Iron Oxides in Selected Brazilian Oxisols: II. Mossbauer StudiesM. P. F. Fontes,* L. H. Bowen, and S. B. Weed

ABSTRACTClay fractions (<2 ion) of Oxisols derived from four different

parent materials, all from the Triangulo Mineiro Region, MinasGerais, Brazil, were analyzed using Mossbauer spectroscopy to pro-vide insight into their Fe mineralogy. As little as 0.5% Fe as hematitein the total clay was quantitatively determined by this technique.Aluminum substitution in the Fe oxides resulted in lowering of thefield of maximum probability (H^) and this Al substitution affectedthe magnetic ordering of goethite more strongly than did poor crys-tallinity. However, Al substitution estimated from #„,„ showed someinconsistencies with varying temperature of Mossbauer analysis aswell as with values obtained from chemical analysis. Apparently,the currently available Mossbauer calibration curves, obtained fromsynthetic samples of hematite and goethite, are not readily applicableto soil samples. Little difference in goethite content was found amongthe soils developed on parent material from mafic rocks. The datasuggest that, in the presence of high levels of Fe in an ustic envi-ronment, the amount of goethite formed is not strongly correlatedwith organic matter. The values of the lower half-width (LHW) ofthe hyperfine field distribution at room temperature showed somehematites to consist of very fine particles (broad LHW), whereasother hematites were better crystallized (sharper LHW).

FOR MICROCRYSTALLINE PARTICLES of Fe Oxideswith ill-defined XRD patterns, an additional

technique that offers great potential is Mossbauer spec-troscopy (Bancroft, 1973; Greenwood and Gibb,1971). The Mossbauer technique gives unique infor-mation about the Fe-oxide constituents of soils, andit is particularly useful in distinguishing the two majornatural forms, hematite («Fe2O3) and goethite(aFeOOH) based on their differing magnetic hyperfineinteractions (Bowen and Weed, 1981; Murad, 1988).

Murad and Schwertmann (1983) considered Moss-bauer spectroscopy as an alternative physical methodto quantify Al substitution in Fe oxides. In the mag-netically ordered state, Fe oxides can be identified bythe internal magnetic hyperfine field (Hi). Substitutionof Fe(III) by Al causes magnetic dilution and lowersthe hyperfine field; however, as Hi is also influencedby crystal size, both Al substitution and crystal sizemust be taken into account to fully explain the vari-ation of Hi in goethites and hematites (Schwertmann,1987).

Well-crystallized Fe oxides order antiferromagnet-ically at the Neel temperature. As particle size de-creases, superparamagnetic relaxation eventuallyoccurs, and the Mossbauer spectrum has the same ap-pearance as in the paramagnetic state (a doublet). Thisphenomenon usually becomes noticeable for particleM.P.F. Fontes, Dep. de Solos, Univ. Federal de Vicosa, Vicosa—MG 36570, Brazil; L.H. Bowen, Dep. of Chemistry, and S.B. Weed,Dep. of Soil Science, North Carolina State Univ., Raleigh, NC27695-7619. Paper no. 12459 of the Journal Series of the NorthCarolina Agricultural Research Service, Raleigh, NC 27695-7643.This work was supported in part by the National Science Foun-dation, Grant EAR-850656. Received 5 Feb. 1990. "Correspondingauthor.

Published in Soil Sci. Soc. Am. J. 55:1150-1155 (1991).

sizes smaller than ?»20 nm. Soil Fe oxides are oftenin this range; therefore, Mossbauer measurements ofsoil materials usually must be carried out at low tem-peratures (77 K or less) to counteract the superpara-magnetic relaxation (Murad, 1977).

Mossbauer spectroscopy has been used to a limitedextent in the study of Brazilian soils. Bigham et al.(1978) showed by Mossbauer studies that the Fe oxidesin the clay fraction of two Oxisols from Brazil weremixtures of finely divided goethite and hematite («a 10-nm effective particle diameter). They also showed thehematite and goethite to contain a substantial amountof structural Al, in the range of 14 to 38 mol %. Pomboet al. (1982), working with a Dusky Red Latosol ofRio Grande do Sol state, reported the presence of he-matite and goethite, which was also verified by XRD.Maghemite was present in the coarse-clay particle-sizeseparate. The XRD peaks of hematite and goethiteindicated Al substitution, as did the values of Hi ob-tained by Mossbauer spectroscopy at room tempera-ture and liquid-N2 temperature.

More recently, Fabris el al. (1985) compared twoOxisols, very similar in chemical and mineralogicalcomposition but markedly different in color, from theBrazilian coastal plain. The Mossbauer spectra at 295,80, and 4.2 K showed the variations in relative goe-thite/hematite content, and Al substitution of hema-tite and goethite was calculated as about 20 mol %.

Our objective was to study the mineralogy and theAl substitution in Fe oxides from some Brazilian Ox-isols, using Mossbauer spectroscopy.

MATERIALS AND METHODSNatural clays (< 2 nm), separated by gravity settling (Jack-

son, 1979) from B horizon material from each soil and fromthe A horizon of one soil, were analyzed by Mossbauer spec-troscopy. The soils are derived from different parent mate-rials (Table 1). Results of XRD analysis (Fontes and Weed,1991) indicated that the samples from sandstone and basalthave only hematite as the Fe-oxide component, whereas theothers have a combination of hematite and goethite. A morecomplete description and chiiracterization of the samplesmay be found in Fontes and Weed (1991). Surface areas (N2)of the soil clays (<2 jttm), before and after citrate-dithionite(CD) extraction (Coffin, 1963) for removal of free Fe oxide,were determined using a Quantasorb Surface Area Analyzer1

(Quantachrome Corp., Syosseu, NY). Surface area of the Fe-oxide fraction was then estimated by difference.

Absorbers of 2 to 10 mg Fe/cm2 were prepared for Moss-bauer spectroscopy by packing the clay in brass rings withsupporting Al foil at the bottom. The sample was then cov-ered by dropwise addition of a benzene-styrofoam slurry andallowed to dry at room temperature.

1 The use of trade names in this publication does not imply en-dorsement by the North Carolina Agricultural Research Service ofthe products named, nor criticism of similar ones not mentioned.Abbreviations: Hmm, field of maximum probability; LHW, lowerhalf-width; XRD, x-ray diffraction; Hi, internal magnetic hyperfinefield; CD, citrate-dithionite; fg, recoil-free (Mossbauer) fraction forgoethite; fh, recoil-free fraction for hematite; e, ratio of fg to fli; RA,relative absorption peak area; QS, quadrupole splitting;mean crystallite diameter; IS, isomer shift.

1150

FONTES ET AL.: IRON OXIDES IN SELECTED BRAZILIAN OXISOLS: II. 1151

Mossbauer spectra were obtained in the constant-accel-eration mode. Velocity calibration of the spectrum was bylaser interferometry. This system has been described by DeGrave et al. (1982a).

The Mossbauer spectra were obtained as a function oftemperature at 16 and 80 K, and at room temperature (295± 3 K), using a closed-cycle cryostat for cooling the absorber.The source (=s 1.85 MBq 57Co/Rh) was at room temperature.Most samples were counted for 18 to 24 h to get reasonablestatistics, usually in the range of 6 X 105 to 106 counts perchannel.

The field distribution method of Wivel and Morup (1981)as modified by Amarasiriwardena et al. (1986) was used forspectral analysis. These authors, as well as Murad (1982),have shown that more realistic fits can be obtained for Fe-oxide spectra using a distribution of hyperfine fields ratherthan only one or two sextets. In this program, the isomershift (mm/s) and quadrupole interaction (mm/s) are iterated.The range of fields and step size must be chosen. The pro-gram is used to obtain the probability distribution for thehyperfine field, and from this is obtained the H^ (T), thefield of average probability, tfav (T), and the LHW (widthat half-maximum probability) of the asymmetric hyperfinefield distribution (T).

RESULTS AND DISCUSSIONMossbauer analyses of selected total clay (<2 /an)

samples are given in Table 2.The Fe-oxide fraction of the samples was assumed

to be composed of hematite or goethite or mixturesof these two minerals. This assumption is valid formost samples, although maghemite may be present insmall amounts; its interference is discussed below. Itwas also assumed that e, the ratio of fg and fh, is equalto one (Amarasiriwardena et al., 1986, 1988).

The relationship between measured RA in the Moss-bauer spectrum and Fe content (wt%) is linear for athin absorber (Bancroft, 1973). Since c = 1 for a mix-ture of goethite-hematite, the ratio of the weight per-cent Fe in goethite to that in hematite is directlyproportional to the ratio of the RA of goethite to thatof hematite. Assuming the formulas FeOOH for goe-thite and Fe2p3 for hematite (neglecting Al substitu-tion), the weight percent of hematite and goethite ineach oxide fraction was calculated (Table 3). Accord-ing to Amarasiriwardena et al. (1986), lower temper-atures than 80 K should be used if a doublet is presentin the 80 K spectrum. As a doublet was present insome samples at 80 K, the 16 K data were preferredto calculate the weight percent of hematite and goe-thite in all samples.

Soils from Clayey SedimentsFigure 1 shows Mossbauer spectra of samples P1B

and P3A, representative of soils from clayey sedi-ments. Sample P1B is from a well-drained environ-ment, whereas P3A is from a poorly drainedenvironment. The room-temperature spectrum of P1Bhas both a magnetic sextet and quadrupole doublet;the magnetic component is identified as hematite. At80 K, the spectrum of P1B shows an additional ru-dimentary magnetic splitting with very broad peaksfor goethite, suggesting very high Al substitution, verysmall particle size, or both. The unreasonably small#max values reported at 80 K for PIB and P2B are due

Table 1. Parent materials, classification, natural vegetation, and or-ganic C contents (surface horizons) of soils studied.

Soil Classification Vegetation Ct

Clayey sedimentsPIP2P3

P4P5P6P7

P8P9P10Pll

Typic AcrustoxTypic AcrustoxAquic Acrustox

SandstoneQuartzipsammentic HaplustoxQuartzipsammentic HaplustoxTypic AcrustoxTypic Acrustox

Typic AcrustoxTypic AcrustoxHumic EutrostoxRhodic Acrustox

Basalt

SchistP12 Typic Acrustox

CerradoCerradoCampo CerradoJ

CerradoCampo CerradoGrassForest

CerradoForestCerradoForest

Cerrado

18.217.015.3

8.54.7

17.814.4

17.529.712.839.7

16.5

t Analyses performed at EMBRAPA Laboratories, using wet oxidationmethod with K2Cr2O7 (EMBRAPA, 1982).

J Fewer trees and shrubs than Cerrado.

to this broad distribution, which, in fact, has no clearmaximum probability. At 16 K, all the Fe-oxide com-ponents should be magnetically ordered and, there-fore, the approximately 5% doublet is attributed to Feother than in Fe oxides. From the quantitative datafor this soil (Table 2), hematite was clearly identified,although present as < 10 g/kg of the P1B sample (Table3). This hematite could not be detected by XRD(Fontes and Weed, 1991).

For Sample P3A at room temperature, the goethitefraction is completely collapsed to a doublet, as inSample P1B, but there is no magnetic component,showing that there is no magnetically ordered hematitein this sample. At 80 K, most of the goethite orderedmagnetically, a somewhat larger fraction than in P1Bat that temperature. Surface-area data (Table 3) showthat the oxides in P3A have higher surface area («110vs. 60 m2/g) and consequently smaller particle size,than those in P1B, while the Al substitution deter-mined from XRD (Table 4) is lower than for P1B («20vs. 36 mol %). At 16 K, all the goethite should bemagnetically ordered. Thus, the 7% doublet remainingis a paramagnetic component.

Soils from SandstoneEach of the soils developed in parent material from

sandstone gave similar results with the exception ofSample P4B, which had much more goethite than theothers. Mossbauer spectra of P4B and P6B are usedto represent these soils (Fig. 2). The room-temperaturespectra consist of a sextet from the magnetically or-dered hematite and a doublet for goethite plus otherparamagnetic components. There is little change in theRA for goethite between 80 and 16 K for any of thesesamples. However, the RA for the doublet does de-crease for P4B from 12 to 5% of the spectrum, whichis attributed to a small amount of goethite not orderedat 80 K. Samples P5B to P7B also show a small de-crease in the doublet from 80 to 16 K, even though

1152 SOIL SCI. SOC. AM. J., VOL. 55, JULY-AUGUST 1991

Table 2. Mossbauer data (//„,„ field of maximum probability; LHW, lower half-width; QS, quadrupole splitting; IS, isomer shift; RA, relativeabsorption peak area) for total clay samples.

Goethite

Sample

P1B

P2B

P3A

P3B

P4B

P5B

P6B

P7B

P8B

P9B

P10B

P11B

Temp.K

RTt8016RT8016RT8016RT8018

RT8016RT8016RT8016RT8016

RT8016RT8016RT8016RT8016

H^

_4.5

47.6_4.6

48.2_

46.348.5

—47.348.9

_47.848.9

—48.649.2

—48.649.2_

48.949.2

_49.049.1_

49.149.1_

48.948.9

—49.049.1

LHW

-T ——————

_2.53.7_

2.73.4_

4.52.6_4.32.0

_

3.72.3_

2.41.7_

2.41.7_2.11.7

_

1.91.6_1.81.8_

2.01.9_1.91.8

QS ISf———— mm/s ———— ————— i —————

fl».x

T

Hematite

LHW

Soils from clayey sediments_

-0.18-0.18

_.-0.15-0.16

_-0.19-0.24

_-0.17-0.26

_-0.17-0.18

—-0.10-0.15

_-0.09-0.03

_-0.01-0.05

_-0.14-0.17

_-0.16-0.16

_-0.12-0.09

_-0.17-0.19

_0.340.35_

0.380.36_

0.380.35

—0.370.35

Soils from_

0.360.37

—0.380.36

—0.410.43_

0.420.40

Soils from_

0.340.33_

0.350.38_

0.340.35

—0.340.33

49.552.652.849.852.853.0-,——

49.052.653.0

sandstone49.652.553.049.952.753.050.052.852.950.052.852.9

mafic rocks48.852.052.649.052.853.049.452.152.349.351.852.8

_1.31.1_

2.60.9___3.01.31.0

2.51.41.02.21.30.91.91.21.11.91.21.1

2.91.41.22.81.51.62.71.41.22.91.21.3

QS———— mm/s

-0.27-0.17-0.15-0.24-0.17-0.17

___

-0.20-0.17-0.20

-0.19-0.19-0.18-0.20-0.20-0.20-0.21-0.19-0.19-0.20-0.19-0.18

-0.17-0.19-0.20-0.16-0.20-0.18-0.15-0.17-0.16-0.15-0.17-0.18

ISf

0.270.360.360.250.360.37___

0.250.360.37

0.250.360.370.240.360.370.250.360.370.240.370.37

0.240.370.370.250.360.360.260.370.380.250.370.36

RA

Goethite

_6883_5770_8693—5053

_3841—2221—1614—2117

_2718—3021_2922—2819

Hematite

121212272626___514243

585054817074878185867382

856980846877796675886979

Soil from schistP12B RT

8016

_

48.348.8

_

4.12.5

_

-0.15-0.20

_

0.350.36

50.052.852.9

2.01.21.1

-0.19-0.18-0.18

0.260.360.37

_5056

423640

t Isomer shift relative to Rh(Fe) source at room temperature. For IS relative to metallic Fe, add 0.11 mm/s to these values.$ RT = room temperature.

the observed RA of goethite shows no increase. Thisanomaly is due to the difficulty of determining preciseRA values from the broad distribution at 80 K. Al-though the samples of this group with low goethitecontent (P5, P6, and P7) have low absolute quadru-pole-splitting (QS) values for goethite, this is probablya fitting artifact due to the weak peak intensities. Thedoublet component at 16 K is <5% of total Fe, in-dicating 95% or more of the Fe is in the form of goe-thite or hematite. Of these samples, P4B has the largestLHW for goethite at both 80 and 16 K (Table 2). TheLHW of goethite in these samples is generally smallerat 16 K than the LHW of the goethites of the clayeysediments, in keeping with their better crystallinities.The //„,„ values for goethite (Table 2) are similar forsamples P5B, P6B, and P7B (48.6-48.9 T at 80 K and49.2 T at 16 K), but Hmax for P4B is lower (47.8 T at80 K and 48.9 T at 16 K), indicating a poorer crys-tallinity for this goethite in agreement with the LHW.

This low //mi0; value for Sample P4B is in agreementwith the relatively large surface area of the Fe-oxidecomponent of this soil (Table 3).Soils from Mafic Rocks

The Mossbauer spectra of the samples from soilsdeveloped in parent material from mafic rocks are sim-ilar (Table 1) and are illustrated by Sample P8B (Fig.3a). The room-temperature spectra show a prominentsextet from hematite (77-88% RA) and a small doubletmostly due to goethite. At 80 K, hematite and goethiteare magnetically ordered, and at 16 K the paramag-netic component for these samples is very low, on theorder of 1 to 3% RA. There is very little differenceamong the samples in terms of goethite RA or weight-percent Fe as goethite (Table 3), although Samples P9Band PUB come from a hijjh organic-C environment(under forest) (Table 1), which should favor goethiteformation over hematite (Schwertmann, 1988).

FONTES ET AL.: IRON OXIDES IN SELECTED BRAZILIAN OXISOLS: II. 1153

Table 3. Free-Fe content, surface area of Fe fraction, and Fe-oxidedistribution in clays (<2 nm) based on Mossbauer spectroscopy(16 K spectra).

Sample

P1BP2BP3AP3BP4BP5BP6BP7BP8BP9BP10BPUBP12B

Surface areaof Fe-oxide

Fef fraction

g/kg41.839.034.756.4

111.4127.3124.0132.5186.1161.0143.4161.892.4

m2/g

6267

1108593674745545156_68

Fe-oxidedistribution

Goethite Hematitewt %Wl 70

8875

10058462415192023252161

1225_42547685818077757939

Fe-oxide in clay$Goethite Hematite

58.045.355.249.575.644.727.936.254.454.951.749.985.7

e/ke ————& **&7.6

15.1—

36.189.6

141.8152.2156.9217.2180.9158.6186.555.1

t Data from Fontes and Weed (1991).j Calculated from the relative peak area for goethite and hematite (Table 2)

and the free-Fe content of the clay, and converting the Fe distribution tothe oxide forms (FeOOH and Fe2O3).

-15 15 -15 15

Velocity(mm/s)Fig. 1. Mossbauer spectra of soil samples (a) P1B (total clay) and

(b) P3A, from clayey sediments, at 16 and 80 K and at roomtemperature (RT).

The #max and LHW for the goethites in each of thesesamples are similar (Hmm «s49.0 T and LHW «<1.9T), indicating low Al substitution and/or larger par-ticle size than P3A and P1B.

As opposed to the other soils for which the goethitecomponent of the Mossbauer spectrum increases asthe temperature is lowered, these samples have higherRA for goethite at 80 than at 16 K. The 16 K RA forhematite is lower than at room temperature, again incontrast to most previous samples. These samples con-tain Al-substituted maghemite, according to the XRDdata (Fontes and Weed, 1991), and maghemite has alower magnetic field than does hematite (Lpngworthand Tite, 1977). The maghemite does not give a sep-arate Mossbauer sextet, but contributes to the assumedgoethite and hematite distributions to a different de-gree at different temperatures. At room temperature,it contributes to the hematite sextets; at 80 and 16 K,it contributes to both hematite and goethite, more so

•I

-16 15-16VelocKy(mm/8)

16

Fig. 2. Mossbauer spectra of soil samples (a) P4B (total clay) and(b) P6B, from sandstone, at 16 and 80 K and at room temperature(RT).

-16 16 -16VelocHyCmm/s)

16

Fig. 3. Mossbauer spectra of soil samples (a) P8B (total clay), frommafic rocks, and (b) P12B, from schist, at 16 and 80 K and atroom temperature (RT).

to the goethite at 80 K than at 16 K, according to theseresults. The magnetic fields of Al-substituted maghem-ite have not been reported, but should be lower thanthose given by Longworth and Tite (1977) for puremaghemite. The discrepancies between hematite RAat various temperatures are thus likely to be due tothe neglect of maghemite as a separate component inthe field distributions. This phenomenon may also ac-count for the lower //max observed for hematite at 16K, especially for Sample PI OB.

Soil from SchistFigure 3b shows the Mossbauer spectra for Sample

P12B. At room temperature, the hematite sextet is

1154 SOIL SCI. SOC. AM. J., VOL. 55, JULY-AUGUST 1991

prominent but of lower RA than for the soils fromsandstone or from mafic rocks. The doublet is de-creased in the 80 K pattern, and even more at 16 K.This decrease corresponds to an increase in the goe-thite RA, which means that some goethite becomesmagnetically ordered below 80 K.

The LHW of goethite in this sample (Table 2) issimilar to that for P4B and P3A, suggesting it also hasa very high goethite surface area. The //max values arealso similar for 16 K, although a little higher for thissample than for P4B and P3A. The goethite in P12Bhas high Al substitution (Fontes and Weed, 1991),which also contributes to the field reduction and LHWspread.

Comparison of HematitesSoil hematites commonly have particle sizes of

about 8 nm (Ibanga et al, 1983; Curi, 1983; Santana,1984; Kampf and Schwertmann, 1985) and should bemagnetically ordered at room temperature (Janot etal., 1973); therefore, comparisons in this study arebased mainly on room-temperature data (Table 2).

No hematite from any sample underwent the Morintransition, as seen by the fact that the quadrupole split-ting did not change sign at any temperature (De Graveet al., 1982b). Lowering of the Morin transition tem-perature is due either to poor crystallinity or Al sub-stitution or both (Kiindig et al., 1966; Srivastava andSharma, 1972; Kren, 1974; De Grave et al., 1982b).

The hematite from sample P3B exhibits the largestLHW (3.0 T) at room temperature, which indicatesthat it consists of smaller particles than most of theother samples. The soils from mafic rocks have a com-parable LHW («2.9 T); however, this broad LHWmay be due, in part, to the maghemite present in thesesoils. The sample from the sandstone rich in goethite(P4B) shows a LHW of 2.5 T, and the more crystalline

Table 4. Aluminum substitution (c) in goethite and hematite as de-termined from the Mossbauer parameters (#„,„, field of maximumprobability; Hi, internal magnetic hyperfine field) obtained fortotal clay samples, t

Goethite HematiteSample

P1BP2BP3AP3BP4BP5BP6BP7BP8BP9BP10BPUBP12B

1

_3224201414121110121116

2

57474145332828282429292935

3

635041_—_—__—_—33

XRDJ——— mol

362820_17____—__25

4

% AA

1916—34202016152624202115

5,1 ——

2815—20221715143614334114

6

2216—22161619192816382219

7

——1262109

1476-

XRDf

12—10101086

1414131411

11 = .//„„ = 50.4 - 0.13 c at 80 K (Amarasiriwardena et al., 1988); 2 =#n,« = 50.6 - 0.051 c at 16 K (Amarasiriwardena et al., 1988); 3 = Hi= 50.65 - 0.042 c - 8.7/MCD,,, (where MCD = mean crystallite di-ameter) at 4.2 K (Murad and Schwertmann, 1983); 4 = #„„ = 51.6 -0.11 c at room temperature (Amarasiriwardena et al., 1988); 5 = /?„» —53.35 - 0.038 c at 80 K (De Grave et al., 1988); 6 = #„„ = 53.54 -0.033 c at 16 K (De Grave et al., 1988); 7 = Hi = 51.72 - 0.076 c -3.2/MCDc at room temperature (Murad and Schwertmann, 1986).

t X-ray diffraction data from Fontes and Weed (1991).

hematites from all the other soils from sandstone andschist show smaller LHW of approximately 2.0 T. Thistrend agrees with the lower surface area of the Fe-oxidecomponent of the hematitic soils from sandstone (Ta-ble 3). An exception is Sample P12B (soil from schist),in which the surface area is comparatively high butthe hematite LHW corresponds to large particle size.This anomaly may be due to the presence of largeramounts of goethite of very small particle size in Sam-ple P12B, which would be included in the determi-nation of the Fe-oxide surface area.

The //max values at room temperature of all hema-tites are compared in Table 2. The soils from maficrocks show lower //max (48.8-49.4 T), although thesevalues may be influenced by the maghemite present,whereas the hematitic soils from sandstone have high-er //max (49.6-50.0 T). In the intermediate range arethe samples from clayey sediments.

Estimation of Aluminum Substitutionfrom Mossbauer Parameters

Since the magnetic hyperfine fields of hematite andgoethite are lowered with Al substitution (Fysh andClark, 1982; Golden et al., 1979; Goodman and Lewis,1981), their fields are suited not only for identifyingthe oxides, but also for characterizing them with re-spect to Al substitution (Murad and Schwertmann,1983). Reductions in magnetic hyperfine fields are alsoassociated with poorly crystallized unsubstituted goe-thites (Golden et al., 1979; Murad, 1982), due to thesmall particle size. For hematites, the different pub-lished relationships betwee:n the magnetic hyperfinefield at room temperature and degree of Al substitutionled Murad and Schwertmann (1986) to conclude thatthis field reduction should tie seen as the simultaneouseffects of Al substitution and poor crystallinity. In ourstudy, the substitutions of Al for Fe in goethite andhematite were calculated from the Mossbauer datausing different calibration curves from those in theliterature.

Aluminum substitution in goethites calculated ac-cording to Amarasiriwardena et al. (1988) for Moss-bauer data obtained at 80 and 16 K gave results thatdiffer markedly (Columns 1 and 2, Table 4) and werehigher than Al substitutions determined from XRD(Fontes and Weed, 1991). According to Amarasiri-wardena et al. (1988), discrepancies between Al sub-stitution from //max obtained at different temperaturesprovide a strong indication of small-particle-size ef-fects. Therefore, a third relationship was used that in-troduces particle size through the MCDm obtainedfrom an independent estimate (Murad and Schwert-mann, 1983), although it should be noted that thisrelationship was developed for spectra at 4 K, ratherthan at 16 K. Column 3 in Table 4 shows that, for thegoethites where MCDU1 was available, and assumingthe fields at 16 K are the same as at 4 K, no improve-ment in the prediction of Al substitution was obtained.The results in Column 3 agree approximately with theresults in Column 2, but they both are well above theAl substitution determined by XRD. The calibrationcurves were obtained from synthetic Fe oxides andour results indicate a discrepancy between these curves

FONTES ET AL.: IRON OXIDES IN SELECTED BRAZILIAN OXISOLS: II. 1155

and soil goethites, even with crystallinity and Al-sub-stitution effects included.

The results for Al substitution in hematites alsoshow discrepancies (Table 4). The room-temperatureand 80 and 16 K data using equations from Amara-siriwardena et al. (1988) and De Grave et al. (1988)gave higher values than XRD. These equations do nottake particle size into consideration. Murad andSchwertmann (1986) proposed an equation where bothAl substitution and crystallinity would be taken intoaccount, provided an independent estimate of particlesize was available. Using this equation, the results(Column 7, Table 4) gave a fair agreement for severalsamples, while for others Al substitution was under-estimated.

CONCLUSIONSMossbauer spectroscopy of these varied Oxisol sam-

ples provides insight not only into their Fe-oxide min-eralogy but also the particle-size effects and effects ofAl substitution of the oxides. The results show thatlow-temperature spectra are required to obtain quan-titative data for Fe oxides in soils. The presence ofmaghemite in the mafic soils is not readily quantifiedby the Mossbauer spectra, but causes error in quan-titative estimates of goethite-hematite ratios under theexperimental conditions used here. The presentlyavailable calibration curves for Al substitution in he-matite and goethite are not uniformly applicable tosoil Fe oxides.