4.1. Seasonal Patterns in Belowground CO2 and Soil Surface CO2 Flux
 Belowground CO2 and soil surface CO2 flux varied seasonally in these cold desert grasslands, ranging from near-atmospheric soil CO2 levels with near-zero fluxes in the winter to maximal, but still low, values in spring and summer when moisture was available. These patterns result from the well-established importance of moisture and temperature as first order controls on biological activity in deserts. The seasonal pattern of soil CO2 contrasts strongly with mesic forests [Jassal et al., 2005, 2008], which maintain much higher CO2 in the soil during the cold season. Our Utah sites share similarities but important differences with Mediterranean grasslands [Tang and Baldocchi, 2005; Xu et al., 2004]. Soil CO2 responded quickly to moisture pulses and decreased continually during dry periods in all these grasslands, but steady rains and milder temperature in the Mediterranean grasslands [Baldocchi et al., 2006] promoted much higher winter activity than in the cold deserts.
 The low soil respiration rates we observed are comparable in magnitude to those from other studies on the Colorado Plateau [Fernandez et al., 2006; Schaeffer and Evans, 2005], and to eddy covariance measurements of net ecosystem C exchange at Corral Pocket [Bowling et al., 2010]. However, they were much lower than soil respiration from Sonoran desert soils [Potts et al., 2006; Sponseller, 2007] and semiarid steppes of Mongolia [Chen et al., 2009] following experimental moisture pulse application. This was likely due to the low organic carbon content of the interspace soils (0.8–1.3%) and the relatively low productivity of these cold desert grasslands compared to other grassland sites [Bowling et al., 2010].
 When moisture was ample, respiration rate increased with seasonal temperature (days 1–150, Figure 4), and generally decreased as soils dried. Peaks in soil respiration in spring were consistent with the timing of maximum net ecosystem C uptake at Corral Pocket [Bowling et al., 2010] and followed the phenology of the grasses. Respiration increased immediately and substantially following moisture pulses in spring and summer, but considerably less so in winter (Figures 4 and 5). The short-term enhancement of soil respiration following moisture pulses, often called the Birch effect, is a general phenomenon observed in most ecosystems, even mesic forests [Borken et al., 2003; Irvine and Law, 2002; Jarvis et al., 2007; Munson et al., 2010; Savage et al., 2009]. In arid regions the increase can be stronger following addition of rain to soils that have been dry for some time [Austin et al., 2004; Cable et al., 2008].
 The respiratory response of dry soils to an initial pulse is often larger than subsequent pulses (see recent review by Borken and Matzner ). Comparing the similar pulses on days 135 (mid-May) and 170 (mid-June) at Corral Pocket (Figure 3) reveals a dramatic difference in the belowground CO2 increase. The peak values were similar, but the kinetics of each decrease were very different. The most likely explanation for this is one of substrate limitation. During long dry periods, respiratory substrates accumulate in the soil, and upon wetting additional organic compounds are released through microbial cell lysis. These substrates are consumed by surviving microbes during the first pulse and incorporated into the growing microbial community, and are thus not available as substrates for subsequent wetting events [Borken and Matzner, 2009; Fierer and Schimel, 2003; Kieft et al., 1987; Saetre and Stark, 2005]. Pulse size also seems to be an important controller of total C mineralized [Cable and Huxman, 2004; Misson et al., 2006; Munson et al., 2010; Sponseller, 2007].
 Rain enhanced not only the rate of soil respiration, but also its temperature sensitivity (Figures 5 and 6). Prior to late-May rain at Squaw Flat on day 154, there was almost no diel variability in soil CO2, but afterward a strong daily pattern was apparent. This involved short-term increases in Q10 from 1.1 to 1.6. While other factors influence the relation between temperature and soil respiration [Davidson et al., 2006a, 2006b], this increase is strong evidence for the highly dynamic involvement of the soil microbial community (including biocrusts). In dry soils, the increase following initial wetting involves activation and growth of a dormant microbial community, which requires time for osmotic adjustment, enzyme production, and diffusion of extracellular enzymes and their substrates [Fierer and Schimel, 2003; Stark and Firestone, 1995]. There can be changes in microbial community composition that occur over a short time frame following wetting, but these do not always occur [Fierer et al., 2003]. The change in temperature sensitivity may reflect changes in the availability of labile C rather than in the intrinsic sensitivity of the pool being consumed [Curiel Yuste et al., 2010].
 Many studies have shown that temperature sensitivity of respiration is a function of soil moisture over longer timescales [Borken et al., 1999; Curiel Yuste et al., 2007; Gaumont-Guay et al., 2006; Jassal et al., 2008; Subke et al., 2003; Xu and Qi, 2001]. These have generally involved data collected over several months that are binned for different moisture classes. To our knowledge, only two other studies have reported enhanced temperature sensitivity of respiration in response to individual wetting events. The first was an analysis using eddy covariance data from the Sonoran desert [Jenerette et al., 2008], and showed that this is a frequent phenomenon at the whole-ecosystem scale for sites with consistent summer precipitation provided by the monsoon. The second was an artificial irrigation experiment of biocrusts in the Kalahari desert [Thomas and Hoon, 2010]. Application of a tiny amount of moisture (1.4 mm) to biocrusts led to a short-term increase in temperature sensitivity of respiration, consistent with our observations. Under conditions of heavy wetting (120 mm water applied), the initial increase showed a Q10 which was much larger than reasonable (12.6) for a solely microbial response [Thomas and Hoon, 2010]. Other factors probably contribute to CO2 exchange following wetting, such as physical displacement of CO2 from soil pores by water, CO2 solubility in water, chemical interaction with carbonate minerals, and possibly adsorption on mineral surfaces [Ball et al., 2009; Parsons et al., 2004; Serrano-Ortiz et al., 2010].
 Net C uptake by biocrusts, and its dependence on sunlight, is unequivocal evidence of the dynamic involvement of the biocrust microbial community following wetting. The automated soil chambers captured several episodes of net C uptake following natural wetting events. These events were generally brief, lasting 2–3 days at most, and in most cases the daily average showed net C loss (Figure 4). The transition from photosynthetic dominance by biocrusts to net respiration by the soil biotic community occurred over a few days as soils dried (Figure 6). Enhanced temperature sensitivity of soil respiration was still apparent in the days following net crust uptake (Figure 6). Overall rates of C gain in biocrusts are determined by their species composition, as moss-lichen-dominated biocrusts from this region fix up to four times more C than cyanobacterially dominated biocrusts [Grote et al., 2010]. Livestock grazing can eliminate lichens and mosses from biocrusts, as they have low resistance to disturbance by compression [Belnap and Eldridge, 2001]. Soils in ungrazed grasslands of the region are covered by moss-lichen crusts [Belnap et al., 2006], and it is likely that the biocrust communities at our sites were of similar composition before the introduction of livestock. Therefore, biocrusts at our sites were possibly more significant C sinks in the past than we have observed because they likely included the late-successional mosses and lichens.
 Our results, and previous studies with biocrusts, have shown rates of net photosynthesis and respiration in biocrusts are strongly moisture dependent [Cable and Huxman, 2004; Lange, 2002, 2003a, 2003b; Lange et al., 1998], including some short-term field studies [Thomas and Hoon, 2010; Thomas et al., 2008; Wilske et al., 2008]. Whereas laboratory and field wetting experiments provide important process information, the in situ wetness in the natural environment is the primary determinant of process rates for biogeochemical cycling associated with crusts. Thus our results, as well as those from the 2 year study of rock lichens by Lange [2002, 2003a, 2003b] provide unique insight into controls on the activity periods and rates in biocrusts in the field. These studies underscore the extreme temporal variability of biocrust and soil community activity and highlight the need for continuous measurements in the field to understand soil ecological processes.
4.2. Belowground CO2 Among Plants and Plant Interspaces
 Using CO2 as an indication of belowground activity, we expected to find phenological differences in the physiology of the C3 (Stipa) and the C4 (Hilaria) grasses across the season. We further expected to find evidence of lower activity in the plant interspaces, as has been found in comparisons of soil respiration and belowground CO2 under C3 shrubs, C4 bunchgrasses, and interspaces between them [Barron-Gafford et al., 2011]. Neither was observed, indicating there is little or no difference in activity periods between these two species in southern Utah, and that the plant/interspace distinction may not be as obvious in our case as we assumed. Irrigation experiments have demonstrated that Stipa and Hilaria in this area can both utilize summer rain pulses, with enhanced photosynthesis and transpiration [Schwinning et al., 2003]. One might expect that these grasses, particularly the C4Hilaria, would respond to summer rains with increased belowground activity, through root growth [e.g., Cui and Caldwell, 1997], ion uptake respiration, or microbial activity (free-living or root symbionts) in response to root exudation.
 Although there were increases in belowground CO2 following summer rain pulses, the level of activity in summer was much lower than in spring (Figure 7). Neither the continuous CO2 measurements nor the flask measurements provided an indication that C4 grasses were more active than C3 grasses in summer. Eddy covariance observations over several years at Corral Pocket [Bowling et al., 2010], and experiments with precipitation exclusion shelters [Schwinning et al., 2005a, 2005b], have provided evidence that cold desert grasslands of Utah are predominantly dependent on winter precipitation. However, during years of abundant late-summer precipitation, both the C3Stipa and the C4Hilaria can be highly productive in late July through September (J. Belnap et al., unpublished data, 2010).
 The δ13C of respiration from C3 and C4 plants generally reflects that of their bulk tissues [Lin and Ehleringer, 1997; Sun et al., 2010]. There are important isotope effects associated with plant C allocation and metabolism [Bowling et al., 2008], but these are small relative to the difference in δ13C of leaf tissue of the C3 and C4 photosynthetic types. Changes in C3 and C4 composition are reflected in δ13C of respiration seasonally in Great Plains grasslands, including the semiarid shortgrass steppe [Lai et al., 2006; Shim et al., 2009; Still et al., 2003], in C3–C4 crop rotations [Rochette et al., 1999], and following conversion of tropical forests to C4 pasture [Neill et al., 1996]. Hence we expected the δ13C of respiration in locations dominated by C3 or C4 bunchgrasses to reflect their physiology, and that δ13CBR would indicate their relative activity at different times of the season. However, the observed δ13CBR was mostly intermediate between the C3 and C4 range for all treatments and did not reflect plant photosynthetic type in the simple fashion predicted (Figure 8). It is possible that the roots of these bunchgrasses overlapped more than we thought, an expectation which was based on the large amount of bare ground and horizontal distance between plants. We are not aware of studies which have examined the horizontal extent of roots in these species. A study of the congeneric Hilaria rigida reported horizontal rooting distributions that were maximally ∼1.5 m2 [Nobel and Franco, 1986], corresponding a radius of 0.7 m. Considering that 86% of our land surface was bare soil, their study provides some evidence for our assumption that the roots of C3 and C4 species were largely separated. We note, however, that vesicular-arbuscular mycorrhizal fungal spores in mixed C3–C4 plant systems are intermediate between the C3 and C4 leaf endpoints [Allen and Allen, 1990], suggesting that respiration from fungal hyphae in the interspaces may also be intermediate in isotope ratio (as we have observed for soil CO2). Fungal metabolism can lead to fractionation in δ13C of up to 4‰ [Hobbie et al., 2004], further complicating interpretation of δ13C of belowground CO2.
 Contrary to our expectations, there was no clear seasonal pattern in δ13CBR. On the sampling dates where δ13CBR was measured in late afternoon and again the next morning (such as days 191–192), there was a large diel change, with substantially more negative δ13CBR in the morning. These patterns were consistent regardless of whether the Keeling or Davidson method was used to calculate δ13CBR. It is highly unlikely that diel changes in δ13C of respiration of such a large magnitude actually occurred, particularly since the changes were quite similar for both C3 and C4 plants, and occurred also in the interspace. The soils at Corral Pocket are high in inorganic carbon [Bowling et al., 2010; Goldstein et al., 2005], and soil carbonates may play a role in short-term CO2 dynamics in arid soils [Serrano-Ortiz et al., 2010]. However, the δ13C of CO2 evolved from calcite precipitation would be more enriched (less negative) than the C4 range [Serna-Perez et al., 2006; Stevenson and Verburg, 2006].
 The most probable explanation for the observed diel pattern in δ13CBR is a temporal change in the relative gradients of 12CO2 and 13CO2 associated with non-steady-state diffusion [Nickerson and Risk, 2009a]. This could occur if the respiration rate alone, and not δ13C of respiration, changed over the diel course, such as driven by soil temperature. The isotope ratio of soil CO2 represents a mixture of CO2 in the air and CO2 from biological respiration, with additional isotope effects due to diffusion [Bowling et al., 2009; Cerling et al., 1991; Davidson, 1995]. Moyes et al.  used a model of soil diffusive transport to demonstrate that when the CO2 production rate is small, and the diel amplitude of the respiratory production rate is large, the δ13C of soil CO2 and of the soil surface flux can exhibit diel oscillations as large as 10‰. In our case, with low flux rates and large diel variability in flux rate, it appears that the isotopic transient associated with diffusion was large enough to obscure even the expected C3 and C4 isotopic patterns in soil CO2. Diel patterns in δ13C of soil respiration have been reported in a variety of ecosystems [Moyes et al., 2010] and are usually attributed to biological processes. A growing body of evidence indicates that the physical details of soil gas transport must be carefully considered to interpret isotopic patterns of soil CO2 [Moyes et al., 2010; Nickerson and Risk, 2009a, 2009b; Phillips et al., 2011; Risk and Kellman, 2008]. These studies suggest caution in interpreting isotopic data for soil CO2 in low flux environments.