Plant Soil DOI 10.1007/s11104-014-2159-9

REGULAR ARTICLE

Temperature sensitivity of soil and root respiration in contrasting soils Alice Thurgood & Balwant Singh & Edward Jones & Margaret M. Barbour

Received: 1 October 2013 / Accepted: 23 May 2014 # Springer International Publishing Switzerland 2014

Abstract Background and aims Positive relationships between temperature and soil respiration rate are widely observed, but it remains unclear if the relationships are due to increases in soil organic matter mineralisation (Rom), or in root and rhizosphere respiration (Rroots), or increases in both. This study aims to determine the relative sensitivity of Rom and Rroots to temperature in soils with differing properties. Methods Taking advantage of the difference in stable carbon isotopic composition provided by C3 and C4 plants, we partitioned soil respiration into Rom and Rroots for two soils with contrasting clay mineralogy, pH and carbon content over a 24 °C temperature range (from 12 to 36 °C). Results The Chromosol (dominated by illite, with near neutral pH and low organic carbon content) showed an increase in the proportion of Rom with temperature, indicating an increase in the decomposition of soil organic carbon. In contrast, the Ferrosol (dominated by hematite and goethite, with acidic pH and high organic carbon)

Responsible Editor: Per Ambus. A. Thurgood : B. Singh : E. Jones Faculty of Agriculture and Environment, The University of Sydney, 1 Central Avenue, Australian Technology Park, Eveleigh, NSW 2015, Australia M. M. Barbour (*) Faculty of Agriculture and Environment, Centre for Carbon, Water and Food, The University of Sydney, 380 Werombi Road, Brownlow Hill, NSW 2570, Australia e-mail: [email protected]

showed no change in the proportion of Rom with warming, and a negative priming effect at the highest temperature. Conclusions The observed positive priming effect for the Chromosol and a negative priming effect for the Ferrosol are consistent with contrasting mineralogy, reflecting the relatively weaker bond strength between soil carbon and illites in the Chromosol compared to the Ferrosol. Keywords Carbon isotope . Soil respiration . Root respiration

Introduction Soil respiration is the second largest terrestrial carbon flux (Lamberty and Thomson 2010), and thus plays a significant role in the global carbon cycle. The relationship between increasing temperature and the rate of soil respiration is a central issue that remains uncertain. Dynamic global vegetation models predict that soil respiration will increase more than total net primary productivity in response to warmer temperatures (Thornley and Cannell 2001) and as a consequence, following the predicted trend in greenhouse gas emissions, the terrestrial carbon sink is expected to decline significantly (Cox et al. 2000). This is supported by the study of Kirschbaum (1995), who found a decrease in soil organic carbon contents with global warming and hence a positive feed-back mechanism in the global carbon cycle (Knorr et al. 2005). Soil mineralogy is a key factor in determining organic carbon contents of soil, its turnover time and fluxes between the soil surface and atmosphere (Torn et al.

Plant Soil

1997). Two mechanisms have been proposed for the stabilization of SOM by the clay fraction: (i) the physical stabilization of soil organic matter by rendering it inaccessible to microbes and enzymes, and (ii) the chemical stabilization by the formation of organomineral complexes (Kaiser and Guggenberger 2003; Six et al. 2004; Wiseman and Püttmann 2006; Jones and Singh 2014). In chemical stabilization, SOM is bound to mineral surfaces involving three types of interactions, i.e. ligand exchange, polyvalent cation bridges and weak interactions (Lutzow et al. 2006). Iron and Al oxides also play a key role in the stabilization of soil organic matter, involving both chemical and physical mechanisms. Additionally, due to their higher specific surface area and poorly ordered structure, Fe and Al oxides have much greater capacity to strongly adsorb organic carbon by ligand exchange mechanisms (Kaiser et al. 1997; Kleber et al. 2004; Lutzow et al. 2006). However a majority of detailed studies concerning stabilisation and sorptive fractionation of organic matter to mineral surfaces have been focussed in sediments rather than soil, limiting our understanding of the effects of mineral associated organic matter in soil environments (Mikutta et al. 2009). Moreover, temperature sensitivity of the mineralorganic interactions in soils has not been evaluated. Soil organic matter (SOM) is made up of different pools varying in rate of decomposition, with the recalcitrant carbon pool having a very slow rate of decomposition (approximately 200–1,500 years; Bol et al. 2003) and release as CO2 to the atmosphere. Respiration from the soil surface is the combined result of respiration from a range of sources, including decomposition of SOM by microbial activity, respiration from plant roots, algae and mosses (Czimczik et al. 2006). SOM decomposition comes from labile carbon and, under some circumstances, recalcitrant carbon. A study by Davidson et al. (2006) concluded that the decomposition of the more recalcitrant soil carbon may have higher temperature sensitivity than that of labile soil carbon. Although the flux of CO2 from the soil to the atmosphere is relatively simple to measure (e.g. Pumpanen et al. 2004), the separation of the total flux into SOM decomposition and root respiration is more difficult, recognised by Rustad et al. (2001). A useful technique for partitioning soil respiration is by measuring the stable carbon isotope composition (δ13C; symbols are defined in Table 1) of the soil respiratory efflux, if there is a systematic difference in δ13C of CO2 between SOM

Table 1 Symbols and abbreviations used in text Symbol or abbreviation

Definition

E0

temperature sensitivity coefficient

fom

proportion of CO2 efflux from soil organic matter decomposition

Rleaf

rate of leaf respiration in the dark

Rom

rate of soil organic matter mineralisation

Rroots

rate of root and rhizosphere respiration

Rsoil

rate of CO2 efflux from pots with soil only

Rsoil+roots

rate of CO2 efflux from pots with plants

R10

rate of respiration at 10 °C

SOM

soil organic matter

T0

reference temperature (227.13 K)

Ts

soil temperature, °C

δroots

δ13C of CO2 respired by roots and rhizosphere separated from the bulk soil

δsoil

δ13C of CO2 respired by pots with soil only

δsoil+roots

δ13C of CO2 respired by pots with plants

decomposition and respiration from roots and their rhizosphere. By growing C4 plants in soil from a C3dominated ecosystem (i.e. many agricultural soils and all forest soils), we can effectively partition soil respiration (Kuzyakov 2004). Using this technique, Uchida et al. (2010) found that respiration from SOM decomposition was insensitive to increasing temperature in an agricultural soil with low carbon content, while root respiration and SOM decomposition had similar temperature sensitivities in a high carbon content soil. That is, old soil carbon was not decomposed in the low carbon soil. In contrast, Vicca et al. (2010) observed a strong temperature sensitivity of SOM decomposition but no increase in root respiration with increasing temperature when exposed to long-term temperature treatments. The priming effect, originally described by Bingeman et al. (1953) as the increase in SOM decomposition as a result of the addition of fresh organic material, has been observed in numerous studies (reviewed by Kuzyakov et al. 2000; Kuzyakov 2010). However, decreases in decomposition rates have also been observed (Cheng et al. 2003; Cheng 1996). The priming effect leads to an increase in the heterotrophic component of soil respiration by enhancing microbial decomposition of mostly labile carbon pools (Blagodatskaya et al. 2011). The priming effect can be affected by plant species and phenology, photosynthetic rate, plant biomass and possibly soil temperature (Scott-Denton et al. 2006).

Plant Soil

The physical and biological processes that regulate soil respiration also contribute to diel variation in soil respiration rates. Short-term changes in rates of soil respiration may be attributed to variables with strong diel cycles, such as temperature, humidity, and photosynthetic rate. For instance, Liu et al. (2006) found that diel variation in soil respiration rate was related to diel variation in soil temperature, but also to inputs of carbon to the root system from canopy photosynthesis. Tang et al. (2005) also highlighted the importance of photosynthetic inputs, as well as temperature and soil moisture, in driving soil respiration. If root respiration and SOM decomposition differ in δ13C (as is the case in partitioning studies like that of Uchida et al. 2010 and Vicca et al. 2010), then diel changes in the proportional contribution of these components should cause considerable diel variability in δ13C of soil respiration. The observed degree of diel variability in δ13C of soil respiration differs considerably between different ecosystems, so that it is difficult to establish any general trends among the small number of studies published to date. Barbour et al. (2011) found no diel trend in δ13C from soils in a mixed shrub-grassland, while Bahn et al. (2009) recorded a diel variation of around 1‰ in a grassland ecosystem which was somewhat related to assimilate supply. A range of studies from woody ecosystems report no significant diel variability in δ13C (Betson et al. 2007; Millard et al. 2010), whereas others have observed either a peak in δ 13C during the night (Marron et al. 2009), or a peak in δ 13C during the day (Kodama et al. 2009). Moyes et al. (2010) concluded that diel variability in the field was driven by non-steady state gas transport effects, a conclusion supported by soil gas transport modelling. Measurements of diel trends in δ13C from soil alone and soil with roots conducted in controlled environments (i.e. to maintain constant temperature in the light and dark) may provide useful information on drivers of variation in δ13C, by allowing separation of gas diffusive and temperaturerelated variability from variability due to diel trends in root metabolism. This study examines the effect of increasing temperature on root and rhizosphere respiration and SOM decomposition in two contrasting soils using stable carbon isotopes to partition total soil respiration. The three aims of this study are; 1) to determine the temperature sensitivity of soil carbon fluxes in soils of contrasting properties including clay mineralogy, 2) to partition total soil

respiration into root (and rhizosphere) and SOM components and quantify these proportions and 3) to quantify diel variation in δ13C and respiration rate of the root (and rhizosphere) and SOM components. An understanding of the temperature sensitivity of soil respiration is essential in determining accurate carbon fluxes from the soil to the atmosphere and increasing the potential for soil to act as a carbon sink.

Materials and methods Soil sampling and preparation Two soils were collected from eastern Australia representing contrasting clay mineral compositions typical of Australia to a depth of 15 cm, including a Red Ferrosol (Ferralsol) from Woolongbar and a Red Chromosol (Luvisol) from Tamworth (Isbell 2002; FAO 2006). Both soils were taken from areas under natural C3 vegetation. The majority of roots were removed from the air-dried soils, the soil then was homogenised and packed into 4 L pots by volume. Soils were not sieved to minimise the level of disturbance to soil structure and microbes. Corn (Zea mays cv. Kelvedon Glory F1), a C4 plant, was germinated in potting mix and four plants transplanted towards the outside of half of the pots 1 week after sowing. PVC collars (14 cm in diameter) were inserted 10 cm into soil (between plants, if present) in all pots to allow for soil respiration measurement. Measurements at each temperature took 5 days, so pots for each temperature treatment were prepared weekly allowing all plants to be the same age when soil respiration measurements were made. All pots were maintained at 30/20 °C day/night temperature, 70 % relative humidity and 14 h light period in a controlled environment facility for 3 weeks prior to movement into treatment temperatures and measurement of respiration rates. All pots were wellwatered daily, initially over the whole surface area of the pot, and from week 2 in the central collar area. Urea fertilizer (46 % N w/w, 0.8 g per pot) was applied to all pots 1 week after transplanting, and a complete fertiliser (Aquasol, NPK—23: 3.95 : 14 % w/w, and micronutrients) applied 2 weeks after transplanting (0.5 g per pot). Both fertilisers were applied to the central collar only, to encourage growth of roots into the soil measured by the soil chamber system described below.

Plant Soil

environment room was unable to maintain 36 °C with the lights off. Soil chambers were sealed to collars of the first replicate pots 24 h after imposition of the treatment temperature, and remained in place for 24 h. That is, eight pots were measured per day; one of each soil type with plants and one each without. Thermocouples were inserted into the soil between the plants to a depth of 5 cm to provide continuous soil temperature measurements. A custom-built, flow-through, eight-chamber system based on the design described by Midwood et al. (2008), was attached to eight pots, four pots with soil only and the other four with plants over a 24-h period to measure the soil carbon flux of soil only (Rsoil) and soil with roots (Rsoil+roots), respectively. The first 2 h of flux and δ13C measurements were discarded to allow for the establishment of steady-state gas transport conditions. The values of Rsoil and Rsoil+roots are expressed as μmol CO2 per m−2 soil surface area inside the chamber per second.

Physical and chemical soil analysis Soil samples were analysed for relevant physical and chemical properties (Table 2) using the procedures described in Rayment and Lyons (2011). Mineral composition of the clay fraction was determined by the Xray diffraction analysis of both random powder and basally oriented samples. X-ray diffraction analysis was done using a GBC MMA diffractometer (CuKα radiation at 35 kV and 28.5 mA, step size 0.020° and speed 1.0° per minute). Dispersed clays were deposited on to ceramic tiles and the oriented samples were analysed following standard pre-treatments (Brown and Brindley 1980). Semi-quantitative estimation of phyllosilicates was done by comparing integrated peak areas of the clay minerals with the known mixtures of standard minerals (Table 2). Total soil respiration measurements Sixteen pots were transferred to a second controlled environment room at the start of each week when plants were 4 weeks old; four pots with plants and four pots without plants for each soil type. The second controlled environment room was maintained at the treatment temperature during the measurement week: 12/12, 18/18, 24/24, 30/30 and 36/30 °C day/night, 70 % relative humidity and 14 h light period. The controlled

δ13C measurement The δ13C of soil respiration from soil only (δsoil) and soil with roots (δsoil+roots) from the chambers were measured by a Picarro 13CO2 laser analyser (G1101-i, Picarro, CA, USA) plumbed in line after the LiCor CO2 concentration analyser (Li-840, LiCor, Lincoln, NE, USA) of the

Table 2 Physical, chemical and mineralogical properties of the two soils Texture

Ferrosol

Chromosol

Sandy loam

Clay loam

pH (1:5 H2O)

4.5

EC (dS m−1)

0.19

6.6

CEC (cmol(+) kg−1)

6.4

18.3

0.06

Exchangeable cations

Ca

3.2

13.0

(cmol(+) kg−1)

Mg

0.9

4.4

Na

0.0

0.2

K

0.3

0.7

OC (%)

7.0

2.1

Total N (%)

0.61

0.20

C/N −1

11.5

10.5

P (mg kg )

18

17

Sand (%)

26.2

40.1

Silt (%)

42.9

28.0

Clay (%)

30.9

31.9

Minerals in the <2 μm fraction

Hematite > gibbsite > kaolinite > goethite

Illite > kaolinite

Plant Soil

chamber system. The δ13C of CO2 respired by roots (δroots) was measured after completion of the efflux measurement. Roots were removed by hand from the soil, excess soil was gently removed, but the rhizosphere remained intact, and the roots sealed in Tedlar® bags. Air was removed from the bags, followed by flushing and filling with CO2-free, dry air. Roots were incubated until the concentration of CO2 reached a range between 400 and 1,000 μmol mol−1. Bags were then attached to the Picarro 13CO2 analyser and δ13C measured for 10– 15 min. These values formed the “root” end member in the partitioning calculations (Kuzyakov 2006). Measuring δ13C of root and rhizosphere respiration, rather than assuming this value is equal to δ13C of root organic matter, removes the need for assumptions regarding fractionation during respiration (Werth and Kuzyakov 2010). We directly assess the use of root organic matter to infer δ13C of root respiration in this experiment. The average δ13C over the last 5 min of measurement was used for further analyses for both chamber measurements and root incubations. The laser was calibrated daily using two standard cylinders of CO2 in air that had been previously measured for δ13C and CO2 concentrations at NIWA, Wellington, New Zealand. The cylinders had similar CO2 concentration but a wide range in σ2 ð f om Þ ¼

h

1 ðδsoil − δroots Þ2

Calculation of temperature sensitivity Temperature response functions of the fluxes were fitted using Origin edition 6.1 (Microcal Software, Northampton, MA) to the Arrhenius model as described by Lloyd and Taylor (1994);  Rom ðor Rroots Þ ¼ R10 e

Calculation of total soil respiration components Soil respiration rates were calculated from measurements of CO2 concentration and flow rates by the automated soil chamber system (Midwood et al. 2008). Total soil efflux (Rsoil+roots) was partitioned into respiration of soil organic matter (C3 soil carbon, Rom) and root and rhizosphere respiration (C4 carbon Rroots) components using a mixing model (Kuzyakov 2004); f om ¼

δsoilþroots − δroots δsoil − δroots

 1 56:02

−

1 T s þ273:15 − T 0

ð3Þ

where R10 (respiration at 10 °C) and E0 (temperature sensitivity coefficient) are fitted parameters, Ts

ð1Þ

where fom is the proportion of respiration from soil organic matter decomposition and δsoil+roots, δroots and δsoil represent the δ13C value for the total soil respiration, the root respiration and soil organic matter respiration (respiration from the pots with no plants), respectively. Variance in fom (σ2(fom)) was calculated as (Uchida et al. 2010);

σ2 ðδsoilþroots Þ þ f 2om  f 2om  σ2 ðδsoil Þ þ ð1 − f om Þ2  σ2 ðδroot Þ

where σ2 is the variance of the individual component as indicated. Variances of the partitioned fluxes were calculated from the sum of the variance of fom, the total flux, and the covariance between the two. Errors are presented as the standard errors of the mean.

E0

the δ13C values; with the values of −8.7‰ and 390.2 ppm, and −29.5‰ and 398.4 ppm, respectively, in the two cylinders.

i

ð2Þ

is soil temperature and T0 is a reference temperature (227.13 K). Photosynthesis and leaf respiration measurements The photosynthesis and dark-adapted leaf respiration rates were measured for one of the four plants per pot with a portable photosynthesis system (Li6400xt, LiCor, Lincoln, NE, USA) fitted with a standard 6 cm2 leaf chamber and blue-red light source. Leaf chamber light levels were set at 200 μmol m−2 s−1 to match measured photosynthetically active radiation two thirds of the way from the uppermost leaves (400 μmol m−2 s−1). Leaf temperature was controlled to match the five treatment temperatures and leaf chamber inlet CO2 was set at 400 μmol m−1; water vapour concentration was not controlled. Respiration measurements were taken under the same conditions with the light off. The leaf chamber

Plant Soil

enclosed a section of the leaf approximately half way along the length of the youngest fully-expanded leaf. For respiration measurements, the measured section of the leaf was dark-adapted by covering completely with tin foil for 30–60 min prior to placing in the chamber. Measurements were recorded when the coefficient of variation was less than 0.1 % (5–10 min after leaves were first enclosed) and the infrared gas analysers were matched every 20 min.

Ferrosol is highly acidic and had much higher organic carbon and total nitrogen (7.0 % and 0.61 %, respectively) than the Chromosol (2.1 and 0.20 %, respectively), although the C/N ratios are similar (11.5 and 10.5). The clay fraction of the Ferrosol is dominated by Fe and Al oxides (hematite, goethite and gibbsite) and kaolinite, whereas the Chromosol is dominated by illite followed by kaolinite.

Plant biomass

Carbon input

After completion of 24 h of continuous soil respiration measurements all four plants were destructively sampled for incubation of roots (for δroot measurement) and above ground biomass. Plants were cut at the soil surface and above-ground biomass was oven-dried at 60 °C and weighed. Biomass of roots was not quantified due to difficulties in ensuring sampling of all roots in all soils, i.e. roots were extremely difficult to remove from the Chromosol. Note that strong effects of treatment temperature on above-ground biomass are not expected because all plants were grown under the same conditions for 4 weeks before being exposed to the treatment temperatures for 2–6 days only. After incubation for δroot measurement, the sampled roots were also oven-dried at 60 °C. Dried root and leaf samples were ground to a fine powder and analysed for δ13C using a stable isotope mass spectrometer (Delta V, Thermo Finnigan).

Photosynthetic rate generally increased with temperature, as expected (Fig. 1). Plants grown in the Chromosol had the highest photosynthetic rate; however the rate was significantly lower at 36 °C compared to 30 °C. Plants grown in the Ferrosol had a relatively low photosynthetic rate at all temperatures, although the

8

bc

A

-2

-1

mol m s )

bc

ab

6

ab

bc

4

abc

ab ab

Chromosol Ferrosol

2 2.0 0

B

e

1.5

de

-2

-1

Rleaf ( mol m s )

Significant differences between the two soils and between temperatures were determined using 2-way ANOVA in Genstat (15th Edition, VSN International, UK). The temperature response function (Eq. 3) was fitted using non-linear fitting in Origin (version 6, Microcal Software, Northampton, MA).

c

A

a

bcd

1.0

abc

cde

bcd 0.5

ab

abcd

abc

a

0.0

C

c

c

12

Above-ground biomass (g)

Statistical analysis

10

c

bc ab

8

abc abc

4

ab

a

ab

Results 0

Soil physical and chemical properties

10

15

20

25

30

35

40

o

Soil temperature ( C)

Table 2 shows the chemical and physical analysis of the two soils. Soil pH is close to neutral for the Chromosol and the cation exchange capacity of this soil (18.3 cmol(+) kg−1) is much greater than the Ferrosol. The

Fig. 1 Photosynthetic rate (a), leaf respiration rate (b), and aboveground biomass (c) for corn plants grown at 30 °C and exposed to a range in temperatures for 2–5 days prior to measurement and harvest. Values are means ± standard errors, n=4, and letters indicate significant differences between means (P<0.05)

Plant Soil

photosynthetic rate was significantly higher at 30 and 36 °C than at 12 °C. As expected, leaf respiration rate (Rleaf) generally increased with temperature, and was generally higher in the plants grown in the Chromosol compared to the Ferrosol (Fig. 1). The balance between photosynthesis and respiration strongly influences the amount of carbon available for growth. We expect differences between growth rates (as measured by above ground biomass) of plants growing in different soils, but only limited effects of measurement temperature because of the short length of time pots were exposed to the treatment temperatures. Significant differences in above-ground biomass were found for plants growing in the two soils (P<0.001). Figure 1c shows that plants grown in the Chromosol had higher biomass production, suggesting higher carbon input. Above-ground biomass was significantly decreased at the lowest temperature (12 °C, P<0.001) for plants grown in the Chromosol. However, biomass production was not affected by low temperature in the Ferrosol. This resulted in a significant treatment temperature by soil type interaction (P=0.020). To see whether measuring the δ13C of plant organic matter can provide an adequate end-member for partitioning below-ground respiration, the 13C of root and shoot organic matter was compared with δ13C of

root-respired CO2 (Fig. 2). The δ13C of root organic matter was not systematically related to δ13C of CO2 respired by the same roots in either soil (P>0.05). δ13C of shoot organic matter was also unrelated to the δ13C of root-respired CO2 for the Ferrosol, but the two were positively and significantly (P<0.001) related in the Chromosol. The slope of the relationship was greater than 1, and the intercept much higher than zero. δ13C of shoot organic matter explained only 59 % of the variability in δ13C of root-respired CO2 for the Chromosol, and had no explanatory power for the Ferrosol (Fig. 2).

Temperature sensitivity of respiration As expected, the CO2 efflux from both soil only and soil with roots increased with increasing temperature (Fig. 3). The respiration rate of the Ferrosol was the greater of the two soils at all temperatures, and had higher temperature sensitivity than the Chromosol, E0 value of 108.0 vs 71.6 for the Chromosol. The respiration rate of the soils with roots was always higher than the same soil without roots (Fig. 3). However, the temperature sensitivity of the respiration rate in the presence of roots was statistically significant for the Ferrosol only; the Chromosol had non-significant differences in the E0 with and without roots.

A Root organic matter

B Shoot organic matter

-10

-12

root

(‰)

-14

-16

-18

Chromosol Ferrosol o 12 C o 18 C o 24 C o 30 C o 36/30 C

-20 -20

-18

-16

-14

-12

13

C root OM (‰)

Fig. 2 The relationships between δ C of root-respired CO2 and δ13C of root (a) and leaf (b) organic matter for corn plants grown at 30 °C and exposed to a range in temperatures for 2–5 days prior to 13

-10

Chromosol = 4.05 + 1.28 root

13

C shoot OM

2

r = 0.59, P < 0.001

-20

-18

-16

-14

-12

-10

13

C shoot OM (‰)

measurement and harvest. In b the line is a least squares regression for the relationship among the samples from the Chromosol with slope, intercept and closeness of fit as indicated

Plant Soil

C

E

Ferrosol 6

0.8 0.6

R

0.4 0.2 0.0

0 Soil and roots 2 R10 = 4.01, E0 = 71.6, r = 0.70 Soil only 2 R10 = 2.09, E0 = 71.3, r = 0.86

F

0.8

2 0.2

10

15

20

25

Chromosol

ROM

2

R10 = 2.27, E0 = 100.2, r = 0.63 Roots only SOM only

-1

0.6

Soil and roots Soil only

0.4

0

2

6

-2

4

4

Chromosol

fOM

-2

-1

R ( mol m s )

6

2

R10 = 3.40, E0 = 135.8, r = 0.81

0

D

Chromosol

R ( mol m s )

B

Ferrosol

ROM

-1 -2

-2

Soil and roots 2 R10 = 7.97, E0 = 108.0, r = 0.91 Soil only 2 R10 = 4.30, E0 = 98.4, r = 0.93

fOM

2

-1

mol m s )

4

R

6

1.0 Ferrosol

mol m s )

A

30

35 o

Soil temperature ( C)

40

0.0

10

15

20

25

30

35 o

Soil temperature ( C)

40

4

2

0

10

15

20

25

30

35

40

o

Soil temperature ( C)

Fig. 3 Respiration rates and their temperature sensitivities in pots with soil only and soil with roots (a and b), the proportion of organic matter decomposition (fom; c and d), and the respiration rates of partitioned total soil respiration rate (Rom and Rroots; e and f) for the two soils. The curves in a, b, e and f represent fitted

Lloyd and Taylor (1994) respiration responses using Eq. 3, for soil only and soil with roots (a and b), and for organic matter decomposition (e and f) with values for fitted parameters and closeness of fit indicated

Carbon isotope composition of soil- and root-respired CO2

not significant. δsoil+roots was generally more depleted for the Ferrosol over all temperatures, while the Chromosol had more enriched respiration over all temperatures. For the Chromosol, δsoil+roots increased slightly with increasing temperature, while there was no clear trend in δsoil+roots with temperature for the Ferrosol. Significant effects of soil type and temperature were found for δ13C of CO2 respired by roots (P<0.001, Table 3), but there was no significant interactive effect. Overall δroots became more depleted as temperature increased for the plants grown in the Chromosol, while for the Ferrosol, δroots was more depleted for 36 °C compared to cooler temperatures. The CO2 respired by roots from plants grown in the Ferrosol was generally more enriched than those from the Chromosol. So overall, δsoil and δroots were more similar at higher temperatures than at lower temperatures. Significant differences between δroots with temperature are likely to be due to changes in the δ13C of recently-fixed carbohydrates. These differences may reflect differences in δ13C of source air in the growth room (which was not measured), or differences in

Average δ13C of CO2 respired from pots with soil only (δsoil) and soil with roots (δsoil+roots) was calculated between 5 pm and midnight for each pot. These times were chosen because both the soil temperatures and the isotope compositions were stable for all pots. δsoil varied significantly between the two soils (P<0.001; Table 3); δsoil was depleted for the Ferrosol (−24.5‰) over all temperatures, while δsoil from the Chromosol was more enriched (−20.6‰). The interactive effect of soil and temperature was significant for δ soil (P<0.001), indicating that δsoil responded to temperature in different ways for the two soils. For the Chromosol, δsoil was significantly (P<0.001) more enriched at 30 and 36 °C than at lower temperatures, while δ s o i l was significantly (P<0.001) more enriched at 12 °C than at higher temperatures for the Ferrosol. Similar effects were found for δ13C of CO2 respired from pots with soil and roots, although differences were

Plant Soil Table 3 Average soil carbon flux (μmol m−2 s−1) from pots with soil only (Rsoil) and soil with roots (Rsoil+roots), and average stable carbon isotope composition (‰) of the respiratory efflux of soil difference (P<0.05) Soil

Temp

Rsoil

Ferrosol

12

0.5±0.0ab

Rsoil+roots

δsoil

δsoil+roots

δroots

0.4±0.0a

−22.9±0.3b

−22.3±2.2

−11.0±0.5f

a

18

0.8±0.1

1.5±0.1

−25.3±0.2

−16.2±0.7

−12.9±0.2e

24

1.4±0.0cd

1.9±0.2ab

30 36 Chromosol

only (δsoil), soil with roots (δsoil+roots) and roots only (δroots) for two soils at a range of temperatures. Values are averages between 5 pm and midnight ± standard errors, n=4, letters indicate significant

12

abc

d

1.9±0.1

e

3.2±0.2

a

0.3±0.0

ab

−24.6±0.1a

−21.2±0.6

−13.9±0.4cde

bc

−25.2±0.1

−19.1±0.6

−12.7±0.3e

c

−24.4±0.2

−21.1±0.6

−15.0±0.2bcd

a

−23.2±0.7

−16.6±0.1

−13.3±0.3de

ab

de

3.6±0.3 2.7±0.9 0.4±0.1

a a bc

18

0.5±0.1

1.4±0.2

−20.4±0.8

−15.9±0.5

−14.2±0.1bc

24

0.8±0.1 abc

1.7±0.1ab

30 36

ab

bc

1.1±0.1

abc

0.8±0.3

−21.4±0.4cd

−17.2±0.3

−15.5±0.1bc

b

−19.2±0.1

−18.1±0.3

−15.9±0.1b

ab

−19.7±0.2

−18.0±0.7

−17.6±0.2a

2.8±0.5 2.0±0.9

e e

Soil effect

<0.001

0.015

<0.001

NS

<0.001

Temp. effect

<0.001

<0.001

0.010

NS

<0.001

Interaction

<0.001

0.003

<0.001

NS

NS

photosynthetic discrimination (Farquhar 1983). Specifically, changes in leakiness of the leaf bundle sheath cells to CO2 during photosynthesis can occur as temperature changes (Kubasek et al. 2007), and could have resulted in differences in δ13C of recently-fixed carbohydrates which were used as respiratory substrates in the roots. Partitioning total soil respiration into root and rhizosphere and SOM components Respiration from pots with roots was partitioned using δ13C values and Eq. 1. Figure 3c shows that the Ferrosol showed no consistent trends in variation in partitioning between root and soil microbial respiration with increasing temperature. However in the Chromosol, fom increased with temperature (Fig. 3d), indicating an increase in the proportional contribution from the decomposition of organic matter with increasing temperature. Both Rom and Rroots (i.e. rates of organic matter decomposition, Rom, and root and rhizosphere respiration, Rroots, calculated for pots with plants) generally increased with increasing temperature as expected (Fig. 3e and f). General increases in Rroots with temperature are consistent with general increases in leaf photosynthesis and respiration with increasing temperature. However, the increase

in fom at higher temperatures for the Chromosol resulted in lower Rroots at 30 and 36 °C than the lower temperatures. The temperature sensitivity of Rom was higher than the temperature sensitivity of Rsoil for both the Ferrosol and the Chromosol. That is, E0 for Rom is higher than E0 for Rsoil; 135.8 compared to 98.4 μmol m−2 s−1 for the Ferrosol (comparing Fig. 3a with Fig. 3e), and 100.2 compared to 71.3 μmol m−2 s−1, for the Chromosol (comparing Fig. 3b to Fig. 3f). Priming effect Figure 4 shows evidence of a positive priming effect in the Chromosol as Rom (partitioned from total soil respiration using Eq. 1) was higher than Rsoil at higher temperatures. However, the Ferrosol showed a negative priming effect at 36 °C, falling below the 1:1 line. That is, more ‘older’ C 3 carbon was progressively decomposed as temperature increased when plants were present for the Chromosol, and the reverse was observed for the Ferrosol. Diel variability There was no significant variability in δ13C of soil-respired CO2 in pots with soil only. This is the case even for pots measured at 36 °C in the

Plant Soil

3

Temperature sensitivity of Rsoil and Rsoil+roots Negative priming effect

-2

-1

Rom ( mol m s )

Discussion

Chromosol Ferrosol

2

o

30 C

Positive priming effect

o

o

36 C

30 C

o

36 C

1 o

18 C

o

12 C

o

o

o

12 C

0 0

o

24 C

24 C

18 C

1

2 -2

3 -1

Rsoil ( mol m s ) Fig. 4 Comparison of respiration rates from pots without plants (Rsoil) with the soil organic matter decomposition rates calculated from the partitioned flux from pots with plants (Rom) for the two soils reveals an increasingly positive priming effect with increasing temperature for the Chromosol, and negative priming effect at the highest temperature for the Ferrosol

light and 30 °C in the dark; the change in temperature did not result in significant change in δsoil. There were small but significant increases in fom for the Ferrosol at 24 °C (from 0.14 to 0.23 before and after midnight), and for the Chromosol at both 30 °C continuously (from 0.66 to 0.76 before and after midnight) and 36/30 °C day/ night temperature (from 0.53 to 0.81 before and after midnight). Further, there was a significant reduction in soil-only respiration rate after the lights were turned off for both soils at 36/30 °C and 30 °C, and also for the Ferrosol at 24 and 18 °C. This response was most pronounced for the Ferrosol at 36/30 °C. The rate of root and rhizosphere respiration for the Chromosol (calculated from the partitioned flux) showed little diel variability at 12 and 18 °C, but tended to be lower after the lights were turned off at the higher temperatures, particularly at 36/30 °C. For the Ferrosol, root and rhizosphere respiration increased from midafternoon until the lights turned off, then decreased significantly for the pots at 36/30 °C, while at constant 30 °C Rroots decreased from mid-afternoon until the lights turned off then remained stable (Fig. 5). There was little diel variability in Rroot at any temperature lower than 30 °C for the Ferrosol.

As expected, both Rsoil and Rsoil+roots increased exponentially with temperature, so that the widely-used Lloyd and Taylor function (Lloyd and Taylor 1994) fits well for both Rsoil and Rsoil+roots for both soil types (r2 of 0.70–0.98). This is consistent with many studies (Kirschbaum 1995) and confirms the suggestion that any substantial temperature increase may lead to a positive feedback phenomenon and result in an intensification of global warming (Cox et al. 2000). The Ferrosol had higher respiration rates (R10) and higher sensitivity to temperature (E0) both for soils alone and soils with roots, as expected given the generally positive relationship between organic carbon content and soil respiration rate (Schlesinger and Andrews 2000). Soil clay mineralogy and temperature response of fom The Ferrosol consists primarily of variable charge minerals including goethite, gibbsite, hematite and kaolinite. SOM binds strongly to the iron and Al oxides in the Ferrosol via the highly reactive singly coordinated OH groups on the edges of Fe and Al oxides (Six et al. 2004; Lutzow et al. 2006). Similar reactions involving the acid Lewis sites on the edges of kaolinite may occur as well. However, Fe and Al oxides have much greater specific surface area (Eusterhues et al. 2005), which makes their contribution more significant for SOM interactions. They also have sites of increased reactivity known as micropores (<2 nm) or small mesopores (2–10 nm) (Kaiser and Guggenberger 2003), which are best suited for sorptive interactions with organic acids (Mikutta et al. 2004). This is where organic matter can form strong multiple bindings that limit the desorbability and the oxidative removal of SOM (Kaiser and Guggenberger 2007; Wattel-Koekkoek et al. 2003). In contrast, the Chromosol is dominated by illite where SOM is adsorbed via cation bonding to the siloxane surfaces with permanent layer charge (Sposito et al. 1999). While the minerals that make up the Chromosol form relatively strong bonds with SOM, the minerals that make up the Ferrosol lead to an even stronger binding with SOM (Jackson 1963). Illite has been found to release SOM almost completely (Kahle et al. 2004). The stronger bonds between clay minerals in the Ferrosol may have been partly responsible for the

Plant Soil A Ferrosol Rsoil

C Ferrosol Rom

E Ferrosol Rroots

B Chromosol Rsoil

D Chromosol ROM

F Chromosol Rroots

3

-2

-1

R ( mol m s )

4

2

1

3

-2

-1

R ( mol m s )

4 0

2

Temperature o 36/30 C o 30 C o 24 C o 18 C o 12 C

1

0 18:00

24:00

06:00

Time (hours)

18:00

24:00

06:00

Time (hours)

18:00

24:00

06:00

Time (hours)

Fig. 5 Diel variation in the rate of respiration from the pots with soil only (a and b), soil organic matter decomposition (c and d) and root respiration (e and f) for the two soils. Shaded sections

indicate the dark period. Values in c, d, e and f are calculated from the partitioned flux from pots with plants using Eq. 1. Values are means, n=4

higher organic carbon content in this soil compared to the Chromosol. Kaiser and Guggenberger (2007) suggested that SOM is bound to goethite by strong multiple bonds involving carboxyl groups and there may be a partial penetration of SOM into mineral structure due to ligand exchange process. Similar observations were also made by Jones and Singh (2014) in a recent study where a positive correlation was observed between Fe concentration and the proportion of carboxylic functional groups in various density fractions of different soil types. The authors also found a greater proportion of carboxylic functional groups in Fe oxides dominated density fractions in the Ferrosol (the same soil as used in this study); chemical bonding of carboxylic groups on to Fe oxides surfaces thus resulted in a greater stabilization of organic matter in this soil. With increased temperature, the δ13C measurements suggested that an increasing proportion of Rsoil+roots came from ‘older’ soil carbon as opposed to ‘new’ plant carbon for the Chromosol, i.e. fom increased with temperature (Fig. 4), indicating an increase in the decomposition of previously stored soil carbon (Fig. 5). This response is likely due to primary production providing the organic fuel that drove the soil metabolic activity

(Raich 1992). Further, the relatively weaker bonding between organic matter and siloxane surfaces with permanent layer charge in the Chromosol may have resulted in the release of organic matter. In contrast, the proportional contribution from organic matter decomposition remained relatively constant with increasing temperature in the Ferrosol. This is consistent with the strong binding of organic matter involving ligand exchange reaction and singly coordinated OH groups on the edge of Fe/Al oxides (KögelKnabner et al. 2008; Sollins et al. 1996; Mikutta et al. 2006; Kaiser and Guggenberger 2007), and likely produced the negative priming effect at 36 °C. Diel variability in respiration rate and δ13C No diel variability in δ13C of respired CO2 was observed for pots with soil only for either of the two soils, at any temperature, even when temperature varied from 36 to 30 °C over 24 h and soil respiration rate varied considerably. We assume no temporal variability in δroots, so that any diel variability in δsoil+roots will be due to changes in the relative contributions of Rom and Rroots to Rsoil+roots. δsoil+roots varied by at most 1‰ (at 36/

Plant Soil

30 °C, comparing an average of the 5 h before the lights were turned off to an average of 5 h after the lights were off) for the Chromosol, but showed little consistent variability in the Ferrosol at any temperature. Significant variability in δsoil+roots has been reported in some (Marron et al. 2009; Bahn et al. 2009; Moyes et al. 2010; Kodama et al. 2009; Grossiord et al. 2012), but not all (Betson et al. 2007; Millard et al. 2010; Barbour et al. 2011), studies in the field. Further, Moyes et al. (2010) report seasonal variation in the degree of diel variability in δsoil+roots, with variability of up to 5‰ in conditions of low soil CO2 fluxes and high diel fluctuations in soil temperature at the beginning of the growing season, and no diel variation in δ13C later in the growing season for the same ecosystem when soil CO2 effluxes were higher. Interestingly, high levels of diel variability were observed by Moyes et al. (2010) at low fluxes both with roots and in trenched plots without roots. Bathellier et al. (2009) found no significant variability in δ13C of root-respired CO2, in contrast to the strong variability found in δ13C of leaf-respired CO2 (Tcherkez et al. 2003). Diel variability in δsoil+roots, when it has been observed, has been interpreted as diel changes in the δ13C of carbohydrates supplied to the roots (Bahn et al. 2009), or variability in the proportional contribution of root and rhizosphere respiration compared to SOM decomposition (Marron et al. 2009), or both depending on the degree of water stress (Barthel et al. 2011), but also to the effects of non-steady state gas transport on diffusive fractionation (Moyes et al. 2010). By holding temperature constant, waiting for at least 2 h after the soil chamber was placed on the soil surface before making δ13C measurements, and keeping the chamber in place for a full 24 h, we greatly reduced the opportunity for non-steady state gas transport effects. Consistent with this, we do not find a negative curvilinear relationship between soil respiration rate and the diel range in δ13C of respiration, as previously reported by Moyes et al. (2010). We found no diel variability in δ13C of soil-only respiration and only very small diel variability in δ13C for soil with roots at higher temperatures. This suggests that proportional changes in the contribution of Rroots to Rsoil+roots were either nonexistent or very small for these soils. Our results support the suggestion by Moyes et al. (2010) that diel variability may be largely due to non-steady state effects, and that great care must be taken in interpreting field observations of diel variability in δsoil+roots.

Advantages and limitations of the δ13C natural abundance method The natural abundance stable isotope partitioning method is limited by the accuracy of carbon isotope measurements, and requires a large difference between δsoil and δroots. The difference between δsoil and δroots depends on previous overlaying vegetation. C3 vegetation growing on C4 soil (or vice versa) indicates a landuse change or significant ecological change. So the application of this method is restricted to locations where these changes have occurred (or can be experimentally created), or to controlled environment studies as described here. High-resolution measurements of δ13C are required due to a maximal range of only 14‰ available for variations of δ13C (Cheng 1996). Here the differences between δsoil and δroots varied between 9.9‰ (for the Chromosol at 12 °C) and 2.1‰ (for the Chromosol at 36 °C) making partitioning successful in all cases. An important limitation is that this technique cannot effectively partition microbial respiration in the rhizosphere from actual root respiration. However, a number of previous studies have suggested that microbial respiration in the rhizosphere is so strongly coupled to root activity that, in effect, the root respiration (using ‘new’ carbon) extends beyond the direct root zone (Heilmeier et al. 1997; Uchida et al 2010). The main advantage of this method is that it requires minimal system disturbance during measurement, that is, the fluxes are measured in situ. If an ecosystem has had a natural or experimentally induced change in photosynthetic type resulting in large enough difference in δ13C, there is no need for artificial labelling of plant carbon. However, labelling studies have been shown to increase precision of partitioning in the short term (e.g. Sapronov and Kuzyakov 2007). Using a null-balance chamber means that separation of the sample δ13C signal from the atmospheric δ13C signal is also not required (as stated by Midwood et al. 2008), reducing the error in the partitioning estimate. Non-steady state gas transport effects on δsoil+roots (Moyes et al. 2010), which would introduce considerable errors in estimating Rom and Rroots, could also be minimised using this null-balance chamber approach, if the chamber were to remain in place for several hours. Keeling plot approaches using closed chambers have been shown to introduce diffusional artefacts (Nickerson and Risk 2009), which may have been interpreted as temporal variability in previous field studies.

Plant Soil

Another important point regarding partitioning using natural abundance stable isotopes is implied in Fig. 2. If we had used δ13C of root organic matter (average δ13C of −17.5‰ for both Ferrosol and Chromosol) in Eq. 1 to partition soil and root respiration, fom would have been underestimated by between 0.23 and 0.44. In fact, the partitioning would not have been possible (i.e. negative values for fom) if δ13C of root organic matter had been used for the Ferrosol at 24 °C and the Chromosol at 12, 18 and 24 °C. This effect is somewhat exaggerated in this experiment, due to the wide range in temperatures compared to the plant growth temperature, which would have meant recently-fixed carbohydrates were very different isotopically to stored carbohydrates. Nevertheless, this demonstrates that it is important to accurately determine the end members, and avoid using fractionation values simply taken from the literature (Werth and Kuzyakov 2010).

Implications of the temperature response of Rom and fom to climate change Losing an increasing proportion of older carbon from the soil (i.e. fom increases) as temperature increases leads to concerns that CO2-driven global warming may accelerate (Cox et al. 2000; Zimmermann et al. 2009). Radiocarbon measurements indicate that the majority of organic matter is found in labile fractions that will easily decompose should the climate experience warming (Chapman and Thurlow 1998: Lindroth et al. 1998). However, this concern is not supported by all studies to date. For example, Luo et al. (2001) found a decrease in the temperature sensitivity of soil respiration with warming. This may have been the result of a reduction in plant production leading to less root respiration due to substrate limitation (Rustad and Fernandez 1998), but may also be due to enhanced soil drying at higher temperatures which reduced root and microbial activity. A recent study by Uchida et al. (2010) investigated the temperature sensitivity of soil respiration components on two pastoral soils with high and low fertility. In the high fertility soil both Rom and fom increased significantly with temperature, while in the low fertility soil Rom remained constant and fom decreased with increasing temperature. The latter effect was suggested to be the result of microbial utilization of root-derived carbon and limited availability of soil carbon. The presence of plant roots increases the total flux, but also

influences the temperature sensitivity Rom (Zhu and Cheng 2011; Uchida et al. 2010). Rom increased with increasing temperature for both soils studied here, although the proportion contribution of Rom to Rsoil+roots remained constant with increasing temperature for the Ferrosol. Both soils studied here had moderate to high fertility, and both had relatively high temperature sensitivity of Rom (E0 =100 and 136, for the Chromosol and the Ferrosol, respectively), compared to the temperature sensitivity of respiration from pots with soil only (E0 =71 and 98 for the Chromosol and the Ferrosol, respectively). Despite the relatively stable values for fom with increasing temperature, and the observed negative priming effect for the Ferrosol (likely due to strong bonds between Fe/Al oxides and organic carbon), concerns regarding the acceleration of the carbon cycle and reinforcement of greenhouse gas-related global warming (as in Cox et al. 2000) are relevant for both the Ferrosol and the Chromosol, because the Ferrosol has a high carbon content and overall high respiration rates which increase with increasing temperature. Acknowledgments M.M.B. was supported by an Australian Research Council Future Fellowship (FT0992063). We thank T Winters, S Ryazanova and S Bachmann for technical support, and two anonymous reviewers for comments that greatly improved this paper.

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