Photochemistry of Saturn's Atmosphere

Icarus 145, 166–202 (2000) doi:10.1006/icar.1999.6320, available online at http://www.idealibrary.com on

Photochemistry of Saturn’s Atmosphere II. Effects of an Influx of External Oxygen Julianne I. Moses Lunar and Planetary Institute, 3600 Bay Area Boulevard, Houston, Texas 77058-1113 E-mail: [email protected]

Emmanuel Lellouch and Bruno B´ezard DESPA, Observatoire de Paris, 92195 Meudon, France

G. Randall Gladstone Space Sciences Department, Southwest Research Institute, San Antonio, Texas 78228-0510

Helmut Feuchtgruber Max-Planck Institut f¨ur Extraterrestrische Physik, 85740 Garching, Germany

and Mark Allen1 Earth and Space Science Division, Jet Propulsion Laboratory/Caltech, 4800 Oak Grove Drive, Pasadena, California 91109 Received April 7, 1999; revised August 18, 1999

We use a one-dimensional diurnally averaged model of photochemistry and diffusion in Saturn’s stratosphere to investigate the influence of extraplanetary debris on atmospheric chemistry. In particular, we consider the effects of an influx of oxygen from micrometeoroid ablation or from ring-particle diffusion; the contribution from cometary impacts, satellite debris, or ring vapor is deemed to be less important. The photochemical model results are compared directly with Infrared Space Observatory (ISO) observations to constrain the influx of extraplanetary oxygen to Saturn. From the ISO observations, we determine that the column densities of CO2 and H2 O above 10 mbar in Saturn’s atmosphere are (6.3 ± 1) × 1014 and (1.4 ± 0.4) × 1015 cm−2 , respectively; our models indicate that a globally averaged oxygen influx of (4 ± 2) × 106 O atoms cm−2 s−1 is required to explain these observations. Models with a locally enhanced influx of H2 O operating over a small fraction of the projected area do not provide as good a fit to the ISO H2 O observations. If volatile oxygen compounds comprise one-third to one-half of the exogenic source by mass, then Saturn is currently being bombarded with (3 ± 2) × 10−16 g cm−2 s−1 of extraplanetary material. To reproduce the observed CO2 /H2 O ratio in Saturn’s stratosphere, some of 1 Also at Division for Geological and Planetary Sciences, Caltech 170-25, Pasadena, CA 91125.

the exogenic oxygen must arrive in the form of a carbon–oxygen bonded species such as CO or CO2 . An influx consistent with the composition of cometary ices fails to reproduce the high observed CO2 /H2 O ratio, suggesting that (i) the material has ices that are slightly more carbon-rich than is typical for comets, (ii) a contribution from an organic-rich component is required, or (iii) some of the hydrogen–oxygen bonded material is converted to carbon–oxygen bonded material without photochemistry (e.g., during the ablation process). We have also reanalyzed the 5-µm CO observations of Noll and Larson (Icarus 89, 168–189, 1990) and determine that the CO lines are most sensitive to the 100- to 400-mbar column density for which we derive a range of (0.7–1.5) × 1017 cm−2 ; the CO observations do not allow us to distinguish between an external or internal source of CO on Saturn. If we assume that all the extraplanetary oxygen derives from a micrometeoroid source, then the unfocused dust flux at 9.5 AU must be (i) (1 ± 0.7) × 10−16 g cm−2 s−1 if interstellar grains are the source of the external oxygen on Saturn, (ii) (4 ± 3) × 10−17 g cm−2 s−1 if IDPs on randomly inclined, highly eccentric orbits (e.g., Halleytype comet grains) are the source of the external oxygen, or (iii) (2 ± 1.4) × 10−18 g cm−2 s−1 if IDPs on low inclination, low eccentricity orbits (e.g., Kuiper-belt grains) are the source of the external oxygen. These estimates can be used in combination with future Cassini dust detection data to determine the ultimate source of the dust at Saturn’s distance from the Sun. °c 2000 Academic Press

166 0019-1035/00 $35.00 c 2000 by Academic Press Copyright ° All rights of reproduction in any form reserved.

EFFECTS OF EXTRAPLANETARY OXYGEN ON SATURN PHOTOCHEMISTRY

Key Words: Saturn, atmosphere; photochemistry; interplanetary dust; infrared observations; meteoroids.

1. INTRODUCTION

Recent Infrared Space Observatory (ISO) observations have revealed the presence of H2 O in the stratospheres of Jupiter, Saturn, Uranus, and Neptune and CO2 on Jupiter, Saturn, and Neptune (Feuchtgruber et al. 1997, 1999, de Graauw et al. 1997, Lellouch et al. 1998). Carbon dioxide is clearly a disequilibrium molecule in these hydrogen-dominated atmospheres, suggesting that the CO2 may have an external or photochemical source. Alternatively, rapid transport of thermochemically produced CO2 from hotter levels deep in Saturn’s troposphere could supply some CO2 to the stratosphere. However, Lellouch et al. (1998) estimate that internal sources could support a CO2 mixing ratio only ∼10 times smaller than what is observed. Similarly, water vapor originating from Saturn would be confined to the deep troposphere where it would condense before it could traverse the tropopause cold trap to be observed in the stratosphere. The ISO observations therefore imply an external supply of oxygen to the outer planets (Feuchtgruber et al. 1997). Possible external sources include direct atmospheric ablation of interplanetary dust particles (IDPs) or interstellar dust, an influx of materials from rings or satellites, or a deposition of material following cometary impacts. No matter what the origin of the external source, an influx of extraplanetary debris can have conspicuous effects on the chemistry and physics of outer-planetary atmospheres. Small dust particles that are introduced after cometary or micrometeoroid impacts or that diffuse inward from ring systems can cause localized heating of the upper atmosphere; the heating can in turn affect atmospheric dynamics (e.g., Rizk and Hunten 1990, West 1996). Such particles can also affect atmospheric photochemistry by attenuating solar ultraviolet radiation, and they can facilitate stratospheric haze formation by providing condensation nuclei in the upper atmosphere (e.g., West 1996, Moses 1992, 1996, Moses et al. 1995b, Hunten et al. 1980). Vapor species introduced from extraplanetary sources can alter stratospheric or ionospheric chemistry (e.g., Prather et al. 1978, Strobel and Yung 1979, Lellouch 1996, Moses 1996, Lyons 1996, Connerney and Waite 1984, Majeed and McConnell 1991). The chemical effects of oxygen compounds, which are abundant in comets, IDPs, icy satellites, and certain planetary rings but are rare in the upper atmospheres of the outer planets, are particularly interesting. The effects of an external oxygen influx on the neutral photochemistry of reducing atmospheres have been considered for Jupiter (Prather et al. 1978, Strobel and Yung, 1979), Titan (Samuelson et al. 1983, Yung et al. 1984, Toublanc et al. 1995, Lara et al. 1996), and Uranus and Neptune (Lyons 1996), but not for Saturn (until a very recent Saturn study by Ollivier et al. 2000). In this paper, we examine the effects of exogenic oxygen on the photochemistry of Saturn’s upper atmosphere. Saturn is a

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good focus for theoretical modeling at this point in time because the ISO observations of Saturn not only demonstrate high signalto-noise detections of CO2 and H2 O, they also provide highquality spectra in which the emission features of CH3 , C2 H2 , C2 H6 , CH3 C2 H, and C4 H2 are readily identified (de Graauw et al. 1997, B´ezard et al. 1998). The extensive wavelength coverage of the ISO data and the simultaneous observations of multiple hydrocarbon compounds help minimize uncertainties in abundance determinations and provide many new constraints for photochemical models (see Moses et al. 2000). We have developed a one-dimensional diurnally averaged model of photochemistry and diffusion in Saturn’s upper atmosphere. Along with traditional hydrocarbon photochemistry, our model includes the photolysis and kinetics of oxygen-bearing species. Our hydrocarbon reaction list (Moses et al. 2000) is loosely based on the Jupiter photochemical model of Gladstone et al. (1996), but we update their reaction rates and pathways, add oxygen photochemistry, include a source of oxygen from the ablation of IDPs (or from incoming ring material), and explicitly consider condensation. Our oxygen reaction list and cross sections are taken from a study of the photochemical evolution of the Comet Shoemaker–Levy 9 (SL9) impact debris on Jupiter (Moses et al. 1995a, Moses 1996). The photochemical model results are used to create synthetic spectra that are then directly compared with ISO observations. The primary goals of our modeling are to elucidate the quantitative details of oxygen photochemistry in Saturn’s stratosphere, to put tighter limits on the abundance of CO2 and H2 O on Saturn, to place constraints on the flux of extraplanetary material into Saturn’s atmosphere, and to determine the possible molecular form and origin of the incoming oxygen. We also discuss implications for the micrometeoroid dust flux in the outer Solar System and for the possibility of an internal source of CO on Saturn. Our results regarding the photochemistry of hydrocarbon compounds and implications for the strength of atmospheric mixing on Saturn are discussed in a companion paper (Moses et al. 2000). 2. PHOTOCHEMICAL MODEL DESCRIPTION

We use the Caltech/JPL chemical kinetics and diffusion code (e.g., Gladstone et al. 1996, Yung et al. 1984, Allen et al. 1981) to solve the coupled diurnally averaged one-dimensional continuity equations for H, He, and all the carbon- and oxygenbearing compounds in Saturn’s atmosphere. Condensation and evaporation have been added to the kinetics code; technical details of our procedure for including condensation are given in Appendix A. Our photochemical model for Saturn contains the stable oxygen-bearing molecules O2 (molecular oxygen), H2 O (water), CO (carbon monoxide), CO2 (carbon dioxide), H2 CO (formaldehyde), CH3 OH (methanol), H2 CCO (ketene), CH3 CHO (acetaldehyde), as well as the oxygen-bearing radicals O, O(1 D), OH, HCO, CH2 OH, CH3 O, HCCO, CH3 CO, and C2 H4 OH. Water in the condensed phase (H2 O(s) ) is also included. Full details of our Saturn photochemical model,

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TABLE I Oxygen Photolysis Reactions

Reaction R245

hν

Reference

4.0 × 10−9 , 7.2 × 10−10 3 ≤ λ ≤ 253 nm 1.3 × 10−8 , 1.1 × 10−8 117 ≤ λ ≤ 178 nm

Yoshino et al. (1987, 1988, 1992); Lewis et al. (1988a, 1988b); Wang et al. (1987); Hayashi et al. (1986); Black et al. (1985); Johnston et al. (1984); Kley (1984); Nicolet (1984); Gibson et al. (1983); Allen and Frederick (1982); Samson et al. (1982, 1977); Kirby et al. (1979); Carver et al. (1977); Lee et al. (1977); Ogawa and Ogawa (1975); Ackerman (1971); Matsunaga and Watanabe (1967); Ditchburn and Young (1962)

hν

7.3 × 10−8 , 3.7 × 10−8 91 ≤ λ ≤ 193 nm

Nee and Lee (1984); van Dishoeck and Dalgarno (1984); van Dishoeck et al. (1984)

hν

5.3 × 10−8 , 2.5 × 10−8 61 ≤ λ ≤ 198 nm 3.9 × 10−9 , 6.0 × 10−13 80 ≤ λ ≤ 143 nm

Haddad and Samson (1986); Lee and Suto (1986); Pravilov and Shul’pyakov (1986); Slanger and Black (1982); Wu and Judge (1981); Stief et al. (1975); Thompson et al. (1963); Watanabe and Zelikoff (1953)

O2 → 2O hν

→ O + O(1 D)

R246

Photolysis rate J (s−1 ) at 10−8 mbar, at 10−3 mbar

R247

OH → O + H

R248

H2 O → H + OH

R249

→ 2H + O

R250

→ H2 + O(1D)

hν

3.3 × 10−9 , 7.0 × 10−12 80 ≤ λ ≤ 143 nm

hν

2.8 × 10−9 , 1.5 × 10−18 39 ≤ λ ≤ 112 nm

Stark et al. (1991); Samson and Haddad (1988); Letzelter et al. (1987); Wright et al. (1976)

hν

6.5 × 10−11 , 6.2 × 10−11 63 ≤ λ ≤ 208 nm 2.4 × 10−9 , 3.8 × 10−10 63 ≤ λ ≤ 208 nm

Lewis and Carver (1983); Hitchcock and Brion (1980); Okabe (1978); Slanger and Black (1978); Lawrence (1972a,b); Shemansky (1972); Nakata et al. (1965)

hν

R251

CO → C + O

R252

CO2 → CO + O hν

→ CO + O(1D)

R253

hν

5.2 × 10−6 , 5.2 × 10−6 193 ≤ λ ≤ 610 nm

K¨onig et al. (1984); Hochanadel et al. (1980); Okabe (1978)

hν

4.2 × 10−7 , 4.2 × 10−7 238 ≤ λ ≤ 345 nm 5.4 × 10−7 , 5.1 × 10−7 61 ≤ λ ≤ 360 nm

DeMore et al. (1992, 1990, 1987, 1985); Suto et al. (1986); Okabe (1978); Glicker and Stief (1971); Mentall et al. (1971); Gentieu and Mentall (1970)

R254

HCO → H + CO

R255

H2 CO → HCO + H

R256

→ H2 + CO

R257

→ 2H + CO

hν

1.8 × 10−7 , 1.6 × 10−7 61 ≤ λ ≤ 238 nm

hν

2.8 × 10−9 , 1.2 × 10−9 105 ≤ λ ≤ 198 nm

hν

R258

CH3 OH → CH3 + OH

R259

→ H2 CO + H2

hν

1.0 × 10−8 , 3.7 × 10−9 105 ≤ λ ≤ 198 nm

R260

→ CH3 O + H

hν

4.0 × 10−8 , 1.5 × 10−8 105 ≤ λ ≤ 198 nm

hν

1.5 × 10−5 , 1.5 × 10−5

Estimate J = J262

1.5 × 10−5 , 1.5 × 10−5 270 ≤ λ ≤ 375 nm

Okabe (1978)

hν

3.8 × 10−8 , 3.8 × 10−8 105 ≤ λ ≤ 198 nm

Atkinson et al. (1992)

hν

2.9 × 10−7 , 2.9 × 10−7 105 ≤ λ ≤ 198 nm

R261 R262 R263 R264

HCCO → CH + CO H2 CCO

hν → 1 CH2

+ CO

CH3 CHO → CH4 + CO → CH3 + HCO

including a description of our inferred eddy diffusion coefficient profile and a discussion of the model boundary conditions, are presented in Moses et al. (2000). The photodissociation reactions for the oxygen-containing molecules in the photochemical model are shown in Table I along with (i) photolysis rates (J values) at the top of Saturn’s atmosphere and at 1.1 × 10−3 mbar, (ii) the wavelength range

Satyapal et al. (1989); Nee et al. (1985); Person and Nicole (1971); Hag`ege et al. (1968)

in which the reactions are important, and (iii) the references for the adopted cross sections. The cross sections are taken from the Moses (1996) model of the photochemical evolution of the Shoemaker–Levy 9 impact sites on Jupiter and are similar to the cross sections adopted by Yung et al. (1984) for their Titan model. Table II lists the complete set of oxygen reactions in our Saturn model. The extensive databases of Baulch et al. (1992,

EFFECTS OF EXTRAPLANETARY OXYGEN ON SATURN PHOTOCHEMISTRY

TABLE II Oxygen Reactions Reaction M

R265 R266 R267 R268 R269 R270 R271 R272 R273 R274 R275 R276 R277 R278 R279 R280 R281 R282 R283 R284 R285 R286 R287

O+H O + H2 O + CH O + 3 CH2

R288 R289 R290 R291 R292 R293 R294 R295 R296 R297 R298 R299 R300 R301 R302 R303 R304 R305 R306 R307 R308 R309 R310 R311 R312 R313

O + HCO → → O + H2 CO → O + CH2 OH → O + CH3 O → → O + CH3 OH → O + HCCO → O + H2 CCO → → → O + CH3 CO → → O + CH3 CHO → O + C2 H4 OH → O(1 D) + H2 → O(1 D) + CH4 → → O(1 D) + H2 O → O(1 D) + CO2 → O2 + C → O2 + CH → → O2 + 3 CH2 → → M OH + H →

O + CH3 O + CH4 O + C2 H O + C2 H2 O + C2 H3 O + C2 H4

O + C2 H5 O + C2 H6 2O O + OH O + CO

→ → → → → → → → → → → → → → → → → → → → M → → M →

OH OH + H CO + H CO + 2H CO + H2 H2 CO + H OH + CH3 CO + CH CO + 3 CH2 HCCO + H H2 CCO + H OH + C2 H2 HCO + 3 CH2 CH3 CO + H HCO + CH3 H2 CO + 3 CH2 H2 CCO + H2 CH3 CHO + H H2 CO + CH3 OH + C2 H5 O2 O2 + H CO2 OH + CO CO2 + H OH + HCO OH + H2 CO O2 + CH3 OH + H2 CO OH + CH2 OH 2CO + H CO + H2 CO 2HCO CO + HCO + H CO2 + CH3 OH + H2 CCO OH + CH3 CO OH + CH3 CHO OH + H OH + CH3 H2 CO + H2 2OH O + CO2 O + CO OH + CO O + HCO OH + CO + H H2 O + CO H2 O

R314 R315 R316 R317

OH + H2 → H2 O + H OH + 3 CH2 → H2 CO + H OH + CH3 → H2 O + 1 CH2 M → CH3 OH

R318

OH + CH4 → H2 O + CH3

Rate constant

Reference

k0 = 1.3 × 10−29 T −1 8.49 × 10−20 T 2.67 e(−3160/T ) 6.6 × 10−11 1.2 × 10−10 8.0 × 10−11 1.4 × 10−10 1.5 × 10−15 T 1.56 e(−4270/T ) 1.7 × 10−11 1.5 × 10−11 e(−1600/T ) 1.5 × 10−11 e(−1600/T ) 2.0 × 10−11 2.0 × 10−11 2.0 × 10−11 2.0 × 10−18 T 2.08 3.45 × 10−18 T 2.08 1.5 × 10−19 T 2.08 1.5 × 10−19 T 2.08 8.3 × 10−11 1.7 × 10−11 1.66 × 10−15 T 1.5 e(−2920/T ) k0 = 5.2 × 10−35 e(900/T ) 2.3 × 10−11 e(110/T ) k0 = 1.7 × 10−33 e(−1510/T ) k∞ = 2.66 × 10−14 e(−1459/T ) 5.0 × 10−11 5.0 × 10−11 6.85 × 10−13 T 0.57 e(−1390/T ) 1.5 × 10−10 3.55 × 10−11 e(−239/T ) 3.0 × 10−12 1.63 × 10−11 e(−2267/T ) 1.6 × 10−10 1.3 × 10−12 e(−680/T ) 1.3 × 10−12 e(−680/T ) 1.3 × 10−12 e(−680/T ) 2.4 × 10−10 7.0 × 10−11 1.8 × 10−11 e(−1100/T ) 1.5 × 10−10 1.1 × 10−10 1.35 × 10−10 1.5 × 10−11 2.2 × 10−10 7.4 × 10−11 e(120/T ) 1.6 × 10−11 2.75 × 10−11 2.75 × 10−11 1.0 × 10−12 4.0 × 10−13 k0 = 6.1 × 10−26 T −2 k∞ = 2.69 × 10−10 e(−75/T ) 7.7 × 10−12 e(−2100/T ) 3.0 × 10−11 1.0 × 10−12 k0 = 6.4 × 10−29 e(1033/T ) k∞ = 1.44 × 10−10 T 0.1 3.9 × 10−12 e(−1885/T )

Tsang and Hampson 1986 Baulch et al. 1992 Baulch et al. 1992 Baulch et al. 1992 Baulch et al. 1992 Atkinson et al. 1992 Baulch et al. 1992 Baulch et al. 1992 DeMore et al. 1987 DeMore et al. 1987 Baulch et al. 1992 Baulch et al. 1992 Baulch et al. 1992 Baulch et al. 1992 Baulch et al. 1992 Baulch et al. 1992 Baulch et al. 1992 Baulch et al. 1992 Baulch et al. 1992 Baulch et al. 1992 Tsang and Hampson 1986 Atkinson et al. 1992 Tsang and Hampson 1986 Simonaitis and Heicklen 1972 Baulch et al. 1992 Baulch et al. 1992 Baulch et al. 1992 Grotheer et al. 1989 Cobos and Troe 1985 Baulch et al. 1992 Failes et al. 1982 Baulch et al. 1992 Baulch et al. 1992 Baulch et al. 1992 Baulch et al. 1992 Miyoshi et al. 1989 Miyoshi et al. 1989 DeMore et al. 1987 Herron 1988 Atkinson et al. 1992 Atkinson et al. 1992 Atkinson et al. 1992 DeMore et al. 1987 Atkinson et al. 1992 Baulch et al. 1992 Baulch et al. 1992 Baulch et al. 1992 Tsang and Hampson 1986 Tsang and Hampson 1986 Baulch et al. 1992 Cobos and Troe 1985 Atkinson et al. 1992 Tsang and Hampson 1986 Est., Oser et al. 1992 Oser et al. 1992 Fagerstr¨om et al. 1993 Atkinson et al. 1992

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TABLE II—Continued Reaction R319 R320 R321

OH + C2 H → O + C2 H2 → CO + 3 CH2 M OH + C2 H2 → CH3 CO

R322 R323

OH + C2 H3 → H2 O + C2 H2 M OH + C2 H3 → CH3 CHO

R324

OH + C2 H4 → C2 H4 OH

R325 R326 R327 R328 R329 R330 R331 R332 R333 R334 R335 R336 R337 R338 R339 R340 R341

M

OH + C2 H5 OH + C2 H5 OH + C2 H6 OH + C3 H8 2OH OH + CO OH + HCO OH + H2 CO OH + CH2 OH OH + CH3 O OH + CH3 OH OH + H2 CCO OH + CH3 CO OH + CH3 CHO H2 O + CH

→ → → → → → → → → → → → → → → → M →

H2 O + C2 H4 O + C2 H6 H2 O + C2 H5 H2 O + C3 H7 O + H2 O CO2 + H H2 O + CO H2 O + HCO H2 O + H2 CO H2 O + H2 CO H2 O + CH2 OH H2 O + CH3 O CO + CH2 OH HCO + H2 CO H2 O + H2 CCO H2 O + CH3 CO CH2 OH

M

R342

CO + H → HCO

R343

CO + CH3 → CH3 CO

M

R344 R345 R346 R347 R348 R349

CO2 + CH CO2 + 3 CH2 HCO + H HCO + 3 CH2 HCO + CH3 HCO + CH3

→ → → → → M →

CO + HCO CO + H2 CO CO + H2 CO + CH3 CO + CH4 CH3 CHO

R350 R351 R352 R353 R354 R355 R356 R357 R358 R359 R360 R361 R362 R363 R364 R365 R366 R367 R368 R369 R370

HCO + C2 H HCO + C2 H3 HCO + C2 H5 2HCO HCO + CH2 OH

→ → → → → → → → → → → → → → → → → → → → →

CO + C2 H2 CO + C2 H4 CO + C2 H6 CO + H2 CO 2H2 CO CO + CH3 OH CO + CH3 OH CO + CH3 CHO HCO + H2 CO + CH3 H2 CCO + H HCO + CH4 HCO + C2 H4 HCO + C2 H6 HCO + CH3 OH HCO + CH3 OH OH + CH3 H2 CO + H2 OH + C2 H4 H2 CO + CH3 H2 CO + CH4

HCO + CH3 O HCO + CH3 CO H2 CO + H H2 CO + CH H2 CO + CH3 H2 CO + C2 H3 H2 CO + C2 H5 H2 CO + CH2 OH H2 CO + CH3 O CH2 OH + H CH2 OH + 3 CH2 CH2 OH + CH3

Rate constant

Reference

3.0 × 10−11 3.0 × 10−11 k0 = 2.6 × 10−26 T −1.5 k∞ = 1.0 × 10−17 T 2 5.0 × 10−11 k0 = 1.0 × 10−31 k∞ = 5.0 × 10−11 k0 = 9.59 × 10−27 T −0.8 k∞ = 8.79 × 10−12 4.0 × 10−11 1.66 × 10−40 T 8.8 e(−250/T ) 7.8 × 10−12 e(−1020/T ) 9.8 × 10−12 e(−640/T ) 4.2 × 10−12 e(−240/T ) 1.5 × 10−13 1.7 × 10−10 8.8 × 10−12 e(25/T ) 4.0 × 10−11 3.0 × 10−11 5.0 × 10−12 e(−600/T ) 1.68 × 10−12 e(−600/T ) 1.0 × 10−11 7.0 × 10−12 2.0 × 10−11 5.6 × 10−12 e(310/T ) k0 = 1.0 × 10−31 k∞ = 9.49 × 10−12 e(380/T ) k0 = 1.4 × 10−34 e(−100/T ) k∞ = 1.96 × 10−13 e(−1366/T ) k0 = 1.26 × 10−33 e(−1636/T ) k∞ = 2.63 × 10−13 e(−3007/T ) 5.7 × 10−12 e(−345/T ) 3.9 × 10−14 1.5 × 10−10 3.0 × 10−11 4.4 × 10−11 k0 = 1.0 × 10−31 k∞ = 5.0 × 10−11 1.0 × 10−10 1.5 × 10−10 2.0 × 10−10 5.0 × 10−11 3.0 × 10−10 2.0 × 10−10 1.5 × 10−10 1.5 × 10−11 3.8 × 10−14 T 1.05 e(−1650/T ) 8.0 × 10−11 e(260/T ) 8.0 × 10−11 e(260/T ) 6.8 × 10−12 e(−4450/T ) 9.01 × 10−21 T 2.81 e(−2950/T ) 9.13 × 10−21 T 2.81 e(−2950/T ) 9.1 × 10−21 T 2.8 e(−2950/T ) 1.7 × 10−13 e(−1500/T ) 1.6 × 10−10 1.0 × 10−11 4.0 × 10−11 2.0 × 10−12 4.0 × 10−12

Tsang and Hampson 1986 Tsang and Hampson 1986 Atkinson et al. 1992 Atkinson et al. 1992 Tsang and Hampson 1986 Estimate Tsang and Hampson 1986 DeMore et al. 1990 DeMore et al. 1990 Tsang and Hampson 1986 Cohen and Westerberg 1991 Atkinson et al. 1992 Atkinson et al. 1992 DeMore et al. 1987 Atkinson et al. 1992 Baulch et al. 1992 Atkinson et al. 1992 Tsang 1987 Tsang and Hampson 1986 DeMore et al. 1990 DeMore et al. 1990 Baulch et al. 1992 Baulch et al. 1992 Tsang and Hampson 1986 Atkinson et al. 1992 Estimate Zabarnick et al. 1988 Wagner and Bowman 1987 Arai et al. 1981 Anastasi and Maw 1982 Watkins and Word 1974 Baulch et al. 1992 Tsang and Hampson 1986 Baulch et al. 1992 Tsang and Hampson 1986 Mulenko 1987 Estimate Callear and Cooper 1990 Tsang and Hampson 1986 Tsang and Hampson 1986 Tsang and Hampson 1986 Baulch et al. 1992 Tsang 1987 Tsang 1987 Tsang and Hampson 1986 Tsang and Hampson 1986 Baulch et al. 1992 Est., Baulch et al. 1992 Est., Baulch et al. 1992 Baulch et al. 1992 Tsang 1987 Tsang 1987 Tsang 1987 Tsang and Hampson 1986 Tsang 1987 Tsang 1987 Tsang 1987 Tsang 1987 Tsang 1987

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TABLE II—Continued Reaction R371 R372 R373 R374 R375 R376 R377 R378 R379 R380 R381 R382 R383 R384 R385 R386 R387 R388 R389 R390 R391 R392 R393 R394 R395 R396 R397 R398 R399 R400 R401 R402 R403 R404 R405 R406 R407 R408 R409 R410 R411 R412 R413 R414 R415 R416 R417 R418 R419 R420

CH2 OH + C2 H → → CH2 OH + C2 H3 → → CH2 OH + C2 H5 → → 2CH2 OH → CH2 OH + CH3 O → CH3 O + H → → CH3 O + 3 CH2 → CH3 O + CH3 → CH3 O + C2 H → CH3 O + C2 H3 → CH3 O + C2 H5 → 2CH3 O → CH3 O + CH3 OH → CH3 O + CH3 CO → → CH3 O + CH3 CHO → CH3 OH + H → → CH3 OH + 3 CH2 → → CH3 OH + CH3 → → CH3 OH + C2 H → → CH3 OH + C2 H3 → → CH3 OH + C2 H5 → → HCCO + H → H2 CCO + H → H2 CCO + 3 CH2 → → CH3 CO + H → → CH3 CO + 3 CH2 → CH3 CO + CH3 → → CH3 CO + C2 H → 2CH3 CO → CH3 CHO + H → CH3 CHO + 3 CH2 → CH3 CHO + CH3 → CH3 CHO + C2 H3 → CH3 CHO + C2 H5 → C2 H4 OH + H → C2 H4 OH + CH3 →

OH + C3 H3 H2 CO + C2 H2 OH + C3 H5 H2 CO + C2 H4 CH3 OH + C2 H4 H2 CO + C2 H6 H2 CO + CH3 OH H2 CO + CH3 OH OH + CH3 H2 CO + H2 H2 CO + CH3 H2 CO + CH4 H2 CO + C2 H2 H2 CO + C2 H4 H2 CO + C2 H6 H2 CO + CH3 OH CH2 OH + CH3 OH H2 CCO + CH3 OH H2 CO + CH3 CHO CH3 CO + CH3 OH H2 O + CH3 CH2 OH + H2 CH2 OH + CH3 CH3 O + CH3 CH2 OH + CH4 CH3 O + CH4 CH2 OH + C2 H2 CH3 O + C2 H2 CH2 OH + C2 H4 CH3 O + C2 H4 CH2 OH + C2 H6 CH3 O + C2 H6 CO + 3 CH2 CO + CH3 CO + C2 H4 HCCO + CH3 H2 CCO + H2 HCO + CH3 H2 CCO + CH3 H2 CCO + CH4 CO + C2 H6 H2 CCO + C2 H2 H2 CCO + CH3 CHO CH3 CO + H2 CH3 CO + CH3 CH3 CO + CH4 CH3 CO + C2 H4 CH3 CO + C2 H6 CH3 CHO + H2 CH3 CHO + CH4

Rate constant

Reference

2.0 × 10−11 6.0 × 10−11 2.0 × 10−11 5.0 × 10−11 4.0 × 10−12 4.0 × 10−12 8.0 × 10−12 4.0 × 10−11 7.52 × 10−11 e(−375/T ) 3.38 × 10−11 e(−375/T ) 3.0 × 10−11 4.0 × 10−11 4.0 × 10−11 4.0 × 10−11 4.0 × 10−11 1.0 × 10−10 5.0 × 10−13 e(−2050/T ) 1.0 × 10−11 1.0 × 10−11 6.0 × 10−15 1.8 × 10−17 T 2.1 e(−2450/T ) 1.8 × 10−18 T 2.1 e(−2450/T ) 5.29 × 10−23 T 3.2 e(−3609/T ) 2.39 × 10−23 T 3.1 e(−3490/T ) 5.29 × 10−23 T 3.2 e(−3609/T ) 2.39 × 10−23 T 3.1 e(−3490/T ) 1.0 × 10−11 2.0 × 10−12 5.29 × 10−23 T 3.2 e(−3609/T ) 2.39 × 10−23 T 3.1 e(−3490/T ) 5.29 × 10−23 T 3.2 e(−4610/T ) 2.39 × 10−23 T 3.1 e(−4500/T ) 2.5 × 10−10 3.0 × 10−11 e(−1700/T ) 2.1 × 10−10 1.0 × 10−17 1.92 × 10−11 3.57 × 10−11 3.0 × 10−11 1.0 × 10−11 4.9 × 10−11 3.0 × 10−11 1.49 × 10−11 2.23 × 10−11 e(−1661/T ) 2.76 × 10−12 e(−1768/T ) 3.3 × 10−30 T 5.64 e(−1240/T ) 1.35 × 10−13 e(−1852/T ) 2.09 × 10−12 e(−4277/T ) 8.3 × 10−11 4.0 × 10−11

Tsang 1987 Tsang 1987 Tsang 1987 Tsang 1987 Tsang 1987 Tsang 1987 Tsang 1987 Tsang 1987 D´ob´e et al. 1991 D´ob´e et al. 1991 Tsang 1987 Tsang and Hampson 1986 Tsang and Hampson 1986 Tsang and Hampson 1986 Tsang and Hampson 1986 Tsang and Hampson 1986 Tsang 1987 Tsang and Hampson 1986 Tsang and Hampson 1986 Estimate Tsang 1987 Tsang 1987 Tsang 1987 Tsang 1987 Tsang 1987 Tsang 1987 Tsang 1987 Tsang 1987 Tsang 1987 Tsang 1987 Tsang 1987 Tsang 1987 Baulch et al. 1992 Baulch et al. 1992 Canosa-Mas et al. 1984 Banyard et al. 1980 Ohmori et al. 1990 and Bartels et al. 1991 Tsang and Hampson 1986 Hassinen et al. 1990 and Adachi et al. 1981 Tsang and Hampson 1986 Hassinen et al. 1990 Whytock et al. 1976 B¨ohland et al. 1985 Baulch et al. 1992 Scherzer et al. 1987 H¨ohlein and Freeman 1970 Bartels et al. 1982 Estimate

Note. Two-body rate constants for reaction i (denoted ki ) and high-pressure limiting rate constants for three-body reactions (k∞ ) are in units of cm3 s−1 . Low-pressure limiting rate constants for three-body reactions (k0 or ki,0 ) are in units of cm6 s−1 . M represents any third body such as H2 .

1994), Mallard et al. (1994), Atkinson et al. (1992), DeMore et al. (1985, 1987, 1990, 1992), and Tsang (1987) were useful in acquiring the necessary rate-constant information. Because of the relevance of such reactions to the Earth’s stratosphere, the low-temperature rate coefficients for the oxygen reactions

are generally better known than the corresponding case for the hydrocarbon reactions. As a result, the oxygen chemistry in our model is better constrained than the hydrocarbon chemistry. To simulate the effects of micrometeoroid bombardment on Saturn, we add an ablation source of oxygen compounds to

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FIG. 1. The ablation rate (g per km altitude) of small spherical interplanetary dust particles entering Saturn’s atmosphere at an assumed initial velocity of 37 km s−1 and an impact angle of 45◦ for pure silicate grains (top) and pure water ice grains (bottom) (after Moses 1998). The different curves are for different assumed initial particle masses (in micrograms) as labeled. The numbers on the right axis indicate altitude (in km) above the 1-bar level as determined by solving the hydrostatic equilibrium equation for 30◦ N latitude (see Moses et al. 2000).

EFFECTS OF EXTRAPLANETARY OXYGEN ON SATURN PHOTOCHEMISTRY

the upper atmosphere. Moses (1998) has calculated the ablation rates of small (3 to 400 µm) micrometeoroidal dust and ice particles that enter Saturn’s atmosphere at velocities ranging from 35 to 38 km s−1 (i.e., for relative encounter velocities at infinity that range from ∼1 to 15 km s−1 ), and we use these results in our model (Fig. 1). Ablation temperatures for icy grains are low, and the molecules will evaporate intact; that is, H2 O will be ablated from water ice, CO2 will be ablated from carbon dioxide ice, etc. The ablated molecules will still be traveling at a large fraction of their initial entry velocity, and although subsequent collisions with atmospheric molecules or other ablated vapor molecules could alter the composition of the volatile ablation products, we consider the oxygen to be introduced only as H2 O, CO, CO2 , and/or CH3 OH—all of which are ices known to be present in comets (e.g., Greenberg 1998)—or as atomic oxygen. In our nominal model, these molecules are introduced in the 10−6 to 10−4 mbar region, with relative fluxes (molecules cm−2 s−1 ) of H2 O : CO : CO2 : CH3 OH equal to 100 : 50 : 5 : 5 and a deposition rate that is uniform within the ablation region. Note that the CO/H2 O ratio here is higher than that implied by the composition of cometary ices (e.g., Crovisier 1993, Greenberg 1998); this issue is discussed more fully in Section 5.2. We also examine the sensitivity of our results to changes in the ablation altitude, relative fluxes, and molecular form of the oxygen-bearing material. Carbon monoxide is assumed to have an internal source in our nominal model; the CO mixing ratio at the lower boundary is fixed at 1.0 × 10−9 (Noll and Larson 1990). 3. PHOTOCHEMICAL MODEL RESULTS

The important pathways for synthesis and interconversion of the oxygen species in our nominal model are illustrated in Fig. 2. Oxygen is introduced to the atmosphere in the form of H2 O, CO, CO2 , and CH3 OH; however, the photolysis of H2 O, in particular, initiates the subsequent oxygen photochemistry in Saturn’s upper atmosphere. Photolysis lifetimes at the top of the model atmosphere for H2 O (6 months), CO (11 years), CO2 (13 years), and CH3 OH (7 months) are relatively long (see Table I); in addition, these molecules are shielded below the methane homopause on Saturn by hydrogen, methane, and other hydrocarbons. As a result, the relevant time scales for the photochemistry of oxygen compounds in Saturn’s atmosphere tend to be longer than those for hydrocarbon compounds. The mixing ratios (species concentration divided by total atmospheric density) for all the interesting oxygen compounds in our nominal model are shown in Figs. 3 and 4. The most important radical oxygen species in the model are OH and O. These radicals are often critical to the formation of more complex oxygen compounds. From the standpoint of the total column density, CO is the most abundant stable oxygen compound in our model (in the stratosphere as well as the entire model), followed by H2 O, CO2 , CH3 OH, H2 CO, and CH3 CHO. Carbon monoxide is abundant because we assume it has an internal source, because

173

FIG. 2. A schematic diagram illustrating the important reaction pathways for oxygen species in our nominal model. The symbol hν corresponds to a solar ultraviolet photon. Radical species are outlined as ovals and stable molecules as rectangles.

we assume that much of the ablated debris is either released as CO or quickly forms CO (either by subsequent impacts or by chemistry in the meteor trail), and because CO is the most stable of the oxygen compounds and is the end product of much of the oxygen photochemistry in Saturn’s stratosphere (see also Prather et al. 1978, Strobel and Yung 1979). Water is abundant because we assume it is a major product from meteoroid ablation and because it is very efficiently recycled in Saturn’s stratosphere.

FIG. 3. The mixing ratios of important oxygen-bearing radicals and molecules in our nominal model as a function of pressure and altitude. The data point (square) is from our reanalysis of the ISO H2 O observations (see Table IV and Section 4.1).

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FIG. 4. The mixing ratios of important oxygen-bearing molecules in our nominal model as a function of pressure and altitude. The data points and associated error bars are from our reanalysis of ISO and IRTF observations (see Table IV). The CO2 data point is marked by a circle, and the CO point by a triangle.

Much of the CO2 and CH3 OH (39 and 77%, respectively) and all of the H2 CO observed in our nominal model is generated from photochemical processes rather than resulting from a meteoric source. The sharp decrease in the mixing ratio of H2 O at ∼0.4 mbar in Fig. 3 is a result of condensation. Chemical loss processes for H2 O do not permanently remove water from Saturn’s upper atmosphere, so that condensation becomes an important loss mechanism for stratospheric H2 O (see also Feuchtgruber et al. 1997). Evaporation of H2 O ice particles at pressure levels greater than ∼300 mbar reintroduces water vapor to the lower atmosphere. Although CH3 OH also becomes supersaturated near the tropopause in our nominal model, we have neglected methanol condensation because the effects will be minor both with regard to changes in the CH3 OH mixing ratio profile and to the inferred condensation flux. In Saturn’s upper atmosphere, it is easier to break hydrogen– oxygen bonds than carbon–oxygen bonds. The efficiency of H2 O recycling helps maintain O–H bonds. Figure 5 shows the top five reactions that are responsible for the production, destruction, and

FIG. 5. Reaction rate profiles for the top five reactions responsible for the production (left), loss (middle), and exchange (right) of carbon–oxygen bonds in our nominal model. The reactions are listed in order of decreasing column reaction rate in the atmosphere.

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EFFECTS OF EXTRAPLANETARY OXYGEN ON SATURN PHOTOCHEMISTRY

TABLE III Important Net Production and Loss Schemes Species H2 O

CO2

CO

CH3 OH

H2 CO

Dominant reaction schemes

Notes

Production

Ablation

Loss

H2 O + C2 H2 + H → CO + CH3 + H2 (7) H2 O + C2 H2 → CO + 1 CH2 + H2 (8)

Efficiently recycled.

Production

Ablation H2 O + CO → CO2 + 2 H (9)

Ablation is 61% of the total production.

Loss

CO2 + 2 H2 → CO + H2 O + 2 H (11) CO2 + CH3 → 2 CO + H2 + H (12)

Recycling is inefficient. Photolysis dominates loss.

Production

Ablation H2 O + C2 H2 + H → CO + CH3 + H2 (7) H2 O + C2 H2 → CO + 1 CH2 + H2 (8) H2 O + C2 H4 + H → CO + CH3 + 2 H2 (14)

Ablation represents 35% of the total stratospheric production. The CO chemistry is complex.

Loss

OH + CO → CO2 + H R330 M CO + H → HCO R342

Difficult to destroy. Efficiently recycled.

Production

CH3 + OH → CH3 OH R317 Ablation

Ablation represents 23% of the total production.

Loss

CH3 OH + H2 → H2 O + CH3 + H (22) CH3 OH → CO + 2H2 (23)

Both H2 O and CO are ultimate products.

Production

CO + 2H → H2 CO (26) H2 O + CH3 → H2 CO + H2 + H (27), (28)

Loss

H2 CO → H2 + CO R256

M

hv

CO is ultimate product.

hv

H2 CO → HCO + H R255 H2 CCO

CH3 CHO

Production

H2 O + C2 H2 → H2 CCO + H2 (29)

Loss

H2 CCO → 1 CH2 + CO R262

Production

H2 O + C2 H4 → CH3 CHO + H2 (30)

Loss

CH3 CHO → CH3 + HCO R264

hv

CO is ultimate product.

hv

CO is ultimate product.

hv

CH3 CHO → CH4 + CO R263

exchange of carbon-oxygen bonds in our nominal model. The reaction of OH with hydrocarbons (mainly C2 H2 , CH3 , and C2 H4 ) is responsible for the bulk of the photochemical production of C–O bonds in our model, although reactions of atomic O with hydrocarbons could become important if ground-state oxygen atoms are a major initial source of the extraplanetary oxygen. Photolysis and cracking by atomic hydrogen help break C–O bonds in Saturn’s atmosphere. Numerous exchange reactions control the abundances of C–O-bearing species in our model; reactions involving CO, HCO, and H2 CO dominate. Complete details of the production and loss mechanisms for the important oxygen compounds can be found in Appendix B. A summary of the important net production and loss mechanisms for these compounds is presented in Table III (where the scheme numbers are from Appendix B).

millimeter observations of CO (Noll and Larson 1990, Rosenqvist et al. 1992). 4.1. ISO Observations of CO2 and H2 O Our photochemical model results are used to create synthetic spectra that are then directly compared with 2.4- to 45.2-µm spectra of Saturn obtained with the grating mode of the ShortWavelength Spectrometer (SWS)1 of the Infrared Space Observatory.2 A description of the SWS instrument is provided in de Graauw et al. (1996), and a description of the ISO satellite is presented in Kessler et al. (1996). Details of the observations relevant to the hydrocarbon modeling are given in Moses et al. (2000); here, we provide information on the observation of oxygen compounds. A spectrum of Saturn covering the 12.5- to 16-µm range was recorded in the AOT 06 observing mode on 6 December 1996

4. COMPARISONS WITH OBSERVATIONS 1

We now compare the predictions from our photochemical models with ISO observations of CO2 and H2 O (e.g., de Graauw et al. 1997, Feuchtgruber et al. 1997, 1999) and with IRTF and

The SWS instrument is a joint project of the SRON and the MPE. ISO is an ESA project with instruments funded by ESA Member States (especially the PI countries France, Germany, the Netherlands, and the United Kingdom) with the participation of NASA and ISAS. 2

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FIG. 6. Comparison of the CO2 ν2 emission band as observed by ISO on 6 December 1996 (LW band; AOT 06) with a synthetic spectrum based on the nominal photochemical model. All emission features besides CO2 are from the C2 H2 ν5 band.

(revolution 386 of ISO) with a resolution of 1600 at ∼15 µm, where emission due to the ν2 band of CO2 was detected (Fig. 6). The total integration time was 32 min. Rotational lines of water were detected in spectra of Saturn on several different dates, particularly on 11 December 1996 (revolution 391) by using the AOT 02 mode (grating line profile scan), which allows deep integration on single lines. Water emissions were clearly detected at 30.90, 35.94, 39.37 (which actually includes two unresolved features at 39.375 and 39.379 µm and a nearby weak emission at 39.399 µm), 40.34, 43.89, 44.19, and 45.11 µm (Fig. 7). A doublet near 35.45 µm was also marginally detected, as well as emission near 33.0 µm. The latter feature is not usable, however,

due to strong contamination by structure in the instrumental response spectral function at this wavelength. Note also that with the AOT 02 mode, the local shape of the continuum in the vicinity of the emission lines is not always reliable, as was pointed out by Feuchtgruber et al. (1997). 00 00 The SWS aperture (14 × 27 at 12–16.5 µm and 20 × 33 at 29–45.2 µm) was centered on Saturn’s disk (equatorial di00 ameter 18 ) with the long axis oriented along celestial north. The ISO observations are thus representative of disk-averaged conditions. The absolute accuracy of the flux scale for the AOT 06 measurements is estimated to be ∼20% at 15 µm and ∼30% at 29–45.2 µm. The absolute calibration of the AOT 02 measurements is even more uncertain for our particular case, so that in practice, only the line-to-continuum ratio of the H2 O features can be used to test the models. A line-by-line radiative transfer code was used to create synthetic spectra from the predicted concentrations of our photochemical model. Details of these calculations and model assumptions (e.g., temperature profile) are given in Moses et al. (2000). In the wavelength region in which the CO2 emission is observed, the molecular opacity derives from C2 H2 , CO2 , and collision-induced H2 –He absorption. In Fig. 6, we compare the observed ISO CO2 emission feature with synthetic spectra generated from our nominal model. The calculations are in excellent agreement with the data. As indicated by the contribution functions (as defined in Moses et al. (2000) and shown in Fig. 8), emission at the peak of the CO2 band (14.98 µm) derives from the 1.2-mbar pressure region on Saturn (full width half maximum from the 0.14- to 17-mbar region). The CO2 emission is optically thin, and its intensity is to first order proportional to the CO2 column density above the ∼10-mbar level. Below 10 mbar, the atmosphere is too cold to contribute significantly to the emission. No vertically resolved information can be retrieved from the data. The observed and calculated CO2 mixing ratios and column densities are given in Table IV. The uncertainties attached to the CO2 values were determined the same way as in Moses et al. (2000): we assigned a 10% uncertainty due to propagation of temperature errors and another 10% uncertainty from the radiative transfer model itself, leading to a ±15% combined error bar. Longward of 30 µm, the gaseous opacity is due to spectral features of water, ammonia, and the continuum collision-induced H2 –He absorption. However, Voyager IRIS observations indicate that cloud opacity also influences the spectral flux in this

TABLE IV Constraints from ISO Spectra Column density (molecules cm−2 )

Mixing ratio

Molecule

Model prediction

Observed

Model prediction

Observed

H2 O CO2 CO

0–100 mb: 1.4 × 1015 0–10 mb: 6.5 × 1014 100– 400 mb: 1.3 × 1017

(1.4 ± 0.4) × 1015 (6.3 ± 1) × 1014 (1.1 ± 0.4) × 1017

0.1 mb: 1 × 10−8 1 mb: 4.0 × 10−10

(1 ± 0.3) × 10−8 (4 ± 0.6) × 10−10

EFFECTS OF EXTRAPLANETARY OXYGEN ON SATURN PHOTOCHEMISTRY

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FIG. 7. ISO/SWS AOT 02 observations showing the detection of at least seven H2 O lines at 30–45.2 µm. The thick solid histograms represent the data, the thin solid lines represent our nominal model, and the dashed lines represent a model with enhanced water influx of 1 × 108 cm−2 s−1 restricted to 28% of the projected disk (see text).

wavelength range (Conrath and Pirraglia 1983, Courtin et al. 1984). We included an absorbing cloud (meant to represent the NH3 cloud) in our model by specifying a cloud base at 1 bar and a cloud scale height equal to the atmospheric scale height. The cloud absorption optical depth τcl was estimated by fitting the model to a large average of well-calibrated Voyager IRIS spectra (specifically the sample B of Courtin et al. 1984, average airmass = 1.196). In this manner, we obtained τcl = 0.4. Applying the same model to the geometry of the ISO observations yielded fluxes at 29–45 µm, differing from the ISO (AOT 06) values by at most 10%, comfortably within the calibration uncertainty. We thus conservatively estimate that the overall accuracy of our continuum model is 10%. Ammonia lines belonging to the J = 11 multiplet are seen in the ISO spectrum near 42.3 µm and can be fit with a NH3 profile undersaturated by a factor of 4. Ammonia was included in the model although its influence at wavelengths where the water lines are detected is very small.

In Fig. 7, the synthetic spectrum based on the nominal photochemical and radiative model (thin solid line) is compared to the observations (thick solid histogram) in the vicinity of the 7 detected water features. A good agreement is obtained, as all features can be reproduced within the noise level, although the model tends to slightly overpredict the 40.34-µm line and underpredict the 45.11-µm line. For this model, the H2 O lines are optically thin or moderately thick (optical depth in vertical viewing and at infinite spectral resolution: 0.19, 0.36, 1.25, 0.42, 0.70, 2.60, 0.85, and 1.52 at 30.90, 35.94, 39.375, 39.379, 40.34, 43.89, 44.19, and 45.11 µm, respectively). In Fig. 8, we show that the contribution functions for the weakest (30.90 µm) and the strongest (43.89 µm) of the well-detected lines considerably overlap. This overlap, combined with the noise level in the data, prevents reliable vertical information from being extracted; explicitly, a profile with a uniform H2 O mixing ratio of 7 × 10−9 at P ≤ 0.5 mbar provides an identical match to the data.

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FIG. 8. Contribution functions (as defined in Moses et al. 2000) calculated at the peak of the ν2 CO2 band and the peaks of the 30.9- and the 43.9-µm H2 O lines. The maximum contribution to the water-vapor line emissions occurs where the H2 O contribution function is a maximum in the stratosphere (at 0.2 and 0.02 mbar, respectively). As can be seen from the broad plateau in the CO2 contribution function, the CO2 emission originates from the 0.1- to 10-mbar region. Deeper in the troposphere, the contribution functions exhibit peaks due to H2 –He collision-induced opacity.

Uncertainties in the derived H2 O mixing profile are due to (i) the noise level in the data, (ii) continuum model uncertainties, and (iii) thermal profile uncertainties (uncertainties due to molecular parameters for water are negligible). Tests in which the nominal H2 O profile was doubled or halved at all levels suggest that for a given line, the observational noise induces a ∼40% uncertainty in the mixing ratio profile (weak lines are relatively more affected by noise than strong lines, but being more optically thin, they show increased sensitivity to the H2 O profile). With 7 well-detected lines, the resulting uncertainty is about 15%. Thermal profile uncertainties were assessed by testing thermal profile changes as in B´ezard et al. (1998). The resulting uncertainty in the H2 O column is 20%. The effect is larger than for CO2 (see above) mainly because H2 O probes a pressure range (∼0.1 mbar) in which the temperature is less constrained than at 1 mbar, the pressure sounded by CO2 . Finally, as the modeled continuum is itself estimated to be uncertain by 10%, the observed line contrasts are (regardless of noise) also uncertain by 10%. As the lines are not exactly optically thin, a conservative estimate of the effect is 15% on the abundance. Adding quadratically the three sources of errors, the global uncertainty on the H2 O mixing profile and column density is 30%. Thus, we conclude that based on the nominal photochemical profile of this paper, the H2 O stratospheric column density is (1.35 ± 0.4) × 1015 cm−2 , and the mixing ratio at 0.1 mbar is (1.0 ± 0.3) × 10−8 (see Table IV). If we were to assume a mixing ratio profile that is constant with altitude above 0.5 mbar, the column density would be (1.2 ± 0.35) × 1015 cm−2 and the

mixing ratio would be (0.7 ± 0.2) × 10−8 . All these numbers are fully consistent with Feuchtgruber et al. (1997). Far-ultraviolet Hubble Space Telescope (HST) observations (Fouchet et al. 1996, Prang´e et al. 1997, 1998) indicate a broad ˚ in Saturn’s spectrum (see also absorption feature below 2000 A the earlier observations of this spectral region by the International Ultraviolet Explorer (IUE); e.g., Winkelstein et al. 1983). The main contribution to this spectral feature comes from C2 H2 , but other absorbers need to be added to account for the shape ˚ Absorbers such of the absorption between 1550 and 1800 A. as H2 O, CH3 C2 H, C4 H2 , C2 H4 , and/or aerosols are particularly promising candidates for contributing to this spectral feature. The disk-resolved HST observations further indicate that the absorption is deeper at mid- and high latitudes (planetocentric latitudes 33◦ S, 41.5◦ S, and 52◦ S) compared to a reference spectrum taken at 15◦ N. When the C2 H2 abundance, constrained by ˚ and by a series of discrete the spectral gradient above 1800 A absorption bands, is accounted for, Prang´e et al. (1997, 1998) can fit the additional absorption by including water alone, minor hydrocarbons alone, or a combination of the two. If the extra absorption is assumed to be entirely due to water, a column density of 3 × 1016 H2 O molecules cm−2 is suggested for the spectrum taken at the latitude magnetically connected to the C and the inner edge of the B ring (33◦ S), whereas upper limits of a few times 1015 molecules cm−2 are suggested elsewhere (Prang´e et al. 1998). Thus, these observations may show the signature of water precipitating from the rings and entering the atmosphere at specific latitudes.

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Can the H2 O emission observed by ISO be caused by enhanced amounts of water in regions covering a fraction of the observed disk of Saturn? We have tested this scenario against the ISO data by computing a synthetic spectrum at 29–45 µm for a water distribution obtained with an anomalously high H2 O flux of 1 × 108 cm−2 s−1 (with the fluxes of all other oxygenbearing species kept the same as our nominal model). Such a flux, while being ∼67 times higher than our earlier estimate of the globally averaged H2 O influx, is 20 times lower than that advocated by Connerney and Waite (1984) at the inner edge of Saturn’s B ring. An influx of 108 H2 O molecules cm−2 s−1 has a column density ∼1.5 times larger than that reported from the HST observations and may therefore represent a typical ring precipitation case. Because this water flux, if pertaining to the entire planet, would obviously grossly overestimate the observed ISO emissions, an adjustable scaling factor was applied to the line contrasts, in essence representing the fraction of Saturn’s projected disk (“filling factor”) subject to this flux. For models like these, a filling factor of 0.28 provides the best overall match to the entire dataset; the resulting model is presented as the dashed line in Fig. 7. Note, however, that as compared to the nominal model and the observations, this model tends to overestimate the weak lines and to underestimate the strong lines. In addition, the enhanced-flux model increases the discrepancy (already present for the nominal model) in the red wing of the 39.37-µm multiplet, where the discrepancy here is due to the appearance of a shoulder corresponding to the weak 39.399-µm line becoming optically thick. Another argument against this model may be the large value of the necessary filling factor (28% of the projected area) compared with the spatial extent of the latitudes that map back to the regions of instability within the ring system—the dark bands interpreted by Connerney (1986) as bearing the signature of precipitating water are narrowly confined to only a few degrees in latitude. On these bases, we regard the localized precipitation scenario as less likely than the horizontally uniform model to explain the ISO observations. A detailed comparison between the infrared and ultraviolet observations remains to be conducted. However, such a comparison will require full latitudinal coverage by HST (notably with STIS). For the time being, our results do not imply a misinterpretation of the HST data in terms of enhanced water at latitudes connected to the rings. Although the H2 O column density suggested by the HST observations (Prang´e et al. 1997, Fouchet et al. 1996) is ∼20 times larger than our favored globally averaged value based on the ISO observations, such enhancements, if localized to a few degrees in latitude, may be invisible in the disk-averaged ISO data. Conversely, “background” column densities of ∼1 × 1015 water molecules cm−2 may be undetectable in the HST spectra. 4.2. 5-Micrometer and Millimeter Observations of CO Carbon monoxide was detected in Saturn’s atmosphere by Noll et al. (1986). Through a continued analysis of a series of infrared observations, Noll and Larson (1990) have concluded

that the CO mixing ratio on Saturn is 1 × 10−9 if the CO is uniformly distributed through the troposphere and stratosphere (i.e., if the CO has an internal source, see Fegley and Lodders 1994) or 2.5 × 10−8 if the CO is concentrated in the stratosphere at pressures less than 80 mbar (indicating an external source). The observations thus do not allow for an unambiguous determination of the origin of the CO. We tested our CO profiles against the high-resolution (0.28 cm−1 ) spectrum recorded by Noll and Larson (1990) in the 4.5- to 5-µm region at the NASA Infrared Telescope Facility (IRTF). About 10 CO-lines from the (1–0) band can be clearly identified in this spectrum; the other lines are strongly contaminated by various molecular absorptions (see Fig. 9). Opacity in the 5-µm region derives from H2 –He collision-induced absorption and from molecular bands of NH3 , PH3 , CH3 D, GeH4 , H2 O, CO, and AsH3 , and Saturn’s spectrum here consists of thermal emission coming from deep hot layers (3–5 bars) and of sunlight reflected by cloud particles in the upper troposphere (∼400 mbar). Our thermal emission model includes an opaque cloud at 4.5 bar (B´ezard et al. 1989) and an attenuating cloud in the 400-mbar region. The solar reflected component is calculated via a Lambertian reflecting layer located at the 400-mbar pressure level (Noll and Larson 1990) and having a constant I /F reflectivity. The transmittance and reflectivity of this cloud were determined by fitting the flux levels in the Noll and Larson spectrum in the 2100–2125 cm−1 region where the reflected component is prominent and in micro-windows (e.g., 2086, 2130, 2135, 2142 cm−1 ) where thermal emission is significant. Our best fit model has a cloud transmittance of 0.26 and an I /F reflectivity of 0.30. Most of the CO lines occur in regions dominated by the solar reflected component. Line formation in that case depends only on the CO profile above the ∼400-mbar level. Sensitivity tests using either constant-with-height mixing ratios or photochemical profiles (Models A–D from Section 5.2 and Table V) indicate that the line intensities are mostly sensitive to the CO column abundance between 100 and 400 mbar. One exception, noted by Noll and Larson, is the P14 line at 2086.3 cm−1 , which primarily probes the CO abundance in the troposphere. Unfortunately, this line is entangled with phosphine absorption and is not well reproduced by the model spectra, which may be due to inaccurate TABLE V Magnitude of the Oxygen Influx to Saturn Model A B C D

O

H2 O

CO

CO2

CH3 OH

Total

8.0 × 105

7.5 × 104 1.0 × 105

7.0 × 104

2.0 × 106

1.5 × 106 1.5 × 106 1.5 × 106 5.0 × 105

2.5 × 106 1.7 × 106 5.4 × 106 5.2 × 106

3.9 × 106 2.7 × 106

Note. Fluxes in the first and last columns are given in terms of equivalent oxygen atoms cm−2 s−1 . The fluxes in the other columns are in molecules cm−2 s−1 . Model A is our nominal model. The final estimate from numerous models is a total oxygen influx (4 ± 2) × 106 O atoms cm−2 s−1 .

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FIG. 9. Saturn’s spectrum in the 4.7-µm region as observed by Noll and Larson (1990) at a resolution of 0.28 cm−1 (squares) compared with synthetic spectra including different CO profiles (solid lines). The data were digitized directly from Fig. 11 of Noll and Larson (1990); the digitization process introduces an estimated error of ∼ < 4%. Calculations (solid lines), showing increasing absorption features, include: (i) no CO (thinnest line), (ii) a CO profile with our nominal-model oxygen fluxes but with no internal source of CO, and (iii) our nominal CO profile (thickest line).

correction of telluric opacity or to the presence of an unknown absorber. From these tests, we estimate that column densities of (0.7–1.5) × 1017 cm−2 in the 100- to 400-mbar interval are consistent with the Noll and Larson observations. If CO were present only in the stratosphere—a test case considered by Noll and Larson—larger column densities of ∼6 × 1017 cm−2 in the 0- to 63-mbar interval would be needed to accommodate the observations. Our nominal CO profile, having a 100- to 400-mbar column abundance of 1.2 × 1017 cm−2 , satisfactorily reproduces the Noll and Larson observations (Fig. 9). A constant mixing ratio profile of (1–1.5) × 10−9 , a value consistent with the Noll and Larson analysis, fits these observations equally well. For Models A–D in Table V, we have assumed that an internal source of CO exists. Although most of the CO in the stratosphere of our model derives from an external source, a tropospheric (internal) CO source is certainly consistent with all the available observations. Figure 10 shows the results of two models that do not have an internal CO source. In the first model (dashed line), we simply changed the lower boundary condition for CO in our nominal model (Model A) from a fixed mixing ratio to a zero concentration gradient (as with the other photochemically

FIG. 10. Sensitivity of the CO mixing ratio to assumptions about the presence of an internal source of CO. The solid curve represents our nominal model (with an internal CO source), while the dashed curve represents a model with oxygen influxes assumed to be the same as our nominal model (Model A) but with no assumed internal source of CO, and the dotted curve represents a model with oxygen influxes as in Model C but with no internal CO. The triangle with associated error bars indicates the CO mixing ratio derived from the IRTF observations of Noll and Larson (1990) (see Table IV) and Section 4.2).

EFFECTS OF EXTRAPLANETARY OXYGEN ON SATURN PHOTOCHEMISTRY

produced species); this latter boundary condition allows the CO vapor that was produced in the stratosphere to diffuse downward through the lower boundary at a maximum possible rate, and no internal source of CO exists. All other model parameters are kept the same as our nominal model. The stratospheric column abundance of the CO in this model (7.8 × 1016 cm−2 above 63 mbar) is not sufficient to explain the Noll and Larson (1990) observations, and neither is the 100- to 400-mbar abundance (3 × 1016 cm−2 ) (see Fig. 9). Therefore, an internal source of CO is required if the external oxygen fluxes from our nominal model (Model A from Table V) are realistic. A similar result was obtained for the oxygen influxes from Model B. The second model with no internal CO source (dotted line in Fig. 10) has oxygen fluxes identical to Model C in Table V (a greatly enhanced influx of CO), but no internal CO source. The resulting species profiles from this model reproduce the ISO observations of H2 O, CO2 , and the hydrocarbons; the CO abundance in the 100- to 400-mbar region (8 × 1016 cm−2 ) is also large enough to yield CO absorption lines consistent with the infrared observations of Noll and Larson (1990). Therefore, if external oxygen influxes similar to Model C best represent reality, no internal CO source on Saturn is required to explain the Noll and Larson (1990) observations. A similar result was obtained for Model D with no internal CO source. Our conclusion is that the observed CO must have an internal source if the oxygen influxes typical of Models A and B are responsible for bringing oxygen to Saturn, but no internal CO source is required if oxygen influxes typical of Models C and D are responsible. We should also note that the tropospheric eddy diffusion coefficient on Saturn is very uncertain, and changes in our adopted value could alter these conclusions. Microwave observations also do not help resolve the internal vs external CO debate. Rosenqvist et al. (1992) unsuccessfully searched for CO on Saturn from millimeter-wave observations of the CO(2–1) line at 230 GHz. Reanalyzing this observation with our adopted thermal profile and assuming that CO is restricted to the stratosphere (P < 63 mbar), we find that the 2-σ upper limit for stratospheric CO (i.e., the value of the mixing ratio that would produce an absorption line contrast of 0.4%) is 5 × 10−8 . For this uniform-with-altitude model, most of the CO column density (1.0 × 1018 cm−2 ) occurs in the lower stratosphere. In contrast, the physically based models of this paper exhibit a strong decrease in the CO mixing ratio at pressures >0.1 mbar, resulting in a lower total stratospheric column density. With or without an internal source, Models A–D are consistent with the CO upper limit provided by the microwave observations. 5. SENSITIVITY STUDIES

5.1. Sensitivity to the Altitude of the Oxygen Influx Because the chemistry of the ablated oxygen compounds might depend on the presence of hydrocarbons in the ablation region, we now examine the sensitivity of our photochemical model results to the assumed ablation altitude. In one extreme

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case, we assume that the external oxygen is introduced at very high altitudes in the atmosphere so that the oxygen flows in through the upper boundary of our model, as would happen if the volatile oxygen compounds derive from vaporization of Saturn’s rings or if very tiny micrometeoroid grains (mass less than 10−11 g) make up most of the mass flux of micrometeoroids at 9.5 AU. In another extreme case, we assume that all the ablated oxygen is introduced at altitudes below the methane homopause (∼1.5 × 10−5 mbar), where methane and other hydrocarbons are present. In this model the ablation occurs in the 2 × 10−3 - to 3 × 10−5 -mbar region, as would happen if most of the ablated oxygen compounds derive from dust grains with masses greater than 1 × 10−6 g. Although changes in the assumed altitude of the oxygen influx do affect the abundances of oxygen compounds at very high altitudes in Saturn’s atmosphere, the differences are limited to pressures less than ∼10−3 mbar, and the total stratospheric column abundances of the major oxygen compounds are unaffected by the altitude changes. For example, the stratospheric column abundances of H2 O, CO, CO2 , CH3 OH, H2 CO, and CH3 CHO differ by less than 5% among the three cases described above. The primary reason for this lack of sensitivity to the ablation altitude is that diffusion time scales are shorter than the chemical lifetimes of the ablated oxygen species in Saturn’s upper stratosphere. H2 O photolysis, which is primarily responsible for initiating the oxygen photochemistry, peaks in the middle-lower stratosphere (at ∼0.3 mbar), and most of the three-body reactions that synthesize carbon-oxygen bonds (e.g., R321, R316, R324) do not begin to be important until the atmospheric pressure exceeds 10−4 mbar. The lack of sensitivity of the lower-stratospheric CO2 and H2 O abundances to the ablation altitude prevents us from using the ISO data to differentiate a possible ring source from an IDP source as Feuchtgruber et al. (1997) hoped would be possible. Neither the physical state of the incoming material (gas vs solid) nor the altitude of deposition can be inferred from the ISO data. Furthermore, the globally averaged H2 O influx derived from our analysis is insufficient for reactions with H2 O to be the dominant mechanism for the removal of H+ ions in Saturn’s ionosphere (Bass and Moses 1998), so it would be difficult to use the Voyager radio occulation electron-density profiles to distiguish between a possible ring and micrometeoric source of the H2 O. 5.2. Sensitivity to the Molecular Form of the Incoming Oxygen The ISO observations cannot be explained by the influx of any single oxygen compound. For instance, although H2 O is photolyzed to produce OH, and the OH can react with an internal source of CO (with a lower boundary mixing ratio of 1 × 10−9 in our nominal model and all other models presented in this subsection) to form CO2 (e.g., see the Titan models of Samuelson et al. 1983, Yung et al. 1984, Toublanc et al. 1995, and Lara et al. 1996), the flux of H2 O required to reproduce the ISO water observations is insufficient to produce the observed amount of CO2 . Similarly, although CO2 can be photochemically converted

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FIG. 11. The results from our nominal model (solid curves) compared with a model that assumes an external oxygen influx from H2 O molecules alone (dotted curves) and with a model that assumes an external oxygen influx from CO2 alone (dashed curves). The dash-triple-dot curves show the saturation mixing ratios for the different species. The data points with associated error bars indicate the H2 O and CO2 mixing ratios derived from ISO observations (this work; Table IV) and the CO mixing ratio derived from ground based infrared observations of Noll and Larson (1990) (see Table IV and Section 4.2).

to H2 O through reaction scheme (11) and other similar schemes (see Appendix B), the CO2 flux required to reproduce the ISO CO2 observations does not allow enough H2 O to be created to explain the ISO water observations. Figure 11 demonstrates this point. Figure 11 compares the results of our nominal model with two models that presuppose that the external supply of oxygen arrives in the form of H2 O alone or CO2 alone. For simplicity, both models consider the H2 O or CO2 to flow in from the top boundary of the model. For the model in which water is the sole source of the external oxygen, an influx of 1.5 × 106 H2 O molecules cm−2 s−1 reproduces the ISO water observations quite well; however, the resulting photochemically produced CO2 stratospheric column abundance is then a factor of 7 too small to explain the observed CO2 emission from ISO. Feuchtgruber et al. (1997) derived a much higher factor of ∼40 deficit in the CO2 column abundance for a pure H2 O flux model because they neglected several photochemical schemes that lead to the production of CO and eventually CO2 in the stratosphere

and because they did not include an external source of CO. For the model in which carbon dioxide is the sole source of external oxygen, an influx of 1.2 × 105 CO2 molecules cm−2 s−1 produces results consistent with the ISO CO2 observations, but the resulting H2 O stratospheric column abundance is a factor of ∼40 too small to explain the H2 O observations. One conclusion we can make from these and other modeling efforts is that the exogenic oxygen must be introduced in the form of both H2 O (or OH, O(1 D), O+ , OH+ , or H3 O+ , all of which would rapidly form H2 O on Saturn) and a species with a carbon–oxygen bond such as CO, CO2 , or HCO+ . None of these species alone can explain the ISO observations of CO2 and H2 O. If the external source introduces oxygen in the form of CO2 and H2 O only, then the ISO observations can be explained by an influx of 1.5 × 106 H2 O molecules cm−2 s−1 and 1.0 × 105 CO2 molecules cm−2 s−1 . These fluxes compare well with the earlier range of values derived by Feuchtgruber et al. (1997). The corresponding CO2 /H2 O ratio inferred for the exogenic oxygen source is 7% by volume or 16% by mass. This fraction is

EFFECTS OF EXTRAPLANETARY OXYGEN ON SATURN PHOTOCHEMISTRY

slightly larger than the estimated CO2 /H2 O ratio in cometary nuclei, 7–15% by mass according to Crovisier (1993) or 10– 12% by mass according to Greenberg (1998). However, the organic material in cometary grains is likely to release CO when heated (Greenberg 1998), and much of the carbon–oxygen bonded species introduced from the ablation of IDPs could derive from an organic component rather than directly from the icy component. In that case, the final CO2 /H2 O ratio could be higher than one would expect from the composition of the cometary ices alone. Alternatively, some H2 O could be converted to CO as the ablated vapors impact with each other and with atmospheric molecules. For our nominal model, we have assumed that the CO2 /H2 O and CH3 OH/H2 O flux ratios are consistent with the estimated abundances of ices in cometary nuclei (Greenberg 1998). We then allow the CO flux to increase until the ISO observations of CO2 and H2 O are both well reproduced. The resulting CO flux is ∼10 times higher than one would expect on the basis of an icy cometary component alone. It is not clear whether such a high CO/H2 O ratio is reasonable. A lot will depend on the presence of an organic component in the grains and on the chemistry that occurs within the meteor trail or during collisions with atmospheric molecules following ablation. Further modeling or experiments of the ablation of comet-like material is warranted. The inference from the above models is that the material that supplies Saturn with oxygen has an important component from carbon-oxygen bonded molecules. This result is consistent with comets and may be consistent with IDPs derived from comets. It is not clear that the result is consistent with interstellar dust (which may have been devolatilized during its history) or with Saturn’s rings (for which CO or CO2 ices have not been identified). Our nominal model, of course, does not represent a unique fit to the observations. Table V lists the total oxygen fluxes required for different assumptions about the composition of the incoming vapor. All the models in Table V provide a good fit to the ISO observations of CO2 and H2 O (and of the hydrocarbons, Moses et al. 2000). Our final estimate from numerous such studies is that the total oxygen influx to Saturn is (4 ± 2) × 106 O atoms cm−2 s−1 , and that some of the incoming material must come from a carbon–oxygen bonded species. 5.3. Sensitivity to Aerosol Properties The model is potentially sensitive to the adopted expression for the vapor pressure of water ice. We find that the vaporpressure expression of Washburn (1924) (as presented in Atreya 1986) leads to a column abundance of H2 O that is within 5% of the case using the vapor-pressure expression of Marti and Mauersberger (1993). However, water-ice vapor pressures are not well known at the very low temperatures relevant to Saturn’s stratosphere, and this parameter represents an unknown source of modeling uncertainty. The condensation scheme using a fixed particle radius (see Appendix A) leads to results that are mildly sensitive to particle

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size. We examined the results for particle radii of 0.1, 0.15, and 0.2 µm. These particle sizes were chosen to roughly bracket the range of observed mean radii of the stratospheric hazes on Saturn (West et al. 1983, Karkoschka and Tomasko 1993). All three of these models reproduce the ISO CO2 and H2 O observations to within observational uncertainties. The results are also only mildly sensitive to the preexisting aerosol concentration. The results for fixed-radius aerosol models compare well with models in which the aerosol radius is allowed to vary with condensation and evaporation (see Appendix A). In all, the model results are relatively insensitive to aerosol properties compared with the sometimes dramatic effects that result from changes in the composition or magnitude of the incoming oxygen. 5.4. Sensitivity to Temperature Because condensation affects the loss of water vapor from Saturn’s stratosphere, the temperature in the condensation region can play a role in the retrieved column abundance of H2 O. A 5-K change in atmospheric temperature at all altitudes in our model leads to a moderate change in the abundance of water vapor. An increase in temperature causes the H2 O to condense at a lower altitude; the column abundance of the vapor is correspondingly increased. However, the OH abundance is affected in the opposite way. An increase in temperature increases the loss of OH via the large exponential temperature dependence in the reaction R314 (OH + H2 → H2 O + H). The reduction in the OH abundance inhibits the reaction of OH with hydrocarbons and so inhibits the production of carbon–oxygen bonded species. Increasing the atmospheric temperature by 5 K results in a 65% increase in the stratospheric H2 O column abundance and a 4% decrease in the CO2 column. A 5-K decrease in atmospheric temperature leads to a 46% decrease in the H2 O column abundance and a &1% increase in the CO2 abundance. The observed water vapor abundance will thus be sensitive to latitudinal and seasonal temperature changes. 5.5. Sensitivity to Enhanced Water Influx Rates As we will discuss in Section 6 (below), an enhanced flux of oxygen-bearing ions or very tiny charged ice grains from Saturn’s rings may be entering Saturn’s atmosphere at latitudes where the magnetic field lines map back to certain instability regions within the ring plane (e.g., Connerney 1986, Northrop and Hill 1982, 1983). To simulate this effect, we now examine the consequences of an enhanced external water flux on the photochemistry of Saturn’s stratosphere (the influx of all other oxygen compounds is kept the same magnitude as in our nominal model). Figure 12 demonstrates the effects of increasing the incoming water flux at the top boundary of our model from the standard 1.5 × 106 molecules cm−2 s−1 in to an enhanced 107 or 108 molecules cm−2 s−1 for models in which the particle size is allowed to vary (see Appendix A). In all three models, the preexisting condensation nuclei are assumed to have a fixed radius of 0.15 µm. We find that an increase in the H2 O flux leads to an increase in the particle size, which in turn leads to

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FIG. 12. Sensitivity of the H2 O, CO, CO2 , CH3 OH, H2 CO, CH3 CHO, C2 H6 , C2 H2 , and C4 H2 mixing ratios to changes in the H2 O influx for a variableparticle-radius model. The solid curves represent a model in which the H2 O influx is 1.5 × 106 cm−2 s−1 (as with our nominal model), the dotted curves represent a model in which the H2 O influx is 107 cm−2 s−1 , and the dashed curves represent a model in which the H2 O influx is 108 cm−2 s−1 . Note that water is supersaturated in the middle stratosphere in the latter two cases. Other symbols in this plot are explained in the caption to Fig. 11. The data points for C2 H6 , C2 H2 , and C4 H2 are from ISO measurements (see Moses et al. 2000).

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increased condensation rates. The increased condensation rates help reduce the vapor density so that an increase in H2 O flux does not lead to an exact linear increase in the H2 O abundance. The increase from 1.5 × 106 to 107 cm−2 s−1 results in a factor of 6 increase in the stratospheric H2 O column abundance. The CO stratospheric column abundance is increased by only a factor of ∼3, while CO2 , which is formed from both CO and H2 O (Scheme 9 in Appendix B), is increased by a factor of 7. Although Fig. 12 demonstrates that the H2 O profile still follows the saturation vapor mixing-ratio curve in the lower stratosphere for the increased H2 O flux, the H2 O saturation ratio is higher than that of the lower-influx model at all levels of the condensation region. The increased abundance of H2 O in the middle and lower stratosphere provides a photolysis source of atomic hydrogen; atomic H in turn can affect the abundances of the observable hydrocarbon molecules. For instance, the model with a water influx of 1 × 107 cm−2 s−1 exhibits a C2 H2 stratospheric column abundance that is decreased 12% from our nominalinflux model. The abundance of C2 H6 is less affected while the abundances of complex hydrocarbons are even more affected. The enhanced water influx from 1.5 × 106 to 1 × 107 cm−2 s−1 leads to a 14% decrease in the CH3 C2 H column abundance or a 29% decrease in the C4 H2 column abundance. The increase in H2 O flux from 1.5 × 106 to 108 cm−2 s−1 results in even more dramatic changes. The stratospheric column abundances of H2 O, CO, and CO2 increase by factors of 25, 8.3, and 74, respectively. The column abundances of C2 H2 , CH3 C2 H, and C4 H2 are reduced by factors of 2.2, 3.2, and 8.9, respectively. If some latitude regions on Saturn are experiencing large water influx rates, changes in H2 O, CO2 , and hydrocarbon abundances at certain latitudes might be observable with either the Composite Infrared Spectrometer (CIRS) or the Ultraviolet Imaging Spectrograph (UVIS) aboard Cassini. Water vapor is greatly supersaturated in the upper stratosphere for the case in which the H2 O flux is 1 × 108 cm−2 s−1 . Despite the allowed increase in particle size, the deposition and diffusion rates in these regions are much greater than the condensation rate, and condensation cannot remove the vapor quickly enough to keep the vapor from becoming supersaturated. The maximum particle sizes are 0.16, 0.24, and 0.75 µm for the models with H2 O influx rates of 1.5 × 106 , 107 , and 108 cm−2 s−1 . The corresponding sedimentation rates at the tropopause are 11, 6.8, and 1.8 years. Therefore, consistent with the hypothesis of Connerney (1986), a large H2 O vapor influx at any particular latitude could lead to a preferential clearing of stratospheric aerosols at that latitude.

affected by changes in the slope of the eddy diffusion coefficient profile. In Fig. 13, we show the results from three models that have different K -profile slopes. The stratospheric eddy diffusion coefficient is assumed to obey the relation µ K (n) = 1.838 × 10

7

The eddy diffusion coefficient profile is one of the free parameters of our photochemical model. We use Voyager UVS occultation data (Smith et al. 1983) to constrain the eddy diffusion coefficient K in the upper stratosphere and use ISO observations of C2 H6 to constrain K in the lower stratosphere (Moses et al. 2000). We now consider how the major oxygen molecules are

¶β

,

where n is the atmospheric density in cm−3 , and the slope β equals 0.3, 0.5, or 0.7. The stratospheric eddy diffusion coefficient is also constrained to be ≥7 × 102 cm2 s−1 in these models. The increase in slope from β = 0.3 to β = 0.7 leads to a factor of 6 increase in the stratospheric column abundance of H2 O and a factor of 37 increase in the stratospheric column abundance of CO2 . None of these K profiles provides a good fit to the ISO CO2 , H2 O, or C2 H6 observations (e.g., Section 4; see also Moses et al. 2000); however, the figure demonstrates that CO2 , like C2 H6 , is very sensitive to the adopted diffusion coefficients in the lower stratosphere. The stratospheric abundances of H2 O, CO2 , and the hydrocarbons are much less affected by changes in the eddy diffusion coefficient in the upper stratosphere. The results of two models that both fit the available ISO data are shown in Fig. 14. The two models have a similar K profile in the lower stratosphere and similar assumed chemistry and oxygen influxes; however, the assumed eddy diffusion coefficients in the upper atmosphere differ greatly between the two models. The model with our nominal K profile has a much greater K h , the eddy diffusion coefficient at the methane homopause. The fact that both models are able to reproduce the ISO observations demonstrates that the midinfrared ISO observations of complex hydrocarbons or oxygen compounds (with the possible exception of CH3 observations, B´ezard et al. 1998) cannot alone be used to constrain K h . 6. DISCUSSION: ORIGIN OF EXTERNAL OXYGEN

6.1. Possible Exogenic Sources of Oxygen We now discuss the possible exogenic sources that could be supplying Saturn with oxygen, and we make some order-ofmagnitude estimates as to the potential oxygen influx rates from these different sources (Table VI). TABLE VI Estimates of the Oxygen Influx from Different External Sources Source

5.6. Sensitivity to Eddy Diffusion Coefficients

7.213 × 1011 n

Small comets Interstellar dust Interplanetary dust Ring vapor Solid ring particles

Oxygen flux (g cm−2 s−1 ) ∼8 × 10−19 2 × 10−21 1 × 10−16 7 × 10−19 Unknown

(H2 O), ∼7 × 10−18 (CO) to 4 × 10−17 to 2 × 10−15 to 1 × 10−17

Note. For the cometary and dust sources, 33% of the mass of the incoming material is assumed to be released as volatile oxygen.

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FIG. 13. Sensitivity of the CH4 , C2 H6 , H2 O, CO, CO2 , CH3 OH, H2 CO, and CH3 CHO mixing ratios to the slope of the eddy diffusion coefficient profile. The first panel shows the CH4 molecular diffusion coefficient profile (dash-triple-dot curve) along with K profiles that have slopes of β = 0.3 (dashed curve), β = 0.5 (solid curve), and β = 0.7 (dotted curve). The mixing ratio profiles in the subsequent panels correspond to these three cases. The data points in the CH4 panel corresponds to Voyager UVS and IRIS measurements (Smith et al. 1983, Festou and Atreya 1982, Courtin et al. 1984). The data point for C2 H6 is from ISO measurements (see Moses et al. 2000). Other symbols are explained in the captions to Figs. 10 and 12. Note that none of these models provides a good fit to the CO2 and C2 H6 abundances derived from ISO observations (Moses et al. 2000, de Graauw et al. 1997, this work).

EFFECTS OF EXTRAPLANETARY OXYGEN ON SATURN PHOTOCHEMISTRY

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FIG. 14. Sensitivity of the CH4 , C2 H6 , H2 O, CO, CO2 , CH3 OH, H2 CO, and CH3 CHO mixing ratios to the upper-atmospheric eddy diffusion coefficient profile. The first panel shows the CH4 molecular diffusion coefficient profile (dash-triple-dot curve) along with K profiles from our nominal model (solid curve) and a model that has a much lower eddy diffusion coefficient at the methane homopause (dotted curve). The mixing ratio profiles in the subsequent panels correspond to these two cases. Other symbols are explained in the captions for Figs. 10, 11, and 13.

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6.1.1. Large comets. As pointed out by Feuchtgruber et al. (1997), large comets are not likely culprits for delivering the observed amount of H2 O and CO2 to Saturn. Although the collision of Shoemaker–Levy 9 with Jupiter demonstrates that large comets do impact the outer planets, the frequency of such events is low compared with the diffusion rate of comet-derived material through the planet’s upper atmosphere. Estimates of the interval between impacts of 1-km diameter comets on Jupiter range from ∼90 years (Zahnle et al. 1998) to ∼240 years (Kary and Dones 1996). Simulations of Kuiper Belt objects evolving into ecliptic (short-period) comets by Levison and Duncan (1997) indicate that the interval between large cometary impacts on Saturn would be a factor of ∼3.2 larger than that for Jupiter, or ∼290–770 years between 1-km cometary impacts on Saturn. Compare these numbers with the ∼14 years it would take a water molecule to diffuse from 10−1 mbar (where most of the cometary material was deposited in the SL9 plume splashback according to the observations of Lellouch et al. (1997); see also the splashback simulations of Zahnle (1996)) to 1 mbar (roughly where the water would condense in Saturn’s stratosphere). The probability that the oxygen compounds that ISO observed on Saturn were supplied by a large cometary impact within the past ∼14 years is small (but not negligible). Furthermore, as discussed by Feuchtgruber et al. (1997), a dominant portion of the (presumably) cometary oxygen deposited in Jupiter’s upper atmosphere after the Shoemaker–Levy 9 impacts ended up (after thermochemical processing in the initial entry and plume-splashback shocks) as photochemically stable CO molecules rather than as more photochemically active H2 O. Given the plume-splachback thermochemical models of Zahnle (1996) and the observed abundances of H2 O and CO during the SL9 impacts (Lellouch 1996), a reasonable assumption for the final composition of the oxygen vapor after an SL9type impact might be 90% CO and 10% H2 O, although these fractions are uncertain. The total mass of water currently on Saturn as implied by the globally averaged H2 O column abundance determined by ISO observations (see Section 4) is 2 × 1013 g. If this water were derived from the impact of one large comet in which 10% of the oxygen ended up as H2 O, then the corresponding cometary diameter would be 0.9 km. Although observers could have easily missed detecting a 0.9-km-diameter comet at 9.5 AU if it were not actively producing dust, the effects from the impact of such a comet on Saturn should have been apparent. As an aside, we should mention that a rare large atmospheric storm appeared in the equatorial regions of Saturn in September 1990. The cloud-top heights of the visible tropospheric clouds appear to have risen from ∼1.4 bar for the undisturbed case to 300–400 mbar during the early and mature phases of the storm (Acarreta and S´anchez-Lavega 1999). Convective penetration into the stratosphere may have occurred during the storm, resulting in the injection of NH3 and other condensible volatiles into the stratosphere. For water ice particles to have been injected into the stratosphere during this storm, the particles would have to have been carried over 200 km in altitude from the estimated

water cloud-top pressure of ∼5 bar (Prinn et al. 1984) to the tropopause (∼60 mbar). For comparison, the atmospheric scale height at the tropopause is 34 km. Although it is possible that some H2 O was injected into the stratosphere in this case, it is unlikely that the storm could have carried water-ice particles to the ∼1-mbar level (400 km from the 5-bar level) where they could have evaporated. For this reason and others related to diffusion and sedimentation time scales, the 1990 atmospheric outburst is not likely to be the source of the observed water vapor on Saturn. 6.1.2. Small comets. What about small comets, if any such comets exist? One can extrapolate the Zahnle et al. (1998) jovian impact rate of 1-km comets to smaller comets assuming that the cumulative number of comets with diameter greater than D follows a power-law distribution. According to Zahnle et al. (1998), the impact rate (year−1 ) on Jupiter from objects with diameter greater than D is µ N˙ (>D) = 0.011

D 1.0 km

¶−γ

,

(1)

where γ = 1.97 is recommended by Shoemaker and Wolfe (1982); note that γ is very uncertain, especially for small objects. The rate on Saturn is a factor of ∼3.2 smaller than Jupiter (Levison and Duncan 1997). It can then be shown that the total mass of comets with diameters between dmin and dmax that impact Saturn per year is µ ¶ Z dmax 0.011 π 15 ˙ D 2−γ d D, (10 ) M = ργ 6 3.2 dmin

(2)

where ρ is the comet bulk density in g cm−3 , and dmax and dmin are given in kilometers. For comets with sizes between dmin = 0 and dmax = 0.22 km (the size of a comet for which the cumulative impact interval on Saturn would equal the 14-year diffusion time scale) and for an assumed comet density of 0.5 g cm−3 , then Eq. (2) can be used to estimate that 4 × 1011 g of cometary material strikes Saturn every year from comets with diameters less than 0.22 km. If we assume that the comets have a volatile oxygen content that is 30% by mass, that 10% of the final oxygen vapor is in the form of H2 O (as opposed to CO), and that the cometary debris is spread over the entire planet before it can diffuse to its condensation region, then the supply of incoming water vapor from small comets would be ∼8 × 10−19 g cm−2 s−1 . This value is uncertain by more than an order of magnitude. 6.1.3. Micrometeoroids. Very small meteoroids are prevalent in interplanetary space (e.g., Gustafson and Hanner 1996), and these particles are continuously raining down into planetary atmospheres. Due to the decreasing importance of asteroidal debris fragments beyond the orbit of Jupiter, IDPs at Saturn’s distance from the Sun should derive almost entirely from comets (included in this category are long-period comets, Halley-type

EFFECTS OF EXTRAPLANETARY OXYGEN ON SATURN PHOTOCHEMISTRY

comets, ecliptic comets, Centaurs, and Kuiper-belt objects). Unless they are already in orbit about Saturn, such particles will enter Saturn’s atmosphere at speeds greater than or equal to the escape velocity from Saturn, ∼35 km s−1 . At these high velocities, small dust particles composed of either silicates or water ice (or a mixture of the two) will completely ablate in Saturn’s upper atmosphere (Moses 1998); the most refractory meteoric debris will recondense to form very small haze particles while the volatile material can remain as vapor and can participate in atmospheric photochemistry. The mass flux of IDPs at Saturn’s orbital distance from the Sun is not certain (see Cuzzi and Estrada 1998 for a recent discussion of this issue). Pioneers 10 and 11 were the only spacecraft to directly measure the dust flux past the orbit of Jupiter. The Pioneer 10 and 11 dust-detection data (Humes 1980) are best explained by a model in which the IDP spatial density remains constant between 1 and 18 AU but in which the IDP population is on randomly inclined, highly eccentric orbits in the outer Solar System. This orbital distribution differs from that in the inner Solar System where most of the IDPs are on lowinclination, prograde orbits of low eccentricity (e.g., Gr¨un et al. 1997). Gr¨un et al. (1994) have suggested that Pioneers 10 and 11 may have been recording impacts from interstellar grains rather than IDPs, despite the fact that the resulting implied interstellar dust flux would exceed the mass flux of interstellar gas flowing into the Solar System and despite the fact that the direction of the Pioneer dust is not consistent with the interstellar grain direction inferred from Ulysses and Galileo. However, the plasma-wave instrument aboard Voyagers 1 and 2 provides further evidence that interplanetary rather than interstellar dust dominates the dust flux in the outer Solar System (Gurnett et al. 1997). The recorded impact rate remained steady far out in the Solar System and then suddenly dropped off; no impacts were recorded past 51 AU for Voyager 1 or 33 AU for Voyager 2. If the impacts were caused by interstellar grains, then the impact rate would be expected to be constant throughout the outer Solar System. Dust detectors aboard Galileo and Ulysses (and eventually Cassini) will give us the best determination of the mass flux of IDPs in the outer Solar System. However, at this point in time, the data from the impact of large dust particles has not been fully analyzed. Since these large particles are responsible for the bulk of the mass, estimates of the dust mass flux in the outer Solar System must await this analysis. To get a crude estimate, we use the Ulysses data presented in Fig. 11 of Gr¨un et al. (1997). This figure demonstrates that the detector recorded impacts from two &10−6 g particles at times when the spacecraft was near its maximum ecliptic latitude. The origin of these particles (interstellar vs high-inclination IDPs) is not clear. The combined mass of these particles is ∼1.3 × 10−6 g, and they impacted a ∼1000 cm2 area over the course of ∼360 days (Fig. 11 of Gr¨un et al. 1997). The implied mass flux is ∼4 × 10−17 g cm−2 s−1 (with very large statistical error bars). Another way to estimate the mass flux of micrometeoroids at Saturn’s orbital distance from the Sun is to use the impact

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rate onto the Earth and extrapolate to 9.5 AU, after taking into account information from Pioneers 10 and 11. This method was used by Feuchtgruber et al. (1997). Analyses of lunar microcraters, meteor data, zodiacal light observations, spacecraft data, and LDEF satellite data indicate that the Earth is collecting interplanetary dust at a rate of ∼30 × 106 kg year−1 (cf. Gr¨un et al. 1985, Love and Brownlee 1993). If the gravitational focusing factor of the Earth is ∼1.5, then the IDP mass flux at 1 AU but away from the Earth’s gravitational influence is ∼1.2 × 10−16 g cm−2 s−1 . If the dust spatial density is constant between 1 and 10 AU (Humes 1980), and if the velocities scale as one over the square root of the heliocentric distance (which would be true for objects with similar orbital shapes but would not be true if dust in the inner Solar System is on low-eccentricity, low-inclination orbits while the dust in the outer Solar System is in random inclination, high-eccentricity orbits), then the unfocused IDP flux at Saturn’s distance from the Sun is roughly 4 × 10−17 g cm−2 s−1 , a result perhaps fortitiously identical to one obtained above from the Ulysses data. Interstellar dust has been identified by dust detectors aboard the Ulysses and Galileo spacecraft (Gr¨un et al. 1993, 1994, Baguhl et al. 1995). Although interstellar grains are thought to dominate the dust impact rate (i.e., number of impacts per day) measured by the Ulysses and Galileo dust detectors while both spacecraft were beyond 3 AU (Gr¨un et al. 1994, 1997, Baguhl et al. 1995, 1996), the interstellar particles detected by the spacecraft tend to be small, so that the corresponding mass flux of impacting grains is low. Specifically, the mass flux of these small interstellar particles (average mass ∼3 × 10−13 g) in the outer Solar System is estimated to be 5 × 10−21 g cm−2 s−1 according to Gr¨un et al. (1994) and Baguhl et al. (1995)—too small to allow interstellar dust to make a major contribution to the volatile oxygen input at Saturn. However, the potential contribution from larger interstellar grains, which could supply most of the mass flux despite their smaller numbers, is not clear from the data presented in Gr¨un et al. (1993, 1994, 1997) or Baguhl et al. (1995, 1996). Impacts by large particles are rare events, and no clear analysis of the largest grains is presented in the above papers. Conservation of energy and angular momentum ensure that dust particles (or larger objects) that approach Saturn from infinity will be accelerated and have their trajectories deflected toward the planet; this gravitational focusing effect causes the planet to have a larger apparent cross section to the incoming ¨ particles (Opik 1951). The flux of micrometeoroids entering Saturn’s atmosphere is therefore larger than the 9.5-AU micrometeoroid flux away from Saturn’s gravitational influence. If F∞ is the unfocused 9.5-AU flux, then the flux entering Saturn’s ¨ atmosphere at a distance Rp from the planet’s center is (Opik 1951) µ ¶ v 2 (Rp ) F(R p ) = F∞ 1 + esc 2 , v∞

(3)

where vesc is the escape velocity from the planet at radius Rp

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and v∞ is the relative velocity between the planet and the micrometeoroid at infinity. The escape velocity in Saturn’s upper atmosphere is ∼35 km s−1 , and the magnitude of the gravitational focusing effect depends largely on v∞ . Particles moving slowly with respect to Saturn will experience a greater focusing effect. To get a rough estimate of v∞ for low inclination, low eccentricity, prograde particles impacting the planets, one can multiply the particle eccentricity by the planet’s mean orbital velocity (∼9.64 km s−1 for Saturn). If we assume that F∞ is 4 × 10−17 g cm−2 s−1 (see above) and that the micrometeoroids contain 33% volatile oxygen by mass, then we can use Eq. (3) to estimate the flux of oxygen-bearing material to Saturn; the corresponding oxygen influx rates for interstellar and interplanetary dust are given in Table VI. 6.1.4. Ring and satellite debris. Planetary rings are constantly eroded by mechanisms such as sputtering by charged particles, sputtering by energetic photons, meteoroid bombardment, interparticle collisions, radiation pressure, and plasma drag (e.g., Durisen 1984, Cuzzi and Durisen 1990). Planetary satellites are also eroded by sputtering and meteoritic impacts. The debris liberated from the erosion process can range from vaporized atoms, ions, and molecules to large solid particles. Some of the debris will escape to space, some will reimpact the rings or satellites, and some will impact the planet’s atmosphere. Banaszkiewicz and Krivov (1997) have modeled the fate of dust particles ejected from Saturn’s satellite Hyperion. They find that Saturn is not the ultimate depository for these particles except for particle radii &1 µm (i.e., radiation pressure effects cause the orbits of small particles to intersect Saturn). Larger particles tend to remain in orbit about Saturn or impact Titan during the time range of the Banaszkiewicz and Krivov simulation (10–100 saturnian years). We estimate that if micrometeorite impacts are responsible for the ejection of dust from satellites, then the satellite source is unlikely to supply Saturn with as much material as the direct ablation of micrometeoroids because the satellites represent small targets compared with Saturn and because only a small fraction of the satellite impact debris mass will reach the planet. Saturn’s rings, however, represent a larger target to IDPs. To estimate the water influx to Saturn from the vaporization of ring particles following micrometeoritic bombardment, we determine the ratio η of the direct ablation source to the indirect ring-vaporization source, Ã η=

2 + v∞ 2 + v∞

GM Rring 2 vesc

!µ

Aring Aatm

¶ f vap f imp ,

(4)

where v∞ is the velocity at infinity of the IDPs relative to Saturn, G is the universal gravitational constant, M is the mass of Saturn. Rring is an average radial distance of the rings from the center of Saturn, Vesc is the escape velocity from Saturn’s upper atmosphere, Aring is the surface area of the rings, Aatm is the surface area of the atmosphere, f vap is the mass fraction of

vapor released after an impact in terms of the mass of an incoming IDP, and f imp is the fraction of vaporized ring material that impacts Saturn’s atmosphere. In the above equation, we assume that the IDPs and ring particles have the same composition. The first term in parentheses on the right-hand side represents the ratio of the focusing factor for the two-dimensional ring system at distance Rring to the focusing factor for the planet in the upper atmosphere. (Note that Cuzzi and Estrada (1998) argue that the focusing factor of isotropic particles onto a flat disk or ring system may not be as simple as the numerator in this term indicates). We can then use estimates given in Morfill et al. (1983) for this same problem (i.e., f vap = 2, f imp = 0.1, Aring /Aatm = 0.14, and G M/Rring = 3.2 × 1012 cm2 s−2 ) to arrive at the conclusion that the indirect ring vapor source would supply 100 to 130 times less oxygen to Saturn than the direct IDP ablation source (or 70 times less if interstellar grains are causing the impacts). Solid debris kicked up from IDP impacts on the rings could also find its way into Saturn’s atmosphere. The flux from this source is difficult to estimate. The problem can be set up similar to Eq. (4), but f vap is replaced by the mass fraction of solid ejecta released after an impact, which can be as large as 104 for rocky or metallic projectiles hitting an icy target (Lange and Ahrens 1987, see also Morfill et al. 1983) but may be be smaller for an icy particle hitting an icy target. Most of the ejecta will reimpact the rings (Durisen et al. 1989), but some uncertain fraction may enter Saturn’s atmosphere on fairly short time scales. If that encounter fraction is *0.2%, then the ring impact source could supply a similar dust influx to Saturn as the direct IDP source. If very small particulate ejecta (or vapor) become ionized, the chances of impacting Saturn increase. Theoretical studies of the stability of ions or very small charged grains in Saturn’s ring plane (Ip 1983, Northrop and Hill 1982, 1983) predict that such particles located within a certain stability distance from Saturn can spiral along magnetic field lines to directly impact Saturn’s atmosphere. Enhanced erosion of the rings (and enhanced influx into the atmosphere) would be expected within these stability limits. Because Saturn’s magnetic field is aligned with the planet’s rotation axis, different latitude regions on Saturn can be mapped back to corresponding regions in the ring plane (e.g., Northrop and Hill 1982, 1983, Connerney 1986). As Connerney (1986) suggests, certain dark latitude bands on Saturn may correspond to regions of instability within the ring system; those planetary latitudes may be experiencing an enhanced water influx, helping to clear stratospheric aerosols from the atmosphere. From an analysis of the effects of water vapor on ionospheric electron-density profiles, Connerney and Waite (1984) estimate that the water influx at the inner edge of Saturn’s B ring could be as high as 2 × 109 molecules cm−2 s−1 , corresponding to a mass influx of 6 × 10−14 g cm−2 s−1 . An enhanced influx of oxygen would be found at very localized latitudes across the planet. 6.1.5. Conclusions regarding exogenic sources of oxygen. The conclusions from the rough estimates given above is that micrometeoroid ablation could be the dominant globally averaged

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mechanism for delivering oxygen to Saturn (see Table VI). The relative importance of interstellar dust versus interplanetary dust is not clear; a resolution of this issue will have to await the analysis of impacts from the most massive dust particles (masses greater than 10−8 g) detected by the Ulysses, Galileo, and eventually Cassini dust detectors. Cometary impacts appear to be too infrequent, their deposits too short-lived, and the resulting CO/H2 O ratio too large to explain the observations, although the mass distribution of cometary nuclei is too poorly known to completely rule out small comets as a major contributor. Satellitederived material and ring vapor will be a less important source of oxygen than direct IDP ablation. However, solid debris ejected from impacts of IDPs on Saturn’s rings may be as important or perhaps even more important than the direct ablation of IDPs. Further modeling of this process is warranted. Due to a ring– magnetosphere–atmosphere connection, certain latitude regions on Saturn could be experiencing an anomalously high oxygen influx rate.

TABLE VIII Approximate Focusing Factors for the Outer Planets

Planet

Escape velocity from the upper atmosphere vesc (km s−1 )

Jupiter Saturn Uranus Neptune

60 35 21 23

IDPs: random i, high e focusing factor 12 7.7 5.9 10

IDPs: low i, low e focusing factor

Interstellar grains focusing factor

200 130 90 170

6.3 2.8 1.7 1.8

2 /v 2 , where v Note. The gravitational focusing factor is 1 + vesc esc is the escape ∞ velocity from the planet’s upper atmosphere and v∞ is the relative velocity between the dust particle and the planet at infinity. Rough estimates for the values of v∞ for the low e, low i IDPs were obtained by multiplying the eccentricity of the dust particles by the planet’s mean orbital velocity. Isotropic, highly eccentric IDPs were assumed to have v∞ equal to the escape velocity from the Solar System at Saturn’s orbit. The interstellar grains are assumed to have v∞ = 26 km s−1 . All values are approximate: two significant digits are included for comparison purposes only.

6.2. Implications Regarding a Micrometeoroid Source Because of the potential importance of small dust grains in supplying Saturn with oxygen (Table VI), we now focus exclusively on the micrometeoroid source and discuss some of the implications of our modeling for the amount and origin of the dust in the outer Solar System. In Section 5.2 and Table V, we demonstrated that different assumptions about the molecular form of the incoming oxygen can result in different oxygen influxes required to explain the CO2 and H2 O emission features observed by ISO. From numerous modeling attempts, we concluded that Saturn is currently experiencing an influx of exogenic oxygen equivalent to (4 ± 2) × 106 O atoms cm−2 s−1 or (1 ± 0.5) × 10−16 g cm−2 s−1 of oxygen. If volatile oxygen compounds comprise one-third to one-half of the exogenic source by mass, then Saturn is currently being bombarded with (3 ± 2) × 10−16 g cm−2 s−1 of extraplanetary material. If the oxygen influx to Saturn derives from a micrometeoroid source (either interstellar or interplanetary dust), then that material must have been gravitationally focused into Saturn’s atmosphere (see Section 6.1.3), and we can use the above estimates of the mass influx to Saturn to determine the unfocused micrometeoroid dust flux at 9.5 AU. Table VII presents the reTABLE VII Unfocused 9.5-AU Dust Flux Required to Explain ISO Observations Source

Flux (g cm−2 s−1 )

Interstellar grains IDPs: random i, high e IDPs: low i, low e

(1 ± 0.7) × 10−16 (4 ± 3) × 10−17 (2 ± 1.4) × 10−18

Note. The symbol i corresponds to inclination, e corresponds to eccentricity. The gravitational focusing factors were assumed to be 2.8 for interstellar grains, 7.7 for IDPs with random i and high e, and 130 for IDPs with low i and low e.

sulting unfocused fluxes, F∞ , for different assumptions about the orbital properties of the dust particles. Note that some overlap between the interstellar source and the randomly inclined, highly eccentric IDP source exists, and a continuum of IDP orbital populations could exist in between the two extreme IDP populations listed in the table. The estimates in Table VII can be used in conjunction with the Cassini dust detection data to help determine the source of the dust in the outer Solar System. If the Cassini dust detectors measure dust fluxes in excess of 4 × 10−18 g cm−2 s−1 , then low-inclination low-eccentricity dust cannot dominate the dust flux at 9.5 AU or Saturn would exhibit much larger H2 O and CO2 abundances than were seen by ISO. Similarly, if the dust flux measured by the Cassini dust detectors is less than 3 × 10−17 g cm−2 s−1 , then gravitational focusing of interstellar grains is insufficient to explain the observed amount of CO2 and H2 O on Saturn. Approximate gravitational focusing factors for all the outer planets are provided in Table VIII for different assumptions about the source of the dust in the outer Solar System. If the Humes (1980) Pioneer 10 and 11 analysis is correct in that the spatial density of IDPs is roughly constant from 1 to 18 AU, if that relation holds out to the orbit of Neptune, and if the dust has similar orbital properties throughout the outer Solar System, then F∞ will fall off as one over the square root of the heliocentric distance. Combining this information with the focusing factors listed in Table VIII, we can estimate that the flux of IDPs into the atmospheres of the outer planets will follow the relative relation 2 : 1 : 0.5 : 0.7 for Jupiter : Saturn : Uranus : Neptune. If the flow of interstellar grains into the heliosphere is uniform, then the relative influx rates for the outer planets will roughly follow the relation in the last column of Table VIII (although we should also note that gravitational focusing of interstellar dust by the Sun will increase the flux at Jupiter relative to planets farther from the Sun).

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TABLE IX Water Fluxes Inferred from ISO Observations Celestial body Jupiter Saturn Uranus Neptune Titan

Inferred H2 O flux (molecules cm−2 s−1 )

Reference

(0.6–1.8) × 106 (0.3–5) × 106 (0.5–2) × 106 (0.06–0.16) × 106 (0.12–15) × 106 (0.8–2.8) × 106

Feuchtgruber et al. 1999 Feuchtgruber et al. 1997 This work Feuchtgruber et al. 1997 Feuchtgruber et al. 1997 Coustenis et al. 1998

Note. Our preferred value for the H2 O influx rate at Saturn is 1.5 × 106 molecules cm−2 s−1 .

Table IX lists the external water influx rates inferred from the ISO observations of H2 O on the outer planets and Titan. Note that the inferred water fluxes are similar for Jupiter, Saturn, Neptune, and Titan (Feuchtgruber et al. 1997, 1999, Encrenaz 1998, Lellouch et al. 1998, Coustenis et al. 1998), while the inferred H2 O influx for Uranus is roughly a factor of 10 smaller. Part of the discrepancy with Uranus may be caused by uncertainties in the stratospheric temperature profile and eddy diffusion coefficients for this planet; the upper stratosphere of Uranus is colder than that of Neptune, Saturn, or Jupiter, causing water to condense out at higher altitudes and preventing large column densities of water vapor from accumulating. However, those issues were considered by Feuchtgruber et al. (1997) in determining the fluxes, so Uranus may truly be receiving a lower influx of oxygen-bearing material than the other outer planets. In fact, at this point, the uncertainties in the H2 O flux at Neptune are such that it could be consistent with the flux at Uranus. The similarity between the other outer planets and Titan bolsters the conclusion that micrometeoroids (either interplanetary or interstellar) are the dominant source of the oxygen influx on Saturn (see also Coustenis et al. 1998)—if ring erosion were the dominant supply mechanism, then one might expect Saturn to have much higher influx rates than Titan or than planets without such extensive ring systems. At the level of uncertainty of the H2 O fluxes listed in Table IX, the relative H2 O influx rates inferred for Jupiter, Saturn, and Neptune are consistent with either an IDP or interstellar source. However, the relative H2 O influx rates for Saturn and Titan are not consistent with an IDP population on low inclination, low ecentricity orbits—in that situation, one would expect Saturn to encounter a mass flux ∼16 times that of Titan on average. Note that the gravitational focusing by Saturn at Titan’s radial distance (focusing factor 1.1–7.9 depending on v∞ of the dust) is an important factor in determining focusing rates at Titan. In fact, the inferred water influx rate to Titan (0.8–2.8 × 106 cm−2 s−1 according to Coustenis et al. 1998) is only 0.5–2 times our preferred influx rate of ∼1.5 × 106 H2 O molecules cm−2 s−1 on Saturn. Such a high inferred relative rate for Titan is surprising and seems to also rule out an IDP population on randomly inclined, highly eccentric orbits as the

source of the external debris (where one would expect Saturn to encounter ∼6 times the mass influx as Titan) or an interstellar source (where Saturn would collect ∼2.6 times the mass influx as Titan). Perhaps the exterior satellites Hyperion, Iapetus, and Phoebe are providing a preferential local source to Titan (e.g., Banaszkiewicz and Krivov 1997, Coustenis et al. 1998). Further conclusions regarding the Titan–Saturn issue will have to await future modeling. Interstellar grains can be viable sources of oxygen to Saturn and the other outer planets only if large (10−8 –10−5 g) grains exist and if these grains contain volatile material. A survey of very faint meteor data recorded by the AMOR radar in New Zealand (Taylor et al. 1996) provides support for the idea that large interstellar grains are penetrating into our Solar System. Taylor et al. (1996) have analyzed data from meteors with measured speeds of at least 100 km s−1 ; the dust grains producing these meteors cannot be bound to the Solar System. At 100 km s−1 , the grains must have radii of at least 20 µm (mass 3 × 10−8 g) in order to be recorded by their detectors. Taylor et al. (1996) determine that there may be two distinct sources of the interstellar grains, one in the direction of the solar galactocentric orbit, and one perhaps in the direction of motion of the Solar System relative to nearby late A-type stars (the latter direction is also close to the direction in which interstellar helium is flowing into the Solar System (Witte et al. 1993)). Although no mass flux estimates are provided by Taylor et al. (1996), their analysis suggests that large interstellar grains do exist. On the other hand, interstellar grains might be expected to have lost most of their volatile ices through several shock episodes in their lifetimes and thus would not be good candidates for supplying the outer planets with CO2 and H2 O. Our required F∞ for randomly inclined, highly eccentric IDPs from Table VII is consistent with the estimate we made in Section 6.1.3 by extrapolating the dust flux at Earth to the flux at 9.5 AU using information obtained from Pioneers 10 and 11 (Humes 1980). Furthermore, the orbits of these assumed randomly inclined, highly eccentric particles fit the Pioneer data quite well. Where would randomly inclined, highly eccentric dust particles come from? Dust particles that begin in orbits of low inclination seldom evolve into orbits of high inclination without being ejected from the Solar System (Liou et al. 1996), and it would be even more difficult to produce retrograde orbits from initially prograde orbits (see the similar situation of the evolution of Kuiper belt objects into short-period comets by Levison and Duncan (1997) and Duncan et al. (1988)). A randomly inclined dust population would have to be produced from a randomly inclined parent source such as long-period comets or Halley-type comets. Despite their large dust production rates during perihelion passages, long-period comets (from the Oort-cloud) are unlikely to be important sources of IDPs in the Solar System. Most of the dust that is released from long-period comets will escape the Solar System in hyperbolic trajectories; the fraction that does not escape the Solar System will travel (at least initially) in

EFFECTS OF EXTRAPLANETARY OXYGEN ON SATURN PHOTOCHEMISTRY

orbits of periods ∼106 years, ensuring that the dust spends only a short time in the planetary region. Randomly inclined, highly eccentric IDPs observed today most likely derive from Halleytype comets that release dust during perihelion passages. Some fraction of the particles released from these comets will also have hyperbolic trajectories (due to solar radiation pressure and other forces); however, the fraction that remains bound to the Solar System can slowly spiral in through the outer Solar System to the inner Solar System, encountering the giant planets along the way (e.g., see the simulations of Liou et al. (1999)). In the past two centuries, 23 Halley-type comets have been identified (Marsden and Williams 1995); at least 13 of these comets have been observed on more than one apparition, and several (including Halley itself) are known to be active dust producers (see Hughes 1996). From the sparse data we have, the dust flux required to explain the ISO observations of CO2 and H2 O at Saturn appears to be consistent with randomly inclined, highly eccentric dust particles released from Halley-type comets. Further refinement of this claim must await data from the Cassini dust detectors. Our modeling indicates that the dust flux in the outer Solar System must be no more than (2 ± 1.4) × 10−18 g cm−2 s−1 if low inclination, low eccentricity IDPs are the dominant source of oxygen into Saturn. Although this small flux is the easiest to support from the standpoint of depleting potential reservoirs, IDP populations with these properties were found to be at most a minor component of the dust in the outer Solar System at Jupiter’s distance from the Sun (Humes 1980, Baguhl et al. 1985, 1986, Gr¨un et al. 1994). The lack of low-inclination low-eccentricity dust is surprising considering that dust particles evolving from the Kuiper Belt maintain low e and low i as they travel inward (Liou et al. 1996) and that mutual collisions among Kuiper-Belt objects (Stern 1995) or impacts of interstellar grains on Kuiper-Belt objects (Yamamoto and Mukai 1998) can readily produce dust. The existing spacecraft data seem to indicate that Kuiper-Belt dust is apparently not a major source of meteoroidal material at &10 AU. Interesting new models by Landgraf et al. (1999) suggest that Kuiper-Belt grains may dominate the dust flux at Uranus and Neptune, Halley-type comet dust might dominate at Jupiter, and a combination of the two could be important for Saturn. Future Cassini data could resolve this issue. 7. SUMMARY AND CONCLUSIONS

The ISO observations of CO2 and H2 O on Saturn (de Graauw et al. 1997, Feuchtgruber et al. 1997) imply that oxygen is currently being supplied to the planet from an external source. We use a one-dimensional photochemistry/diffusion model to study the influence of the extraplanetary oxygen on atmospheric chemistry. Photolysis of extraplanetary water initiates much of the oxygen photochemistry on Saturn. CH4 and other hydrocarbons help shield H2 O from ultraviolet photons; that shielding combined with the efficiency of water recycling allows observable

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amounts of water vapor to build up. Water is removed from the stratosphere through condensation as well as through photochemical loss mechanisms that lead to the formation of CO and CO2 . Water vapor will condense at ∼0.4 mbar in Saturn’s stratosphere, and water ice contributes ∼4% to the total column mass of the global stratospheric haze in our model. Carbon monoxide is a dominant end product of the oxygen photochemistry on Saturn (see Fig. 2). In the stratosphere, CO is synthesized primarily through reactions of OH with C2 H2 , C2 H4 , and CH3 ; the reaction products undergo photolysis or subsequent reactions to eventually produce CO. Once formed, CO is very stable. The only effective photochemical loss process in the stratosphere is the reaction of CO with OH to form CO2 and H at a slow but steady rate. The CO2 thus formed is destroyed by photolysis, resulting in the recycling of CO. Other molecules produced by oxygen photochemistry in moderate but currently unobservable quantities include CH3 OH, H2 CO, H2 CCO, and CH3 CHO. These species are important intermediaries in the conversion of H2 O to CO in Saturn’s stratosphere. The ISO observations indicate a CO2 column abundance of (6.3 ± 1) × 1014 cm−2 above 10 mbar (or a mixing ratio of (4 ± 0.6) × 10−10 at 1 mbar) and an H2 O column abundance of (1.4 ± 0.4) × 1015 cm−2 above 100 mbar (or a mixing ratio of (1 ± 0.3) × 10−8 at 0.1 mbar). These observations can be reproduced in our models by an influx of 1.5 × 106 H2 O molecules cm−2 s−1 and 1.0 × 105 CO2 molecules cm−2 s−1 (cf. Feuchtgruber et al. 1997). However, different assumptions about the molecular form of the incoming oxygen lead to different estimates of the total oxygen influx to Saturn (see Table V). From the results of several models, we conclude that Saturn is currently receiving the equivalent of (4 ± 2) × 106 O atoms cm−2 s−1 . This amount of exogenic oxygen has little effect on the background hydrocarbon photochemistry. We further determine that although CO2 can be produced from the reaction of OH (from water photolysis) with an internal saturnian source of CO, an influx of H2 O alone cannot account for the ISO observations. To reproduce the observed CO2 /H2 O ratio from ISO, some of the extraplanetary oxygen must arrive in the form of a carbon–oxygen bonded species such as CO, CO2 , HCO, or HCO+ . This conclusion is not unexpected. The composition of IDPs in the outer Solar System should mimic the parent comets, and carbon–oxygen bonded ices such as CO, CO2 , H2 CO, and CH3 OH are known to be present in comets. In addition, CO could be an important volatile vapor released from the organic component of IDPs during atmospheric ablation (cf. Greenberg 1998). Some combination of Models A and B from Table V may best represent the oxygen influx on Saturn. The possibility of an internal (tropospheric) source of CO cannot be ruled out from our modeling. In fact, the CO observations of Noll and Larson (1990) are easiest to reproduce if we allow the existence of an internal CO source, although models with high external oxygen fluxes and no internal CO are also consistent with all available observations. High-resolution spectroscopy of CO lines in the 4.7-µm band could help constrain

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the origin of this molecule. A spectral resolution higher than ∼0.05 cm−1 would allow us to recover the line profiles and to separate the pressure-broadened line wings due to tropospheric CO from a putative unresolved line core originating from enhanced stratospheric CO. Such observations would, however, require a sensitive instrument and long integration times to ensure sufficient signal-to-noise ratios. If volatile oxygen compounds make up one-third to onehalf of the extraplanetary material entering Saturn’s atmosphere, then our modeling indicates that Saturn is currently collecting a global average of (3 ± 2) × 10−16 g cm−2 s−1 of exogenic material (for comparison, the Earth receives (2 ± 1) × 10−16 g cm−2 s−1 of micrometeoroidal material according to Love and Brownlee (1993) and Gr¨un et al. (1985)). The implied mass flux to Saturn is greater than our estimates of the possible contribution from small comets or ring-derived vapor but seems consistent with a micrometeoroid ablation source. Similarities between the inferred water influx rates for Jupiter, Saturn, and Titan provide further support for micrometeoroids being the responsible party. The micrometeoroid population in the outer Solar System consists of cometary dust and/or interstellar grains; the relative importance of two sources is currently uncertain, although given the possible depletion of volatile material in interstellar grains and the lack of evidence for low-inclination low-eccentricity IDPs at Jupiter’s distance from the Sun, we favor Halley-type comets as the origin of the grains at Jupiter and Saturn. The dust grains in this case are most likely released from the comets during perihelion passages. Because of gravitational focusing, the amount of dust entering Saturn’s atmosphere is greater than the nearby flux away from Saturn’s gravitational influence. For different assumptions about the orbital properties of dust in the outer Solar System, we can use our determination of the current mass flux of material to Saturn to estimate the unfocused dust flux at Saturn’s distance from the Sun (see Table VII). These estimates can be used in conjunction with future data from the Cassini dust detectors to help determine the primary source of debris in the outer Solar System.

stable, however, and they use realistic boundary conditions for the condensing constituents at the lower boundary to avoid conservation-of-mass problems. Romani and Atreya (1989), on the other hand, use a more physically realistic description of the condensation process but use a less realistic lower boundary condition. Romani and Atreya’s technique can lead to problems with mass conservation, can become unstable, and is also sensitive to initial conditions (e.g., the assumed particle size). We have developed a technique that is fast yet more stable than either of the above models, but care must still be exercised to avoid problems with conservation of mass. In essence, we treat the condensed phase as a separate molecular species that is produced by condensation and lost by evaporation. We assume that condensation nuclei (CN) already exist in the atmosphere and material can condense on these CN. We also assume that the CN radius and the altitude profile of the CN concentration remain fixed with respect to time. Condensation and evaporation are then represented by the two equations that describe the production and loss of the condensed phase. In the case of water condensation in Saturn’s atmosphere, the equations become H2 O + dust = H2 O(s)

R421

H2 O(s) + V = H2 O

R422

where “dust” represents the pre-existing CN particles, “V” represents a dummy molecule whose number density is 1 cm−3 at all atmospheric levels, and the subscript (s) refers to the condensed phase. The boundary conditions for the condensed phase are the same as those for the vapor; that is, we assume zero flux at the top of the atmosphere and a constant concentration at the lower boundary (so that the species flows through the lower boundary at a maximum possible rate). Note that our lower boundary is well below the tropopause so that condensation is not important at the lower boundary, and the boundary condition correction suggested by Summers and Strobel (1989) is not necessary. For steady-state diffusion-limited condensation, the rates of the above reactions become [cf. Seinfeld 1986] k421 = 4πrp Dβ

Ã

k422 = 4πrp Dβ

where rp is the particle radius (i.e., the total radius of the CN plus the consensed material), rN is the radius of the CN, D is the diffusion coefficient of the condensible molecules through a predominantly H2 atmosphere (e.g., Reid et al. 1987 for H2 O), µ is the mean molecular mass of the condensible molecule, ρ is the bulk density of the condensed phase, and β is the correction factor added to account for the fact that the condensation can occur in the transition regime between free-molecular and continuum flow. The correction factor β can be determined through a flux-matching approach (e.g., Seinfeld 1986),

APPENDIX A: CONDENSATION The addition of condensation to photochemical models can be numerically challenging. Condensation and evaporation operate on time scales that are frequently much shorter than those for vapor transport or gas-phase chemistry; the continuity equations become “stiff,” and numerical instabilities can result. Several techniques have been developed to parameterize condensation in photochemical models (e.g., Yung et al. 1984, Summers and Strobel 1989, Romani and Atreya 1989). These parameterization schemes, which are based on standard theories for vapor condensation (e.g., Pruppacher and Klett 1980), operate only when the partial pressure of the condensible vapor exceeds the saturation vapor pressure at any level in the atmosphere. For stratospheric condensation, condensible vapor is supplied through photochemical reactions within the condensation region or through diffusion from higher altitude levels. When supply rates are much different from condensation rates, small time steps are required to achieve convergence in the models. Summers and Strobel (1989) alleviate this problem by using artificially long condensation time scales, and their results are often sensitive to the choice of condensation time scale. Their method is relatively

! 3µ ¡ ¢ n sat , 4πρ rp3 − rN3

β=

1 + KnD 1+

2KnD (1+KnD ) δ

,

where δ is the sticking coefficient (the probability that a condensible molecule will stick when it encounters the aerosol particle), KnD =

2D , rp hci

√ where hci = 8kT /π µ is the mean speed of the condensible molecules, k is the Boltzmann constant, and T is the temperature. We assume that δ = 1. In one adaptation of this condensation scheme, the particle radius is allowed to vary as the vapor condenses or evaporates. The particle radius is then µ rp =

3µ[H2 O(s) ] + rN3 4πρ[dust]

¶1 3

,

195

EFFECTS OF EXTRAPLANETARY OXYGEN ON SATURN PHOTOCHEMISTRY where [H2 O(s) ] and [dust] are respectively the concentration of ice in the condensed phase (variable) and the concentration of condensation nuclei (fixed). For a condensation scheme of increased stability, one can assume that rN = 0 and rp is fixed. This technique is used for our nominal model. The advantages of the second scheme are its stability and speed; the main disadvantage is that the results will depend on assumptions about the particle size, and one must exercise care to avoid problems with conservation of mass. Because the aerosol size is fixed, the condensed phase will never completely evaporate and can become a vapor source in the model. For our choice of aerosol size (0.15 µm) and concentration (for pressures P (in mbar) less than 20 mbar, [dust] = P cm−3 ; for P > 20 mbar, [dust] = 20 cm−3 ), evaporation is not a significant net source of water to the model, and the oxygen influx from meteoroid ablation is balanced to within 0.1% of the oxygen flow through the tropopause. Violations of mass conservation are therefore not a significant problem, but different choices could lead to larger errors, and our condensation scheme should be used with caution. To derive the above expression for the evaporation rate coefficient for water (k422 ), we have assumed that the condensible portion of the aerosols is composed of a single phase (pure water ice) and that the number density of water molecules in the condensed phase can be described as the number density of aerosols multiplied by the mass of the condensed phase divided by the mean molecular mass of a water molecule. Even in regions where the vapor is not supersaturated, molecules can condense on and evaporate from the preexisting nuclei. However, if the particle radius decreases such that it becomes less than or equal to the radius of the condensation nuclei, we assume that the evaporation rate constant k422 becomes zero and that rp = rN . The results for the scheme with variable particle sizes are not very sensitive to the assumed initial size of the CN. For instance, if the vapor becomes supersaturated, condensation is enhanced, and the particle size grows, allowing condensation to become even more efficient, helping to reduce vapor supersaturations. If the particle size remains fixed (as with our nominal model), an increase in vapor concentration is not always accommodated by efficient condensation, and supersaturations can persist. The fixed-radius scheme is thus more sensitive to assumptions about particle size. For the vapor pressure of water ice at low temperatures, we use the following expression from Marti and Mauersberger (1993):

log10 P = 12.537 −

and, to a lesser extent, hν

H2 O → H + OH OH + C2 H6 → H2 O + C2 H5 Net: C2 H6 → C2 H5 + H.

hν

H2 O → H + OH M

APPENDIX B: PHOTOCHEMISTRY DETAILS Water introduced to Saturn’s atmosphere from the ablation of IDPs or ring particles or from the diffusion of ring vapor is very stable in the upper stratosphere. Although H2 O is readily dissociated at wavelengths less than 185 nm, shielding by methane, ethane, and acetylene increases the lifetime of a water molecule from 6 months at the top of the atmosphere to 15 months at 10−3 mbar (see Table I). More importantly, H2 O is very efficiently recycled once it is dissociated. For example, in our nominal model, 96% of the water that is photolyzed in Saturn’s stratosphere and upper troposphere is recycled through schemes such as

Net: H2 → 2 H,

R248

OH + C2 H2 → CH3 CO

R321

CH3 CO + H → HCO + CH3

R408

HCO + H → CO + H2

R346

Net: H2 O + C2 H2 + H → CO + CH3 + H2 ,

(7)

and hν

H2 O → H + OH M

R248

OH + C2 H2 → CH3 CO

R321

CH3 CO + H → H2 CCO + H2

R407

hν 1

H2 CCO → CH2 + CO Net: H2 O + C2 H2 → CO + CH2 + H2 , 1

hν

where the pressure P is in Pa, the temperature T is in K, and the expression is valid in the region 170 < T < 273 K. Note that we dangerously extrapolate this expression to temperatures below which the vapor pressures have been measured in the laboratory. See Moses et al. (2000) for a discussion of the condensation of hydrocarbon molecules.

OH + H2 → H2 O + H

(6)

Scheme (5) is responsible for 97% of the recycling of H2 O. Although the above schemes represent a means for the catalytic destruction of H2 and C2 H6 , their effects on hydrocarbon chemistry or on the production of atomic H are minor for the globally averaged water influx rates inferred from the ISO observations. The remaining 4% of the total column photochemical loss of H2 O above 5 bar in our model results in the permanent conversion of H2 O into CO and CO2 through schemes such as

H2 O → H + OH

hν

R327

R262 (8)

and

2663.5 , T

H2 O → H + OH

R248

R248 R314 (5)

OH + CO → CO2 + H Net: H2 O + CO → CO2 + 2 H.

R248 R330 (9)

Photochemical schemes such as these (and others involving reactions R317 and R324) account for ∼70% of the permanent removal of H2 O from the stratosphere; condensation accounts for the other ∼30%. Although condensation represents an effective loss mechanism for H2 O in the stratosphere, evaporation of water-ice grains in the troposphere balances the stratospheric condensation, and there is no net column condensation loss of H2 O in our model. Because the column-integrated production rate for H2 O exceeds the column loss rate, water vapor diffuses downward through the lower boundary of our model at a rate of ∼5 × 105 cm−2 s−1 . The meteoroid ablation source of CO2 in our nominal model (7.5 × 104 molecules cm−2 s−1 ) represents 61% of the total column production rate of carbon dioxide in Saturn’s upper atmosphere. The remaining 39% of the CO2 is produced through reaction R330 (OH + CO → CO2 + H) via reaction scheme (9). Although the rate coefficient for reaction R330 is low (∼1.5 × 10−13 cm3 s−1 at temperatures relevant to Saturn’s stratosphere according to Atkinson et al. 1992 and Frost et al. 1993), CO and OH (from water photolysis) exist in sufficient quantities that CO2 is produced at a slow but steady rate in Saturn’s upper atmosphere. Scheme (9) is also believed to be the dominant mechanism for synthesizing CO2 on Titan (Samuelson et al. 1983, Yung et al. 1984, Toublanc et al. 1995, Lara et al. 1996).

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Carbon dioxide is lost by photolysis (R253 dominates over R252) and, to a lesser extent, by reaction with hydrocarbon radicals. Recycling schemes such as hν

CO2 → CO + O(1 D) O( D) + H2 → OH + H

R253

1

R303

OH + CO → CO2 + H

R330

Net: H2 → 2 H

(10)

Our scheme for the fate of incoming oxygen on Saturn more closely follows that of Prather et al. (1978) than Strobel and Yung (1979), primarily because atomic oxygen is not our dominant product from the ablation process. In our model, roughly 35% of the total stratospheric column production of CO results from meteoroid ablation, the other 65% is due to photochemical processes. The photochemical production of CO is very complex—nearly every oxygen species shown in Fig. 2 has some pathway leading to the production of CO. Roughly 78% of the photochemical production of CO in the stratosphere derives from the photolysis of water through reaction schemes such as schemes (7), (8), and (in decreasing order of importance)

are very inefficient and recover less than 1% of the total column loss rate of CO2 . The chemical destruction of CO2 results in the production of H2 O and CO (∼ 65% of the time) through reaction schemes such as hν

CO2 → CO + O( D) 1

hν

H2 O → H + OH

R248

M

OH + C2 H4 → C2 H4 OH C2 H4 OH + H → CH3 CHO + H2

R253

hν

R324 R419

1

R303

CH3 CHO → CH3 + HCO

R264

OH + H2 → H2 O + H

R314

HCO + H → CO + H2

R346

O( D) + H2 → OH + H

Net: CO2 + 2 H2 → CO + H2 O + 2 H,

Net: H2 O + C2 H4 + H → CO + CH3 + 2 H2 ,

(11)

or results in the production of 2 CO molecules (∼35% of the time) through schemes such as

and hν

hν

CO2 → CO + O O + CH3 → H2 CO + H hν

H2 CO → H2 + CO Net: CO2 + CH3 → 2 CO + H2 + H,

(14)

H2 O → H + OH

R252

M

R270 R256

OH + C2 H2 → CH3 CO

R321

CH3 CO + CH3 → CO + C2 H6

R411

Net: H2 O + C2 H2 + CH3 → CO + C2 H6 + H,

(12)

R248

(15)

and

and hν

CO2 → CO + O

R252

O + C2 H4 → HCO + CH3

R279

HCO + H → CO + H2

R346

Net: CO2 + C2 H4 + H → 2 CO + CH3 + H2 .

(13)

hν

H2 O → H + OH M

OH + CH3 → CH3 OH hν

CH3 OH → CH3 O + H CH3 O + H → H2 CO + H2 hν

Photolysis of CO2 initiates one of the few pathways that can convert carbon– oxygen bonds into hydrogen–oxygen bonds (e.g., through schemes such as (11)). In general, carbon–oxygen bonds are quite stable in Saturn’s upper atmosphere. Carbon monoxide has long been identified on Saturn (Noll et al. 1986, Noll and Larson 1990). The observations do not indicate whether the CO is concentrated in the stratosphere (which would indicate an external source) or whether the CO mixing ratio is constant throughout the troposphere and stratosphere (which would indicate an internal source), although Noll and Larson (1990) favor the tropospheric source. Two decades ago, investigators suggested that the CO observed on Jupiter could be due to an external source of oxygen flowing in from the jovian satellites (Strobel and Yung 1979) or from meteoroid bombardment (Prather et al. 1978). In the scheme of Strobel and Yung (1979), oxygen ions that stream into Jupiter’s atmosphere from the Io torus would be quickly converted to O and OH, which could then react with CH3 to produce H2 CO. The H2 CO is then photolyzed to form CO or HCO, and the HCO can in turn react with H or CH3 to form CO. Carbon monoxide would be the dominant end product of all the oxygen photochemistry. The scheme of Prather et al. (1978) is similar, except that H2 O is the initial form of the incoming oxygen, and water photolysis would lead predominantly to the production of OH. The OH then reacts with C2 H2 and C2 H4 to eventually produce CO through a complex series of intermediate reactions involving species such as H2 CCO, C2 H4 OH, CH3 CHO, and HCO. Reactions of atomic oxygen with hydrocarbons would be less important in synthesizing CO.

H2 CO → H2 + CO Net: H2 O + CH3 → CO + 2 H2 + H.

R248 R317 R260 R380 R256 (16)

Note that, as with the Prather et al. (1978) model, reactions of OH with C2 H2 , C2 H4 , and CH3 are the most important for converting water to CO in Saturn’s stratosphere. Reactions involving R321 (OH + C2 H2 + M → CH3 CO + M) dominate the photochemical production of CO (∼53%), followed distantly by R324 (OH + C2 H4 + M → C2 H4 OH + M) at ∼10%, R317 (OH + CH3 + M → CH3 OH + M) at ∼10%, and direct CO2 photolysis (at ∼9%). Reaction schemes involving R270 (O + CH3 → H2 CO + H) that were considered dominant in the Strobel and Yung (1979) Jupiter model are less important in our model (&3%). However, if it turns out that atomic O is an important ablation product, then reaction schemes involving R270 will also operate effectively. Direct recycling accounts for ∼13% of the photochemical column production of CO in the stratosphere. The predominant recycling scheme is M

CO + H → HCO HCO + H → CO + H2 Net: 2 H → H2 .

R342 R346 (17)

EFFECTS OF EXTRAPLANETARY OXYGEN ON SATURN PHOTOCHEMISTRY Carbon dioxide photolysis (R252, R253) and other indirect recycling schemes involving CO2 account for another ∼9% of the photochemical production rate of CO in the strtosphere. Although CO can be dissociated by ultraviolet radiation at wavelengths less than 112 nm, shielding by hydrogen and methane in Saturn’s stratosphere makes CO very long lived. The lifetime of CO against photolysis is 120 years at the methane homopause (1.5 × 10−5 mbar) and virtually infinite at 10−3 mbar. Chemical loss processes are inefficient—CO can react with OH to form CO2 (reaction R330) or with H to form HCO (reaction R342); however, in both cases, the CO is efficiently recycled. The important CO recycling schemes in the stratosphere are scheme (17), and schemes that involve CO2 : OH + CO → CO2 + H hν

CO2 → CO + O(1 D) Net: OH → O(1 D) + H,

here is hν

CH3 OH → CH3 O + H

R260

CH3 O + H → OH + CH3

R379

OH + H2 → H2 O + H

R314

Net: CH3 OH + H2 → H2 O + CH3 + H.

R330 hν

R260

CH3 O + H → H2 CO + H2

R380

CH3 OH → CH3 O + H

R253 (18)

hν

H2 CO → H2 + CO Net: CH3 OH → CO + 2 H2 ,

hν

CO2 → CO + O Net: OH → O + H.

R330

M

2 HCO → CO + H2 CO hν

H2 CO → H2 + CO Net: 2 H → H2 ,

M

hν

hν

H2 CO → H2 + CO Net: CH3 OH → CO + 2 H2 ,

R353 R256 (20)

hν

CH3 OH → CH3 O + H CH3 O + H → H2 CO + H2 H2 CO → HCO + H

Net: Nothing.

(21)

The net column production rate of CO in the upper atmosphere of our model is balanced by the diffusion of CO (at a rate of 2.0 × 106 cm−2 s−1 ) downward through the lower boundary. Deeper in the troposphere, the CO will react with H2 to be converted to the equilibrium forms of carbon and oxygen on Saturn, CH4 , and H2 O. The stratospheric column density (above 63 mbar) of CO in our nominal model is 9.8 × 1016 cm−2 . Methanol (CH3 OH) is produced from the three-body association of CH3 and OH (reaction R317), where the OH results predominantly from H2 O photolysis (reaction R248). Reaction R317 is effective throughout the middle and upper stratosphere of Saturn. Photochemical production of CH3 OH accounts for 77% and meteoric ablation for 23% of the total column production of CH3 OH in our nominal model. Photolysis is the main CH3 OH loss mechanism: reaction R260 dominates, followed by R259 and R258. Slightly over half of the time, CH3 OH photolysis leads to the production (recycling) of H2 O. The dominant scheme

R256 (24)

R260 R380 R255 R346

Net: CH3 OH → CO + 2 H2 .

(25)

The column abundance of CH3 OH in our nominal model is 1.6 × 1013 cm−2 . Although formaldehyde (H2 CO), ketene (H2 CCO), and acetaldehyde (CH3 CHO) are not very abundant in our model, they are important intermediaries in converting water to CO. In the troposphere, H2 CO is formed from CO, with the predominant scheme being M

R254

R259

HCO + H → CO + H2

R342

HCO → H + CO

(23)

and

hν

and CO + H → HCO

hν

CH3 OH → H2 CO + H2

(19)

R342

R256

and

R252

In the troposphere, the internal source of CO allows large CO concentrations at higher pressures (i.e., the downward flux of CO is reduced from what would be the case with no internal CO). Scheme (17) is prevalent, as are other reaction schemes that continuously recycle the abundances of CO, HCO, and H2 CO (e.g., R342, R346, R353, R254, R352, R351, and R256). Schemes (18) and (19) are less important in the troposphere and are replaced by schemes such as 2(CO + H → HCO)

(22)

The remaining half the time, CH3 OH photolysis results in the production of CO through numerous schemes that all have H2 CO as an intermediate. The top three schemes that convert CH3 OH to CO are

and OH + CO → CO2 + H

197

2(CO + H → HCO) 2 HCO → CO + H2 CO Net: CO + 2 H → H2 CO.

R342 R353 (26)

The increase in the mixing ratio of H2 CO with depth in the troposphere (Fig. 4) is due to the internal source of CO. In the stratosphere, H2 CO is formed from H2 O 85% in the time and from CO2 15% of the time. The dominant schemes are (in decreasing order of importance) hν

H2 O → H + OH M

OH + CH3 → CH3 OH hν

CH3 OH → CH3 O + H

R248 R317 R260

CH3 O + H → H2 CO + H2

R380

Net: H2 O + CH3 → H2 CO + H2 + H,

(27)

198

MOSES ET AL.

and hν

H2 O → H + OH M

OH + CH3 → CH3 OH hν

CH3 OH → H2 CO + H2 Net: H2 O + CH3 → H2 CO + H2 + H,

R248 R317 R259 (28)

and scheme (12). Formaldehyde is lost by photolysis to form CO, either directly (i.e., R256, R257) or indirectly (e.g., R255 followed by R346). The column abundance of H2 CO in our model is 2×1011 cm−2 in the stratosphere and 6×1012 cm−2 above 5 bar—quantities that are unobservable with current technology. Ketene is produced in our model predominantly through hν

Anastasi, C., and P. R. Maw 1982. Reaction kinetics in acetyl chemistry over a wide range of temperature and pressure. J. Chem. Soc. Farad. Trans. 1 78, 2423–2433. Arai, H., S. Nagai, and M. Hatada 1981. Radiolysis of methane containing small amount of carbon monoxide—Formation of organic acids. Radiat. Phys. Chem. 17, 211–216. Atkinson, R., D. L. Baulch, R. A. Cox, R. F. Hampson, Jr., J. A. Kerr, and J. Troe 1992. Evaluated kinetic and photochemical data for atmospheric chemistry. Supplement IV. IUPAC subcommittee on gas kinetic data evaluation for atmospheric chemistry. J. Phys. Chem. Ref. Data 21, 1125–1568. Atreya, S. K. 1986. Atmospheres and Ionospheres of the Outer Planets and Their Satellites. Springer-Verlag, Berlin. Baguhl, M., E. Gr¨un, D. P. Hamilton, G. Linkert, R. Riemann, P. Staubach, and H. A. Zook 1995. The flux of interstellar dust observed by Ulysses and Galileo. Space Sci. Rev. 72, 471–476.

R248

Baguhl, M., E. Gr¨un, and M. Landgraf 1996. In situ measurements of interstellar dust with the Ulysses and Galileo spaceprobes. Space Sci. Rev. 78, 165–172.

OH + C2 H2 → CH3 CO

R321

CH3 CO + H → H2 CCO + H2

R407

Banaszkiewicz, M., and A. V. Krivov 1997. Hyperion as a dust source in the saturnian system. Icarus 129, 289–303.

H2 O → H + OH M

Net: H2 O + C2 H2 → H2 CCO + H2 .

(29)

Photolysis to form CO (R262) removes H2 CCO from Saturn’s atmosphere. Acetaldehyde is produced mainly by the reaction scheme hν

H2 O → H + OH M

OH + C2 H4 → C2 H4 OH C2 H4 OH + H → CH3 CHO + H2 Net: H2 O + C2 H4 → CH3 CHO + H2 .

R248 R324 R419

Banyard, S. A., C. E. Canosa-Mas, M. D. Ellis, H. M. Frey, and R. Walsh 1980. Ketene photochemistry. Some observations on the reactions and reactivity of triplet methylene. J. Chem. Soc. Chem. Commun. 1156–1157. Bartels, M., J. Edelb¨uttel-Einhaus, and K. Hoyermann 1991. The detection of CH3 CO, C2 H5 , and CH3 CHO by REMPI/mass spectrometry and the application to the study of the reactions H + CH3 CO and O + CH3 CO. Symp. Int. Combust. Proc. 23, 131–138. Bartels, M., K. Hoyermann, and R. Sievert 1982. Elementary reactions in the oxidation of ethylene: The reaction of OH radicals with ethylene and the reaction of C2 H4 OH radicals with H atoms. Symp. Int. Combust. Proc. 19, 61–72.

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Bass, S. F., and J. I. Moses 1998. The chemistry and structure of Saturn’s ionosphere. Bull. Am. Astron. Soc. 30, 1099.

Photolysis is the dominant loss mechanism, and CO is the ultimate final product.

Baulch, D. L., C. J. Cobos, R. A. Cox, C. Esser, P. Frank, T. Just, J. A. Kerr, M. J. Pilling, J. Troe, R. W. Walker and J. Warnatz 1992. Evaluated kinetic data for combustion modeling. J. Phys. Chem. Ref. Data 21, 411–734.

ACKNOWLEDGMENTS The Caltech/JPL KINETICS code was developed by Y. L. Yung and M. Allen, with assistance from many people over the years. We thank Y. L. Yung, J.-L. Ollivier, J. Cuzzi, E. Gr¨un, H. Zook, M. Landgraf, L.-M. Lara, and an anonymous referee for useful advice and suggestions. This work was supported by NASA Contract NAG5-6915 and by the Lunar and Planetary Institute, which is operated by the Universities Space Research Association under NASA Contract NASW-4574. This paper represents LPI Contribution 983.

REFERENCES Acarreta, J. R., and A. S´anchez-Lavega 1999. Vertical cloud structure in Saturn’s 1990 equatorial storm. Icarus 137, 24–33. Ackerman, M. 1971. Ultraviolet solar radiation related to mesospheric processes. In Mesospheric Models and Related Experiments (G. Fiocco, Ed.), pp. 149– 159. Reidel, Dordrecht, Holland. Adachi, H., N. Basco, and D. G. L. James 1981. The acetyl radicals CH3 CO· and CD3 CO· studied by flash photolysis and kinetic spectroscopy. Int. J. Chem. Kinet. 13, 1251–1276. Allen, M., and J. E. Frederick 1982. Effective photodissociation cross sections for molecular oxygen and nitric oxide in the Schumann–Runge bands. J. Atoms. Sci. 39, 2066–2075. Allen, M., Y. L. Yung, and J. W. Waters 1981. Vertical transport and photochemistry in the terrestrial mesosphere and lower thermosphere (50–120 km). J. Geophys. Res. 86, 3617–3627.

Baulch, D. L., C. J. Cobos, R. A. Cox, P. Frank, G. Hayman, T. Just, J. A. Kerr, T. Murells, M. J. Pilling, J. Troe, R. W. Walker, and J. Warnatz 1994. Evaluated kinetic data for combustion modeling. J. Phys. Chem. Ref. Data 23, 847–1033. [See also Errata, J. Phys. Chem. Ref. Data 24, 1609–1630] B´ezard, B., P. Drossart, E. Lellouch, G. Tarrago, and J. P. Maillard 1989. Detection of arsine in Saturn. Astrophys. J. 346, 509–513. B´ezard, B., H. Feuchtgruber, J. I. Moses, and T. Encrenaz 1998. Detection of methyl radicals (CH3 ) on Saturn. Astron. Astrophys. 334, L41–L44. Black, G.,a R. L. Sharpless, T. G. Slanger, and M. R. Taherian 1985. The 1150– 1300 A absorption spectrum of O2 at 930 K. Chem. Phys. Lett. 113, 311–313. B¨ohland, T., S. D´ob´e, F. Temps, and H. G. Wagner 1985. Kinetics of reaction between CH2 (X 3 B1 )-radicals and saturated hydrocarbons in the temperature range 296 K ≤ T ≤ 707 K. Ber. Bunsenges. Phys. Chem. 89, 1110–1116. Callear, A. B., and I. A. Cooper 1990. Formation of CH3 CHO in the combination of CH3 with CHO. J. Chem. Soc. Farad. Trans. 86, 1763–1775. Canosa-Mas, C. E., H. M. Frey, and R. Walsh 1984. Studies of methylene chemistry by pulsed laser-induced decomposition of ketene. 1. Ketene in the presence of noble gases. J. Chem. Soc. Farad. Trans. 2 80, 561–578. Carver, J. H., H. P. Gies, T. I. Hobbs, B. R. Lewis, and D. G. McCoy 1977. Temperature dependence of the molecular oxygen photoabsorption cross section near the H Lyman α line. J. Geophys. Res. 82, 1955–1960. Cobos, C. J., and J. Troe 1985. The influence of potential energy parameters on the reaction H + CH3 * ) CH4 . Chem. Phys. Lett. 113, 419–424. Cohen, N., and K. R. Westerberg 1991. Chemical kinetic data sheets for hightemperature reactions. Part II. J. Phys. Chem. Ref. Data 20, 1211–1311.

EFFECTS OF EXTRAPLANETARY OXYGEN ON SATURN PHOTOCHEMISTRY Connerney, J. E. P. 1986. Magnetic connection for Saturn’s rings and atmosphere. Geophys. Res. Lett. 13, 773–776.

199

Connerney, J. E. P., and J. H. Waite 1984. New model of Saturn’s ionosphere with an influx of water from the rings. Nature 312, 136–138.

Encrenaz, T. 1998. Jupiter and Saturn: Chemical composition in the stratosphere and the upper troposphere from ISO 6–180 µm observations. In International Symposium: The Jovian System after Galileo, the Saturnian System before Cassini-Huygens, May 11–15, Nantes, France, pp. 15–16.

Conrath, B. J., and J. A. Pirraglia 1983. Thermal structure of Saturn from Voyager infrared measurements: Implications for atmospheric dynamics Icarus 53, 286–293.

Fagerstr¨om, K., A. Lund, G. Mahmoud, J. T. Jodkowski, and E. Ratajczak 1993. Kinetics of the cross reaction between methyl and hydroxyl radicals. Chem. Phys. Lett. 204, 226–234.

Courtin, R., D. Gautier, A. Marten, B. B´ezard, and R. Hanel 1984. The composition of Saturn’s atmosphere at northern temperate latitudes from Voyager IRIS spectra: NH3 , PH3 , C2 H2 , C2 H6 , CH3 D, CH4 , and the saturnian D/H isotopic ratio. Astrophys. J. 287, 899–916.

Failes, R. L., D. L. Singleton, G. Paraskevopoulos, and R. S. Irwin 1982. Rate constants for the reaction of ground-state oxygen atoms with methanol from 297 to 544 K. Int. J. Chem. Kinet. 14, 371–379.

Coustenis, A., A. Salama, E. Lellouch, T. Encrenaz, G. L. Bjoraker, R. E. Samuelson, T. de Graauw, H. Feuchtgruber, and M. F. Kessler 1998. Evidence for water vapor in Titan’s atmosphere from ISO/SWS data. Astron. Astrophys. 336, L85–L89.

Fegley, B. F., Jr., and K. Lodders 1994. Chemical models of the deep atmospheres of Jupiter and Saturn. Icarus 110, 117–154. Festou, M. C., and S. K. Atreya 1982. Voyager ultraviolet stellar occultation measurements of the composition and thermal profiles of the saturnian upper atmosphere. Geophys. Res. Lett. 9, 1147–1150.

Crovisier, J. 1993. Molecular abundances in comets. In Asteroids, Comets, Meteors 1993 (A. Milani, M. Di Martino, and A. Cellino, Eds.), pp. 313– 326. Kluwer, Dordrecht.

Feuchtgruber, H., E. Lellouch, T. de Graauw, B. B´ezard, T. Encrenaz, and M. Griffin 1997. External supply of oxygen to the atmospheres of the giant planets. Nature 389, 159–162.

Cuzzi. J. N., and R. H. Durisen 1990. Bombardment of planetary rings by meteoroids: General formulation and effects of Oort-cloud projectiles. Icarus 84, 467–501.

Feuchtgruber, H., E. Lellouch, T. Encrenaz, B. B´ezard, A. Coustenis, P. Drossart, A. Salama, T. de Graauw, and G. R. Davis 1999. Oxygen in the stratospheres of the giant planets and Titan. In The Universe as Seen by ISO (P. Cox and M. F. Kessler, Eds.), ESA Special Publication SP-427, pp. 133–136.

Cuzzi, J. N., and P. R. Estrada 1998. Compositional evolution of Saturn’s rings due to meteoroid bombardment. Icarus 132, 1–35. de Graauw, T., and 61 colleagues 1996. Observing with the ISO short-wavelength spectrometer. Astron. Astrophys. 315, L49–L52. de Graauw, T., H. Feuchtgruber, B. B´ezard, P. Drossart, T. Encrenaz, D. A. Beintema, M. Griffin, A. Heras, M. Kessler, K. Leech, E. Lellouch, P. Morris, P. R. Roelfsema, M. Roos-Serote, A. Salama, B. Vandenbussche, E. A. Valentijn, G. R. Davis, and D. A. Naylor 1997. First results of ISO-SWS observations of Saturn: Detection of CO2 , CH3 C2 H, C4 H2 and tropospheric H2 O. Astron. Astrophys. 321, L13–L16. DeMore, W. B., D. M. Golden, R. F. Hampson, C. J. Howard, M. J. Kurylo, J. J. Margitan, M. J. Molina, A. R. Ravishankara, and R. T. Watson 1985. Chemical Kinetics and Photochemical Data for Use in Stratospheric Modeling, Evaluation 7, JPL Publication 85–37. Jet Propulsion Laboratory, Pasadena, CA. DeMore, W. B., D. M. Golden, R. F. Hampson, C. J. Howard, M. J. Kurylo, M. J. Molina, A. R. Ravishankara, and S. P. Sander 1990. Chemical Kinetics and Photochemical Data for Use in Stratospheric Modeling, Evaluation 9, JPL Publication 90–1. Jet Propulsion Laboratory, Pasadena, CA. DeMore, W. B., S. P. Sander, D. M. Golden, R. F. Hampson, M. J. Kurylo, C. J. Howard, A. R. Ravishankara, and M. J. Molina 1987. Chemical Kinetics and Photochemical Data for Use in Stratospheric Modeling, JPL Publication 87–41. Jet Propulsion Laboratory, Pasadena, CA. DeMore, W. B., S. P. Sander, D. M. Golden, R. F. Hampson, M. J. Kurylo, C. J. Howard, A. R. Ravishankara, C. E. Kolb, and M. J. Molina 1992. Chemical Kinetics and Photochemical Data for Use in Stratospheric Modeling, Evaluation 10, JPL Publication 92–20. Jet Propulsion Laboratory, Pasadena, CA. Ditchburn, R. W., and P. A. a Young 1962. The absorption of molecular oxygen between 1850 and 2500 A. J. Atmos. Terr. Phys. 24, 127–139. D´ob´e, S., T. B´erces, and I. Szil´agyi 1991. Kinetics of the reaction between methoxyl radicals and hydrogen atoms. J. Chem. Soc. Farad. Trans. 87, 2331– 2336. Duncan, M., T. Quinn, and S. Tremaine 1988. The origin of short-period comets. Astrophys. J. 328, L69–L73. Durisen, R. H. 1984. Transport effects due to particle erosion mechanisms. In Planetary Rings (G. Greenberg and A. Brahic, Eds.), pp. 416–446. Univ. of Arizona Press, Tucson. Durisen, R. H., N. L. Cramer, B. W. Murphy, J. N. Cuzzi, T. L. Mullikin, and S. E. Cederbloom 1989. Ballistic transport in planetary ring systems due to particle erosion mechanisms. Icarus 80, 136–166.

Fouchet, T., R. Prang´e, J. E. C. Connerney, R. Courtin, L. Ben-Jaffel, K. Noll, and J. C. McConnell 1996. HST spectro-imaging of Saturn and search for water from the rings. Bull. Am. Astron. Soc. 28, 1129. Frost, M. J., P. Sharkey, and I. W. M. Smith 1993. Reaction between OH (OD) radicals and CO at temperatures down to 80 K: Experiment and theory. J. Phys. Chem. 97, 12254–12259. Gentieu, E. P., and J. E. Mentall 1970.a Formaldehyde absorption coefficients in the vacuum ultraviolet (650–1850 Angstroms). Science 169, 681–683. Gibson, S. T., H. P. F. Gies, A. J. Blake, D. G. McCoy, and P. J. Rogers 1983. Temperature dependence in the Schumann-Runge photoabsorption continuum of oxygen. J. Quant. Spectrosc. Radiat. Trans. 30, 385–393. Gladstone, G. R., M. Allen, and Y. L. Yung 1996. Hydrocarbon photochemistry in the upper atmosphere of Jupiter. Icarus 119, 1–52. a Glicker, S., and L. J. Stief 1971. Photolysis of formaldehyde at 1470 and 1236 A. J. Chem. Phys. 54, 2852–2857. Greenberg, J. M. 1998. Making a comet nucleus. Astron. Astrophys. 330, 375– 380. Grotheer, H.-H., G. Riekert, D. Walter, and T. Just 1989. Reactions of hydroxymethyl and hydroxymethyl radicals with molecular and atomic oxygen. Symp. Int. Combust. Proc. 22, 963–972. Gr¨un, E., B. Gustafson, I. Mann, M. Baguhl, G. E. Morfill, P. Staubach, A. Taylor, and H. A. Zook 1994. Interstellar dust in the heliosphere. Astron. Astrophys. 286, 915–924. Gr¨un, E., P. Staubach, M. Baguhl, D. P. Hamilton, H. A. Zook, S. Dermott, B. A. Gustafson, H. Fechtig, J. Kissel, D. Linkert, G. Linkert, R. Sama, M. S. Hanner, C. Polanskey, M. Horanyi, B. A. Lindblad, I. Mann, J. A. M. McDonnell, G. E. Morfill, and G. Schwehm 1997. South–north and radial traverses through the interplanetary dust cloud. Icarus 129, 270–288. Gr¨un, E., H. A. Zook, M. Baguhl, A. Balogh, S. J. Bame, H. Fechtig, R. Forsyth, M. S. Hanner, M. Horanyi, J. Kissel, B.-A. Lindblad, D. Linkert, I. Mann. J. A. M. McDonnell, G. E. Morfill, J. L. Phillips, C. Polanskey, G. Schwehm, N. Siddique, P. Staubach, J. Sveska, and A. Taylor 1993. Discovery of jovian dust streams and interstellar grains by the Ulysses spacecraft. Nature 362, 428–430. Gr¨un, E., H. A. Zook, H. Fechtig, and R. H. Giese 1985. Collisional balance of the meteoritic complex. Icarus 62, 244–272. Gurnett, D. A., J. A. Ansher, W. S. Kurth, and L. J. Granroth 1997. Micron-sized dust particles detected in the outer Solar System by Voyager 1 and 2 plasma wave instruments. Geophys. Res. Lett. 24, 3125–3128.

200

MOSES ET AL.

Gustafson, B. A. S., and M. S. Hanner (Eds.) 1996. Physics, Chemistry, and Dynamics of Interplanetary Dust. Astronomical Society Pacific, San Francisco. Haddad, G. N., and J. A. R. Samson 1986. Total absorption and photoionization a cross sections of water between 100 and 100 A. J. Chem. Phys. 84, 6623– 6626.

Lara, L. M., E. Lellouch, J. J. L´opez-Moreno, and R. Rodrigo 1996. Vertical distribution of Titan’s atmospheric neutral constituents. J. Geophys. Res. 113, 23,261–23,284. Lawrence, G. M. 1972a. Photodissociation of CO2 to produce CO(a 3 II). J. Chem. Phys. 56, 3435–3442.

Hag´ege, J., P. C. Roberge, and C. Vermeil 1968. Methanol photochemistry: Collision-induced phenomena. Ber. Bunsenges. Phys. Chem. 72, 138–141.

Lawrence, G. M. 1972b. Production of O(1 S) from photodissociation of CO2 . J. Chem. Phys. 57, 5616–5617.

Hammel, H. B. 1996. HST imaging of Jupiter shortly after each impact: Plumes and fresh sites. In The Collision of Comet Shoemaker–Levy 9 and Jupiter (K. S. Noll, H. A. Weaver, and P. D. Feldman, Eds.), pp. 111–120. Cambridge Univ. Press, Cambridge.

Lee, L. C., and M. Suto 1986. Quantatative photoabsorption and fluorescence study of H2 O and D2 O at 50–190 nm. Chem. Phys. 110, 161–169.

Hassinen, E., K. Kalliorinne, and J. Koskikallio 1990. Kinetics of reactions between methyl and acetyl radicals in gas phase produced by flash photolysis of acetic anhydride. Int. J. Chem. Kinet. 22, 741–745. Hayashi, T., Y. Iida, Y. Morioka, M. Sasanuma, E. Ishiguro, and M. Nakamura 1986. Photoionisation mass spectrometry of O2 in the VUV region. Atom. Molec. Phys. 19, 2861–2870. Herron, J. T. 1988. Evaluated chemical kinetic data for the reactions of atomic oxygen O(3 P) with saturated organic compounds in the gas phase. J. Phys. Chem. Ref. Data 17, 967–1026. Hitchcock, A. P., and C. E. Brion 1980. Absolute oscillator strengths for valenceshell ionic photofragmentation of N2 O and CO2 (8–75 eV). Chem. Phys. 45, 461–478. Hochanadel, C. J., T. J. Sworski, and P. J. Ogren 1980. Ultraviolet spectrum and reaction kinetics of the formyl radical. J. Phys. Chem. 84, 231–235. H¨ohlein, G., and G. R. Freeman 1970. Radiation-sensitized pyrolysis of dimethyl ether. Free-radical reaction rate parameters. J. Am. Chem. Soc. 92, 6118– 6125. Hughes, D. W. 1996. The size, mass and evolution of the Solar System dust cloud. Q. J. R. Astron. Soc. 37, 593–604. Humes, D. H. 1980. Results of Pioneer 10 and 11 meteoroid experiments: Interplanetary and near Saturn. J. Geophys. Res. 85, 5841–5852. Hunten, D. M., R. P. Turco, and O. B. Toon 1980. Smoke and dust particles of meteoric origin in the mesosphere and stratosphere. J. Atmos. Sci. 37, 1342–1357. Ip, W. H. 1983. On plasma transport in the vicinity of the rings of Saturn: A siphon flow mechanism. J. Geophys. Res. 88, 819–822. Johnston, H. S., M. Paige, and F. Yao 1984. Oxygen absorption cross sections in the Herzberg continuum and between 206 K and 327 K. J. Geophys. Res. 89, 11661–11665. Karkoschka, E., and M. G. Tomasko 1993. Saturn’s upper atmospheric hazes observed by the Hubble Space Telescope. Icarus 106, 428–441.

Lee, L. C., T. G. Slanger, G. Black, and R. L. Sharpless 1977. Quantum yields a for the production of O(1 D) from photodissociation of O2 at 1160–1770 A. J. Chem. Phys. 67, 5602–5606. Lellouch, E. 1996. Chemistry induced by the impacts: Observations. In The Collision of Comet Shoemaker–Levy 9 and Jupiter, (K. S. Noll, H. A. Weaver, and P. D. Feldman, Eds.), pp. 243–268. Cambridge Univ. Press, Cambridge. Lellouch, E., B. B´ezard, R. Moreno, D. Bockel´ee-Morvan, P. Colom, J. Crovisier, M. Festou, D. Gautier, A. Marten, and G. Paubert 1997. Carbon monoxide in Jupiter after the impact of Comet Shoemaker–Levy 9. Planet. Space Sci. 45, 1203–1212. Lellouch, E., H. Feuchtgruber, T. de Graauw, T. Encrenaz, B. B´ezard, M. Griffin, and G. Davis 1998. D/H ratio and oxygen source: A Jupiter-Saturn comparison. In International Symposium: The Jovian System after Galileo, the Saturnian System before Cassini–Huygens, May 11–15, Nantes, France, pp. 21–22. Letzelter, C. M. Eidelsberg, F. Rostas, J. Breton, and B. Thieblemont 1987. Photoabsorption and photodissociation cross sections of CO between 88.5 and 115 nm. Chem. Phys. 114, 273–288. Levison, H. F., and M. J. Duncan 1997. From the Kuiper belt to Jupiter-family comets: The spatial distribution of ecliptic comets. Icarus 127, 13–32. Lewis, B. R., and J. H. Carver 1983. Temperature dependence ofa the carbon dioxide photoabsorption cross section between 1200 and 1970 A. J. Quant. Spectrosc. Radiat. Trans. 30, 297–309. Lewis, B. R., S. T. Gibson, M. Emami, and J. H. Carver 1988a. Resonances in the photodissociation of isotopic molecular oxygen. I. The longest band. J. Quant. Spectrosc. Radiat. Trans. 40, 1–13. Lewis, B. R., S. T. Gibson, M. Emami, and J. H. Carver 1988b. Resonances in the photodissociation of isotopic molecular oxygen. II. The second and third bands. J. Quant. Spectrosc. Radiat. Trans. 40, 469–477. Liou, J.-C., H. A. Zook, and S. F. Dermott 1996. Kuiper belt dust grains as a source of interplanetary dust particles. Icarus 124, 429–440. Liou, J.-C., H. A. Zook, and A. A. Jackson 1999. Orbital evolution of retrograde interplanetary dust particles and their distribution in the Solar System. Icarus 141, 13–28.

Kary, D. M., and L. Dones 1996. Capture statistics of short-period comets: Implications for Comet D/Shoemaker–Levy 9. Icarus 121, 207–224. Kessler, M. F., J. A. Steinz, M. E. Anderegg, J. Clavel, G. Drechsel, P. Esteria, J. Faelker, J. R. Riedinger, A. Robson, B. G. Taylor, and S. Ximenz de Ferr´an 1996. The Infrared Space Observatory (ISO) mission. Astron. Astrophys. 315, L27–L30.

Love, S. G., and D. E. Brownlee 1993. A direct measurement of the terrestrial mass accretion rate of cosmic dust. Science 262, 550–553.

Kirby, K. E. R. Constantinides, S. Babeu, M. Oppenheimer, and G. A. Victor 1979. Photoionization and photoabsorption cross-sections of He, O, N2 and O2 for aeronomic calculations. Atomic Data Nucl. Data Tables 23, 63–81.

Mallard, W. G., F. Westley, J. T. Herron, R. F. Hampson, and D. H. Frizzell 1994. NIST Chemical Kinetics Database: Version 6.0. National Institute of Standards and Technology, Gaithersberg, MD.

Kley, D. 1984. Ly (α) absorption cross-section of H2 O and O2 . J. Atmos. Chem. 2, 203–210.

Marsden, B. G., and G. V. Williams 1995. Catalogue of Cometary Orbits 1995. International Astronomical Union, Central Bureau for Astronomical Telegrams, and Minor Planet Center, Cambridge, MA.

K¨onig, R., S. Mory, A. Rosenfeld, J. Lademann, R. Grundwald, and G. Winkelmann 1984. Laser absorption and fluorescence spectroscopy of excited molecules in nanosecond time scale. J. Mol. Struct. 114, 101–104. Landgraf, M., J.-C. Liou, H. A. Zook, and E. Gr¨un 1999. Multi spacecraft data on dust in the outer solar system. Lunar Planet. Sci. Conf. 30. [CD ROM] Lange, M. A., and T. J. Ahrens 1987. Impact experiments in low-temperature ice. Icarus 69, 506–518.

Lyons, J. R. 1996. Atmospheric Chemistry of the Outer Solar System: From 40 K to 4000 K. Ph.D. thesis, California Institute of Technology, Pasadena. Majeed, T., and J. C. McConnell 1991. The upper ionospheres of Jupiter and Saturn. Planet. Space Sci. 39, 1715–1732.

Marti, J., and K. Mauersberger 1993. A survey and new measurements of ice vapor-pressures at temperatures between 170 and 250 K. Geophys. Res. Lett. 20, 363–366. Matsunaga, F. M., and K. Watanabe 1967. Total and photoionization coefficients a and dissociation continua of O2 in the 580–1070 A region. Sci. Light 16, 31–42.

EFFECTS OF EXTRAPLANETARY OXYGEN ON SATURN PHOTOCHEMISTRY Mentall, J. E., E. P. Gentieu, M. Krauss, and D. Neumann 1971. Photoionization and absorption spectrum of formaldehyde in the vacuum ultraviolet. J. Chem. Phys. 55, 5471–5479. Miyoshi, A., H. Matsui, and N. Washida 1989. Reaction of acetaldehyde and acetyl radical with atomic and molecular oxygen. J. Phys. Chem. 93, 5813– 5818. Morfill, G. E., H. Fechtig, E. Gr¨un, and C. K. Goertz 1983. Some consequences of meteoroid impacts on Saturn’s rings. Icarus 55, 439–447. Moses, J. I. 1992. Meteoroid ablation in Neptune’s atmosphere. Icarus 99, 368–383. Moses, J. I. 1996. SL9 impact chemistry: Long-term photochemical evolution. In The Collision of Comet Shoemaker–Levy 9 and Jupiter (K. S. Noll, H. A. Weaver, and P. D. Feldman, Eds.), pp. 243–268. Cambridge Univ. Press, Cambridge. Moses, J. I. 1998. Micrometeoritic bombardment of Saturn. I. Ablation profiles. Lunar Planet. Sci. Conf. 29, 1427. [CD ROM] Moses, J. I., M. Allen, and G. R. Gladstone 1995a. Nitrogen and oxygen photochemistry following SL9. Geophys. Res. Lett. 22, 1601–1604. Moses, J. I., K. Rages, and J. B. Pollack 1995b. An analysis of Neptune’s stratospheric haze using high-phase-angle Voyager images. Icarus 113, 232– 266. Moses, J. I., B. B´ezard, E. Lellouch, G. R. Gladstone. H. Feuchtgruber, and M. Allen 2000. Photochemistry of the upper atmosphere of Saturn. I. Hydrocarbon photochemistry and comparisons with ISO observations. Icarus 143, 244–298. Mulenko, S. A. 1987. The application of an intracavity laser spectroscopy method for elementary processes study in gas-phase reactions. Rev. Roum. Phys. 32, 173–178. Nakata, R. S., K. Watanabe, and F. M. Matsunaga 1965. Absorption and phoa toionization coefficient of CO2 in the region 580–1670 A. Sci. Light 14, 54– 71. Nee, J. B., and L. C. Lee 1984. Photoabsorption cross sections of OH at 115– 183 nm. J. Chem. Phys. 81, 31–36. Nee, J. B., M. Suto, and L. C. Lee 1985. Photoexcitation processes of CH3 OH: Rydberg states and photofragment fluorescence. Chem. Phys. 98, 147–155. Nicolet, M. 1984. On the molecular scattering in the terrestrial atmosphere: An empirical formula for its calculation in the homosphere. Planet. Space Sci. 32, 1467–1468. Noll, K. S., and H. P. Larson 1990. The spectrum of Saturn from 1990 to 2230 cm−1 : Abundances of AsH3 , CH3 D, CO, GeH4 , NH3 , and PH3 . Icarus 89, 168–189. Noll, K. S., R. F. Knacke, T. R. Geballe, and A. T. Tokunaga 1986. The detection of carbon monoxide in Saturn. Astrophys. J. 309, L91–L94. Northrop, T. G., and J. R. Hill 1982. The stability of negatively charged dust grains in Saturn’s ring plane. J. Geophys. Res. 87, 6045–6051. Northrop, T. G., and J. R. Hill 1983. The inner edge of Saturn’s B ring. J. Geophys. Res. 88, 6102–6108. 3 − Ogawa, S., and M. Ogawa 1975. a Absorption cross sections of O2 (X 6g ) in the region from 1087 to 1700 A. Can. J. Phys. 53, 1845–1852. Ohmori, K., A. Miyoshi, H. Matsui, and N. Washida 1990. Studies on the reaction of acetaldehyde and acetyl radicals with atomic hydrogen. J. Phys. Chem. 94, 3253–3255. Okabe, H. 1978. Photochemistry of Small Molecules. Wiley, New York. Ollivier, J. L., M. Dobrij´evic, and J. P. Parisot 2000. New photochemical model of Saturn’s atmosphere. Planet. Space Sci., in press. ¨ Opik, E. J. 1951. Collision probabilities with the planets and the distribution of interplanetary matter. Proc. R. Irish Acad. A 54, 165–199. Oser, H., N. D. Stothard, R. Humpfer, and H. H. Grotheer 1992. Direct measurement of the reaction CH3 + OH at ambient temperature in the pressure range 0.3–6.2 mbar. J. Phys. Chem. 96, 5359–5363.

201

Person, J. C., and P. P. Nicole 1971. Isotope effects in the photoionization yields and in the absorption cross sections for methanol, ethanol, methyl bromide, and ethyl bromide. J. Chem. Phys. 55, 3390–3397. Prang´e, R., T. Fouchet, J. E. P. Connerney, and R. Courtin 1997. FUV remote sensing of the stratosphere of Saturn with the Hubble Space Telescope: Search for the signature of water originating from the rings. In Proc. of Les rencontres de Blois, June 22–27, Blois, France. Prang´e, R., T. Fouchet, and R. Courtin 1998. FUV remote sensing of the atmosphere of Saturn with HST: Polar haze and latitudinal dependence of water abundance. In International Symposium: The Jovian System after Galileo, the Saturnian System before Cassini-Huygens, May 11–15, Nantes, France, pp. 25–26. Prather, M. J., J. A. Logan, and M. B. McElroy 1978. Carbon monoxide in Jupiter’s upper atmosphere: An extraplanetary source. Astrophys. J. 223, 1072–1081. Pravilov, A. N., and I. O. Shul’pyakov 1986. Processes of H2 O and NO2 photodissociation in the spectral region of λ > 155 nm. High Energy Chem. USSR 20, 31–35. Prinn, R. G., H. P. Larson, J. J. Caldwell, and D. Gautier 1984. Composition and chemistry of Saturn’s atmosphere. In Saturn (T. Gehrels and M. S. Matthews, Eds.), pp. 88–149. Univ. of Arizona Press, Tucson. Pruppacher, H. R., and J. D. Klett 1980. Microphysics of Clouds and Precipitation. Reidel, Dordrecht, Holland. Reid, R. C., J. M. Prausnitz, and B. E. Poling 1987. The Properties of Liquids and Gases. McGraw-Hill, New York. Rizk, B., and D. M. Hunten 1990. Solar heating of the uranian mesopause by dust of ring origin. Icarus 88, 429–447. Romani, P. N., and S. K. Atreya 1989. Stratospheric aerosols from CH4 photochemistry on Neptune. Geophys. Res. Lett. 16, 941–944. Rosenqvist, J., E. Lellouch, P. N. Romani, G. Paubert, and T. Encrenaz 1992. Millimeter-wave observations of Saturn, Uranus, and Neptune: CO and HCN on Neptune. Astrophys. J. 392, L99–L102. Samson, J. A. R., and G. Haddad 1988. Absolute cross-sections for molecular photoabsorption, partial photoionization, and ionic photofragmentation processes. J. Phys. Chem. Ref. Data 17, 48. Samson, J. A. R., J. L. Gardner, and G. N. Haddad 1977. a Total and partial photoionization cross-sections of O2 from 100 to 800 A. J. Electron Spect. Relat. Phenom. 12, 281–292. Samson, J. A. R., G. H. Rayborn, and P. N. Pareek 1982. a Dissociative photoionization cross sections of O2 from threshold to 120 A. J. Chem. Phys. 76, 393–397. Samuelson, R. E., W. C. Maguire, R. A. Hanel, V. G. Kunde, D. E. Jennings, Y. L. Yung, and A. C. Aikin 1983. CO2 on Titan. J. Geophys. Res. 88, 8709– 8715. Satyapal, S., J. Park, R. Bersohn, and B. Katz 1989. Dissociation of methanol and ethanol activated by a chemical reaction or by light. J. Chem. Phys. 91, 6873–6879. Scherzer, K., U. Loser, and W. Stiller 1987. BSBL-rechnungen zu wasserstoffabspaltungsreaktionen durch alkenylradikale; vinylradikale. Z. Chem. 27, 300–301. Seinfeld, J. H. 1986. Atmospheric Chemistry and Physics of Air Pollution. Wiley, New York. a Shemansky, D. E. 1972. CO2 extinction coefficient 1700–3000 A. J. Chem. Phys. 56, 1582–1587. Shoemaker, E. M. 1962. Interpretation of lunar craters. In Physics and Astronomy of the Moon (Z. Kopal, Ed.), pp. 283–359. Academic Press, New York. Shoemaker, E. M., and R. F. Wolfe 1982. Cratering time scales for the Galilean satellites of Jupiter. In Satellites of Jupiter (D. Morrison, Ed.), pp. 277–339. Univ. of Arizona Press, Tucson. Simonaitis, R., and J. Heicklen 1972. Kinetics and mechanism of the reaction of O(3 P) with carbon monoxide. J. Chem. Phys. 56, 2004–2011.

202

MOSES ET AL.

Slanger, T. G., and G. Black 1978. CO2 photolysis revisited. J. Chem. Phys. 68, 1844–1849. a Slanger, T. G., and G. Black 1982. Photodissociative channels at 1216 A for H2 O, NH3 , and CH4 . J. Chem. Phys. 77, 2432–2437. Smith, G. R., D. E. Shemansky, J. B. Holberg, A. L. Broadfoot, B. R. Sandel, and J. C. McConnell 1983. Saturn’s upper atmosphere from the Voyager 2 EUV solar and stellar occultations. J. Geophys. Res. 88, 8667–8678. Stark, G., K. Yoshino, P. L. Smith, K. Ito, and W. H. Parkinson 1991. High resolution absorption cross sections of carbon monoxide bands at 295 K between 91.7 and 100.4 nanometers. Astrophys. J. 369, 574–580. Stern, S. A. 1995. Collisional time scales in the Kuiper disk and their implications. Astron. J. 110, 856–868. Stief, L. V., W. A. Payne, and R. B. Klemm 1975. A flash photolysis-resonance fluorescence study of the formation of O(1 D) in the photolysis of water and the reaction of O(1 D) with H2 , Ar, and He. J. Chem. Phys. 62, 4000– 4008. Strobel, D. F., and Y. L. Yung 1979. The Galilean satellites as a source of CO in the jovian upper atmosphere. Icarus 37, 256–263. Summers, M. E., and D. F. Strobel 1989. Photochemistry of the atmophere of Uranus. Astrophys. J. 346, 495–508. Suto, M., X. Wang, and L. C. Lee 1986. Fluorescence from VUV excitation of formaldehyde. J. Chem. Phys. 85, 4228–4233. Taylor, A. D., W. J. Baggaley, and D. I. Steel 1996. Discovery of interstellar dust entering the Earth’s atmosphere. Nature 380, 323–325. Thompson, B. A., P. Harteck, and R. R. Reeves 1963. Ultraviolet absorption coefficients of CO a 2 , CO, O2 , H2 O, N2 O, NH3 , NO, SO2 , and CH4 between 1850 and 4000 A. J. Geophys. Res. 68, 6431–6436. Toublanc, D., J. P. Parisot, J. Brillet, D. Gautier, F. Raulin, and C. P. McKay 1995. Photochemical modeling of Titan’s atmosphere. Icarus 113, 2–26. Tsang, W. 1987. Chemical kinetic data base for combustion chemistry. 2. Methanol. J. Phys. Chem. Ref. Data 16, 471–508. Tsang, W., and R. F. Hampson 1986. Chemical kinetic data base for combustion chemistry. Part 1. Methane and related compounds. J. Phys. Chem. Ref. Data 15, 1087–1279. van Dishoeck, E. F., and A. Dalgarno 1984. Photodissociation of OH in interstellar clouds. Astrophys. J. 277, 576–580. van Dishoeck, E. F., M. C. van Hemert, A. C. Allison, and A. Dalgarno 1984. Resonances in the photodissociation of OH by absorption into coupled 2 II states—Adiabatic and diabatic formulations. J. Chem. Phys. 81, 5709– 5724. Wagner, A. F., and J. M. Bowman 1987. The addition and dissociation reaction H + CO * ) HCO. 1. Theoretical RRKM studies. J. Phys. Chem. 91, 5314– 5324. Wang, J., D. G. McCoy, A. J. Blake, and L. Torop 1987. Effects of the close approach of potential curves in photoabsorption by diatomic molecules. II. Temperature dependence of the O2 cross section in the region 130–160 nm. J. Quant. Spectrosc. Radiat. Trans. 38, 19–27. Washburn, E. W. 1924. The vapor pressure of ice and of water below the freezing point. Monthly Weather Rev. 488–490.

Watanabe, K., and M. Zelikoff 1953. Absorption coefficients of water vapor in the vacuum ultraviolet. J. Opt. Soc. Am. 43, 753–755. Watkins, K. W., and W. W. Word 1974. Addition of methyl radicals to carbon monoxide: Chemically and thermally activated decomposition of acetyl radicals. Int. J. Chem. Kinet. 6, 855–873. West, R. A. 1996. Particles in Jupiter’s atmosphere from the impacts of Comet P/Shoemaker–Levy 9. In The Collision of Comet Shoemaker–Levy 9 and Jupiter (K. S. Noll, H. A. Weaver, and P. D. Feldman, Eds.), pp. 269–292. Cambridge Univ. Press, Cambridge. West, R. A., M. Sato, H. Hart, A. L. Lane, C. W. Hord. K. E. Simmons, L. W. Esposito, D. L. Coffeen, and R. B.a Pomphrey 1983. Photometry and polarimetry of Saturn at 2640 and 7500 A. J. Geophys. Res. 88, 8679–8697. Whytock, D. A., J. V. Michael, W. A. Payne, and L. J. Stief 1976. Absolute rate of the reaction of atomic hydrogen with acetaldehyde. J. Chem. Phys. 65, 4871–4875. Winkelstein, P., J. Caldwell, S. J. Kim, M. Combes, G. E. Hunt, and V. Moore 1983. A determination of the composition of the saturnian stratosphere using the IUE. Icarus 54, 309–318. Witte, M., H. Rosenbauser, M. Banaszkiewicz, and H. Fahr 1993. The Ulysses neutral gas experiment: Determination of the velocity and temperature of interstellar neutral helium. Adv. Space Res. 13(6), 121–130. Wright, G. R., M. J. Van der Wiel, and C. E. Brion 1976. Dipole excitation and fragmentation of N2 and CO in the 10–60 eV region. J. Phys. B 9, 675–689. Wu, C. Y. R., and D. L. Judge 1981. Lyman-α fluorescence from hydrogen photofragments of CH4 and H2 O. J. Chem. Phys. 75, 172–178. Yamamoto, S., and T. Mukai 1998. Dust production by impacts of interstellar dust on Edgeworth-Kuiper Belt objects. Astron. Astrophys. 329, 785–791. Yoshino, K., A. S.-C. Cheung, J. R. Esmond, W. H. Parkinson, D. E. Freeman, S. L. Guberman, A. Jenouvrier, B. Coquart, and M. F. Merienne 1988. Improved absorption cross-sections of oxygen in the wavelength region 205– 240 nm of the Herzberg continuum. Planet. Sci. 36, 1469–475. Yoshino, K., J. R. Esmond, A. S.-C. Cheung, D. E. Freeman, and W. H. Parkinson 1992. High resolution absorption cross sections in the transmission window region of the Schumann–Runge bands and Herzberg continuum of O2 . Planet. Space Sci. 40, 185–192. Yoshino, K., D. E. Freeman, J. R. Esmond, and W. H. Parkinson 1987. High resolution absorption cross-sections and band oscillator strengths of the Schumann–Runge bands of oxygen at 79 K. Planet. Space Sci. 35, 1067–1075. Yung, Y. L., M. Allen, and J. P. Pinto 1984. Photochemistry of the atmosphere of Titan: Comparison between model and observations. Astrophys. J. Suppl. Ser. 55, 465–506. Zabarnick, S., J. W. Fleming, and M. C. Lin 1988. Temperature dependence of CH radical reactions with H2 O and CH2 O. Symp. Int. Combust. Proc. 21, 713–719. Zahnle, K. 1996. Dynamics and chemistry of SL9 plumes. In The Collision of Comet Shoemaker–Levy 9 and Jupiter (K. S. Noll, H. A. Weaver, and P. D. Feldman, Eds.), pp. 183–212. Cambridge Univ. Press, Cambridge. Zahnle, K., L. Dones, and H. Levison 1998. Cratering rates on the Galilean satellites. Icarus 136, 202–222.