ClayMinerals (1982) 17, 359-363
NOTE
QUANTITATIVE
HEMATITE
DETERMINATION
IN KAOLINITIC
SOILS
OF GOETHITE
BY X - R A Y
AND
DIFFRACTION
In the past, two factors have impeded the quantitative estimate of Fe-oxides in soils by
X-ray diffraction. First, Fe-oxides are still quite often considered X-ray amorphous,
although numerous results, e.g. a low ratio of oxalate- to dithionite-soluble Fe, have
indicated the opposite. Second, even if crystalline, the concentration of Fe-oxides in many
soils is low, thereby complicating their identification by XRD. Recently, however, more
sensitive methods such as M6ssbauer spectroscopy and Differential-XRD (Schulze, 1981)
have been introduced, which substantially reduce the lower limit of detection. Because
these two methods are not generally available and, especially in the case of M6ssbauer
spectroscopy, are rather time consuming, ordinary X R D should be adapted for
quantitative estimation of Fe-oxides.
Determination can be facilitated by using samples in which the Fe-oxides are
concentrated by particle-size separation and a 5 M N a O H boiling treatment (Norrish &
Taylor, 1961). The latter treatment is particularly suitable for kaolinitic soils as the
Fe-oxides are unaffected--provided certain precautions are taken (K~impf &
Schwertmann, 1982a). This paper gives details of a procedure for the quantitative
estimation of goethite (Gt) and hematite (Hm) by XRD.
Soil material
Gt and H m contents of 49 clay fractions from kaolinitic Inceptisols, Ultisols and Oxisols
formed from basalt in southern Brazil (K/impf, 1981) with an Feo/F%* ratio of <0.05 were
concentrated by boiling in 5 M N a O H for 60 min. (K~impf & Schwertmann, 1982a). After
this treatment the samples consisted of 27-88% Fe-oxides, 9 - 6 9 % residual silicates, 0 - 5 %
quartz and 0 - 5 % anatase. The concentration factor for Fe-oxides varied between 3 and 20.
Iron in these samples was extracted by the dithionite/bicarbonate/citrate procedure (Mehra
& Jackson, 1960).
X-ray diffraction techniques
The sample was gently pressed into a round perspex holder (Klug & Alexander, 1974, p.
373) against filter paper to minimize orientation. A Philips diffractometer was used with a
vertical goniometer, a graphite monochromator, C o - K radiation (25 mA, 35 kV) and a
scanning speed of 89176
20 min. Other instrument constants were: 2 sec. t.c., 1~ divergence slit,
0.2 mm receiving slit and 1 ~ scatter slit.
The Gt (110) and Hm (012) lines were scanned three times and the areas under these
measured with a planimeter. If there was an adjacent quartz (100) peak, the integrated
* Feo and Fea represent the oxalate- and dithionite-soluble Fe respectively.
9 1982 The Mineralogical Society
360
Note
intensity of the Gt (110) peak was found by doubling the area of the unaffected
higher-angle half of the peak. The percentage Gt or Hm was calculated from the equation
Ix U*
% Gt or Hm = - - . - - .
is u*
I x and I S represent, respectively, the line intensities of the mineral to be measured and
standard; p* and p* represent the respective mass absorption coefficients of sample and
standard.
The value of Px* was obtained from the initial X R D approximation of quartz (p* = 54.8)
and anatase (/~* = 191-9) contents, from the Fe d value for the Fe-oxides (average/~* for
Hm and G t = 4 1 . 0 ) and for the residual kaolinite ( p * = 46.2) (not dissolved by 5 M
NaOH) taken as the weight difference between 100 and the sum of quartz, anatase and
Fe-oxides. The effect of possible Al-substitution in the Fe-oxides was not considered when
calculating p* because of its negligible influence when Co-radiation is used. The calculated
values for p* for 45 samples only varied between 43 and 45 which does not differ
significantly from a 1 : 1 mixture of H m and Gt (p* = 41). This is because anatase was only
a very minor fraction of the samples and p* of all other components is similar to that of
H m and Gt (for Co-radiation). Consequently, the calibration curvet showing the
relationship between peak intensity and Gt or Hm concentration was essentially linear
(r = 0.9996). It was, therefore, assumed to be linear for the samples also. Neglecting the
variation of p* (i.e. taking Px* = P*) did not significantly change the G t / G t + H m ratio.
This means that for a set of samples of similar mineralogy p* can be neglected. From the
Gt and Hm contents determined by XRD, Fe present in these two minerals was calculated
(FexRD) and compared with Fe d.
Reference samples
The main limitation to accurate quantitative estimation of Gt and H m lies in achieving
similarity between the Gt and H m in the soils and the reference samples used for
calibration. For this reason a number of different Gt and Hm samples were tried. Natural
standards were isolated from soils by the boiling 5 M N a O H treatment. Whereas kaolinitic
soils with Gt as the only Fe-oxide were readily available, soils rich in H m were found never
to be completely free of Gt. A soil clay with a H m / H m + Gt ratio of 0.94 thus had to be
used to produce the Hm standard and a 6% correction for the Gt present was applied. In
addition, Hm 39, a hematite specimen of unspecified locality, was tested. Synthetic
standards were produced as follows: Gt G N K 2 by storing ferrihydrite at pH 13 and 25~
for 14 days, Gt G2B3II by slow oxidation of a mixed FeClz-A1C13 solution at pH 11.2,
and hematite H N K 3 by storing a mixed FeC13-A1C13 solution at pH 7 and 50~ for 90
days.
Table 1 shows the large variation of the Gt (110) and Hm (012) line intensity of the
various reference samples. D. G. Schulze from this laboratory has recently shown that the
Gt (110) line is particularly sensitive to structural defects. The (130) line of Gt is difficult to
use because of interference with the (104) line of Hm, whereas the Gt (111) line is too weak
in samples low in Gt and difficult to measure in the presence of larger amounts of H m
and/or quartz.
t The calibration curve was derived from mixtures of 70% (Gt + Hm) standards with various Gt/Hm ratios
+ 25% kaolinite (from Rosenthal, FRG), 2.5% anatase (Merck TiO2) and 2.5% quartz.
Note
361
TABLE 1. Properties of various standard goethites and hematites.
(110) line
Goethite
GNK2, synth.
GFe 3/1, synth.
G2B3II, synth.
G3Bj/2 Fe, nat.
G4Va/2 Fe, nat.
(130) line
Al-substitution*
labs.
IreL
labs.
Irel.
(mole %)
900
750
453
625
664
100
83
50
69
74
354
284
162
172
283
100
80
46
48
80
0
12
31
23
17
(012) line
Al-substitution'~
Hematite
labs.
Irel.
HNK2, synth.
HNK3, synth.
HM39, nat.
H6Pf/26Fe, nat.
275
195
218
180
100
71
79
65
(mole %)
2
10
0
5
* After Thiel 0963). t After v. Steinwehr 0967).
A
i
A
,"
oJ
UL.
: ", A 9
1.2
A
1,0
:
9
A
A
-A
4
llA
9
A
0.8
0,6
0,2
0.4
0.6
0,8
1.0
Gt/GI,Hm
FIG. 1. Plot of the ratio FeXRD/Fed against Gt/Gt + Hm using a synthetic (upper) and a soil
(lower) goethite-hematite pair. FexR o = Fe derived from Gt + Hm as determined by XRD;
Fe = dithionite-extract~ible Fe.
Note
362
To select the most suitable pair of standards, a plot of FeXRD/Fe d against G t / G t + H m
was used. If the standard is reasonably suitable, all data should lie around a horizontal line
at FeXRD/F%----l. Values below or above the horizontal line indicate under- or
overestimation of the two minerals.
The upper part of Fig. 1 shows an example where synthetic standards were used. Here
the proportion of Gt is overestimated whereas that of Hm is underestimated. In contrast,
no such trend is visible (lower part of Fig. 1) when two natural standards were used,
namely a Gt and a Hm isolated from an Ultisol B horizon and an Oxisol C horizon,
respectively. These standards were then used for the results reported in the next section.
Validity of results
In contrast to clay silicates, the quality of the procedure can be checked by comparing
the X R D results with the amount of 'free Fe-oxides' independently determined chemically
(dithionite) if no other Fe-oxides besides Gt and Hm are present. This is the case in many
soils. We have used this test twice before. Once for some Spanish soils containing Gt and
H m where the relationship between Fe from Gt + Hm as determined by X R D (FexRD) and
the Fe extracted by dithionite (Fed) was FeXRD = 0.83F% + 2.5 (r = 0.96, n = 11)
(Torrent et al., 1980) and also in a group of soils from South Africa with Gt and
lepidocrocite where the relationship was F e X R D = 0 . 9 4 F % + 2.6 ( r = 0 . 8 7 , n = 32)
(Schwertmann & Fitzpatrick, 1977).
In the present study the relationship between FexR" and Fe d was again significant (Fig.
2) and the regression coefficient of 0.97 is statistically not different from 1-0. This proves
that the method yields correct results. The accuracy of FexR o determination is indicated in
60,
(o1
o~ ~C
yr
9
12J
L.1--
20
gr
(o)
:/o
(o)
II
I
-
e
-
( r = 0.96. n = 6,5 I
4'o
'
6'o
Fed (%)
FIG. 2. Plot of Fexa D against Fe0. Broken lines indicate 95% confidence interval of the
regression (symbols as in Fig. 1). The four points in () were not used in calculating the
regression line.
Note
363
Fig. 2 b y the 9 5 % confidence interval. C o n s i d e r i n g the n u m e r o u s uncertainties, a m o n g
which an u n s a t i s f a c t o r y m a t c h i n g o f the reference samples is p r o b a b l y the m o s t limiting
one, the m e t h o d seems a c c u r a t e e n o u g h to be used for m o s t soil m i n e r a l o g i c a l studies (see
e.g. K~impf & S c h w e r t m a n n , 1982b).
ACKNOWLEDGMENT
The assistence of Dr R. M. Taylor, CSIRO Div. of Soils, Adelaide, S. Australia, in preparing the manuscript is
gratefully acknowledged. One of us (NK) acknowledges the financial support by MEC-CAPES, Brazil (Grant
4617/76).
Institut f/ir B o d e n k u n d e ,
T e c h n i s c h e Universit~it M / i n c h e n ,
8050 Freising- W e i h e n s t e p h a n ,
FRG
1 N o v e m b e r 1981; revised 29 D e c e m b e r 1981
N. K.~MPF*
O. SCHWERTMANN
REFERENCES
K.~MPF N. (1981) Die Eisenoxidmineralogie einer Klimasequenz von Btden aus Eruptiva in Rio Grande do
Sul, Brasilien. Dissertation Techn. Univ. M/inchen.
Ks
N. & SCHWERTMANNU. (1982a) The 5 M NaOH concentration method for iron oxides in soils. Clays
Clay Min. (in press).
Ks
N. & SCHWERTMANNU. (1982b) Goethite and hematite in a soil climosequence from volcanic rocks in
southern Brazil and its usefulness for differentiating kaolinitic soils. Geoderma. (in press).
KLUG H.P. & ALEXANDERL.E. (1974) X-ray diffraetionproeedures. Wiley, New York.
MEHRA O.P. & JACKSONM.L. (1960) Iron oxide removal from soils and clays by a dithionite-citrate system
buffered with sodium bicarbonate. Clays Clay Miner. 7, 317-327.
NORmSH K. & TAYLORR.M. (1961) The isomorphous replacement of iron by aluminium in soil goethites. J.
Soil Sci. 12, 294-306.
SCHULZED.G. (1981) Identification of soil iron oxide minerals by differential X-ray diffraction. Soil Sci. Soe.
Am. J. 45,437-440.
SCHWERTMANNO. & FITZPATRICKR.W. (i977) Occurrence of lepidocrocite and its association with goethite
in Natal soils. Soil Sei. Soe. Am. J. 41, 1013-1018.
STEINWEHR H.E.v. (1967) Gitterkonstanten im System tt-(AI,Fe,Cr)203 und ihr Abweichen yon der
Vegardregel. Z. Kristallogr. Mineral. 125, 377-403.
THIEL R. (1963) Zum System a'-FeOOH-a'-AIOOH. Z. anorg, allg. Chem. 326, 70-78.
TORRENT J., SCHWERTMANNO. t~ SCHULZED.G. (1980) Iron oxide mineralogy of some soils of two river
terrace sequences in Spain. Geoderma 23, 191-208.
* On leave from: Depto. Solos, Fac. Agronomia-UFRGS, Porto Alegre, Brazil.