The Yatir Forest (31∘20′49′′ N, 35∘03′07′′ E; 650 m a.s.l.) is situated in the transition zone between subhumid and arid Mediterranean climates (Fig. S1 in the Supplement) on the edge of the Hebron mountain ridge. The ecosystem is a semiarid pine afforested area established in the 1960s and covering approximately 18 km2. The average air temperatures for January and July are 10.0 and 25.8 ∘C, respectively. Mean annual potential evapotranspiration (ET) is 1600 mm, and mean annual precipitation is 287 mm. Only winter (December to March) precipitation occurs in this region, creating a distinctive wet season, while summer (June to October) is an extended dry season. There are short transition periods between seasons, with a wetting season (i.e., fall) and a drying season (i.e., spring). The forest is dominated by Aleppo pine (Pinus halepensis Mill.), with smaller proportions of other pine species and cypress and little understory vegetation. Tree density in 2007 was 300 trees ha-1; mean tree height was 10.0 m; diameter at breast height (DBH) was ∼15.9 cm, and the leaf area index (LAI) was ∼1.5. The native background vegetation was sparse shrubland, which is dominated by the dwarf shrub Sarcopoterium spinosum (L.) Spach, with patches of herbaceous annuals and perennials reaching a total vegetation height of 0.30–0.50 m (Grünzweig et al., 2003, 2007). The root density range is 30–80 roots m-2 at the upper 0.1 m soil depth, falling to the minimum value (∼0 roots m-2) at 0.7 m soil depth (Preisler et al., 2019). Biological soil crust (BSC) is evident in the forest but is less than in the surrounding shrub by ∼40 % (Gelfand et al., 2012).

The soil at the research site is shallow (20–40 cm), reaching only 0.7–1.0 m; the stoniness fraction for the soil depth (0–1.2 m) is 15 %–60 %, and the rock cover of the surface ranges between 9 % and 37 %, as recently described in detail (Preisler et al., 2019); the soil is eolian-origin loess with a clay–loam texture (31 % sand, 41 % silt, and 28 % clay; density is 1.65±0.14 g cm-3) overlying chalk and limestone bedrock. Deeper soils (up to 1.5 m) are sporadically located at topographic hollows. While the natural rocky hill slopes in the region are known to create flash floods, the forested plantation reduces runoff dramatically to less than 5 % of annual rainfall (Shachnovich et al., 2008). Groundwater is deep (> 300 m), reducing the possibility of groundwater recharge due to negative hydraulic conductivity or of water uptake by trees from the groundwater.

An instrumented eddy covariance (EC) tower was erected in the geographical center of Yatir Forest, following the EUROFLUX methodology (Aubinet et al., 2000). The system uses a three-dimensional (3-D) sonic anemometer (Omnidirectional R3, Gill Instruments, Lymington, UK) and a closed path LI-7000 CO2/H2O gas analyzer (LI-COR Inc., Lincoln, NE, USA) to measure the evapotranspiration flux (ET) and net CO2 flux (NEE). EC flux measurements were used to estimate the annual scale of NEP by integrating half-hourly NEE values. The long-term operation of our EC measurement site (since 2000; see Rotenberg and Yakir, 2010) provides continuous flux and meteorological data with about 80 % coverage, which are subjected to U* nighttime correction and quality control, and gap filling is based on the extent of the missing data, as recently described in more detail in Tatarinov et al. (2016). A site-specific algorithm was used for flux partitioning into Re and GPP. Daytime ecosystem respiration (Re-d; in µmol m-2 s-1) was estimated based on measured nighttime values (Re-n; i.e., when the global radiation was < 5 W m-2), averaged for the first 3 half hours of each night. The daytime respiration for each half hour was calculated according to Eq. (3) (Maseyk et al., 2008a; Tatarinov et al., 2016): 3Re-d=Re-nα1βsdTs+α2βwdTa+α3βfdTa, where βs, βw, and βf are coefficients that correspond to soil, wood, and foliage, respectively; dTs and dTa are soil and air temperature deviations from the values at the beginning of the night; and α1, α2, and α3 are partitioning coefficients fixed at 0.5, 0.1, and 0.4, respectively. The βs, βw, and βf coefficients were calculated as follows: βs values were based on Q10 from the Grünzweig et al. (2009) study at the same site, where βs = 2.45 for wet soil (i.e., SWC in the upper 30 cm above 20 % vol.); βs = 1.18 for dry soil (i.e., SWC in the upper 30 cm equal to or below 20 % vol.); βf=3.15–0.036 Ta; and βw=1.34+0.46exp⁡(-0.5((DoY-162)/66.1)2), where DoY is the day of the hydrological year starting from 1 October. Finally, GPP was calculated as GPP = NEE–Re. Negative values of NEE and GPP indicated that the ecosystem was a CO2 sink.

Half-hourly auxiliary measurements used in this study included photosynthetic activity radiation (PAR; mol m-2 s-1), vapor pressure deficit (VPD; kPa), wind speed (m s-1), and relative humidity (RH; %), with additional measurements as described elsewhere (Tatarinov et al., 2016). Furthermore, the soil microclimatology half-hourly measurements were measured and calculated with soil chamber measurements, using the LI-8150-203 (LI-COR, Lincoln, NE), as described below, namely air temperature (Ta; ∘C) and relative humidity at 20 cm above the soil surface and soil temperature (Ts; ∘C) at a 5 cm soil depth using a soil temperature probe, as well as volumetric soil water content (SWC0-10; m3 m-3) in the upper 10 cm of the soil near the chambers, using the ThetaProbe model ML2x (Delta-T Devices Ltd., Cambridge, UK), which was calibrated to the soil composition based on the manufacturer's equations.

Soil CO2 fluxes (Rs) were measured with automated non-steady-state systems, using 20 cm diameter opaque chambers and a multiplexer to allow for simultaneous control of several chambers (LI-8150, -8100-101, -8100-104; LI-COR, Lincoln, NE). The precision of CO2 measurements in the chambers' air is ±1.5 % of the measurements' range (0–20 000 ppm). The chambers were closed on preinstalled PVC collars of 20 cm diameter, allowing for a short measurement time (i.e., 2 min), and positioned away from the collars for the rest of the time. Data were collected using a system in which air from the chambers was circulated (2.5 L min-1) through an infrared gas analyzer (IRGA) to record CO2 (µmol CO2 mol-1 air) and H2O (mmol H2O mol-1 air) concentrations in the system logger (1 s-1). Gap filling of missing data due to technical problems (i.e., 27 % of the data across the study period between November 2015 and October 2016) was based on the average diurnal cycle of each month.

The rates of soil CO2 flux, Rs (µmol CO2 m-2 s-1), were calculated from chamber data using a linear fit of change in the water-corrected CO2 mole fraction using Eq. (4) as follows: 4Rs=dCdt⋅vPsTaR, where dC/dt is the rate of change in the water-corrected CO2 mol fraction (µmol CO2 mol-1 air s-1), v is the system volume (m3), P is the chamber pressure (Pa), s is the soil surface area within the collar (m2), Ta is the chamber air temperature (K), and R is the gas constant (J mol-1 K-1). A measurement period of 2 min was used, based on preliminary tests to obtain the most linear increase of CO2 in the chambers with the highest R2.

Soil CO2 fluxes in the experimental plot were measured between November 2015 and October 2016 by means of three measurement chambers using 21 collars grouped in seven sites in the forest stand, with three locations (i.e., three collars) per site, based on different distances from the nearest tree (Dt). The collars were inserted 5 cm into the soil. Data were recorded on a half-hourly basis (48 daily records). The three chambers were rotated between the seven sites every 1–2 weeks to cover all sites and to assess spatial and temporal variations.

Upscaling of the collar measurements to plot-scale soil CO2 flux was carried out by grouping collars based on three locations (i.e., under trees (< 1 m from nearest tree; UT), in gaps between trees (1–2.3 m; BT), and in open areas (> 2.3 m; OA)), with one chamber taking measurements at each location, and estimating the fractional areas (∅) of the three locations based on mapping the sites according to the distances noted above, as previously done by Raz-Yaseef et al. (2010): 5Rs=RsOA⋅∅OA+RsBT⋅∅BT+RsUT⋅∅UT,6∅OA+∅BT+∅UT=1. The annual scale of Rs was derived from the upscaled chamber measurements (Eq. 5) based on daily records (48 half-hourly values) of spatially upscaled Rs.

Estimating the temperature sensitivity of Rs (Q10) was performed as described by Davidson and Janssens (2006) using a first-order exponential equation (see also Xu et al., 2015): 7Rs=aebTs, where Rs represents the half-hourly spatially upscaled time series of soil respiration flux (µmol m-2 s-1), Ts (∘C) is soil temperature at a 5 cm depth (upscaled spatially and temporally using the same method as for Rs), and a and b are fitted parameters. The b values were used to calculate the Q10 value according to the following equation: 8Q10=e10b.

The determination of different sources of soil CO2 efflux was based on linear mixing models (Lin et al., 1999) to estimate proportions of three main sources (autotrophic, heterotrophic, and abiotic), using isotopic analysis of soil CO2 profiles and soil incubation data from eight campaigns (January to September) during 2016, according to Eqs. (9)–(11). Partitioning of the monthly Rs values into components was done using a three-end-member triangular model for interpreting the δ13C and Δ14C values of CO2 flux; the three-end-member triangular corners are the autotrophic (Rsa), heterotrophic (Rh), and abiotic (Ri) sources of Rs. The δ13C and Δ14C isotope signatures of monthly Rs locate it inside the triangle (Fig. S2): 9δ13CRs=fsa⋅δ13Csa+fh⋅δ13Ch+fi⋅δ13Ci,10Δ14CRs=fsa⋅Δ14Csa+fh⋅Δ14Ch+fi⋅Δ14Ci,111=fsa+fh+fi, where f indicates the fraction of total soil flux (e.g., fh=Rh/Rs), while subscripts sa, h, and i indicate autotrophic, heterotrophic, and inorganic components, respectively. The three-equations system was used to solve the three unknown f fractions of the total soil flux based on empirical estimates of the isotopic end-members. Additionally, δ13C and Δ14C are the stable and radioactive carbon isotopic ratios, where δ13C = [([13C/12C]sample/ [13C/12C]reference) -1] ⋅1000 ‰, and the reference is the Vienna international standard (VPDB). Radiocarbon data are expressed as Δ14C in parts per thousand or per mil (‰), which is the deviation of a sample 14C/12C ratio relative to the OxI standard in 1950 (see Taylor et al., 2015), that is, Δ14C = [([14C/12C]sample/ (0.95⋅ [14C/12C]reference⋅exp⁡ [(y-1950)/8267])) -1] ⋅1000 ‰, where y is the year of sample measurements.

The δ13CRs was estimated monthly using the Keeling plot approach (Figs. S3 and S4; Pataki et al., 2003; Taneva and Gonzalez-Meler, 2011). Soil air was sampled using closed-end stainless-steel tubes (6 mm diameter) perforated near the tube bottom at four depths (30, 60, 90, and 120 cm). Samples of soil air were collected in pre-evacuated 150 mL glass flasks with high-vacuum valves, the dead volume in the tubing and flask necks having been purged with soil air using a plastic syringe equipped with a three-way valve.

Note that the Keeling plot approach is based on the two-end-member mixing model (see Review of Pataki et al., 2003), which often does not hold in soils because of variations in the δ13C values of source material with depth (see a recent example in Joseph et al., 2019). However, probably because of the very dry conditions at our study site, no change in δ13C with depths in the root zone is observed (±0.1 ‰ across the 35 cm depth profiles; Fig. S5), providing an opportunity to avoid this caveat; we must also conclude of course that the variations among the contributions of Rsa, Rh, and Ri do not change significantly with depth, permitting the use of the single set of isotopic signatures in Table 2. The soil CO2 samplings carried out therefore represented predominantly the mixing of atmospheric CO2 with a single integrated soil source signal, consistent with the Keeling plot approach.

The autotrophic (δ13Csa) end-member was estimated based on incubations during the sampling periods of excised roots, following Carbone et al. (2008). Fine roots (< 2 mm diameter) were collected, rinsed with deionized water, and incubated for 3 h in 10 mL glass flasks connected with Swagelok Ultra-Torr tee fittings to 330 mL glass flasks equipped with Louwers high-vacuum valves. The flasks were flushed with CO2-free air at room temperature close to field conditions. The CO2 was allowed to accumulate to at least 2000 ppm (∼2 h).

The heterotrophic (δ13Ch) end-member was estimated as in Taylor et al. (2015), and, similar to the root-incubation experiment, soil samples from the top 5 cm of the litter layer or 10 cm below the soil surface were collected, and roots were carefully removed to isolate heterotrophic components. Root-free soils were placed in 10 mL glass flasks and allowed to incubate for 24 h before being transferred to evacuated 330 mL glass flasks. The inorganic source (δ13Ci) end-member was estimated using 1 g of dry soil (ground to pass through a 0.5 mm mesh) placed in a 10 mL tube with a septum cap; then, 12 mL of 1 M HCl was added to dissolve the carbonate fraction, and the fumigated CO2 withdrawn from each tube was collected using a 10 mL syringe and injected into a 330 mL evacuated flask for isotopic analysis.

Radiocarbon estimates were based on the work of Carmi et al. (2013) at the same site, adjusted to the measured atmospheric 14C values during the study period (49.5 ‰; Carmi et al., 2013). The Δ14Csa and Δ14Ch end-members were estimated based on the assumption that they carry the 14C signatures of 4 and 8.5 years, respectively, older than the 14C signature of the atmosphere at the time of sampling, based on mean ages previously estimated (Graven et al., 2012; Levin et al., 2010; Taylor et al., 2015). The ratio Δ14Ci was obtained from Carmi et al. (2013). Monthly values of Δ14CRs were obtained using the linear equation of the regression line of the measured δ13C values of Rsa, Rsh, and Ri and the corresponding estimated Δ14C values (Fig. S2) and monthly δ13C values of Rs.

Isotopic analysis followed the methodology described in Hemming et al. (2005). The δ13C of CO2 in the air was analyzed using a continuous-flow mass spectrometer connected to a 15-flask automatic manifold system. An aliquot of 1.5 mL of air was expanded from each flask into a sampling loop on a 15-position valve (Valco, Houston, TX, USA). CO2 was cryogenically trapped from the air samples using helium as a carrier gas; it was then separated from N2O with a Carbosieve G (Sigma-Aldrich) packed column at 70 ∘C and analyzed on a Europa 20-20 isotope ratio mass spectrometer (IRMS; Sercon, Crewe, UK). The δ13C results were quoted in parts per thousand (‰) relative to the VPDB international standard. The analytical precision was 0.1 ‰. To measure [CO2], an additional 40.0 mL subsample of air from each flask was expanded into mechanical bellows and then passed through an infrared gas analyzer (LI-6262; LI-COR, Lincoln, NE, USA) in an automated system. The precision of these measurements was 0.1 ppm. Flasks filled with calibrated standard air were measured with each batch of 10 sample flasks; five standards were measured per 10 samples for δ13C analyses and four standards per 10 samples for [CO2] analyses.

Organic matter samples were dried at 60 ∘C and milled using a Wiley mill fitted with a size 40 mesh, and soil samples were ground in a pestle and mortar. Soils containing carbonates were treated with 1 M hydrochloric acid. Between 0.2 and 0.4 mg of each dry sample was weighed into tin capsules (Elemental Microanalysis Ltd., Okehampton, UK), and the δ13C of each was determined using an elemental analyzer linked to a Micromass Optima IRMS (Manchester, UK). Three replicates of each sample were analyzed, and two samples of a laboratory working standard cellulose were measured for every 12 samples. Four samples of the acetanilide (Elemental Microanalysis Ltd.) international standard were used to calibrate each run, and a correction was applied to account for the influence of a blank cup. The precision was 0.1 ‰.

TBCA (g C m-2 yr-1) was calculated following Giardina and Ryan (2002) for the study year (November 2015–October 2016) as follows: 12TBCA=Rs-Ralp+ΔCsoil, where Ralp is the annual aboveground litter production between November 2014 and October 2015, and ΔCsoil is the annual change in belowground total soil organic C. Litter production, not measured during the present study, was estimated based on values obtained by Masyk et al. (2008) for 2000–2006 (56 g C m-2 yr-1) and assumed to have increased in the study period (2014–2015) proportionally to the measured increase in leaf area index (LAI; 1.31 to 1.94; i.e., Ralp= [(1.94⋅56)/1.31] = 83 g C m-2 yr-1). For herbaceous litter production, three plots of 25 m2 were randomly selected in 2002 and harvested at the end of the growing season, total fresh biomass was weighed, and subsamples were used to determine dry weight and C content. Grünzweig et al. (2007) found that herbaceous litter production was close to the average rainfall for the specific year; this method was adapted in the current study for the period between November 2014 and October 2015. Since aboveground litter (Ralp; the sum of tree litter and herbaceous litter production) of a given year was mainly produced during that year but decayed during the following hydrological year, TBCA was based on the current year's Rs (2015–2016) and the previous year's Ralp (2014–2015). ΔCsoil was set constant as the average annual belowground carbon increase since afforestation (Qubaja et al., 2019).

Two-way ANOVA tests were performed at a significance level set at p=0.001 to detect significant effects of locations (OA, BT, and UT), sites, and their interactions on Rs and metrological parameters. Pearson correlation analysis (r) was used to detect the correlation between Rs and meteorological parameters. To quantify spatiotemporal variability in Rs, the coefficient of variation (CV %) was calculated as [(STDEV / Mean) ⋅ 100 %]. Heterogeneity was considered weak if CV % ≤ 10 %, moderate if 10 % < CV % ≤ 100 %, and strong if CV % > 100 %. All the analyses were performed using MATLAB software, Version R2017b (MathWorks, Inc., MA, USA).

The spatial variations in Rs across locations (distance from nearest tree) and sites (across the study area) are reported in Table 1, together with other measured variables. The results indicated an overall mean Rs value of 0.8±0.1 µmol m-2 s-1, with distinct values for the three locations. Rs was greater at UT locations than at the BT and OA locations by a factor of ∼2. The spatial variability among the locations was also apparent in the Rs daily cycle (Fig. 1), with clear differences between the wet season (November to April), when the UT showed consistently higher Rs values than at other locations by a factor of about 1.6, and the dry season, when the equivalent values differed by a factor of approximately 2.6. Note that the daily peak in Rs remained at midday in both the wet and dry seasons. Overall, the 21 collars showed moderate variations (CV = 55 %; Table 1); Rs was negatively correlated with distance from trees (Dt; r=-0.62; p<0.01) and with soil and air temperatures (Ts and Ta; r=-0.45; p<0.05) and positively correlated with soil water content and relative humidity (SWC and RH; r=0.50; p<0.05). The inverse correlation between Rs and distance from the nearest tree could be useful in considering the expected decline in stand density due to thinning and mortality (e.g., associated with a drying climate). For a first approximation, the results indicate that decreasing from the present stand density of 300 to 100 trees ha-1 and the resulting increase in mean distance among trees could result in decreasing ecosystem Rs by 11 %.

On the diurnal timescale, CO2 fluxes showed typical daily cycles (Fig. 1). As expected, on average, all CO2 fluxes were higher during the wet period compared to the dry season by a factor of ∼2. However, Rs and Re peaked around midday in both the wet and dry seasons, while the more physiologically controlled NEE and GPP showed a shift from midday (around 11:00–14:00 LT) to early morning (08:00–11:00 LT) in the dry season, with a midday depression and a secondary afternoon peak (Fig. 1d).

The temporal variations across the seasonal cycle are reported in Fig. 2, based on monthly mean values and exhibiting sharp differences between the wet and dry seasons. As previously observed in this semiarid site, all CO2 fluxes peak in early spring between March and April. The corresponding high-resolution data are reported in Fig. S6, which show also that the high winter (February) Rs rates were associated with clear days when photosynthetic active radiation (PAR) increased with air temperature, Ta. These data also show that, following rainy days, daily Rs values could reach 6.1 µmol m-2 s-1 (i.e., in the UT microsite; data not shown), although the average was 1.1±0.2 µmol m-2 s-1 during the wet period, which diminished by ∼55 % in the dry season to mean daily values of 0.5±0.1 µmol m-2 s-1. In spring (April), all CO2 fluxes peaked during the crossover trends of decreasing soil moisture content and increasing temperature and PAR (Fig. S6).

The temporal variations in the half-hourly values of Rs reflected changes in soil moisture at 0–5 cm depth and PAR (r=0.5 and 0.2, respectively; p<0.01) and negative correlations with Ts and RH (r=0.2 and 0.1, respectively; p<0.01). The variations in the integrated Rs showed a CV of 71 %, with the temporal variations dominated strongly by PAR (CV > 100 %), moderately by SWC (CV ∼85 %), and weakly by RH (CV ∼40 %; correlations and CV values were not included in figures and tables). Repeating the models applied by Grünzweig et al. (2009), the potential climatic factors that best predicted daily Rs shifted from SWC and PAR in the dry season to Ts and PAR in the wet season (Table S2). These equations explained 43 % and 70 % of the variation in Rs in the dry and wet seasons, respectively (Table S2). A reasonable forecast of the temporal variations in Rs (µmol m-2 s-1) at half-hourly values (R2=0.60, p<0.0001) was obtained based on SWC0-10 and Ts values across the entire seasonal cycle, based on 13Rs=0.05126⋅exp⁡(0.04274⋅Ts+28.51⋅SWC-74.44⋅SWC2). At the ecosystem scale, Re was characterized by high fluxes in the wet season and peak values of ∼2.4 µmol m-2 s-1 in February to April (Fig. 2; Table S1). Re fluxes rapidly decreased after the cessation of rain and reached the lowest values in the fall (September to October), with mean dry-period values of 0.5±0.1 µmol m-2 s-1 (Fig. 2, Table S1). GPP had a mean value of -1.8±0.4 µmol m-2 s-1, and daily NEE had a mean value of -0.5±0.3 µmol m-2 s-1 (Table S1 and Fig. S6), with the same seasonality for both (Fig. 2).

Figure 3 (see also Table 2) summarizes the seasonal variations in Rs and Re partitioning. The monthly Rsa and Rh were not significantly different but were significantly different from Ri (p<0.05). The Rsa/Rs ratios ranged from 0.32 to 0.46, the largest contribution occurring in early spring from February to April. The Rh/Rs fraction ranged between 0.33 and 0.45, being the highest during the wet season. The Ri/Rs fraction – the fraction of inorganic sources from the total soil respiration – ranged from 0.09 to 0.35, the largest contribution being in the driest period. The mean relative contributions of these components to Rs over the sampling campaigns are presented in Fig. 3a, but, on average, soil biotic fluxes were higher than abiotic fluxes by a factor of ∼4. Re partitioning showed an average increase in Rf/Re from 25 % in the wet season to 54 % in the dry season and a decline in Rs/Re from 75 % to 46 % on average from the wet to the dry season, respectively, which reflected a seasonal change of Rf in the wet season to peak values in the dry season (Fig. 3b). Both the highest and lowest Rs fractions (∼0.74 and nearly 0.34) along the seasonal cycle were associated with low total Re fluxes, that is, in the fall before the Rf peak in the spring and in the summer, when physiological controls limited water loss (Fig. 2).

On an annual timescale, estimates of CO2 flux components based on EC measurements resulted in annual values of GPP, NPP, Re, and NEP of 655, 282, 488, and 167 g C m-2 yr-1, respectively (Tables 3 and S1). On average across the measurement period, Rs was the main CO2 flux to atmosphere, making up 60±6 % of Re (295±4 g C m-2 yr-1; Tables 3 and S1), and Rf was another significant component accounting for 40±6 % of Re (Fig. 3b), which reflected the low-density (300 trees ha-1) nature of the semiarid forest. As indicated above, Re partitioning showed a decrease in Rs/Re and an increase in Rf/Re from winter to summer, which is clearly apparent in Fig. 3b. On an annual scale, during the study period, estimates of Rf, Rsa, Rh, and Ri values were 194±36, 119±21, 115±20, and 61±6 g C m-2 yr-1, respectively. These rates of respiration fluxes translated at the ecosystem scale to Re / GPP of ∼75 %, lower than observed in other ecosystems (Table S3) and leading, in turn, to high ecosystem CUE of 0.43.

Using the site records of nearly 20 years, long-term trends in GPP, NPP, Re, and NEP were examined. Soil respiration and its partitioning could not be similarly monitored continuously, but combining the present results with the 2001–2006 values obtained by Grünzweig et al. (2009) and Maseyk et al. (2008a) provided a basis for estimating the long-term trends in soil respiration. Notably, no clear or significant trend over time was observed in any of the canopy-scale continuously monitored fluxes, but, because of relatively large interannual variations, associated mainly with those in precipitation (see Qubaja et al., 2020), it is likely that the relative contributions of the different fluxes, expressed as ratios in Table 3, provide a more robust perspective of the long-term temporal changes in the ecosystem functioning. The results presented in Table 3 reflect the long-term growth of the forest, with a relatively large increase in LAI, but the TBCA remained around 40 %. The results also indicated little change in the total soil respiration, Rs, component, (as a fraction of Re or GPP) but a general shift from the autotrophic components to the heterotrophic component (i.e., Rh). This was reflected in the decreasing ratio of the autotrophic components (i.e., Rsa, Rl, and Rw) and the increasing ratio of Rh to Re (Table 3) across the 13-year observation period (2003 to 2016).