mineralogy of three slightly palagonitized basaltic tephra samples from the summit of mauna kea,...

JOURNAL OF GEOPHYSICAL RESEARCH; VOL. 98, NO. E2, PAGES 3401-3411, FEBRUARY 25, 1993

Mineralogy of Three Slightly Palagonitized Basaltic Tephra Samples From the Summit of Mauna Kea, Hawaii

D.C. Gom•.s, R. V. Mom•IS, • D. W. Milo

NASA Johnson Space Center, Houston, Texas

H. V. L^u•R JR., ̂ N• S. R.

Lockheed ESC, Houston, Texas

Certain palagonites from Hawaii are considered to be among the best analogs for Martian fines, based upon similar spectral properties. For this study, three distinctly colored layers were sampled from slightly palagonitized basaltic tephra just below the summit of Mauna Kea at 4145 m elevation. The mineralogy of size fractions of these samples was examined by diffuse reflectance (visible and near-IR) and far-IR spectroscopy, optical microscopy, X ray diffraction, M6ssbauer spectroscopy, magnetic analysis, electron microprobe analysis (EMPA), and transmission and scanning electron microscopy. For the 20-1000 grn size fraction, sample HWMK11 (red) is essentially completely oxidized and has a hematite (Ti-hematite) pigment dispersed throughout the silicate matrix. The alteration is present throughout particle volumes, and only a trace amount of glass is present; no palagonitic rinds were detected. In addition to ferric Fe-Ti oxides, other phases detected were plagioclase feldspar and a trace of olivine. Sample HWMK12 (black) has the lowest proportion of ferric- bearing phases and is thus least weathered. It consists mostly of unaltered glass with embedded plagioclase and minor amounts of pyroxene, olivine, and Ti-magnetite. In some grains, a thin palagonitic rind is visible, indicating some surface alteration. The mineralogy for sample HWMK13 (yellow) is the same as that for HWMK12, except that it has distinct, well-developed palagonitic rinds consisting of erionite and smectite. For all samples, the amount of glass and plagioclase decreases and the amount of smectite increases with decreasing particle size for size fractions <20 gin. For H3VMK11, the amount of hematite is essentially constant, and mica is present only in the coarse clay-sized fraction; smectites are low in structural Fe. For HWMK12 and HWMK13, the zeolite erionite is present along with smectites and nanophase ferric oxides (np-Ox). Erionite abundance decreases and np-Ox abundance increases with decreasing particle size. The smectite in both black and yellow samples contains some Fe 3+ in octahedral layers. There were only two mineral phases containing iron in the fine day fraction, namely, smectites and iron oxides. For HWMK11, relatively large iron oxide particles (0.1 to 0.4 grn) were dispersed on clay surfaces; for HWMK12 and HWMK13, much finer np-Ox particles were present in lesser concentrations. Formation of the zeolite erionite is consistent with the arid climate zone where these samples were collected. However, transient hydrothermal processes that occurred during the eruption of Mauna Kea volcano under its permanent ice cap during the Pleistocene may have resulted in minerals such as zeolites and smectites which may persist as relicts over a long period of time. Most of the iron released during weathering of basaltic tephra precipitated as poorly crystalline iron oxides and some of the Fe has substituted for the octahedral cations in the structure of authigenic smectites. The Ti-hematite in HWMK11, however, is the result of high-temperature oxidation of Ti-magnetite and exsolution from iron- bearing silicate phases. Visible and near-IR reflectivity spectra for the <20 grn size fraction of HWMK11 is dominated by well-crystalline Ti-hematite. Corresponding spectra for HWMK12 and H3VMK13, whose ferric mineralogy is dominated by np-Ox particles, are more similar to Martian bright-region spectra.

INTRODUC•ON

Physical and chemical measurements were performed on Martian surface materials during the Viking missions in the 1970s. Bulk chemical compositions of fine materials (<2 ram), magnetic properties, and other physical data were obtained at two sites on the Martian surface [e.g., Hargraves et al., 1977, 1979; Clark et al., 1982]. Spectral reflectance data from spacecraft observations and Earth-based telescopes provide some chemical and mineralogical information for the Martian surface [Burns, 1989; Singer et al., 1979; Singer, 1982]. Collectively, these data provide the physical and chemical constraints available to date for the mineralogy of Martian surface materials. Theoretical modeling of mineral stabilities and weathering mechanisms and rates under Martian surficial conditions [Gooding, 1978] may also help constrain mineralogies.

Until Martian samples are returned to Earth, terrestrial analog

Copyright 1993 by the American Geophysical Union.

Paper number 9ZIE02590. 0148-0227/93/9ZIE-02590505.00

soils will serve as an option available to scientists to infer the properties of the Martian regolith. Certain palagonitic soils from Mauna Kea, Hawaii, are good spectral and magnetic analogs of Martian surface materials [Singer et al., 1979;Adams et al., 1986; Evans and Adams, 1979; Allen et al., 1981, 1982; Morris et al., 1990]. Palagonite is a yellow or orange mineraloid formed by hydrafion and devitrification of basaltic glass [Bates and Jackson, 1984]. Throughout the world, there are many palagonites formed under different weathering conditions, out of which only some are known to exhibit Mars-like spectral properties. For example, reflectivity spectra of <20 Ixm size separates of palagonitic soil HWMK1 resemble spectra of Martian bright regions, and reflectivity spectra of its coarser separates have some spectral characteristics common with Martian dark regions [Morris et al., 1990]. Identification of minerals in palagonitic materials is important in the interpretation of genetic processes. Although some work has been done on the mineralogy of palagonites [e.g., Ugolini, 1974; Ming et al., 1988; Morris et al., 1990; Bell et al., this issue], detailed mineralogical information for Martian spectral analog palagonites is rather scarce. The objective of this study is to characterize the mineralogy and spectral properties of three

3401

3402 Got•m,• fir AL.: iVhhrm•mc•y oF• SAM•, MArinA YmA

slightly palagonitized basaltic tephra samples collected near the summit of Mauna Kea in order to contribute to the basis for

inferring mineralogy and processes for Martian surface materials.

SAMPLES AND METHODS

Three distinctly colored layers (upper red, HWMK11; middle black, HWMK12; and lower yellow, HWMK13) were sampled from basaltic tephra just below the summit of Mauna Kea at 4145 m (13,600 feet) elevation. During sample collection, a stainless- steel screen with 2-mm holes was used to exclude larger-diameter tephra particles. The cinder cone from which the samples were taken consists of ejecta of strombolian type [Porter, 1972a] and the particle size is commonly 1 to 10 cm. The present-day elimate is cool to cold and extremely dry with nocturnal freezing temperatures throughout the year [Ugolini, 1974]. Pleistocene ice caps have repeatedly covered the summit area eausing three or four episodes of glaciation. The cinder cones at the summit were penecontemporaneous with next to the last glaciation [Porter, 1972b].

Size Separates Size separates of <1 mm bulk soil were obtained by two

methods. For method 1, -200 g of bulk soil were wet sieved with freon into seven size separates using 1000, 500, 250, 150, 90, 45, and 20 grn sieves. The sieve screens were rhodium-plated nickel with square holes. No ultrasonication was used to remove loosely bound surface rinds. For method 2, ~2 g of soil were dispersed in 40 mi., of deionized water in a 50-mL polypropylene centrifuge tube and brought to pH 9.5 using 0.1 M NaOH. An ultrasonicator equipped with a microtip operated at 22 •tm amplitude was used for 5 rain to disperse the sample. The sonication is disruptive and will remove loosely bound surface rinds from particles. After sonication, the suspension was transferred to a 100-mL measuring cylinder, and the total volume was brought to 100 mL mark using pH 9.5 water. The water temperature was 22øC, and the sedimentation time for 2 Ixm size particles for a 20 cm depth was 15 hours. The top 20 cm was carefully siphoned and the process repeated until no suspended particles were visible after the dispersion (-4 dispersions). The suspension containing <2 grn fraction was flocculated by dissolving sufficient solid NaC1 in it. After flocculation of the clay, the supernatant was discarded, and the sediment was transferred to•dialysis bags and dialyzed against deionized water for a 2-day period. The salt free clay sample was quick frozen in liquid nitrogen and freeze dried. The free silt fraction (2-5 Ixm) was separated by sedimentation from the >2 Ixm residue. The 5-20 Ixm fraction was obtained by wet sieving of the >5 Ixm fraction using a 20-1xm sieve.

To obtain the free clay (<0.2 I•m) fraction, the freeze-dried clay fraction was dispersed in 40 mL of pH=9.5 water. The resulting suspension was centrifuged in an International Equipment centrifuge (head 253) operated at 3000 rpm. The time for sedimentation of >0.2 Ixm was calculated by the formula given by Jackson [1974]. The resulting clay fractions (<0.2 and 0.2-2.0 I•m) were flocculated and freeze-dried.

X Ray Diffraction (XRD)

The freeze-dried powders were mounted on cavities carved on glass slides and X rayed by a $cintag XDS 2000 X ray diffractometer using CuKa radiation; a software routine was used for KI• stripping. Untreatecl clay fractions (Na-saturated) were X rayed using oriented mounts. Mg saturation, followed by

glyceration, and K saturation followed by step-wise heating fxom 25øC to 300øC to 550øC of the oriented mounts were used as

pretreatment prior to XRD analysis [Jackson, 1974].

Electron Microscopy and Electron Probe Microanalysis (EPMA)

For scanning electron microscopy (SEM), individual tephra grains were mounted on silver tape and glued onto A1 stubs using epoxy resin. The grains were coated with Au-Pd prior to observation in a JEOL JSM 35 scanning electron microscope. For transmission electron microscopy (TEM), the fine clay size fraction was dispersed in distilled water, and a drop of the resulting suspension was evaporated on a holey carbon film. The dried sample was coated with carbon prior to examination with a JEOL JEM 100 CX scanning transmission electron microscope.

Polished petrographic thin sections of epoxy embedded particles (<1.0 ram) were carbon coated and analyzed using a Cameca Camebax microbeam electron microprobe operated at 15 kV and 10 nA beam current. National Institute of Standards and

Technology mineral standards were used for calibration.

M6ssbauer , Diffuse Reflectance (Visible and Near JR ) and FarJR Spectroscopy

M6ssbauer spectra at 293 K were obtained with Elscint and Ranger Scientific (MS-1200) M6ssbauer spectrometers. M6ssbauer parameters were calculated from fits of the spectra to theoretical line shapes using an in-house computer program (JSCFIT) which allows for the asymmetric line shapes. All spectra were fit with skewed-Lorentzian line shapes which are two half Lorentzians that have the same peak value but different widths.

Diffuse reflectivity spectra at 293 K in the visible and near-IR (0.35-2.0 gm) were obtained relative to a Halon standard using a Cary-14 spectrophotometer configured with a 9-inch diameter integrating sphere. The reflectivity relative to Halon was converted to absolute reflectivity using the data of Weldher and Hsia [1981]. Far-IR data were obtained on 2-ram-diameter KBr pellets containing 0.3 wt % clay using a BIORAD FTIR

Compositional and Magnetic Analysis Concentrations of iron as Fe were detemfined by instrumental

neutron activation analysis. Saturation magnetizations (Js) at 293 K were calculated by a linear extrapolation to zero applied field of the descending branch of the hysteresis curve measured between 1.3 and 2.1 T with a vibrating sample magnetometer (Princeton Applied Research Model 151). Water concentrations were determined using a DuPont 902 moisture evolution analyzer.

SIZE DISTRmUTION AND Mn•E•Y

Particle Size Distribution

Mass percentages of soil in each size separate (Table 1) show that all three soils are very coarse grained. The amount of mass in the <20 Ixm size fraction is 0.2 to 1.3 wt %. Ultrasonic disaggregation produced significantly more material (3.3 to 7.7%) in size fractions <20 [tin, which implies individual soil particles have easily disrupted regions. M6ssbauer and other data discussed later show that this material is highly weathered because most of the Fe is in ferric form.

XRD and Far-lR Analysis

Fine clay (<0.2 pm) fraction. XRD patterns of the Na- saturated fine clay fraction at 56% relative humidity indicated

GOLDEN • At,.: •¾ Or • SAMPLF. S, MAtrSA KEA 3403

TABLE 1. Particle Size Distribution and Selected Physieochemical Propenes (Total, Adsorbed and Bound H20, Total Iron

Concentration as Fe, and Saturation Magnetization (.Is) at 293 K) for Palagonitie Soils

HWMK11 HWMK12 HWMK13

(Red) (Black) (Yellow)

Size Fraction (ltm) and Relative Mass (%) From Wet Sieving (Method 1) 500-1000 26.8 46.5 36.2

250-500 35.7 29.9 32.9 150-250 18.6 13.4 15.2

90-15 10.2 6.2 7.3 45-90 4.8 2.7 4.5 2045 2.5 0.9 Z9

<20 1.3 0.2 1.0

Size Fraction (ton) and Relative Mass (%) From Ultrasonic Disaggregation, Wet Sieving, and Sedimentation (Method 2)

20-1000 94.50 96.69 92.31 5-20 3.22 2.03 5.86

2-5 1.04 0.48 0.67 0.2-2 0.93 0.31 0.57 <0.2 0.31 0.49 0.59

Physicochemical Data for <1 mm Size Fraction (Water-Free Basis for Fe and Js)

H20, total, % 1.0 1.8 5.7 H20, adsorbed, % 0.6 0.7 2.7 H20, bound, % 0.4 1.2 3.0

Fe, % 7.75 7.88 7.93

Js, Am2/kg of Fe 3.2 12. 4 11.2

only a single basal peak at 1.25 nm for the HWMK 11 and a peak at 1.21 nm for HWMK12 and HWMK13 (Figure 1). All three samples upon Mg saturation gave a layer spacing of 1.5 nm which upon glycerol saturation expanded further to 1.8 nm, conf'u'ming all three samples contained smectites. The K-saturated fine clay from HWMK11 sample did not collapse to 1.0 nm utxm heating to 300øC, but at 550øC it collapsed to 1.0 nm. Failure to collapse to 1.0 nm at 300øC indicates the presence of some hydroxy interlayered material in this clay. The K-saturated fine clay from HWMK12 and HWMK13 samples, however, collapsed to 1.0 nm at 300øC. Far-infrared (far-IR) spectra indicated that the OH fundamental vibration of the fine clay material in the HWMK11 sample was at 3618 cm-1 compared to the one at 3566 cm-1 for the fine clay from HWMK12 and HWMK13 samples (Figure 2). The (A1,A1)OH and (A1,Fe)OH dioctahedral OH stretching vibrations generally occur in the range of 3620 to 3630 cm -1 [Fanning et al., 1989]. Due to the low abundance of octahedral Fe in the sample (see later), the peak near 3620 cm- 1 in HWMK11 sample is assigned to (Ai,Ai)OH stretching (Figure 2). The presence of an absorption band at 917 cm-1 resulting from (AI,A1)OH and the absence of an absorption band at 879 cm-1 due to (Ai, Fe)OH indicates the presence of mostly Al in a dioctahedral layer. HWMK12 and HWMK13 samples both had a distinct OH- bending band at 879 crn-1 (Figure 2) resulting from an OH group coordinated to A1 and Fe atoms [Beutelspacher and van der Marel, 1968], which implies appreciable Fe substitution in octahedral layers of the smectite in these samples. Unit cell formulae for the smectites in all three samples were derived from the EMPA data, assuming <0.2 }xm fraction was pure smectite. Fe oxides were removed prior to analysis by the procedure of Jackson [1974]. The unit cell formulae were (Si7.39A!0.61)(Alzn3Mg1.23Fe0.58)O20(OH)4X1.11 for HWMK11,

1.2 nm XRD Analysis (<0.2 um)

HWMK13

HWMK12

HWMK11

44.1 7.37 4•04 2.79 2 15

d (A)

Fig. 1. XRD patterns of oriented fine-clay (<0.2 gm) fractions (Na- saturated) from Hawaiian basaltic tephra samples HWMK11, HWMK12, and HWMK13. The principal clay mineral in the fine clay fraction of all three soil samples is smectite.

3.5

3.0

3566

2.0

1.5

lO

IR SI3ectra (< 0.2 pm)

HWMK13

HWMK12

0.5

0.0

3618

HWMK 1 I

879cm 1

3500 3000 2500 2000 ! 5•30 ! 0 O0 500 WavenumOers (cm '1 )

Fig. 2. IR absorption spectra of fme clay fractions of red (HWMK11), black (HWMK12), and yellow (HWMK13) samples. An absorption band at 879 em '1 in the black and yellow samples suggests Fe substitution in the octahedral layer of smecfite. Note the OH stretching fundamental bands for HWMK12 and HWMK13 are similar but different from that of HWMK11.

(Si7.49A!0.51)(Alz06Mg0.77Fe1.27)O20(OH)4X0.85 for HWMK12, and (S i7.62AI0.37)(All.81Mg0.99Fe1.30)O20(OH)4X •. 0 6 for HWMK13, where X represents a monovalent exchangeable cation. The latter two samples were rich in octahedral Fe3+ in line with the IR data. They were all dioctahedral smectites, but none of them qualify to be nontronite where the predominant octahedral cation should .be Fe3+ and the charge should originate mainly in the tetrahedral layer.

Coarse clay (02-2 I•m), fine silt (2-5 lzm), and medium silt (5-20 lzm) pactions. The coarse clay fractions contained smectite in all three samples. However, erionite (a zeolite) and a trace of feldspar (plagioclase) was observed in HWMK12 and HWMK13 samples (Figures 3 and 4). Plagioclase, hematite, and some mica were observed in the coarse clay fraction from the HWMK11 sample (Figure 3). In coarser size fractions (>5 gm), more primary minerals were present (e.g., plagioclase) and the contribution of phyllosilicates to total mass diminished (Table 1).

EMPA and Petrographic Observations of Thin Section Backscattered electron (BSE) images of the thin sections of all

three bulk samples (Figure 5) indicated that the surface material

3404 GOLDE•T m' AL.: •Y OF• S•Mpt•, M•tr•A I•A

e

i XRD Analysis (0.2-2 iJm)

J • e p iPY py I

•'4 HWMK 12

i i h h HWMK1 1

44.1 7.37 4.04 2.79 2.15

d(•) Fig. 3. XRD pattems of oriented sample mounts of Na-saturated, coarse- clay fractions (Cu-K• radiation) indicating the presence of smectite (s), and erionite (e) in the yellow HWMK13, and black (HWMK12) samples, hematite Oa) in the red sample HWMK11, and plagioclase feldspar (p) and pyroxene (py) in all three samples. A trace of mica (m) and smectite were in sample HWMK11.

XRD Analysis

e (5-20•Jm) j P HWMK 13

I i e J p py J

s e e Pep o e P• h h t e Y •j•,.,•l••,•,•,.,• 1 J s HWMK 12

, ij h j HWMK11

, Iv , h h h 44.1 7.37 4.04 2.79 2.15 1.76

Fig. 4. XRD patterns of powder sample mounts of coarse silt (5-20 fractions indicate the presence of plagioclase (,p) in al.l samples, crionite (c) in black sample HWMK12 and in yellow sample HWMK13, smectite (s) in yellow and black samples, and hematite Oa) in the red sample HWMK11. Pyroxene (py) and olivine (o) were present in all samples. x is an unassigned peak.

Fig. 5. Backscattered electron (BSE) images of bulk (<lmm) tephra samples: (a) red sample CHWMK11) without a distinct rind (r); (b) black sample (HWMK12) with a slightly developed rind; (c) yellow sample (HWMK13) with a well-developed palagonitic rind; (d) BSE of a thin section of HWMK12 sample heated to 600øC for 4 hours in the air. Note the similarity of Figures 5a and 5d which have numerous small hematite particles Oa) dispersed in the glass matrix. Dispersed in the glass matrix of black and yellow samples are microphenocrysts of plagioclase (p) and olivine (o). Alteration rinds consisted of mainly Fe-rich smectites and np- oxides.

GOLDEN Er AL.: MINERALOGY OF TF. PHRA SAMre23, MAUNA KIiA 3405

had been'altered to various degrees. Sample HWMKll when observed under the petrographic microscope had red-colored iron pigment dispersed throughout the silicate matrix along with plagioclase microphenocrysts. No rind was visible around the particles of HWMK11. The BSE image of the HWMK11 sample clearly indicated the distribution of Fe and Fe-Ti oxides within the altered glass matrix in concurrence with the optical observations. Concentrations of Fe and Ti were measured for 18 and 12 oxide

grains (5-50 •m in size) in samples HWMK11 and HWMK12, respectively. As shown in Figure 6, the range of Ti concentrations in the oxides of HWMK11 is 4.2 to 24.6 wt % (0.07 to 0.4 mole fraction). The larger Fe-oxide grains were rich in Ti and the finer grains distributed in the glass matrix were poor in Ti. XRD data discussed above and M6ssbauer and reflectivity data discussed below indicate that the mineralogy of the oxides in HWMK11 is Ti-hematite. Ti-magnetite grains from sample HWMK12 had a narrow compositional range (10.5 to 14.5 wt % Ti). Ti-hematite compositions in HWMK11 are spread above and below this range which indicates that they formed by oxidation and phase separation from Ti-magnetites equivalent to those found in HWMK12. In addition, low-Ti hematites in HWMK11 result

from exsolution from silicate phases during oxidation. Ti/Fe weight ratio of 0.43 corresponds to either ulv6spinel or pseudobrookite, and a Ti/Fe weight ratio of 0 corresponds to either pure hematite or magnetite (Figure 6).

Samples HWMK12 and HWMK13 both had palagordtic rinds resulting from the aqueous alteration of particle surfaces (Figures 5b and 5c). Sample HWMK13 had a thicker rind than sample HWMK12 so that the former has been more altered than the latter.

The glass matrix in both of these samples contain large grains of magnetite (approximately 1-50 IJJn) similar in appearance to the large hematite particles in HWMKll. The glass matrix in HWMK12 and HWMK13, however, is free from fine hematite particles. The fine hematite particles in HWMKll must have separated from iron-bearing silicate phases during a high- temperature thermal event during or subsequent to their ejection from the volcano. This hypothesis was tested by heating sample HWMK12 (the least altered sample) to 600øC in air for 4 hours. The BSE image of heated HWMK12 (Figure 5d) is similar to the BSE of HWMK11 (Figure 5a) in that small hematite crystals have formed from the glass matrix of the HWMK12 sample, supporting the above hypothesis. If the temperature is high enough (e.g., 800øC), Fe 2+ containing pyroxenes can also yield magnetite and hematite as oxidation products [Phillips et al., 1991]. Evidence for pyroxene alteration can be seen in some of the particles in HWMKll.

M•ssbauer Spectroscopy, Magnetic Data, and Diffuse Reflectance (Visible, and Near-lR) Spectroscopy

The interpretations of XRD and EMPA data are supported by M6ssbauer, magnetic, and reflectivity data. M6ssbauer spectra for 500-1000 and <20 gm size fractions are shown in Figure 7; derived M6ssbauer parmeters and mineralogical assignments are compiled in Table 2. The doublet having IS -0.35 mm/s and QS -0.66 mm/s results from octahedrally coordinated ferric iron and is apparently a characteristic of palagonitic samples. The M6ssbauer parameters are consistent with those for small-particle iron oxides/oxyhydroxides like ferrihydrite [e.g., Murad arm Johnson, 1987], nanophase (superpararnagnetic) hematite [e.g., Kundig et al., 1966; Morris et al., 1989], and other superparamagnetic ferric oxides [e.g., Murad and Johnson, 1987; Murad, 1989]. Because we cannot identify the mineralogy of the small oxide particles and because we know from TEM data discussed below that small (<20 nm) iron oxide particles are

25

o• 16-

- 13-

I- - _

10- _

_

7- _

_

4 0

HWMK11 and HWMK12

ß

ß '•, 14 HWMK12

12 , ,

11 ,,'

10 0.20 .... o.5 .... o.o

012 ' 014 ' 016 Ti/Fe (weight ratio)

Fig. 6. Plot of weight percentage Ti versus Ti/Fe weight ratio for Ti-Fe oxide grains contained within samples HWMKll (Ti-hematites) and HWMK12 (Ti-magnetites). Inset shows data for Ti-magnetites in HWMK12.

500-1000 um <20 um

WMK12 (Black)

r•w•HWMK13 (Yellow)

1'•''''''' '•''' '•'" '1'0 •(•''' •''' '•)' ' ' ' ' ' ' ' '1'0 - -5 - - 5

Velocity (mm/s)

Fig. 7. M6ssbauer spectra (293 K) for 500-1000 and <20 g.m size fractions (method 1) of HWMK11 (red), WMK12 (black), and WMK13 (yellow). Mbssbauer parameters and mineralogies are summarized in Table 2. The Mbssbauer spectra of both size fractions of HWMK11 and the <20 gm fractions of HWMK12 and HWMK13 are dominated by ferric mineralogies. For <20 Itm size fractions, well-crystalline hematite dominates the spectral area relative to np-Ox for HWMK11 and vice versa for HWMK12 and HWMK13.

3406 GoLt)•N Lzr AL.: •Y ov 21UnmA SAMPtJ•, MArinA I•A

TABLE 2. M6ssbauer Parameters (Isomer Shift (IS), Quadrupole Splitting (QS), Hyperfine Field (Bhf), And Relative Area (A)) at 293 K

Sample Fraction, gm IS, mm/s QS, mm/s Bhf, T A, % Mineralogy

HWMK 11 500-1000

<20

HWMK12 500-1000

<20

HWMK13 500-1000

<20

1.14 2.90 4 Olivine

0.38 0.66 26 rip-Oxide 0.38 -O.21 51.6 70 Hematite

0.38 0.67 33 np-Oxide 0.38 -0.21 51.6 67 Hematite

[0.74/0.26] [-0.08/0.06] [46.2/48.6] 8 Magnetite [1.07] [0.67] 8 l]menite 1.14 2.94 20 Olivine

[0.35] [0.65] 29 np-Oxide 1.04 2.10 35 Glass/Pyroxene

1.14 2.91 4 Olivine

1.13 1.97 8 Glass/Pyroxene 0.38 -0.22 51.8 18 Hematite

0.35 0.65 71 np-Oxide

[1.07] [0.67] 4 I1menite [0.74/0.26] [-0.08/0.06] [46.2/48.6] 7 Magnetite

1.13 2.93 19 Olivine

1.07 1.93 25 Glass/Pyroxene 0.32 0.80 45 np-Oxide

[0.74/0.26] [-0.08/0.06] [46.2/48.6] 5 Magnetite 1.15 2.95 7 Olivine

1.11 1.92 7 Glas s/Pyroxene 0.35 0.64 81 np-Oxide

Size fractions refer to method 1. Numbers in brackets are constrained parameters. Components are listed in order of increasing area.

present, we use the generic phrase "nanophase ferric oxide (np- Ox)" to refer to iron oxide particles having nanoscale (<20-50 nm) parfide dimensions. Ferrihydrite, superparamagnetie (at room temperature) parfides of hematite and goethite, and nanometer- sized particles of inherently paramagnefie (at room temperature) lepidoeroeite are all nanophase oxides that could contribute to the ferric doublet.

For sample HWMKll, M6ssbauer data show that Fe exists predominantly as ferric iron. There is only a small ferrous doublet probably resulting from olivine in coarse size fractions. M6ssbauer parameters for the sextet, which dominates the spectra of HWMKll, are consistent with those for well-crystalline hematite; the skew in the peak shapes toward zero velocity may result from Ti substitution as observed in the EMPA data. The

percentage area resulting from Ti-hematite is 2-3 times larger than that from the np-Ox doublet.

M6ssbauer spectra for HWMK12 and HWMK13 are very different from HWMK11, but they are similar to those reported by Morris et al. [1990] for another Hawaiian palagonitic soil. The 500-1000 !xm fractions are dominated by ferrous iron with, respectively, 71 and 55% of the total Mfssbauer peak area contributed by ferrous-bearing phases (magnetite, ilmenite, pyroxene, olivine, and glass). The ferric doublet (np-Ox) is the balance of the peak area. In contrast, M6ssbauer data for <20 }an fractions (method 1) show that np-Ox predominates; a weak doublet from olivine is the only ferrous phase evident. A hematite sextet is also present in sample HWMK12. A very weak sextet is also present for HWMK13; it was fit with the parameters for cation-deficient magnetite, but this assignment is uncertain.

As shown in Table 1, the saturation magnetizations (Js) for samples HWMK12 and HWMK13 (12.4 and 11.2 A m2/kg of Fe)

are about a factor of four larger than that for HWMK11 (3.2 A m2/kg of Fe). This is consistent with the M6ssbauer data which shows the presence of strongly magnetic magnetite in only the first two samples.

Diffuse reflectivity spectra for the 500-1000 and <20 gtm size fractions (method 1) are shown in Figure 8. The spectra for the HWMKll (red) sample are similar to those for bulk-hematite [e.g., Morris et al., 1985; Sherman and Waite, 1985], except that the minimum of the 6A 1 --> 4T lg electronic transition, which is near 860 nm in bulk hematite, is not well der'reed, and there is no

relative reflectivity maximum near 750 nm. Because hematites in these samples contain Ti, these differences may result from Ti substitution in the hematite structure. The 450-nm feature evident

in HWMK12 (black) and HWMK13 (yellow) sample spectra is due to the 6A1 --> 4E, 4A electronic transition of ferric iron in np- oxide particles and/or a mineralogy containing paramagnetic ferric iron. Morris and Lauer [1990] have shown that np-hematite can have a 450-nm feature very similar to the one observed here for HWMK12 and HWMK13.

Electron Microscopy

SEM analysis. Surface morphologies of the samples are shown in Figure 9. The surface of HWMK11 is rather featureless except for an occasional plagioclase particle protruding from the surface (Figure 9a). For HWMK12, a few cavities and vesicles were lined with smectite (Figure 9b), where a honeycomb-like smectite morphology could be seen at high magnification (Figure 9c). Particles from the yellow sample HWMK13, which is more palagonitized than HWMK12, were covered with erionite (a zeolite) and smectite (Figure 9d). The smectite has a crumpled paper-like morphology and formed from alteration of glass

Gox•m• [rr m..: •Y OF • SAM]'Lm, M•tT•A • 3407

0.8

0.6

0.4

0.2

0.0

.

.

.

.

0.0

.

.

.

0.2-

._> 0.4

n' 0.2

HWMK11 (Red) <20 um

500-1000 um

HWMK12 (Black) <20 um

500-1000 um - ,,

HWMK13 (Yellow) <20 um

0.4 ' •

500-10,••..,.u m I I I I I I I I I I I I I I I I I I I I I

0.0 0.2 0.4 0.6 0.8 1.0 1.2 1.4 1.6 1.8 2.0 2.2 Wavelength (x1000 nm)

Fig. 8. Diffuse reflectivity spectra (293K) for 500-1000 and <20 grn size fractions (method 1) HWMK11 (red), HWMK12 (black), and HWMK13 (yellow).

surfaces (Figure 9e). A magnified view of hexagonal rods of erionite is shown in Figure 9f.

TEM analysis of clay fractions. The fine clay (<0.2 !xrn) and coarse clay (0.2-2.0 !xm) fractions of HWMKll contained numerous iron oxide particles (0.1 to 0.4 grn) dispersed over smectite surfaces (Figure 10a). EDS analysis revealed that iron oxide particles often contained Ti. Np-Ox particles are present in both clay fractions of HWMK12 and HWMK13 (Figures 10b, 10c, and 10d). EDS analysis confirmed the presence of iron oxide particles on smectite surfaces and also some structural Fe in the clays. The morphology and composition of the clay fraction of sample HWMK13 was similar to that of sample HWMK12.

Summary

The mineralogy of the three palagonitic soils is summarized in Table 3. Size fractions listed in the table refer to those obtained

by method 2 (ultrasonic disruption followed by wet sieving and sedimentation). Because rinds were present and removed from HWMK12 and HWMK13 by the ultrasonic disruption, the mineralogy of size fractions <20 grn for those two samples is representative of the mineralogy of rinds still present on particles in size fractions >20 Ixm obtained by method 1 (wet sieving only).

The 20-1000 Imi size fractions. The HWMK11 (red) sample is essentially completely oxidized and has a hematite (Ti-hematite) pigment dispersed throughout the altered glass matrix. The alteration is present throughout particle, and only a trace amount of glass is present. In addition to ferric Fe-Ti oxides, other phases detected are plagioclase feldspar and a trace of olivine. Sample

HWMK12 (black) has the lowest proportion of ferric-bearing phases and is thus least weathered. It consists mostly of unaltered glass with embedded plagioclase and minor amounts of pyroxene, olivine, and Ti-magnetite. In some gains a thin palagonific rind is visible (Figure 5b), indicating some surface alteration. The mineralogy for sample HWMK13 (yellow) is the same as that for HWMK12. It has, however, distinct well-developed palagonitic rinds consisting of erionite and smecfite.

The <20 Ism size fractions. For sample HWMK11, the amount of glass and plagioclase decreases and the amount of smectite increases with decreasing particle size. The amount of hematite is essentially constant, and mica is present only in the coarse clay- sized fraction. The small amount of smectite present in the <0.2 gm fraction of HWMK11 sample is low in structural Fe. For samples HWMK12 and HWMK13, the same trends for glass, plagioclase, and smectites are observed. In addition, the zeolite erionite is also present, and its abundance decreases with decreasing particle size. The smectite for both black (HWMK12) and yellow (HWMK13) samples contains an appreciable amount of Fe in octahedral layers. There are only two mineral phases containing iron in the fine clay fraction, namely, smectites and iron oxides. In the red (HWMK11) sample, relatively large iron oxide particles (0.1 to 0.4 grn) were dispersed on clay surfaces; in black (HWMK12) and yellow (HWMK13) samples, much finer np-Ox particles were present in lesser concentrations.

DISCUSSION

Extent and Modes of Alteration The red, black, and yellow samples (HWMK11, HWMK12, and

HWMK13) show different extents and different modes of alteration. The highly oxidized nature and very low water content for the red sample implies that the alteration occurred in a dry environment and at high temperatures, possibly as a result of its proximity to a volcanic vent or as a result of primary volcanic alteration in the eruption. The occurrence of mica in sample HWMK11 is well size- sorted (0.2-2 Iaxn), suggesting an eolian origin which is in concurrence with earlier observations [Rex et al., 1969; Dymond et al., 1974]. The small amount of smectite may have completely or in part resulted from alteration of eolian mica. Fine particles of mica can easily alter into smectite [Fanning et al., 1989]. The low abundance of Fe in octahedral layers of the smectites in association with the red sample suggests a different origin (perhaps eolian) for these smectites as compared to Fe-substimted smectites in the clay fraction from black and yellow samples.

The high proportion of ferrous iron and low water contents for samples HWMK12 and HWMK13 indicate that they are relatively unaltered materials. This is especially the case for HWMK12, which has only 1.8% total water. This compares with 5.7% for HWMK13 and 14.0% for the palagonitic soil studied by Morris et al. [1990]. Clearly, HWMK12 and HWMK13 were not subjected to high-temperature and oxidizing conditions which was the alteration path for HWMK11. Most likely, samples HWMK12 and HWMK13 were subjected to aqueous alteration at relatively low temperatures (palagonitization). Palagonitization could result from either mild hydrothermal action of steam from volcanic vents or hot melt water from glaciers which were covering Mauna Kea during its past eruptions under a permanent ice cap or from a slow alteration process under an arid cold climate similar to what is observed today near the summit of Mauna Kea. Palagonifization as experienced by these samples indicate at least two alteration paths.

If the palagonitization occurred during past eruptions of the volcano under a permanent ice cap, then the smectite and zeolite

3408 GOL•.N ET AL.: •Y OF• SAMPLF. S. MAUNA I•A

Fig. 9. Scanning electron micrographs showing surface morphology of minerals: (a) red sample (HWMK11) with a plagioclase particle on its surface; (b) black sample (HWMK12) with authigenic smectite (s) in a cavity; (c) close-up of cavity in Figure 9b showing smectite having honeycomb-like morphology; (d) hexagonal rods of (twinned) efionite (e) and spongy mass of smecfite on the surface of yellow sample (HWMK13); (e) high magnification of smecfite in yellow sample; and if) efionite in the yellow sample.

GOLt)F_• ET At.: •¾ OF• SAMPt., MAtr•A KEA 3409

Fig. 10. Transmission electron micrographs of the fine clay fractions: (a) red sample (HWMK11) showing submicron sized iron oxide particles dispersed on clay surface; (b) thin section of black sample (HWMK12) showing smecfite; (c) clay from black (HWMK12) sample with some iron oxide particles (10 nm) on the surface;and (d) smecfite from yellow (HWMK13) sample.

erionite are relicts from that environment. However, the formation

of smecfite and the zeolite erionite is possible under an arid climatic regime, where there is little leaching of bases. Transient water action from snow melting, occasional rains, or atmospheric condensation in these soils has caused the dissolution of basaltic

glass and precipitation of zeolites. Basaltic glass dissolution under a limited moisture supply causes accumulation of bases under high pH conditions releasing A1 and Si from the basaltic tephra• This A1 and Si along with the alkali metal ions cause the precipitation of

structure. The above alterations under moderate to high pH conditions is contrary to low pH conditions on the Martian surface as discussed by Burns [1987, 1992a,b]. However, the incipient alteration of basaltic tephra provides a good process analog for Martian weathering. Absence of def'mitive mineralogical data for the Martian surface makes it very difficult to constrain the conditions of alteration to a high degree of precision, but one of the purposes of this study is to see how well the palagonite model fits the currently available Martian chemical data.

zeolite and smectite. Although alkali metals, A1, and Si are soluble under highpH conditions, any Fe2+ released from the dissolution of Application to Mars glass will be immediately oxidized and hydrolyzed. High concentration of Si in solution inhibits the crystallization of well- ordered iron oxides, and preserves the poorly crystalline phases like ferrihydrite [Schwertman and Taylor, 1989]. These iron phases may coexist with smecfite. Some Fe3+ may also be incorporated into the phyllosilicate structures. The smectite and erionite are both consistent with aqueous alteration. Color of the black sample (bulk) results mainly from Ti-magnetite in the unaltered basaltic tephra. Color of the yellow sample results from ferric np-Ox particles with perhaps some contribution from smecfite containing Fe 3+ in its

Various authors have suggested the possible presence of hematite, nontronite, smecfite/iron-oxide mixtures, jarosite, zeolite, etc., as possible Martian minerals which have formed by aqueous alteration [Banin and Marguiles, 1983; Burns, 1987; Ming and Gooding, 1988; Morris et al., 1989; Morris and Lauer, 1990]. If the weathering conditions encountered by Martian basalts were analogous to those encountered in Hawaiian palagonites, then the formation of minerals similar to what we have observed here may be a possibility on Mars. If minerals like zeolites and smectite are predominant on the Martian surface, they will play an important

3410 GOLDSN BT tin.: MlmmaJ. x•Y oF• S•. MAtr•AIQ•

TABLE 3. Mineralogy of Hawaiian Basaltic Tephra Samples Based on XRD (All Fractions), ED (Clay Fractions), and EMPA (Sand Fraction) Data

Size Fraction (gm)

Sample <0.2 0.2-2 2-5 5-20 20-1(X)0* Fine Clay Coarse Clay Fine Silt Medium Silt Sand

HWMK11

IT•MK12

HWMK13

smecfite smecfite plagioclase plagioclase plagioclase hematite hematite hematite hematite hematite

np-Ox (?) plagioclase smectite glass (?) glass (?) mica

np-Ox (?)

smecfite smectite erionite glass glass hemarite erionite smectite erionite plagioealse np-Ox plagioclase plagioclase plagioclase erionite

np-Ox magnetite smectite olivine pyroxene

Ti-magnetite

smectite smectite erionite glass glass Fe-oxides erionite smecfite efionite plagioclase

np-Ox plagioclase plagioclase plagioclase erionite np-Ox magnetite? smecfite olivine

pyroxene Ti-magnefite

Size fractions refer to method 2 (ultrasonic disaggregation followed by wet sieving and sedimentation). Components are fisted in approximate order of decreasing abundance. Np-Ox particles are too small to be detected by our XRD and EMPA procedures; thus np-Ox is not reported for silt and sand fractions.

* The sand fraction includes the coarse silt (20-50 grn) fraction.

role in the storage of volatiles (e.g., H20 and CO2) in the Martian issue]) from, for example, oxidation of magnetite, oxidative regolith [Gooding, 1986]. exsolufion from silicate phases, and transformation of np-Ox.

Iron oxides, structural iron in phyllosilicates, and other minerals found in the <20 Ixm fraction will be the most influential in Ac•owledgments. Authors wish to express their gratitude to determining the spectral properties of a soil. These materials have Cassandra Coorobs for her help in obtaining the basaltic tephra samples and John Mustard and James Bell 1II for their constructive review an influence on spectral properties that is inversely proportional to their particle diameter. Also, if they separate from parent particles, they will constitute the bulk of the eolian sediments. In the discussion below, reflectivity spectra for <20 Ixm size fractions are compared to that for Martian bright regions.

The reflectivity specmnn of HWMKll (Figure 8), which is dominated by well-crystalline Ti-hematite, is not similar to spectra for Martian bright regions even though it is red in color. The ferric crystal field transitions are too strong. On the other hand, the spectra for HWMK12 and HWMK13, which contain mostly np-oxides with minor well-crystalline oxides and small amounts (~1 wt %) of smecfite (not nontronite), are more similar because they are characterized by a relatively featureless ferric absorption edge as observed for Mars. However, Martian spectral data, especially the high spatial and spectral resolution data from the 1988 opposition [Bell et al., 1990a,b; Singer, 1990], do show evidence for weak ferric bands characteristic of well-

crystalline ferric oxides. In agreement with the results of this and other studies [Morris et al., 1989; Morris and Lauer, 1990; Bell et al., this issue], spectral data for Martian bright regions seem to be consistent with minor proportions of well-crystalline ferric oxides (probably hematite) in a matrix of palagonitic material like that of the <20 I•m size fractions of HWMK12 and HWMK13. In this scenario, np-Ox is formed, for example, hydrothermally at relatively low temperatures during eruption under ice or permafrost or subaerially at ambient Martian conditions. Smectite clays could also be formed during this weathering. Well- crystalline hematite is formed during transient and local heating events (volcanic activity, meteoritic ira. pact, etc. [Bell et al., this

comments.

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(Received January 31, 1992; revised October 19, 1992;

accepted October 29, 1992.)