The response of heterotrophic CO2 flux to soil warming

Global Change Biology (2005) 11, 167–181, doi: 10.1111/j.1365-2486.2004.00878.x

The response of heterotrophic CO2 flux to soil warming P E T E R E . E L I A S S O N *, R O S S E . M C M U R T R I E w , D AV I D A . P E P P E R w , M O N I K A ˚ GREN* ¨ M G R E N z, S U N E L I N D E R z and G O ¨ RAN I. A STRO *Department of Ecology and Environmental Research, Swedish University of Agricultural Sciences, PO Box 7072, SE 750 07 Uppsala, Sweden, wSchool of Biological, Earth and Environmental Sciences, University of New South Wales, Sydney NSW 2052, Australia, zSouthern Swedish Forest Research Centre, Swedish University of Agricultural Sciences, PO Box 52, SE-230 53 Alnarp, Sweden

Abstract In a forest ecosystem at steady state, net carbon (C) assimilation by plants and C loss through soil and litter decomposition by heterotrophic organisms are balanced. However, a perturbation to the system, such as increased mean soil temperature, will lead to faster decay, enhancing CO2 release from decomposers, and thus upsetting the balance. Recent in situ experiments have indicated that the stimulation of soil respiration following a step increase in annual average soil temperature declines over time. One possible explanation for this decline may be changes in substrate availability. This hypothesis is examined by using the ecosystem model G’DAY, which simulates C and nitrogen (N) dynamics in plants and soil. We applied the model to observations from a soil-warming experiment in a Norway spruce (Picea abies (L.) Karst.) stand by simulating a step increase of soil temperature. The model provided a good qualitative reproduction of the observed reduction of heterotrophic respiration (Rh) under sustained warming. The simulations showed how the combined effects of faster turnover and reduced substrate availability lead to a transient increase of Rh. The simulated annual increase in Rh from soil was 60% in the first year after perturbation but decreased to 30% after a decade. One conclusion from the analysis of the simulations is that Rh can decrease even though the temperature response function for decomposition remains unchanged. G’DAY suggests that acclimation of Rh to soil warming is partly an effect of substrate depletion of labile C pools during the first decade of warming as a result of accelerated rates of mineralization. The response is attributed mainly to changing levels of C in pools with short time constants, reflecting the importance of high-quality soil C fractions. Changes of the structure or physiology of the decomposer community were not invoked. Therefore, it becomes a question of definition whether the simulated dynamics of the declining response of CO2 release to the warming should be named acclimation or seen as a natural part of the system dynamics. Keywords: acclimation, carbon storage, century, ecosystem model, feedback, G’DAY, global warming, Q10, soil respiration, soil warming

Received 18 December 2003; revised version received and accepted 2 June 2004

Introduction The dynamics of plant production and decomposition are of crucial importance for understanding the terrestrial carbon (C) cycle. A major concern is that future global warming may lead to loss of soil C. Firstly, the C content of soil organic matter (SOM) in the Correspondence: Peter Eliasson, tel. 1 46 18 672458, fax 1 46 18 673430, e-mail: [email protected]

r 2005 Blackwell Publishing Ltd

terrestrial biosphere is about twice the content of the atmosphere (Ciais et al., 1995; Schimel, 1995). Small changes in the size of the global soil C pool can therefore have a considerable impact on the atmospheric [CO2]. Secondly, rates of decomposition determine the availability of nutrients to plants and can limit net primary production (NPP) in a nutrient-limited environment. Thus, rates of accumulation of SOM in the boreal zone, where NPP is normally limited by nutrient availability (Tamm, 1991), are sensitive to 167

168 P. E . E L I A S S O N et al . factors limiting decomposition. Soils in northern biomes are important because 24% of global terrestrial C is stored in the cool soils of tundra and boreal forest alone (Schlesinger, 1977). Some key factors regulating net soil respiration on the forest floor are temperature, moisture and nutrient availability, with temperature being the most significant factor in boreal forests. Temperature directly and indirectly regulates the activity of a range of interacting individual processes (Rustad et al., 2000). By definition, net C assimilation by plants and decomposition of SOM by heterotrophic organisms are balanced in a steady state forest ecosystem. By the term steady state, we refer to a quasiequilibrium state where the long-term average of litter C input from vegetation to the soil equals long-term average C output by heterotrophic respiration. It has been suggested that C assimilation is less sensitive to temperature than soil heterotrophic respiration (Kirschbaum, 1995, 2000) so that future global warming may induce a positive feedback from terrestrial ecosystems (Woodwell et al., 1995; Cox et al., 2000; Davidson et al., 2000). Other findings, which may appear contradictory, have indicated that decomposition of old SOM depends weakly on temperature (Liski & Westman, 1997) and that decomposition in the mineral soil does not vary with mean annual temperature (Giardina & Ryan, 2000). These observations were made over geographical temperature gradients on stabilized soils in established ecosystems and therefore give little insight into the transient response to a temperature perturbation. During the transient phase, responses of heterotrophic respiration to altered temperature may vary in magnitude and even direction. Such developments may depend on the initial quality and quantity of SOM and litter, complicated by different characteristics of the temperature signal and variable distribution of sub˚ gren, strate quality over the gradients in question (A ˚ gren & Bosatta, 2002). 2000; Davidson et al., 2000; A Radiocarbon data have shown that the average age of C in SOM is greater than the average age of the CO2 lost in respiration (Trumbore, 2000). This indicates that new plant C entering soil decomposes faster than older fractions of soil C. Results from litter bag experiments point in the same direction: decomposing plant litter gradually becomes enriched in recalcitrant chemical components (Berg et al., 2000). This difference between young and old soil C may have implications for the temperature response of soil respiration. For instance, Christensen et al. (1999) claim that the response of CO2 release to temperature decreases with increasing soil depth. They proposed that this decline could be explained by the age of SOM, based on the observation that the proportion of recalcitrant C increases with

increasing soil depth. Other experiments have shown that temperature response on decomposing wheat straw in incubated forest soils depends on biodegradability of SOM remaining during the process of decomposition (Dalias et al., 2001). Recent findings from soil-warming experiments have shown that the response of soil CO2 efflux to a step increase in temperature declines over time (McHale et al., 1998; Luo et al., 2001; Stro¨mgren, 2001; Melillo et al., 2002). Such experiments have shown that soil respiration declines under prolonged warming and eventually returns to rates similar to unwarmed soil. Consequently, extrapolations of the relationship of soil CO2 efflux to temperature on a given soil, as measured prior to warming, will overestimate the flux from the same soil if the mean temperature has been elevated. Carbon budget models that ignore this so-called acclimation of soil respiration may therefore overestimate soil C loss in response to global warming. The aim of this study is to investigate the mechanisms underlying the observed acclimation of total soil CO2 efflux in a boreal ecosystem to elevated soil temperature. We use the mechanistic plant–soil model G’DAY as a tool to simulate C and nitrogen (N) dynamics in the ecosystem as it shifts from steady state on unwarmed soil towards a new steady state on warmer soil. By simulating a steady-state system, we avoid temperature responses from pre-experimental perturbations that still may be in effect. Observations of CO2 efflux from a soil-warming experiment in forest stands of Norway spruce in Northern Sweden (Stro¨mgren, 2001) serve as a reference for our modelling study.

Hypothesis The above experiments indicate that the temperature dependence of decomposition is sensitive to changes in substrate quality. They suggest that changes in substrate availability may partly explain why the soil CO2efflux rates appear to acclimate to elevated soil temperature and that the acclimation process should be linked to the quantity of substrate in labile C pools. We hypothesize that all soil C fractions respond to increased temperature by accelerated rates of decomposition and enhanced soil respiration. However, the CO2 release will gradually decrease because there is a depletion of substrate available to decomposers, with the more labile fractions decreasing faster. As a consequence of the substrate depletion the CO2 evolution will, in spite of the elevated temperature, decrease to levels comparable with those at ambient temperatures. In the long term, the accelerated decomposition r 2005 Blackwell Publishing Ltd, Global Change Biology, 11, 167–181

S O I L WA R M I N G rate leads to increased NPP because the loss of soil C is associated with a soil nutrient release. This enhanced production generates increased litter input, which subsequently will counteract substrate loss. Key issues that we will address are: (1) Can the simulated decline in C substrate availability under increased rates of decomposition reproduce the acclimation of soil CO2 flux as observed experimentally (Stro¨mgren, 2001)? (2) To what extent is acclimation associated with labile vs. nonlabile soil pools and litter vs. SOM pools? (3) Do rates of litter C input change in response to soil warming and does altered litter input contribute to the acclimation process?

Materials and methods

Reference experiment Observations used in this study are published data from a soil-warming experiment (Stro¨mgren, 2001; Stro¨mgren & Linder, 2002) commenced in 1995 on the growth of Norway spruce (Picea abies (L.) Karst.) at Flakaliden (64107 0 N; 19127 0 E, alt. 310 ma.s.l.) in northern Sweden (Bergh & Linder, 1999). The soil at the site is a thin podzolic, sandy, glacial till with an average depth of 120 cm and a mean thickness of the organic layer of 4.3 cm. The site was planted in 1963 with ca. 2500 seedlings ha1 after the harvest of a naturally generated stand, followed by prescribed burning and soil scarification. The present stand is most likely the first generation of a managed forest from start on this site. Stocks of litter and soil C to the depth of 25 cm have been estimated to ca. 45 t ha1 (Andersson et al., 2002) and 58 t ha1 to a depth of 30 cm (Persson T & EU-

June

July

Control Aug.

Sep.

Oct.

169

FORCAST database, 7/7 2004, personal communication). To our knowledge, there are no available data on pool quantities reflecting decomposability on the studied site that can verify the balances and transitional changes of pools from steady state to the present state of possible disequilibrium. The Flakaliden research site started as a nutrient optimization study in 1987 (Bergh & Linder, 1999). In 1995, a soil-warming experiment with two replicates was applied on one fertilized and one nonfertilized treatment. The warming treatment was applied by maintaining soil on warmed plots at 5 1C above the temperature in unwarmed plots by means of heating cables (Bergh et al., 1999; Stro¨mgren, 2001; Stro¨mgren & Linder, 2002). The warming was extended in spring and autumn as described in Stro¨mgren & Linder (2002) to lengthen the period of unfrozen soil. All plots were irrigated to keep soil water potential above 100 kPa in order to eliminate the effects of soil moisture on respiration rates. Starting from the fourth year of warming (1998), net soil CO2 efflux (Rnet) on the soil surface and above surface structural litter was measured continuously by an automated flux system and monthly by a portable respiration system. The measurements thus include heterotrophic, root and mycorrhizal respiration plus net flux from understorey, if any. In the fifth year of warming (1999), rates of soil CO2 efflux on heated plots were not significantly different from the control plots (Stro¨mgren, 2001). Daily fluxes of nocturnal (22:00– 02:00 hours) Rnet vs. soil temperature over a 5-month period in the fifth year of warming are shown in Fig. 1 for both unwarmed and warmed soil. Measured annual increment of stem wood production (m3 ha1 a1) responded to warming by an increase of 57% on fertilized plots and 115% on nonfertilized

June

July

Heated Aug.

Sep.

Oct.

µmol CO2 m−2 s−1

8 6 4 2 0

140 160 180 200 220 240 260 280 140 160 180 200 220 240 260 280 Day of year

Fig. 1 Daily night-time soil-surface CO2 flux (Rnet), taken from Stro¨mgren (2001), and daily average heterotrophic CO2 flux (Rh) during 5 months in the fifth year of soil warming treatment. Closed symbols show measured Rnet and open symbols show simulated Rh. r 2005 Blackwell Publishing Ltd, Global Change Biology, 11, 167–181

170 P. E . E L I A S S O N et al .

Modelling approach The G’DAY model simulates Carbon (C) and Nitrogen (N) dynamics in plant–soil ecosystems (Comins & McMurtrie, 1993). The plant production processes are described in McMurtrie (1991) and Medlyn et al. (2000). The decomposition model is based on the CENTURY model of soil C and nutrient dynamics (Parton et al., 1987, 1993). The model is briefly described below; further details can be found in the above references. In G’DAY, gross photosynthesis is calculated by taking into account the separate contributions of beam and diffuse radiation (Medlyn et al., 2000) and the biochemically based model of photosynthesis of Farquhar & von Caemmerer (1982). Thus, growth depends on canopy leaf area, radiation incident on the canopy, [CO2] and leaf N content. Plant respiration is assumed to be a constant fraction of gross primary production (GPP) and is not further separated into above- and belowground fractions. Biomass is accumulated in five plant pools: foliage, stem, branch, coarse roots and fine roots. The rate of plant N uptake depends on N mineralized from litter and SOM. Minimum values of C : N ratio in leaf, fine root and wood limit plant uptake of mineralized N. Plant N is accumulated in the biomass pools. Stem N is further distributed to mobile and immobile fractions so that mobile N in stem wood and N in

foliage can be retranslocated. These processes affect the C : N ratio in foliage, and hence the photosynthetic capacity as described in McMurtrie et al. (2000). The outflow of N from the trees into the soil is coupled to litter production and occurs at the current C : N ratio of each tissue, minus a proportion that is retranslocated prior to senescence. The soil model tracks C and N dynamics in four litter pools (surface structural (u), belowground structural (v), surface metabolic (m), and belowground metabolic (n)) and three SOM compartments (active (a), slow (s) and passive (p)) with contrasting decomposition time constants that have been derived empirically (Parton et al., 1987). The soil compartments, representing fractions of substrate C depending on degradability, are interlinked by C flows. The amounts of C transferred from one pool to the next are proportional to the amounts of C resident in each donor pool. A temperature activity factor f(Tsoil), shown in Fig. 2, scales all fluxes between the compartments (Parton et al., 1987; Comins & McMurtrie, 1993): Tsoil 7:19 : fðTsoil Þ ¼ 0:0326 þ 0:00351ðTsoil Þ1:652 41:748 Litter input to the soil, which is proportional to plant biomass pool sizes with different litterfall rates, occurs through the litter compartments (u, v, m and n). Carbon transfers from litter pools to SOM pools (a, s and p) depend on lignin and N concentrations in foliage and roots.

1.0

0.8

Activity

plots over a 6-year period (1995–2000) of warming (Stro¨mgren & Linder, 2002). After the 6-year period, N content of aboveground biomass in trees on heated plots was 50 kg ha1 higher than those on unheated plots, corresponding to increases of 50% and 10% on nonfertilized and fertilized plots, respectively. The uptake of N by plants was thus higher on warmed plots, which may explain why N concentrations in the soil solution were not affected by soil warming (Stro¨mgren, 2001). In the reference experiment, much data from the automated measurements on the nonfertilized plots were lost. Monthly measurements, however, indicated that the responses to soil warming on nonfertilized plots were no different from the response on fertilized plots when CO2 flux is seen as a function of soil temperature (Stro¨mgren, 2001). In the present study, we therefore used daily averages from the automated measurements on fertilized unheated and fertilized heated plots under the assumption that the daily data reflect the fluxes of nonfertilized plots similar to monthly data. We will refer to fertilized control as ‘control’ and fertilized heated as ‘warming’ from here on. For further details about the experiment, see Stro¨mgren (2001) and Stro¨mgren & Linder (2002).

0.6

0.4

0.2

0.0

0

Fig. 2

10 20 30 Soil temperature (°C)

40

Temperature activity factor used in G’DAY: Tsoil 7:19 : fðTsoil Þ ¼ 0:0326 þ 0:00351ðTsoil Þ1:652 41:748

r 2005 Blackwell Publishing Ltd, Global Change Biology, 11, 167–181

S O I L WA R M I N G The C flow diagram in Fig. 3 is derived from the CENTURY model (Parton et al., 1987, 1993). The rate of inflow to the active SOM pool (a) is proportional to the outflows of the nonlignified C compounds from structural pools (u and v) and the metabolic litter pools (m and n) plus recirculation of C from the slow (s) and passive (p) pools. Output from the active pool is largely allocated to the slow pool under the influence of a soil texture factor. Lignified C compounds from the structural litter pools enter the slow pool (fluxes Lfl and Lrl in Fig. 3), which represent an intermediate fraction of C that is physically protected and/or chemically resistant to decomposition. Output from the slow pool is largely allocated to the active pool with a small flux into the passive pool. Carbon transfers out of each pool, including heterotrophic respiration, are proportional to the amount of C resident in that pool. Total heterotrophic respiration (Rh) is calculated by adding the sum of respiratory fluxes of CO2 out of each pool. Mineralization of N depends on the quality and quantity of litter and SOM, decomposition, soil temperature and soil moisture content. Rates of gross N mineralization and N immobilization in soil are calculated

171

from soil C fluxes combined with equations for changes in C : N ratios of soil pools (McMurtrie et al., 2001).

Parameters, model initialization and characteristics of the simulated stand In this study, we have chosen to initialize the simulations with equilibrated C pools. This is a simplification aiming to separate the transient response of C release to soil warming from other possible disrupted inputs of C and N pools following harvest, site preparation and planting seedlings on the reference experiment. The parameters are derived for the site (see the ECOCRAFT database; Linder, 1995; Roberntz & Stockfors, 1998; Medlyn & Jarvis, 1999) and have previously been used in simulations of the Flakaliden site. Responses of NPP and N uptake to increased [CO2] on C storage in trees and soil on different time scales were examined in McMurtrie et al. (2000), showing the importance of soil N constraints to the long-term responses of CO2. The interactive effects of elevated [CO2] and increased temperature (in air and soil) on NPP on The Biology of Forest Growth (BFG) site in Canberra, Australia were compared with the Flakaliden site in Medlyn et al. (2000).

Rh Wood(w )

Roots(r )

Foliage(f ) Plant

Surf.struct.(u )

Belowstruct.(v )

L fl

L rl

Litter Surf.met.(m )

Belowmet.(n)

Active(a)

Passive(p)

SOM

Slow(s)

Fig. 3 Carbon flows in litter and soil in the G’DAY model. Plant compartments are unshaded, litter compartments are light grey and soil organic matter (SOM) compartments are dark grey. r 2005 Blackwell Publishing Ltd, Global Change Biology, 11, 167–181

172 P. E . E L I A S S O N et al . McMurtrie et al. (2001) examined the role of increased N mineralization and assumptions on plant N uptake, temperature and [CO2] to the response of ecosystem C storage. Key parameters used in the above three studies, as well as the present study are shown in Table 1.

Weather variables G’DAY uses the mean of maximum and minimum daily air temperatures as soil temperature. We denote this soil surface temperature because this input act on the temperature activity function. An 8-year sequence of mean air temperature recorded at the Flakaliden site in the years 1992–1999 was used as control soil temperature. The warming scenario was simulated by elevating the control temperature sequence by 5 1C. Growing season was hereby extended in autumn and spring, similar to the reference experiment (Table 2). The model was initiated by looping the climate sequence until the 8-year sum of net ecosystem production (NEP 5 NPPRh) was zero. We denote this as steady state, although it is technically quasiequilibrium because of the daily and seasonal variability within the 8-year sequence. The steady state thus simulates a mature unwarmed (control) stand on the experimental site, where soil inputs and outputs balance, and was run in parallel with the simulation of soil warming. The simulation start date in the first loop was set to January 1995 (shown as time zero in Figs 7 and 8). Climatic conditions during the initial 5 years in the simulations thereby correspond with the first 5 years of the experiment (Stro¨mgren, 2001) and then loop back to the first year in the climate sequence.

Results The simulated response of daily heterotrophic CO2 flux (Rh) and the observed net soil CO2 flux (Rnet) (Stro¨mgren, 2001), in the fifth year after warming commenced, is plotted vs. temperature in mineral soil for control and warmed plots in Fig. 4a. Figure 4b shows the simulation output of Rh plotted vs. surface soil temperature shown over the same time sequence of 5 months as in Fig. 4a. The simulated fluxes in Figs 4a and b are thus the same as the simulated fluxes shown in Fig. 1 but plotted vs. the temperature at different depths of the soil profile. Measured data in Figs 1 and 4 represent average nocturnal (22:00–02:00 hours) net soil CO2 flux (Rnet), whereas simulated data show average daily (24 h) heterotrophic CO2 flux (Rh). Measurements thus incorporate soil heterotrophic and autotrophic respiration from live coarse and fine roots, plus any net assimilation by mosses and other understorey if present

Table 1 Parameters of the G’DAY model for the Flakaliden site Parameter Nitrogen input through deposition and fixation Specific leaf area Light extinction coefficient m2 leaf (m2 ground)1 Slope of Jmax/N relationship Slope of Vcmax/N relationship Ratio of intercellular CO2/ atmospheric CO2 NPP/GPP ratio Carbon fraction of biomass Plant allocation to Foliage Stem Roots Plant turnover Foliage Wood (stem, branch and coarse roots) Fine roots Maximum value of C : N in foliage Maximum value of C : N in fine roots Fraction of leaf nitrogen translocated before senescence Fraction of lignin in senescent foliage biomass g lignin (g DM)1 Fraction of lignin in senescent foliage biomass g lignin (g DM)1 Litter turnover* Metabolic litter (m 1 n) Structural litter (u 1 v) C : N ratio Metabolic litter (m 1 n) Structural litter (u 1 v) SOM turnover* Active fraction (a) Slow fraction (s) Passive fraction (p) Fine soil fraction (soil texture factor) C : N ratio Active fraction (a) Slow fraction (s) Passive fraction (p) Fraction of mineralized nitrogen lost through leaching

Unit

Value 2

gN m

1

yr

0.4

m2 (kg DM)1

2.5

0.6 mmol (g N)1 s1 mmol (g N)1 s1 Dimensionless

50.6 25.9 0.57

Dimensionless Dimensionless

0.4 0.49

Dimensionless Dimensionless Dimensionless

0.2 0.45 0.35

Years Years

11 50

Years g C (g N)1

1 29

g C (g N)1

29

Dimensionless

0.4

0.25

0.25 Years Years

0.05 0.5

g C (g N)1 g C (g N)1

150 22

Years Years Years Dimensionless

0.3 40 1000 0.72

g C (g N)1 g C (g N)1 g C (g N)1 Dimensionless

5 22 10 0.1

*Estimated turnover times for litter and SOM carbon fractions with a mean growing season soil temperature of 13.1 1C (McMurtrie et al., 2000). NPP, net primary production; GPP, gross primary production.


S O I L WA R M I N G Table 2

173

Differences between observations & simulations

Origin Carbon and nitrogen pool Age of trees at start of soil warming Interval of CO2-flux estimate CO2 flux Length of growing season* Time period Control Soil warming

Observations

Simulations

Planted with seedlings Perturbed by harvest and site preparation 35 years Nocturnal (22:00–02.00 hours) Surface net: Rnet

Naturally generated Steady state Mature 100 years 24 h average Heterotrophic: Rh

1995–1999 138 200

1992–1999 138 196

*Defined as the number of days in 1 year when the 5-day running mean air temperature is above 5 1C.

June

July

August

September

October

Observed R net

6 4 2 8

Simulated R h


8

6 4 2 0

0

5

10 15 20

0

5

10 15 20

0

5

10 15 20

0

5

10 15 20 0

5

10 15 20 25

Temperature in mineral soil (°C)

June

July

August

September

October

Simulated R h


8 6 4 2 0

0

5

10 15 20

0

5

10 15 20

0

5

10 15 20

0

5

10 15 20

0

5

10 15 20 25

Soil surface temperature (°C) Fig. 4 (a) Measured daily night-time soil-surface CO2 flux (Rnet) and simulated daily average heterotrophic CO2 flux (Rh) vs. temperature in mineral soil during 5 months in the fifth year of soil warming. Data are shown for both unwarmed (open symbols) and warmed (filled symbols) soil. (b) Simulated daily average heterotrophic CO2 flux (Rh) vs. soil surface temperature during 5 months in the fifth year of soil warming. Open symbols show flux from unwarmed soil and filled symbols from warmed soil.

inside the respiration chambers. See Table 2 for a comparison of differences between the experimental and the simulated systems. There is a seasonal variation in both the measured and simulated soil respiration rates but the seasonality is less pronounced in the simulations. Measured and simulated rates r 2005 Blackwell Publishing Ltd, Global Change Biology, 11, 167–181

deviate more from each other at mid-summer than in spring and autumn. In early spring, simulated Rh even exceed measured Rnet, indicating that Rh is overestimated. The simulated contributions, month by month, from different soil C pools to the simulated daily CO2

174 P. E . E L I A S S O N et al . effluxes are shown in Fig. 5. The differences in temperature response between control and warming treatments are most apparent in the litter pools and to a lesser extent in the active pool. Differences between treatments in the slow and passive (not shown) pools are hardly noticeable. There is also a seasonal trend in Fig. 5, showing more obvious differences in the summer and less in spring and autumn. One way of analysing the apparent acclimation is to calculate the difference between simulated Rh for control and warmed plots. This difference, DRh, is graphed on a daily basis in Fig. 6 vs. soil surface temperature on control plots. The difference varies seasonally and declines with increasing temperature in months June–September. Positive DRh indicate that warmed plots respire more than control plots, which occurs at cooler temperatures and early in the season. Heterotrophic respiration, Rh, from litter pools in warmed plots is less than in control plots on warm days

May

June

in June–September. This can be understood as follows. The respiration from a pool is a function of the pool size (C) and a specific rate constant (r), which is modified by the temperature activity factor f(Tsoil). The difference in respiration rate between control and warmed plots is then DRh ¼ rCw fðTsoil þ DTsoil Þ rCc fðTsoil Þ; where subscripts c and w indicate control warmed plots, respectively, and DTsoil represents the temperature elevation of 5 1C. Faster decomposition rates, caused by soil warming, have reduced the C pools on the warmed plots (Cw) over 5 years. Because more C has been lost on the warmed plots, Cw will always be less than Cc. At low temperatures, the temperature activity function is so steep that respiration from the warmed plots exceeds that from the control plots. At higher temperatures, on the other hand, the temperature activity function is less steep (or even decreases; see Fig. 2) so that the control plots have higher

July

August

September

October

Litter below (v+n)

0.8 0.6 0.4 0.2 0.0 0.8 0.6 0.4 0.2 0.0 0.8

Active (a)

R h (µmol CO2 m−2 s−1)

Litter above (u+m)

1.0

0.6 0.4 0.2 0.0 1.2

Slow (s)

1.0 0.8 0.6 0.4 0.2 0.0

0 5 10 15 20 25 0 5 10 15 20 25 0 5 10 15 20 25 0 5 10 15 20 25 0 5 10 15 20 25 0 5 10 15 20 25 30 Soil surface temperature (°C)

Fig. 5 Source of simulated heterotrophic flux, Rh, from selected pools as a function of soil surface temperature during 6 months in the fifth year of soil warming. Open symbols show control and filled symbols show warmed soil. r 2005 Blackwell Publishing Ltd, Global Change Biology, 11, 167–181

S O I L WA R M I N G

the belowground litter is faster than that of the surface litter because turnover rates for fine and coarse root litter are faster than those for foliage and wood. For active SOM, balance is also resolved between in- and outfluxes within approximately 10 years (Fig. 7c). This is not the case for slow SOM (Fig. 7d), which has a much longer time constant and therefore declines slowly (Fig. 7h). The initial amount of organic soil and litter C per unit of surface area in the simulated system at equilibrium was 109 t ha1. The effects of warming on the C storage in different pools are seen in Figs 7e–h. Litter pools have a strong annual amplitude because their decomposition is a function of temperature, which has a pronounced annual cycle. Influxes to litter pools, on the other hand, mainly depend on inputs from NPP in contrast to the outfluxes and has a weaker within-year variability. SOM pools have weaker annual amplitudes because both input and output fluxes have the same temperature dependence.

Discussion

Implications of comparing simulated Rh vs. measured Rnet The comparisons between simulated heterotrophic soil CO2 flux (Rh) and measured net soil CO2 flux (Rnet) made in this study are qualitative rather than

May

June

July

August

September

October

0 5 10 15 20

0 5 10 15 20

0 5 10 15 20

0 5 10 15 20

0 5 10 15 20

0 5 10 15 20

0.10 0.05 0.00

− 0.05

Litter below (v+n)

∆R h (µmol CO2 m−2 s−1)

Litter above (u+m)

respiration rates than the warmed plots. A complicating factor is the variability in C pools both within a single year and in the long term. It is only for those pools that have lost most C (i.e. the litter pools) (Figs 7e–f) that the control plots can have higher respiration than warmed plots on a given day (seen as negative values in Fig. 6). As discussed above, it is important to explain why litter and SOM change over time. These changes can be investigated by comparing the fluxes into and out of specific pools. Simulated total annual fluxes into and out of litter and SOM pools are shown in Figs 7a–d for the warmed plots relative to the control plots. Highlighted areas in the figures show the fifth growing season after warming commenced in comparison with Figs 1, 4 and 6. Simulations, with and without warming, start at the same quasiequilibrium prior to the soil warming that commence at time zero in Figs 7 and 8. The influxes (Cin) and outfluxes (Cout 1 Rh), where Cin and Cout are organic C transfers to and from pools, respectively, are simulated fluxes in the warmed plots minus fluxes in the control plots and therefore represent net response to warming. For the litter pools, the short-term response of outfluxes (Cout 1 Rh) to the elevated temperature lasts for approximately 10 years. After 10 years under soil warming, the in- and outfluxes of the litter pools closely balance (Figs 7a and b), and their C contents decline by approximately 65% relative to control (Figs 7e and f). The response of

175

0.10 0.05 0.00

− 0.05

Active (a)

0.10 0.05 0.00 − 0.05

Soil surface temperature (°C)

Fig. 6 The difference DRh in daily simulated heterotrophic flux (Rh of warmed soil minus Rh of control) for selected pools during 6 months in the fifth year of soil warming plotted against soil surface temperature in control plots. The difference in the slow pool (not shown) is little on the time scale of interest here (see Fig. 5). r 2005 Blackwell Publishing Ltd, Global Change Biology, 11, 167–181

176 P. E . E L I A S S O N et al .

(a)

Cout +Rh

80 60 40 20

C in 0

(e)

700 600 500 400 300 200 100 0

0

1

2

3

4

5 6 Years

7

8

9

Net responses (g C m−2 a−1)

Soil litter (v+n)

C in pool (v+n) (g C m−2)

C in pool (u+m) (g C m−2)


Surface litter (u+m)

60 Cout +Rh

40 20

C in 0

(f)

700 600 500 400 300 200 100 0

10

(b)

80

0

1

2

3

(c)

80 Cout +Rh

60 40

C in

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Fig. 7 Simulated differences in annual fluxes and pool sizes of litter, active SOM and slow SOM between warmed and control plots during the first decade of soil warming (a–d) and the resulting C store on simulated warmed plots (e–h). Highlighted areas show the fifth growing season after warming commenced. Cin is C entering, Cout is C transferred and Rh is heterotrophic CO2 release from respective soil pools.

quantitative because there are important differences between the modelled and the simulated systems, as shown in Table 2. The modelled stand represents a mature, steady-state system whereas measurements are from plots of 35-

year-old stands of planted trees after prescribed burning and soil scarification. These management practices may have caused transient changes in soil C (Johnson & Curtis, 2001; Yanai et al., 2003). When strictly quantitative comparisons between simulations r 2005 Blackwell Publishing Ltd, Global Change Biology, 11, 167–181

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Fig. 8 Components of simulated total Rh shown as the difference between warmed plots and control plots during the first decade of soil warming. Highlighted area shows the fifth growing season after warming commenced.

and measurements are made, it is important to include deviations from steady state in the initial conditions (Bruun & Jensen, 2002; Gu et al., 2004). However, estimates of how the soil C pools at Flakaliden may have changed before the soil-warming experiment are poorly known. Therefore, and also because it has been difficult to relate soil pools to functionally meaningful and experimentally verifiable fractions (Trumbore, 1997; Six et al., 2002; Wang & Hsieh, 2002), it would be necessary to estimate pools by simulating the preexperimental history of the site to reflect the present state. However, to simulate the distribution of SOM and litter pools at the time when the experiment started would introduce additional uncertainties. Instead, for the sake of comparability and clarity, we initialized the model using steady-state conditions from previous applications to the experimental site. Another difference between measurements and simulations is that the measured fluxes are from fertilized plots while the simulated fluxes represent a nonfertilized system. The response on nonfertilized plots were no different from the response on fertilized plots when CO2 flux is seen as a function of soil temperature, however. N uptake and stem wood production on nonfertilized plots increased in response to warming, also indicating that the fertilized stand still was nutrient limited (Stro¨mgren, 2001; Stro¨mgren & Linder, 2002). The qualitative responses can therefore be expected to be similar. Measured net soil CO2 fluxes (Rnet) include nocturnal heterotrophic and autotrophic respiration whereas the heterotrophic CO2 fluxes (Rh) represent the daily average of only the heterotrophic component. Comparisons of responses of Rnet with Rh are problematic because the components of Rnet may respond differr 2005 Blackwell Publishing Ltd, Global Change Biology, 11, 167–181

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ently to soil temperature than Rh (Boone et al., 1998). Furthermore, the fraction of Rnet that is autotrophic respiration in forest ecosystems is uncertain (Ho¨gberg et al., 2002). A comprehensive review of available methods over a range of forest tree species showed that the average contribution of root/rhizosphere respiration to soil CO2 flux was 49% (Hanson et al., 2000). Live root respiration tends to be the dominant fraction of soil CO2 flux in cold, northern biomes (Raich & Tufekciogul, 2000), where ratios of root respiration to total soil CO2 flux range from 62% to 89% in boreal forests. An estimate of root/rhizosphere respiration contribution in a pine (Pinus sylvestris (L.)) stand, adjacent to our reference site, was 54% (Ho¨gberg et al., 2001, 2002). Applying this ratio of belowground autotrophic respiration to our simulated fluxes of Rh makes simulated fluxes comparable with measured Rnet in our study, where simulated Rh over the 5-month growing season was 41% and 47% of measured Rnet in the control and warmed plots, respectively. This approximate belowground autotrophic respiration, seen as the difference between Rnet and Rh, is higher in mid-season as also found in the studies by Ho¨gberg et al. (2001), although the model overestimates Rh early in the season and possibly underestimates fluxes in late season because of the use of mean air temperature acting on the decay rates (Fig. 1). We acknowledge that a quantitative assessment of autotroph respiration based on simulated Rh must be carried out with caution, because tree species are likely to respond differently (Atkin et al., 2000), and also because the components of the total soil CO2 flux respond differently to raised temperatures (Boone et al., 1998).

Consistency with observations in the field Given the above uncertainties, there are features of the simulations that are qualitatively similar to the observations of acclimation. One similarity between measured Rnet and simulated Rh is that the response of the fluxes to soil temperature was reduced on warmed soil compared with fluxes on control after 5 years of warming (Figs 1 and 4). A key dissimilarity remains, however, because the magnitude and dynamic response of root respiration are included in measured Rnet but not in simulated Rh. There was no significant difference between Rnet on control and heated plots in the fifth year of warming (Stro¨mgren, 2001), suggesting that initial Rnet was restored to pre-experimental rates, whereas G’DAY over 5 years predicted a 50% reduction of the initial increase of Rh in the same time. Similar results are reported from the soil-warming experiment on mixed hardwood stands at Harvard forest in central Massachusetts, USA (Melillo et al., 2002). The soil

178 P. E . E L I A S S O N et al . temperature at Harvard forest was elevated approximately equally and accomplished with similar methods as in Flakaliden. Total soil respiration at Harvard forests increased by an average of 28% over the first 6 years of soil warming (Melillo et al., 2002). Root respiration at Harvard forests was estimated to account for 20% of total soil respiration in this temperate system. The increase of soil CO2 flux decreased to approximately 5%, averaged over the years 7–10, with no significant response to warming by the tenth year of warming (Melillo et al., 2002). The corresponding response of Rh in G’DAY was a 44% increase over the first 6 years, decreasing to 30% increase over the last 3 years of a 10-year simulation (Fig. 8). The overestimation of Rh at the start of the vegetation period seen in Fig. 1 can likely be explained by a lag in the soil temperature relative to the air temperature that is used to drive the simulation. The response may also be affected both because the temperature amplitude decreases with increasing soil depth (Rayment & Jarvis, 2000; Stro¨mgren, 2001) and because the response of C release to temperature may vary in the soil profile (Goulden et al., 1998; MacDonald et al., 1999).

The relative contribution of litter and SOM pools to the acclimation process The largest contribution to the apparent acclimation seen in the simulations was a result of responses of above- and belowground litter input (Cin) compared with output (Cout), plus CO2 release (Rh), as shown in Fig. 7. After 5 years of warming, the CO2 release from the litter pools was lower than under ambient temperatures at a given temperature during the warmer months of the growing season (Fig. 6). One reason that the active pool contributed less to the acclimation, in spite of its high turnover rate, is because in- and outfluxes to SOM-pools scale with the soil temperature activity factor, f(Tsoil). Accordingly, C input rates increase simultaneously with output rates in the active pool. Outfluxes from the litter pools are also proportional to f(Tsoil) but not the influxes, which are proportional to plant C, and thus respond to climatic variables and nutrient availability affecting tree growth. In terms of direct CO2 release, however, the slow and passive pools did not contribute to the acclimation (Fig. 8). The explanation is that C content of these pools decline little on the timescale of interest here, and hence Rh declines little. In fact, input to slow SOM declined faster than output (Fig. 7d) because input to the slow pool is coupled to the relatively smaller sizes of source pools, which have faster turnover times and therefore are more sensitive to temperature than the slow pool. Carbon output from the slow pool, despite its large C

content, does not change much in absolute terms because of its slow turnover. The reduction of slow soil C was only approximately 6% during the first 10 years of warming, compared with a loss of 38% in the most responsive aboveground litter pools over the same period (Fig. 7h).

Function and composition of microbial communities G’DAY does not take into account any possible shift in the function and composition of microbial communities related to soil warming. Experiments from a northern American hardwood ecosystem (Lakes States Spodosols) on stands of sugar maple (Acer saccharum Marsh.) indicate that warming may favour microbial populations that can metabolise substrate not used by populations at lower temperature (Zogg et al., 1997). It was thus suggested that the substrate utilisation increased and that the active microbial biomass pool decreased in response to warming during the incubation. Estimates of pool quantities to support the expanded utilisation of substrate C compounds were not measured directly, but were inferred from the firstorder kinetics relationship by assuming that the decomposition rate constant was less influenced by temperature than the amount of substrate (MacDonald et al., 1995; Zogg et al., 1997). The reference experiment used in this study, and thus simulations, were designed to exclude the influence of water limitation to soil respiration. Water limitation is known to influence soil respiration (Rustad et al., 2001) and also appears to influence the kinetics of microbial respiration differently depending on soil temperature: laboratory experiments with the Lakes States Spodosols mentioned above (Zak et al., 1999) suggest that the substrate demand of microbial activity is limited by the diffusion of soluble substrate when temperatures in incubated soil are high.

Model features A feature of the G’DAY model is that the temperature dependence has a maximum. The results presented in Fig. 5 can be less pronounced with other temperature activity functions (e.g. a Q10 function). Another feature is that all soil C fluxes are scaled by the same temperature activity factor. If the soil-decomposer system is seen as a discrete number of soil pools as in the G’DAY model, no parameters or properties in the description of the system need to be changed in order to describe qualitatively the dynamic response to soil warming. On the other hand, if the soil-decomposer system is seen as a single unit, it is likely that it is necessary to introduce some type of mechanism of r 2005 Blackwell Publishing Ltd, Global Change Biology, 11, 167–181

S O I L WA R M I N G downregulation to be able to describe the acclimation (e.g. a Q10 function varying with time). However, the assumption of how different fluxes respond to tem˚ gren (2000) perature is not trivial, as was shown by A ˚ and Agren & Hyvo¨nen (2002) and the relation between the temperature response of a substrate and its quality (turnover time) is far from being resolved. The problems of correctly identifying temperature responses experimentally as a consequence of varying substrate composition on an annual time scale have also been illustrated by Gu et al. (2004).

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patterns of Rh – acclimation may instead represent a natural system response by a multicompartment soil environment leading to a transient increase in heterotrophic respiration followed by a decrease. If this is so, then it is not necessary to invoke any changes in the structure or physiology of the decomposer community. Whether the dynamics of the response to the warming should actually be named an acclimation or seen as a natural part of the system dynamics therefore becomes a question of definition.

Future research Conclusions We have simulated a step increase in mean soil temperature in an N-limited boreal ecosystem where aboveground climate was kept constant and soil moisture was nonlimiting. The absence of relevant soil pool data has prevented us from making quantitative assessment without introducing additional uncertainties. The simulations were initialised in steady state. Under these circumstances, the simulations show that the model is consistent with the observed short-term changes in soil respiration induced by elevated soil temperatures. Our findings are similar to those of recent experiments (McHale et al., 1998; Luo et al., 2001; Stro¨mgren, 2001; Melillo et al., 2002) insofar that these changes – when viewed at the ecosystem or stand level – can be described as a downregulation, or acclimation, of the response of heterotrophic respiration to soil temperature. The simulation of a more complete history of the experimental stand is beyond the scope of this study as it would introduce the complication of initializing simulations of soil warming 35 years after the removal of stems and the burning of foliage and branches, while drastically increasing litter pools by leaving dead coarse and fine roots from a virgin preharvest stand. In most multicompartment ecosystem models, the initial distribution of C and N in pools is important for the simulated results. Initial sizes of slow pools are of particular importance because of their inertia induced by large quantities in combination with long turnover times (Bruun & Jensen, 2002). The simulated response suggests that what can be interpreted as an acclimation partly in fact is an effect of substrate depletion of labile C pools during the first decade of warming as a result of accelerated rates of mineralization. The response of Rh to raised soil temperature in the model is attributed mainly to changing levels of C in pools with short time constants, reflecting the importance of high-quality soil C fractions. Using this approach, no downregulation of temperature activity is required to explain observed r 2005 Blackwell Publishing Ltd, Global Change Biology, 11, 167–181

The simulations have shown that the model accounts for a reduced sensitivity of heterotrophic soil CO2 release to soil temperature in a nonwater-limited plant– soil system in equilibrium. Quantification of the reduction relies on comparisons of simulations with a stand where pools presumably are in a transitional stage. The complications of simulating the present conditions highlight the uncertainties surrounding sizes and turnover times of soil C pools over time. More precise assessments of the role of heterotrophic respiration in terrestrial ecosystems call for development of standard methods to relate turnover times to specific organic compound classes found in soils. Such experimental development has the potential to improve extrapolations of temperature responses of belowground autotrophic and heterotrophic respiration, contributing to assessments of consequences of global climate change.

Acknowledgements This work was financed by the Faculty of Forestry at the Swedish University of Agricultural Sciences (SLU). Part of the research was carried out during two visits to the School of Biological Earth & Environmental Sciences at the University of New South Wales in Sydney, Australia. The Bro¨derna Edlund Foundation and the Knut and Alice Wallenberg foundation are acknowledged for financial support. Thanks are due to Dr Michael Freeman for collaborative support, discussions and comments on earlier versions of the manuscript. Climate data were kindly provided by Chister Degermark at the Vindeln Experimental Forests at SLU. We also acknowledge support from the Australian Research Council.

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