Small altitudinal change and rhizosphere affect the SOM light fractions but not the heavy fraction in European beech forest soil

Catena 181 (2019) 104091

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Small altitudinal change and rhizosphere affect the SOM light fractions but not the heavy fraction in European beech forest soil

T

M. De Feudisa, , V. Cardellib, L. Massaccesia, S.E. Trumborec, L. Vittori Antisarid, S. Coccob, G. Cortib, A. Agnellia,e ⁎

a

Department of Agricultural, Food and Environmental Sciences, Università degli Studi di Perugia, Perugia, Italy Department of Agricultural, Food and Environmental Sciences, Università Politecnica delle Marche, Ancona, Italy c Department of Biogeochemical Processes, Max Planck Institute for Biogeochemistry, Jena, Germany d Department of Agricultural and Food Sciences, Alma Mater Studiorum, Università di Bologna, Bologna, Italy e Institute of Ecosystem Study (ISE-CNR), Sesto Fiorentino, FI, Italy b

ARTICLE INFO

ABSTRACT

Keywords: Rhizosphere effect Organic C pools Mountain soils Density fractionation Climate change

We evaluated the influence of the rhizosphere, soil depth, and altitude on the amount and nature of the density separated soil organic matter (SOM) pools. Samples were collected from the A and AB horizons of European beech (Fagus sylvatica L.) forest soils located at two altitudes (800 and 1000 m) along 1° of latitudinal gradient in central Italy, by using altitude as a proxy for air temperature change. Specifically, we hypothesized that: i) larger amount of labile organic matter, comprising of fresh plant and organism residues and easily degradable molecules, was allocated in the rhizosphere than in the non-rhizosphere soil, and ii) the temperature had a stronger effect on the C pools of the rhizosphere than in that of the non-rhizosphere soil. At both altitudes, we found more organic C (OC) associated with the light fractions of the rhizosphere than in those of the non-rhizosphere soil and, specifically in the rhizosphere free light fraction, larger OC concentrations were observed at 1000 m than at 800 m above sea level. These higher amounts of OC have been attributed to roots, which are one of the main source of particulate organic matter, and their activity and turnover increase when the environmental conditions become more restrictive, as it happens at higher altitude. Conversely, no effect related to rhizosphere and altitude on the OC associated to the heavy fraction was found. The recalcitrance of the OC of the heavy fraction has been ascribed both to its protection due to the tight bounds to mineral particles and to its degradation degree, as indicated by δ13C values, which were greater than those of the light fractions. The similar 14C signature and the presence of recent C in all the density fractions of rhizosphere and non-rhizosphere soil of both A and AB horizons suggested the occurrence of a rapid incorporation of fresh organic matter into the mineral horizons, followed by occlusion into aggregates and adsorption on mineral surfaces. Further, the lack of different Δ14C values between the fractions at 800 and 1000 m could indicate that a temperature change of 1 °C is not sufficient to induce marked changes in SOM cycling.

1. Introduction

occlusion within organo-mineral complexes that make organic C less accessible to microbial attack (Han et al., 2016). Schrumpf et al. (2013) studied SOC stability across a range of European soils that varied in vegetation, soil type, parent material, and land use, and concluded that organics are more stable when they are occluded within aggregates. Further, besides aggregate occlusion, the interaction of organic compounds with di- and trivalent cations promotes the preservation of organic matter over decades or longer timescales (Kaiser et al., 2016). At larger spatial scale, SOC content is strongly correlated with the mean annual air temperature that, in turn, is controlled by physiographic factors such as latitude, altitude, and exposure (Soethe et al., 2007; Tsui

Soil comprises the largest pool of terrestrial C (Jackson et al., 2017) and, through the soil organic matter (SOM) cycling, it represents either an important sink of atmospheric CO2 or a possible source of greenhouse gases (Bispo et al., 2017). SOM includes a wide range of compounds at different stage of decomposition derived from litter, root turnover, and microorganisms, and its dynamics is controlled by substrate quality, biological activity, and environment (Dungait et al., 2012). Furthermore, the stability of soil organic C (SOC) is enhanced by interactions with minerals having large specific surface area and by ⁎

Corresponding author at: Dipartimento di Scienze Agrarie, Alimentari ed Ambientali, Università degli Studi di Perugia, Borgo XX Giugno 72, 06121 Perugia, Italy. E-mail address: [email protected] (M. De Feudis).

https://doi.org/10.1016/j.catena.2019.104091 Received 12 February 2018; Received in revised form 21 May 2019; Accepted 27 May 2019 Available online 05 June 2019 0341-8162/ © 2019 Elsevier B.V. All rights reserved.

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et al., 2004). Temperature can influence SOC cycling both directly, through its impact on microbial metabolism (Blume et al., 2002; Chen et al., 2016; Xu et al., 2014), and indirectly, through long-term effects on soil properties such as the mineralogy (Rech et al., 2001). However, the effect of temperature changes on SOC dynamics is still poorly known because of the scarce information on how, and how fast, changes affect different stabilization mechanisms (Trumbore, 2009). The rhizosphere, the soil in proximity to the roots, is the edaphic compartment most influenced by plant roots (Agnelli et al., 2016; Massaccesi et al., 2015) and where most of the soil chemical, biogeochemical and biological reactions take place (e.g., Lambers et al., 2009). The rhizosphere can be considered as a C cycle hotspot that, in stable conditions such as those of an old-growth forest, can influence size and rate of SOM storage (Agnelli et al., 2016; Tefs and Gleixner, 2012). Through the process of rhizodeposition, roots can release a wide range of C compounds, including root caps and border cells, dead root cells (cortex, root hairs, etc.), and root exudates (Jones et al., 2009). The fate of these rhizodeposits, which represent 15–40% of the C photosynthetically fixed by plants (Gunina and Kuzyakov, 2015), might differ from that of SOC found far from the roots. According to Gunina and Kuzyakov (2015), most of the compounds released through exudation are decomposed to CO2 by microorganisms, and only a small fraction of those substances (2–5%) accumulates in the SOC pools (Helal and Sauerbeck, 1989; Hutsch et al., 2002). Conversely, Strickland et al. (2012) found that most of the glucose-C released by roots in forest and pasture soils was fixed as microbial biomass and SOC, with lesser amounts lost as CO2. The compounds released by roots can also become part of SOM stable pool (Han et al., 2016). Indeed, low molecular weight organic acids can easily bind to mineral surfaces whereas polysaccharides can form strong complexes with minerals and organic substances (Rasse et al., 2005). Also, aliphatic molecules with long C chain (for example, suberin) can persist for a long time due to their hydrophobicity, so contributing to slow the cycling of SOC pools (Ji et al., 2015). However, the release of easily degradable compounds in the rhizosphere also affects the decomposition rate of more stable organic C by enhancing the activity of the rhizosphere microbial community through a “priming effect” (Kuzyakov et al., 2007). Since the temperature influences both rhizodeposition (Yin et al., 2013) and SOC pools size and dynamic (Lehmann and Kleber, 2015), we tested the SOM pool sensitivity to temperature in the rhizosphere and non-rhizosphere soil of a forest environment by a study protocol that uses altitude as a proxy for temperature change (Gutiérrez-Girón et al., 2015). Specifically, we contrasted the organic C pools of rhizosphere and non-rhizosphere soil from the A and AB horizons of limestone-derived soils under European beech (Fagus sylvatica L.) forest at 800 and 1000 m above sea level along 1° of latitudinal gradient in central Italy (from 42°28′ N to 43°28′ N). The sites at the two altitudes differed in their mean annual air temperature by on average 1 °C, corresponding to the expected global increase of the air temperature to the year 2050 (IPCC, 2013). We hypothesized that i) larger amount of labile organic C was allocated in the rhizosphere than in the non-rhizosphere soil, and ii) the temperature had a stronger effect on the C pools of the rhizosphere soil than in that of the non-rhizosphere soil. To test these hypotheses, rhizosphere and non-rhizosphere soil samples were fractionated by density and the obtained pools were characterized by chemical and isotopic (13C and 14C) means.

(a.s.l.). At the three sites, the soils developed from limestone and were classified as Mollisols (Hapludoll or Haprendoll) or Inceptisols (Humudept) according to Soil Survey Staff (2014). The pH ranged between 6.6 and 6.9 at 800 m a.s.l. and between 6.9 and 7.7 at 1000 m a.s.l. In the study sites the mean annual air temperature (MAAT) was 10 °C at 800 m and 9 °C at 1000 m, with January and July as the coldest and warmest months, respectively. The mean annual precipitation varied between 825 and 1430 mm, but for each site was similar at the two altitudes. Conversely to other works that considered wide temperature ranges (e.g., Bu et al., 2012; Siles et al., 2017), we deliberately chose to evaluate the effect of a narrow difference (1 °C) in air temperature between the two elevations to compare similar ecosystems (same lithology, soil types, mineralogy, vegetation, exposure, and forest management) differing for MAAT. Details on the general features (e.g., vegetation and forest management, soil description and classification) of the investigated areas are reported in De Feudis et al. (2016). 2.2. Soil sampling During early March 2014, at both altitudes of Mount Terminillo, Mount San Vicino and Mount Acuto, we selected one area of about 0.4–0.5 ha with homogeneous vegetation where to evaluate the soil spatial variability by taking into consideration the thickness of the solum (A plus B horizons). Within each study area, by manual auger holes we assessed that the solum thickness at Mount Terminillo ranged from 15 to 75 cm at 800 m a.s.l., and from 18 to 45 cm at 1000 m a.s.l.; at Mount San Vicino, it ranged from 30 to 91 cm at 800 m, and from 35 to 55 cm at 1000 m; at Mount Acuto it ranged from 90 to 115 cm at 800 m, and from 68 to 85 cm at 1000 m. In each of the six study areas the soil survey was then refined by digging mini-pits and additional auger holes so to define two plots of about 100 m2 representative of the solum thickness extremes. In each plot a soil pit was dug, for a total of 12 pits (3 sites × 2 altitudes × 2 pits). The sampled soils had a solum thickness of 19 and 66 cm at 800 m a.s.l., and of 24 and 38 cm at 1000 m a.s.l. for Mount Terminillo; of 37 and 85 cm at 800 m, and of 40 and 46 cm at 1000 m for Mount San Vicino; of 102 and 108 cm at 800 m, and of 75 and 79 cm at 1000 m for Mount Acuto (De Feudis et al., 2017b). Each pit was dug at about 50–60 cm from the stem of the biggest beech tree located within the plot. The age of these trees varied from 45 to 65 years. The sampling was accomplished in late winter with the aim to reduce possible influences due to the intense root activity and biochemical processes occurring at the soil-root interface in the warmer seasons (Buée et al., 2005; Calvaruso et al., 2014; Ruehr and Buchmann, 2010) and, hence, have a more stable picture of the rhizosphere C pools. The study was conducted on the mineral topsoil where most of the SOM accumulates. In particular, both the A and the underlying AB horizons were sampled separately in a large amount (at least 3 kg) from each pit, and the samples were stored in a portable refrigerator. The mean thickness of the A and AB horizons was, respectively, 9.0 and 8.5 cm at 800 m, and 17.0 and 9.5 cm at 1000 m on Mount Terminillo; 9.5 and 6.5 cm at 800 m, and 6.5 and 9.0 cm at 1000 m on Mount San Vicino; 9.0 and 16.0 cm at 800 m, and 19.0 and 7.5 cm at 1000 m on Mount Acuto. Once in laboratory, the roots with a diameter < 2 mm together with the adhering soil were picked up from each bulk sample (Cocco et al., 2013; Massaccesi et al., 2015). The roots with a diameter > 2 mm were discarded. The remainder of the sample represented the soil not strictly adhering to the roots and was called non-rhizosphere soil (NRS). Following the procedure reported in Cocco et al. (2013) and Massaccesi et al. (2015), the collected roots were gently shaken to detach the weakly adhering soil particles, which were then added to the NRS. The soil material firmly adhering to the roots was considered as rhizosphere soil (RS) and recovered by further shaking and soft brushing. Both RS and NRS samples were air-dried and sieved through a

2. Materials and methods 2.1. Study sites The study sites were all from three calcareous massifs of the central Apennines (Italy): Mount Terminillo (42°28′ N, 12°56′ E), Mount San Vicino (43°19′ N, 13°03′ E), and Mount Acuto (43°28′ N, 12°41′ E). For each mountain, European beech (Fagus sylvatica L.) forests were chosen on the northern facing slopes at about 800 and 1000 m above sea level 2

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2-mm mesh.

(HE) and fulvic extract (FE) were separated by centrifugation at 2000g for 10 min. The OC content of HE and FE (HE–C and FE–C, respectively) was obtained by K-dichromate digestion method, heating the suspension at 180 °C for 30 min. The amount of non-extractable organic C (NEOC) contained in the Hf was calculated by subtracting HE–C and FE–C from the organic C content of Hf.

2.3. Chemical analysis of rhizosphere and bulk soil samples To prevent inorganic C detection, the content of total organic C (TOC) was estimated by K-dichromate digestion, heating the suspension at 180 °C for 30 min. Total N content was determined by a Carlo Erba EA1110 dry combustion analyzer (Carlo Erba Instruments, Milan, Italy). Water-extractable organic matter was obtained with the following procedure (Agnelli et al., 2014): 10 mL of distilled water were added to 1 g of soil sample (solid:liquid ratio 1:10) and shaken overnight on orbital shaker (140 rpm). The suspension was centrifuged at 1400g for 10 min, and then filtered through a 0.45 μm cellulose membrane filter. The water-extractable organic C (WEOC) in the resulting solution was analyzed with a TOC-500A (Shimatzu, Tokyo, Japan) analyzer after the addition of a few drops of concentrated H3PO4 to eliminate carbonates. Available P was estimated according to Olsen et al. (1954).

2.6. Isotopic analyses of the soil density fractions Stable carbon isotope (13C) natural abundance was determined on aliquots of purified CO2 obtained from the sample combustion. Aliquots of FLf, OLf, and Hf were weighed into tin capsules, treated with 2 M H2SO4 solution to remove carbonates, and dried at 60 °C. Then, the aliquots were combusted in an O2 stream using an elemental analyzer (NA 1110, CE Instruments, Milan, Italy). The evolved CO2 was analyzed using an isotope ratio mass spectrometer (IRMS Delta C or DELTA+XL, Thermo Finnigan MAT, Bremen, Germany). The obtained values were expressed as δ13C: 13C

2.4. Density fractionation

= [( 13C/ 12Csample –13C/ 12Cstandard)/(13C/ 12Cstandard)] x 103

relative to the international reference standard v-PDB using NBS19 (Werner and Brand, 2001). Radiocarbon sample preparation and analysis were performed at the 14 C laboratory of the Max Plank Institute, Jena, Germany (Steinhof et al., 2004). As for 13C analyses, aliquots were weighed into tin capsules and submitted to acid treatment to remove carbonates. The samples were then combusted by the NA 1110 elemental analyzer and the evolved CO2 was transferred into a glass tube cooled by liquid N2 and containing an iron catalyst. Reduction of CO2 to graphite was carried out at 600 °C under H2 atmosphere. The obtained graphite was analyzed for its radiocarbon content by 14C accelerator mass spectrometry (AMS; 3MV Tandetron 4130 AMS 14C system; High Voltage Engineering Europe, The Netherlands). The 14C values were expressed as Δ14C, namely the ‰ deviation from 14C/12C ratio of oxalic acid standard in 1950 (oxalic acid standard NBS SRM 4990C), after normalization to a δ13C of −25‰ to correct for mass-dependent isotopic fractionation effects (Stuiver and Polach, 1977), and correction for decay between 1950 and time of analysis (2015).

The fractionation was performed on both the RS and the NRS following the method used by Schrumpf et al. (2013) with some modifications. In 250 mL centrifugation tubes, 12 g of soil and 50 mL of sodium polytungstate (SPT) solution with a density (ρ) of 1.6 g cm−3 were mixed. The tubes were gently shaken on a horizontal shaker so to release the free light fraction (FLf), comprising of plant and animal residues at different stages of decomposition. Suspensions were allowed to rest for 1 h and the floating FLf (ρ ≤ 1.6 g cm−3) was aspirated together with part of the SPT solution and filtered on C-free glass fiber filters. The SPT solution recovered from this step was checked for its density, added with fresh SPT to bring it to the wanted density (ρ = 1.6 g cm−3), and added to the same sample for further fractionation. To obtain the occluded light fraction (OLf), the organic debris physically protected inside the aggregates, the suspension with the remnant of the previous extraction was sonicated into the tubes with output energy of 300 J mL−1, and centrifuged at 5500g for 30 min. All the material with ρ ≤ 1.6 g cm−3 released from the aggregates disintegrated after sonication (OLf) was aspirated from the suspension and filtered by glass fiber filters as described earlier. Both FLf and OLf were washed using deionized water to remove SPT until the conductivity of the rinsed solution was < 50 μS. The residue with ρ > 1.6 g cm−3, called heavy fraction (Hf) and containing organo-mineral complexes, was collected on a glass fiber filter and washed with deionized water until the conductivity of the rinsing solution was < 200 μS. The conductivity cut-off of 200 μS is considered acceptable for the Hf obtained from calcareous soils due to the dissolution of carbonates (Schrumpf et al., 2013). The mean mass recovery achieved with the density fractionation ranged between 93 and 99% of the total sample masses. All the fractions were freeze-dried and ground for further analyses. The organic C (OC) content of FLf, OLf, and Hf (Table S1 of the Supplementary materials) was estimated by K-dichromate digestion method, heating the suspension at 180 °C for 30 min.

2.7. Statistical analysis Three-way ANOVA was performed to assess the effect of altitude, soil horizons, and soil fractions (RS and NRS) on the soil properties. The graphical analysis of residuals was used to verify the normality and homoscedasticity of the data, which were transformed if necessary. The transformation was selected by the maximum likelihood procedure suggested by Box and Cox (1964), as implemented in the boxcox function of the package MASS (Modern Applied Statistics with S). Since no interaction effect (altitude × soil horizon, altitude × soil fractions, soil horizon × soil fractions, altitude × soil horizon × soil fractions) occurred, Fisher's least significant difference (LSD) was used at the 95% probability level to compare means for the main effects (altitude, soil horizons, soil fractions) identified to be significant by ANOVA. The data were analyzed using R software (R Core Team, 2014).

2.5. Extraction and fractionation of humic and fulvic C from the heavy fraction

3. Results 3.1. General properties of the RS and the NRS

The Hf isolated by density separation was submitted to the extraction and fractionation of humic C. Briefly, 5 g of Hf were added of 25 mL of NaOH 0.1M solution (solid:liquid ratio 1:5) and the mixture, maintained under N2 atmosphere, was shaken for 24 h on an orbital shaker (140 rpm) at room temperature. The suspension was centrifuged at 2000g for 10 min and the supernatant filtered through a glass fiber filter. The resulting solution was acidified to about pH 1.5 with 6 M H2SO4 solution, and allowed to settle overnight. Finally, humic extract

TOC concentration was higher in the RS than in the NRS, with exception of the A horizon at 800 m where no differences occurred between the RS and the NRS (Fig. 1a). Moreover, TOC content decreased with depth and was greater at 1000 m than at 800 m for both fractions and horizons. The WEOC content did not show any significant differences between the RS and the NRS at 800 m, whereas at 1000 m the RS 3

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Fig. 1. Total organic C (a), water-extractable organic C (b), total N (c), and available P (d) contents of rhizosphere soil (white bars) and non-rhizosphere soil (grey bars) of European beech (Fagus sylvatica L.) forests at 800 and 1000 m above sea level (a.s.l.), central Apennines, Italy. The standard error (SE) and the least significant difference (LSD) were calculated on the original or transformed means (n = 3) when the effects from the ANOVA were significant. Total organic C – significant effects: altitude, soil horizon, RS vs NRS; SE = 0.082, LSD = 0.231. Water–extractable organic C – significant effects: soil horizon, RS vs NRS; SE = 0.041, LSD = 0.114. Total N – significant effects: altitude, soil horizon, RS vs NRS; SE = 1.22, LSD = 3.43. Available P – significant effects: altitude, soil horizon, RS vs NRS; SE = 0.026, LSD = 0.082.

had a greater WEOC content than the NRS in both A and AB horizons (Fig. 1b). However, no WEOC content differences occurred between horizons and altitudes, except for the NRS of the AB horizon at 1000 m, which displayed the lowest WEOC content. The total N content was similar between the RS and the NRS but declined from A to AB horizon, with a higher total N content at 1000 than at 800 m (Fig. 1c). At both altitudes, available P had a larger concentration in the RS than in the NRS and, for both fractions, decreased from A to AB horizon (Fig. 1d).

no significant difference occurred (Table 2). For both the RS and the NRS of the two altitudes, the OC in the OLf decreased from A to AB horizon. However, whereas in the RS the OLf had similar OC content at 800 m and 1000 m, in the NRS the OLf contained more OC at 1000 m than at 800 m in both horizons. The Hf represented most of the total mass recovered after the density fractionation (758 to 973 g kg−1 soil) and made up 59.9 to 82.6% of TOC (Tables 1 and 2). The OC concentration in the Hf did not show significant differences among the samples from both altitudes (Table 2). The distribution of HE–C and FE–C and NEOC in the Hf was similar between soil fractions, horizons, and altitudes, with the exception of FE–C that showed a general decline from A to AB horizon (Table 3).

3.2. Distribution of the density fractions and their organic carbon contents In the two horizons, the recovered FLf mass accounted from 16 to 154 g kg−1 soil (Table 1) and contained an amount of OC that ranged from 8.3 to 21.7% of TOC (Table 2). At both altitudes, the OC of FLf was larger in the RS than in the NRS and, with exception of the NRS at 1000 m, decreased from A to AB horizon (Table 2). Moreover, on a horizon by horizon comparison, the FLf had a greater OC content at 1000 than at 800 m with exception of the NRS of the AB horizon where no difference occurred between the two altitudes. The recovered OLf mass accounted from 11 to 87 g kg−1 soil (Table 1) and its contribution to TOC ranged from 7.7 to 19.4% (Table 2). The OC content of the OLf was larger in the RS than in the NRS, with the exception of the AB horizon at the higher altitude, where

3.3. Δ14C and δ13C of the density fractions The Δ14C values of FLf, OLf, and Hf showed no significant difference between soil fractions, horizons, and altitudes (Fig. 2). The mean radiocarbon signatures of all density fractions exceeded the atmospheric Δ14C level in 1950, indicating the incorporation of increased 14 C atmospheric concentrations in recent decades due to nuclear bomb testing occurred after 1950 (Levin and Kromer, 2004; Trumbore, 2009). The δ13C of the FLf (Fig. 2) did not significantly differ between soil fractions and between horizons. However, while the FLf was always 4

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Table 1 Density fractions recovery from rhizosphere soil (RS) and non-rhizosphere soil (NRS) of European beech (Fagus sylvatica L.) forests at 800 and 1000 m above sea level, central Apennines, Italy. For each fraction, the standard error (SE) and the least significant difference (LSD) were calculated on the original or transformed (in parentheses) means (n = 3) when the effects from the ANOVA were significant. Horizon

Free light fraction 800 m a.s.l. RS

Occluded light fraction 1000 m a.s.l.

800 m a.s.l.

Heavy fraction 1000 m a.s.l.

800 m a.s.l.

NRS

RS

NRS

RS

NRS

RS

NRS

53 (0.14) 16 (0.25)

154 (0.10) 52 (0.16)

86 (0.14) 27 (0.21)

68 (4.06) 26 (3.17)

53 (3.87) 11 (2.15)

87 (4.34) 34 (3.29)

48 (3.60) 20 (2.84)

RS

1000 m a.s.l. NRS

RS

NRS

g kg−1 soil A AB

72 (0.12) 42 (0.16)

860

894

758

866

932

973

914

953

Free light fraction – significant effects: soil horizon, RS vs NRS; SE = 0.02, LSD = 0.07. Occluded light fraction – significant effect: soil horizon; SE = 0.42, LSD = 1.27. Heavy fraction – significant effect: soil horizon; SE = 57, LSD = 171.

more enriched in 13C at 1000 m than at 800 m, δ13C of the OLf and Hf did not differ between the altitudes in the A horizon, but higher 13C values were observed at 1000 than at 800 m in the AB horizon (Fig. 2). Contrasting the 13C signatures of the three density fractions, the FLf and OLf of both the RS and the NRS from 800 and 1000 m had similar mean values ranging between about −27.8 and − 26.6‰. Conversely, a greater fractionation effect occurred for the Hf (Fig. 3), which showed values ranging between −26.6 and −25.4‰.

(e.g., litter quality, soil microbial community) factors (Garten Jr. and Hanson, 2006). In our case the higher TOC content at 1000 than at 800 m was attributed to the lower MAAT at 1000 m that, by reducing the soil microbial activity, promotes the organic matter accumulation (Blume et al., 2002; Xu et al., 2014). In agreement with other studies (e.g., Agnelli et al., 2016; Otero et al., 2015), the larger TOC concentration in the RS compared to the NRS at both altitudes was ascribed to rhizodeposition processes that release exudates and decaying root tissues in the root-surrounding soil (Sokolova, 2015). The availability of easily degradable compounds due to rhizodeposition should produce an enrichment of WEOC in the RS via: i) direct inputs of labile compounds such as carbohydrates (sugars and mucilage), aliphatic and aromatic organic acids, phenols, and fatty acids into the soil solution (ColinBelgrand et al., 2003; Toal et al., 2000), or/and ii) solubilization of SOM favoured by the exudates (Nardi et al., 2000). Since the general greater microbial biomass in the RS than in the NRS (Massaccesi et al., 2015; Tang et al., 2014), the rhizodeposition processes should also support microbial and enzymatic activities (Brzostek et al., 2013), promoting the mineralization of stable organic matter (Kuzyakov, 2010) with the consequent further release of water-soluble organic compounds. However, we found a larger content of WEOC in the RS than in the NRS only at 1000 m, where more severe conditions occurred. In fact, the lower MAAT at 1000 m than at 800 m slowed soil weathering and organic decomposition, as suggested by the reduced profile development and

4. Discussion 4.1. General properties of the RS and the NRS at 800 and 1000 m of altitude TOC and total N contents declined with soil depth and increased with altitude. The high TOC content in the A horizon was consistent with the findings of many studies on forest soils (e.g., Agnelli et al., 2016; Song et al., 2016), as well as the increased TOC concentrations at high altitude (e.g., De Feudis et al., 2016; Fernández-Romero et al., 2014). Although at high elevations a reduced aboveground net primary production and the amount of litterfall may occur (Joshi et al., 2003), the different TOC contents at the two altitudes could indicate a modification in the balance between soil C inputs and soil C losses due to changes in both abiotic (e.g., temperature, precipitation) and biotic

Table 2 Organic carbon content in the density fractions and its percentage contribution to the total organic carbon (TOC) of rhizosphere soil (RS) and non-rhizosphere soil (NRS) of European beech (Fagus sylvatica L.) forests at 800 and 1000 m above sea level (a.s.l.), central Apennines, Italy. For each fraction, the standard error (SE) and the least significant difference (LSD) were calculated on the original or transformed (in parentheses) means (n = 3) when the effects from the ANOVA were significant. A horizon g C kg Free light fraction

800 m a.s.l. 1000 m a.s.l.

Occluded light fraction

800 m a.s.l. 1000 m a.s.l.

Heavy fraction

800 m a.s.l. 1000 m a.s.l.

RS NRS RS NRS RS NRS RS NRS RS NRS RS NRS

−1

soil

14.0 7.9 27.8 12.7 17.3 (5.5) 8.0 (3.4) 21.8 (6.3) 13.6 (4.7) 54.8 (24.6) 56.0 (25.0) 64.6 (28.1) 62.7 (27.4)

Free light fraction – significant effects: altitude, soil horizon, RS vs NRS; SE = 1.0, LSD = 3.0. Occluded light fraction – significant effects: altitude, soil horizon, RS vs NRS; SE = 0.3, LSD = 0.9. Heavy fraction – no significant effect. 5

AB horizon % of TOC

g C kg−1 soil

% of TOC

20.3 15.8 21.7 14.1 19.4 19.0 17.4 14.6 59.9 64.8 60.6 71.0

10.2 4.0 13.5 5.1 8.4 (3.5) 3.3 (1.9) 9.9 (3.8) 6.8 (3.1) 44.2 (20.7) 38.0 (18.4) 52.5 (23.8) 52.5 (23.8)

17.6 9.0 14.3 8.3 14.2 7.7 14.0 9.7 67.5 82.6 71.2 81.7

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Table 3 Amounts of C of fulvic and humic extracts (FE–C and HE–C, respectively), and non-extractable organic matter (NEOC) in the heavy fraction of rhizosphere soil (RS) and non-rhizosphere soil (NRS) of European beech (Fagus sylvatica L.) forests at 800 and 1000 m above sea level (a.s.l.), central Apennines, Italy. For each fraction, the standard error (SE) and the least significant difference (LSD) were calculated on the transformed (in parentheses) means (n = 3) when the effects from the ANOVA were significant. Horizon

FE-C

HE-C

800 m a.s.l. RS

1000 m a.s.l.

NEOC

800 m a.s.l.

1000 m a.s.l.

800 m a.s.l.

1000 m a.s.l.

NRS

RS

NRS

RS

NRS

RS

NRS

RS

NRS

RS

NRS

7.7 (11.48) 5.6 (7.14)

8.1 (14.91) 7.1 (12.38)

8.2 (12.92) 6.6 (8.64)

12.5 (0.97) 20.6 (1.11)

20.3 (1.16) 16.7 (1.02)

26.0 (1.27) 17.1 (1.06)

22.5 (1.19) 23.9 (1.18)

34.7 (2.53) 18.3 (1.98)

28.0 (2.28) 15.7 (1.89)

30.5 (2.43) 28.3 (2.34)

32.0 (2.49) 22.0 (2.08)

g C kg−1 soil A AB

7.6 (11.75) 5.3 (7.11)

FE–C – significant effect: soil horizon; SE = 1.30, LSD = 3.91. HE–C – no significant effect. NEOC – no significant effect.

the FE–C, which generally were in greater proportion in the Hf of the A than in that of the AB horizon, likely because of the higher flux of fresh plant-debris in the A horizon (Guimarães et al., 2013). Because of the key role played by fulvic acids on nutrients mobilization, the decline of FE-C with depth could also explain the lower available P content in the AB than in the A horizon (Yang et al., 2013). Hence, while the OC of the light fractions responded to a small difference in altitude (corresponding to 1 °C of MAAT), the stabilised organic matter forming Hf showed a low sensitivity to external factors (Baisden et al., 2002). These results are supported by the findings of several authors. Huang et al. (2011) reported that while OC concentration in the light fractions increased after afforestation of grasslands, no changes occurred for Hf (ρ > 1.6 g cm−3). Grüneberg et al. (2013) observed variation of OC content in the light fractions but not in Hf (ρ > 1.7 g cm−3) of beech forest soils differently managed, and concluded that the lack of changes of OC in the denser fraction was mainly controlled by the formation of organo–mineral complexes. Díaz-Pinés et al. (2011), in a study carried out on Scots pine and Pyrenean oak forest soils, found that despite differences in soil TOC content among vegetation types, the OC associated to minerals remained stable, both across vegetation and with increasing soil depth.

higher pH values, and the larger TOC and total N contents at the higher altitude (De Feudis et al., 2017b). Hence, plants may release larger amounts of root exudates to overcome the more severe soil conditions at higher altitude so triggering the microbial community activity and, consequently, enhancing the nutrient cycling (Celi et al., 2013; De Feudis et al., 2016) in the RS. 4.2. Organic C distribution in the density fractions of the RS and the NRS at 800 and 1000 m of altitude In the two soil horizons at both altitudes, the most abundant OC pool was in the Hf, indicating that the association between organics and minerals provides a great contribution to OC sequestration in forest soils (e.g., John et al., 2005; Schrumpf et al., 2013). This condition has been attributed to the stabilization provided by both the physicochemical interactions with soil minerals (through cation bridges and, under acid to neutral pH conditions, positively charged colloids) and the chemical recalcitrance of the organic molecules (Grünewald et al., 2006; von Lützow et al., 2008). The enrichment of recalcitrant moieties in the Hf of calcareous soils have been also reported by Rovira and Vallejo (2003) and Grünewald et al. (2006), who found a reduction of easily decomposable organic compounds and even lignin in the denser fractions compared to the lighter ones. Our results revealed different effects of altitude, soil horizons, and rhizosphere on the OC content of the fractions obtained by density fractionation. In fact, we found a general influence of these factors on FLf and OLf, whereas no effect was detected for Hf. With regard to both light fractions, the greater OC concentration in the RS than in the NRS has been attributed to the high and frequent supply of fresh particulate organic matter occurring in the rhizosphere due to root activity and turnover (Angst et al., 2016). The larger presence of light fraction in the RS than in the NRS could also explain the greater amount of available P in the soil close to the roots. Indeed, according to O'Hara et al. (2006), P released by the degradation of the light organic matter fraction can provide an important contribution to the biological P cycling and supply, and plays a significant role in the tree nutrition. Further, the increasing OC content in the FLf of the RS with altitude may reflect an augmented rhizodeposition process due to the lower temperature and to the reduced soil development at the higher sites (De Feudis et al., 2017a). Unlike the FLf, the altitude affected OC content only in the OLf from the NRS. We were unable to provide an explanation for this as OLf is controlled by formation/disruption of soil structure, which depends on complex interactions among climatic and edaphic factors and not only from litter or root inputs (Barto et al., 2010; Bronick and Lal, 2005). Conversely to the two light fractions, altitude, soil horizons and rhizosphere did not affect the OC content of Hf and its distribution in HE, FE, and NEOC. The only significant different distribution regarded

4.3. Δ14C and δ13C in the density fractions of the RS and the NRS at 800 and 1000 m of altitude The positive Δ14C values of light and heavy fractions did not significantly differ between the RS and the NRS, horizons, and altitudes (Fig. 3). The similar 14C signature and the presence of “bomb 14C” in all the density fractions from the RS and the NRS of both A and AB horizons suggested that SOM underwent degradation and rapid resupplying because of an efficient incorporation and mixing of litter floor into the mineral horizons (Hall et al., 2015; Schnecker et al., 2016), followed by occlusion into aggregates and formation of organomineral complexes (Herold et al., 2014; McFarlane et al., 2013). Further, the lack of statistically different Δ14C values between the fractions at 800 and 1000 m might indicate that 1 °C of MAAT difference is not sufficient to induce marked changes in the 14C signature of SOM. Similar results were reported by Schnecker et al. (2016) after seven year of forest soil warming experiment where the soil was kept at 4 °C higher temperature than the control. Indeed, these authors found no difference in the Δ14C signatures of bulk and three density fractions between heated and control soils. Although the radiocarbon data did not show differences among the density fractions, in agreement with several studies on organic matter dynamic in forest soils (e.g., Heckman et al., 2014; McFarlane et al., 2013; Schrumpf et al., 2013), a greater content of 13C was found in Hf than in the light fractions (P < 0.05). Since OLf is generally considered more protected than FLf (von Lützow et al., 2007), the lack of difference 6

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Fig. 2. δ13C versus Δ14C of organic C in the free light fraction, occluded light fraction, and heavy fraction of rhizosphere soil (RS) and non-rhizosphere soil (NRS) of European beech (Fagus sylvatica L.) forest at 800 and 1000 m above sea level, central Apennines, Italy. For each fraction, the standard error (SE) and the least significant difference (LSD) were calculated on the original or transformed means (n = 3) when the effects from the ANOVA were significant. δ13C free light fraction – significant effect: altitude; SE = 0.009, LSD = 0.028. δ13C occluded light fraction – significant effect: altitude; SE = 0.013, LSD = 0.040. δ13C heavy fraction – significant effect: altitude; SE = 0.047, LSD = 0.016. Δ14C –no significant effects for the free light fraction, occluded light fraction and heavy fraction.

in the δ13C between the two light fractions was possibly due to the active relationship that occurs between FLf and OLf due to the disruption and reorganization of aggregates during the seasonal climatic changes (Heckman et al., 2014). With regard to Hf, the greater δ13C of the mineral-bound organic matter was due to its high degree of degradation and the preferential use of the lighter C isotope during the microbial C mineralization (Gunina and Kuzyakov, 2014). This latter fact might also explain the similar δ13C values found for all the fractions of the RS and the NRS. The low δ13C of the fresh rhizodeposits released in the RS would be counterbalanced by the greater microbial organic matter degradation in the RS compared to the NRS, where the main source of organics is the degradation of the litter (Sayer, 2006). Further, the similar 13C content in each fraction of the A and AB horizons showed weak or negligible δ13C enrichment with depth, which was

attributed mainly to soil mixing produced by soil faunal disturbance and slope movements (Agnelli et al., 2016; Hall et al., 2015; Vesterdal et al., 2008). The absence of δ13C depth-trend in the upper soil horizons of forest soils has been reported by McCorkle et al. (2016), who found a marked increase with depth of the 13C content in the density fractions only when the whole soil profile was considered. Concerning with the altitude, a greater 13C enrichment occurred at 1000 m than at 800 m in the FLf of the A horizon and in all the three fractions of the AB horizon. As a strong positive correlation between foliar and soil δ13C occurs (Peri et al., 2012), the greater δ13C of FLf at 1000 m might be due to the lower MAAT and air pressure, which induce a weaker photosynthetic C isotope discrimination and a consequent higher accumulation of 13C in plant tissues (Cernusak et al., 2013; Hultine and Marshall, 2000; Zhou et al., 2011). Since the organics of 7

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OLf and Hf are generally more degraded than those of FLf (Schrumpf et al., 2013), their similar signatures at 800 and 1000 m found in the A horizon was attributed to a combined effect of the 13C concentration in the organic debris and the different C-cycling intensity at the two elevations, which was enhanced at lower altitude because of the warmer temperature (Powers and Schlesinger, 2002; Zimmermann et al., 2012). The greater δ13C of OLf and Hf from the AB horizon of the soil at 1000 m than those at 800 m suggested that the counterbalancing effect played by the microbial activity was lesser effective in this sub-superficial horizon. Likely, in the AB horizon the difference of 1 °C air temperature is not sufficient to affect the intensity of SOM cycling to trigger a differential effect between the two altitudes, as it occurred for the uppermost horizon. Hence, at 1000 m altitude, the higher δ13C of FLf, OLf, and Hf from the AB horizon should be due more to the initial 13C content of the organic residues than to a different fractionation effect occurring at the two altitudes.

Surface and subsurface microbial biomass, community structure and metabolic activity as a function of soil depth and season. Appl. Soil Ecol. 20, 171–181. Box, G.E.P., Cox, D.R., 1964. Analysis of transformations. J. R. Stat. Soc. Ser. B Stat Methodol. 26, 211–252. Bronick, C.J., Lal, R., 2005. Soil structure and management: a review. Geoderma 124, 3–22. Brzostek, E.R., Greco, A., Drake, J.E., Finzi, A.C., 2013. Root carbon inputs to the rhizosphere stimulate extracellular enzyme activity and increase nitrogen availability in temperate forest soils. Biogeochemistry 115, 65–76. Bu, X., Ruan, H., Wang, L., Ma, W., Ding, W., Yu, X., 2012. Soil organic matter in density fractions as related to vegetation changes along an altitude gradient in the Wuyi Mountains, southeastern China. Appl. Soil Ecol. 52, 42–47. Buée, M., Vairelles, D., Garbaye, J., 2005. Year-round monitoring of diversity and potential metabolic activity of the ectomycorrhizal community in a beech (Fagus sylvatica) forest subjected to two thinning regimes. Mycorrhiza 15, 235–245. Calvaruso, C., Collignon, C., Kies, A., Turpault, M.-P., 2014. Seasonal evolution of the rhizosphere effect on major and trace elements in soil solutions of Norway Spruce (Picea abies Karst) and beech (Fagus sylvatica) in an acidic forest soil. Open J. Soil Sci. 4, 323–336. Celi, L., Cerli, C., Turner, B.L., Santoni, S., Bonifacio, E., 2013. Biogeochemical cycling of soil phosphorus during natural revegetation of Pinus sylvestris on disused sand quarries in Northwestern Russia. Plant Soil 367, 122–134. Cernusak, L.A., Ubierna, N., Winter, K., Holtum, J.A.M., Marshall, J.D., Farquhar, G.D., 2013. Environmental and physiological determinants of carbon isotope discrimination in terrestrial plants. New Phytol. 200, 950–965. Chen, L.-F., He, Z.-B., Du, J., Yang, J.-J., Zhu, X., 2016. Patterns and environmental controls of soil organic carbon and total nitrogen in alpine ecosystems of northwestern China. Catena 137, 37–43. Cocco, S., Agnelli, A., Gobran, G.R., Corti, G., 2013. Changes induced by the roots of Erica arborea L. to create a suitable environment in a soil developed from alkaline and finetexture marine sediments. Plant Soil 368, 297–313. Colin-Belgrand, M., Dambrine, E., Bienaimé, S., Nys, C., Turpault, M.-P., 2003. Influence of tree roots on nitrogen mineralization. Scand. J. Forest Res. 18, 260–268. De Feudis, M., Cardelli, V., Massaccesi, L., Bol, R., Willbold, S., Cocco, S., Corti, G., Agnelli, A., 2016. Effect of beech (Fagus sylvatica L.) rhizosphere on phosphorous availability in soils at different altitudes (central Italy). Geoderma 276, 53–63. De Feudis, M., Cardelli, V., Massaccesi, L., Hofmann, D., Berns, A.E., Bol, R., Cocco, S., Corti, G., Agnelli, A., 2017a. Altitude affects the quality of the water-extractable organic matter (WEOM) from rhizosphere and bulk soil in European beech forests. Geoderma 302, 6–13. De Feudis, M., Cardelli, V., Massaccesi, L., Lagomarsino, A., Fornasier, F., Westphalen, D.J., Cocco, S., Corti, G., Agnelli, A., 2017b. Influence of altitude on biochemical properties of European beech (Fagus sylvatica L.) forest soils. Forests 8, 213. Díaz-Pinés, E., Rubio, A., Van Miegroet, H., Montes, F., Benito, M., 2011. Does tree species composition control soil organic carbon pools in Mediterranean mountain forests? For. Ecol. Manag. 262, 1895–1904. Dungait, J.A.J., Hopkins, D.W., Gregory, A.S., Whitmore, A.P., 2012. Soil organic matter turnover is governed by accessibility not recalcitrance. Glob. Change Biol. 18, 1781–1796. Fernández-Romero, M.L., Lozano-García, B., Parras-Alcántara, L., 2014. Topography and land use change effects on the soil organic carbon stock of forest soils in Mediterranean natural areas. Agric. Ecosyst. Environ. 195, 1–9. Garten Jr., C.T., Hanson, P.J., 2006. Measured forest soil C stocks and estimated turnover times along an elevation gradient. Geoderma 136, 342–352. Grüneberg, E., Schöning, I., Hessenmöller, D., Schulze, E.-D., Weisser, W.W., 2013. Organic layer and clay content control soil organic carbon stocks in density fractions of differently managed German beech forests. For. Ecol. Manag. 303, 1–10. Grünewald, G., Kaiser, K., Jahn, R., Guggenberger, G., 2006. Organic matter stabilization in young calcareous soils as revealed by density fractionation and analysis of ligninderived constituents. Org. Geochem. 37, 1573–1589. Guimarães, D.V., Gonzaga, M.I.S., da Silva, T.O., da Silva, T.L., da Silva Dias, N., Matias, M.I.S., 2013. Soil organic matter pools and carbon fractions in soil under different land uses. Soil Till. Res. 126, 177–182. Gunina, A., Kuzyakov, Y., 2014. Pathways of litter C by formation of aggregates and SOM density fractions: implications from 13C natural abundance. Soil Biol. Biochem. 71, 95–104. Gunina, A., Kuzyakov, Y., 2015. Sugars in soil and sweets for microorganisms: review of origin, content, composition and fate. Soil Biol. Biochem. 90, 87–100. Gutiérrez-Girón, A., Díaz-Pinés, E., Rubio, A., Gavilán, R.G., 2015. Both altitude and vegetation affect temperature sensitivity of soil organic matter decomposition in Mediterranean high mountain soils. Geoderma 237-238, 1–8. Hall, S.J., McNicol, G., Natake, T., Silver, W.L., 2015. Large fluxes and rapid turnover of mineral-associated carbon across topographic gradients in a humid tropical forest: insights from paired 14C analysis. Biogeosciences 12, 2471–2487. Han, L., Sun, K., Jin, J., Xing, B., 2016. Some concepts of soil organic carbon characteristics and mineral interaction from a review of literature. Soil Biol. Biochem. 94, 107–121. Heckman, K., Throckmorton, H., Clingensmith, C., González Vila, F.J., Horwath, W.R., Knicker, H., Rasmussen, C., 2014. Factors affecting the molecular structure and mean residence time of occluded organics in a lithosequence of soils under ponderosa pine. Soil Biol. Biochem. 77, 1–11. Helal, H.M., Sauerbeck, D., 1989. Carbon turnover in the rhizosphere. J. Plant Nutr. Soil Sci. 152, 211–216. Herold, N., Schoning, I., Michalzik, B., Trumbore, S., Schrumpf, M., 2014. Controls on soil carbon storage and turnover in German landscapes. Biogeochemistry 119, 435–451. Huang, Z., Davis, M.R., Condron, L.M., Clinton, P.W., 2011. Soil carbon pools, plant

5. Conclusions This study showed that the organic carbon pools obtained by density fractionation have a different sensitivity to the roots activity (rhizosphere effect) and the altitude (and hence, temperature). The light fractions (FLf and OLf) were markedly affected by both rhizosphere and altitude, whereas the similar amount and quality (distribution of HE–C, FE–C, and NEOC) of OC in the Hf between the RS and the NRS and between the two altitudes indicated low sensitivity to rhizosphere and climate effects. The relative low sensitivity of the OC comprising Hf was attributed to its advanced degradation stage, as also suggested by δ13C values, and the close association to mineral particles. The similar positive Δ14C values observed in all soil density fractions, indicating the predominance of recent C in the organic pools, were probably due to i) an efficient incorporation of the organic matter derived from decayed litter floor into the mineral horizons at all sites, ii) the incapability of 1 °C of MAAT difference between the two altitudes to differently impact on 14C fractionation, and iii) an unavoidable spatial variation among sites. Unlike the Δ14C, the greater δ13C values found in the three SOM pools at 1000 than at 800 m have been mainly attributed to the impact of the lower MAAT and air pressure occurring at the higher altitude, which may promote the accumulation of 13C in the plant tissues and, in turn, into the soil. In summary, in the studied conditions, our results indicated that 1 °C warmer MAAT reduces the size of the OC pool associated to the light fractions of the rhizosphere, but it is ineffective on the stabilised organic pool, which represents the great part of SOM. Supplementary data to this article can be found online at https:// doi.org/10.1016/j.catena.2019.104091. References Agnelli, A., Bol, R., Trumbore, S.E., Dixon, L., Cocco, S., Corti, G., 2014. Carbon and nitrogen in soil and vine roots in harrowed and grass-covered vineyards. Agric. Ecosyst. Environ. 193, 70–82. Agnelli, A., Massaccesi, L., De Feudis, M., Cocco, S., Courchesne, F., Corti, G., 2016. Holm oak (Quercus ilex L.) rhizosphere affects limestone-derived soil under a multi-centennial forest. Plant Soil 400, 297–314. Angst, G., Kögel-Knabner, I., Kirfel, K., Hertel, D., Mueller, C.W., 2016. Spatial distribution and chemical composition of soil organic matter fractions in rhizosphere and non-rhizosphere soil under European beech (Fagus sylvatica L.). Geoderma 264, 179–187. Baisden, W.T., Amundson, R., Cook, A.C., Brenner, D.L., 2002. Turnover and storage of C and N in five density fractions from California annual grassland surface soils. Glob. Biogeochem. Cycles 16, 1117. Barto, E.K., Alt, F., Oelmann, Y., Wilcke, W., Rillig, M.C., 2010. Contributions of biotic and abiotic factors to soil aggregation across a land use gradient. Soil Biol. Biochem. 42, 2316–2324. Bispo, A., Andersen, L., Angers, D.A., Bernoux, M., Brossard, M., Cécillon, L., Comans, R.N.J., Harmsen, J., Jonassen, K., Lamé, F., Lhuillery, C., Maly, S., Martin, E., Mcelnea, A.E., Sakai, H., Watabe, Y., Eglin, T.K., 2017. Accounting for carbon stocks in soils and measuring GHGs emission fluxes from soils: do we have the necessary standards? Front. Environ. Sci. 5, 41. Blume, E., Bischoff, M., Reichert, J.M., Moorman, T., Konopka, A., Turco, R.F., 2002.

8

Catena 181 (2019) 104091

M. De Feudis, et al. biomarkers and mean carbon residence time after afforestation of grassland with three tree species. Soil Biol. Biochem. 43, 1341–1349. Hultine, K.R., Marshall, J.D., 2000. Altitude trends in conifer leaf morphology and stable carbon isotope composition. Oecologia 123, 32–40. Hutsch, B.W., Augustin, J., Merbach, W., 2002. Plant rhizodeposition – an important source for carbon turnover in soils. J. Plant Nutr. Soil Sci. 165, 397–407. IPCC, Stocker, T.F., Qin, D., Plattner, G.-K., Tignor, M., Allen, S.K., Boschung, J., Nauels, A., Xia, Y., Bex, V., Midgley, P.M., 2013. Climate Change 2013: The Physical Science Basis. Contribution of Working Group I to the Fifth Assessment Report of the Intergovernmental Panel on Climate Change. Cambridge University Press, Cambridge, United Kingdom and New York, NY, USA. Jackson, R.B., Lajtha, K., Crow, S.E., Hugelius, G., Kramer, M.G., Piñeiro, G., 2017. The ecology of soil carbon: pools, vulnerabilities, and biotic and abiotic controls. Annu. Rev. Ecol. Evol. Syst. 48, 419–445. Ji, H., Ding, Y., Liu, X., Li, L., Zhang, D., Li, Z., Sun, J., Lashari, M.S., Joseph, S., Meng, Y., Kuzyakov, Y., Pan, G., 2015. Root-derived short-chain suberin diacids from rice and rape seed in a paddy soil under rice cultivar treatments. PLoS One 10, e0127474. John, B., Yamashita, T., Ludwig, B., Flessa, H., 2005. Storage of organic carbon in aggregate and density fractions of silty soils under different types of land use. Geoderma 128, 63−79. Jones, D.L., Nguyen, C., Finlay, R.D., 2009. Carbon flow in the rhizosphere: carbon trading at the soil–root interface. Plant Soil 321, 5–33. Joshi, A.B., Vann, D.R., Johnson, A.H., Miller, E.K., 2003. Nitrogen availability and forest productivity along a climosequence on Whiteface Mountain, New York. Can. J. For. Res. 33, 1880–1891. Kaiser, M., Zederer, D.P., Ellerbrock, R.H., Sommer, M., Ludwig, B., 2016. Effects of mineral characteristics on content, and stability of organic matter fractions separated from seven forest topsoils of different pedogenesis. Geoderma 263, 1–7. Kuzyakov, Y., 2010. Priming effects: interactions between living and dead organic matter. Soil Biol. Biochem. 42, 1363–1371. Kuzyakov, Y., Hill, P.W., Jones, D.L., 2007. Root exudate components change litter decomposition in a simulated rhizosphere depending on temperature. Plant Soil 290, 293–305. Lambers, H., Mougel, C., Jaillard, B., Hinsinger, P., 2009. Plant-microbe-soil interactions in the rhizosphere: an evolutionary perspective. Plant Soil 321, 83–115. Lehmann, J., Kleber, M., 2015. The contentious nature of soil organic matter. Nature 528, 60–68. Levin, I., Kromer, B., 2004. The tropospheric (CO2)-C-14 level in mid-latitudes of the Northern Hemisphere (1959–2003). Radiocarbon 46, 1261–1272. von Lützow, M.V., Kögel-Knabner, I., Ekschmitt, K., Flessa, H., Guggenberger, G., Matzner, E., Marschner, B., 2007. SOM fractionation methods: relevance to functional pools and to stabilization mechanisms. Soil Biol. Biochem. 39, 2183–2207. von Lützow, M., Kögel-Knabner, I., Ludwig, B., Matzner, E., Flessa, H., Ekschmitt, K., Guggenberger, G., Marschner, B., Kalbitz, K., 2008. Stabilization mechanisms of organic matter in four temperate soils: development and application of a conceptual model. J. Plant Nutr. Soil Sci. 171, 111–124. Massaccesi, L., Benucci, G.M.N., Gigliotti, G., Cocco, S., Corti, G., Agnelli, A., 2015. Rhizosphere effect of three plant species of environment under periglacial conditions (Majella massif, Central Italy). Soil Biol. Biochem. 89, 184–195. McCorkle, E.P., Berhe, A.A., Hunsaker, C.T., Johnson, D.W., McFarlane, K.J., Fogel, M.L., Hart, S.C., 2016. Tracing the source of soil organic matter eroded from temperate forest catchments using carbon and nitrogen isotopes. Chem. Geol. 445, 172–184. McFarlane, K.J., Torn, M.S., Hanson, P.J., Porras, R.C., Swanston, C.W., Callaham, M.A., Guilderson, T.P., 2013. Comparison of soil organic matter dynamics at five temperate deciduous forests with physical fractionation and radiocarbon measurements. Biogeochemistry 112, 457–476. Nardi, S., Concheri, G., Pizzeghello, D., Sturaro, A., Rella, R., Parvoli, G., 2000. Soil organic matter mobilization by root exudates. Chemosphere 41, 653–658. O'Hara, C.P., Bauhus, J., Smethurst, P.J., 2006. Role of light fraction soil organic matter in the phosphorus nutrition of Eucalyptus globulus seedlings. Plant Soil 280, 127–134. Olsen, S.R., Cole, C.V., Watanabe, F.S., Dean, L.A., 1954. Estimation of Available Phosphorus in Soils by Extraction with Sodium Bicarbonate. US Department of Agriculture, Washington Circ. 939. Otero, X.L., González-Guzman, A., Souza-Junior, V.S., Pérez-Alberti, A., Macías, F., 2015. Soil processes and nutrient bioavailability in the rhizosphere of Bolax gummifera in a subantarctic environment (Martial Mountains, Ushuaia - Argentina). Catena 133, 432–440. Peri, P.L., Ladd, B., Pepper, D.A., Bonser, S.P., Laffan, S.W., Amelung, W., 2012. Carbon (δ13C) and nitrogen (δ15N) stable isotope composition in plant and soil in Southern Patagonia's native forests. Glob. Chang. Biol. 18, 311–321. Powers, J.S., Schlesinger, W.H., 2002. Geographic and vertical patterns of stable carbon isotopes in tropical rain forest soils of Costa Rica. Geoderma 109, 141–160.

R Core Team, 2014. R: A Language and Environment 538 for Statistical Computing. R Foundation for Statistical Computing, Vienna, Austria URL. http://www.R-project. org/. Rasse, D.P., Rumpel, C., Dignac, M.-F., 2005. Is soil carbon mostly root carbon? Mechanisms for a specific stabilisation. Plant Soil 269, 341–356. Rech, J.A., Reeves, R.W., Hendricks, D.M., 2001. The influence of slope aspect on soil weathering processes in the Springerville volcanic field, Arizona. Catena 43, 49–62. Rovira, P., Vallejo, V.R., 2003. Physical protection and biochemical quality of organic matter in mediterranean calcareous forest soils: a density fractionation approach. Soil Biol. Biochem. 35, 245–261. Ruehr, N.K., Buchmann, N., 2010. Soil respiration fluxes in a temperate mixed forest: seasonality and temperature sensitivities differ among microbial and root–rhizosphere respiration. Tree Physiol. 30, 165–176. Sayer, E.J., 2006. Using experimental manipulation to assess the roles of leaf litter in the functioning of forest ecosystems. Biol. Rev. 81, 1–31. Schnecker, J., Borken, W., Schindlbacher, A., Wanek, W., 2016. Little effects on soil organic matter chemistry of density fractions after seven years of forest soil warming. Soil Biol. Biochem. 103, 300–307. Schrumpf, M., Kaiser, K., Guggenberger, G., Persson, T., Kӧgel-Knabner, I., Schulze, E.-D., 2013. Storage and stability of organic carbon in soils as related to depth, occlusion within aggregates, and attachment to minerals. Biogeosciences 10, 1675−1691. Siles, J.A., Cajthaml, T., Filipová, A., Minerbi, S., Margesin, R., 2017. Altitudinal, seasonal and interannual shifts in microbial communities and chemical composition of soil organic matter in Alpine forest soils. Soil Biol. Biochem. 112, 1–13. Soethe, N., Lehmann, J., Engels, C., 2007. Carbon and nutrient stocks in roots of forests at different altitudes in the Ecuadorian Andes. J. Trop. Ecol. 23, 319–328. Soil Survey Staff, 2014. Keys to Soil Taxonomy, 12th ed. USDA–Natural Resources Conservation Service, Washington. Sokolova, T.A., 2015. Specificity of soil properties in the rhizosphere: analysis of literature data. Eurasian Soil Sci 48, 968–980. Song, B.-L., Yan, M.-J., Hou, H., Guan, J.-H., Shi, W.-Y., Li, G.-Q., Du, S., 2016. Distribution of soil carbon and nitrogen in two typical forests in the semiarid region of the loess plateau, China. Catena 143, 159–166. Steinhof, A., Adamiec, G., Gleixner, G., van Klinken, G.J., Wagner, T., 2004. The new 14C analysis laboratory in Jena, Germany. Radiocarbon 46, 51–58. Strickland, M.S., Wickings, K., Bradford, M.A., 2012. The fate of glucose, a low molecular weight compound of root exudates, in the belowground foodweb of forests and pastures. Soil Biol. Biochem. 49, 23–29. Stuiver, M., Polach, H., 1977. Reporting of 14C data. Radiocarbon 19, 355–363. Tang, X.Y., Bernard, L., Brauman, A., Daufresne, T., Deleporte, P., Desclaux, D., Souche, G., Placella, S.A., Hinsinger, P., 2014. Increase in microbial biomass and phosphorus availability in the rhizosphere of intercropped cereal and legumes under field conditions. Soil Biol. Biochem. 75, 86–93. Tefs, C., Gleixner, G., 2012. Importance of root derived carbon for soil organic matter storage in a temperate old-growth beech forest - evidence from C, N and 14C content. For. Ecol. Manag. 263, 131–137. Toal, M.E., Yeomans, C., Killham, K., Meharg, A.A., 2000. A review of rhizosphere carbon flow modeling. Plant Soil 222, 263–281. Trumbore, S., 2009. Radiocarbon and soil carbon dynamics. Annu. Rev. Earth Planet. Sci. 37, 47–66. Tsui, C.-C., Chen, Z.-S., Hsieh, C.-F., 2004. Relationships between soil properties and slope position in a lowland rain forest of southern Taiwan. Geoderma 123, 131–142. Vesterdal, L., Schmidt, I.K., Callesen, I., Nilsson, L.O., Gundersen, P., 2008. Carbon and nitrogen in forest floor and mineral soil under six common European tree species. For. Ecol. Manag. 255, 35–48. Werner, R.A., Brand, W.A., 2001. Referencing strategies and techniques in stable isotope ratio analysis. Rapid Commun. Mass Spectrom. 15, 501–519. Xu, M., Li, X., Cai, X., Li, X., Christie, P., Zhang, J., 2014. Soil microbial community structure and activity along a montane elevational gradient on the Tibetan plateau. Eur. J. Soil Biol. 64, 6–14. Yang, S., Zhang, Z., Cong, L., Wang, X., Shi, S., 2013. Effect of fulvic acid on the phosphorus availability in acid soil. J. Soil Sci. Plant Nutr. 13 (3), 526–533. Yin, H., Xiao, J., Li, Y., Chen, Z., Cheng, X., Zhao, C., Liu, Q., 2013. Warming effects on root morphological and physiological traits: the potential consequences on soil C dynamics as altered root exudation. Agric. For. Meteorol. 180, 287–296. Zhou, Y., Fan, J., Zhang, W., Harris, W., Zhong, H., Hu, Z., Song, L., 2011. Factors influencing altitudinal patterns of C3 plant foliar carbon isotope composition of grasslands on the Qinghai-Tibet plateau, China. Alpine Bot 121, 79–90. Zimmermann, M., Leifeld, J., Conen, F., Bird, M.I., Meir, P., 2012. Can composition and physical protection of soil organic matter explain soil respiration temperature sensitivity? Biogeochemistry 107, 1–14.

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