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

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1 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 Sciences, California Institute of Technology 1200 E. California Blvd., Pasadena, CA To whom all correspondence should be addressed. mcl@gps.caltech.edu A long-standing problem involving C is its isotopic composition in the middle atmosphere. In a three-isotope plot of oxygen in C, a non-mass dependent slope of 1.7 is found in the stratosphere, that decreases to 1.2 in the mesosphere. Here we show that the exchange with atmospheric ( D) can explain this trend. In the stratosphere, the major source of ( D) is photolysis, while photolysis dominates in the mesosphere. Because of its high abundance ( 350 ppmv), C isotopic composition is a powerful and unique tool for tracing the chemical and dynamical history of the upper atmosphere, where other common tracers, such as CH and N, are no longer available. Long lived molecules are useful tracers of atmospheric processes such as stratospheretroposphere exchange and mixing in the upper atmosphere. Carbon dioxide is potentially the most useful species among trace molecules (e.g., CH, N, SF, and the CFCs) thanks to its high abundance ( 350 ppmv in the stratosphere, dropping only to 100 ppmv at the homopause) and stability in the upper atmosphere. Indeed, while the reactions of trace molecules 1

2 with ( D) usually lead to their destruction, the C + ( D) reaction regenerates carbon dioxide. The unusual isotopic effects produce signatures that are unique in their ability to trace the chemical and dynamical processes in the upper stratosphere and mesosphere. In particular, the mass-independent fractionation (MIF) first discovered in stratospheric ozone ( ) (1) has since been extended to long lived trace molecules such as N and C (2). Spurred on by the first stratospheric/mesospheric measurements of (C ) and (C ) at 30 N by Thiemens et al. (3), it has been suggested that the MIF in stratospheric C is caused by isotopic exchange between C and, mediated by the reaction with ( D) (4). The observed depletion of ( ) at 53.3 and 59.5 km (3) also provided the first evidence of mesosphere-stratosphere exchange, which is believed to be due to the downward transport of lighter atoms from altitudes greater than 80 km (5 6). Subsequent stratospheric measurements by Lämmerzahl et al. (7) at latitudes of 43.7 N and 67.9 N revealed larger fractionation in both (C ) and (C ) than those seen at 30 N. Yung et al. (4) 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 C 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 C data sets referred to above give = and , respectively. The latter one has been successfully reproduced in laboratory under a condition similar to that in the stratosphere (8). Subsequent atmospheric measurements yield still more variability: slopes of for the Alexander et al. (9) balloon data collected in the Arctic vortex and for the Boering et al. (10) ER-2 data from the lower stratosphere. Within the framework of Yung et al. s mechanism (4), the range of the measured slopes reflects the different transport history of air parcels and sources of ( D). The magnitude of (C ) or (C ) can, in principle, be used to determine how the air parcels are trans- 2

3 ported, but only if all sources of ( D) are accounted for. As we will discuss below, while ozone photolysis is the dominant source of ( 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 (C ) or (C ) alone tends to be ambiguous if the full suite of ( D) sources are not taken into account. This ambiguity can be largely removed when both (C ) and (C ) are analyzed. The isotopic anomalies of C are caused by the following reactions ( D) + C ( D) + C ( D) + C ( D) + C ( D) + C C C C C C where is either ( D) or ( P) and - are the isotopic fractionation in the rate coefficients (3). The resulting isotopic composition of C in equilibrium is then determined simply by the isotopic composition of ( D), i.e., [C ]/[C ] = ( )/( )[ ( D)]/[ ( 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 (C ) and (C ). For example, at an altitude of 45 km, where ( 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 (11). During the time air ascends from the tropopause to this altitude, vertical mixing acts to dilute the isotopic fractionation of C. Thus, as the system approaches steady state the isotopic composition of C is determined by the combination of the isotopic composition of ( D) and transport. Limits to the slope can be estimated by assuming isotopic equilibrium between C and 3

4 ( D), since the transport will only dilute the magnitude of al. s mechanism, and can be approximated by and. Based on Yung et (C ) (C ) ( D) ( D) (C ) (1) (C ) (2) where (C ) and which are (C ) (C ) are the enrichments of C relative to the tropospheric values, 21 and (C ) 41 per mil relative to atmospheric. If the reaction rate with ( D) is scaled by the reduced mass of each colliding pair, - are -21.8, -3.0, -41.6, -5.8 per mil, respectively. However, the abundance and values of ( D) must also include quenching reaction with (and N ), which provide additional enrichments of 19.8(18.9) and 37.7(36.0) to ( D) and ( 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 ( ( D) - (C ) )/( ( D) - (C ) ). For ( D) = ( D) = 100 per mil, which are about the magnitude of isotopic fractionation in ozone, the approximate is 1.3. In order to account for the observed slope of 1.7, there must be a source of ( D) with ( D) ( D). More quantitative calculations of the two isotope slope can be obtained via recent kinetic calculations modeling the isotopic fractionation of ozone (12) versus altitude. The contributions for 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 gives an explanation of the observed altitudinal variation of the enrichments. Using the best fit model to the observed enrichments in a three-isotope plot of (12), where the temperature is 200 K at 20 km and 250 K at 45 km, the computed equilibrium values of at altitudes between 35 and 60 km (where most of ( D) resides) range from The slope at the ( D) peak (at 46 km) is 4

5 1.61. ver the same altitude range of the C measurements in the stratosphere, the slope is 1.60, which is good agreement with the measured value of 1.7 (7). In order to provide better agreement with the measurements or to predict the isotopic variation with position and season, two- or three-dimensional atmospheric models are needed since the exchange time constant of C is generally larger than the transport time. At altitudes greater than 70 km, the dissociation of becomes a more important source of ( D) than that from photolysis. Exchange of ( D) with C at these altitudes could therefore modify the slope if the ( D) from oxygen photolysis is isotopically distinct from that generated in the stratosphere. Using a semi-analytical calculation of the photolysis-induced fractionation (13) in the Schumann-Runge band of, the calculated enrichment of heavy ( D) is less than 100 per mil. However, recent measurements of dissociation spectrum near Lyman- line ( Å) showed that the cross section is a strong function of wavelength (14). Quantum calculations of the dissociation processes of predict that the isotopic dependence is extremely large. Even though the cross sections near Å are 2-3 orders of magnitude less than those in the Schumann-Runge band, the solar flux near Lyman- is correspondingly enhanced. Taking the calculated cross sections near Lyman- and the solar spectrum, we calculate that the ( D) and ( D) resulting from photolysis peak at about 80 km and have values of and per mil, respectively! These fractionations are enormous, 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. (3). We stress that this source of ( D) has not been considered in previous models. To provide a more quantitative assessment of the probable impact of the multiple sources of ( D), the results of a one-dimensional atmospheric model are summarized in Fig. 1. The dominant slope of is produced by the ( D) from ozone photolysis, and reproduces 5

6 well the stratospheric data. The change in slope of the calculation (depicted by 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. ver this range, molecular diffusion dominates, and the slope becomes mass dependent, that is 0.5. As the insert in Fig. 1 shows, the heavy atoms from Lyman- photolysis can greatly modify the isotopic composition of C at altitudes 40 km. To provide a better agreement with the stratospheric measurements (7), the eddy coefficients below 40 km in this calculation have been reduced by 30% compared to those in Allen et al. (15). The disagreement of the model results and the measurements of Thiemens et al. (3) is most likely due to circulation cells between the tropics and 30 N, where the age of air is significantly younger than that at higher latitudes and similar altitudes (4, 11). This 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 C isotopic fractionation. Box-1 (fresh air from the troposphere) has (C ) = 21 and (C ) = 41 per mil. Box-2 (the stratosphere) has (C ) = ( D) = 115 and (C ) = ( D) = 100 per mil, values defined by chemical equilibrium with ( D). Box-3 (the mesosphere) has (C ) = ( D) = 3137 and (C ) = ( D) = per mil, that predicted from the Lyman- photolysis of molecular oxygen. The isotopic composition of C can be constrained by mixing air parcels from these three boxes. (C ) (C ) (C ) (C ) (C ) (C ) (C ) (C ) (C ) (3) (C ) (4) 6

7 where,, and are the fractions of air from box-1, 2, and 3, and + + = 1. For atmospheric C,. Fig. 2 gives an illustration of this simple model. It is clear that the magnitude of (C ) increases with the age of the air parcel, i.e., more C will exchange with ( D) from box-2 and 3. The solid line in this figure shows the results of (C ) by varying while = 0. The symbols represent (C ) with different degrees of mixing with box-3 after mixing of air from boxes 1 and 2. The mixing of box-1 and 2 produces a slope of 1.6, as shown in 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). The extreme data points from Fig. 1 are overplotted by asterisks. nly 0.02% mixing with box-3 is needed to explaining the atmospheric measurements. Therefore, we conclude that more measurements of (C ) and (C ) in the mesosphere will be able to quantify mesosphere-stratosphere dynamical and chemical processes. In summary, we have demonstrated that the isotopic composition of C is potentially the most useful tracer in studying the dynamical and chemical processes in the upper atmosphere. Experimentally, laboratory measurements of the dissociation cross sections of isotopically substituted near Lyman- line ( Å) are urgent. bservationally, more measurements of C are needed. When combining with two- or three-dimensional atmospheric modelings, it will refine our understanding of dynamical and chemical history of trace molecules in the upper atmosphere. 7

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

9 13. M. C. Liang, G. A. Blake, and Y. L. Yung, J. Geophys. Res. 109, (2004). 14. J. Lacoursière, S. A. Meyer, G. W. Faris, T. G. Slanger, B. R. Lewis, and S. T. Gibson, J. Chem. Phys. 110, 1949, (1999). 15. M. Allen, Y. L. Yung, and J. Waters, J. Geophys. Res. 86, 3617, (1981). 16. Special thank to B.R. Lewis for sharing the cross sections of near the solar Lyman- line, and G. R. Gladstone for the solar Lyman- flux. We also thank B. C. Hsieh, X. Jiang, H. M. Pickett, and R. L. Shia for helping us on two-dimensional model. 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 C 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). Insert: vertical profiles of (C ). Dotted line represents the results excluding the fractionation of ( D) from photolysis over the solar Lyman- emission. Fig. 2. Three-box mixing model. Solid line represents the mixing of box-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 box-2 and 3. 10

11 Figure 1: 11

12 Figure 2: 12

13 Figure 3: 13

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