and temporal variation (Moore et al.
1990; Whalen and Reeburgh 1988, 1992; Dise 1993; Kettunen 2003). The
spatial variation is due to the fact that the basic processes are affected by
site-specific factors, such as average hydrological conditions (Svensson
and Rosswall 1984; Sebacher et al. 1986; Roulet et al.1992, 1993; Moore et
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al.1994; Fiedler and Sommer 2000; Bosse and Frenzel 2001; Christensen et
al. 2004),soil nutrient contents (Svensson and Rosswall,1984; Dise,1993),
substrate concentration and quality (Morrissey and Livingston 1992;
Whiting and Chanton 1992; Valentine et al.1994; Fiedler and Sommer
2000), and vegetation type (Torn and Chapin 1993; Shannon and White
1994; Bubier 1995; Bubier et al.1995a, b). The connection between CH
4
flux and hydrology and vegetation at the microsite scale has been estab-
lished in numerous studies (Roulet et al.1992, 1993; Bubier et al. 1993a,b,
1995a, b; Christensen 1993; Torn and Chapin 1993; Vourlitis et al.1993;
Moore et al.1994; Shannon and White 1994; Bubier 1995; Schimel 1995;
Bellisario et al.1999; van den Pol-van Dasselaar et al.1999; Tuittila et al.
2000; Granberg et al. 2001; Nilsson et al. 2001; Heikkinen et al. 2002;
Nyk‰nen et al. 2003; Christensen et al. 2004). Interannual variation
(Mattson and Likens 1990; Whalen and Reeburgh 1992; Frolking and Crill
1994; Shurpali and Verma 1998; Nyk‰nen et al. 2003), seasonal variation
(Dise et al. 1993; Shurpali et al.1993; Frolking and Crill 1994; Alm et al.
1999a; Mast et al.1998; Panikov and Dedysh 2000; Silvola et al.2003),diur-
nal cycles (Chanton et al.1993; Mikkelä et al.1995; Thomas et al.1996; van
der Nat et al.1998),and episodic fluxes (Mattson and Likens 1990; Wind-
sor et al.1992; Christensen 1993; Frolking and Crill 1994; Grant and Roulet
2002; Christensen et al. 2003) have been related to effects of temporally
changing environmental factors, like weather conditions, on the basic
processes that affect CH
4
fluxes (Conrad 1989).Segers et al.(2001),Walter
et al.(2001a, b), Kettunen (2003), and Zhuang et al.(2004) have modeled
the CH
4
fluxes from peatlands. The model of Kettunen (2003) connects
CH
4
fluxes to microsite vegetation cover and water level throughout the
growing season.
The efflux of CH
4
C corresponds to less than 1 to more than 100 % of
the annual total C efflux from boreal peatland ecosystems (Svensson et al.
1975; Svensson and Rosswall 1984; Crill et al. 1988; Nykänen et al. 1998;
Heikkinen et al. 2002). The values exceeding 100% also include “old”CH
4
released from pools, mud-bottoms, and other wet or sparsely vegetated
sites that can lose C in many ways (Karofeld 2004).The logarithmic rela-
tionship between water level and CH
4
flux has been shown in many stud-
ies (Moore 2001; Heikkinen et al. 2003), while Nykänen et al. (1998)
showed this relationship to differ between ombrotrophic,minerotrophic,
and drained peatlands.As with CO
2
, an important part of the annual CH
4
flux (up to 40%) from boreal peatlands also occurs during winter (Wind-
sor et al. 1992;Dise 1993; Melloh and Crill 1996; Alm et al.1999a). The pro-
portion of winter flux increases toward drier site types (summer dryness)
and toward the north (long winters).
Kettunen et al.(1999) noticed that when the water level showed a down-
ward shift, CH
4
production and oxidation potentials in the layers that no
longer were water-saturated very slowly decreased toward rates more
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characteristic of permanently unsaturated layers.The decrease was faster
in wet microsites, where a 2-week period of unsaturation eliminated the
potentials,while in dry microsites, significant CH
4
production and oxida-
tion potentials were found after 6 weeks of unsaturation. These findings
imply that methanogens and methanotrophs are well adapted to natural
conditions where the water table shows both a seasonal cycle and short-
term fluctuation (von Fischer and Hedin 2002). Under laboratory condi-
tions, both methanogenic Archaea (Huser et al.1982) and methanotrophic
bacteria (Roslev and King 1994) have been shown to retain their viability
during periods of unsaturation and nutritional starvation. The results of
Kettunen et al.(1999) also indicate that methanogens and methanotrophs
are attached to peat particles and are not transported by vertical water
movements (van den Pol-van Dasselaar and Oenema 1999).The reactiva-
tion of potentials in Kettunen et al.(1999) and simulated CH
4
flux in Ket-
tunen (2003) depended on the length of the period of unsaturation, being
slow after a rise in water level if the populations of CH
4
-producing and
CH
4
-oxidizing bacteria had considerably decreased during the period of
unsaturation (Freeman et al.2002).
In addition to the direct effect on moisture and oxygen concentrations
in the peat profile, changes in water level may affect substrate levels.
After a decrease in water table position,increased aerobic degradation in
the unsaturated layers consumes the C compounds that in anoxic condi-
tions could promote CH
4
production. The reduction in substrate then
decreases the CH
4
production potential. The CH
4
oxidation potential
may also be reduced by a decrease in CH
4
concentration in the unsatu-
rated peat layers. On the other hand, a temporary rise followed by a
downward shift in water level may liberate CH
4
in deep layers (Moore
and Roulet 1993), so CH
4
oxidation may be reactivated by the substrate
peak (Kettunen et al. 1999).
Pore water CH
4
concentration builds up only gradually and causes a lag
before the CH
4
formed in the peat is released to the atmosphere (Chris-
tensen 1993; Shurpali et al.1993). The result in Kettunen et al.(1996) that
CH
4
fluxes did not correlate with a differentiated temperature series dur-
ing the early season strengthens the hypothesis that during the early sea-
son substrate availability is a dominant control for CH
4
fluxes. Later dur-
ing the season, the correlation between a differentiated temperature series
and CH
4
fluxes in Kettunen et al. (1996) and the higher production poten-
tials toward late summer in Kettunen et al. (1999) suggest that the higher
photosynthetic activity of plants had supplied methanogenesis with sub-
strates, so temperature effects became more evident. The methanogenic
activity that is found to be strongly temperature dependent (Dunfield et
al.1993; Segers and Kengen 1998) is argued to decrease at low autumn
temperature, so excess substrate may accumulate in peat (Saarnio et al.
1997). In the laboratory, the CH
4
production potential measured at a tem-
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perature higher than the in situ temperature activated the CH
4
-producing
bacteria, resulting in a high CH
4
production potential (Yavitt et al.1988;
Valentine et al.1994). The population dynamics of methanogenic popula-
tion in response to substrate supply may also contribute to the observed
increase in CH
4
production potentials late in the growing season (Svens-
son and Rosswall 1984; Dunfield et al.1993;Westerman 1993; Whiting and
Chanton 1993;Valentine et al.1994; Kettunen et al.1999).
9.5 Conclusions
Northern peatlands play an important role in the global C cycle. Further-
more, the CO
2
and CH
4
exchange rates between peatlands and the atmos-
phere show large variation (see review tables in Crill et al. 2000; Blodau
2002; Heikkinen 2003). Consequently, C cycling of northern peatlands is
crucial to global C cycle dynamics under both current and future condi-
tions. Although peatlands have shown positive long-term average net
ecosystem exchange sequestration of large amounts of atmospheric C
during the past few thousand years, the actual present-day C sequestra-
tion may be positive during a wet year and negative during a subsequent
dry one (Alm et al. 1999a, b; Griffis et al. 2000; Waddington and Roulet
2000; Heikkinen et al. 2002; Bubier et al. 2003a; Lafleur et al. 2003). This
variability suggests that the peatland C cycle is susceptible to environ-
mental change (Moore et al. 1998; Malmer and Wallén 2004) and in the
short term, the C exchange in mires depends on the rates of photosynthe-
sis and respiration, each of which is affected by short-term variation in
environmental factors (Bubier et al.2003b).
Even though peatlands on average have acted as C sinks, they simulta-
neously are the most important single CH
4
source, globally. High-latitude
northern peatlands, most of which lie within the boreal zone, are sug-
gested to contribute 34–60 % of the global wetland CH
4
emissions
(Matthews and Fung 1987; Cicerone and Oremland 1988; Aselman and
Crutzen 1989; Bartlett and Harriss 1993). Because CH
4
has a global warm-
ing potential (GWP) factor of 23 in relation to CO
2
over a 100-year time
horizon (IPCC 2001), boreal peatlands may have a considerable warming
influence over a 100-year time horizon. Over a longer (e.g., 500-year) time
horizon, with a GWP of 7 for CH
4
, the same peatlands may act predomi-
nantly as net sinks owing to the dominance of long-lived CO
2
molecules.
Thus, in the long term, the development of peatlands contributes as a
mediator of or even as a positive feedback for atmospheric trace gas con-
centrations (Prinn 1994; Gajewski et al. 2001; Whiting and Chanton 2001;
Friborg et al. 2003). Considering human-induced climate change, ecohy-
drological changes may be considered to be the primary driving force in
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the changes of the peatland C cycle (Gorham 1991; Roulet et al. 1992; Laine
et al. 1996; Moore 2002; Turetsky et al. 2002; Belyea and Malmer 2004;
Christensen et al. 2004; Malmer and Wallén 2004; Shurpali et al. 2004;
Strack et al. 2004; Fig.9.2).
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