1 Supplementary Information: 2 SI- Methods 3 A slice (~1g) of ALH84001-214 was crushed (0.1- 100 μm) and divided into two subsections. The two portions of

4 comminuted samples were reacted separately with concentrated H3PO4 in the side arm of reaction vessels that were -6 o 5 evacuated for 4 days to ~10 Torr. Owing to small amount of CO2 released (~0.5 μmole) at 25 +1 C, a third cut of 6 ALH84001C (~1g) was processed as above and surface impurities were removed by allowing acid to react with the 7 sample for 1 hour at ambient temperature (25 +1oC). Gas released was collected at -196oC after passing through two 8 traps (ethanol slush at -60oC to remove traces of water). The 1-hr extraction was not performed on the first two samples

9 (ALH84001A and ALH84001B) only on the third sample (ALH84001C) (Table 1). CO2 was extracted from all of o o 10 the samples after acid reaction for 12 h at 25 + 1 C. Finally, the samples were heated for 3h at 150 + 1 C and CO2 11 collected following the above procedure. The sequential extraction technique at two temperatures allowed us to 12 separate Ca-rich phase from the Fe-Mg rich phase. Laboratory experiments with various minerals and grain sizes, Al-

o 13 asam et al .,(19) noted that < 10% CO2 is released from dolomite at 25 C after 12h which resulted in <1 ‰ shift in o 14 C and O isotopes of calcite mineral. The contribution of CO2 from siderite and magnesite can be ruled out at 25 C as o 15 no CO2 was released upon heating these mineral to 50 C with phosphoric acid even after 10h (36). CO2 was purified 16 using gas chromatography and C-isotopes were measured using Mat 253 Isotope Ratio Mass spectrometer. For O-

o 17 triple isotope measurement, CO2 gas was reacted with BrF5 at 780 + 10 C for 45h and O2 gas purified 18 chromatographically to allow multi-oxygen isotopic measurements(37). Carbon and O triple isotopic compositions 19 were measured using isotope ratio mass spectrometer. For replicate mass spectrometric analysis 1 SD standard

18 17 2 2 20 deviations is 0.01‰ for δ O and δ O. Overall analytical error (Stot) of the procedure is determined as (Stot) = (S1) 2 2 2 21 + (S2) + (S3) + (S4) where S1= acid digestion, S2= gas chromatography, S3 = fluorination step and S4 = isotope ratio 22 measurement for 17O and 18O= 0.5‰. All these processes follow mass dependent fractionation, therefore, uncertainty 23 on Δ17O = + 0.06‰ based on laboratory standards (n = 5). 24 Ion microprobe oxygen isotope measurements were obtained from two separate thin sections of ALH 84001 25 (302 and 303) using the CAMECA ims 6f at Arizona State University and the CAMECA ims 1270 at the University 26 of California at Los Angeles. Samples and standards were coated with ~20-30 nm of carbon. 27 Section ALH 84001,302 was analyzed at Arizona State University. During analyses, negative secondary 28 ions were sputtered by a Cs+ primary beam with a beam current of 25 nA focused to a spot size of ~20 m diameter. 29 The analysis area was flooded with low energy electrons for charge compensation (as in Leshin et al. (8)). Samples 30 were pre-sputtered for 210 seconds to remove the effects of coating. Each measurement was comprised of 55 cycles 31 of counting for ~1.5 seconds on mass16 and ~5.5 seconds on mass 18. Typical count rates for 16O in the calcite 32 standard were ~ 50 million counts per second. 16O was measured using a Faraday Cup (FC) and 18O was measured 33 using an electron multiplier (EM). Peak intensities were corrected for background (FC) and deadtime (EM). 34 The IMF for each unknown was calculated from measurements of calcite, magnesite and siderite standards. 35 Standards were mounted together on one section, separate from the unknown sample. These standards included Joplin 36 calcite, ZS magnesite, MS-siderite, DM dolomite and 2594 Breunnerite. Instrumental mass fractionation (IMF) varied 37 from ~-18‰ to ~+2‰. Uncertainties for individual analyses are ~2‰.

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38 Section ALH 84001,303 was analyzed on the UCLA ims 1270. During this run, a Cs+ beam with a beam 39 current of ~2 nA was focused to a spot size of ~20 m by ~30 m. Secondary ions were analyzed in multi-collector 40 mode using two Faraday cups. Samples were once again pre-sputtered for 210 seconds. The IMF for each unknown 41 was again calculated from measurements of the same calcite, magnesite and siderite standards listed above. 42 Instrumental mass fractionation (IMF) varied from ~-12‰ to ~-4‰. Uncertainties for individual analyses are again 43 ~2‰. 44 The carbonates measured above were described in the main text, as well as thoroughly discussed in Corrigan 45 and Harvey (21). Similar occurrences were found by 49. Petrographic relationships amongst these carbonates were 46 described in Corrigan and Harvey (21) as well. Specifically, in regard to the carbonates shown in Figure 3, Corrigan 47 and Harvey (21) interpreted that a number of steps were required to form the assemblages seen in the thin section. 48 First, the slab/rosette carbonates (“slab”) were deposited by fluids supersaturated in carbonate components. Rosettes 49 would only have had one nucleation point, while slabs would likely have formed from the coalescence of carbonates 50 growing from multiple nucleation sites in rare, large fractures. The Ca-rich portions of the slab and rosette carbonates 51 were the first to form, with compositions becoming more Mg- and Fe-rich over time. Visible zones were formed as 52 the occasional recharge changed the composition of fluids slightly. Magnesite and siderite rims likely formed on 53 rosette/slab carbonates by high temperature thermal decomposition of carbonate during an impact event(21). The Mg- 54 rich carbonates (termed “post-slab magnesites”(21) but represented in Figure 3 by the label “Cb”) were formed next, 55 intruding into spaces not filled with slab carbonate(21). Feldspathic glasses (“Fs”) were intruded, and, finally, 56 carbonate and silica glasses were formed interstitial to the Mg-rich carbonate and feldspathic glasses produced the 57 final assemblage seen in the meteorite today. 58 59 60 SI-DISCUSSION:

61 We propose three different pathways to generate the O-isotopic anomaly in Martian CO3 1 62 i) CO2–O3 isotope exchange via excited oxygen atom O( D) (R1-R2) (38-41).

63 ii) Catalytic reaction of hydroxyl radicals (HOx) with CO to produce CO2 (R3-R5)(42)

64 iii) Interaction of CO2 with surface adsorbed water on the regolith and aerosol dust particles 65 in the presence of ozone (R6-R8) as suggested by (37).

1 66 OOQ+ h →Q( D) + O2 (R1) 1 3 67 Q( D) + CO2 ↔ CO2Q* ↔ COQ + O( P) (R2) 68 Here Q denote the heavy isotopes of oxygen (17O, 18O) in ozone and the enrichment is transferred

69 to CO2 via short live CO2Q*. 1 70 Q( D) + H2O →OH + QH (R3)

71 CO+ QH → COQ + H (R4)

72 CO+ O → CO2 (R5)

2

73

74 OOQ+ (H2O)ads → (H2OQ)ads + O2 (R6)

75 X(SiO3) + (H2OQ)ads → X(QH)(OH) + SiO3 (R7)

76 CO2+ X(HQ)(HO) → X(QH)---COQ----OH →XCOQO + H2O (R8a)

77 COQ+ X(HQ)(HO) → X(QH)---COQ----OH →XCOQ + H2Q (R8b) 78 Here X= Mg, Fe, Mn and Ca rich silicates such as enstatite, ferrosilite, rhodonite,wellostonite etc.

79 The generation of O-isotope anomaly in the CO2 via excite oxygen atoms (route 1) has 80 been extensively studied (40). On Earth oxygen isotope anomaly produced in the stratospheric

81 CO2 (R1-R3) is removed in the troposphere due to the O-isotope exchange between water and CO2

82 by the hydrosphere and biosphere (43). The interaction of anomalous CO2 generated via R1-R2 83 and its precipitation as carbonate to preserve the O-isotope anomaly would depend on the ratio of

84 CO2 /H2O reservoirs. The lack of data on the O-isotopic composition and magnitude of paleo

85 hydrosphere and atmosphere on Mars, however, does not allow us to define this ratio. During CO2- - 18 86 H2O equilibration processes HCO3 acquires O-isotopic composition of water and hence  O

87 values are dictated by the equilibration temperature. The O-triple isotope measurements on both 88 carbonate phases in ALH84001 suggests that source water from which carbonates were

89 precipitated may possess higher oxygen isotope anomaly, provided they were formed after CO2-

90 H2O equilibration processes as CO2 acquires the O-isotopic composition of water.

91 Photolysis of CO2 to yield CO and its reaction with hydroxyl radicals (R3-R5) also

92 produce O-isotope anomaly in product CO2 (44, 45). Numerous pathways have been proposed for

93 the production of hydro peroxy radicals (OHx= H2O2, OH, H O2) on Mars, such as electric 94 discharge on dust devils, photolysis of water and reaction of O(1D) with water vapor (46). Peroxy 1 95 radicals (OHx= H2O2, OH, H O2) produced by the interaction of ozone with water vapor via O( D)

96 has shown to inherit ozone isotopic signature (R3) (47). An anti-correlation of O3 with H2O at the 97 equator and summer pole of Mars suggests the role of O(1D) in the production of hydro peroxy 98 radicals (48). Recombination reaction of CO with O atoms is known to produce enrichments in

99 CO2 comparable to ozone (R5) (49, 50). The reaction of OH +CO is also known to produce mass 100 independent fractionation with remaining CO progressively enriched in 17O (51). On Earth, the 17 101 Δ O of OHx is preserved in the stratosphere but the original ozone signal is erased due to rapid 102 isotope exchange of the hydroxyl radical with tropospheric water vapor (52).

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103 The combined mass independent O-isotopic anomaly produced via pathways i) and ii) will

104 result in a steady state isotopically anomalous CO2 reservoir which upon equilibration with water 17 105 would yield CO3 with a component of the ∆ O of water. Laboratory experiments have

106 demonstrated that CO3 acquires the O-isotopic composition of H2O during CO2-H2O equilibration 107 process (53). Additionally, there would be no microscopic heterogeneity in the δ18O of carbonates 108 precipitated after equilibration with surface water reservoirs and the value of the isotopic anomaly 109 would be constant and the small δ18O value differences would simply reflect differential

110 temperature and reactivity chemistry. Pathway iii) involves interaction of CO2 with surface

111 adsorbed water on the regolith and/or dust particles in a CO2-O3-H2O-H2O2 reaction system that 112 generates microscopic heterogenity (spatial and mineral specific) due to the kinetic isotope effects 113 in the processes of adsorption and sublimation of gas-liquid layers (39, 40, 54, 55).

114 Thermodynamic equilibrium and kinetic processes such as condensation and sublimation of CO2

17 18 115 and H2O, however, fractionate O-isotopes in a mass dependent fashion with  O ~ 0.5 O (56).

116 The O-isotopic anomaly (Δ17O > 0 ‰) is only generated in processes involving interaction with

117 ozone and consequently are a measure of odd oxygen cycling (O, O3) in the atmosphere of Mars, 118 especially the ozone/water ratio. At present, the oxygen isotopic composition of Martian

119 molecular oxygen and bulk surface water mostly stored as CO2-H2O at the poles or subsurface

120 water is unknown. If the molecular oxygen 17O value is somewhere near bulk oxygen of the

121 silicate, then the observed carbonate and water values in the SNC meteorites reflect change in the

122 ozone isotopic composition and water levels. The observation of similar 17O values in both Ca-

123 rich and Fe-Mg rich carbonate phases in the present measurements reflect no significant change in

124 odd oxygen cycle (O, O3) and hydroxy radical reactions. These measurements begin to show that 125 the multi isotope approach on different carbonate phases can advance recognition of atmospheric 126 and surficial changes, allowing for full atmospheric modelling efforts in the future. 127 128

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129

130

131 Fig. S1a. Oxygen triple isotopic composition (a) O- carrying reservoirs, CO3 (ALH84001 this study), SNC CO3 and 132 silicate (16), mineral water from SNC(57). The insert shows slight offset of water O-isotopic composition from

17 18 133 terrestrial fractionation line with δ O ~ 0.52 δ O. Here red circles denote CO3, blue square = silicate, open triangle= 134 mineral water released during pyrolysis at 600oC, filled triangle = mineral water released at 1000 oC. Here 17O= 103 135 ln(1+ 17O/103) and 18O= 103 ln(1+ 18O/103). 136

137

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138 Fig. S1b. (b) Excess 17O (Δ17O) versus 18O of Martian O-carrying compounds as in Fig. 1a. 139

140 141 Fig. S2: The oxygen triple isotopic composition of carbonates in martian meteorites, closed blue symbols= Fe-Mg

o 142 rich phase in ALH84001 (CO2 released at 150 C). Open blue square= Ca-rich phase in AH84001 (CO2 released at 143 25oC after removal of terrestrial contamination (green cross). Magenta open triangle= Fe-Mg rich phase in Lafayatte, 144 magenta closed symbol = Fe-Mg rich phase in Nakhla (16). Brown open circle = Oxygen triple isotopic composition

145 of the rock in ALH84001, closed red square = Atmospheric CO2 measured by MSL (webester et al., 2013). 146 147 148 TERRESTRIAL CONTAMINATION: 149 Prolonged residence time of meteorites on ice resulted in surface contamination, possibly due to partial melting of ice 150 and seepage of water to the rock over the 13,000 yrs the rocks laid in Antarctica. To determine and compare the effect 151 of surface weathering on the O-isotopic composition of the ALH84001 Martian meteorite during its residence time in 152 Antarctic, surface crust from an igneous rock in the Dry Valley (DVC) rock was also analyzed. Water in equilibrium

153 with atmospheric CO2 produces mildly acidic conditions (pH <5) whereby causing mineral weathering and release of

154 cations to increase the pH of the solution to less acidic values (>6) and causing CO3 precipitation (Fig. S3). Carbonates 155 formed by surface weathering are enriched in both C and O-isotopes (13C = 11‰, 17O =14‰ and 18O = 28‰),

17 17 156 however, no excess O (∆ O ≈ 0) is observed. By using the measured fractionation factors for pure CO2-H2O system

157 (53) at a range of temperature (0-20oC), the equilibrium values ((13C = +3-1‰ , 18O = -20 to -23‰) is obtained

13 18 158 using isotopic composition of preindustrial CO2 ( C ~ -7‰,  O ~ 41‰) (58) and Standard Light Arctic 159 Precipitation (SLAP 18O = -55.5‰). These values are much lower than measured C and O values for the ADV

160 carbonate crust. Isotopic fractionation of DIC due to CO2 outgassing or multiple freeze thaw cycles may be the primary

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161 cause of enrichment of precipitated carbonates at low temperature. Terrestrial contamination in the CaCO3 fraction of 14 162 the carbonates based on the C activity of CO2 extracted from EET79001 has also been reported (59). Previous studies 163 of SNC meteorites have not measured O-triple isotopic composition of the calcite fraction owing to small sample 164 size(16). We have reported these values for ALH84001 for the first time after isolating surface contaminants. 165

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