Isotopic Composition of Carbon Dioxide in the Middle Atmosphere

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1 Isotopic Composition of Carbon Dioxide in the Middle Atmosphere Mao-Chang Liang, Geoffrey A. Blake, Brenton R. Lewis, and Yuk L. Yung Division of Geological and Planetary Sciences, California Institute of Technology 1200 E. California Blvd., Pasadena, CA Research School of Physical Sciences and Engineering, The Australian National University Canberra, ACT 0200, Australia To whom all correspondence should be addressed. Photolysis of O O and O O by solar Lyman- radiation yields O( D) times more enriched in O and O than that from O photolysis. We incorporate these processes into models that include isotope-dependent schemes for the O( D) + CO reactions as well as stratospheric/mesospheric transport to explain the unusual enhancement of O in three-isotope analysis results of oxygen in CO for in situ measurements from the upper stratosphere and lower mesosphere. New laboratory and atmospheric measurements are proposed to test our our model and validate the use of CO isotopic fractionation as a tracer of chemical and dynamical processes in the middle atmosphere. Long-lived molecules are useful tracers of atmospheric processes such as stratospheretroposphere exchange and mixing in the middle atmosphere. Carbon dioxide is potentially the most useful species among trace molecules (e.g., CH, N O, SF, and the CFCs) thanks to its high abundance ( 350 ppmv in the stratosphere, dropping to 100 ppmv at the homopause) and stability in the upper atmosphere. Indeed, while the reactions of trace molecules with O( D) 1

2 usually lead to their destruction, the O( D) + CO reaction regenerates carbon dioxide. This reproduced CO is unique in its potential to trace the chemical and dynamical processes in the upper stratosphere and mesosphere. When transported down to the troposphere, it will produce measurable effects in biospheric cycles involving CO (1). In particular, the mass-independent fractionation (MIF) first discovered in stratospheric ozone (O ) (2) has since been discovered in long-lived trace molecules such as N O and CO (3). The first stratospheric/mesospheric measurements of O(CO ) and O(CO ) at 30 N by Thiemens et al. (4) led to a suggestion that the MIF in stratospheric CO is caused by isotopic exchange between CO and O, mediated by the reaction with O( D) (5). The observed depletion of O(O ) at 53.3 and 59.5 km (4) also provided the first evidence of mesospherestratosphere exchange, which is suggested to be due to the downward transport of lighter O atoms from altitudes greater than 80 km (6 7). Subsequent stratospheric measurements at latitudes of 43.7 N and 67.9 N (8) revealed larger fractionations in both O(CO ) and O(CO ) than those seen at 30 N. Yung et al. (5) suggested upwelling of tropospheric air from the tropics and downwelling at about 30 N could dilute the isotopic signature, further illustrating the critical role that transport plays in establishing the isotopic systematics of CO in the stratosphere and mesosphere. In three-isotope plots, the mass-dependent fractionation of oxygen has a slope of = 0.5, yet least squares fits to the CO data sets referred to above give = and , respectively. The latter has been successfully reproduced in the laboratory under similar stratospheric conditions (9). Subsequent atmospheric measurements yielded still more variability: slopes of for the balloon data collected in the Arctic vortex (10) and for the ER-2 data from the lower stratosphere (11). Within the framework of Yung et al. s mechanism (5), the range of measured slopes reflects the variety of transport histories of air parcels and sources of O( D). The magnitude of O(CO ) or O(CO ) can, in principle, O/ O 2

3 be used to determine how the air parcels are transported, but only if all sources of O( D) are accounted for. As we discuss below, while ozone photolysis is the dominant source of O( D) in the stratosphere, other sources must be considered at higher altitudes. These have not been considered in previous models, and, as a result, the transport history of the air parcels carried by O(CO ) or O(CO ) alone tends to be ambiguous if the full suite of O( D) sources are not taken into account. This ambiguity can be largely resolved when both O(CO ) and O(CO ) are analyzed. The isotopic anomalies of CO are caused by the following reactions O( D) + C O O C O O + O O( D) + C O O C O O + O O( D) + C O O C O O + O O( D) + C O O C O O + O O( D) + C O O C O O + O where O is either O( D) or O( P) and - are the fractionation in the rate coefficients. The resulting isotopic composition of CO in equilibrium is then determined simply by the isotopic composition of O( D), i.e., [C O O]/[C O O] = ( )/( )[ O( D)]/[ O( D)]. However, the time constant for this isotopic exchange is substantially greater than typical atmospheric transport times. As a result, the age of air can also have a significant impact on the magnitude of O(CO ) and O(CO ). For example, at an altitude of 45 km, where O( D) peaks, the chemical exchange time is 10 s, while the vertical mixing time is 10 s. The age of air entering from the troposphere is 10 s (12). During the time air ascends from the tropopause to this altitude, vertical mixing acts to dilute the isotopic fractionation of CO. Thus, as the system approaches a steady state, the isotopic composition of CO is determined by the combination of the isotopic composition of O( D) and transport. 3

4 Limits to the slope can be estimated by assuming isotopic equilibrium between CO and O( D), since the transport will only dilute the magnitude of O and O. Based on Yung et al. s mechanism, O and O can be approximated by O(CO ) O( D) O(CO ) (1) O(CO ) O( D) O(CO ) (2) where O(CO ) and O(CO ) are the enrichments of CO relative to the tropospheric values, which are O(CO ) 21 and O(CO ) 41 per mil relative to atmospheric O. If the reaction rate with O( D) is scaled by the reduced mass of each colliding pair, - values are -21.8, -3.0, -41.6, -5.8 per mil, respectively. However, the abundance and values of O( D) must also include a quenching reaction with O (and N ), and such reaction provides additional enrichments of 19.8(18.9) and 37.7(36.0) to O( D) and O( D), respectively. Since these effects are of opposite sign and similar in magnitude, the reduced mass effect is largely minimized in Eqs.(1) and (2), and the slope can thus be well approximated by ( O( D) - O(CO ) )/( O( D) - O(CO ) ). For O( D) = O( D) = 100 per mil, which are of about the same magnitude as isotopic fractionation of ozone in the stratosphere, the approximate is 1.3. To account for the observed slope of 1.7, there must be a source of O( D) with O( D) O( D). More quantitative calculations of the three-isotope slope can be obtained via recent kinetic calculations modeling the isotopic fractionation of ozone versus altitude (13). Contributions to the enrichments of isotopically heavy ozone follow from two processes: chemical formation and UV photolysis. The formation reactions determine the approximate magnitude of the enrichments and their mass-independent characteristics, while the photolysis explains the observed altitudinal variation of the enrichments. Using the best fit model to the observed enrichments in a three-isotope plot of O (13), where the temperature is 200 K at 20 km and 250 4

5 K at 45 km, the computed equilibrium values of at altitudes between 35 and 60 km (where most of O( D) resides) range from The slope at the O( D) peak (at 46 km) is Over the same altitude range as that of the CO measurements in the stratosphere, the slope is 1.60, which is in good agreement with the measured value of 1.7 (8). At altitudes greater than 70 km, the photodissociation of O becomes the dominant source of O( D). Exchange of O( D) with CO at these altitudes could therefore modify the slope if the O( D) from O photolysis is isotopically distinct from that generated in the stratosphere. Using a semi-analytical calculation of the photolysis-induced fractionation (14) in the Schumann-Runge bands of O, the calculated enrichment of heavy O( D) is less than 100 per mil. However, a recent study of O dissociation near Lyman- ( Å) has shown that the cross section and O( D) yield are strong functions of wavelength, and predicts extremely large isotopic dependence (15). Although the cross section near Lyman- is 2-3 orders of magnitude less than those in the Schumann-Runge bands, the solar flux is correspondingly enhanced. For a range of temperature from K, using the coupled-channel method described in detail by Lacoursière et al. (15), we have computed the isotopic dependence of the O dissociation cross section and O( D) yield near Lyman-. Using these cross sections, together with the solar spectrum, we calculate that the O( D) and O( D) resulting from O photolysis peak at about 80 km and have values of and per mil, respectively. These fractionations are large and give 0.3. With even small amounts of mixing of mesospheric air with the = 1.6 gas that characterizes the stratosphere, it is likely that oxygen photolysis can provide an explanation for the slope of 1.2 or less observed by Thiemens et al. (4). We stress that Lyman- photolysis as the source of O( D) has not been considered in previous models. To provide a more quantitative assessment of the probable impact of the multiple sources of O( D), the results of a one-dimensional atmospheric model are summarized in Fig. 1. The dominant slope of is produced by the O( D) from ozone photolysis, and reproduces the 5

6 stratospheric data well. The change in slope of the solid line at A corresponds to altitudes km. At higher altitudes (and for fractionations greater than these fiducial values), the slope is 0.3 as expected from oxygen photolysis. Another change of slope in the calculation occurs at B for altitudes of 90 km and higher. Over this range, molecular diffusion dominates, and the slope becomes mass dependent, that is 0.5. As the inset in Fig. 1 shows, the heavy O atoms from O Lyman- photolysis can greatly modify the isotopic composition of CO at altitudes 40 km. To provide a better agreement with the stratospheric measurements (8), the eddy coefficients below 40 km in this calculation have been reduced by 30% compared to those commonly used (16). The disagreement of the model results with the measurements of Thiemens et al. (4) is most likely due to circulation cells between the tropics and 30 N, where the air is significantly younger than that at higher latitudes and similar altitudes (5, 12). This difference cannot be resolved satisfactorily without two- or three-dimensional simulations. Indeed, we expect that a global mapping of, when combined with proper models, should be able to refine our understanding of atmospheric transport and chemical processes especially in the remote regions of the mesosphere. Finally, we use a three-box model to evaluate the potential impact of transport on the slope and magnitude of the CO isotopic fractionation. Box 1 (fresh air from the troposphere) has O(CO ) = 21 and O(CO ) = 41 per mil. Box 2 (the stratosphere) has O(CO ) = O( D) = 115 and O(CO ) = O( D) = 100 per mil, values defined by chemical equilibrium with O( D). Box 3 (the mesosphere) has O(CO ) = O( D) = 3137 and O(CO ) = O( D) = per mil, that predicted from the Lyman- photolysis of molecular oxygen. The isotopic composition of CO can be constrained by mixing air parcels from these three boxes. O(CO ) O(CO ) O(CO ) O(CO ) O(CO ) O(CO ) O(CO ) O(CO ) O(CO ) (3) O(CO ) (4) 6

7 where,, and are the fractions of air from boxes 1, 2, and 3, and + + = 1. For atmospheric CO,. Fig. 2 gives an illustration of this simple model. It is clear that the magnitude of O(CO ) increases with the age of the air parcel, i.e., more CO will exchange with O( D) from boxes 2 and 3. The solid line in this figure shows values of O(CO ) obtained by varying while = 0. The symbols represent O(CO ) with different degrees of mixing with box 3 after mixing of air from boxes 1 and 2. The mixing of boxes 1 and 2 produces a slope of 1.6, as shown in the solid line of Fig. 3. When mixing in air from box 3 the slope is modified, and the dotted line represents the cases for which (, ) = (0.82, 0), (0.77, ) and (0.72, 0.001). Two extreme data points from Fig. 1 are overplotted by asterisks. About 0.02% mixing with box 3 is needed to explain the atmospheric measurements. Generally speaking, in the middle atmosphere, we predict that the upwelling air would have CO isotopic fractionations characterized by ozone in the stratosphere, while that by O photolysis would play a part in the downwelling air. In summary, we have demonstrated that the isotopic composition of CO is potentially an exceptionally useful tracer in studying the dynamical and chemical processes in the middle atmosphere. Experimentally, laboratory measurements of the dissociation cross sections of isotopically substituted O near the Lyman- line ( Å) are urgently needed. Observationally, whole-atmosphere mapping of CO is needed. When combined with two- or three-dimensional atmospheric models, these measurements will refine our understanding of the dynamical and chemical history of trace molecules in the middle atmosphere. 7

8 References and Notes 1. B. Luz, E. Barkan, M. L. Bender, M. H. Thiemens, K. A. Boering, Nature 400, 547 (1999). 2. K. Mauersberger, Geophys. Res. Lett. 8, 935 (1981). 3. M. H. Thiemens, Science 283, 341 (1999). 4. M. H. Thiemens, T. Jackson, E. C. Zipf, P. W. Erdman, and C. van Egmond, Science 270, 969 (1995). 5. Y. L. Yung, A. Y. T. Lee, F. W. Irion, W. B. DeMore, and J. Wen, J. Geophys. Res. 102, (1997). 6. R. N. Clayton, T. K. Mayeda, D. E. Brownlee, Earth Planet. Sci. Lett. 79, 235 (1986). 7. F. D. Colegrove et al., J. Geophys. Res. 70, 4931 (1965). 8. P. Lämmerzahl, T. Rockmann, C. A. M. Brenninkmeijer, D. Krankowsky, and K. Mauersberger, Geophys. Res. Lett. 29, 1582 (2002). 9. S. Chakraborty and S. K. Bhattacharya, J. Geophys. Res. 108, 4724 (2003). 10. B. Alexander, M. K. Vollmer, T. Jackson, R. F. Weiss, and M. H. Thiemens, Geophys. Res. Lett. 28, 4103 (2001). 11. K. A. Boering, T. Jackson, K. J. Hoag, A. S. Cole, M. J. Perri, M. Thiemens, and E. Atlas, Geophys. Res. Lett. 31, L03109, (2004). 12. T. M. Hall, D. W. Waugh, K. A. Boering, R. A. Plumb, J. Geophys. Res. 104, 18815, (1999). 8

9 13. M. C. Liang, F. W. Irion, J. D. Weibel, C. E. Miller, G. A. Blake, and Y. L. Yung, in preparation (manuscript available upon request). 14. M. C. Liang, G. A. Blake, and Y. L. Yung, J. Geophys. Res. 109, D10308, (2004). 15. J. Lacoursière, S. A. Meyer, G. W. Faris, T. G. Slanger, B. R. Lewis, and S. T. Gibson, J. Chem. Phys. 110, 1949, (1999). 16. M. Allen, Y. L. Yung, and J. Waters, J. Geophys. Res. 86, 3617, (1981). 17. Special thanks to G. R. Gladstone for the solar Lyman- flux, and S. T. Gibson for providing his coupled-channel code. We thank B. C. Hsieh, X. Jiang, and R. L. Shia for helping us with the model, and M. Gerstell, H. Hartman, A. Ingersoll, C. Miller, H. Pickett, and all of members in our group for their helpful comments. This work was supported by an NSF grant ATM The development of one-dimensional model was supported partially by NASA grant NAG

10 Fig. 1. Three-isotope plot of oxygen in CO. The tropospheric values have been subtracted off. Solid line: model calculation. Dashed line: = 1.7. Line AB: 0.3. Dash-dotted line: 0.5. The atmospheric measurements are from Lämmerzahl et al. (7) (circles) and from Thiemens et al. (3) (asterisks). Inset: vertical profiles of O(CO ). Dotted line represents the results excluding the fractionation of O( D) from O photolysis over the solar Lyman- emission. Fig. 2. Three-box mixing model. Solid line represents the mixing of boxes 1 (troposphere) and 2 (stratosphere) only, while the symbols denote additional mixing with box 3 to a different degree. Squares: no mixing with box 3. Triangles: 0.05% of air from box 3. Diamonds: 0.1% of air from box 3. Two arrows indicate the direction of increase of the age of air. Fig. 3. The notation and symbols are the same as in Figs. 1 and 2. The dotted line is an example illustrating the flattening of the slope by mixing of air from boxes 2 and 3. 10

11 Figure 1: 11

12 Figure 2: 12

13 Figure 3: 13

The Isotopic Composition of Carbon Dioxide: A Tracer of Dynamical and Chemical Processes in the Middle Atmosphere

The Isotopic Composition of Carbon Dioxide: A Tracer of Dynamical and Chemical Processes in the Middle Atmosphere Mao-Chang Liang, Geoffrey A. Blake, and Yuk L. Yung Division of Geological and Planetary