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

GLOBAL BIOGEOCHEMICAL CYCLES, VOL. 17, NO. 1, 1030, doi:10.1029/2002GB001933, 2003Effects of photorespiration, the cytochrome pathway, and thealternative pathway on the triple isotopic composition ofatmospheric O 2Alon Angert 1Institute of Earth Sciences, Hebrew University of Jerusalem, Jerusalem, IsraelShimon RachmilevitchDepartment of Plant Sciences, Institute of Life Sciences, Hebrew University of Jerusalem, Jerusalem, IsraelEugeni Barkan and Boaz LuzInstitute of Earth Sciences, Hebrew University of Jerusalem, Jerusalem, IsraelReceived 21 May 2002; revised 21 November 2002; accepted 21 November 2002; published 25 March 2003.[1] The triple isotopic composition of atmospheric O 2 is a new tracer used to estimatechanges in global productivity. To estimate such changes, knowledge of the relationshipbetween the discrimination against 17 O and the discrimination against 18 O is needed.This relationship is presented as q =ln( 17 a)/ln( 18 a). Here, the value of theta wasevaluated for the most important processes that affect the isotopic composition ofoxygen. Similar values were found for dark respiration through the cytochrome pathway(0.516 ± 0.001) and the alternative pathway (0.514 ± 0.001), and slightly higher valuewas found for diffusion in air (0.521 ± 0.001). The combined effect of diffusion andrespiration on the atmosphere was shown to be close to that of dark respiration. Thevalue we found for photorespiration (0.506 ± 0.005) is considerably lower than that ofdark respiration. Our results clearly show that the triple isotopic composition of theatmosphere is affected by the relative rates of photorespiration and dark respiration.Also, we show that closing the current global isotopic balance will enable the estimationof the current global rate of photorespiration. Using the Last Glacial Maximum as a casestudy, we show that variations in global rate of photorespiration affected the tripleisotopic composition in the past. Strong fractionations measured in illuminated plantsindicated that the alternative pathway is activated in the same conditions that favor highrate of photorespiration. This activation suggests that the global rate of the alternativepathway is higher than believed thus far, and may help to close the gap between thecalculated and measured Dole Effect. INDEX TERMS: 0315 Atmospheric Composition andStructure: Biosphere/atmosphere interactions; 0365 Atmospheric Composition and Structure: Troposphere—composition and chemistry; 1040 Geochemistry: Isotopic composition/chemistry; 1615 Global Change:Biogeochemical processes (4805); KEYWORDS: triple isotope, oxygen isotopes, Dole Effect,photorespiration, LGM, 17 OCitation: Angert, A., S. Rachmilevitch, E. Barkan, and B. Luz, Effects of photorespiration, the cytochrome pathway, and thealternative pathway on the triple isotopic composition of atmospheric O 2 , Global Biogeochem. Cycles, 17(1), 1030,doi:10.1029/2002GB001933, 2003.1. Introduction[2] The isotopic composition of atmospheric O 2 fluctuatedover glacial-interglacial timescales. Variations in theratios of all three stable O 2 isotopes were applied to estimatechanges in global gross production [Luz et al., 1999]. The1 Now at Center for Atmospheric Sciences, University of California,Berkeley, Berkeley, California, USA.Copyright 2003 by the American Geophysical Union.0886-6236/03/2002GB001933$12.00difference between the 18 O/ 16 O ratio of atmospheric O 2 andseawater H 2 O is known as the Dole Effect. Variations in theDole Effect, are used to estimate changes in the ratio ofmarine to terrestrial productivity in the past [Bender et al.,1994]. To improve interpretation of past atmosphericchanges, it is necessary to understand the basic processesaffecting atmospheric isotopic composition.[3] Atmospheric O 2 is produced by photosynthesis withoutisotopic fractionation from the substrate water [Guy etal., 1993]. Thus, marine photosynthesis results in productionof O 2 with identical isotopic composition to seawater.30 - 1

30 - 2 ANGERT ET AL.: BIOLOGICAL EFFECTS ON THREE O 2 ISOTOPESThe substrate water in leaves is enriched in 17 O and 18 Othrough evapotranspiration [Dongmann, 1974]. As a result,terrestrial photosynthesis produces O 2 that, on average, isenriched in these heavy isotopes. In both marine andterrestrial systems, the major cause of the heavy isotopeenrichment is preferential removal of 16 O (over 17 O and18 O) by biological uptake mechanisms [Lane and Dole,1956]. This preferential removal is mass-dependent and therelative increase in 17 O/ 16 O is about half (0.52) of therelative increase in 18 O/ 16 O (Figure 1). In addition toevapotranspiration and biological uptake, atmospheric O 2is also affected by photochemical reactions in the stratosphere[Luz et al., 1999]. These reactions preferentiallytransfer 17 O and 18 O from O 2 to CO 2 in a mass-independentway. Hence, in these reactions the ratio of the increase in17 O/ 16 O over 18 O/ 16 O in CO 2 is equal to 1 or larger[Lämmerzahl et al., 2002; Thiemens, 1999]. As a result,O 2 becomes mass-independently depleted in the heavyisotopes [Bender et al., 1994]. At the surface of the Earth,the 17 O depletion of atmospheric O 2 is removed byexchange through photosynthesis and uptake. Therefore,the net result is depletion in 17 O of atmospheric O 2 incomparison to O 2 affected by biological uptake alone(Figure 1). The magnitude of this depletion depends onthe ratio of production of 17 O depleted O 2 in the stratosphere,and its destruction by biological cycling. Thus, pastvariations in the 17 O depletion can be used to infer changesin global biological productivity.[4] In order to apply this approach, it is necessary that therelationship between 17 O/ 16 O and 18 O/ 16 O to be known withgreat accuracy. However, the triple isotope relationships indifferent physical-chemical mass-dependent processes,varies slightly about 0.52 [Luz and Barkan, 2000; Matsuhisaet al., 1978; Young et al., 2002]. If this is also the casefor O 2 consumption by different biological mechanisms,then a change in the ratio of uptake by the different pathwayswill result in variations in the magnitude of the 17 Odepletion. As a consequence, past variations in this parameterwill be misinterpreted.[5] In interpretations of variations in the Dole Effect it isusually assumed that the major driving force is changes inthe terrestrial to marine production rate (terrestrial producedO 2 is heavier than marine produced, because of the enrichmentof leaf water in evapotranspiration). An impliedassumption is that the enrichment due to biological uptakeremains constant. In the present paper, we explore thepossibility that time variations in Dole Effect were alsoaffected by changes in the relative rates of O 2 uptake bydifferent biological mechanisms, that have different fractionationsassociated with them. Consumption (or uptake) ofO 2 in plants occurs via four mechanisms: ordinary respirationthrough the cytochrome oxidase pathway (COX), respirationby the alternative oxidase pathway (AOX), Mehlerreaction, and photorespiration. The first two processes takeplace in light as well as in dark conditions, while the lattertwo occur only under illumination. In addition, O 2 diffusionin air precedes biological uptake in some cases. The discriminationagainst 18 O associated with the AOX (28%)isstronger than the discrimination associated with other mechanisms(COX 18%, Mehler 15%, photorespiration Figure 1. Schematic plot (not to scale) of 17 O and 18 Orelative to 16 O variations. Oxygen is produced by photosynthesisand it is fractionated by biological uptake in amass-dependent way, and by stratospheric photochemistryin a mass-independent way. The balance between these twotypes of fractionation control the triple isotopic compositionof atmospheric O 2 .22%). Hence, an increase in the proportion of O 2 uptake byAOX will significantly intensify the Dole Effect.[6] To improve the interpretation of past variations in thetriple O 2 isotopes and the Dole Effect, we have investigatedthe effects of biological mechanisms in controlled laboratoryexperiments. In one set of experiments, we studiedtriple isotope variations due to uptake alone by COX andAOX, and due to removal by diffusion. In a second set ofexperiments, conducted in an airtight terrarium, we analyzedthe combined effects of production and consumptionby both dark and light reactions in steady state betweenphotosynthesis and uptake.2. Triple Isotope Systematics2.1. Terminology[7] Atmospheric O 2 is depleted in 17 O due to stratosphericreactions. Hence, O 2 that is produced by photosynthesiswill have 17 O excess with respect to atmospheric O 2 .This 17 O excess was formulated in previous papers [Luz andBarkan, 2000; Luz et al., 1999] as D 17 O=d 17 O 0.521 d 18 O. However, this definition is somewhat problematicsince it contains a linear approximation to the mass-dependentfractionation, and because the parameter 0.521 (thatrepresents mass-dependent processes) depends on thechoice of reference material. Here, we follow the suggestionof Miller [2002] for a new definition of the 17 O excess thatis free of these problems:17 R18 R17 D ln C ln : ð1Þ17R18 ref R ref

ANGERT ET AL.: BIOLOGICAL EFFECTS ON THREE O 2 ISOTOPES 30 - 3[8] The new definition of the 17 O excess will be noted as17 D to distinguish it from the old definition (D 17 O). 17 Rstands for the isotope ratio 17 O 16 O/ 16 O 2 , 18 R stands for theisotope ratio 18 O 16 O/ 16 O 2 , and the subscript ‘‘ref’’ stands forthe reference. It should be noted that 17 D is a calculatedvalue that depends both on the standard used and on thevalue of the parameter C. This parameter represents thetriple isotope relationships in mass-dependent processes.However, as was mentioned in the introduction, theserelationships vary slightly about 0.52 for different processes.As a result, The chosen C value can only correspond to therelationship associated with one of the processes, preferablythe main one. In this study we choose to use a value of C =0.516, which corresponds to ordinary dark respiration(COX), the most wide spread global O 2 uptake mechanism.[9] In the old definition of the 17 Oexcess(D 17 O) thevariations in the isotopic composition were described on ad 17 O versus d 18 O plot. To describe the variation in the isotopiccomposition according to the new definition, two new termsthat will be used for the axes of such plot are suggested,Ln 17 O lnLn 18 O ln17 R17R ref18 R18R ref¼ ln d 17 O þ 1Þ¼ ln d 18 O þ 1Þ:[10] Another advantage of 17 D,Ln 17 O and Ln 18 O is thatthey are dimensionless. Hence, the value of a sample versusthe reference exactly equals to the value of the referenceversus the sample with the opposite sign (this is not true forD 17 O, d 17 O and d 18 O). For convenience, Ln 17 O and Ln 18 Oare expressed as permil (%) and 17 D as per meg deviationsfrom the standard.[11] The slope of a line passing between two points (Aand B) on a Ln 17 O versus Ln 18 O plot, will be defined as l,as in previous work [e.g., Miller, 2002], ln 17 R A = 17 R Bl : ð4Þln 18 R A = 18 R B[12] Following Mook [2000], the relationship between thediscrimination against17 O and 18 O (relative to 16 O) ispresented asq ln ð17 aÞlnð 18 aÞ ; ð5Þwhere 17 a and 18 a are the fractionation factors x R p / x R s . Thesubscripts ‘‘p’’ and ‘‘s’’ stand for ‘‘product’’ and ‘‘substrate,’’respectively (x can be 17 or 18).[13] Blunier et al. [2002] used an additional relationshipbetween the discriminations,g ð2Þð3Þ17 a 118a 1 ¼ 17 e18e ; ð6Þwhere x e is x a 1.[14] The terms q and g describe the inherent relationshipbetween the discriminations, but cannot be measureddirectly. What can be measured is the slope (l) that iscontrolled by these relationships and the processes thattake place in the system. In the special case of a systemat steady state, in which production equals uptake, theslope (l) is equal to q (see section 2.2). In the specialcase of a system where only uptake takes place, the slope(l) is equal to g (see section 2.3). Hence, by conductingexperiments in such systems, q or g could be calculatedfrom the measured l. In other systems (for example, inones where mixing takes place) the slope can be differentfrom both q and g. The difference between the value of qand g is small (0.003), but since the 17 O excessvariations in O 2 are minute such a difference cannot beneglected.[15] This paper deals with the effect of the biosphere onthe triple isotopic composition of the atmosphere. Since thebiosphere-atmosphere system is close to production-uptakesteady state, we chose to use q to report the relationshipbetween discriminations.2.2. Production-Uptake Steady State[16] In this study, we used a terrarium similar to that ofLuz et al. [1999] to estimate the q value associated withphotorespiration. To understand how the q value of aprocess can be evaluated from such experiment, we willdiscuss the triple isotope systematics of a closed system inbiological production-uptake steady state. Such a systemcan also be used as a model for the Earth atmosphere andbiosphere without the effects of stratospheric photochemistryand the hydrological cycle.[17] The starting point for such a system is oxygen that isproduced by photosynthesis in photosystem 2 from water(noted as ‘‘W’’) without isotopic fractionation. Uptake byrespiration causes enrichment in the heavy isotopes. Sincethe fractionation in biological uptake is mass dependent, theenrichment in 17 O will be about half of the enrichment in18 O. In steady state, depending on the value of q, the systemwill reach either the point that is marked as ‘‘BSS1’’(Biological Steady State 1) or ‘‘BSS2’’ in Figure 2. TheLn 18 O value of the O 2 in steady state relative to thesubstrate water is illustrated by the horizontal distancebetween ‘‘BSS1’’ (or ‘‘BSS2’’) and ‘‘W,’’ and is thesystem’s equivalent of the ‘‘Dole Effect.’’[18] A system in production-uptake steady state can beanalyzed graphically as in Figure 2, or, more rigorously, bya model that deals with the fluxes of the three isotopicspecies. Such a one-box model of production-uptake steadystate can be formulated as follows:y P ¼ y U;where y P is production rate of y O 16 O (y can be 16, 17 or18), and y U is uptake rate of y O 16 O.[19] Equation (7) can be rewritten as follows:ð7Þ16 P x R W ¼ 16 U x R BSS x a; ð8Þwhere x can be 17 or 18, the subscript ‘‘W’’ stands foroxygen produced by photosynthesis, and ‘‘BSS’’ stands forthe system air in biological steady state.

30 - 4 ANGERT ET AL.: BIOLOGICAL EFFECTS ON THREE O 2 ISOTOPES[23] According to equation (1), the 17 D of the photosyntheticallyproduced oxygen is17 R W17 R W17 D W ¼ ln C ln : ð11Þ17R17 ref R ref[24] Subtracting equation (11) from equation (10), andusing equation (5) yields17 D 17 BSS D W ¼ ln18 a C q: ð12ÞFigure 2. Schematic plot (not to scale) of Ln 17 O versusLn 18 O for a closed system in production-uptake steadystate. Point ‘‘W’’ represents oxygen that is produced byphotosynthesis in photosystem 2 from water, and ‘‘BSS1’’and ‘‘BSS2’’ represent biological steady states. The slope ofthe line that connects ‘‘BSS1’’ and ‘‘W’’ is equal to q 1 andthe slope of the line connecting ‘‘BSS2’’ and ‘‘W’’ is equalto q 2 (q =ln( 17 a)/ln( 18 a)). When q of the system equals q 1the system reaches the steady state condition indicated by‘‘BSS1’’ with an 17 O excess of 17 D BSS1 . Since C equals toq 1 , 17 D BSS1 equals to the 17 D W . When q of the system equalsq 2 the system reaches steady state indicated by ‘‘BSS2’’with 17 O excess of 17 D BSS2 that is lower than 17 D BSS1 . Thehorizontal distance between ‘‘BSS’’ and ‘‘W’’ is thesystem’s equivalent of the global Dole Effect (in Ln 18 Oterms). The difference between 17 D BSS1 and 17 D BSS2 isgiven by the system’s ‘‘Dole Effect’’ times the differencebetween q 1 and q 2 .[20] Rearranging equation (8) and substituting it intoequation (4) givesl BSS ¼ ln ð17 aÞln 18 ¼ q: ð9Þð aÞ [21] Hence, the slope of the line connecting the O 2 of abiological system in production-uptake steady state, and theO 2 produced from the substrate water, is equal to the valueof q. However, as we will show in section 2.3 below, thevalue of l is not always equal to the value of q.[22] The 17 D in production-uptake steady state can befound by substituting equations (7) and (8) into equation (1),[25] When q equals C, the 17 D value of the air in steadystate equals to that of the oxygen that is produced from thesubstrate water. However, if q is different from C, then17 D BSS is controlled both by q and 18 a.[26] Currently, the value of 17 D W cannot be measureddirectly with sufficient accuracy. However, this value canbe estimated by conducting a terrarium experiment inwhich the q value of the uptake process is known (thisknown q will be noted as q 1 ). In such an experiment, thevalue of 17 D in steady state ( 17 D BSS1 ) will be identical tothat of 17 D W , if we choose C = q 1 . By conducting anadditional experiment in which the q value of the uptakeprocess is unknown (this unknown q will be noted as q 2 ),we can calculate the value of q 2 from the 17 O excess insteady state of this experiment ( 17 D BSS2 ), by rearrangingequation (12),ð17 D 17 BSS 2 D W Þq 2 ¼ q 1ln 18 : ð13Þð aÞ[27] The value of ln( 18 a) for the system can be found byrearranging equation (8),ln 18 a ¼ ln 18 R BSS = 18 R W : ð14Þ[28] The right-hand side of equation (14) is the Ln 18 Ovalue of ‘‘BSS’’ versus ‘‘W,’’ which is equivalent to theterrarium Dole Effect in Ln 18 O terms (the value of the DoleEffect in d 18 O terms is greater by 0.3% if SMOW is thereference and lower by 0.3% if atmospheric O 2 is thereference). Hence, the calculation of q 2 from equation (13)is identical to the solution presented graphically in Figure 2.[29] To estimate q values by the method describe above,the q value of at least one process (q 1 ) must be knownindependently. This value can be found by conductingexperiments in systems in which there is O 2 uptake butno production.2.3. O 2 Removal Only[30] In our dark respiration and binary diffusion experiments,O 2 is only removed from the system. The change inisotopic composition in such experiments follows theRaleigh distillation equation,1717 R W = 17 a18 R W = 18 aD BSS ¼ ln C ln : ð10Þ17R18 ref R refx e ¼ ln x R= x R 0; ð15Þln f

ANGERT ET AL.: BIOLOGICAL EFFECTS ON THREE O 2 ISOTOPES 30 - 5where x R 0 is the initial isotope ratio x O 16 O/ 16 O 2 and f is theremaining 16 O 2 fraction. Substituting equation (15) intoequation (4) yieldsl R ¼17 e¼ g; ð16Þ18ewhere the subscript R stands for Raleigh distillation.Equation (16) shows that in a system where there is onlyuptake of oxygen, the value of l is different from thevalue of q. The value of q in ‘‘removal only’’ experimentscan be calculated by substituting equation (16) intoequation (5),q ¼ ln ð 1 þ l R 18 eÞlnð1 þ 18 eÞ : ð17Þ[31] The range of 18 e in biological uptake is 15 to32% and the value of q in mass-dependent fractionationis about 0.5. Hence, l R will be larger by 0.002–0.004 than qin most systems. For O 2 samples close in composition to thestandard O 2 , using slightly different C values will notsignificantly affect the calculated 17 D. However, for thepurpose of the present study where comparisons betweensteady state O 2 and its substrate water are necessary, a smalldifference in parameter C becomes significant. To illustratethe importance this point, consider a water sample havingLn 18 O= 23% and Ln 17 O= 11.7% with respect to airO 2 . Using equation (1) and C = 0.52, 17 D of this sample iscalculated as 260 per meg. Similar calculation with C =0.524 yields 17 D of 352 per meg.3. Experimental Methods3.1. Dark Respiration Experiments[32] Axenic cultures of Lemna gibba, a small waterplant, were grown in a nutrient solution. Illumination frommixed fluorescent/incandescent lamps was at an intensityof 200 ± 20 mE m 2 s 1 photosynthetic photon flux (PPF),at the level of the fronds, for 12 h d 1 . Air temperaturewas 24°C and 19 ± 0.5°C in the light and dark periods,respectively.[33] The seeds of Orobanche aegyptica, an obligatoryroot parasite, were placed in empty tea bags and suspendedfor sterilization in 1% solution of NaOCl for 5 min. Theseeds were then rinsed thoroughly for 20 min in steriledistilled H 2 O. After the rinse, the seeds were activated by 3days imbibition in 5 ml of H 2 O in 9-cm-diameter Petridishes with one layer of Whatman paper, to which a mixtureof antibiotics was added. The mixture contained 50 mg eachof streptomycin and penicillin G and 25 mg of chloramphenicol,in order to prevent the development of bacterialinfections during the imbibition period.[34] To determine the q value associated with the COX,we inhibited the AOX in the plant samples with 2 mMSalicylhydroxamic acid (SHAM). To determine the valueassociated with the AOX, we inhibited the COX with 1mM NaCN. The inhibitors were added to a closedaerated chamber containing the plant samples and waterfor at least 90 min before measurements were made.Additional control experiments without inhibitors wereconducted with Lemna and natural soil (the resulting qvalue was expected to lie between that of the AOX andthe COX).[35] For the incubation, about 1 g of plant sample wasinserted into 6 cm 3 blood collection tubes (Vacutainers 1 )that were closed with rubber septa. The tubes wereimmersed in water during incubation in order to preventair leaks. After an incubation period of 1 to 10 hours, the airin the tubes was sampled, via a needle, directly to thevacuum preparation line. In other experiments, plant sampleswere incubated in water and the changes in thedissolved oxygen were monitored. At the end of theseexperiments, 120 cm 3 water were sampled in 250 cm 3pre-evacuated flasks closed with a Lowers Happert 1 valve.Sampling and extraction of the dissolved gases were doneaccording to Luz et al. [2002].3.2. Terrarium Experiments[36] We used an airtight terrarium as described by Luz etal. [1999]. The terrarium contained Philodendron plants,soil and natural water from Lake Kinneret, Israel (d 18 O=0.5% versus SMOW). It was illuminated by a fluorescentlamp (100 mE m 2 s 1 ) for 24 h d 1 ,10hd 1 ,4hd 1and 2 h d 1 in different stages of the experiment. Differentlight conditions were used to manipulate the CO 2concentration in the terrarium, which, in turn, affects therelative rate of photorespiration [Badger, 1985]. Theterrarium was also covered and darkened for dark incubations.The CO 2 concentration in the terrarium air wasdetermined by sampling 6 cm 3 air in pre-evacuated bloodcollection tubes (Vacutainers 1 ), and measuring the airsample with an infrared gas analyzer (LI-COR-6252 1 )by a method similar to Davidson and Trumbore [1995].The relative accuracy for CO 2 concentration measurementwas ±5%.3.3. O 2 -N 2 Diffusion Experiments[37] In the diffusion of O 2 in N 2 experiments, a 4-cm 3flask was filled with pure oxygen. The neck of the flask wasfilled with a diffusive medium (plastic sponge) that preventedadvective mixing, and the oxygen that diffused outof the tube was immediately removed by a flow of N 2 . After1–2 hours, the O 2 concentration in the flasks was considerablylowered and the flasks were closed and transferredfor isotopic analysis.3.4. Sample Preparation and Mass Spectrometry[38] Sampling, sample preparation, and mass spectrometrywere according to Angert and Luz [2001], Luz et al.[2002], and Luz and Barkan [2000]. The preparation of thesample included cryogenic removal of water vapor and CO 2,and chromatographic separation of N 2 , which is needed foraccurate 17 D measurements. The samples were frozen at4°K into stainless steel tubes and transferred for analysisin a Finnigan Delta-Plus mass-spectrometer. Correctionswere applied in order to account for the sensitivity ofionization efficiencies of the three isotopic species tovariations in the O 2 /Ar ratio. The analytical error (absolutedifference from the average) for d 18 O, dO 2 /Ar, and D 17 Owas 0.02%, 0.5% and 5 per meg, respectively. All values

30 - 6 ANGERT ET AL.: BIOLOGICAL EFFECTS ON THREE O 2 ISOTOPESThere was no significant difference between the l R valuesin the incubations of Lemna gibba and Orobanche aegyptica,nor between different experiments with the sameplants. The l R value determined for incubation withoutinhibitors of Lemna was 0.517 ± 0.001, and a value of 0.517± 0.002 was found for the natural soil. In the two diffusionexperiments we found l R value of 0.5228 ± 0.0002, with nosignificant difference between the experiments.[40] The 17 D, dO 2 /Ar, Ln 18 O and [CO 2 ] of the terrariumair in the course of the experiment are presented in Table 2and shown in Figure 4. In days 1–58, in which there was a24-hour illumination per day, the CO 2 concentrations in theterrarium were about 150 ppm. In days 127–178, when theillumination was of 10 h d 1 ,CO 2 concentration decreasedfrom 4500 ppm in the beginning of the light period to 150ppm, after about 3 hours. In days 184–235 and 330–365, inwhich there were 2–4 hours illumination, the CO 2 concentrationwere 10,000–40,000 ppm, and in days 249–252, inwhich there were 4 hours of illumination, CO 2 decreased toabout 150 ppm after 2–3 hours of illumination. The isotopicdiscrimination ( 18 e) that was calculated from the darkperiods (days 183–184 and 283–285) is 16.7%.Figure 3. Relative changes in the triple isotopic compositionof O 2 (plotted as Ln 17 O versus Ln 18 O in per mils) asresult of (a) uptake through the COX, (b) uptake through theAOX, and (c) removal by diffusion into N 2 . The slope of thefitted regression line is l, and the reported precisionaccuracy is standard error. The values of q are calculatedaccording to equation (17).are reported with respect to the atmospheric O 2 (HLAstandard).4. Results[39] The results of the dark incubation and the diffusionof O 2 in N 2 experiments, are illustrated in Figure 3, andsummarized in Table 1. The isotopic discrimination ( 18 e)was calculated by equation (15) from the results of the darkincubation experiments. As expected, higher fractionationfor each plant species was measured when the plants wereinhibited with NaCN, and lower fractionation when theplants were inhibited with SHAM. Considerably loweruptake rates were measured when the plant was inhibitedwith NaCN (data not shown). The value of l R in theseexperiments was determined from the Ln 17 O versus Ln 18 Oaccording to equation (4). The l R values that were found are0.518 ± 0.001 for the COX, 0.5179 ± 0.0003 for the AOX.5. Discussion5.1. Dark Incubation Experiments and Diffusion inLiquid Phase[41] The value of q was calculated from the measured l Rvalues by equation (17). The fractionations ( 17 e and 18 e)arestrongly controlled by the effects of slow diffusion to theconsumption site [Angert and Luz, 2001; Guy et al., 1989].However, the value of l R in our experiments can be shownto be independent of such effects. The overall fractionationof a system in which the oxygen supply is limited bydiffusion is given byx e ¼ x e diff þ ð x e x con e diff ÞC i =C a ; ð18Þtaken from Farquhar et al. [1982], where x e is the overallfractionation, x e con is the fractionation in consumption, x e diffTable 1. Summary of Dark Incubation Experiments and DiffusionExperiments a Inhibitor n l SE e SEOrobanche aegyptica SHAM 3 0.517 0.003 14.5 0.9Orobanche aegyptica SHAM 4 0.520 0.006 14.2 0.7Lemna gibba SHAM 7 0.518 0.003 20.7 0.1Lemna gibba SHAM 7 0.518 0.002 19.6 0.2Lemna (in water) SHAM 8 0.518 0.001 14.1 0.4COX 29 0.518 0.001Orobanche aegyptica NaCN 4 0.516 0.002 18.5 0.9Orobanche aegyptica NaCN 5 0.517 0.002 21.5 1.4Lemna gibba NaCN 6 0.5181 0.0003 27.9 0.1Lemna gibba NaCN 6 0.518 0.001 27.1 0.2AOX 21 0.5179 0.0003Lemna gibba - 4 0.517 0.001 20.1 0.3Soil - 4 0.517 0.002 16.6 0.3Diffusion in N 2 - 6 0.5226 0.0004 14.9 0.1Diffusion in N 2 - 6 0.5231 0.0003 - -Diffusion in N 2 12 0.5228 0.0002a Abbreviations: n, number of data points; SE, standard error.

ANGERT ET AL.: BIOLOGICAL EFFECTS ON THREE O 2 ISOTOPES 30 - 7Table 2. Summary of the Terrarium ExperimentDayDO 2 /ArVersus Atm, aD 18 Oversus Atm, aLn 18 Oversus Water b17 Dversus Atm, aLight, CO 2 ,hd 1 ppmCO 2 ,

30 - 8 ANGERT ET AL.: BIOLOGICAL EFFECTS ON THREE O 2 ISOTOPESwill be an intermediate value between the q of COX and thatof diffusion. Such a situation occurs in soils in general (andalso the soil in our terrarium experiments), in whichfractionation of entering O 2 by diffusion in air has beenshown to have strong effects [Angert et al., 2001]. Since soilrespiration is an important component of the oxygen cycle,such an effect should be considered in any attempt toestimate the terrarium or global q.[46] The effective fractionation of a system in which theoxygen supply is limited by diffusion is given by equation(18). We can estimate the effective q for different C i /C aratios by calculating the 17 e and 18 e using this equation.Oxygen concentration in soils vary widely with soil depth,soil type, soil moisture, and rate of respiration in the soil.However, most of the soil respiration takes place at the topof the soil profile. Hence, it will be justified to assume thatthe C i /C a ratio for weighted average soil respiration is notsmaller than 0.9. This value corresponds to about 20,000ppm of CO 2 , which is in the upper range of the concentrationsin soils. Even for this high value, the effective q is0.5164, only 0.0004 larger than that of COX and within themargin of error of its assessment. This effect might becompensated if some of the consumption in soil respirationis through the AOX, in which q is lower than in the COX.Thus, the effect of diffusion on the effective q in soils isvery small. Confirmation for that comes from the l value(0.517 ± 0.001) that was measured in natural soil, in whichdiffusion limitation was indicated by relatively weak fractionation( 18 e = 16%). In summary, effects of diffusion inair through soil profiles can be neglected in the discussionof the global triple isotope balance as well as the isotopicbalance in the terrarium. Additional diffusion limitations insoils occur in liquid phase, but, as shown in the previoussection, the effect on the triple isotopic composition isnegligible.Figure 4. Summary of the terrarium experiment results.Shown in the markers are the variations in (a) 17 D versus theatmosphere in per megs (with C = 0.516), (b) Ln 18 O in permils versus the terrarium water, and (c) dO 2 /Ar in per milsversus the atmosphere. The solid line shows the hours perday in which photorespiration was engaged due to low CO 2concentration (5000 ppm) throughout the day and the average17 D was 215 ± 14 per meg. In the second interval (‘‘Low’’),

ANGERT ET AL.: BIOLOGICAL EFFECTS ON THREE O 2 ISOTOPES 30 - 9the terrarium was illuminated for 24 h d 1 ,CO 2 concentrationwas low (150 ppm), and the average 17 D was 93 ±6 per meg. In the third interval (‘‘Variable’’), the terrariumwas illuminated for 10 h d 1 and the CO 2 concentrationchanged with the illumination. The plant was exposed toCO 2 levels below 500 ppm, while illuminated, for about 7hours per day. The average 17 D in this interval was 145 ± 6per meg. Similar dependence of the D 17 O values inillumination can be seen in the terrarium experiments ofLuz et al. [1999]. Although there is no CO 2 concentrationdata for these experiments, the relationship between illuminationand CO 2 were probably similar to the relationships inthe current experiment.[49] The different D 17 O values in steady state indicatechanges in the average q of the terrarium (Figure 2).These changes indicate variations in the relative rates ofdifferent processes in the terrarium, each with its characteristicq. The difference between the q of COX and AOXwe report here is only 0.002. Thus, even if the entireuptake in the terrarium was shifted from COX to AOX,which is not very likely, it will only cause a lowering of17 D of about 50 per meg (according to equation (12)).Therefore, the change of 122 per meg in 17 D betweenintervals ‘‘High’’ and ‘‘Low’’ cannot be explained bychanges in the ratio between COX and AOX. It can beexplained by an increase in the importance of a process(or processes) in which q is considerably smaller than indark respiration. Because lowering of q occurs when CO 2concentration is low, photorespiration is a likely mechanism.Another process that might be activated in the lightis the Mehler reaction. Whether or not this is an importantmechanism is debatable, but preliminary experiments inour laboratory indicate that the q value in Mehler reactionis relatively high (Y. Helman, unpublished results, 2002),and hence this reaction could not be responsible for thelowering of the terrarium q in the light. In addition, in arecent review, Badger et al. [2000] argue that O 2 uptakeby Mehler reaction is of small importance in higherplants.[50] When CO 2 concentration in the terrarium was highthroughout the day (interval ‘‘High’’), photorespiration wasinhibited and the average q was controlled by COX andAOX alone. Since the q values of COX and AOX are close(0.516 and 0.514, respectively), and usually most of therespiration takes place through the COX, we assume that theaverage q in interval ‘‘High’’ (‘‘q 1 ’’) was 0.516. Since q 1 isequal to C, the 17 D W of the terrarium water is equal to17 D BSS in interval ‘‘High’’ (see equation (12)), 215 per megon average. In interval ‘‘Variable,’’ the CO 2 concentrationwas low enough to allow photorespiration in part of the dayand D 17 O dropped from 215 to 145 per meg, and it wentdown even further to 93 per meg, in interval ‘‘Low,’’ whenlow CO 2 concentration and illumination were present for 24hours per day (Figure 5).[51] By using the value of 215 for 17 D W and 18 e of 24%(see section 5.5) in equation (13), we calculated an averageq for interval ‘‘Low’’ (‘‘q 2 ’’) of 0.511. This value of q 2represents a weighted average for both dark respiration andphotorespiration. This clearly indicates that the q value ofphotorespiration must be significantly less than 0.516. AFigure 5. Analysis of the terrarium experiments (not toscale). In interval ‘‘High,’’ photorespiration was inhibited asa result of high CO 2 concentrations and the relationshipbetween the fractionations was controlled by dark respiration(q 1 = 0.516). In Interval ‘‘Low,’’ both dark respirationand photorespiration (PR) were engaged (q 2 = 0.511). Toaccount for such an effect, the q of photorespiration must belower than 0.511, and it can be calculated from the relativerates of dark respiration and photorespiration as 0.506.quantitative estimate of the q value for photorespiration willbe derived in section 5.6.5.4. Deviations From Steady State[52] While the 17 D of the terrarium air was in steady statein intervals ‘‘Low,’’ ‘‘Variable’’ and ‘‘High,’’ the dO 2 /Arvalues were constant only through interval ‘‘Low.’’ Inintervals ‘‘Variable’’ and ‘‘High,’’ the O 2 concentrationincreased with time, indicating a net production in theterrarium. To study the effect of this net production on the17 D, we used a simple numerical model. This model dealswith the fluxes of production and uptake of the differentisotopic species without the steady state assumption. Thetemporal concentration of isotopic species is given byequation (19) where Dt stands for time interval.x O 16 Ot þDt ¼ x O 16 Ot þ Dt 16 P x R W16 U x R t x a : ð19Þ[53] As expected, running the model with 16 U =16 P and q =C gave identical result to the analytical model (equation(12)); hence, when the value of 17 D reached a plateau(equivalent to 17 D BSS in the analytical model), it was equalto that of 17 D W . However, running the model with production( 16 P) that is higher then uptake ( 16 U) resulted in 17 Dvalues in a plateau slightly higher then 17 D W . For example,running the model with 16 U/ 16 P ratio of 0.5 and 17 D W of211 per meg resulted in 17 D in a plateau of 221 per meg. Toestimate uptake in interval ‘‘High,’’ we assumed that it wasidentical to the uptake rate in the dark periods (the oxygendepletion rate). This assumption is probably true since thehigh CO 2 levels in interval ‘‘High’’ inhibited photorespiration.From this uptake rate and the observed net production(rate of increase in O 2 concentration), we found the ratio ofuptake to production was 0.9 for interval ‘‘High.’’ Using

30 - 10 ANGERT ET AL.: BIOLOGICAL EFFECTS ON THREE O 2 ISOTOPESthis value in the numerical model, we found that thisdeviation from steady state will cause an increase of 4per meg in the value of 17 D, when it reached a plateau.Thus, 17 D W is 211 per meg instead of the 215 per megcalculated based on the steady state assumption. Thiscorrection for 17 D W is small with respect to other uncertaintiesand does not significantly change the value calculatedfor q 2 (0.511).[54] Another nonsteady state effect results from the diurnalillumination cycle in the terrarium. Photosynthesis tookplace in the few hours of illumination, while the uptakethrough dark respiration continued all day, and as a resultthe O 2 concentration fluctuated. By modeling this conditionin the numerical model, we found that it will cause the 17 Dvalue to fluctuate around the value 17 D BSS . The amplitudeof the fluctuations in 17 D depends on the amplitude of thefluctuations in O 2 concentration. For interval ‘‘High,’’ wecan calculate from the dark periods that dark respirationconsumed about 9% of the O 2 reservoir of the terrarium perday. Using this value and the value of 0.9 for uptake toproduction ratio, the calculated magnitude of the 17 Dfluctuations resulting from the light-dark cycle is ±2 permeg. Again, this value is considerably smaller than theanalytical uncertainty, and therefore this effect can be alsoneglected.5.5. Dependence of 18 E on Illumination and [CO 2 ]:Implication for the Dole Effect[55] The weighted-average 18 e of all the processes in theterrarium can be estimated from the terrarium equivalent ofthe global Dole Effect (equation (14)), which is the value ofLn 18 O of ‘‘BSS’’ versus ‘‘W.’’ The Ln 18 O values of theterrarium air versus the value of the substrate water arepresented in Figure 4b. In the three intervals in which the17 D of the terrarium was constant, the Ln 18 O was almostconstant.[56] In interval ‘‘High,’’ the fractionation (e) in theterrarium according to equation (14) is 18.4 ± 1.8%, ininterval ‘‘Low’’ 23.9 ± 0.5%, and in interval ‘‘Variable’’21.7 ± 0.2%. The fractionation in interval ‘‘High’’ is inagreement with the known fractionation for the COX,18% [Guy et al., 1989], and with the fractionation thatwas calculated for the dark periods 16.7%. This agreementindicates that, as was assumed in section 5.3., theuptake in the terrarium was dominated by COX in interval‘‘High’’ in which the CO 2 concentration was high.[57] The fractionation by COX is 18%, in photorespirationit is 21.7%, and that of AOX is about30% [Ribas-Carbo et al., 2000]. Thus, photorespirationand COX alone cannot explain the high Ln 18 O valuesmeasured in interval ‘‘Low’’ and interval ‘‘Variable.’’These high values seem to indicate that a considerableportion of the uptake was through the AOX. Since ininterval ‘‘Variable’’ there was net production that introducedoxygen with light isotopic composition, the fractionationmust have been even stronger than that calculatedabove for the same interval according to the steady stateassumption ( 21.7%). The fractionation in interval‘‘Low’’ was extremely strong. The relative rate of uptakethrough the AOX in this interval was estimated from theobserved Ln 18 O values in section 5.6. as 41–31% ofgross production.[58] Some enrichment of the terrarium leaf water byevapotranspiration might have contributed to the highLn 18 O. However, since the relative humidity in the terrariumwas 100% this effect was probably very small. In fact,no enrichment was found when we compared the d 18 Oofthe terrarium free water and the terrarium leaf water (datanot shown). However, this result might originate frommeasuring total leaf water, which includes depleted veinwater. Even if we assume that the enrichment at the site ofphotosynthesis was as high as 1%, our main conclusionswill remain the same. The relative rate of the AOX ininterval ‘‘Low’’ will be 19–29%, still a very high figure,and the correction to q P (see section 5.6) will be muchsmaller than the other uncertainties.[59] The strong measured fractionation indicates that theAOX was activated in the same conditions that favorhigh rate of photorespiration-illumination and low CO 2 .This finding is in agreement with the indication for highAOX rates in the light inferred from in situ measurementsin a lake [Luz et al., 2002]. The CO 2 concentrationin the terrarium were very low (150 ppm), much lowerthan in most natural environments. However, since therelative humidity in terrarium was 100% stomatal conductancemust have been high. Consequently, the internalCO 2 concentration in the leaves was similar to that ofmidday in many natural environments. Since strongfractionation occurred not only with the 24 h d 1 illuminationbut also with the 10 h d 1 illumination, which iscloser to the natural cycle, we conclude that the engagementof the AOX in the light is likely also in manynatural systems.[60] In previous models of the Dole Effect, the global rateof the AOX was assumed to be very low and was neglected[Bender et al., 1994; Malaize et al., 1999]. This low rate isbased on measurements of the AOX activity in the dark.However, if the AOX activation is enhanced in illuminatedleaves in natural systems, then its global rate should beconsiderably higher. This higher rate may help to close thegap between the calculated value of the Dole Effect (20.8%[Bender et al., 1994]) and the measured one (23.5%[Kroopnick and Craig, 1972]), and compensate for theweak fractionation recently reported for soil respiration[Angert et al., 2001]. The connection between photorespirationand AOX might also explain past changes in the DoleEffect. Increased rate of photorespiration will be coupledwith an increased rate of AOX. Drier and hotter climate isexpected to cause an increased rate of photorespiration, aswell as more evapotranspiration that will result in 18 Oenriched leaf water. Thus, such climate will cause anincreased Dole Effect by both heavier composition of leafwater and increased rate of photorespiration and the AOX,two processes that have high fractionation relative to that ofCOX.5.6. Estimating the Q of Photorespiration[61] The average q in the terrarium in interval ‘‘Low’’ was0.511. This value and the values found for the AOX andCOX, can be used to estimate the q associated with photo-

ANGERT ET AL.: BIOLOGICAL EFFECTS ON THREE O 2 ISOTOPES 30 - 11respiration by assuming production uptake steady state, anda simple weighted average equation,GP ¼ U P þ U COX þ U AOXð20Þ0:511 ¼ ðq P U P þ 0:516U COX þ 0:514U AOX Þ=GP; ð21Þwhere GP is gross production in the terrarium, U stands foruptake, and the subscript P stands for photorespiration. Inorder to solve equation (20) for q P the rates of the differentuptake pathways in the terrarium relative to gross productionmust be estimated.[62] The gross oxygen production in the terrarium, atinterval ‘‘Low,’’ can be estimated from previous experimentsin the same terrarium [Luz et al., 1999]. In one ofthese experiments the d 18 O of the water was 6.6% (onSMOW scale). As a result, the d 18 O value in steady statewas very different from the initial atmospheric value. At thebeginning of the experiment (days 1–37, illumination of24 h d 1 , as at interval ‘‘Low’’), changes in the d 18 O anddO 2 /Ar of the terrarium were observed, when it approachedsteady state. The changes in the d 18 O and dO 2 /Ar weremodeled by the numerical model described in section 5.4.The best fit to the observed data was reached with uptake togross production ratio of 0.99 and gross production rate(GP) that is 19.2% of the terrarium O 2 reservoir per day.[63] The uptake rate measured in the dark periods was45% of the GP estimated above, and was probably mostlythrough the COX. When the terrarium was illuminated,considerable portion of the total dark respiration must havebeen through the AOX (section 5.5.); hence, the rate ofCOX in interval ‘‘Low’’ was lower than 45% of GP. If weassume a COX of 10–30% of GP, the rate of photorespiration(U P ) can be calculated from the known fractionations ofthe different pathways and the terrarium-fractionation ( 18 e)we evaluated for interval ‘‘Low’’ ( 24%) by equation (20)and a weighted average equation similar to equation (21),24 ¼ ðe P U P þ e COX U COX þ e AOX U AOX Þ=GP:ð22Þ[64] Using fractionation values of 30% for AOX[Ribas-Carbo et al., 2000], 18% for COX [Guy et al.,1989], and 21.8% for photorespiration U P was calculatedas 29–59% of GP. The corresponding rate of AOX is 41–31% of GP. According to these rates and the q values of theCOX and the AOX, the q associated with photorespiration iscalculated as 0.506 ± 0.003. The error margin is based onlyon the uncertainty in the rate of the COX, and neglects theuncertainty in the fractionations of the different pathways.The uncertainty in the fractionation in the COX and photorespirationis small [Guy et al., 1989, 1993]; however theuncertainty in the fractionation of the AOX is larger, about±2% for green tissue [Ribas-Carbo et al., 2000; Robinsonet al., 1992]. Including this uncertainty the q associated withphotorespiration is 0.506 ± 0.005.[65] In photorespiration, O 2 is consumed by two enzymes,Rubisco and glycolate-oxidase. Hence, the q value of photorespirationrepresents the weighted average of the two consumptionprocesses. Further study is necessary in order toderive the q values associated with each of the two enzymes.Figure 6. Analysis of triple isotopic composition (not toscale) of the Last Glacial Maximum (LGM). The point thatrepresents the LGM atmosphere (after correction for lowerrate of stratospheric photochemistry) lies 12 per meg lowerthan the present atmosphere. A change in the global averageq (q b ) will affect the triple isotopic composition of theatmosphere in a way that is similar to the way a change in qaffected the terrarium air (Figure 5). Hence, a change in q bwill cause a change in the atmospheric 17 D with amagnitude of the Ln 18 O difference between atmosphericO 2 and photosynthetic oxygen times the difference in q b .Thus, a change of 0.001 in q b will result in a change ofabout 19 per meg in the atmospheric 17 D. The higher globalrates of photorespiration (PR) in the LGM (27% of globalproduction instead of 24% today) caused lowering of q b thatcan explain 5 per meg of the 12 per meg difference. The restof this difference can be attributed to lower global grossproduction.[66] The q value of 0.506 ± 0.005 we report here forphotorespiration is considerably lower than that of both darkrespiration pathways. In the next section, we will use thisvalue to estimate the effect of photorespiration on the tripleisotopic composition of the atmosphere.5.7. Implication for the Triple Isotopic Composition ofAtmospheric O 2[67] In order to estimate how a change in the globalaverage q of biological processes will affect the tripleisotopecomposition of the atmosphere, we can analyze itin a graphical approach presented in Figure 6, or with arigorous analytical model. In this model, the troposphere(which is well mixed relatively to the lifetime of oxygen init) is represented by one box, and O 2 is exchanged betweenthe biosphere, the troposphere, and the stratosphere. Insteady state we can write the following equation:y Sy T þ y Py U ¼ 0;ð23Þwhere y S is the stratosphere to troposphere flux and y Trepresent the flux from the troposphere to the stratosphere

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

ANGERT ET AL.: BIOLOGICAL EFFECTS ON THREE O 2 ISOTOPES 30 - 13change can be explained as the result of higher photorespirationrate in the LGM. In conclusion, small changesin the global rate of photorespiration can have an importantimpact on the atmospheric 17 D. Consequently, any interpretationof triple-isotope ice core data must take thisimpact into account.5.7.2. Global 17 D Budget in the Present Atmosphere[74] To gain some insight on the global 17 D budget, weestimate the parameters of equation (26) that represents theeffects of the biosphere and the stratosphere on the tripleisotopic composition. The current value of q b was calculated,as in the previous section, from the derived q valuesand the estimated global relative rates of photorespiration(24%), COX (68%) and AOX (8%), as 0.513 ± 0.002 (0.24 0.506 + 0.68 0.516 + 0.08 0.514 = 0.513, the errormargin is derived from the uncertainty in the q values).According to Luz et al. [1999] the stratospheric massindependentprocesses result in lowering of the currentatmospheric D 17 O in 117 ± 13 per meg. In the terms ofequation (26) (with 16 P/ 16 S ratio of 0.0097 [Luz et al., 1999]and q s = 1) this 117 ± 13 per meg lowering corresponds tox a s of 0.99999765 ± 0.0000001 assuming 17 D ffi D 17 O and18 e b that was calculated in section 5.7.1 as 19.4%. Usingthese parameters, we calculated by equation (26) that 17 D T-AW (the troposphere versus AW) is 175 ± 50 per meg (theerror margin is based on the error in q b and x a s ). Correspondingly,17 D AW (the average substrate water for photosynthesisversus the troposphere) is 175 ± 50 per meg. Thiscalculation can also be preformed graphically on Figure 7.The calculated 17 D AW value is smaller than either 215 per megobtained for Lake Kinneret water (section 5.3) or 249 per megfor seawater [Luz and Barkan, 2000]. The difference betweenthe calculated 17 D AW and value obtained for seawater can besolved if higher global rates of photorespiration will beassumed in the calculation. However, this will not explainthe difference between the values obtained for Kinneret waterand seawater. Both this difference, and the difference betweenseawater and the calculated 17 D AW , can be explained byvarious isotope effects in the hydrological cycle.[75] Reported l values for meteoric precipitation rangefrom 0.528 [Meijer and Li, 1998] to 0.525 [Miller, 2002].Since q is only slightly lower than l (0.002), it is clear thatthe corresponding q values are higher than the ones wedetermined for biological uptake (0.506–0.516). The sourceof Lake Kinneret water is depleted meteoric precipitationthat later becomes enriched by evaporation. Hence, arelatively low q associated with evaporation can explainwhy the 17 D W of lake Kinneret is lower than that of seawater[Luz and Barkan, 2000]. Similar explanation can show how17 D AW becomes 74 per meg smaller than 17 D W of seawater.If the q of evapotranspiration is also low, then the depletionin the heavy isotopes of meteoric precipitation and theenrichment in evapotranspiration will yield a low 17 D ofaverage leaf water. Such low 17 D of leaf water will contributeto a low 17 D AW (175 per meg, Figure 7).[76] The discussion above demonstrates that rigorousclosure of the triple isotopic balance requires direct measurementsof the q values associated with the hydrologicalcycle. Such estimation will enable independent estimate ofthe isotopic composition of average water ( 17 D AW ). TheFigure 7. Schematic illustration (not to scale) of suggestedtriple isotope budget in the present global atmosphere. Theglobal composition of oxygen produced by photosynthesis(AW) is a weighted average of about 60% O 2 producedfrom leaf water and 40% from seawater. Average leaf wateris formed from seawater that is depleted in 17 O and 18 Oduring meteoric precipitation (q 0.525) and enriched byevapotranspiration (q < 0.525). AW oxygen is enriched dueto biological uptake (q b = 0.513) to form BSS oxygen. BSSsignifies the hypothetical composition of the atmospherethat would have been produced in the absence of stratosphericphotochemistry. Further mass independent depletion(q 1) by stratospheric photochemistry causes 117 per meglowering of atmospheric 17 D.average global biological q (q b ) could be calculated from17 D AW and the measured triple isotopic composition ofatmospheric O 2 . Our study shows that the relative globalrates of dark respiration and photorespiration control q b .Hence, the rates of these globally important processes couldbe evaluated from triple isotope studies.6. Conclusions[77] The q values for the cytochrome pathway, the alternativepathway, photorespiration, and diffusion in air are0.516 ± 0.001, 0.514 ± 0.001, 0.506 ± 0.005, and 0.521 ±0.001, respectively. The combined effect of diffusion andrespiration on the atmosphere was shown to be close to thatof dark respiration. Since the value associated with photorespirationwas found to be considerably lower than that ofdark respiration, the triple isotopic composition of atmosphericO 2 is strongly controlled by the relative global ratesof these two processes. This control makes the tripleisotopic composition a tracer of the global rate of photorespirationin the past, as well as in the present.[78] The engagement of the alternative oxidase pathway(AOX) in the light is larger in low CO 2 than in high CO 2concentrations. Thus, environmental conditions that lowerthe internal CO 2 concentration in leaves are expected tocause an increase in photorespiration rate as well as in theAOX rate. Since the fractionation in both processes is strong,such environmental conditions will cause an increased DoleEffect.

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