Phosphorus dynamics during early soil development in extreme environment

Abstract. At the early stages of pedogenesis, the dynamics of phosphorus (P) in soils are controlled by microbial communities, the physicochemical properties of the soil and the environmental conditions. While various microorganisms involved in carrying out biogeochemical processes have been identified, little is known about the actual contribution of microbial processes, such as organic P hydrolysis and microbial P turnover, to P cycling. We thus focused on processes driven by microbes and how they affect the size and cycling of organic and inorganic soil P pools along a soil chronosequence in the Chamser Kangri glacier forefield (Western Himalayas). The rapid retreat of the glacier allowed us to study the early stages of soil formation under cold arid climate. Biological P transformations were studied with the help of the isotopic composition of oxygen (O) in phosphate (δ18OP) coupled to sequential P fractionation performed on soil samples from four sites of different age spanning 0 to 100–150 years. The mineral P, i.e. 1M HCl-extractable P, represented still 95 % of the total P stock after approximately 100 years of soil development. Its isotopic composition was similar to the parent material also at the most developed site. Primary phosphate minerals, therefore, mostly composed this pool. The δ18OP of the available P and the P bound to Fe and Al oxides instead differed from that of the parent material, suggesting that these pools underwent biological turnover. The isotopic composition of O in of the available P was mostly controlled by the microbial P, suggesting fast exchanges occurred between these two pools possibly fostered by repeated freezing-thawing and drying-rewetting cycles. The release of P from organic P become increasingly important with soil age, constituting one third of the P flux to available P at the oldest site. Accordingly, the lighter isotopic composition of the P bound to Fe and Al oxides at the oldest site indicated that this pool contained phosphate released by organic P mineralization. Compared to previous studies on early pedogenesis under alpine or cold climate, our findings suggest a much slower decrease of the P-bearing primary minerals during the first 100 years of soil development under extreme condition. However, they provide evidence that, by driving short-term P dynamics, microbes play an important role in controlling the redistribution of primary P into inorganic and organic soil P pools.


hypothesized that Po mineralization was driven by the microbial need of C. The microbial P pool could also contribute to available P via the turnover of microbial cells determined by biological (e.g. grazing) or abiotic (e.g. freezing/thawing) factors (Oberson and Joner, 2005). However, little is known about the actual contribution of microbial processes to P 55 availability especially during the early pedogenesis (Schulz et al., 2013).
The study of isotopic composition of oxygen (O) in phosphate (expressed here as δ 18 OP in the delta per mil notation) is a relatively new approach to trace P biogeochemical transformation processes overcoming limitations of approaches that depend on radioisotopes . The P-O bond in phosphate is resistant to inorganic hydrolysis under natural temperature and pressure (O'Neil et al., 2003), therefore negligible O atoms exchange occurs between 60 phosphate and water without biological mediation (Tudge, 1960;Blake et al., 2001). Two main enzyme-mediated processes can alter the δ 18 OP. First, intracellular metabolism of Pi (reversible conversion of pyrophosphate to two phosphate ions mediated by inorganic pyrophosphatase) causes a complete and fast exchange of O atoms between phosphate and water molecules (Cohn, 1958;Blake et al., 2005). This results in a temperature-dependent equilibrium between O in the phosphate molecule and O of the intracellular water (Fricke et al., 1998). The δ 18 OP at equilibrium can 65 be predicted, based on the measured isotopic ratio of oxygen in water (δ 18 OW) and temperature (Longinelli and Nuti 1973;Kolodny et al. 1983;Chang and Blake 2015). Therefore, the difference between the calculated equilibrium value and the measured δ 18 OP in an environmental sample can provide insights into the extent to which Pi has been cycled by the soil microorganisms through intracellular metabolic reactions (Davies et al., 2014). The second process is the hydrolysis of Po compounds by phosphohydrolases-mediated reactions (Liang and Blake, 2006). During Po hydrolysis, 70 P-O bonds are cleaved and one to two O atoms in the phosphate are replaced with O from the water with a specific fractionation factor (ԑ). The ԑ of main phosphohydrolase enzymes is often negative, e.g. alkaline phosphatases have an ԑ of -30‰, while acid phosphatases of -10‰, resulting in depleted δ 18 OP in the released phosphate (Liang and Blake, 2009;Von Sperber et al., 2015).
Glacier forefields are ideal sites to study the initial steps of soil formation as neighboring sites represent a soil 75 chronosequence of different soil developmental stages. The general assumption is the space for time substitution, implying that each site along a glacier forefield chronosequence was subject to the same initial conditions and followed the same sequence of changes. We studied the early stages (0 to approximately 100-150 years) of soil development in a glacier forefield located in the Western Himalayas. Under these conditions, microorganisms may be subjected to drought, intense solar radiation (Blumthaler et al., 1997), and high temperature fluctuations 80 Rehakova et al., 2011). Direct forefield observations on the role of microorganisms in P cycling using O isotopes in phosphate are rare. A study in the forefield of the Damma glacier in the Swiss Alps revealed that at the early stages of https://doi.org/10.5194/soil-2021-65 Preprint. Discussion started: 21 July 2021 c Author(s) 2021. CC BY 4.0 License. soil development (<150 years), the microbial P was the main contributor to available P, which carried the isotopic equilibrium signature (Tamburini et al., 2012). More recently, it has been shown that under P-limiting conditions, instead, available P might show non-equilibrium lower δ 18 OP as a result of tight Po recycling through phosphohydrolase-85 mediated reactions (Pistocchi et al., 2020).
In this study, we applied a sequential fractionation method to identify loosely to strongly bound Pi and Po pools and analyzed the δ 18 OP values in these pools  to quantify biological P contribution to available P during early soil formation. Given that sequentially-extracted P pools differ in their availability and turnover time (Helfenstein et al., 2020), we hypothesized 1) that the influence of biological cycling will increase with increasing soil 90 age; 2) that the main contribution to available P, at the early development stages, will derive from microbial P rather than Po mineralization; and 3) that the contribution to available P from Po mineralization will increase with soil age as observed in an older chronosequence (Roberts et al., 2015). We will therefore: i) quantify sequentially extracted P pools and potential enzymatic activities linked to C, nitrogen (N) and P mineralization in soils along the forefield chronosequence, ii) determine the δ 18 OP values of these pools, including Po and compare them to the calculated isotopic 95 equilibrium and expected isotopic values from Po hydrolysis, in order to assess their relative contribution to available P.

100
The glacier forefield of Chamser Kangri was chosen as it has been studied within a long-term interdisciplinary research development project (Dolezal et al., 2016). Since 2008, the local environmental conditions have been monitored, changes in vegetation and relationships between vascular plants and soil microbial communities studied (Dvorský et al., 2015(Dvorský et al., , 2011Janatková et al., 2013;Řeháková et al., 2017;Rehakova et al., 2011;Čapková et al., 2016;Aschenbach et al., 2013). The Chamser Kangri glacier is located in Ladakh, Northwest India, in the southwestern extension of the 105 Tibetan Plateau on the northern slope of Chamser Kangri peak (6645 m a.s.l.) belonging to the Lungser Range above  The soils had a coarse-grained structure, with 30-50% of gravel and a pH between 7.8 and 8.5 (Rehakova et al., 2011).
The parent rock consists mainly of metasedimentary siliceous rocks, the Tso Morari granite gneiss (Epard and Steck, 2008). The cold steppe vegetation cover is sparse, characterized by alpine grasslands mainly constituted by hemicryptophytes (Dvorský et al., 2015(Dvorský et al., , 2011. The environmental conditions give rise to an extensive development of soil biological crusts dominated by cyanobacteria, which cover up to 40% of the soil surface in this area, and facilitate the initial establishment of vascular plants (Čapková et al., 2016; Řeháková et al., 2017). Cyanobacterial communities (Nostocales, Chroococcales and Oscillatoriales) were observed to be more abundant and species-rich in bare than vegetated soils (Řeháková et al., 2017). Together with Cyanobacteria, Gram-positive spore-forming bacteria (Actinobacteria and Firmicutes) were the main recorded bacterial clades (Řeháková et al., 2015).

Sampling description
Soil samples were collected along the frontal (M1-3) and lateral (M4) moraines of the Chamser Kangri glacier chronosequence (5711, 5710, 5700 and 5598 m a.s.l., M1-M4 respectively). The age of the chosen soils was estimated to be <1, 25, 50 and > 100 years, for M1 to M4, respectively, according to the slope and relative ice cover loss of 0.4% km -2 year -1 (Dvorský et al., 2015 Figure S2). The average soil temperature (5 cm below surface) during the week (23/2017) of sampling was 0.18°C, which corresponded to previous years values (based on long-term data collected from installed microclimatic stations TOMST®, Figure S1A). The average soil moisture for the corresponding week was around 0.25 cm 3 cm -3 and was 25-42% higher in respect to the corresponding period in the previous years ( Figure S1B). The sampling period represented the first moist period with above-zero temperatures after a long dry spell ( Figure S3). 140 Representative samples were obtained by pooling ~100 g of upper soil surface 0-5cm, collected at 10 different spots, i.e. approximately 1 kg of soil, at each locality (M1-4, Figure 1) along the glacier forefield chronosequence.
To evaluate the hydrological and evaporation influence on water oxygen (δ 18 OW) and hydrogen (δ 2 HW) isotopic composition, samples of the glacier snout and the stream water (covered by ice) coming from the glacier (W1-2, Figure   1), were taken. Samples were stored in zip-lock plastic bags/HDPE bottles and a thermo-bag to minimize biological 145 activity during transport to the laboratory. Soils were sieved after removing the gravel material (2-mm mesh) and homogenized. Subsamples for cryo-distillation, dry weight measurement and enzymatic activities were taken and stored at 4°C until processed.

150
The general soil characteristics were measured using standard laboratory procedures on the sieved samples. Soil pH was measured in 1:5 soil: water suspensions (ISO, 2005), total organic carbon and total nitrogen were measured on a TOC/TN analyser (Formacs), total P and micronutrients were analyzed by ICP-OES after sequential digestion by HNO3 and HClO4 (Kopacek et al., 2001). Soil texture was estimated based on particle size analysis (Afnor, 1994).
Bulk density of soil, used to calculate P stocks, was estimated using a pedotransfer function (Leonaviciute, 2000), 155 corresponding to eluvial deposits considering both the soil texture and organic carbon content.

Extraction and purification of phosphate in different pools
Phosphorus was sequentially extracted from the soil following a modified Hedley fractionation, which was upscaled to obtain a sufficient amount of phosphate for isotopic analyses according to . Five different P 160 pools were defined: resin-and hexanol-extractable P (bioavailable and microbial, respectively), NaOH-EDTA extractable P (bound to iron-(Fe) and aluminum-(Al) oxides and to organic P), and HCl extractable P (mineral P, mostly P bound to Ca in apatite). Additionally, parent material samples were crushed and extracted directly with 1M HCl. To measure the δ 18 OP values, all soil extracts were purified following the protocols of Tamburini et al. (2010, https://doi.org/10.5194/soil-2021-65 Preprint. Discussion started: 21 July 2021 c Author(s) 2021. CC BY 4.0 License. 2018). The inorganic P concentration in each pool was determined by the malachite green method (Ohno and Zibilske, 165 1991). For some samples, the soil available was not enough to extract and purify enough P to measure δ 18 OP reliably for all pools. This was the case of microbial P in the samples from sites M1 and M2 and the P bound to oxides in the sample from site M2. Only the organic pool of sample M4 contained enough material for δ 18 OP analyses.

Stable isotope analysis of oxygen in water and phosphates
170 Soil water was extracted by cryogenic vacuum extraction from one sample per sampling site (Orlowski et al., 2013).
Analyses of stable isotopes in samples from glacier snout and glacier stream water (δ 18 OW and δ 2 HW) were performed using the L2120i laser instrument (Picarro Inc.) at the Institute of Hydrology, Slovakia. Hydrogen and oxygen isotope analyses were calibrated against the V-SMOW (Vienna Standard Mean Ocean Water) and were reported in the standard ‰ notation. Typical precisions are better than ±0.1‰ for both δ 18 OW and δ 2 HW, respectively. The oxygen isotopes in 175 soil water were measured by equilibration with CO2 (Seth et al., 2006) using a gas bench (Thermo Scientific Gas Bench II) connected to an isotope ratio mass spectrometer (Thermo Scientific Delta V plus) at the Stable Isotope Laboratory of the Geological Institute of the ETH Zurich. The system was calibrated against V-SMOW, SLAP (Standard Light Antarctic Precipitation) and GISP (Greenland Ice Sheet Precipitation).
To distinguish between evaporated and non-evaporated stream water sources, we used deuterium excess (d-excess) that 180 is associated with kinetic isotopic fractionation and calculated as d-excess = δ 2 H -8 * δ 18 O (Dansgaard, 1964). Samples with d-excess value <10 suggest a deviation from the equilibrium fractionation conditions, indicating that the water may have been subject to evaporation (Dansgaard, 1964).
The stable oxygen isotope signature in phosphate was determined on a Vario Pyro Cube (Elementar, GmbH, Hanau, Germany) coupled in a continuous flow to an Isoprime 100 isotopic ratio mass spectrometer (Isoprime, Manchester, 185 UK). Calibration and corrections for instrumental drifts were done by repeated measurements of a Ag3PO4 internal standard (with a value of +14.20‰) and benzoic acids IAEA 601 and 602. The δ 18 O values are expressed in the standard delta notation with respect to V-SMOW. Reproducibility of the measurements based on repeated measurements of the internal standard was within 0.4‰.

Calculations of isotopic equilibrium values and box model
The equilibrium between oxygen in phosphate and water, given by intracellular P turnover by pyrophosphatase (PPase) was computed using the revised Chang and Blake equation (Chang and Blake, 2015): https://doi.org/10.5194/soil-2021-65 Preprint. Discussion started: 21 July 2021 c Author(s) 2021. CC BY 4.0 License.
where T is the temperature in °C and δ 18 OW is the isotopic composition of water in ‰. To integrate the variability at different time scales, Eq. 1 was solved using values from two reference periods: i) the 10 days preceding the sampling (from June 6, 2017, hereafter equilibrium 1), using on site soil temperatures and the δ 18 OW of soil water as the upper limit and δ 18 OW of June monthly rainfall adjusted for the altitudinal gradient as by Lone (2019) as lower limit; ii) a multi-annual average (hereafter equilibrium 2), calculated using a temperature range spanning the mean soil temperature 200 of warm months (years 2013-2017) and 0°C, considering biological activity as negligible below 0°C. We took a weighted average of seasonal precipitations, again adjusted for altitude (Lone 2019) with or without the evaporative enrichment observed in soil water. In using rainfall δ 18 OW as a proxy for soil water isotopic composition, we assumed that soil water reflects a mass balance of seasonal precipitation short of an evaporative enrichment (Roberts et al., 2015;Sprenger et al., 2017). 205 The expected δ 18 OP values of phosphate mineralized from phosphomonoesters (PME, δ 18 OPME) were calculated as follows (Liang and Blake, 2006): δ 18 OPME = 0.25 (δ 18 OW + ԑ) + 0.75 δ 18 OP-org (2) using the fractionation factor (ԑ) of -30‰ for hydrolysis of Po by alkaline phosphatases, as the pH range of studied soils is alkaline (Table 1), and a δ 18 OP-org value of +12.84‰, as the Po pool in site M4 assuming similar values for the other 210 soils.
To determine the contribution to available P from the mineral, organic and microbial P pools, we estimated P fluxes and calculated the expected isotopic composition of the available P as follows (box model approach, Tamburini et al., 2012): where f-Pmic is the P flux from microbial P turnover (mg P m -2 day -1 , see below) and δ 18 OPmic the isotopic composition 215 of the microbial P; f-Po is the P flux from the mineralization of Po, δ 18 OPME the expected δ 18 OP values of released phosphate; and f-Pmineral is the flux from the mineral P and δ 18 OPmineral the corresponding isotopic composition.
When the expected isotopic signatures matched the measured ones, we assumed P fluxes were estimated correctly. The flux from the hydrolysis of Po was estimated both from microbial respiration rates measured at 10°C in the same area (comparable period of sampling, distance to the glacier, age; pers.com. Capkova et al.) and from potential phosphatase 220 activities (Phillips et al., 2005) (Table 2). The first approach assumes that the mineralized P is proportional to the organic C released as CO2, according to their stoichiometric ratio in non-living organic matter (Achat et al., 2010;Bünemann, 2015): https://doi.org/10.5194/soil-2021-65 Preprint. Discussion started: 21 July 2021 c Author(s) 2021. CC BY 4.0 License.
Where C-CO2 is the carbon released through soil respiration (µmoles C m -2 day -1 ) adjusted for the average soil 225 temperature during the sampling period according to Lloyd and Taylor (1994), microbial efficiency was set at 0.4 (Murphy et al., 2003), C:Po is the molar ratio between the Po (0.25M NaOH-0.05M EDTA extraction) and the soil organic carbon and AMP the atomic mass of P.
The second approach yielded very high Po mineralization rates and consequently the δ 18 OPexpected of available P were strongly depleted compared to the measured values. This confirms that measured phosphomonesterase activities 230 reflected potential rather than actual rates (Bünemann, 2015). This second approach was, therefore, discarded. To estimate P fluxes from microbial P (f-Pmic), we considered two different turnover times for microbial P: 15 days (as reported for microbial N and microbial P in P-depleted forest soil, Fisk et al. 1998;Schmidt et al. 2007)  and 70 days (as reported for microbial P in arable soils, Oehl et al. 2001). Finally, the flux from the mineral P (f-Pmineral) was estimated by dividing the difference in stock concentrations by the difference in age of sites M4 and M1, as 235 in Tamburini et al. (2012).
To account for uncertainties introduced with calculations, the mean and standard deviation of the expected isotopic signatures were obtained with a Monte Carlo error propagation simulation (Anderson, 1976). Calculations were repeated 10 million times by varying the variables (δ 18 O and P fluxes) according to their mean and standard deviation from analytical replicates. When the δ 18 OP of microbial P were missing (site M1 and M2), we used the entire range of values 240 measured at sites M3 and M4.

Soil enzyme activities
To understand the regulation of enzymes activity in our soil chronosequence, we assessed five enzymes involved in mineralization of organic C, N and P (β-glucosidase, cellobiosidase, chitinase, leucine aminopeptidase and acid/alkaline 245 phosphatase) (Marx et al. 2001;Bárta et al. 2014). Briefly, 0.5 g of soil was homogenized in 10 mL of MQ water using microplate reader using the excitation and emission wavelength of 365 and 450 nm, respectively. Enzyme efficiencies were calculated as enzyme activities in nmol h -1 g -1 soil dry weight, using an eight-point calibration curve.

General soil characteristics
Particle size analysis revealed all forefield soils being sandy, with 78.0-85.5% of sand (Table 1). M1 had the highest portion of clay, silt and fine sand, 4.1, 17.8 and 58.6%, respectively. The soil at the M3 site was the poorest in clay with only 0.5%. The estimated bulk density ranged between 1.4 and 1.7 Mg m -3 from the oldest soil to the youngest, respectively. The total organic carbon concentration increased with soil age, from 0.6 g kg -1 in the youngest soil M1 to 260 21.7 g kg -1 at M4. The soil N content was low in the young soils at M1-M3, ranging from 0.03 to 0.6 and was higher at M4 with 2.4 mg kg -1 . The total P did not show a particular trend along the chronosequence, ranging from 0.72 to 0.93 g kg -1 , with the lowest content at M2 and highest at M3. The molar C:N ratio was relatively constant along the chronosequence, averaging 9, except at the youngest site M1, where total nitrogen concentration was very low and the C:N ratio the highest (Table 1). The C:Po and N:Po molar 265 ratios increased along the chronosequence from 527 to 932 and from 24 to 86, respectively. This increase was greater between 0 and 25 years (M1 and M2) and 50 and 100 years (M3 and M4).
The soil pH gradually decreased from 8.7 at the youngest site M1 to 7.7 at the oldest site M4.

Enzymes activity
The activity of all enzymes involved in C, N and P mineralization increased with site age (R 2 = 0.91, 0.97 and 0.88 respectively; Table 1, Table S2). The sum of the C-decomposing (β-glucosidase and cellobiosidase) and N-decomposing enzymes (leucine 275 aminopeptidase and chitinase), was positively correlated with TOC and TN, respectively (R 2 = 0.97 and 0.93). Phosphatase activities were positively correlated with activities of other enzymes, and several soil characteristics (TOC, TN, fine silt content, total P, Ca, Fe, Mg). The activities increased with soil age, and were inversely correlated with the HCl extracted mineral P and positively with all other P pools (Table S2). Overall, the lowest activity was measured for chitinase and cellobiosidase enzymes responsible for the hydrolysis of glycosidic bonds in chitin and cellulose, respectively, ranging from 0.2 to 9.8 nmol h -1 g -1 of soil. The highest enzyme 280 activity was measured for phosphatases, with 208.9 nmol h -1 g -1 at M4. Despite the increasing activity from less to more developed soils, very similar activities were observed at M2 and M3 for β-glucosidase and phosphatases, enzymes responsible for the hydrolysis of glucose from cellobiose, and of phosphate from phosphosaccharides, nucleotides and phospholipids, respectively (Table 1).

Stable isotopes in glacier snout, stream and soil water and isotopic equilibrium
The isotopic values at the glacier snout were -15.8‰ for δ 18 OW and -116.3‰ for δ 2 HW, and -12.4‰ for δ 18 OW and -94.3‰ for δ 2 HW in the glacier stream water. The d-excess at the glacier snout was 10.1% suggesting that there was no evaporation bias. However, d-excess for the stream water was 4.9%, which is indicative of a strong evaporation signal.
The δ 18 OW values of the soil water from the four sites were variably enriched compared to the glacier snout. They ranged 290 from -12.6‰ for site M1, located the glacier snout to around -3.0‰ and -5.6‰ for sites M2-M3 and M4, respectively, more distant from the retreating glacier (Table 2).
The δ 18 OP(eq) values expected at equilibrium (Eq. 1) ranged from +9.8‰ to +23.4‰ for the short-term equilibrium and from +9.0 to +23.4‰ for the long-term equilibrium ( Table 2). The upper range of the isotopic equilibrium varied according to the δ 18 OW values of the soil water and was therefore significantly higher at the intermediate sites M2-M3 295 than that at sites M1 and M4, due to heavier δ 18 Ow (Table 2).

Phosphorus soil pools concentrations and their oxygen isotopic value
The sequentially extracted P pools are shown in Table S1. The mineral P pool accounted for 95.1 to 99.5% of total P.
The concentrations of the other pools were low and increased with soil age. The available P ranged between 0.3 and 2.3 mg P kg -1 , the microbial P between 0.4 and 7.1 mg P kg -1 , the P bound to Fe-and Al-oxides increased from 2.3 to 9.0 305 mg P kg -1 , and the organic P from 3.1 to 60.0 mg P kg -1 .
The δ 18 OP of the available P ranged from +4.66‰ at site M1 to +8.90‰ at site M3. The isotopic value of the microbial P was higher compared to the available P and equal to +12.9‰ at site M3 and +7.

Nutrients limitation and microbial P dynamics during early stages of pedogenesis
Nitrogen is absent from most parent materials, but increases across the stages of ecosystem development through biological N2 fixation (Smittenberg et al., 2012). Increased C and N availability along with soil development may 325 eventually result in P becoming the secondary limiting nutrient .
Accordingly, total N concentration in the Chamser Kangri forefield increased with soil age, along with the potential activity of C-and N-decomposing enzymes. The C:N ratio, consequently, decreased (Table 1) limitation in most soils with a C:N ratio above 13. We concluded that the most limiting nutrient to primary production 330 at the youngest site was N (cf. C:N:P ratios in Table 1). This conclusion is supported by the occurrence of N-fixing cyanobacteria in soil biological crusts at the younger sites (Couradeau et al., 2016). These N-fixing species could provide with a significant N input in subnival soils .
The increasing trend of the C:Po and N:Po molar ratios indicates that C and N accumulated faster than P in the soil organic matter (Table1). On the pedogenesis time scale, assuming inputs to soil organic matter to be mainly derived 335 from microbial products, C and Po accumulate at a similar rate when P is not limiting and the mineralization of Po is driven by the need for carbon rather than the need of P (Wang et al., 2016). The increase in C:Po and N:Po ratios may, therefore, suggest the progressive onset of P limitation or co-limitation. Alternatively, it may point to a shift in the quality (stoichiometry, recalcitrance) of organic matter inputs. However, the available P concentrations along the Chamser Kangri chronosequence were rather low relative to comparably young chronosequences (Tamburini et al., 340 2012;Wang et al., 2016;Zhou et al., 2019). Moreover, the microbial P pool increasingly exceeded the available P pool and the potential phosphatases activity increased with soil age, which can be interpreted as indications of P limitation (Lajtha and Schlesinger, 1988). With the progressive colonization by vascular plants at sites M3 and M4 (20-50% to 70-80% plant cover, respectively, Figure S2), it is likely that microbes started to compete effectively with plants for the available P (Lajtha and Schlesinger 1988;Seeling and Zasoski 1993;Zhou et al. 2013). 345 Contrary to what has been found previously in an alpine glacier forefield (Tamburini et al., 2012), the δ 18 OP of microbial P at sites M3 and M4 were well below the isotopic equilibrium with soil water (Figure 2 and Table 2). However, the δ 18 OP of microbial P at site M3 did fall within the short-term equilibrium range, which includes the isotopic value of rainfall water. This suggests that microbes were only active under favorable temperature and moisture conditions, i.e. after precipitations, and rapidly turned P over, bringing it to isotopic equilibrium. Under such extreme conditions, 350 microbial activity might be disrupted by frequent droughts and occurrence of low temperatures. The isotopic value of microbial P might, therefore, reflect previous favorable conditions rather than those occurring at the moment sampling (Shen et al., 2020). At site M3, this translated in a δ 18 OP somewhere in between the equilibrium with rain water and soil water.
At the oldest site M4, however, the δ 18 OP of microbial P was much lower compared to site M3 and lied below the short-355 term equilibrium range, suggesting that other processes intervened. The difference between the two sites (about 5.2‰) might be determined by differences in soil water dynamics. The lower clay content characterizing site M3 potentially accelerates soil water evaporation. Indeed, at the sampling time, the δ 18 OW at site M4 was lower by approximately 2.6‰ compared to site M3 (Table 2), thus partially explaining the offset between the two microbial P isotopic compositions. https://doi.org/10.5194/soil-2021-65 Preprint. Discussion started: 21 July 2021 c Author(s) 2021. CC BY 4.0 License.
Additionally, the hydrolysis of Po compounds induces the release of Pi with low δ 18 O (Liang and Blake, 2006;von 360 Sperber et al., 2014). Low δ 18 OP have been observed in microbial P of forest soils and related to phosphomonoesterase catalyzed reactions induced by intracellular Po recycling in response to P limiting conditions (Pistocchi et al., 2020).
Alternatively, low metabolic activity and dormancy might induce hydrolysis of Po within cells for maintenance of basic functions, thus potentially leading to low δ 18 OP values (Lebre et al. 2017;Blagodatskaya and Kuzyakov 2013).
Together with differences in evaporative enrichment of soil water, differences in the metabolic status of the microbial 365 community, such as internal Po recycling in response to P limitation or adverse conditions, can, therefore, explain the lower microbial P δ 18 O values observed at the oldest site.

Contributions of biological P cycling to the inorganic P pools: long-term P dynamics
We observed a significant departure from the isotopic value of the parent material at sites M1 and M4 for available P, 370 at M3 and M4 for the P bound to oxides and at M3 for mineral P (Figure 2). As the exchange of O atoms between phosphate and water is negligible in absence of biological processes (Blake et al., 2005;Lecuyer et al., 1999;Winter et al., 1940), a departure from the isotopic value of the parent material suggests that these Pi pools underwent to a certain extent biological transformations.
Conversely, the mineral P maintained an isotopic value similar to the parent material also at the oldest site (˜100 years). 375 The mineral P pool was, therefore, still composed by primary phosphate minerals after 100 years of soil development.
This pool represented still 95% of the total P. The decline in soil mineral P stocks (-20%, see Table 1) was much less pronounced than that observed in other young chronosequences in alpine or other cold environment (Egli et al., 2012;Celi et al., 2013;Zhou et al., 2019). Unlike these studies, along the Chamser Kangri chronosequence the pH decreased only slightly, most likely because of less acidic inputs from rainfall and a slower colonization by vascular plants, which 380 prevented the rapid dissolution of primary apatite (Lajtha and Schlesinger, 1988).
Unlike the classic Walker and Syers (1976), there was no evident decrease of total soil P in the top 0-5 cm over the first 100 years of soil development (Table 1). However, while primary mineral P declined, an accumulation of P in organic and secondary mineral forms associated with metal oxides was observed, as well as a slight increase in the available P (Table 1). The δ 18 OP of P bound to oxides was initially similar to that of the parent material (Figure 2, site M1), possibly 385 reflecting the alteration processes of primary minerals with the formation of oxides-bound P by processes not cleaving the P-O bond, e.g. weathering by organic acids (Brunner et al. 2011;Mitchell et al. 2016). As this pool built up with soil development, its δ 18 OP deviated from that of the parent material and remained within or below the long-term isotopic equilibrium at the intermediate M3 and oldest M4 site, respectively (Figure 2 and Table 2). The exchange of phosphate https://doi.org/10.5194/soil-2021-65 Preprint. Discussion started: 21 July 2021 c Author(s) 2021. CC BY 4.0 License.
ions between the soil solution and the P bound to oxides occur within months to years (Helfenstein et al. 2020). The 390 difference in δ 18 OP of this pool at different sites should, therefore, depend on persistent long-term differences in the contributing processes (e.g. sorption of phosphate from the soil solution vs phosphate from dissolution of parent material or persistent differences in water isotopic composition).
As reported previously, δ 18 OP of sorbed or mineral Pi might approach isotopic equilibrium values via sorption onto soil solid phase or precipitation as secondary minerals of biologically cycled P (Tamburini et al., 2012;Joshi et al., 2016;395 Helfenstein et al., 2018;Roberts et al., 2015). In a study of a costal dunes chronosequence in New Zealand, the NaOHextractable Pi approached equilibrium after few millennia of soil development (Roberts et al., 2015). This has apparently happened much earlier in the Chamser Kangri chronosequence (site M3 is approximately 50 years old). This might be due to the relatively high proportion of the microbial P pool to the P bound to oxides, which additionally increased from 17% to 79% along the chronosequence (Table S1). Considering the large range of the long-term equilibrium (Table 2), 400 we cannot rule out that the P bound to oxides at this development stage (50 to 100 years) was only partially constituted by biologically cycled P and its isotopic value only partially overprinted.
Conversely, the drop of the δ 18 OP value of P bound to oxides at the oldest site might be linked to the increase of Po mineralization rate. Depleted δ 18 OP have been observed for NaOH-extractable Pi at the oldest sites of the mentioned coastal dune chronosequence and related to an increasing contribution of Po mineralization to P cycling as the soil 405 develops (Roberts et al., 2015). The low δ 18 OP of P bound to oxide observed at site M4 might, therefore, result from the sorption onto Fe and Al oxides of 18 O-depleted Pi released after extracellular Po mineralization. This explanation is corroborated by other findings, such as the sharp increase of Po pool, which tripled compared to site M3 (Table S1), the similarly low δ 18 OP values of the available P and the results of the box model (see section 4.3). In the Chamser Kangri chronosequence, this dynamic appeared to be faster than in the dune chronosequence, maybe because it affected a much 410 smaller quantity of the total soil P stock compared to previous studies (Roberts et al., 2015).

Contributions of biological P cycling to the available P: short-term P dynamics
The build-up of the microbial P pool observed across the chronosequence may suggest that this pool would increasingly contribute to the available P. With the help of a box model, we try to elucidate the contributions to the available P from 415 microbial P turnover and Po mineralization. We consider that the box-model P fluxes were estimated correctly, when the calculated δ 18 OP of the available P (δ 18 OPexpected, Eq. 3) closely matched the measured ones, e.g. their difference is less than twice the standard deviation of samples replicates (0.8‰). This is the case for sites M3 and M4, when assuming the lowest microbial P turnover time of 15 days (see Table 3). The difference between the δ 18 OPexpected and the measured https://doi.org/10.5194/soil-2021-65 Preprint. Discussion started: 21 July 2021 c Author(s) 2021. CC BY 4.0 License. δ 18 OP is instead greater for site M1 (+2.7‰) and M2 (+1.7‰). Differences between the measured and estimated δ 18 OP 420 values increase for all sites with increasing microbial P turnover time. For sites M3 and M4, the estimated fluxes from microbial P (f-Pmic) accounted for approximately 60% to 80% of the available P. Therefore, in agreement with our initial hypothesis, microbial P largely controlled the available P at these stages of pedogenesis (50 to 100 years). Accordingly, the ratio of microbial P to available P increased from approximately 1 to 3 along the chronosequence. Microbial P turnover might be accelerated by repeated freezing-thawing and drying-rewetting cycles occurring frequently in the 425 area, thus fostering its transfer to the available P pool, especially in a low-sorbing sandy soils (Chen et al., 2021). In addition, microbial grazing might potentially play a role .
At sites M1 and M2, the measured and estimated δ 18 OP values of the available P do not match, indicating either incorrect assumptions or the occurrence of processes we could not account for in the box model. The δ 18 OP of microbial P at sites M1 and M2 was not measured, so we assumed that these values could vary within the range of values from sites M3 430 and M4. However, the δ 18 OPexpected was poorly sensitive to variations in the δ 18 OP of the microbial P, so we can exclude that this assumption is a source of error. Additionally, although the δ 18 OP values of available P at site M2 was very close to that of mineral P (+8.49 and +7.95‰, respectively), we exclude that fluxes from the mineral P pool could have strongly influenced the isotopic composition of the available P. Indeed, the flux from mineral P (f-Pmineral) should be more than 100 times the one we estimated to account for the observed δ 18 OP of available P. Such a net flux appears 435 unlikely given the very slow exchange time (in the order of millennia) of the mineral P with the soil solution at alkaline soil pH (Helfenstein et al., 2020). The flux from the P bound to soil oxides via sorption/desorption processes was not accounted for in the box model. However, at the youngest sites, this flux can be considered as negligible compared to other contributions, due to the relatively small concentration of the P bound to oxides. Moreover, the measured δ 18 OP of available P at site M1 was clearly lower than that of the parent material and of the P bound to oxides (Figure 2). 440 As the δ 18 OPexpected were lower compared to the measured δ 18 OP of available P at sites M1 and M2, we conclude that the flux from Po mineralization (f-Po), which would carry a low δ 18 OP value, was possibly overestimated. Organic P mineralization was calculated from the CO2 released by microbial respiration, assuming that a stoichiometric proportion of Po was mineralized from non-living soil organic matter (see Eq. 4) (Achat et al., 2010;Bünemann, 2015). Stabilization of Po compounds by adsorption onto soil particles (Zhou et al., 2019), could lead to the mineralization of a lower 445 proportion of Po than C. However, considering the soil sandy texture and the fact that the C:Po ratio increased along the chronosequence, we exclude that abiotic stabilization played a major role. A second possibility is that C:Po was overestimated. The C:Po used for the f-Po calculation corresponded to the C:Po of the bulk soil organic matter (Eq. 4).
However more labile Po pools, such as microbial biomass, might have a narrower C:Po ratio, for example, ranging https://doi.org/10.5194/soil-2021-65 Preprint. Discussion started: 21 July 2021 c Author(s) 2021. CC BY 4.0 License. between 6 and 13 (Bünemann 2015 and references therein). The preferential mineralization of microbial necromass over 450 mineralization of the bulk soil organic matter might, therefore, explain the observed discrepancy.
At the oldest sites (M3 and M4), the good match between δ 18 OPexpected and measured δ 18 OP indicates that the f-Po were estimated correctly and accounted for approximately one third of total P flux to available P. Unlike the youngest sites, Po appeared to be mineralized in stoichiometric proportion to C in the non-living soil organic matter. This finding does not agree with the observed increase in the C:Po ratio between M3 and M4, which indicates a faster depletion of Po than 455 C on the pedogenesis time scale. This inconsistency might be explained by the fact that the box model only captures short-term P dynamics, as it is built on the isotopic composition of the available P, which varies on a seasonal scale and therefore potentially reflects transient conditions (Angert et al., 2011).
As discussed earlier, the increase of the C:Po ratio along the chronosequence suggests that in the long run Po was mineralized faster than C. Concurrently, the colonization by vascular plants might have contributed to modify the soil 460 organic matter stoichiometry through inputs with higher C:P ratio. More investigation on the contribution of plant litter to soil organic matter would be needed to clarify this aspect.

Conclusion
In the Chamser Kangri chronosequence, only a minor fraction of the total P in the top 5 cm of the newly formed soil 465 was affected by biological processes at early soil development stages. However, through a box model approach using stocks and δ 18 O of phosphate pools, we can conclude that the available P was mostly controlled by the microbial P via a rapid microbial P turnover, possibly accelerated by drying-rewetting or freezing-thawing cycles. Organic P mineralization became important in replenishing the available P pool after 50 years of soil development and contributed also to the phosphate sorbed on oxides. Unlike previous studies in other alpine environments, the P associated to primary 470 minerals decreased only by 20% after approximately 100 years of soil development and its isotopic composition reflected negligible biological cycling and secondary minerals precipitation. Finally, although cold arid conditions slowed down the weathering of primary P minerals and controlled biological activity, microbes still played a pivotal role in controlling the P dynamics, which affected P distribution to inorganic and organic soil pools. Our study highlights that also in extreme environments, integrating the analysis of the isotopic composition of oxygen in soil P pools in 475 chronosequence studies can provide insights in the short-and long-term P dynamics. https://doi.org/10.5194/soil-2021-65 Preprint. Discussion started: 21 July 2021 c Author(s) 2021. CC BY 4.0 License. 5.6 (0.6) b 4.38 (0.59) 1.9 a,b microbial P fluxes were calculated assuming a microbial P turnover rate of 15 and 70 respectively.  Figure S1: Long-term temperature and moisture content in soils from January 2013 to June 2018, Figure S2: Photographs of the locations of the individual sampling sites, Figure S3: Daily soil temperature and soil moisture contents from July 2016 to June 2017 Monthly variations in temperature and moisture content in soils, Discussion S1: Water isotopes. 490 Table S1: Sequentially extracted P pools,

500
The authors declare that they have no conflict of interest.

Data availability
The dataset used in this work is available at the following link: https://doi.org/10.15454/TKOUKH 505

Acknowledgment
This study was financially supported by the Czech Science Foundation (GACR (21-26883S, GACR 21-04987S) a longterm research development project no. RVO 67985939 (Czech Academy of Sciences), Biology Centre CAS (Institute of Hydrobiology and Soil Biology) & SoWa (MEYS projects LM2015075 and EF16_013/0001782 -Soil and Water Ecosystems Research). We would like to thank Vilem Ded for assistance with graphical data analysis in R, Iva Tomkova 510 for analyzing soils for TP and Lenka Capkova for technical help with enzyme activity measurements in the laboratory for Ecosystem Biology, at the University of South Bohemia and Institute of Hydrology in Slovakia for the analyses of stable isotopes in samples from glacier snout and glacier stream water. We are grateful to Madalina Jaggi and Stefano Bernasconi for the analysis of oxygen isotopes in soil water at the Stable Isotope Laboratory of the Geological Institute of the ETH Zurich. 515 https://doi.org/10.5194/soil-2021-65 Preprint. Discussion started: 21 July 2021 c Author(s) 2021. CC BY 4.0 License.