Temperature controls of microbial respiration in arctic tundra soils above and below freezing

Soil Biology & Biochemistry 34 (2002) 1785–1795 www.elsevier.com/locate/soilbio

Temperature controls of microbial respiration in arctic tundra soils above and below freezing Carl J. Mikan*, Joshua P. Schimel, Allen P. Doyle Department of Ecology, Evolution, and Marine Biology, University of California Santa Barbara, Santa Barbara, CA 93106, USA Received 14 November 2001; received in revised form 13 June 2002; accepted 19 August 2002

Abstract Winter soil respiration can represent a significant fraction of the annual C cycle of arctic tundra, but the temperature response of microbial CO2 production in frozen soils is poorly understood. We used a short-term laboratory incubation to describe temperature effects on aerobic respiration in four organic tundra soils from Northern Alaska. Respiration conformed closely to simple first-order exponential equations with constant temperature dependence (r 2 ¼ 0.81 – 0.98). Temperature coefficients (Q10) increased abruptly with freezing, varying from 4.6 to 9.4 among the thawed soils (þ 0.5 to þ 14 8C), and from 63 to 237 among the frozen soils (2 10 to 20.5 8C). The single abrupt increase in temperature dependence with freezing suggests a shift in the dominant process controlling respiration below 0 8C. The Q10s of frozen soils are too large to represent the direct kinetic effect of temperature, and more likely reflect extracellular barriers to diffusion and/or intracellular desiccation. In thawed soils, respiration Q10s decreased with increasing soil organic matter quality, as indexed by baseline respiration at 0 8C, consistent with the thermodynamic argument that reactions metabolizing structurally complex, aromatic molecules have higher activation energies and temperature dependence than reactions metabolizing structurally simpler molecules. In frozen soils, Q10s were unrelated to baseline respiration, suggesting that freezing uncouples the direct link between the C chemistry of microbial substrates and the temperature dependence of respiration. Our results underscore the potential for warming to stimulate microbial activity and the turnover of C and nutrients contained in tundra organic soils. q 2002 Elsevier Science Ltd. All rights reserved. Keywords: Soil organic matter; Arctic tundra; Respiration kinetics; Frozen soil

1. Introduction Arctic tundra has historically been a strong C sink because low temperatures and poor soil drainage limit rates of decomposition (Adams et al., 1990; Gorham, 1991). Whether tundra regions are still accumulating C is not clear, however, as estimates of net annual C exchange with the atmosphere vary in direction and magnitude (Hobbie et al., 2000). Uncertainty over the current C balance of the arctic comes amid concerns that warming of C-rich tundra soils may start a positive feedback to the global climate system (Chapin et al., 2000). Climate projections expect warming in the arctic to exceed global average warming by a factor of 2 or 3 (Maxwell, 1992), with the potential to stimulate decomposition and the release of soil C to the atmosphere (Lashof, 1989; Vourlitis et al., 2000). * Corresponding author. Tel.: þ 1-805-893-4543; fax: þ1-805-893-4724. E-mail address: [email protected] (C.J. Mikan).

Until recently, most studies of C cycling in the tundra have focused on the brief growing season, under the assumption that biological activity is minimal during the cold arctic winter. However, the ability of cold-adapted microbes to survive and grow below 0 8C was discovered over a century ago (Gilchinsky, 1995), and respiration in frozen soils has been demonstrated repeatedly in the laboratory (Flanagan and Bunnell, 1980; Coxson and Parkinson, 1987; Clein and Schimel, 1995; Panikov, 1999). Microbial activity is possible in soils below 0 8C because small amounts of water remain unfrozen, allowing the diffusion of microbial substrates and waste products (Ostroumov and Siegert, 1996). Biological decay may be responsible for much of the litter mass loss occurring over the winter in arctic and other seasonally snow-covered systems (Bleak, 1970; McBrayer and Cromack, 1980; Moore, 1983; Hobbie and Chapin, 1996). Recent work has confirmed that, although rates are low, the cumulative winter CO2 flux from tundra soils may account for a significant component of their annual C budget (Zimov et al.,

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1993, 1996; Oechel et al., 1997; Fahnestock et al., 1998; Fahnestock et al., 1999; Grogan et al., 1999; Welker et al., 2000). It is becoming clear that accurate estimates of C storage and turnover in the tundra will require an understanding of the factors controlling respiration rates in the winter (Hobbie et al., 2000; McGuire et al., 2000). Temperature is one factor that is still poorly understood. Many studies that have reported sub-zero respiration can provide limited detail on the temperature response function because measurements were made at only one or two points below 0 8C. For example, Coxson and Parkinson (1987) found a shift in the apparent activation energy of respiration near the middle of a measured range from 2 10 to þ 18 8C; the shift could not be more precisely determined than between þ 1 and þ 6 8C, however, because rates were determined at 5– 6 8C intervals. A more likely transition temperature is 0 8C, below which liquid water content declines rapidly (Ershov, 1998). In other cases, respiration has been modeled using combined data from both frozen and thawed soils, an approach that assumes a single functional relationship that crosses the freezing point (Rosswall et al., 1975; Flanagan and Veum, 1974). More detail exists at higher temperatures, but the results are variable. Numerous studies have observed respiration in arctic soils or litter with strong and/or exponential responses to temperature (Flanagan and Veum, 1974; Bliss, 1975; Clein and Schimel, 1995; Hobbie, 1996). Others have found respiration to be insensitive to temperature, and have suggested that substrate quality is a more important control of respiration over the range of normal summer temperatures (Nadelhoffer et al., 1991; Schmidt et al., 1999). The temperature response of arctic microbes is part of a wider debate over the relationship between the temperature sensitivity of decomposition and the quality of soil organic matter (SOM). On the basis of enzyme kinetics, Bosatta and ˚ gren (1999) hypothesized that the decomposition of lowA quality substrates, i.e. recalcitrant polymers of older SOM, should have a greater temperature dependence than that of higher quality, labile organic matter. In contrast, two recent models of C turnover across gradients of mean annual temperature have concluded that the breakdown of older, recalcitrant organic matter is relatively insensitive to temperature (Liski et al., 1999; Giardina and Ryan, 2000). The assumptions and conclusions of these studies have been ˚ gren, 2000; Davidson et al., 2000), but despite challenged (A the relevance of this issue to the global C cycle, few experimental data exist to address it directly. We report here results of a laboratory study on the temperature dependence of microbial respiration in organic tundra soils of northern Alaska. The objective of our study was to describe respiration kinetics of these soils in sufficient detail to answer the following questions: (1) Is the functional relationship between respiration and temperature significantly different above and below 0 8C, or can respiration be effectively defined by a single equation? (2) Within a modeled range, is temperature dependence

constant or variable? (3) Is the temperature dependence of respiration inversely related to organic matter quality, as defined by respiration rate at a common temperature?

2. Materials and methods We collected soils in August 1998 from three tundra ecosystems at the Toolik Lake Long Term Ecological Research site in northern Alaska (688380 N, 1498380 W, elevation 760 m). Landform and vegetation of the area have been described by Walker et al. (1989). Wet meadow tundra is dominated by rhizomatous sedges and occupies river terraces and lake margins, as well as much of the arctic coastal plain. Moist tussock tundra occurs on flat or gently sloping terrain. The dominant plant Eriophorum vaginatum L. grows in hemispherically shaped clumps, or tussocks, that are elevated 15 –20 cm above the ground surface (Chapin et al., 1979). Intertussock areas between tussocks are dominated by evergreen mosses and deciduous shrubs. Moist shrub tundra occurs on hillslopes, high-centered polygons and water track margins, and are dominated by deciduous shrubs. Eight samples of each soil type were collected with a coring tube (3.8 cm diameter) from random locations approximately 3 m apart in each community. Organic horizons were sampled from three depths in wet meadow tundra (0 – 5, 5– 15, and 15 – 25 cm), corresponding roughly to Oi, Oe, and Oa horizons. In tussocks, Oe and Oa horizons were absent or very weakly developed. The organic horizon of tussock tundra was composed of Oi1 and Oi2 layers formed, respectively, from leaf and root litter, i.e. the ‘structured dead’ (Shaver and Cutler, 1979). We discarded the Oi1 and retained the top 10 cm of the Oi2 of tussock samples. Organic horizons of intertussock and shrub soils varied in depth from 5 to 15 cm and were sampled to mineral soil. In both soils Oi, Oe, and Oa layers were present. Before being placed in storage at 2 20 8C, the samples were composited, briefly hand sorted, and living plant materials were discarded. One week prior to the experiment in February 1999, 20 g sub-samples ðn ¼ 6Þ of each soil were used to determine gravimetric water contents and saturation water-holding capacities; after thawing, the sub-samples were weighed at existing moisture levels, at saturation, and a third time following drying to constant weight at 100 8C. Water contents ranged from 55% (tussock) to 82% (wet meadow 5 – 15 cm) of saturation capacities. We did not manipulate water content prior to the experiment because it was already close to the range considered optimal for respiration in these soils, i.e. 60 –80% of saturation (Flanagan and Veum, 1974), and because respiration within this range is relatively insensitive to moisture (Kielland and Schimel, unpublished data). Frozen samples (n ¼ 6; 30 g wet wt.) of each soil were sealed into serum vials fitted with rubber septa and placed in

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a glycol bath. Over 14 days, the bath temperature was raised by 2 8C every other day, from 0 to þ 14 8C. We measured CO2 production at þ 14, þ 12, þ 10, þ 8, þ 6, þ 4, þ 2, þ 0.5, 2 0.5, 2 2, 2 4, 2 6, 2 8, 2 10, and 2 12 8C. Because reaction rates of partly frozen mixtures can exhibit complex temperature dependencies (Fennema, 1975), we avoided measurements at 0 8C where bulk soil water freezes rapidly. At each temperature, headspace gas was sampled repeatedly with a syringe without removing the samples from the temperature controlled bath. CO2 concentrations were determined with a LI-COR 6252 infrared gas analyzer (Lincoln, Nebraska, USA) fitted with a 1 ml injection loop. We calculated the CO2 production rate at a given temperature as the mean rate over two successive intervals whose rates differed by less than 10%, i.e. lrate1 2 rate2l , 0.1p rate1. Because rates of CO2 production declined rapidly with temperature, samples spent relatively brief periods of time at the higher temperatures. It was necessary to incubate samples for increasingly longer periods as the temperature was lowered in order to measure reliable differences in headspace CO2 concentration. We measured rates from 14 to 8 8C over 2 days, from 6 to 2 8C over 5 days, and at the remaining seven lower temperatures over 20 days. To prevent inhibition of respiration, vial headspaces were periodically flushed with a stream of ambient air for 20 min when CO2 concentrations approached 2%. After the experiment, the samples were dried to constant weight at 100 8C, ground in a Wiley mill, and analyzed for C content with a Fisons NA 1500 C/N analyzer (Milan, Italy). We applied two non-linear regression models to the data. Both models describe an exponential increase in respiration with respect to temperature, but differ in their treatment of temperature dependence. The first model, a simple firstorder exponential equation, R ¼ A eðBTÞ ; has a single, constant temperature dependence over the entire modeled range. We fit this model separately to the respiration data above and below 0 8C (273.15 K), using the pooled data for each soil, i.e. single model parameters A and B were determined for each soil using the data from all six replicates of that soil type. The respiration coefficient, or Q10, is a widely used index of temperature dependence which describes the proportional change in rate given a 108 change of temperature. We used the pooled-data values of B p and the formula Q10 ¼ eð10 BÞ ; derived from the first model (van’t Hoff, 1898), to calculate respiration Q10s for each soil in both temperature ranges. To test the hypothesis that temperature dependence varies inversely with soil quality across soil types, we regressed the exponential parameters B vs. A derived from pooled data for each soil as described above. The exponential constant A was used as a simple index of soil quality, as A is equivalent to the rate of respiration at 0 8C, at which soil microbes are still fully hydrated and active. Tundra soils are near 0 8C for extended periods in spring and fall, and this temperature is close to the annual average

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temperature of the upper active layer (Osterkamp and Romanovsky, 1999). The first-order exponential equation was also fit individually to the respiration data of each sample, i.e. model parameters A and B were derived for each of the six replicates of each soil type. We used these individually derived B values in one-tailed, paired t-tests (n ¼ 6 for each test) to test the hypothesis that the temperature dependence of respiration in a given soil was greater at sub-zero temperatures, i.e. Bsub-zero . Babove-zero. The pooled data were also fit to a second exponential equation in which temperature dependence varies inversely with temperature: 0 R ¼ A0 eð2B =ðT2T0 ÞÞ (Lloyd and Taylor, 1994). The model parameters from both equations, i.e. A, A0 , B, B0 , and T0, were derived using temperature T on the Kelvin scale. For ease of interpretation, all results are presented in the Celsius scale except where otherwise noted.

3. Results Respiration continued down to 2 10 8C, and was highly sensitive to temperature in all soils (Table 1, Figs. 1 and 2). Both exponential equations were statistically significant for each soil and over both temperature ranges ( p , 0.01). Temperature coefficients (Q10) in the range from þ 0.5 to þ 14 8C varied from 4.6 (inter-tussock) to 9.4 (wet meadow 15 – 25 cm, tussock). The exponential parameter B was significantly greater ( p , 0.05) below than above 0 8C in all soils. The mean Q10 below 0 8C was 134 vs. 7.8 above 0 8C (Table 1). The separate temperature control of respiration above and below 0 8C is also apparent in the abrupt change in apparent activation energy, i.e. slopes of the lntransformed data (Figs. 1 and 2 inserts). Variable temperature dependence did not improve predictive ability; for a given soil and temperature range, goodness-of-fits (R 2) for models with constant and variable temperature dependence were identical or within 0.01 of each other (Table 1). A comparison of residual plots for each soil (not shown) indicated that where small under- or over-estimates of respiration occurred, they were generally shared by both models. In frozen soils, indices of SOM quality and temperature dependence (exponential model parameters A and B, respectively) were statistically unrelated (Fig. 3B). In contrast, temperature dependence in thawed soils tend to decrease with increasing SOM quality, both among subsamples of a given soil (small symbols Fig. 3A), and across soil type among the pooled data (large symbols Fig. 3A). Because the sub-samples were taken from composited field samples of each soil, for the regression analyses we used only the six independent estimates of A and B generated by the pooled data (large symbols Fig. 3A). When all six points were included, the negative relationship between the two parameters was significant at p ¼ 0.089 (r 2 ¼ 0.44). Excluding the outlying tussock soil, the relationship was

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Table 1 Model parameters for microbial respiration from exponential equations with constant or variable temperature dependence (Lloyd and Taylor, 1994). The equations were fit to data from each soil separately above and below 0 8C. Units of A are mg CO2-C g21 soil C d21, units of B are T 21 where T is temperature Constant temperature dependence: Respiration ¼ A e(BT ), T in 8C A

B

(þ 0.5 to þ14 8C) Wet meadow 0–5 cm Wet meadow 5–15 cm Wet meadow 15–25 cm Tussock Intertussock Shrub

77.725 67.209 20.107 70.476 90.484 76.941

0.162 0.185 0.224 0.224 0.153 0.159

(210 to 20.5 8C) Wet meadow 0–5 cm Wet meadow 5–15 cm Wet meadow 15–25 cm Tussock Intertussock Shrub

74.735 42.727 21.661 75.540 99.928 64.477

0.525 0.547 0.510 0.450 0.449 0.415

Variable temperature dependence: Respiration ¼ 0 A0 eð2B =ðT2T0 ÞÞ ; T in K

R2

A0

B0

T0

R2

5.1 6.4 9.4 9.4 4.6 4.9

0.94 0.81 0.89 0.98 0.98 0.98

1.633 £ 1012 1.301 £ 1014 2.873 £ 104 2.978 £ 1015 1.573 £ 1015 2.740 £ 1015

2.735 £ 103 1.987 £ 103 1.330 £ 102 3.334 £ 103 5.644 £ 103 5.597 £ 103

161 210 255 169 88 94

0.94 0.80 0.90 0.98 0.98 0.98

191.3 237.0 164.7 89.8 89.1 63.4

0.83 0.92 0.89 0.93 0.93 0.94

6.459 £ 1011 2.492 £ 103 3.387 £ 104 1.089 £ 106 5.170 £ 106 3.473 £ 1013

1.079 £ 103 6.772 £ 101 1.795 £ 102 3.384 £ 102 3.719 £ 102 1.839 £ 103

226 257 249 238 239 205

0.82 0.93 0.90 0.93 0.93 0.94

Q10

highly significant among the remaining five soils (line shown in Fig. 3A; p ¼ 0.004, r 2 ¼ 0.94). Within the overall pattern, the trend was also reflected in the depth sequence of the three wet meadow soils, which exhibited decreasing rates of respiration and increasing temperature dependence with depth (Fig. 3A).

4. Discussion 4.1. Controls of microbial respiration below 0 8C The separate temperature control of respiration in frozen soils is the most striking result from our study (Table 1, Figs. 1 and 2). Respiration was closely described by simple first-order exponential functions of temperature, but unlike thawed soils, the Q10s of frozen soils are too large to reflect the direct kinetic effect of temperature. Rather, the abrupt change in temperature dependence at 0 8C suggests that in frozen soils, temperature controls respiration indirectly, through effects on physical factors. For example, the unfrozen water content of soils below 0 8C declines rapidly as an exponential or power function of temperature (Lovell, 1957; Nakano and Brown, 1971; Romanovsky and Osterkamp, 2000), and likely limits the diffusion of substrates, nutrients, and waste products (Ostroumov and Siegert, 1996). The idea of extracellular diffusion limitations to biological activity in frozen soil is supported by experimental observations of a close correspondence between temperature-dependent thickness of unfrozen water films and microbial growth (Rivkina et al., 2000). Unfrozen water content may control the temperature response of respiration through intracellular mechanisms as well. The disappearance of liquid water with freezing creates an osmotic gradient that rapidly dehydrates

microbial cells (Franks et al., 1990). One possibility is that the resulting increases in intracellular solute concentration and pH raises activation energies by altering the conformation and/or substrate binding efficiencies of microbial enzymes (Raison, 1973; Gutfreund, 1995). Another possibility is that the high Q10s of sub-zero respiration reflect the disappearance of liquid water as a biochemical reactant; water participates directly in numerous reactions essential to life, e.g. the condensation and hydrolysis reactions of ATP and ADP (Lehninger et al., 1993). Whether or to what extent intracellular mechanisms operate together with extracellular diffusion limitations to define respiration kinetics below 0 8C is not known. Stark and Firestone (1995) found that mechanisms controlling nitrification changed along a gradient of declining soil water potential. Freezing may shift the balance of mechanisms controlling respiration kinetics in a similar way, with diffusion important over the first several degrees below 0 8C where unfrozen water content falls most rapidly, and with limitations caused by intracellular dehydration gaining control at lower temperatures. 4.2. Effects of C quality on temperature dependence We observed an inverse relationship between Q10 and baseline respiration rate in thawed soils (Fig. 3A). This finding supports the basic thermodynamic argument that enzymatic reactions metabolizing structurally complex, aromatic molecules have higher activation energies and temperature dependence than reactions metabolizing structurally simpler molecules. Respiration among the thawed soils, normalized for C content, differed by a factor of 4.5 (parameter A in Table 1). We interpret this as representing SOM over a range of C quality, i.e. soilspecific differences in fractions of easily decomposable

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Fig. 1. Rates of soil microbial respiration from 210 to þ 14 8C at three depths of a wet meadow tundra organic soil (n ¼ 6). Lines of best fit are shown separately for frozen and thawed temperature ranges, using the simple first-order exponential equation, Respiration ¼ A e(BT ). Inserts show the ln-transformed data.

molecules such as soluble C and cellulose, and of recalcitrant compounds such as lignin and humified organic matter (McClaugherty and Berg, 1987; Melillo et al., 1989). Our use of respiration as an index of SOM quality is consistent with measures of SOM chemistry at a nearby wet meadow (Gebauer et al., 1996); the lignocellulose index (lignin/(cellulose þ lignin)) of SOM there increased by 41% across the same depth sequence of organic horizons (0 –25 cm) in which we found a large decline in thawed-soil respiration (74%) (Table 1). It is important to note, however, that an inverse relationship between SOM quality and respiration Q10

may only apply to surface organic horizons, where C quality is a function of chemical recalcitrance and not of physical stabilization. Our findings agree with other studies of arctic and northern temperate soils, in which respiration Q10s of surface organic horizons increased with depth and degree of decay (Bunnell et al., 1977; Buchmann, 2000; Leiros et al., 1999). In deeper soils where mineral materials are present, however, the relationship may be reversed or absent (Christensen et al., 1999; Ka¨tterer et al., 1998). An hypothesis proposed by Thornley and Cannel (2001) to explain such diverging results suggests that the physical processes which stabilize organic C through adsorption to

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Fig. 2. Rates of soil microbial respiration from 210 to þ14 8C in three moist upland tundra organic soil (n ¼ 6). Lines of best fit are shown separately for frozen and thawed temperature ranges, using the simple first-order exponential equation, Respiration ¼ A e(BT ). Inserts show the ln-transformed data.

mineral surfaces may have higher temperature dependencies than the biological processes of microbial respiration. If this is the case, the Q10s of organic soils, which lack mineral content, reflect the true temperature dependence of respiration in the absence of processes that compete with microbes for organic substrates. Respiration kinetics of thawed tussock soils were outside of the pattern described by the remaining soils (Fig. 3A), a result which may be linked to the unique temperature and nutrient cycling regimes of tussocks. These have been described in detail by Chapin et al. (1979); the elevated growth form results in an earlier spring thaw, a longer growing season, and average soil temperatures that are 6– 8 8C warmer than surrounding

soils. Warmer conditions in turn favor high rates of productivity, microbial activity, and litter and nutrient turnover. The origin of the tussock soils may also have contributed to their respiration kinetics. In contrast to the other soils which developed from a mixture of aboveand below-ground litter of multiple species, the tussock soil (Oi2 layer) was formed almost exclusively from the dead, annual roots of a single, dominant species, E. vaginatum. Temperature response curves of soil respiration are in fact combined responses of many individual populations having varying temperature dependence and temperature optima (Flanagan and Scarborough, 1974; Heal et al., 1981). The respiration kinetics of tussocks may therefore be linked to characteristics of

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that Q10 becomes uncoupled from SOM chemistry below 0 8C because of strong changes in physical factors. The dissolved organic substrates that are respired by microbes are produced largely through the action of extracellular enzymes on polymers. As soils freeze and active microbes are limited to thin water films, the diffusion of substrates to microbes may become the rate-limiting step that defines respiration kinetics. If this is correct, a shift in substrate use could still occur with freezing, but simply would not affect respiration Q10s. 4.4. Constant vs. variable temperature dependence

Fig. 3. Relationships between the temperature sensitivity of microbial respiration and SOM quality for frozen (panel A) and thawed (panel B) temperature ranges. The indices A and B were obtained from the simple first-order exponential equation Respiration ¼ A e(BT ). The small symbols represent individual fits from each of six replicates of a given soil type. The large symbols represent the single fit parameters obtained for each soil type using all six replicates.

their microbial community, which developed under a warmer temperature regime and homogeneous organic matter pool. 4.3. Temperature dependence and microbial substrate use The C chemistry of SOM appears to influence the temperature sensitivity of respiration in thawed but not in frozen soils (Fig. 3B). One possible mechanism to explain this is that as soils freeze, microbial substrate use shifts from detrital material to dissolved compounds, dead microbial biomass, and products of microbial metabolism, as hypothesized by Clein and Schimel (1995). Since these substrates would likely be similar across soils we would expect the relationship between A and B to collapse. However, the use of a common substrate would also cause Q10s to be less variable below than above zero 0 8C, but in fact we found the opposite. An alternative explanation is

Our results are in agreement with recent studies reporting similar or less predictive ability in equations with a variable Q10 (Leiros et al., 1999; Reichstein et al., 2000; Buchmann, 2000; Fang and Moncrieff, 2001). In contrast, allowing temperature dependence to vary improved model fits and distribution of residuals when data from many studies were expressed in common units (Lloyd and Taylor, 1994; Kirschbaum, 1995). The difference may lie in the range of temperatures considered. The reviews compiled temperature responses over a total range of almost 40 8C, whereas our study and others modeled respiration over smaller intervals of 10– 208. Lloyd and Taylor (1994) attributed variable Q10 to temperature effects on enzyme activation and differences in microbial community composition. Subtle temperature effects on activation energy may only be apparent over larger intervals. Similarly, the short incubation times of our experiment likely prevented microbial community composition from changing in response to temperature. 4.5. Temperature dependence and components of soil organic matter Q10s based on field measurements or long-term laboratory incubations may not reflect the actual temperature dependence of microbial respiration. In the field, soil respiration is the product of both root and microbial respiration, which have different temperature responses (Boone et al., 1998). Other sources of bias include variation in soil moisture and CO2 storage in deep soils (Reiners, 1973). In the laboratory, Q10s are dependent upon incubation time. Long incubations may underestimate Q10 because labile substrate pools are more rapidly depleted at higher temperatures. (Reichstein et al., 2000). Temperature dependence of SOM decay should therefore be based either on short-term laboratory incubations to avoid substrate depletion or analysis of both time and temperature effects over longer intervals (Lomander et al., 1998; Ka¨tterer et al., 1998; Reichstein et al., 2000). Our approach was to measure the shortterm respiration response to temperature before appreciable change in substrate pools could occur. We varied incubation temperature over time, but made rate

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measurements rapidly in order to minimize confounding effects of substrate depletion. One advantage of this method is that all the samples have the same temperature history, allowing for comparison of kinetic parameters among soils. Short-term studies have been criticized in turn, for measuring temperature responses that may not represent the bulk of SOM (Townsend et al., 1992). Models of SOM distribution place most C in intermediate and passive pools which contribute little to respiration relative to a small pool of labile C (Jenkinson and Raynor, 1977; Parton et al., 1987). These models were developed for temperate soils, however, and may be less useful in characterizing the SOM structures of organic tundra soils. Most evidence indicates that the available C pool is a relatively large component of tundra SOM that is not rapidly exhausted. For example, respiration rates continued without systematic change for 13 weeks in the study of Nadelhoffer et al. (1991). Similarly, Weintraub and Schimel (2002) found that up to a third of SOM was respired over the course of an entire year without an appreciable change in rate. In a pilot experiment to the current study, we found that respiration at 5 8C declined rapidly for several days after thawing but by day 10 was nearly constant, having fallen to 32 – 37% of its initial rate (not shown). This pattern suggests that respiration was supplied by a relatively large and uniform component of SOM after the initial disappearance of a very small, hyper-labile substrate. In the current experiment, respiration over the 2 week pre-incubation period consumed a mean 0.47% of soil organic C, which should have exhausted most or all of this small C pool before temperature treatments began.

5. Conclusions Our results underscore the potential for warming to stimulate microbial activity and CO2 efflux from tundra surface soils, in both the thawed and frozen portions of their annual temperature range. Respiration Q10s from 0.5 to þ 14 8C were significantly greater than those of temperate and tropical soils (Kirschbaum, 1995), which may be a common feature of organic soils from cold environments (Flanagan and Veum, 1974; Clein and Schimel, 1995; Niklinska et al., 1999). Moreover, the inverse relationship between respiration rates and Q10 in thawed soils indicates that older, more recalcitrant components of tundra SOM are still sensitive to temperature and should be considered a potential source of respiratory CO2 under warmer conditions. The loss of this relationship and the abrupt increase in temperature sensitivity with freezing suggest a fundamental change in the processing of SOM below 0 8C, such as a shift to diffusion control of microbial activity, or a shift in microbial substrate use. Regardless of

the mechanisms involved, there are several implications of high sub-zero Q10s. One is that the use of equations developed for thawed soils with lower Q10s will lead to large over-estimates of winter C efflux. Another is that the percentage of annual CO2 production occurring over the winter may be increased by a relatively small rise in the temperature of frozen soils. This should be most significant close to 0 8C where respiration rates are still high. As surface temperatures fall below freezing in early winter, the portion of total soil respiration produced by organic horizons should decline in comparison to that of deeper soils, which remain warmer and have lower Q10s (Michaelson and Ping, 2002). Respiration Q10s describe the direct temperature effect on C loss, but may also influence ecosystem C balance through indirect effects on production (Shaver et al., 2000). Net primary production in tundra is limited by the availability of N and P, stocks of which are large but overwhelmingly bound in SOM (Chapin et al., 1980; Shaver and Chapin, 1995). Because nutrient release from decaying litter is largely the result of microbial activity, respiration Q10s should also predict the temperature response of nutrient availability. Respiration Q10s may actually be conservative estimates of temperature effects on the mineralization of N in particular (Nadelhoffer et al., 1991; Hobbie and Chapin, 1996; Rustad et al. 2001). The extremely high Q10s of frozen soils suggest that winter temperatures may be as or more important to nutrient cycling than temperatures over the brief growing season. Mass and N loss of decaying tundra litter appear to occur primarily during the winter when soils are frozen (Hobbie and Chapin, 1996) to the extent that these are the result of soil microbial activity, as opposed to physical processes, they should be strongly influenced by temperature. This is supported by the observation that winter soil temperatures are a key regulator of microbial N mineralization and immobilization patterns in Alaskan tundra (Bilbrough et al., 2002). Climate models predict arctic warming to continue, but the extent of warming is expected to vary spatially, and some areas may even experience cooling (Chapman and Walsh, 1993; Maxwell, 1997). Further, because temperature effects on microbial activity interact with soil moisture (Bunnell and Tait, 1974; Heal et al., 1981), the degree to which decay rates respond to warming will depend on local topography and drainage (Nadelhoffer et al., 1997). Warming may promote C and nutrient turnover more readily in well-drained moist tundra, where favorable soil oxygen status already supports aerobic metabolism. In the saturated soils of wet tundra where microbial metabolism is primarily anaerobic, significant warming effects will likely be contingent upon improvements in soil aeration, whether through increased evapotranspiration or a change in local drainage patterns.

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Acknowledgements This work was supported by a grant from the U.S. National Science Foundation Office of Polar Programs ATLAS Program.

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