Plant effects on soil carbon storage and turnover in a montane beech (Nothofagus) forest and adjacent tussock grassland in New Zealand.
The role of the terrestrial biosphere in the global carbon (C) cycle is still poorly understood because of the complex biology underlying C storage, the spatial variability of vegetation and soils, and the effects of land use. Among reasons for this uncertainty are the differing rates of deforestation and tree regrowth in the tropics and mid-latitudes, the contribution of [CO.sub.2] fertilisation, and the global redistribution of ecosystems in response to anthropogenic changes in atmospheric [CO.sub.2] concentrations, land use, and climate (Melillo et al. 1995). Before C budgets can be extrapolated from patch to regional scales, the various factors that control C allocation above- and below-ground, including ecosystem type, must be identified.
Soils can act as both a source and a sink for atmospheric [CO.sub.2], depending on the balance between litter inputs, decomposition rates, and soil organic matter formation. Globally, the storage of soil organic C is apparently more closely related to factors controlling decomposition of detrital inputs than to net primary production (NPP) (Cebrian and Duarte 1995). However, grasslands generally accumulate more organic matter to greater soil depths than forests because of a more extensive network of fine roots (Sanchez et al. 1982; Tate 1987). This concept, although widely accepted, has not to our knowledge been tested in comparative studies of undisturbed forest and grassland ecosystems in the same climo-edaphic environment, and where the soils are developed from the same parent material.
We discuss here, and also in Ross et al. (1996), the storage of soil C and its exchange between the soil surface and the atmosphere in two adjacent, unmanaged, indigenous ecosystems near the treeline, namely, an old-growth beech (Nothofagus) forest and an ungrazed tussock (Chionochloa) grassland. Our aim was to use experimental observations, supported by modelling, to define the factors that influence the patterns of C accumulation in these ecosystems.
Materials and methods
The two study sites were located near the treeline on the western slopes of the Camp Stream catchment in Craigieburn Forest Park; the Craigieburn Range lies between the Canterbury Plains and the Southern Alps, South Island, New Zealand. Neither site had been disturbed by recent avalanches or erosion, in contrast to most of the land area in the Craigieburn Range (R. B. Allen, pers. comm.).
The forest site (43 [degrees] 7'S, 171 degrees] 42'E) comprised a closed canopy of mature mountain beech [Nothofagus solandri var. cliffortioides (Hook. f.) Poole] with no understorey vegetation, and was on a slope of about 20-25 [degrees] about 100 m below treeline at 1240 m. The stand had similar characteristics to an adjacent mature stand (age [is greater than] 150 years) where tree stem density was 1220 per ha, mean diameter 25 cm at 1.4 m height, basal area 68 [m.sup.2]/ha, and stem biomass 245 Mg/ha (Allen et al. 1997). The pit and mound microtopography of the site reflected the history of tree overturn (Burns et al. 1984). Carpeting the forest floor was a thin continuous cover (about 30 mm) of intact and decomposing fine litter, and much decaying woody debris ([is greater than] 5 mm diam.).
The ungrazed grassland site (43 [degrees] 08'S, 171 [degrees] 42'E was about 100 m distance above treeline at 1315 m elevation on the nose slope (about 20 [degrees]) of a stable ridge protected from avalanches, and with little evidence of fire effects, erosion, of previous forest cover. The dominant vegetation was Chionochloa pallens Zotov, with minor components of Poa colensoi Hook. f. and Celmisia lyallii Hook. f.; about 50% of the area was unvegetated inter-tussock spaces, with a thin continuous cover of fresh and decomposing litter (about 3 cm depth). Much of the tussock litter was erect in the tussock crowns.
Because of the remoteness of the sites, we were unable to collect climate data during this study. However, based on the 7-year data of Noble (1974), and estimates of precipitation during winter and spring (July-September) by using a topoclimate model (Parshotam et al. 1995), annual precipitation was calculated to be about 1700 mm with a more or less even distribution throughout the year. Estimated annual evaporation, using the topoclimate model (Parshotam et al. 1995), was about 970 mm, with a water deficit suggested for January and February only. Net precipitation reaching the soil surface was estimated as the difference between precipitation above the plant canopies and wet canopy evaporation of intercepted precipitation; values of 39% (Rowe 1975) and 21% (Campbell and Murray 1990) for intercepted precipitation were used for the forest and grassland, respectively. Net annual precipitation is then about 1040 mm for the forest and 1340 mm for the grassland.
The mean annual air temperature at the sites is about 6 [degrees] C, with monthly ranges from 0 [degrees] C in July to 12 [degrees] C in February (Noble 1974).
Soils and sampling
The soils were formed from mixed greywacke loess and colluvium over shattered (in situ) rock, and were fine sandy, mixed, frigid, Andic Dystrochrepts. The forest soil was Bealey silt loam and the grassland soil Puketeraki silt loam, with a more strongly developed profile.
Soils were described according to Clayden and Hewitt (1989). Profile horizons and depths (cm) to rock are given in Table 1; the forest soil did, however, have an additional horizon (Cs at 88-95 cm) which was not sampled for property measurements. Roots were most abundant in the uppermost horizons and common throughout both profiles in all but the deepest horizons.
Table 1. Some litter and mineral soil properties by horizon to the bases of the profiles
Depth sample Total C Soil (cm) Horizon pH (g/kg) Bealey silt loam 0-0.5 L 4.1 465 0.5-3 F/H 3.9 361 3-8 Ah 4.1 138 8-19 E/B 4.4 64 19-34 Bs 4.6 33 34-44 Bw 5.1 22 44-54 BC1 5.1 17 54-65 BC2 5.3 10 65-88 C 5.4 8 Puketeraki silt loam 0-3 H 4.7 180 3-13 Ah1 4.6 103 13-20 Ah2 4.8 64 20-24 E/B 4.8 44 24-34 Bs1 5.0 36 34-46 Bw 5.1 26 46-54 BC 5.1 22 54-66 Bs2 5.5 14 Base Total N saturation(A) CEC(A) Alpyr(B) Soil (g/kg) (%) (cmol/kg) (mg/kg) Bealey Forest litter silt loam 6.0 45 66 200 8.5 15 74 1700 Mineral soil 3.9 7 36 3700 2.4 3 31 4900 1.4 2 25 6500 1.0 1 15 6600 0.9 1 12 5800 0.6 3 6 3500 0.5 4 5 2100 Puketeraki Grassland litter silt loam 7.0 15 34 3100 Mineral soil 5.3 29 5500 3.7 24 7100 2.5 21 7600 2.0 19 8000 1.5 14 8100 1.3 12 6500 0.1 10 3700 Bulk density Clay [Delta][sup.14]C Soil (t/[m.sup.3]) (g/kg) ([salinity] [+ or -] s.d.) Bealey silt loam 0.04 0 216 [+ or -] 9 0.12 0 n.d. 0.39 250 147 [+ or -] 9 0.45 360 19 [+ or -] 9 0.62 380 -36 [+ or -] 8 0.85 320 n.d. 1.05 270 -247 [+ or -] 7 1.12 130 n.d. 1.18 120 -445 [+ or -] 5 Puketeraki silt loam 0.04 0 154 [+ or -] 9 0.50 230 48 [+ or -] 8 0.61 280 -59 [+ or -] 7 0.65 330 n.d. 0.75 320 -167 [+ or -] 7 0.81 280 -387 [+ or -] 5 1.08 130 n.d. 0.72 60 -440 [+ or -] 5
n.d., not determined.
(A) Base saturation and CEC (cation exchange capacity) determined at pH 7 using ammonium acetate (Blakemore et al. 1987).
(B) Alpyr, 0.1 M [Na.sub.4][P.sub.2][O.sub.7] - extractable Al (=Al bound to humus).
Samples were collected from both sites in March 1993 by 2 procedures. In the first procedure, litter and mineral soil were sampled by horizon in freshly dug soil pits (about 1 by 1.3 m; 1 per site) down to the base of the soil profiles. The pit locations were selected by an experienced pedologist to be morphologically representative of the soil within a 100-[m.sup.2] area at each site. Bulk samples for chemical analyses were collected immediately adjacent to cores that were taken for physical analyses.
In the second procedure, litter (L and FH material) and mineral soil (2.5 cm diam. cores; 0-10 and 10-20 cm depth) samples were taken in the forest within this same 100-[m.sup.2] area as described by Ross et al. (1996). In the grassland, litter and mineral soil were sampled similarly from under the tussock canopy and between tussock plants; grab samples of C. pallens standing dead material were also collected (Ross et al. 1996). Five replicate pooled samples of litter and mineral soil from each depth were thereby obtained, with each replicate sample consisting of 10 collections of litter or mineral soil. Mineral soil was also collected at 20-50 cm depth with a knife and trowel from the pits used in the first sampling procedure to provide 3 replicate samples, each being obtained from a different face.
Litter and soil analyses
The litter and mineral soil samples from the pits were used for soil chemical, physical, and radiocarbon analyses; total carbon (C) and nitrogen (N) were also determined in the multiple-core samples (Ross et al. 1996). Samples were air-dried and sieved ([is less than] 2 mm; [is less than] 0.25 mm for total C and N determinations) before chemical analyses.
Determinations of pH ([H.sub.2]O), total C and N, CEC, % base saturation, exchangeable (ammonium acetate) Ca, Mg, K, Na, and Al, and pyrophosphate-extractable Al were made according to Blakemore et al. (1987). After removal of fine twigs and roots from the samples of litter and sieved mineral soil by hand-picking, subsamples were analysed for their [sup.14]C contents by Accelerator Mass Spectrometer dating as previously described (Tate et al. 1993). Total soil C and microbial C, measured on litter and mineral soil samples collected to 50 cm depth by the second procedure (Ross et al. 1996), were used to estimate total annual input of plant residues by a modelling approach (Jenkinson and Coleman 1994).
Dry bulk density of the litter layers was measured on 20 replicate samples of known volume, collected within a radius of about 20 m of the soil pits; woody debris larger than about 5 mm diam. was excluded. One of two techniques was used to sample soil for dry bulk density determinations of the mineral soil horizons. Where possible, duplicate undisturbed core samples were collected from each mineral soil horizon (Gradwell 1972). An excavation/sand replacement method (Blake and Hartge 1986) was used for the 2 upper-most mineral soil horizons at the forest site, where abundant roots and the light texture of the soil prevented undisturbed core samples being taken. The excavation/sand replacement method was also used for the bottom-most soil horizons at both sites where abundant stones again prevented undisturbed core samples being taken.
Particle size distributions (Claydon 1989), pore size, and moisture release characteristics (Gradwell 1972) were also determined in duplicate for each horizon. Water-filled pore space (WFPS) (%) was estimated for each soil horizon as volume of soil water x 100/total pore volume for the different combinations of bulk density and water potential (Linn and Doran 1984).
In situ respiration measurements
An in situ chamber method using alkali trapping of respired [CO.sub.2] was employed for forest floor and intertussock grassland measurements on 9-10 March 1993. Details of the respirometers, and their method of use, are given elsewhere (Tate et al. 1993). Twelve respirometers were arrayed at the forest site, and 10 at the grassland site. They were positioned about 4 m apart and within 5-10 m of the soil pits.
Ten further in situ soil respiration measurements were made in each ecosystem on 16 March 1994, using a portable infrared gas analyser system (PP Systems, UK) (Jensen et al. 1996). The remoteness of the sites, and their inaccessibility for up to 4 months because of snow cover, prevented more frequent measurements being taken. Annual soil [CO.sub.2]-C effluxes were, therefore, estimated using mean monthly air temperatures and a soil respiration-temperature relationship (Tate et al. 1993).
Estimation of total soil C inputs, ecosystem NPP, and inert organic matter
A modelling approach based on the Rothamsted soil C turnover model (ROTHC) (Jenkinson 1990) was used to estimate annual inputs of plant residue C to 23 cm depth from the mineral soil surface (Jenkinson and Coleman 1994), and also ecosystem NPP (Jenkinson et al. 1992; Tate et al. 1993). This was achieved by assuming steady-state conditions (i.e. detrital C input = soil [CO.sub.2]-C production) under the climate and soil conditions prevailing at both sites. The annual inputs of plant residue C and size of the inert organic matter (IOM) pool were estimated by varying them iteratively until soil C and [sup.14]C contents matched the measured values. The replicate core samples (to 20 cm depth of mineral soil) and replicate pit-face samples (20-50 cm depth of mineral soil) were used for the C measurements to take account of spatial variability. Predicted and measured microbial C values (Ross et al. 1996) were then compared to assess model performance.
Inputs were assumed to be evenly distributed throughout the year in monthly time steps (Coleman and Jenkinson 1996). Further sources of uncertainty were the values selected for the ratios of readily decomposable plant material (DPM) to resistant plant material (RPM). These were set at 0.25 for the forest (Tate et al. 1993) and 0.67 for the grassland (Tate et al. 1995). The principal source of uncertainty in estimating NPP was the assumption that the rate-modifying factors in the model for the top 23 cm of the mineral soil were also applicable for estimating C inputs to the lower soil depths.
Changes in soil chemical and physical properties with mineral soil depth generally followed the same patterns for the forest and grassland profiles (Table 1). When comparisons were made by horizon, however, some differences were apparent down to about 40 cm; below this depth, most soil properties were similar. Clay contents were slightly higher at some depths in the forest than in the grassland soil, with a clay bulge occurring in both profiles at about 25 cm.
Soil chemical properties
Beech litter and mineral soil were more acidic than those of comparable grassland samples to the base of the Bs horizon. At deeper profile depths, pH values were similar in both ecosystems.
The profile distributions of base saturation and cation exchange capacities for the two ecosystems were similar, but concentrations of Al-humus (pyrophosphate-extractable Al) were generally higher in the grassland than in the forest soil (Table 1). On an area basis, however, the grassland and forest soils had similar Al-humus contents (3.4 and 3.0 kg Al/[m.sup.2], respectively), although there was a pronounced bulge at about 46 cm depth in the grassland soil.
Soil total C and N concentrations generally declined with depth at both sites (Table 1); at comparable depths, no major differences were apparent between the ecosystems, but C to N ratios were higher at all depths in the forest, especially in the litter (data not shown). On an area basis (Table 2), however, soil C contents to 53 cm depth were greater in the grassland soil than in the forest soil and, in the core samples, were 18.4 [+ or -] 0.8 (s.d.) and 13.0 [+ or -] 0.3 kg/[m.sup.2], respectively. The between-soil difference to the base of the profiles was more modest, with C contents being 19.9 kg/[m.sup.2] for the grassland soil and 16.7 kg/[m.sup.2] for the forest soil. Ecosystem differences in soil N to 53 cm depth were even more apparent, with 0.89 and 0.53 kg N/[m.sup.2] in the grassland and forest soils, respectively. The agreement between the C contents of samples collected from the pits and from cores was generally good, with no consistent differences between the forest and grassland samples at different depths (Table 2).
Table 2. Total C distribution in soil profiles, expressed on an area basis, from samples collected in pits and by multiple coring
Total C Soil (kg/[m.sup.2]) depth(A) (cm) Pit Core(B) Forest 0-26 8.4 7.2 [+ or -] 0.3 26-53 5.3 5.8 [+ or -] 0.3 53-88 3.7 3.7 Total 17.4 16.7 Grassland 0-26 10.1 9.6 [+ or -] 0.9 26-53 6.4 8.8 [+ or -] 0.7 53-66 1.5 1.5 Total 18.0 19.9
(A) Sample depths were from the top of the litter to a depth of 26 cm (0-23 cm of mineral soil).
(B) From 26-53 cm depth, samples were collected from pit faces and analysed separately, with C values given as means [+ or -] s.d.; below 53 cm, C analyses were from a single face of each pit.
Soil [Delta] [sup.14]C
The distribution of radiocarbon in both profiles indicated progressively more negative [Delta] [sup.14]C values with depth, and no discontinuities that could be attributed to past erosion events (Table 1). The [Delta] [sup.14]C values were more negative in the grassland than in the forest profile at all comparable depths. A value of -440 [+ or -] 5 [salinity] was measured at 54-66 cm in the grassland, with a comparable [Delta] [sup.14]C value being attained at 65-80 cm in the forest profile.
Soil bulk density, pore size distribution, and water contents
Dry bulk densities ranged between 0.06 and 0.19 t/[m.sup.3] (mean 0.12 [+ or -] 0.04) for the forest FH material and from 0.02 to 0.08 t/[m.sup.3] (mean 0.04 [+ or -] 0.01) for the grassland litter. Standing dead litter in the C. pallens tussock crowns was not included in the litter compartment for bulk density estimates or C modelling purposes, but comprised the major litter C pool (0.83 kg C/[m.sup.2] v. 0.47 kg/[m.sup.2] for surface litter) in the grassland ecosystem (R. B. Allen and P. H. Harcombe, pers. comm.).
Dry bulk densities of the mineral soil progressively increased with depth in both profiles, except at 54-66 cm depth in the grassland (Table 1).
Total porosity (%) was similar in the profiles of the forest and grassland soils (Fig. 1a), whereas major differences were apparent in the distribution of macropores (Fig. 1 b). At each of the 3 depths where direct comparisons were made, the percentage of macropores ([is greater than] 60 [micro]m diam.) was at least twice as high in the forest soil as in the grassland soil, and 9 times higher at about 30 cm depth (36% v. 4%).
[Figure 1 ILLUSTRATION OMITTED]
Intact cores of the grassland mineral soil retained much more water between -5 and -20 kPa than the forest soil (Table 3), with differences least in the bottom-most depth range. Water-filled pore space was generally much higher at all soil water potentials and depths in the grassland than in the forest soil (Table 3).
Table 3. Relationships between soil water potential, soil water content, and water-filled pore space (WFPS)
Water content % (w/w) WFPS % (v/v) Depth (cm) -5 kPa -10 kPa -20 kPa -5 kPa -10 kPa -20 kPa Forest 3-8 n.d. n.d. n.d. n.d. n.d. n.d. 8-19 n.d. n.d. n.d. n.d. n.d. n.d. 19-34 66 60 55 54 49 45 34-44 56 53 50 72 67 63 44-54 43 41 38 74 72 62 Grassland 3-13 115 110 99 72 72 62 13-20 108 105 97 87 84 78 20-24 107 105 97 93 91 84 24-34 91 89 85 95 93 89 34-46 75 73 71 88 85 82 46-54 51 50 48 91 89 85
n.d., not determined; abundant roots and light soil texture at the two uppermost depths of the soil profile of the forest site prevented cores being taken.
In situ soil respiration
Rates of soil [CO.sub.2]-C efflux (mg [CO.sub.2]-C/[m.sup.2].h) measured by the static chamber method in March 1993 were similar at the grassland (69 [+ or -] 13) and the forest (84 [+ or -] 19) sites; mean soil temperatures over the 24-h period at the 2 sites were 8.2 and 8.4 [degrees] C, respectively. At similar soil temperatures in March 1994, soil [CO.sub.2]-C efflux (mg [CO.sub.2]-C/[m.sup.2].h) measured by the infrared gas analyser was 115 [+ or -] 97 at the grassland site and 226 [+ or -] 24 at the forest site. When the 1993 values were adjusted to those that would have been obtained with the infrared gas analyser, using the relationship:
Static chamber flux = 23.63 In (dynamic flux) - 39.65
(Jensen et al. 1996), the corresponding mean values were 117 [+ or -] 80 and 254 [+ or -] 217 mg [CO.sub.2]-C/[m.sup.2].h at the grassland and forest sites, respectively. Although these mean values differed markedly between the two ecosystems, the difference is not significant because the relationship used to adjust the 1993 soil [CO.sub.2]-C efflux amplified the uncertainty associated with the forest-soil value.
The respiration of the decomposing standing dead litter in the crown of the tussock grasses was not included in soil [CO.sub.2] flux measurements, but was estimated from the mass of standing litter (1.8 kg/[m.sup.2]) and its respiration at 25 [degrees] C (40 mg [CO.sub.2]-C/kg.h; Ross et al. 1996) using a [Q.sub.10] of 2. When respiration at the grassland site was adjusted to include the contribution from standing dead litter, [CO.sub.2]-C production values were 141 [+ or -] 96 and 139 [+ or ] 115 mg/[m.sup.2].h for 1993 and 1994, respectively.
As the water balance at both sites was positive for most months of the year, an annual soil [CO.sub.2]-C efflux was calculated from a soil temperature-soil respiration relationship established for a non-water-limited beech forest site (Tate et al. 1993). Annual soil [CO.sub.2]C efflux at our Craigieburn forest site would then be about 0.5 kg [CO.sub.2]-C/[m.sup.2].year. A comparable annual value appears likely at the grassland site, given the large variability associated with the respiration for this site, and the similar basal soil respiration observed for both ecosystems (Ross et al. 1996).
Mountain beech is a relatively long-lived species (about 300 years; Wardle 1984) that is not subject to major disturbance by browsing animals and was the only canopy species present at the forest site. The alpine tussock-grassland site was also dominated by one species, Chionochloa pallens. These grasslands have `many of the characteristics of a forest and few of those of a short-rotation pasture', with slow growth rates and extreme longevity of individual species being typical (Mark 1993). The grassland site had not recently been disturbed by either fire or avalanche effects. Furthermore, the absence of major discontinuities in the soil properties with depth in both ecosystems (Table 1) suggests no obvious past disturbance from erosion events. Consequently, the two sites appear ideal for this comparative study of below-ground C storage processes. We have accordingly assumed steady-state conditions (i.e. annual C inputs equal C losses) when estimating the soil C balance. We do, however, recognise that inter-site comparisons would have been strengthened if replicate pits had been sampled. Nevertheless, the fact that the morphological characteristics of the pit profiles were modal for the sites, together with the similarity of the total C values measured in the pit and multiple-core samples, strongly suggest that our data are representative of the sampling areas selected.
At comparable soil depths, considerably more soil C was stored at the grassland site than at the forest site. However, differences in soil C contents in the 2 ecosystems were less marked when the deeper forest soil profile was taken into account. Our grassland site had not been subjected to direct human interference, and we cannot invoke management effects (Nepstad et al. 1994) to explain these results. The explanation must instead be found in a difference in NPP between the grassland and the forest, or in the edaphic response to plant growth and root activity.
Carbon assimilated by the vegetation reaches the soil directly from litterfall, and indirectly from roots as soluble exudates, mucilage, dead root material, and [CO.sub.2].
At two mountain beech stands at 1210 m, and a few hundred metres from our site, annual litterfall averaged 0.25 (range 0.24-0.25) kg/[m.sup.2] and twig fall 0.08 (range 0.082-0.085) kg/[m.sup.2] (RB Allen, pers. comm.); total litterfall C was, therefore, about 0.17 kg C/[m.sup.2]. Large woody debris for an old ([is greater than] 150 years) mountain beech stand was 2.4 kg/[m.sup.2] (Allen et al. 1997). These estimates of fine litterfall and large woody debris can be regarded as only approximate as they are subject to large inter-annual variability caused by the effects of occasional severe winter storms (Allen et al. 1997). No litterfall data were available for our tussock grassland site, but annual litterfall averaged 0.26 (range 0.21-0.29) kg/[m.sup.2] for C. macra at comparable altitudes in Otago, New Zealand (Meurk 1978).
Total above- and below-ground production was estimated for both our sites by a soil C-turnover modelling approach (Jenkinson et al. 1992), which had been used successfully for estimating NPP in an old growth lowland beech forest (Tate et al. 1993). Rough estimates of below-ground production can then be obtained by subtracting the above-ground litterfall from NPP. The close agreement between the measured and modelled soil microbial C values (Table 4) suggests that the steady-state assumption is reasonable (Jenkinson 1990). The simulated annual C inputs of 0.41 and 0.54 kg/[m.sup.2] for the forest and grassland sites, respectively, provide approximate estimates of NPP for the two ecosystems (Jenkinson et al. 1992).
Table 4. Model estimates of total annual inputs to soil, and inert organic matter (IOM)
Soil Total depth soil C Radiocarbon(B) (cm) (kg/[m.sup.2]) ([salinity]) Forest 0-26(A) 7.2 [+ or -] 0.3 +51.5 26-53 5.8 [+ or -] 0.3 -145 53-85 3.7 -398 Est. NPP Grassland 0-26(A) 9.6 [+ or -] 0.9 -23.3 26-53 8.8 [+ or -] 0.7 -348 53-63 1.5 -438 Est. NPP Microbial C Soil (kg/[m.sup.2]) depth Annual C (cm) Measured(C) Modelled input(D) Forest 0-26(A) 0.14 [+ or -] 0.01 0.14 0.21 26-53 0.07 [+ or -] 0.01 0.09 0.14 53-85 n.d. 0.04 0.06 Est. NPP 0.41 Grassland 0-26(A) 0.22 [+ or -] 0.03 0.18 0.31 26-53 0.10 [+ or -] 0.01 0.11 0.19 53-63 n.d. 0.02 0.04 Est. NPP 0.54 Soil Inert IOM-C depth organic C (as % total (cm) (kg/[m.sup.2]) soil C) Forest 0-26(A) 0.15 2 26-53 1.19 21 53-85 1.62 44 Est. NPP Grassland 0-26(A) 1.00 10 26-53 3.55 40 53-63 0.76 51 Est. NPP
n.d., not determined.
(A) Litter + 23 cm depth of mineral soil.
(B) Estimated by depth-weighting the sum of [Delta][sup.14]C values for all horizons in Table 1. Where depth increments straddle soil horizons, the [Delta][sup.14]C for that part-horizon was apportioned to each depth increment.
(C) Microbial C values based on those presented in Ross et al. (1996).
(D) Units: kg/[m.sup.2].year. Litterfall C (kg/[m.sup.2].year) was about 0.17 at the forest site, and, as estimated from Meurk (1978), was about 0.13 at the grassland site.
An apparent inconsistency exists, however, when the modelled C input rates and litter and soil [Delta][sup.14]C values are compared with rates of decomposition in the two ecosystems. The modelled results suggest that higher, rather than lower, concentrations of `modern' C should have been observed in the grassland soil than in the forest soil, i.e. the [Delta][sup.14]C values for the grassland litter and mineral soil layers should have been more positive. Our experimental results, by contrast, indicate that (modern) [sup.14]C concentrations were lower in the grassland litter and soil layers than at comparable depths in the forest soil (Table 4). This seeming anomaly can best be explained by the markedly slower turnover rate of grassland-litter C (approx. 50 years) than forest-litter C (approx. 27 years), when estimated from the [Delta][sup.14]C data (Parshotam and Tate 1997); this slow turnover rate of the tussock litter is consistent with previous litter-bag decomposition experiments at a similar altitude (Williams et al. 1977). In 1992, when our samples were taken, there would therefore have been a higher proportion of `pre-bomb' (before 1956) [sup.14]C entering the soil at the grassland site.
We have previously suggested (Tate et al. 1993) that annual C inputs integrated over centuries are unlikely to be exactly synchronised with [CO.sub.2]-C efflux estimates based on measurements and modelling over just 1 year. The estimate of annual soil [CO.sub.2] efflux of about 0.5 kg [CO.sub.2]-C/[m.sup.2] from the forest site nevertheless agrees quite well with the NPP for this site. This estimate includes [CO.sub.2] from root respiration, which contributed about 23% of the total forest floor respiration in a lowland old-growth beech forest (Tate et al. 1993). If roots had contributed a comparable percentage of the total respiration from the Craigieburn forest floor, then about 0.4 kg [CO.sub.2]-C/[m.sup.2].year would have originated from the soil. Subtraction of the annual fine litterfall C [about 0.2 kg C/[m.sup.2] for the forest site (RB Allen, pers. comm.), and about 0.1 kg C/[m.sup.2] for the grassland site] from the simulated annual C inputs suggests that about 0.2 kg [CO.sub.2]-C/[m.sup.2] of the total annual C input in the forest, and about 0.4 kg [CO.sub.2]-C/[m.sup.2] in the grassland, came from roots. This root C-input value for the forest is equivalent to about 40% of the annual total soil [CO.sub.2]-C efflux, and falls within the range reported for other forest ecosystems (Raich and Nadelhoffer 1989).
Although the greater soil C content in the grassland may be partly explained by a higher annual C input than in the forest ecosystem, the difference between annual C inputs for the two ecosystems is small, given the uncertainty of at least [+ or -] 0.03 kg C/[m.sup.2] (Tate et al. 1995). Other factors that could also have been responsible for the greater soil C content in the grassland than in the forest system are now considered.
Influence of soil Al, and physical factors on soil C content
In order to balance soil C budgets and explain the existence of `old' (based on radiocarbon dating) soil C, models like ROTHC include a highly recalcitrant C fraction; in ROTHC it is inert organic matter (IOM) (Jenkinson 1990). Several factors could account for the existence of this recalcitrant soil C fraction, including its chemical structure (Theng et al. 1992; Golchin et al. 1995), physical protection (Theng et al. 1986), distribution in the soil profile (Elzein and Balesdent 1995), and location in the soil matrix (Bonde et al. 1992; Theng et al. 1992; Golchin et al. 1995). Although all these mechanisms could contribute to the existence of IOM at both our sites, they cannot satisfactorily explain the greater content of total and recalcitrant (IOM) C in the grassland soil, especially at 26-53 cm depth (Table 4).
Al-humus can impede C mineralisation (Bruner and Blaser 1989) and its accumulation may explain the very long soil C turnover times observed in some New Zealand tussock grassland soils (Tate 1992). In this study, Al-humus concentrations were higher in the grassland soil (Table 1), and complexing of humus by Al may, therefore, explain some of the between-soil differences observed in C contents. Soil acidity and clay contents, however, are unlikely to have been responsible for the higher C content in the grassland soil because they were slightly lower, rather than higher, in the grassland than in the forest soil at most depths (Table 1).
Nitrogen concentrations in the mineral soil were also much higher in the grassland than in the forest. Although a higher N content may explain greater soil C accumulation in some forest ecosystems (Berg et al. 1996), higher N concentrations and availability in grassland soils may, instead, favour decomposition and lead to reduced soil C content (Tate et al. 1991). We are, therefore, unable to provide an unequivocal explanation of the higher soil C in our grassland soil based solely on the ecosystem differences in N content.
Forest soils can contain a higher proportion of macropores than grasslands (Davis 1994; R. J. Jackson, pers. comm.), and one of the most striking differences between our two soils is in pore-size distribution (Fig. 1b). As the proportion of macropores decreases, more water is retained in the soil profile, increasing the number of anaerobic microsites which could favour organic C accumulation. This mechanism could be operating particularly strongly in the grassland soil (Fig. 1b). In addition, although both sites receive the same total precipitation, a greater proportion of the intercepted precipitation reaches the soil surface in the grassland because of lower wet-canopy evaporation. Greater snow accumulation, and lower water loss by transpiration in the grassland than in the forest (Fahey and Rowe 1992) are also likely. The water content in the grassland soil is, therefore, generally higher than in the forest soil.
Thus, recalcitrant C formation could have been influenced by soil pore-size distribution and the percentage of water-filled pore space. A similar mechanism may have been responsible for the spatial, 3-fold variation in IOM observed previously over just a few metres in the soil of a lowland old-growth beech forest (Tate et al. 1993); in areas with the highest IOM contents, macropores comprised only 4% of total porosity. These results together indicate that drainage class may be a more useful criterion than clay content for predicting soil C distribution in some ecosystems. A similar conclusion was reached by Davidson and Lefebvre (1993) to account for large-scale spatial variations in soil C in cool regions of northern grasslands and temperate and boreal forests.
Inert organic matter
IOM-C as a proportion of total C increased markedly with depth in the grassland soil (Table 4), with only marginal changes in the high WFPS values (Table 3). This strongly suggests that the grassland soil would be frequently water-saturated, even at shallow depths, and that aeration status and decomposition could thereby be restricted, especially during prolonged periods in spring during snow melt. In the forest soil, IOM-C comprised a lower percentage of total C than in the grassland soil at comparable depths. At -20 kPa, WFPS values in the forest soil ranged between 45% and 63% (Table 3), well below those where anaerobic conditions might be expected (Linn and Doran 1984). One or more of the other mechanisms discussed earlier could, therefore, be mainly responsible for the IOM accumulation in the forest soil.
Estimation and interpretation of the IOM fraction in soil C turnover models such as ROTHC continue to present a formidable challenge (Falloon et al. 1998). The perception that IOM is a discrete fraction (e.g. Golchin et al. 1995; Theng et al. 1992) may need to be revised if further studies suggest that recalcitrant C can be, to some extent, a consequence of environmental constraints on decomposition. The corollary to this hypothesis is that if the environmental constraint is removed, then the `recalcitrant' C fraction would no longer be recalcitrant.
Plant effects on C inputs, soil Al, and soil moisture and other physical characteristics, rather than differences in soil mineral weathering or texture, can best explain the patterns of soil C accumulation observed in these adjacent forest and grassland ecosystems at treeline. The greater accumulation of soil C in the grassland appears to have resulted mainly from the markedly lower macroporosity and greater soil water retention in the grassland soil at depths below 20 cm. Interpretation of the inert organic matter fraction in soil C turnover models may need to be revised in light of this study, which suggests that recalcitrant C can be a consequence of an environmental constraint on decomposition rather than a discrete fraction. Further studies at other comparable undisturbed areas of indigenous forest and grassland are now needed to confirm the generality of our observations.
We thank Alan H. Nordmeyer, Forest Research, Christchurch, and Rob B. Allen, Landcare Research, Lincoln, for their assistance in site selection and description of the vegetation; Tony McSeveny, Landcare Research, Lincoln, for providing the climate data; Trevor Webb, Landcare Research, Lincoln, for selecting the pit locations and describing and classifying the soils; and Rick Jackson, Landcare Research, Lincoln, for constructive comments on the soil physics. The research was supported by the New Zealand Foundation for Research, Science and Technology.
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Manuscript received 27 July 1999, accepted 29 November 1999
K. R. Tate(AC), N. A. Scott(A), D. J. Ross(A), A. Parshotam(A), and J. J. Claydon(B)
(A) Landcare Research, Private Bag 11052, Palmerston North, New Zealand.
(B Landcare Research, Private Bag 3127, Hamilton, New Zealand.
(C) Corresponding author; email: email@example.com
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|Author:||Tate, K. R.; Scott, N. A.; Ross, D. J.; Parshotam, A.; Claydon, J. J.|
|Publication:||Australian Journal of Soil Research|
|Article Type:||Statistical Data Included|
|Date:||May 1, 2000|
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