Application of a laser-based spectrometer for continuous in situ measurements of stable isotopes of soil CO2 in calcareous and acidic soils

The short-term dynamics of carbon and water fluxes across the soil–plant–atmosphere continuum are still not fully understood. One important constraint is the lack of methodologies that enable simultaneous measurements of soil CO2 concentration and respective isotopic composition at a high temporal resolution for longer periods of time. δ13C of soil CO2 can be used to derive information on the origin and physiological history of carbon, and δ18O in soil CO2 aids in inferring the interaction between CO2 and soil water. We established a real-time method for measuring soil CO2 concentration, δ13C and δ18O values across a soil profile at higher temporal resolutions (0.05–0.1 Hz) using an off-axis integrated cavity output spectroscopy (OA-ICOS). We also developed a calibration method correcting for the sensitivity of the device against concentration-dependent shifts in δ13C and δ18O values under highly varying CO2 concentration. The deviations of measured data were modelled, and a mathematical correction model was developed and applied for correcting the shift. By coupling an OA-ICOS with hydrophobic but gas-permeable membranes placed at different depths in acidic and calcareous soils, we investigated the contribution of abiotic and biotic components to total soil CO2 release. We found that in the calcareous Gleysol, CO2 originating from carbonate dissolution contributed to the total soil CO2 concentration at detectable degrees, potentially due to CO2 evasion from groundwater. The C-CO2 of topsoil at the calcareous soil site was found reflect δ13C values of atmospheric CO2, and the δ13C of topsoil CO2 at the acidic soil site was representative of the biological respiratory processes. δ18O values of CO2 in both sites reflected the δ18O of soil water across most of the depth profile, except for the 80 cm depth at the calcareous site where a relative enrichment in 18O was observed.

topic composition is available for short-term variations on a diurnal scale (Werner and Gessler, 2011). The lack of a proper understanding of the diurnal fluctuations in soil CO 2 release might introduce uncertainty in estimating the soil carbon budget and the CO 2 fluxes to the atmosphere. The isotopic composition of soil CO 2 and its diel fluctuation can be a critical parameter for the partitioning of ecosystem gas exchange into its components (Bowling et al., 2003;Mortazavi et al., 2004) and for disentangling plant and ecosystem processes (Werner and Gessler, 2011). By assessing the δ 13 C of soil CO 2 , it is possible to identify the source for CO 2 (Kuzyakov, 2006) and the coupling between photosynthesis and soil respiration when taking into account postphotosynthetic isotope fractionation (Werner et al., 2012;Wingate et al., 2010). δ 13 C soil CO 2 reflects, however, not only microbial and root respiration but also abiotic sources from carbonate weathering (Schindlbacher et al., 2015).
Soil water imprints its δ 18 O signature on soil CO 2 as a result of isotope exchange between H 2 O and CO 2 (aqueous). The oxygen isotopic exchange between CO 2 and soil water is catalysed by microbial carbonic anhydrase (Sperber et al., 2015;Wingate et al., 2009). Thus, soil CO 2 can give information on the isotopic composition of both soil water resources and carbon sources. The oxygen isotope composition of plant-derived CO 2 is both a tracer of photosynthetic and respiratory CO 2 and gives additional quantitative information on the water cycle in terrestrial ecosystems (Francey and Tans, 1987). To better interpret the δ 13 C and δ 18 O signals of atmospheric CO 2 , the isotopic composition and its variability in the different sources need to be better understood (Werner et al., 2012;Wingate et al., 2010).
The conventional method for estimating the δ 13 C and δ 18 O of soil CO 2 efflux is by using two end-member mixing models of atmospheric CO 2 and CO 2 produced in the soil (Keeling, 1958). The conventional methods for sampling soil produced CO 2 are chamber-based (Bertolini et al., 2006;Torn et al., 2003), "mini-tower" (Kayler et al., 2010;Mortazavi et al., 2004), and soil-gas-well-based (Breecker and Sharp, 2008;Oerter and Amundson, 2016) methods. In conventional methods, air sampling is done at specific time intervals, and δ 13 C and δ 18 O are analysed using isotope ratio mass spectrometry (IRMS; Ohlsson et al., 2005). Such offline methods have several disadvantages, like high sampling costs, excessive time consumption for sampling and analysis, and increased sampling error and low temporal resolution. Kammer et al. (2011), showed how error prone the conventional methods could be while calculating δ 13 C and δ 18 O (up to several per mil when using chamber and mini-tower-based methods; Kammer et al., 2011). In chamber-based systems, nonsteady-state conditions may arise within the chamber due to increased CO 2 concentrations, which in turn hinders the diffusion of 12 CO 2 more strongly than that of heavier 13 CO 2 (Risk and Kellman, 2008). Moreover, it has been found that δ 18 O of the CO 2 inside a chamber is significantly influenced by the δ 18 O of the surface soil water, as an equilibrium iso-topic exchange happens during the upward diffusive movement of soil CO 2 (Mortazavi et al., 2004). The advent of laser-based isotope spectroscopy has enabled cost-effective, simple, and high precision real-time measurements of δ 13 C and δ 18 O in CO 2 (Kammer et al., 2011;Kerstel and Gianfrani, 2008). This technique opened up new possibilities for faster and reliable measurements of stable isotopes in situ, based on the principle of light absorption, using laser beams of distinct wavelengths in the near-and mid-infrared range (Bowling et al., 2003). Recently, several high-frequency online measurements of δ 13 C and δ 18 O, of soil CO 2 and 2 H, and of the 18 O of soil water vapour across soil depth profiles were reported by coupling either hydrophobic but gaspermeable membranes (installed at different depths in soil) or automated chamber systems with laser spectrometers (Bowling et al., 2015;Jochheim et al., 2018;Stumpp et al., 2018). Such approaches enable detection of vertical concentration profiles, temporal dynamics of soil CO 2 concentration, and the isotopic signature of soil CO 2 across different soil layers, thus aiding in identifying and quantifying various sources of CO 2 across the depth profile.
In 1988, O'Keefe and Decon introduced cavity ring-down spectroscopy (CRDS) for measuring the isotopic ratio of different gaseous species based on laser spectrometry (O'Keefe and Deacon, 1988). With the laser-based spectrometry techniques, measuring sensitivities up to parts per trillion (ppt) concentrations is achieved (von Basum et al., 2004;Peltola et al., 2012). In CRDS, the rate of change in the absorbed radiation of the laser light that is temporarily "trapped" within a highly reflective optical cavity is determined. This is achieved using resonant coupling of a laser beam to the optical cavity and active locking of laser frequency to cavity length (Parameswaran et al., 2009). Another well-established technique similar to CRDS is off-axis integrated cavity output spectroscopy (OA-ICOS). It is based on directing narrowband and continuous-wave lasers in an off-axis configuration to the optical cavity (Baer et al., 2002).
Even though OA-ICOS can measure concentration and isotope signature of various gaseous species at a high temporal resolution, we found pronounced deviations in δ 13 C and δ 18 O measurements from the absolute values when measured under changing CO 2 concentrations. So far, to our knowledge, no study detailing the calibration process of OA-ICOS CO 2 analysers correcting for fluctuations of both δ 13 C and δ 18 O values under varying CO 2 concentrations has been made available. Most of the OA-ICOS CO 2 analysers are built for working under stable CO 2 concentrations, so periodical calibration against in-house gas standards at a particular concentration is sufficient. However, as there are pronounced gradients in CO 2 levels in soils (Maier and Schack-Kirchner, 2014), CO 2 -concentration-dependent shifts in measured isotopic values have to be addressed and corrected. Such calibration is, however, also relevant for any other OA-ICOS application with varying levels of CO 2 (e.g. in chamber measurements). Hence the first part of this work comprises the establishment of a calibration method for OA-ICOS. The second part describes a method for online measurement of CO 2 concentrations and stable carbon and oxygen isotope composition of CO 2 in different soil depths by coupling OA-ICOS with gas-permeable hydrophobic tubes (membrane tubes, Accurel ® ). The use of these tubes for measuring soil CO 2 concentration (Gut et al., 1998) and the δ 13 C of soil CO 2 (Parent et al., 2013) has already been established, but the coupling to an OA-ICOS system has not been performed, yet.
We evaluated our measurement system by assessing and comparing the concentration of the δ 13 C and δ 18 O of soil CO 2 for a calcareous and an acidic soil system. The primary foci of this study are to (1) introduce OA-ICOS in online soil CO 2 concentration and isotopic measurements, (2) calibrate the OA-ICOS to render it usable for isotopic analysis carried out under varying CO 2 concentrations, and (3) analyse the dynamics of δ 13 C and δ 18 O of soil CO 2 at different soil depths in different soil types at a higher temporal resolution.

Instrumentation
The concentration of δ 13 C and δ 18 O values of CO 2 were measured with an OA-ICOS, as described in detail by Baer et al. (2002) and Jost et al. (2006). In this study, we used an OA-ICOS, (LGR CCIA 36-d) manufactured by Los Gatos Research Ltd in San Francisco, USA. The LGR CCIA 36d measures CO 2 concentration and δ 13 C and δ 18 O values at a frequency up to 1 Hz. The operational CO 2 concentration range was 400 to 25 000 ppm. The operating temperature range was +10-+35 • C, and the sample temperature range (gas temperature) was between −20 and 50 • C. The recommended inlet pressure was < 0.0689 MPa. The multiport inlet unit (MIU), an optional design that comes along with LGR CCIA 36-d, had a manifold of eight digitally controlled inlet ports and one outlet port. It presented the user with an option of measuring eight different CO 2 samples at the desired time interval. Three standard gases with distinct δ 13 C and δ 18 O values were used for calibration in this study (see Table S1 in the Supplement). The standard gases used in this study were analysed for absolute concentration and respective δ 13 C and δ 18 O values. δ values are expressed based on Vienna Pee Dee Belemnite (VPDB) CO 2 scale and were determined by high precision IRMS analysis.

Calibration set-up and protocol
We developed a two-step calibration procedure to (a) correct for concentration-dependent errors in isotopic data measurements and (b) correct for deviations in measured δ values from absolute values due to the offset (other than concentration-dependent error) introduced by the laser spectrometer. Also, we used Allan variance curves for determining the time interval to average the data (Nelson et al., 2008) in order to achieve the highest precision that can be offered by the LGR CCIA 36-d (Allan et al., 1997).
The first part of our calibration methodology was developed to correct for the concentration-dependent error observed in preliminary studies for δ 13 C and δ 18 O values measured using OA-ICOS. Such a calibration protocol was used in addition to the routine three-point calibration performed with in-house CO 2 gas standards of known δ 13 C and δ 18 O values. We developed a CO 2 dilution set-up (see Fig. 1) in which each of the three CO 2 standard gases was diluted with synthetic CO 2 -free air (synthetic air) to different CO 2 concentrations. By applying a dilution series, we identified the deviation of the measured (OA-ICOS) from the absolute (IRMS) δ 13 C and δ 18 O values depending on CO 2 concentration (see Fig. 4). The δ 13 C and δ 18 O values of our inhouse calibration gas standards were measured via cryoextraction and dual-inlet IRMS. δ 13 C and δ 18 O of the standard gases (see Table S1) across a wide range of CO 2 concentrations are measured using OA-ICOS. The deviation of the measured δ 13 C, and δ 18 O from absolute values with respect to changing CO 2 concentrations was mathematically modelled and later used for data correction (see Fig. 5). A standard three-point calibration was then applied to correcting for concentration-dependent errors (see Fig. 7). The standards used covered a wide range of δ 13 C and δ 18 O values, including the values observed in the field of application.
Standard gases were released to a mass flow controller (ANALYT-MTC, series 358, MFC1) after passing through a pressure controller valve (see Fig. 1) with safety bypass (TESCOM, D43376-AR-00-X1-S, version 5). A Swagelok filter, (Stainless Steel All-Welded In-Line Filter; Swagelok, SS-4FWS-05, F1) was installed at the inlet of the flow controller (ANALYT-MTC, series 358, MFC1). Synthetic air was released and passed to another flow controller (ANALYT-MTC, series 358, MFC2) through a Swagelok filter (F2 in Fig. 1). CO 2 and synthetic air leaving the flow controllers (MFC1 and MFC2 respectively) were then mixed and drawn through a Teflon tube (P8) with a 6.35 mm outer diameter (OD), which was kept in a gas thermostat unit (see Fig. 1). The thermostat unit contained (a) a thermostat-controlled water bath (Kottermann, 3082) and (b) an Isotherm flask containing liquid nitrogen. The water bath was used to raise the temperature above room temperature and also to bring the temperature down to +5 • C by placing ice packs in the water bath. To reach low temperatures (−20 • C), we immersed the tubes in the isotherm flask filled with liquid N 2 . Leaving the thermostat unit, the gas was directed to the multiport inlet unit of the OA-ICOS. By using the thermostat unit, we introduced a shift in the reference gas temperature, and the aim was to test the temperature sensitivity of the OA-ICOS in measuring δ 13 C and δ 18 O values. The third CO 2 standard gas (which is used for validation) was produced by mixing the other two gas standards in equal molar proportions in a 10 L volume plastic bag with an inner aluminum foil coating and welded seams LGR CCIA 36-d). I (1, 2) represents CO 2 standards, CO 2 mix denotes gas standards mixed in equal molar proportion, I3 represents synthetic air, MFC (1, 2) denotes mass flow controller, F (1, 2) represents PTFE filter, V (1, 2, 3) denotes pressure-reducing valves, V4 shows three-way ball valve, V (5, 6) stands for pressure controller valve with safety bypass, P (1-7) denotes steel pipes, and P (8-11) represents Teflon tubing.
(CO 2 mix: Linde PLASTIGAS ® ) under 0.03 MPa pressure by diluting to the required concentration using synthetic air. The mixture was then temperature adjusted and delivered to the MIU by using a 6.35 mm (OD) Teflon tube (P10). From the multiport inlet unit, calibration gases were delivered into the OA-ICOS for measurement using a 6.35 mm OD Teflon tube (P9) at a pressure < 0.0689 MPa, with a flow rate of 500 mL min −1 . The gas leaving the OA-ICOS through the exhaust was fed back to the 6.35 mm (OD) Teflon tube (P8) by using a Swagelok pipe tee (Stainless Steel Pipe Fitting, Male Tee, 6.35 mm OD, Male NPT), intersecting the P8 line before entering the thermostat unit. Thus, the gas fed was looped in the system until steady values were reported by the OA-ICOS based on CO 2 (ppm), δ 13 C, and δ 18 O measurements. CO 2 gas standards were measured at 27 different CO 2 concentration levels ranging between 400 and 25 000 ppm. Every hour before sampling, synthetic air gas was flushed through the system to remove CO 2 to avoid memory effects. The calibration gases were measured in a sequence, one after the other, four times. During each round of measurement, every calibration gas was diluted to different concentrations of CO 2 (400-25 000 ppm), and the respective isotopic signature and concentration were determined. For each measurement of δ 13 C and δ 18 O at a given concentration, the first 50 readings were omitted to avoid possible memory effects of the laser spectrometer, and the subsequent readings for the next 256 s were taken and averaged to get maximum precision for δ 13 C and δ 18 O measurements. When switching between different calibration gases at the multiport inlet unit, synthetic air was purged through the systems for 30 s to avoid cross contamination.

Experimental sites
In situ experiments were conducted to measure δ 13 C, δ 18 O, and concentrations of soil CO 2 in two different soil types (calcareous and acidic soil). The measurements in a calcareous soil were conducted during June 2014 in cropland cultivated with wheat (Triticum aestivum) in Neuried, a small village in the upper Rhine Valley in Germany, situated at 48 • 26 55.5 N, 7 • 47 20.7 E, 150 m a.s.l. The soil type described as calcareous Fluvic Gleysol IUSS Working Group WRB (2015) developed on gravel deposits in the upper Rhine Valley. Soil depth was medium to deep, with high contents of coarse material (> 2 mm) up to 30 %-50 %. Mean soil organic carbon (SOC) content was 1.2 %-2 %, and SOC stock ranged between 50 and 90 t ha −1 . The average pH was found to be 8.6. The study site receives an annual rainfall of 810 mm and has a mean annual temperature of 12.1 • C.
In situ measurements in an acidic soil were conducted by the end of July 2014 in the model ecosystem facility (MODOEK) of the Swiss Federal Research Institute WSL in Birmensdorf, Switzerland (47 • 21 48 N, 8 • 27 23 E; 545 m a.s.l.). The MODOEK facility comprises 16 model ecosystems, split below ground into two lysimeters with an area of 3 m 2 and a depth of 150 cm. The lysimeters used for the present study were filled with acidic (Haplic Alisol) forest soil IUSS (2014) and planted with young beech trees (Arend et al., 2016). The soil pH was 4.0, with a total SOC content of 0.8 % (Kuster et al., 2013).

Experimental set-up
The OA-ICOS was connected to gas-permeable, hydrophobic membrane tubes (Accurel ® tubing, 8 mm outer diameter) of 2 m length, placed horizontally in the soil at different depths. Tubes were laid in six different depths (4,8,12,17,35, and 80 cm) for calcareous soil and three depths (10, 30, and 60 cm) for acidic soil.
Technical details of the measurement set-up are shown in Fig. 2. Both ends of the membrane tubes were extended vertically upwards, reaching the soil top by connecting them to gas impermeable Synflex ® tubing (8 mm OD) using Swagelok tube fitting union (Swagelok: SS-8M0-6, 8 mm tube OD). One end of the tubing system was connected to a solenoid switching valve (Bibus: MX-758.8E3C3KK), by using a stainless-steel reducing union (Swagelok: SS-8M0-6-6M), to the outlet of the LGR CCIA 36-d by using 6.35 mm (OD) Teflon tubing. The other end was connected via the multiport inlet unit to the gas inlet of the LGR CCIA 36-d.
This way, a loop was created in which the soil CO 2 drawn into the OA-ICOS was circulated back through the tubes and in and out of the OA-ICOS and measured until a steady state was reached. We experienced no drop in cavity pressure while maintaining a closed loop (see Fig. S2). Each depth was selected and continuously measured for 6 min at specified time intervals by switching to defined depths at the multiport inlet unit and also at the solenoid valve.

Instrument calibration and correction
The highest level of precision obtained for δ 13 C and δ 18 O measurements at the maximum measuring frequency (1 Hz) was determined by using Allan deviation curves (see Fig. 3). The maximum precision of 0.022 ‰ for δ 13 C was obtained when the data were averaged over 256 s, and the maximum for δ 18 O, 0.077 ‰, was obtained for the same averaging interval as for δ 13 C.
To correct for CO 2 concentration-dependent errors in raw δ 13 C and δ 18 O data, we analysed data obtained from the OA-ICOS to determine the sensitivity of δ 13 C and δ 18 O measurements against changing concentrations of CO 2 . We observed a specific pattern of deviance in the measured isotopic data from the absolute values (both for δ 13 C and δ 18 O) across CO 2 concentration ranging from 25 000 to 400 ppm (see Fig. 4). Uncalibrated δ 13 C and δ 18 O measurements showed a standard deviation of 6.44 ‰ and 6.80 ‰ respectively, when measured under changing CO 2 concentrations.
The dependency of δ 13 C and δ 18 O values on the CO 2 concentration was compensated by using a non-linear model. The deviations (Diff-δ) of the measured delta values (δ (OA-ICOS) ) from the absolute value of the standard gas (δ (IRMS) ) at different concentrations of CO 2 were calculated (Diff-δ = δ (OA-ICOS) − δ (IRMS) ). Several mathematical models were then fitted to Diff-δ as a function of changing CO 2 concentration (see Fig. 5). The mathematical model with the best fit for Diff-δ data was selected using the corrected Akaike information criterion (AIC C ; Glatting et al., 2007;Hurvich and Tsai, 1989;Yamaoka et al., 1978). The nonlinear model fits applied for Diff-δ 13 C and Diff-δ 18 O measurements are given in Tables 1 and 2, respectively. For Diffδ 13 C, a three-parameter exponential model fitted best with r 2 = 0.99 (see Table 3 for the values of the parameters; see Fig. S3a for model residuals), and a three-parameter power function model (see Table 2) with r 2 = 0.99 showed the best fit for Diff-δ 18 O (see Table 3 for the values of the parameters; see Fig. S3b for model residuals). The best fit was then introduced into the measured isotopic data (δ 13 C and δ 18 O) and corrected for concentration-dependent errors (see Fig. 6). After correction, the standard deviation of δ 13 C was reduced  to 0.08 ‰, and the deviation of δ 18 O to 0.09 ‰, for all measurements across the whole CO 2 concentration range.
After correcting the measured δ 13 C and δ 18 O values for the CO 2 concentration-dependent deviations, a three-point calibration (Sturm et al., 2012) was made by generating linear regressions with the concentration-corrected δ 13 C and δ 18 O values against absolute δ 13 C and δ 18 O values (see Fig. 7; see Fig. S4 for linear regression residuals). Using the linear regression lines, we were able to measure the validation gas standard, with standard deviations of 0.0826 ‰ for δ 13 C and 0.0941 ‰ for δ 18 O.
For the LGR CCIA 36-d, we found that routine calibration (correction for concentration-dependent error plus threepoint calibration) was necessary for obtaining the required accuracy, in particular under fluctuating CO 2 concentrations. The LGR CCIA-36d offers an option for calibration against a single standard, a feature which was already in place in a predecessor model (CCIA DLT-100; Guillon et al., 2012). This internal calibration is sufficient when LGR CCIA-36d is operated only under stable CO 2 concentrations. To correct for the concentration dependency, we introduced mathematical model fits, which corrected for the deviation pattern found for both δ 13 C and δ 18 O. We assume that these deviations are instrument specific and that the fitting parameters need to be adjusted for every single device. Experiments conducted to investigate the influence of external temperature fluctuations on OA-ICOS measurements did not show any significant changes in the temperature inside the optical cavity of the OA-ICOS (see Fig. S1). The previous version of the Los Gatos CCIA was strongly influenced by temperature fluctuations during sampling (Guillon et al., 2012). The lack of temperature dependency as observed here with the most recent model can be mostly due to the heavy insulation provided with the system, which was not found in the older models. Guillon et al. (2012) found a linear correlation between CO 2 concentration and respective stable isotope signatures with a previous version of the Los Gatos CCIA CO 2 stable isotope analyser. In our experiments with the OA-ICOS, the best fitting correlations between CO 2 concentration and δ 13 C and δ 18 O measurements were exponential and power functions, respectively. We assume that measurement accuracy is influenced by the number of CO 2 molecules present inside the laser cavity of the particular laser spectrometer, as we observed large standard deviation in isotopic measurements at lower CO 2 concentrations. This behaviour of an OA-ICOS can be expected, as it functions by sweeping the laser along an absorption spectrum, measuring the energy transmitted after passing through the sample. Therefore, energy transmitted is proportional to the gas concentration in the cavity. The laser absorbance is then determined by normalising against a SOIL, 5, 49-62, 2019 www.soil-journal.net/5/49/2019/ reference signal, finally calculating the concentration of the sample measured by integrating the whole spectrum of absorbance (O'Keefe et al., 1999).

Variation in soil CO 2 concentration, carbon, and oxygen isotope values
Figures 9 and 10 show the CO 2 concentration and the δ 13 C and δ 18 O measurements of soil CO 2 in the calcareous as well as in the acidic soil across the soil profile with a subdaily resolution and as averages for the day, respectively. We observed an increase in the CO 2 concentration across the soil depth profile for both the calcareous and the acidic soil. Moreover, there were rather contrasting δ 13 C values across the profile for the two soil types. In the calcareous soil, CO 2 was relatively enriched in 13 C in the surface soil (4 cm) as compared to the 8 cm depth. Below 8 cm down to 80 cm depth, we found an increase in δ 13 C values. At 80 cm depth, the δ 13 C in soil CO 2 ranged between −7.15 ‰ and −3.35 ‰ (see Fig. 9), with a daily average of −6.19±1.45 ‰ (see Fig. 10), hence being clearly above atmospheric values (≈ −8.0 ‰). For δ 18 O values of calcareous soil, the depth profile showed no specific pattern, except for the δ 18 O values at 80 cm depth, which were found to be less negative than the values at the other depths. The δ 18 O value in the top 4 cm was found to be slightly more enriched that the 8 cm depth, and between 8-35 cm, δ 18 O values showed little variation relative to each other. For the sub-daily measurements, we observed a sharp decline in δ 18 O values at around 02:00 CET, which is also observed but less pronounced for the δ 13 C signal. We assume that the reason for such aberrant values is a technical issue rather than a biological process. It could be due to the fact that the internal pump in the OA-ICOS was not taking an adequate amount of gas into the optical cavity, thereby creating a negative pressure inside the cavity resulting in the observed aberrant values. The patterns observed for the δ 13 C values of CO 2 in the calcareous soil with 13 C enrichment in deeper soil layers can be explained by a substantial contribution of CO 2 from abiotic origin to total soil CO 2 release as a result of carbonate weathering and subsequent outgassing from soil water (Schindlbacher et al., 2015). According to Cerling (1984), the distinct oxygen and carbon isotopic composition of soil carbonate depends primarily on the isotopic signature of meteoric water and on the proportion of C 4 biomass present at the time of carbonate formation (Cerling, 1984) but also on numerous other factors that determine the 13 C value of soil CO 2 . CO 2 released as a result from carbonates in calcareous soil site have a distinct δ 13 C value of −9.3 (mean value across soil profile 0-80 cm depth; Fig. 8c), while CO 2 released during biological respiratory processes has δ 13 C values around −24 ‰, as observed in the acidic soil (Fig. 10e). The δ 13 C values of soil  CO 2 observed in the deepest soil layer in the calcareous soil site most likely indicate the presence of carbonate sources of a pedogenic and geologic origin. Even though the contribution of CO 2 from abiotic sources to soil CO 2 is often considered to be low, several studies have reported significant proportions ranging between (10 %-60 %), emanating from abiotic sources (Emmerich, 2003;Plestenjak et al., 2012;Ramnarine et al., 2012;Serrano-Ortiz et al., 2010;Steven-son and Verburg, 2006;Tamir et al., 2011). Bowen and Beerling (2004) showed that isotope effects associated with soil organic matter (SOM) decomposition can cause a strong gradient in δ values of soil organic matter with depth but are not always reflected in the δ 13 C values of soil CO 2 . We have measured soil samples for bulk soil δ 13 C, carbonate δ 13 C, and δ 18 O values and have also determined the percentage of total carbon in the soil across a depth profile of (0-80 cm; see    We observed an increase in δ 13 C values for bulk soil in deeper soil layers (see Fig. 8a, c). Moreover, the carbonate δ 13 C values also got more positive in the 60-80 cm layer. Since total organic carbon content decreases with depth, it can be assumed that the CO 2 derived from carbonate weathering, having less negative δ 13 C values, more strongly contributed to the soil CO 2 (especially since we see an increase in soil CO 2 concentration with depth). This is accordance with the laser-based measurements which showed a strong increase in the δ 13 C of soil CO 2 in the deepest soil layer, leading us to the hypothesis that this signal indicates a strong contribution of carbonate-derived CO 2 . Water content, soil CO 2 concentration, and the presence of organic acids or any other source of H + are the major factors influencing carbonate weathering, and variations in soil CO 2 partial pressure, moisture, temperature, and pH can cause degassing of CO 2 which contributes to the soil CO 2 efflux (Schindlbacher et al., 2015;Zamanian et al., 2016). CaCO 3 solubility in pure H 2 O at 25 • C is 0.013 g L −1 , but in weak acids like carbonic acid, the solubility is increased up to 5 fold (Zamanian et al., 2016). The production of carbonic acid due to CO 2 dissolution will convert carbonate to bicarbonates, resulting in exchange of carbon atoms between carbonates and dissolved CO 2 .We assume that at our study site, the topsoil is decar-bonated due to intensive agriculture for a longer period, thus the soil CO 2 there originates primarily from autotrophic and heterotrophic respiratory activity. In contrast to the deeper soil layers, where the carbonate content is high, CO 2 from carbonate weathering is assumed to be a dominating source of soil CO 2 . Also, outgassing of CO 2 from the large groundwater body underneath the calcareous Gleysol might contribute to the inorganic CO 2 sources in the deeper soil, as we found the groundwater table to be 1-2 m below the soil surface. Relative 13 C enrichment of the CO 2 in the topsoil (4 cm) compared to that at 8 cm depth is probably due to the invasive diffusion of atmospheric CO 2 , which has a δ 13 C value close to −8 ‰ (e.g. Levin et al., 1995). The δ 18 O patterns for CO 2 between 4 and 35 cm might reflect the δ 18 O of soil water with stronger evaporative enrichment at the top and 18 O depletion towards deeper soil layers. In comparison, the strong 18 O enrichment of soil CO 2 towards 80 cm in the calcareous Gleysol very likely reflects the 18 O values of groundwater lending further support to the high contribution of CO 2 originating from the outgassing of groundwater. We, however, need then to assume that the oxygen in the CO 2 is not in full equilibrium with the precipitation-influenced soil water. Since mainly microbial carbonic anhydrase mediates the fast equilibrium between CO 2 , and water in the soil and the microbial activity is low in deeper soil layers (Schmidt et al., 2011), we speculate that in deep layers with a significant contribution of groundwater derived CO 2 to the CO 2 pool, a lack of full equilibration with soil water might be the reason for the observed δ 18 O values. Soil CO 2 concentration in the acidic soil showed a positive relationship with soil depth as CO 2 concentration increased along with increasing soil depth (Figs. 9 and 10). CO 2 concentrations were distinctly higher than in the calcareous soil, very likely due to the finer texture than in the gravel-rich calcareous soil. δ 13 C values amounted to approximately −26 ‰ in 30 and 60 cm depth, indicating the biotic origin from (autotrophic and heterotrophic) soil respiration (Schönwitz et al., 1986). In the topsoil, δ 13 C values did not strongly increase, pointing towards a less pronounced inward diffusion of CO 2 in the acidic soil site, most likely due to more extensive outward diffusion of soil CO 2 , as indicated by the still very high CO 2 concentration at 10 cm creating a sharp gradient between soil and atmosphere. Moreover, the acidic soil was rather dense and contained no stones, strongly suggesting that gas diffusivity was rather small. δ 18 O depth patterns of soil CO 2 in the acidic soil most likely reflected δ 18 O values of soil water as CO 2 became increasingly 18 O depleted from top to bottom. The δ 18 O of deeper soil layers CO 2 (30-60 cm) was close to the values expected when full oxygen exchange between soil water and CO 2 occurred (Kato et al., 2004). Assuming an 18 O fractionation of 41 ‰ between CO 2 and water (Brenninkmeijer et al., 1983), this would result in an expected value for CO 2 of ≈ −10 ± 2 ‰ vs. VPDB CO 2 . Corresponding results have been shown for δ 18 O of soil CO 2 using similar hydrophobic gas-permeable membrane tubes used when measuring δ 18 O of soil CO 2 and soil water in situ (Gangi et al., 2015).

Conclusions
During our preliminary tests with the OA-ICOS, we found that the equipment was highly sensitive to changes in CO 2 concentrations. We found a non-linear response of the δ 13 C and δ 18 O values against changes in CO 2 concentration. Given the fact that laser-based CO 2 isotope analysers are deployed on site in combination with different gas sampling methods like automated chambers systems (Bowling et al., 2015) and hydrophobic gas-permeable membranes (Jochheim et al., 2018) for tracing various ecosystem processes, it is important to address this issue. Therefore, we developed a calibration strategy for correcting errors introduced in δ 13 C and δ 18 O measurements due to the sensitivity of the device against changing CO 2 concentrations. We found that the OA-ICOS measures stable isotopes of CO 2 gas samples with a precision comparable to conventional IRMS. The method described in this work for measuring CO 2 concentration and δ 13 C and δ 18 O values in soil air profiles using an OA-ICOS and hydrophobic gas-permeable tubes is promising and can be applied for soil CO 2 flux studies. As this set-up is capable of measuring continuously for longer time periods at a higher temporal resolution (0.05-0.1 Hz), it offers greater potential to investigate the isotopic identity of CO 2 and the interrelation between soil CO 2 and soil water. By using our measurement set-up, we could identify abiotic as well as biotic contributions to the soil CO 2 in the calcareous soil. We infer that degassing of CO 2 from carbonates due to weathering and evasion of CO 2 from groundwater may leave the soil CO 2 with a specific and distinct δ 13 C signature, especially when the biotic activity is rather low.