Effects of photorespiration, the cytochrome pathway, and the ...

30 - 12 ANGERT ET AL.: BIOLOGICAL EFFECTS ON THREE O 2 ISOTOPESof any isotopic species ( y O 16 O, y = 16,17,18). Flux y Prepresents global photosynthetic production and y U isglobal uptake of any isotopic species. In the model, thetransfer of O 2 to the stratosphere involves no fractionationand O 2 reentering the troposphere is fractionated by x a s (asin the work of Blunier et al. [2002]). The relationshipbetween 18 a s and 17 a s is given by q s . The parameter x a sdescribes the sum of stratospheric effects on O 2 isotopes,and does not represent any single physical process. Theglobal average isotopic ratio of oxygen produced fromwater by photosynthesis will be noted as x R AW . The isotopiccomposition of leaf water (which is the substrate forphotosynthesis on land) is controlled by the hydrologicalcycle and evapotranspiration. However, since informationon the different q’s in the hydrological cycle is scarce[Meijer and Li, 1998; Miller, 2002] and since there is noinformation about the q of evapotranspiration, we willassume an average isotopic ratio for all the oxygenproduced by photosynthesis and will not deal directly withdifferences in the substrate water. This average isotopic ratiois assumed to be constant in time. The average globalfractionation in biological uptake is given by x a b and q b .Wealso assume 16 T = 16 S and 16 P = 16 U. Including theseformulations in equation (23) gives16 S x R T ð1Þþ 16 P x R 16 AW P x a b x R T ¼ 0; ð24Þx a swhere x can be 17 or 18, and the subscript ‘‘T’’ stands fortropospheric oxygen.[68] According to equation (1), the 17 D of tropospheric O 2relative to oxygen produced by photosynthesis is 17 D T AW ¼ ln 17 R T = 17 R AWC ln 18 R T = 18 R AW 10 6 :ð25Þ[69] The value of 17 D T-AW is always negative, sincetropospheric O 2 is depleted in 17 O. Equations (24) and(25) give17 D T AW ¼ ln16 P16P 17 a16 b S ð17 a s 1Þ16 PC ln16P 18 : ð26Þa16 b S ð18 a s 1Þ[70] The average global q associated the biological uptakewill be noted as q b . The difference in the tropospheric 17 Dresulting from a change of the global average q b from q b1 toq b2 is given by substituting equation (26) with equation (5),17 D T W ðq b1 Þ17 D T W ðq b2 Þ¼ln!16 P ð 18 a b Þ qb2 16 S ð17 a s 1Þ:16P ð18 a b Þ qb1 16 S ð17 a s 1Þð27Þ5.7.1. Implications for the 17 D in Ice Cores[71] The new q values reported here may change theinterpretation of 17 D measured in ice cores. To demonstratethis, we will use the Last Glacial Maximum (LGM) as acase study. According to Blunier et al. [2002], the ratio ofphotorespiration to terrestrial production was 43% in theLGM relative to 38% at the modern biosphere. This differenceis due to relatively low atmospheric CO 2 concentrationin the LGM (since the residence time of O 2 in the atmosphereis 1200 years, the preindustrial CO 2 levels are usedfor present), which was partially compensated by anincrease in the relative rate of photosynthesis in C4 plants.We can apply equation (27) or Figure 6 to estimate theeffect of this different photorespiration rate on the LGM’s17 D T-AW . By assuming that the ratio of terrestrial productionto marine production is 1.7 [Blunier et al., 2002] and that nophotorespiration takes place in the oceans, we estimate thatthe ratio of photorespiration to global production is 24% atpresent and was 27% in the LGM. Using the q values fordark respiration and photorespiration we are reporting here,and assuming that 10% of dark respiration is through theAOX, we calculated that the q b in the LGM was lower byabout 0.00025 from the present value. Assuming the AW is40% seawater and 60% leaf water and using an averageLn 18 O value of 23% versus the atmosphere for the formerand +6% versus SMOW for the later [Gillon and Yakir,2001], we calculated the average global fractionation ( 18 e b )as 19.4 (0.4 [ 23] + 0.6 [ 23 + 6] = 19.4). Inaddition, we used a 16 P/ 16 S ratio of 0.0097 [Luz et al.,1999].[72] Using equation (27), with the above parameters wecalculated that the effect of higher photorespiration rates inthe LGM is expected to cause the 17 D of the LGM troposphereto be lower by 5 ± 2 per meg relative to the presenttroposphere (the ±2 per meg error margin is derived fromthe uncertainty in the q values). This result agrees with thesimple graphical solution presented at Figure 6 (0.00025 19.4 = 5 per meg). This 5 per meg estimate is insensitive tomost of the assumptions in the calculation, except theassumption on the variations in the global rate of photorespiration.Hence, this calculation gives the order ofmagnitude of the variation in 17 D that are caused byvariations in global photorespiration. To estimate whetheror not this 5 per meg signal is significant, it should becompared to the 17 D signal generated by changes in globalproductivity.[73] The D 17 O of the LGM troposphere relative to thepresent troposphere was measured in air from ice cores andwas found to be +38 per meg [Blunier et al., 2002; Luz etal., 1999]. This corresponds to a 17 D value of +43 per meg(with C of 0.516). This value is affected by a lower rate ofmass-independent stratospheric processes that resultedfrom lower CO 2 concentration in the LGM atmosphere.This lower rate resulted in 17 D of the LGM tropospherehigher by 55 per meg than the present value (under theassumptions of linear dependence of the stratosphericprocesses rate on CO 2 concentration and constant ozoneconcentration [Luz et al., 1999]. There is a 12 per megdifference between the +55 per meg expected from thelower rate of stratospheric processes and the +43 per megactually measured. This difference was considered inprevious papers [Blunier et al., 2002; Luz et al., 1999]to indicate lower global productivity in the LGM. However,as shown above, about half of this 12 per meg