the temperature of formation of carbonate in martian meteorite alh84001: constraints from cation...

PII S0016-7037(00)00528-7

The temperature of formation of carbonate in Martian meteorite ALH84001:Constraints from cation diffusion

A. J. R. KENT,1,* ,† I. D. HUTCHEON,1 F. J. RYERSON,2 and D. L. PHINNEY1

1Analytical and Nuclear Chemistry Division, Lawrence Livermore National Laboratory, Livermore CA 94550, USA2Institute of Geophysics and Planetary Physics, Lawrence Livermore National Laboratory, Livermore CA 94550, USA

(Received December8, 1999;accepted in revised form July18, 2000)

Abstract—We have measured the rates of chemical diffusion of Mg in calcite and Ca in magnesite and usedthese new data to constrain the formation temperature and thermal history of carbonates in the Martianmeteorite ALH84001. Our data have been collected at lower temperatures than in previous studies and provideimproved constraints on carbonate formation during relatively low-temperature processes (#400°C). Mea-sured logD0 values for chemical diffusion of Mg in calcite and Ca in magnesite are216.06 1.1 and27.864.3 m2/s and the activation energies (EA) are 766 16 and 2146 60 kJ/mol, respectively. Measured diffusionrates of Mg in calcite at temperatures between 400 and 550°C are substantially faster than expected fromextrapolation of existing higher-temperature data, suggesting that different mechanisms may govern diffusionof Mg at temperatures above and below;550°C.

We have used these data to constrain thermal histories which will allow the;1 mm variations in Ca-Mgcomposition in ALH84001 carbonates to survive homogenization by diffusion. Our results are generallyconsistent with models for formation of carbonates in ALH84001 at low temperatures. For initial cooling ratesof between;1021 and 103°/Ma our results demonstrate that carbonates formed at temperatures no higher than400°C and most probably less than 200°C. This range of cooling rates is similar to those observed within theterrestrial crust, and suggests that formation of the carbonates by igneous, metamorphic or hydrothermal (orother) processes in the Martian crust most plausibly occurred at temperatures below 200 to 400°C. Models thatsuggest ALH84001 carbonates formed during a Martian impact event are also constrained by our data. Thethermal histories of terrestrial impact structures suggest that cooling rates sufficiently rapid to allow preser-vation of the observed carbonate zoning at formation temperatures in excess of 600°C (.;107°C/Ma) occuronly within the uppermost, melt-rich portions of an impact structure. Material deeper within the impactstructure (where cooling is dominated by uplifted crustal material and where much of the formation ofhydrothermal minerals is concentrated) cools much slower, typically at rates of;102 to 103°/Ma. Carbonatesformed within this region would also only preserve;1 mm compositional zoning at formation temperaturesof less than;200 to 400°C. Copyright © 2001 Elsevier Science Ltd

1. INTRODUCTION

An important test of the hypothesis that Martian meteoriteALH84001 contains the fossil remnants of an ancient Martianbiota (McKay et al., 1996) is posed by the formation temper-ature and thermal history of the carbonate minerals associatedwith the proposed biomarkers. Terrestrial microbial life isknown to occur at temperatures between210 to 113°C (Gold-en et al., 2000 and references therein); if it can be demonstratedthat carbonates in ALH84001 formed at temperatures outsidethis range, then it is unlikely that these minerals formed inassociation with a terrestrial-like biota. However, before thetemperature of carbonate formation can to be used to assess thepossible presence of ancient Martian life in ALH84001, there isa need for robust and unequivocal estimates of that tempera-ture. Unfortunately, there is little present agreement on thetemperature of formation of carbonate minerals in ALH84001.Current estimates, based on a variety of techniques includingtextural interpretation, phase equilibrium, stable isotope geo-thermometry, paleomagnetism, and diffusion modeling range

from ;0 to .650°C (see summary in Table 1). In addition,many temperature estimates are model-dependent and rely onassumptions about chemical and isotopic equilibrium andclosed versus open-system behaviour.

In this study we have measured the rates of Mg and Cadiffusion in carbonate minerals and used these new data toconstrain the formation temperature and thermal history ofALH84001carbonates. The carbonate “rosettes” in ALH84001that host the putative biomarkers are chemically zoned andexhibit compositional heterogeneity on the submicron scale(e.g., Cooney et al., 1999; Schwandt et al., 1999). Zonation isevident as Mg-rich rims grading to Ca- and Fe-rich cores, andalso by more marked changes in Mg, Fe and Ca concentrationsover shorter length scales (e.g., Harvey and McSween Jr., 1996;McKay and Lofgren, 1997; Schwandt et al., 1999; Scott et al.,1998; Treiman, 1995; Treiman and Romanek, 1998). Shortlength-scale compositional variations occur adjacent to theMg-rich rims of the rosettes as well as in intermediate locationsand represent dramatic compositional gradients—with changesexceeding 10 mol cation % over distances of;1 mm (Scott etal., 1997; Scott et al., 1998). Recent high-resolution field emis-sion SEM images and micro-Raman studies confirm the indi-cations from electron probe analyses that carbonates are heter-ogeneous over very short length scales (Cooney et al., 1999;

*Author to whom correspondence should be addressed ([email protected]).†Present address:Danish Lithosphere Centre, Øster Voldgade 10, 1350Copenhagen K, Denmark.

Pergamon

Geochimica et Cosmochimica Acta, Vol. 65, No. 2, pp. 311–321, 2001Copyright © 2001 Elsevier Science LtdPrinted in the USA. All rights reserved

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311

Schwandt et al., 1999) and demonstrate that the variations incomposition measured by the electron probe are not due tosmall inclusions of magnetite or other minerals.

Given an appropriate thermal history (i.e., if temperatures arehot enough for long enough), cation diffusion will act to ho-mogenize chemical variations (Fisler and Cygan, 1998; Shenget al., 1992). Thus the observation that compositional variationsin ALH84001 carbonates occur at the;1 mm scale can be usedin combination with the rates of cation diffusion to constrainthe formation temperature and thermal history of ALH84001carbonates.

We have experimentally measured the rates of chemicaldiffusion of Ca in magnesite and Mg in calcite and the rate ofself-diffusion of Mg in magnesite. Although our approach issimilar to that used in a recent study by Fisler and Cygan(1998), one important difference is that our experimental de-sign enables us to measure much slower diffusion rates, as lowas 10225 m2/s, and to perform experiments at temperatures aslow as 400°C (compared to 550°C in previous studies; Fislerand Cygan, 1998; 1999). This approach reduces errors associ-ated with the down-temperature extrapolation required for cal-culations in the 100 to 400°C temperature interval and isparticularly important for calculations involving Mg diffusionin carbonates, as this is unexpectedly fast at temperatures below;550°C. Overall, our results provide a substantial refinement

of the temperature constraints that can be placed on the forma-tion of carbonates in ALH84001.

We also note that although ALH84001 carbonates also con-tain a substantial amount of Fe, experimental difficulties (pre-dominantly the instability of the Fe-rich carbonates siderite andankerite under our experimental conditions) prevented us fromobtaining data on cation diffusion rates in Fe-rich carbonates.These data are a future goal for our experimental program.

To avoid confusion, we have applied the following usage inour discussion of diffusion in carbonates (cf. Watson, 1994):Chemical diffusionrefers to that which occurs in the presenceof a chemical gradient, whereasself-diffusion is that whichoccurs without a chemical gradient.

2. EXPERIMENTAL METHOD

The chemical diffusion rates of Ca and Mg in carbonates weremeasured using a simple powder source technique similar to thatdescribed by Cherniak (1997; 1998). Selected cleavage fragments ofnatural calcite (CH-1 from Chihuahua, Mexico) and magnesite (fromBrumado, Brazil) were placed in platinum capsules together with driedanalytical-grade powder of the appropriate composition (CaCO3 formagnesite, MgO and MgCO3 for calcite). The chemical compositionsof magnesite and calcite used for our experiments are given in Table 2.All diffusion measurements were made using natural (1011̄) cleavagesurfaces of calcite and magnetite. Carbonate cleavage fragments werenot preannealed before experiments, and although no specific effortwas made to determine if crystallographic orientation has an observable

Table 1. Estimates of the formation temperatures of carbonates in ALH84001.

Study and technique Estimated formation temperature Mechanism of carbonate formation

Romanek et al. (1994)O and C isotope measurement of bulkcarbonate sample

0–80°C Open system precipitation of carbonatefrom aqueous fluids duringhydrothermal circulation

Mittlefehldt (1994)Chemical composition of carbonates andequilibrium thermodynamic arguments

700°C Carbonates formed during hydrothermalcirculation

Harvey and McSween (1996)Chemical composition of carbonates andequilibrium thermodynamic arguments

;650°C Impact-related metasomatism of ultramaficrocks by CO2-rich fluids

Hutchins and Jakosky (1997)Re-evaulation of data fromRomaneck et al. (1994)

40–250°C

Kirschvink et al. (1997)paleomagnetic studies

,325°C

Scott et al. (1997); Scott et al. (1998)Petrographic observations

High temperature (,;600°C?) Crystallization of carbonate fromshock-melted material

Valley et al. (1997)Ion-probe O and C isotope measurementsand mineralogical arguments

,300°C Low-temperature non-equilibriumcrystallization

Saxton et al. (1998)Ion-probe O isotope measurements

Initially ,;400°C, final stagesof precipitation, 150°C

Precipitation from hydrothermal fluids attemperatures,;400°C, the final stagesof carbonate deposition may haveoccurred at,;150°C

Leshin et al. (1998)Ion-probe O isotope measurements

1. ;125°C with fluctuations up to250°C (open system)

2. .500°C (closed system)

1. Low-temperature precipitation fromopen system with variable temperatures

2. High-temperature closed systemprecipitation from limited amount ofCO2-rich fluid

Fisler and Cygan (1998)Modelling of cation diffusion

,300°C

Golden et al. (2000)Experimental preparation of analogouscarbonate globules

;150°C Precipitation from Mg-rich supersaturatedsolutions

This StudyModelling of cation diffusion

,200°C over the range ofterrestrial crustal cooling rates

312 A. J. R. Kent et al.

effect on Mg and Ca diffusion rates, other studies suggest that there areno detectable crystallographic orientation effects in calcite for C and Oself-diffusion between 500 to 800°C (Kronenberg et al., 1984) and Caself-diffusion between 700 to 900°C (Farver and Yund, 1996). Forcalcite1 MgO experiments a small amount of CaCO3 powder was alsoadded to the capsule to enhance stability of the calcite surface (Cher-niak, 1997; Cherniak, 1998). Packed capsules were dried in a vacuumoven at 60°C overnight and sealed in evacuated silica glass tubes beforebeing annealed in ThermolyneTM muffle furnaces for periods rangingfrom 24 h to 150 d. Type-K chrome/alumel thermocouples were usedto monitor temperatures throughout the experiments, and the temper-ature uncertainty is estimated at less than63°C.

After annealing, samples were quenched by removal of the silicatubes from the furnace. After cooling, carbonate fragments were re-moved from the silica tubes and the platinum capsules and rinsedseparately in both distilled H2O and in ethanol in an ultrasonic bath for5 to 10 min to remove any adhering powdered source material. An-nealed minerals were mounted with double-sided carbon tape on glassdiscs and gold coated. Mineral surfaces were first examined for anyevidence of breakdown or residual source material using the SEM andoptical microscope and then analyzed with the ion microprobe. Withina given experimental system the upper temperature bounds mark thelimit of stability of the carbonate cleavage faces (see below). The lowertemperature bounds in CaCO3 1 magnesite and MgO1 calcite exper-iments represent the minimum temperatures at which measurable dif-fusion gradients could be generated under convenient laboratory timescales.

We report data from a total of 39 diffusion experiments. Experimentswere annealed at temperatures of 400, 450 and 500°C for magnesite1CaCO3 powder and 400, 450, 500, 550, and 600°C for calcite1 MgOpowder. To examine the possible effects of differences in diffusantcomposition, experiments were also run at 400, 450 and 500°C usingMgCO3 powder as the diffusant source for Mg chemical diffusion incalcite and at 450°C using26MgO powder to measure the rate of Mgself-diffusion in magnesite. In addition, as a test of the veracity of ourexperimental technique, both time series, where experiments are runover a range of times for a given temperature, and zero time experi-ments for MgO1 calcite and CaCO3 1 magnesite were also per-formed. For the zero time experiments diffusion couples were broughtup to the experimental temperature for;5 min and then quenched.

Diffusion profiles were analyzed with a modified Cameca ims-3f ionmicroprobe, using a16O2 primary ion beam with 17 keV impact energyand ;1 to 10 nA current. To provide adequate depth resolution andensure uniform sputter rates, the primary ion beam was rastered over a1003 100mm area. To avoid crater-edge effects an aperture, insertedin the sample image plane, allowed only secondary ions originatingfrom an;30 mm diameter region in the center of the rastered area toenter the mass spectrometer. The mass spectrometer was tuned toaccept only high energy secondary ions by offsetting the acceleratingvoltage by280 V relative to the voltage at which the intensity of16O1

dropped to 10% of its maximum value on the low energy side of theenergy distribution. The width of the energy slit was adjusted to acceptsecondary ions over a;40 eV bandpass. The energy distribution of16O1 was re-measured every 5 to 10 cycles to compensate for samplecharging during an analysis; the magnitude of charging usually rangedbetween;2 to 5 V. A typical analysis consisted of 200 to 1000individual scans over the masses25Mg1, 26Mg1, 40Ca1, 42Ca1, 44Ca1

(with count times of 1–2 s per mass) using low mass resolving power(m/Dm of ;500). The measured ion intensities were corrected forcounting system deadtime, taking into account the instantaneous countrate.

Analysis times were converted to progressive depth from the mineralsurface via measurement of the depth of the final rastered crater with aDektakTM stylus profilometer and assuming a constant sputter rate. Theoverall accuracy of this measurement varied with surface roughnessand crater depth but is estimated to be better than65 to 10% relative;the uncertainty in crater depth is not a significant contributor to theuncertainty in the calculated diffusion coefficients.

A slightly modified analytical protocol to that described above wasused for the analysis of the26Mg isotopic tracer experiments. Toresolve the isobaric interference of24MgH1 on 25Mg1 and to obtainbetter precision of Mg isotopic ratios, samples were analysed using amass-resolving power of;2600 and no energy filtering, i.e., secondaryions energies of 156 40V.

Diffusion coefficients for each experiment were calculated by per-forming a least-squares regression using the normalized25Mg, 26Mg or42Ca depth profiles to the solution of Fick’s diffusion equation fordiffusion from a semi-infinite medium (1)

~Cx 2 C1!

~C0 2 C1!5 erfS x

2ÎDtD , (1)

whereCx is the concentration at depthx, Co is the initial concentrationin the calcite,C1 is the concentration at the crystal surface,D is thediffusion coefficient andt is the duration of the anneal. For theMg-in-calcite experiments25Mg1/42Ca1 and 26Mg1/42Ca1profileswere inverted through Eqn. 1 and then fitted, while for the Ca-in-magnesite experiments the40Ca1/25Mg1 profile was used. For the Mgself-diffusion experiments in magnesite the26Mg1/24Mg1 profile wasused. The formal solution to Eqn. 1 requires the sample surface to bewell defined in terms of both the concentration of the diffusant andposition relative to the measured depth profile. In practice the surfaceconcentration of the diffusant is approached asymptotically, making itdifficult to determine accurately. Small amounts of adhering diffusantsource material can also give erroneously high surface concentrations.To minimize these problems, the surface concentration was estimatediteratively by adjusting the value to minimize the reduced chi-squaredstatistic for the least-squares fit; typically, the chi-squared valuesranged from 1 to 4. A more extensive discussion of this protocol isgiven by Ryerson and McKeegan (1994).

The surfaces of all mineral fragments used for diffusion experimentswere examined after annealing and ultrasonic cleaning by opticaland/or scanning electron microscopy for evidence of dissolution ordegradation. Alteration of the mineral surface was apparent as small,,1 to 5 mm diameter, regularly spaced pits in the cleavage surface.Based on these observations, we established upper limits to the tem-peratures at which the host mineral remained stable in the presence ofthe diffusant source. The limit of surface stability for calcite1 MgOexperiments was;600°C, for magnesite1 CaCO3 experiments;550°C, and for magnesite1 26MgO experiments;500°C. At thetemperatures for which surface alteration was apparent in a givenexperimental system, the degree of alteration appeared to be an ap-proximate function of anneal time, with short-duration experimentsshowing little or no sign of alteration, whereas longer duration exper-iments produced significantly altered mineral surfaces. For one cal-cite 1 MgO experiment annealed for 26.4 h at 600°C, several parts ofthe host crystal appeared to have experienced no detectable alteration,although elsewhere on the same crystal fragment alteration was readilyvisible. Data from the analysis of one of these “unaltered” regions isgiven in Table 3 and used herein under the proviso that this experimentmay have been affected by incipient surface alteration.

Table 2. Compositions of carbonate minerals used for diffusionexperiments.

ChihuahuaCalciteCH-1 61s

BrumadoMagnesite

BR-1 61s

MgO 0.20 0.01 41.55 0.42CaO 57.52 0.46 0.30 0.02MnO 0.19 0.02 0.19 0.02FeO 0.10 0.03 0.19 0.02CO2 42.00 0.49 57.77 0.43Number of

analyses21 27

Compositions given in oxide wt.%. Carbonate compositions mea-sured with a JEOL-733 electron microprobe, using a 10 nA, 20mmdiameter electron beam and an accelerating voltage of 15 kV. Counttimes for all elements were 60 s. Natural carbonates were used forelemental standards. CO2 contents were calculated by difference as-suming an oxide total of 100%.

313Formation temperature of ALH84001 carbonate

3. RESULTS

Data from two typical SIMS depth profiles are shown inFigure 1. The results of the diffusion experiments are given inTable 3 together with the corresponding experimental condi-tions and are shown graphically in Figures 2 and 3. The depthof the analysis craters varied between;100 and 5000 nm.Replicate analyses show that the reproducibility of a measureddiffusion coefficient for a given experiment is6 ; 0.2 logunits (Table 3). However, the overall reproducibility for allexperiments conducted at a given temperature is generally6; 0.4 log units (1s; Table 3). This uncertainty is taken to berepresentative of all the individual diffusion measurements, andis similar to the variability reported in other experimental

studies of diffusion in calcite (e.g., (Cherniak, 1997; 1998).One exception is the measured diffusion rate of Ca-in-magne-site at 450°C, which has a higher standard deviation of 0.82 logunits (Table 3). The reason for the higher standard deviation forthese particular experiments is uncertain but may be partly dueto shorter (and thus more difficult to measure) diffusion profilesin Ca diffusion experiments.

The results of the zero time experiments are also shown inFigure 1. In this figure the measured zero time profiles for the25Mg1/42Ca1 and 40Ca1/25Mg1 ratios are compared to thediffusion gradients from the shortest duration and lowest tem-perature Mg and Ca diffusion experiments for which we reportdata (i.e., the experiments with the shortest diffusion profiles).

Table 3. Data for the diffusion of Mg and Ca in calcite and magnesite.

ExperimentaTemperature

(°C)Time

(hours)D

(m2/sec)log D

(Ave. 6 1s)c

Mg diffusion in calciteA17 400 1440 6.803 10222 221.756 0.39A2 400 2256 1.433 10222

A15 400 3648 1.003 10222

E4b 400 1608 1.043 10222

A11 450 744 1.103 10221 221.516 0.45M47A 450 1080 8.063 10222

M47B 450 1080 3.263 10222

A12 450 1440 1.103 10222

A18 450 2256 8.573 10223

A16 450 3648 1.043 10222

E1b 450 1608 8.813 10222

A5 500 240 8.143 10222 221.016 0.21A13 (i) 500 240 2.073 10221

A13 (ii) 500 240 1.043 10221

A4 (i) 500 576 3.103 10222

A4 (ii) 500 576 5.903 10222

M13 (i) 500 600 8.973 10222

M13 (ii) 500 600 8.623 10222

A8 500 1632 4.503 10222

M22b 500 600 1.483 10221

A19 555 362 2.073 10221 220.926 0.03M48 (i) 555 384 1.173 10221

M48 (ii) 555 384 1.053 10221

M49 (i) 555 504 1.003 10221

M49 (ii) 555 504 7.703 10222

M14 600 26.4 1.973 10220

26Mg diffusion in magnesiteC2 (i) 450 1104 7.303 10223 221.966 0.25C2 (ii) 450 1104 1.673 10222

Ca diffusion in magnesiteB16 400 2256 1.493 10225 224.596 0.33B5 400 3648 4.403 10225

B8 450 720 7.543 10224 223.066 0.82B3 450 1440 5.003 10225

B4 (i) 450 2256 3.503 10223

B4 (ii) 450 2256 6.603 10223

B1 450 3648 5.903 10224

B6 500 240 1.583 10223 222.436 0.43B7 500 816 1.093 10222

B9 500 1632 3.833 10223

a Small Roman numerals denote repeat analyses of the same experiment.b Experiment performed with MgCO3 powder and calcite.c Averages for Mg-diffusion include data from MgCO3 and MgO-source experiments.


Although these results show that detectable amounts of residualdiffusant material do remain on the sample surfaces after clean-ing (as shown by the slightly elevated25Mg1/42Ca1 and40Ca1/25Mg1 ratios at the shallowest depths in zero timeprofiles), the degree to which the25Mg1/42Ca1 and 40Ca1/25Mg1 ratios are elevated is small compared to the size of themeasured diffusion profiles (Fig. 1). We are thus confident ourdiffusion measurements are not significantly effected by resid-ual diffusant material present on the sample surface duringanalysis. The zero time experiments also allow us to estimatethe smallest diffusion gradients that we can measure with ourexperimental protocol: for MgO1 calcite experiments thesmallest measurable profiles are on the order of;10 to 15 nm,whereas for CaCO3 1 magnesite profiles as small as 5 nm canbe measured (Fig. 1).

The results of time series experiments are shown in Figure 2.

For all temperatures examined, and for anneal times that varyby up to a factor of;5 for the 400, 450 and 500°C experimentsand ;1.5 for the 550°C experiments, there is no consistent

Fig. 1. Results from Mg and Ca-diffusion experiments. A. Measureddiffusion gradient for chemical diffusion of Mg in calcite for experi-ment A17, annealed at 400°C for 1440 h (60 d). Triangles representindividual ion-probe measurements of the25Mg1/42Ca1 ratio as afunction of depth. The thick solid line shows the error function fitted tothese data and corresponds to a diffusion coefficient of 6.803 10222

m2/s. Conversion of analysis time to depth was done by measuring thedepth of the ion probe crater. To emphasize the quality of the errorfunction fitting procedure, the dashed lines show diffusion profilescalculated for a factor of two variation in the diffusion coefficient (D53.403 10222 and D5 13.603 10222 m2/s). Diamonds represent the25Mg1/42Ca1 ratios measured in a zero time experiment conducted at400°C (see text for details). B. Measured diffusion gradient for chem-ical diffusion of Ca in magnesite for experiment B16, annealed at400°C for 2256 h (94 d). Triangles represent individual ion-probemeasurements of the40Ca1/25Mg1 ratio as a function of depth. Thethick solid line shows the error function fitted to these data andcorresponds to a diffusion coefficient of 1.493 10225 m2/s. The dashedlines show diffusion profiles calculated for a factor of two variation inthe diffusion coefficient (D5 2.98 3 10225 and D5 0.75 3 10222

m2/s). Diamonds represent the40Ca1/25Mg1 ratios measured in a zerotime experiment conducted at 400°C.

Fig. 2. Time series plots for Mg and Ca diffusion experiments incalcite and magnesite. Hollow and gray squares represent data for Mgdiffusion in calcite measured using MgO and MgCO3 respectively. Thestar represent the rate of self diffusion of26Mg in magnesite. Circlesrepresent the rate of Ca diffusion in magnesite. Solid lines show theaverage log Do value for Mg diffusion in calcite and dashed lines showthe average value for Ca diffusion in magnesite. Averagesdo notinclude data for MgCO3-source diffusion experiments. Error bars onthe individual symbols represent estimated61s uncertainty. A. 400; B.450; C. 500; and D. 550°C.


change in diffusion coefficient with time for either Ca or Mgdiffusion (Fig. 2). This absence of variation suggests that thetransport of Ca and Mg is dominated by a single mechanism,lattice diffusion, over the range of time scales investigated.

Our data for Mg and Ca diffusion are plotted on an Arrheniusplot in Figure 3. Collectively, data for Ca diffusion in magne-site and Mg diffusion in calcite form separate, negatively-sloped arrays on this plot, suggesting that the rates of diffusionof Mg and Ca in carbonate minerals are substantially differentover this temperature interval. The Arrhenius parameters, acti-vation energy (EA) and preexponential factor (D0), were deter-

mined by a weighted least-squares fit to the individual mea-sured diffusivities (as opposed to the average values) at eachtemperature. The arrays for Mg-in-calcite and Ca-in-magnesitediffusion correspond to respective activation energies (EA) of76 6 16 and 2146 60 kJ/mol and logDo values of216.061.1 and27.8 6 4.3 m2/s. The quality of the fit of the data tolinear arrays on the Arrhenius plot again suggests that a singletransport mechanism dominates the chemical diffusion of Caand Mg over the examined temperature range. The Arrheniusparameters obtained for Mg-in-calcite diffusion remain essen-tially unchanged if the single data point from the diffusionexperiment at 600°C is removed (logDo 5 216.9 6 2.6 andEA5 63 6 14 kJ/mol). As discussed above, this experimentmay have been subject to incipient surface alteration.

The measured Mg diffusion rates appear independent of thecomposition of the diffusant source, as Mg diffusion ratesdetermined using MgO and MgCO3 as diffusant sources arewithin analytical uncertainty of each other at 500, 450 and400°C (Fig. 2, 3). In addition, the array defined by the threeMgCO3-source experiments corresponds to EA and Do values(117 6 36 and212.86 2.6) that are within error (1s) of theEA and Do values defined by the regression of data fromMgO-source experiments (766 16 and216.06 1.1). The rateof 26Mg self-diffusion in magnesite at 450°C (with an averagevalue of log DMg 5 221.966 0.25 cm2/s) is also within theestimated 1s analytical uncertainty of chemical diffusion ratesof Mg in calcite measured at the same temperature (Table 3,Fig. 3).

4. DISCUSSION

4.1. Diffusion Rates of Ca and Mg in Calcite andMagnesite

Our measured rates of chemical diffusion of Mg in calciteagree well with the data reported by Fisler and Cygan (1998;1999) for self-diffusion of Mg in calcite between 600 and550°C (Fig. 3A). However, at lower temperatures our diffusionrates are substantially faster (between 2–3 orders of magnitudeat 400°C) than predicted from simple down-temperature ex-trapolation of their results. Our calculated activation energy forMg diffusion in calcite between 600 to 400°C (766 14 kJ/mol)is also substantially different from the 2846 74 kJ/mol valuereported by Fisler and Cygan (1998; 1999) for Mg diffusion incalcite between 550 and 800°C.

The distinct change in the slope of the Arrhenius array forMg diffusion in calcite between our data and those of Fisler andCygan (1998; 1999) indicates that a substantial difference inthe measured activation energy exists between the two data sets(as well as significantly different rates of Mg diffusion whenthe data of Fisler and Cygan (1998; 1999) are extrapolated to400°C). This difference could be related to differences inexperimental technique (i.e., measurement of the rates of self-diffusion using isotopic tracers sputtered onto the mineral sur-face, compared to measurement of chemical diffusion from apowder source). However, given that the diffusion rates for Mgmeasured at 600 and 550°C by both studies are within analyt-ical error of each other (Fig. 3), the observed difference inmeasured activation energies may also indicate that differenttransport mechanismsdominate Mg diffusion in calcite at tem-peratures above and below;550°C. Fisler and Cygan (1998)

Fig. 3. A. Arrhenius plot showing our data for diffusion of Mg incalcite and magnesite and Ca in magnesite. Symbols are the same as forFigure 2 and are also shown in the accompanying legend. Representa-tive 61s error bars are shown at the bottom left. Data for Ca and Mgself-diffusion in calcite from Farver and Yund (1996) (“F&Y 96”) andFisler and Cygan (1998), Fisler and Cygan (1999) (“F&C 98”) are alsoshown by the labeled thin solid lines. Thin dashed line shows the leastsquares best fit for MgCO3-source Mg diffusion data, solid dashed lineshows best fit for all our Mg diffusion data collected in this study andsolid thick line shows the best fit for our Ca diffusion data. B. Com-parison of our data for Mg and Ca diffusion with data for diffusion ofREE, Sr, Pb, and O (additional data from Cherniak, 1997; 1998; Farver,1994; Kronenberg et al., 1984).


speculated that higher diffusivity of Mg in calcite at temperaturesbelow 550°C (their lowest experimental temperature) would beconsistent with a change from an intrinsic diffusion process to anextrinsic diffusion process (or processes) at lower temperatures.While our results are consistent with this suggestion, we note that,similar changes in activation energy at;550°C are not observedfor diffusion of Ca in magnesite (Fig. 3) or for several othercations (Sr, Pb, REE) or O in calcite (Cherniak, 1997; 1998;Farver, 1994).

One potential explanation for the differences in diffusivitiesand activation energies between our Mg diffusion data andthose of Fisler and Cygan (1998; 1999) is that they reflectdifferences in the composition of the diffusant source material(in this case oxide for the majority of our Mg diffusion exper-iments and for the experiments of Fisler and Cygan (1998;1999) and carbonate for our Ca diffusion measurements). Thecomposition of the source material used in diffusion experi-ments can potentially influence surface defect characteristicsand thus may affect the measured rates of diffusion. Althoughthe majority of our data for Mg diffusion in calcite is fromMgO-source experiments, we have tested the possible effect ofvariations in diffusant source composition by also conductingexperiments using MgCO3 powder as the diffusant for Mgdiffusion in calcite. Our results suggest that no significantdifference exists between Mg diffusion data derived fromMgO-source and MgCO3-source experiments (Table 2; Figs. 2,3). Experiments using MgCO3 powder as a source for Mgdiffusion have measured Mg diffusivities that are within ana-lytical uncertainty of measured MgO-source Mg diffusivities at400, 450 and 500°C (Table 2; Figs. 2, 3) and theregression ofdata from the three MgCO3 experiments produces EA and Do

values that are within analytical error of the values calculated fromthe regression of MgO-source diffusion data. In addition, thediffusivities for Mg measured at 450 and 400°C using MgCO3 asa diffusant source are substantially faster than both the measuredrates of Ca diffusion (using CaCO3 as a source) from our studyand the extrapolated diffusivities of Mg from the study of Fislerand Cygan (1998; 1999). These data suggest that the rate of Mgdiffusion is independent of the choice of oxide or carbonatepowder for sourcematerial and that the differences in activationenergy and diffusivity that we observe between Ca and Mg maybe a fundamental feature of diffusive transport in carbonateminerals.

An alternate explanation for the differences between our dataand those of Fisler and Cygan (1998; 1999) is the possibilitythat these could be due to different concentrations of minorelements in the calcites used for experiments. This possibilityrequires further investigation. Variations in the concentrationsof Mn in calcite have been observed to have a marked effect onthe rate of self diffusion of O (Kronenberg et al., 1984),although the same study found that the diffusion rate of Cappeared to be unaffected by the Mn concentration. We notethat although we have used calcite from the same locality(Chihuahua, Mexico) as Fisler and Cygan (1998; 1999) for ourexperiments, the MnO contents measured by us(0.19 6 0.02wt.%) are much greater than the;0.02 wt.% MnO reported byFisler and Cygan (1999). In addition, the concentration of FeO inthe calcite used for our experiment is also greater (0.10 comparedto 0.02 wt.%) than that reported by Fisler and Cygan (1999). Thiscould also have an effect on rates of cationdiffusion. If Mg

diffusion below 550°C is an extrinsic process governed by thedensity of Fe-based defects, diffusion would be expected to befaster at higher FeO concentrations.

The rates of chemical diffusion of Ca in magnesite deter-mined here are essentially indistinguishable from the rates ofself-diffusion of Ca in calcite indicated by the down-tempera-ture extrapolation of the diffusivities measured by Fisler andCygan (1998; 1999) (Fig. 3). The activation energy for chem-ical iffusion of Ca in magnesite between 400 to 500°C deter-mined here, 2026 51 kJ/mol, is within analytical uncertaintyof the 2716 80 kJ/mol activation energy for self-diffusion ofCa in calcite reported by Fisler and Cygan (1998; 1999). Boththe activation energy and the diffusivity of Ca measured by usand by Fisler and Cygan (1998; 1999) distinct from the corre-sponding values for self-diffusion of Ca in calcite measured attemperatures of 700 to 900°C by Farver and Yund (1996)[EA 5 382 6 37kJ/mol]. Fisler and Cygan (1998) suggestedthat this discrepancy may be the result of Ca diffusion occur-ring via a different mechanism at the relatively high tempera-tures (700–900°C) investigated by Farver and Yund (1996).

4.2. Constraints on the Thermal History of ALH84001Carbonates

Application of our data to investigate the thermal history ofALH84001 requires that the chemical diffusion rates that wehave measured in calcite and magnesite are representative ofthe diffusion rates of Ca and Mg in the carbonate minerals inALH84001. The removal of zoning and heterogeneity in theCa-Mg composition of carbonates by diffusive transport re-quires that both Ca and Mg diffuse simultaneously within thecarbonate lattice. Assuming that this is a binary diffusionprocess and that diffusion occurs with thermodynamically idealmixing behaviour, the rate at which binary diffusion occurs isrelated to the self-diffusion rates of Mg and Ca by the followingequation (Chakraborty, 1995):

DMgCa 5D*MgD*Ca

XMgD*Mg 1 XCaD*Ca, (2)

whereDMgCa is the binary diffusion coefficient,XMg, XCa arethe respective mole fractions of Ca and Mg andD*Mg, D*Ca arethe respective self-diffusion coefficients for Mg and Ca.

Inspection of Eqn. 2 shows that for calcite, whereXCa

approaches 1, the Ca-Mg binary diffusion coefficient will be' D*Mg. This approximation is confirmed by the similaritybetween the measured rates for the self-diffusion of Mg inmagnesite and chemical diffusion of Mg in calcite at 450°C(Table 3, Fig. 3). Similarly, for magnesite the Ca-Mg binarydiffusion coefficient (which for a Mg-rich carbonate is the sameas the rate of chemical diffusion of Ca) will be'D*Ca. Thus, therates of chemical diffusion of Mg in near-pure calcite and Ca innear-pure magnesite (which are measures of the rates of Mg-Cainterdiffusion in each mineral) are approximately the same asthe rates of self-diffusion of Ca and Mg. For carbonates con-taining both Ca and Mg, the binary diffusion coefficient willhave an intermediate value, and the rate at which cation diffu-sion will remove Mg-Ca compositional differences in carbon-ates may reasonably be expected to fall within the boundscalculated from the measured rates of chemical diffusion of Cain magnesite and Mg in calcite (Fisler et al., 1998).


For ALH84001, we believe that the rate of chemical diffu-sion of Mg in calcite provides the most stringent constraint onthe formation temperature of carbonates. Mg diffusion is muchmore rapid than that of Ca and thus in Ca-rich carbonates,where the binary diffusion coefficient is dominated by the Mgdiffusivity, the most stringent temperature constraints on theformation temperature of ALH84001 carbonates will be pro-vided by Mg diffusivity. Carbonate compositions inALH84001 vary considerably. Ca-rich carbonates are known tooccur (Harvey and McSween, 1996; Scott et al., 1998) and,importantly, are associated with fine-scale compositional zon-ing and are juxtaposed with other composition carbonates onthe ;1 mm scale (e.g., Fig. 2 in Scott et al., 1997).

A simple treatment of our data shows that the diffusion rateof Mg in calcite is sufficiently fast at temperatures of 400°Cthat it is unlikely that short-length scale differences in Mg or Caconcentration could survive in Ca-rich carbonates that formedat or above this temperature. This is shown in Figure 4, wherewe plot the measured diffusion profile for Mg in calcite an-nealed at 400°C for 94 d (experiment A2). Using the diffusioncoefficient calculated from this experiment (Do 5 1.43 310222 m2/s) and Eqn. 1, we have calculated and plotted thediffusion profiles that would result from continued annealing ofthis experiment at 400°C for 10, 100 and 1000 yr. The initialstep function in Mg composition (i.e., the t5 0 Mg concen-tration profile) is quickly flattened and removed over the 1mmdistance (similar to the observed length scale of zoning ob-served in ALH84001) used for the x-axis, even over the geo-logically short durations (10–1000 yr) used for the calculation.The only assumption required in this treatment is that thechemical diffusion rate of Mg in calcite determined experimen-tally reflects the rate at which Mg-Ca compositional differenceswould be removed by diffusion at 400°C in Ca-rich carbonatein ALH84001. This simple analysis suggests that carbonates inALH84001 have not experienced temperatures of 400°C orgreater for periods of time exceeding 100 yr. We conclude thatif the ALH84001 carbonates formed at temperatures greater

than 400°C, they must have cooled extremely rapidly to pre-serve the observed variations in Mg concentration.

To investigate further the relationship between formationtemperature, cooling rate and survival of short length-scalechemical zonation in ALH84001 carbonates over a more real-istic range of thermal histories, we consider the case of mono-tonic cooling from a maximum initial temperature (Ganguly etal., 1994; Kaiser and Wasserburg, 1983; Sheng et al., 1992).From the formulation of (Sheng et al., 1992):

x2 >RT0

2D~T0!

r 0EA, (3)

wherex is the minimum distance over which carbonate willpreserve Mg-Ca compositional differences,To is the initialformation temperature,D(To) is the diffusivity atTo, ro is theinitial cooling rate atTo, EA is the activation energy andR is thegas constant. Application of Eqn. 3 requires that the termE/RTo

is .. 1 (for Mg diffusion in calcite withT0 5 600°C this termis ;10) and that the cooling history following formation attemperatureTo is linear in 1/T. For conductive cooling follow-ing an initial formation temperature or decay of a transientthermal spike, the latter assumption will be approximately true(Dodson, 1973).

The results of our calculations are shown in Figure 5, wherewe plot the length scale over which Ca-Mg heterogeneitywould survive as a function of the cooling rate and temperatureat the time of formation of the host mineral. For calculationsbased on the diffusion of Mg in calcite we have plotted threecurves, representing formation temperatures of 200, 300 and600°C. For comparison, we have also plotted two curves usingthe Mg diffusion data of Fisler and Cygan, (1998; 1999) andformation temperatures of 450 and 600°C. For calculationsbased on the rate of Ca diffusion in magnesite we have plotteda curve representing a formation temperature of 400°C. As

Fig. 4. Measured diffusion profile for chemical diffusion of Mg incalcite, annealed at 400°C for 94 d (experiment A2). Squares mark theindividual ion-probe measurements of25Mg1/42Ca1 for this experi-ment and the solid line shows the error function fitted to these data.Dashed lines show the calculated diffusion profiles that would resultfrom heating durations of 10, 100, 1000 yr.

Fig. 5. Plot of cooling rate versus the length scale of compositionalzoning (that would survive diffusional homogenization) for a range ofcarbonate formation temperatures, calculated following the approach ofSheng et al. (1992). Thick solid lines represent formation temperaturesof 600, 300 and 200°C, calculated using our data for Mg diffusion incalcite. The thin solid line represents an initial formation temperature of400°C, calculated using our data for Ca diffusion in magnesite. Dashedlines represent formation temperatures of 450 and 600°C, calculatedfrom the Mg self-diffusion data of Fisler and Cygan (1998; 1999). Thearea below each line represents the conditions under which carbonatecompositional heterogeneity would survive homogenization from dif-fusion during cooling from the given formation temperature.


discussed above, the rate of Mg diffusion in calcite controls therate at which Mg-Ca compositional variations in Ca-rich car-bonates will be removed by diffusion and, as Mg diffusion isconsiderably faster than Ca, provides the most stringent con-straint on the formation temperature. In Mg-rich carbonate therate at which Mg-Ca compositional variations can be homog-enized by diffusion is determined by the slower rate of Cadiffusion, and calculations based on Ca diffusion in magnesitewill provide an upper limit for the formation temperature. Thecurves in Figure 5 illustrate the relationship between the initialcooling rate and the characteristic diffusion distance for a largerange of potential cooling rates (10 orders of magnitude) andform a basis for evaluating the role of different processes in theformation of ALH84001 carbonates.

In Figure 5 we have highlighted the range of cooling ratesknown from rocks that cool within the terrestrial crust (;0.1–500°C/Ma), with the slowest rates equivalent to those experi-enced by stable Archaean crustal regions (e.g., Kent, 1994) andthe fastest rates from rapidly uplifted terranes such as meta-morphic core complexes (Baldwin et al., 1993). We believe thatthis range approximates the range of cooling rates that wouldexpected to be associated with igneous, metamorphic, tectonic,hydrothermal or other processeswithin the Martian crust. Thespecial case of carbonate formation during and after an impactevent is dealt with separately below.

From Figure 5 it is apparent that the rate of Mg-Ca diffusionin Ca-rich carbonates is sufficiently rapid that preservation of;1 mm scale Mg-Ca compositional heterogeneity in carbonatesformed at temperatures of 600°C or greater would require aninitial cooling rate of at least;107°C/Ma. Lowering the for-mation temperature to;300°C still requires cooling rates con-siderably faster than those known from within the terrestrialcrust. Formation temperatures of;200°C or less are requiredto preserve the 1mm scale Mg-Ca zoning in Ca-rich carbonatesin ALH 84001 for the range of cooling rates known from withinthe terrestrial crust. For Mg-rich carbonates, calculations basedon the rate of Ca diffusion indicate that formation temperaturesof ;400°C or less are required to preserve 1mm Ca-Mgzoning. We thus conclude that if ALH84001 carbonates formedwithin the Martian crust, the carbonate formation temperaturesdid not exceed 400°C and were most plausibly below 200°C.These conclusions strongly favor models that involve low-temperature processes for carbonate formation and place tighterconstraints on the formation temperature of ALH84001 carbon-ates than previous studies (Fisler and Cygan, 1998; 1999). Wenote that the Mg diffusion data of Fisler and Cygan (1998;1999), if used for a similar analysis, only require formation ofALH84001 carbonates at temperatures below;450°C (Fig. 5).

4.2.1. Carbonates formed during an impact event

Several authors have suggested that carbonate rosettes inALH84001 formed at relatively high temperatures (;600–700°C) during heating and/or high temperature fluid circulationassociated with a Martian impact event (Harvey and McSween,1996; Scott et al., 1997; Scott et al., 1998). Although thecalculations discussed in the preceding section indicate that theformation of carbonates at temperatures of.600°C requiresinitial cooling rates in excess of 107°C/Ma (Fig. 5), cooling

rates following impact events may be substantially higher thanthose related to terrestrial tectonic processes (Crossey et al.,1994), and it is possible that fine-scale compositional zonationin carbonates formed at temperatures of$600°C during animpact event could survive. In addition, petrographic observa-tions of ALH84001 indicate that fine-scale chemical zoning incarbonateshas survived at least one impact event, as finely-zoned carbonate grains are disrupted and offset in ALH84001by impact-produced, plagioclase-composition melt glass(Shearer et al., 1999; Treiman, 1995; 1998). Cooling ratesfollowing any impact or thermal event must have been fastenough to preserve the fine-scale carbonate zonation.

Our data on cation diffusivities can still be used within thecontext of a monotonic cooling history to place some con-straints on models that argue for formation of ALH84001carbonates during impact events. Cooling rates (and thus thepotential survival of fine-scale chemical zonation in carbon-ates) within impact zones during and after impact vary signif-icantly depending upon the geometry of the structure andlocation within the impact structure (Crossey et al., 1994;McCarville and Crossey, 1996). Studies of terrestrial impactsites have shown that rocks within impact structures, includingthose that like ALH84001 contain diaplectic and melt glasses(and thus reached peak temperatures of up to 800°C or moreduring impact; Grieve, 1987), experienced relatively slow cool-ing after impact (Crossey et al., 1994). This slow cooling is aresult of postimpact rebound and uplift of basement rockswithin the central region of an impact site. This causes lowercrustal isotherms to move closer to the surface, and the postim-pact decay of the resulting thermal anomaly then promotesrelatively slow cooling compared to the uppermost melt-richportions of the impact structure (Crossey et al., 1994). Forexample, cooling rates within the lower portions of the Mansonimpact structure (peak temperatures of 600–1000°C) were;500 to 800°C/Ma (Crossey et al., 1994)—similar to thefastest known terrestrial cooling rates associated with rapidtectonic uplift shown in Figure 5. Only the uppermost, melt-rich parts of the Manson structure experienced substantiallyfaster cooling (; 106–107°C/Ma). Carbonate minerals formedwithin an impact structure, with the exception of those that mayhave formed within the relatively thin and rapidly cooled uppermelt sheet, would be subject to the same thermal constraintsshown in Figure 5 and are thus are unlikely to have formed attemperatures greater than 200 to 400°C.

In addition, carbonate minerals that form during an impactevent are more likely to occur within the lower, more slowly-cooled portions of an impact structure because it is this regionthat experiences the most intense hydrothermal activity (Allenet al., 1982; McCarville and Crossey, 1996). In terrestrialenvironments the formation of carbonate minerals is commonin hydrothermal circulation zones and many workers havesuggested that carbonate rosettes in ALH84001 are also theresult of hydrothermal processes (Table 1). Studies of hydro-thermal minerals and fluid inclusions from the Manson impactstructure show that hydrothermal circulation within the upliftedcentral region occurred continuously from temperatures of.300°C, down to ambient temperatures (Boer et al., 1996;McCarville and Crossey, 1996).


5. CONCLUSIONS

We have measured the rates of chemical diffusion of Ca inmagnesite and Mg in calcite and applied these new data to astudy of the formation temperature and thermal history ofcarbonates in Martian meteorite ALH84001. Measured logD0

values for chemical diffusion of Mg in calcite and Ca inmagnesite are216.0 6 1.1 and27.8 6 4.3 m2/s and theactivation energies (EA) are 766 16 and 2146 60 kJ/mol,respectively. Diffusion rates of Mg in magnesite at tempera-tures between 400 to 550°C are substantially faster than ex-pected from extrapolation of existing data, suggesting thatdifferent mechanisms may govern diffusion of Mg at temper-atures above and below;550°C.

Considerations of isothermal and slow cooling thermal his-tories for carbonates suggest that, for a range of cooling ratessimilar to those observed within the terrestrial crust,ALH84001 carbonates are likely to have formed at tempera-tures at least,400°C and most likely,200°C. This supportsmodels involving low-temperature processes for carbonate for-mation.

In addition, evaluation of the range of thermal historiesevident within impact structures suggests that carbonatesformed within the lower parts of an impact structure would alsohave cooled at rates slow enough to restrict carbonate forma-tion to temperatures below;200 to 400°C.

Acknowledgments—This work was funded by Lawrence LivermoreNational Laboratory under the Laboratory Directed Research and De-velopment (LDRD) program, and by NASA through work orderW-18845. Reviews by A. Treiman and two anonymous reviewerssignificantly improved the quality of this manuscript. We are alsograteful to Haley Ryerson for helping to choose the prettiest carbonatesamples. Work performed under the auspices of the U.S. Department ofEnergy by Lawrence Livermore National Laboratory under ContractW-7405-ENG-48.

Associate editor:H. E. Newsom

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