Determination of respiration and photosynthesis fractionation 1 coefficients for atmospheric dioxygen inferred from a vegetation-soil-2 atmosphere analog of the terrestrial biosphere in closed chambers 3 4

17 The isotopic composition of dioxygen in the atmosphere is a global tracer which depends on the 18 biosphere flux of dioxygen toward and from the atmosphere (photosynthesis and respiration) as well 19 as exchanges with the stratosphere. When measured in fossil air trapped in ice cores, the relative 20 concentration of O, O and O of O2 can be used for several applications such as ice core dating and 21 past global productivity reconstruction. However, there are still uncertainties about the accuracy of 22 these tracers as they depend on the integrated isotopic fractionation of different biological processes 23 of dioxygen production and uptake, for which we currently have very few independent estimates. 24 Here we determined the respiration and photosynthesis fractionation coefficients for atmospheric 25 dioxygen from experiments carried out in a replicated vegetation-soil-atmosphere analog of the 26 terrestrial biosphere in closed chambers with growing Festuca arundinacea. The values for O 27 discrimination during soil respiration and dark respiration in leave are equal to -12.3 ± 1.7 ‰ and -19.1 28 ± 2.4 ‰, respectively. We also found a value for terrestrial photosynthetic fractionation equal to +3.7 29 ± 1.3 ‰. This last estimate suggests that the contribution of terrestrial productivity in the Dole effect 30 may have been underestimated in previous studies. 31 32 https://doi.org/10.5194/bg-2021-324 Preprint. Discussion started: 16 December 2021 c © Author(s) 2021. CC BY 4.0 License.


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In addition to the use of δ 18 Oatm, the combination of δ 17 O and δ 18 O of O2 provides a way to quantify 76 variations in past global productivity (Luz et al., 1999). This method relies on the fact that O2- In these studies, the underlying assumption is that the fractionation factor associated with O2

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The sealing of the closed chamber was checked before each experiment using helium.

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To ensure atmospheric pressure stability in the closed chamber, a pressure compensation system, 161 made of two connected 10 liters gas tight bags (multi-layers foil bags, Restek, USA), was installed. Each 162 bag was half full of atmospheric air, the first one was installed in the closed chamber while the second 163 one was outside this chamber. This way, each bag inflates or deflates in response to pressure variation 164 either due to O2 or CO2, uptake or release. The pressure difference between the closed chamber and 165 the atmosphere was regularly measured using a differential sensor (FD A602-S1KAlmemo, Ahlborn,

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Finally, the enclosed air was mixed and considered homogeneous using seven brushless fans.  separate the dioxygen and the argon from the dinitrogen. After separation of the dioxygen and argon 208 from helium, the gas was collected in a stainless-steel manifold immersed in liquid helium at -269°C.

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After collection, the samples were analyzed by the IRMS previously mentioned for leaf water analyses.

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The following ratios were measured: 18      We detail below how we used the results from our experiments to quantify the associated 261 fractionation coefficients. Notations used below are gathered in Table 1.

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The isotopic fractionation coefficient of oxygen is expressed through the fractionation coefficient α.

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Since argon is an inert gas, we can link

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During the 4 sequences, the respiration activity led to a decreasing level of the O2 concentration 395 measured by the optical sensor or through the δO2/Ar evolution from IRMS measurements (Fig. S1).

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The comparison of the evolution of the O2 concentration during the different sequences showed that 397 respiratory fluxes were different with a maximum factor of 4 between the different sequences ( Fig.   398   S1). In parallel to the decrease in O2 concentration, the  18 O increased as expected since respiration 399 preferentially consumes the lightest isotopes: over the 51 days of the 2 nd soil respiration sequence, we 400 observed a linear decrease of oxygen concentration by more than 5% while δ 18 O increased by 8 ‰ 401 (Fig. 2). A Mann-Kendall test (95%) showed that the Δ 17 O of O2 does not show any trend within 95% 402 over the 4 sequences (Fig. S2).

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following Eq. (10) (Fig. 3). The slope of the corresponding regression line provided the isotopic 407 discrimination 18 _ and hence the fractionation coefficient 18 _ for each sequence. (Table S2). It could be observed that despite differences in respiratory fluxes for the different sequences, the relationship between δ 18 O of O2 and O2 concentration (or O2/Ar) and hence the

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During the night periods, when only respiration occurred, we observed a decrease in O2 concentration 432 by 1% within 3 days and a δ 18 O increase by 1 ‰ during the same period (Fig. 4). The evolution was  (Table S3).  (Table S6).

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The value of isotopic fractionation associated with the light period of period 1 of sequence 1 appeared 487 clearly out of range. Following the Dixon's outlier detection test (Dixon, 1960), this value was 488 considered an anomaly (likelihood > 99 %) and was removed from further analysis (Table S7). The

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individual determination is presented on Table S7.

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The average 18 ℎ ℎ is + 3.7 ± 1.3‰ which goes against the classical assumption that 537 terrestrial photosynthesis does not fractionate (Guy et al., 1993). This value for the isotopic 538 discrimination is smaller than the photosynthetic fractionation in marine environment 18  = + 6 ‰