ORIGINAL PAPER

Determination of Fe3+/RFe ratios in chrome spinels usinga combined Mossbauer and single-crystal X-ray approach:application to chromitites from the mantle section of the Omanophiolite

Davide Lenaz • Jacob Adetunji • Hugh Rollinson

Received: 17 June 2013 / Accepted: 12 December 2013 / Published online: 3 January 2014

� Springer-Verlag Berlin Heidelberg 2014

Abstract We present the results of a comparative study

in which we have measured Fe3?/RFe ratios in chromites

from mantle chromitites in the Oman ophiolite using

Mossbauer spectroscopy and single-crystal X-ray diffrac-

tion. We have compared these results with ratios calculated

from mineral stoichiometry and find that mineral stoichi-

ometry calculations do not accurately reflect the measured

Fe3?/RFe ratios. We have identified three groups of sam-

ples. The majority preserve Fe3?/RFe ratios which are

thought to be magmatic, whereas a few samples are highly

oxidized and have high Fe3?/RFe ratios. There is also a

group of partially oxidized samples. The oxidized chrom-

ites show anomalously low cell edge (a0) values and their

oxygen positional parameters among the lowest ever found

for chromites. Site occupancy calculations show that some

chromites are non-stoichiometric and contain vacancies in

their structure randomly distributed between both the T and

M sites. The field relationships suggest that the oxidation of

the magmatic chromitites took place in association with a

ductile shear zone in mantle harzburgites. Primary mag-

matic Fe3?/RFe ratios measured for the Oman mantle

chromitites are between 0.193–0.285 (X-ray data) and

0.164–0.270 (Mossbauer data) and preserve a range of

Fe3?/RFe ratios which we propose is real and reflects

differences in the composition of the magmas parental to

the chromitites. The range of values extends from those

MORB melts (0.16 ± 0.1) to those for arc basalts

(0.22–0.28).

Keywords Cr-spinel chromitite � X-ray single-crystal

diffraction � Mossbauer non-stoichiometry � Oxidation

processes � Oman

Introduction

Understanding the oxidation state of the Earth’s mantle is a

major quest for igneous petrology for it has significant

implications for our understanding of the genesis of mag-

mas, their evolution and the origin of associated volatile

species. This is particularly relevant to the current debate

about the mantle source of basaltic lavas and the possible

contrasting source regions of MORB and arc lavas (Car-

michael 1991; Lee et al. 2010; Cottrell and Kelley 2011). A

major tool in understanding the oxidation state of the

Earth’s mantle has been the study of Fe-bearing spinels in

which the Fe is present in both the Fe2? and Fe3? oxidation

states. From this, the Fe3?/RFe (Fe3 ?/(Fe2? ? Fe3?)

ratio can be determined and used as a measure of the

degree of oxidation. This approach has been widely used to

characterize the oxidation state of the subcontinental

mantle using spinel-peridotite xenoliths (Dyar et al. 1989;

Luth and Canil 1993; Canil et al. 1994; Canil and O’Neill

1996; McCammon and Kopylova 2004). A slightly dif-

ferent approach has been taken in understanding the nature

of the suboceanic mantle. In this case, chrome-spinel from

MORB melts and from chromitites in the mantle, thought

to have formed in melt channels in the mantle, is used to

record the oxidation state of melt fluxing through the upper

mantle (Rollinson et al. 2012; Ballhaus et al. 1991).

Communicated by C. Ballhaus.

D. Lenaz (&)

Department of Mathematics and Geosciences, University

of Trieste, Via Weiss 8, 34122 Trieste, Italy

e-mail: lenaz@units.it

J. Adetunji � H. Rollinson

School of Science, University of Derby, Kedleston Road,

Derby DE22 1GB, UK

123

Contrib Mineral Petrol (2014) 167:958

DOI 10.1007/s00410-013-0958-2

However, accurately measuring the Fe3?/RFe ratio in

spinels is not trivial. The most common approach is to

estimate the ratio from electron microprobe analyses using

a charge balance calculation such as that of Droop (1978).

Yet this is based upon the untested assumption of mineral

stoichiometry. A more accurate approach is to measure

directly the Fe3?/RFe ratio in spinels using the technique of

Mossbauer spectroscopy (Wood and Virgo 1989; Rollinson

et al. 2012). This technique is time consuming for it

requires the preparation of pure chromite mineral separates

and the spectroscopy itself may take one or more days per

sample and for this reason is not commonly used. More

importantly, however, the method does not always provide

a unique solution, for there can be a number of possible

ways of fitting the spectrum which can lead to a variety of

possible ‘measured’ Fe3?/RFe ratios. Further in our pre-

vious studies of chromitites from the mantle section of the

Oman ophiolite (Rollinson et al., 2012; Rollinson and

Adetunji 2013a, b), we have accurately determined Fe3?/

RFe ratios using the Mossbauer technique and yet find a

large variability in the results, even when samples are

collected from a small geographic area. For example, we

have found that chromitites preserve a wide range of oxi-

dation ratios, a feature also noted in the study of other

ophiolitic chromitites (Quintiliani 2005; Quintiliani et al.

2006). Some samples show extreme oxidation with Fe3?/

RFe ratios close to 1.0, a feature also noted in Bushveld

chromitites (Adetunji et al. 2013) and attributed to post-

magmatic processes; further some of the oxidized samples

are not stoichiometric a feature also noted in the Bushveld

chromitites by Nell and Pollack (1998). This variability

leads to large uncertainties in estimating the primary

magmatic Fe3?/RFe ratios of mantle chromitites.

Our purpose here therefore is to combine a number of

different mineral-chemical techniques to find a robust

approach to measure Fe3?/RFe ratios in chrome spinels.

The samples are from an area which is well characterized

both geologically and petrologically so that any variations

in Fe3?/RFe ratio can be contextualized. We present a

method which uses the techniques of electron microprobe

analysis, Mossbauer spectroscopy and single-crystal X-ray

diffraction to determine the crystal chemistry, site occu-

pancies and Fe3?/RFe ratios in mantle chrome spinels. Our

objectives therefore are to identify

(a) the true range of variability of Fe3?/RFe ratios in

mantle chromitites. This has major implications for

magmatic processes within the upper mantle and also

has profound implications for the tectonic setting

within which ophiolites are created;

(b) those later, post-magmatic oxidation processes which

they appear to record.

This is in order that we might more precisely assess the

variable oxidation state of the suboceanic mantle.

Geological setting and petrography

We use as a case study samples from Wadi Rajmi from the

mantle section of the Oman ophiolite in the north of Oman.

These samples have been described in some detail both

with respect to their field occurrence and their geochem-

istry by Rollinson (2008). They were further investigated

by Mossbauer spectroscopy by Rollinson et al. (2012) and

Rollinson and Adetunji (2013b). Here, we use an expanded

data set in which we have refined the fitting of Mossbauer

Fig. 1 Geological map of the

Wadi Rajmi area of the Oman

ophiolite, showing the sample

localities and the oxidation zone

identified by Rollinson and

Adetunji (2013b)

958 Page 2 of 17 Contrib Mineral Petrol (2014) 167:958

123

spectra and constrain our data using the results of single-

crystal X-ray diffraction for the same samples.

The samples are from chromitite pods hosted in harz-

burgite, in the mantle section of the Oman ophiolite in the

Wadi Rajmi area in the north of Oman. Sample localities

are shown in Fig. 1. A previous study of their geochemistry

and field relationships showed that the chromitites may be

divided into two groups. Chromitites which formed close to

the Moho form concordant pods, have relatively low cr#

[Cr/(Cr ? Al)] (ca 0.5–0.6), are associated with gabbroic

interstitial minerals and appear to have been derived from a

slightly fractionated melt, thought to be the product of

mixing between MORB-like magmas and depleted harz-

burgite. Chromitites from deeper in the mantle section are

discordant with respect to banding in the mantle harz-

burgite, have higher cr# ([0.7), are associated with inter-

stitial olivine and are thought to be derived from a boninitic

parent magma (Rollinson 2008). Some of these samples are

located close to a high-temperature shear zone described in

detail by Michibayashi et al. (2006).

Samples from close to the Moho are from the Maharra

and Rajmi sites (Fig. 1). These samples have chromite with

cr# of between 0.52 and 0.60. The samples from Maharra

(04–25, 04–26, 85 and 75 % chromite by volume, respec-

tively) are from a layered sill and contain chromite grains

between 3 and 5 mm in diameter containing olivine

inclusions and interstitial olivine grains (Fo94–95) up to

4 mm in diameter. In sample 04-26, some of the olivine is

altered to serpentine, calcite and tremolite. SEM images

show both light and dark patches indicating some alteration

of the chromite and the creation of more and less alumi-

nous regions with cr# rising from 0.52 to 0.63 (Fig. 2a).

Sample 12–15 (95 % chromite) with interstitial olivine and

clinopyroxene is from a massive domain within an anti-

nodular chromitite containing elliptical olivine nodules

2–3 cm long.

Chromitite from the Rajmi North (05–15, 90 % chro-

mite) comprises large (5 mm) subhedral grains, free of

inclusions with smaller (2 mm) interstitial clinopyroxene

(mg = 0.95–0.97), altered in places to magnesiohorn-

blende and plagioclase (An70). There are small areas of

alteration 50 lm across in some chromite grains in which

the cr# increases from 0.58 to 0.68 and the fe# from 0.38 to

0.45. At Rajmi South (03–18, 85 % chromite), chromite

forms large (5 mm) elliptical grains with irregular grain

boundaries rimmed with magnesiohornblende. This local-

ity is serpentinized, and chlorite is also present.

At about 3 km below the Moho (horizontal distance) are

chromitites from the Shamis and Jabri pits (Fig. 1). At

Shamis (03–12, 04–05, 85 and 90 % chromite, respectively),

chromite (cr# between 0.58 and 0.64) forms grains 4–5 mm

Fig. 2 Backscattered images of selected chromites. a sample 04-26:

SEM image shows both light and dark patches indicating some

alteration of the chromite and the creation of more and less aluminous

regions with cr# rising from 0.52 to 0.63; b sample 04-07: from a

zone of intense brecciation where the chromite is reduced to angular

fragments with a grain size of a few microns; c sample 03-11: grains

showing occasional alteration to ferritchromite along fine fractures

which are up to about 5 lm wide and on grain boundaries in which

the cr# increases to 0.82

Contrib Mineral Petrol (2014) 167:958 Page 3 of 17 958

123

in diameter with interstitial olivine (Fo95–96). The chromite

contains inclusions of olivine, and the olivine contains small

inclusions of chromite. Sample 12-01 is from the Jabri pit

and contains chromites (cr# = 0.63–0.65) with inclusions of

an aluminous amphibole and interstitial olivine (Fo95–97)

which is serpentinized in part.

Deeper in the mantle at 5–6 km below the Moho (hor-

izontal distance), the chromitites of the Mining Camp

locality (Fig. 1) form irregular bands and veins in the host

dunite and are located close to a high-temperature shear

zone in the mantle harzburgites. These have higher cr# of

between 0.71 and 0.78 and contain interstitial olivine

which is highly magnesian (Fo95–97). Sample 04-07

([90 % chromite) from north of the Mining Camp locality

contains grains up to 3 mm in diameter with inclusions of a

Ca-poor amphibole. This sample is strongly brecciated in

places and contains zones of angular fragments of chromite

in which the grain size is reduced to a few microns

(Fig. 2b). In addition, adjacent to these zones, there are

patches in the chromite grains of up to about 100 lm

across where the chromite is altered to ferritchromite.

Samples 04-11 and 04-12 are from a smaller pit to the

north of the main pit. Sample 04-11 (85 % chromite)

contains two different sizes of chromite grains. There are

larger subhedral grains up to 4 mm across, some with

olivine inclusions intergrown with interstitial olivine,

which in places contains small rounded chromite grains up

to about 100 lm across (Fig. 2b). Both grain sizes have the

Table 1 Results of structure refinement

Sample locality 05-15A

Rajmi

03-18A

Rajmi

04-25A

Maharra

04-26A

Maharra

12-15A

Maharra

04-05A

Shamis

03-12A

Shamis

12-01A

Jabri

04-07A

Mining camp

a0 8.2600 (2) 8.2535 (3) 8.2464 (2) 8.2405 (5) 8.2485 (3) 8.2511 (1) 8.2675 (4) 8.2709 (2) 8.2542 (3)

u 0.26286 (8) 0.26286 (9) 0.2628 (1) 0.2629 (1) 0.26286 (7) 0.2628 (1) 0.2626 (1) 0.2627 (1) 0.2608 (2)

T–O 1.972 (1) 1.971 (1) 1.968 (1) 1.969 (1) 1.970 (1) 1.969 (1) 1.971 (2) 1.972 (1) 1.941 (3)

M–O 1.9645 (6) 1.9630 (7) 1.9617 (7) 1.9593 (7) 1.9618 (5) 1.9628 (7) 1.9679 (9) 1.9685 (7) 1.979 (1)

m.a.n.T 17.8 (2) 17.5 (2) 16.8 (3) 18.1 (2) 17.6 (1) 16.7 (2) 17.0 (2) 17.2 (2) 16.8 (5)

m.a.n.M 19.7 (2) 19.5 (3) 19.1 (4) 19.3 (2) 19.4 (1) 19.6 (2) 20.3 (3) 20.7 (3) 21.5 (8)

U (M) 0.0040 (9) 0.0054 (1) 0.0049 (2) 0.0048 (1) 0.0037 (8) 0.0043 (1) 0.0042 (1) 0.0048 (1) 0.0052 (2)

U (T) 0.0060 (2) 0.0075 (2) 0.0074 (3) 0.0070 (2) 0.0060 (2) 0.0070 (2) 0.0067 (3) 0.0067 (2) 0.0061 (4)

U (O) 0.0061 (2) 0.0072 (3) 0.0072 (3) 0.0066 (2) 0.0056 (2) 0.0062 (2) 0.0063 (3) 0.0063 (2) 0.0077 (5)

N. refl. 163 148 151 153 163 148 144 163 133

R1 2.09 2.71 2.40 2.54 2.13 2.50 2.78 2.54 3.30

wR2 4.35 4.57 5.21 4.85 3.94 4.81 5.40 5.57 6.82

GooF 1.252 1.256 1.254 1.295 1.452 1.279 1.212 1.315 1.145

Diff. peaks 1.74; -1.25 1.46; -1.32 1.27; -1.30 2.72; -1.09 2.02; -0.53 1.04; -0.96 1.63; -1.51 2.48; -1.09 2.59; -1.70

Sample locality 04-11A

Mining camp

04-11E

Mining camp

04-12A

Mining camp

03-11A

Mining camp

05-13A

Mining camp

05-10A

Mining camp

12-09A

Mining camp

12-12A

Mining camp

04-18A

a0 8.2830 (5) 8.2558 (3) 8.2845 (5) 8.2703 (4) 8.2912 (3) 8.2886 (6) 8.2855 (6) 8.2726 (6) 8.2885 (2)

u 0.2622 (1) 0.2612 (2) 0.2621 (1) 0.2616 (1) 0.2622 (1) 0.2623 (1) 0.2621 (1) 0.2621 (1) 0.2624 (1)

T–O 1.969 (2) 1.948 (2) 1.967 (2) 1.957 (1) 1.971 (2) 1.971 (1) 1.968 (2) 1.964 (2) 1.972 (2)

M–O 1.9746 (7) 1.976 (1) 1.9763 (9) 1.9761 (7) 1.9766 (9) 1.9756 (7) 1.976 (1) 1.9734 (8) 1.9752 (9)

m.a.n.T 16.6 (2) 16.6 (3) 17.9 (2) 16.9 (2) 17.1 (1) 16.7 (2) 16.9 (2) 16.9 (2) 17.3 (2)

m.a.n.M 21.4 (3) 21.1 (4) 21.5 (3) 21.6 (2) 21.5 (2) 21.6 (3) 21.5 (3) 21.5 (3) 21.5 (3)

U (M) 0.0049 (1) 0.0061 (1) 0.0073 (1) 0.0051 (1) 0.0031 (1) 0.0034 (1) 0.0039 (1) 0.0043 (1) 0.0054 (1)

U (T) 0.0070 (3) 0.0071 (3) 0.0090 (2) 0.0064 (2) 0.0062 (3) 0.0052 (2) 0.0061 (3) 0.0062 (3) 0.0074 (3)

U (O) 0.0070 (3) 0.0085 (4) 0.0091 (3) 0.0071 (2) 0.0047 (3) 0.0050 (2) 0.0059 (3) 0.0063 (3) 0.0067 (3)

N. refl. 152 136 162 155 139 167 144 156 154

R1 2.77 2.96 3.06 2.09 3.09 2.49 2.94 2.74 2.93

wR2 5.38 5.77 6.59 4.42 4.29 5.32 5.21 5.70 5.95

GooF 1.301 1.019 1.174 1.211 1.189 1.347 1.158 1.281 1.161

Diff. peaks 3.57; -1.94 1.96; -0.95 1.94; -1.08 1.54; -1.16 2.62; -1.71 2.46; -1.29 2.71; -1.61 3.50; -2.05 1.98; -1.81

a0: cell parameter (A); u: oxygen positional parameter; T–O and M–O: tetrahedral and octahedral bond lengths (A), respectively; m.a.n.T and M: mean atomic

number; U(M), U(T), U(O): displacement parameters for M site, T site and O; N. Refl.: number of unique reflections; R1 all (%), wR2 (%), GooF as defined in

Sheldrick (2008). Diff. peaks: maximum and minimum residual electron density (± e/A3). Space Group: Fd-3 m. Origin fixed at -3 m. Z = 8. Reciprocal space

range: -19 B h B 19; 0 B k B 19; 0 B l B 19. Estimated standard deviations in brackets

958 Page 4 of 17 Contrib Mineral Petrol (2014) 167:958

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Table 2 Chemical analyses, cation distribution and comparison between X-ray and Mossbauer data

Sample 05-15

N.O.

03-18

N.O.

04-25

N.O.

04-05

N.O.

03-12

N.O.

12-01

N.O.

05-13

N.O.

05-10

N.O.

04-11c

P.O.

MgO 13.2 (2) 14.4 (2) 15.6 (1) 16.1 (2) 14.6 (2) 14.5 (1) 14.4 (2) 14.2 (1) 14.7 (2)

Al2O3 22.2 (2) 23.4 (2) 25.3 (4) 23.7 (3) 19.4 (2) 19.1 (3) 12.6 (1) 13.7 (1) 14.1 (3)

TiO2 0.29 (4) 0.39 (2) 0.33 (2) 0.22 (1) 0.25 (1) 0.23 (1) 0.18 (1) 0.20 (2) 0.18 (2)

Cr2O3 45.5 (4) 42.2 (3) 42.5 (3) 44.1 (5) 48.6 (4) 48.4 (4) 56.9 (3) 56.0 (4) 55.3 (4)

MnO 0.27 (3) 0.25 (3) 0.21 (3) 0.18 (2) 0.21 (2) 0.22 (3) 0.22 (3) 0.21 (4) 0.21 (2)

FeOtot 18.9 (2) 19.0 (2) 16.9 (1) 15.65 (7) 17.0 (1) 17.5 (2) 15.6 (2) 15.6 (2) 15.5 (2)

NiO 0.09 (3) 0.13 (3) 0.15 (3) 0.18 (3) 0.16 (3) 0.13 (3) 0.11 (4) 0.13 (3) 0.15 (3)

V2O3 0.12 (4) 0.17 (3) 0.17 (4) 0.18 (3) 0.14 (4) 0.15 (3) 0.13 (3) 0.12 (4) 0.14 (3)

ZnO 0.07 (4) 0.07 (4) 0.06 (4) 0.07 (5) 0.05 (4) 0.06 (3) 0.05 (4) 0.06 (3) 0.05 (4)

Sum 100.6 100.1 101.1 100.4 100.3 100.3 100.1 100.2 100.3

FeO 15.6 (2) 13.8 (2) 12.6 (1) 11.16 (7) 12.8 (1) 13.0 (2) 12.0 (2) 12.4 (2) 11.8 (2)

Fe2O3 3.7 5.8 4.8 5.0 4.7 5.1 4.0 3.6 4.1

T site

Mg 0.576 (6) 0.627 (8) 0.649 (5) 0.691 (7) 0.609 (8) 0.625 (6) 0.642 (6) 0.659 (6) 0.675 (7)

Al 0.0167 (8) 0.0007 (2) 0.027 (2) 0.0038 (5) 0.025 (1) 0.029 (2) 0.0060 (6) 0.0133 (7)

Mn 0.0070 (8) 0.0064 (8) 0.0054 (8) 0.0046 (5) 0.0055 (5) 0.0054 (8) 0.0058 (8) 0.005 (1) 0.0056 (5)

Fe2? 0.362 (5) 0.293 (4) 0.284 (3) 0.238 (3) 0.324 (4) 0.305 (4) 0.289 (5) 0.293 (4) 0.251 (5)

Fe3? 0.038 (4) 0.073 (3) 0.035 (4) 0.062 (5) 0.033 (4) 0.035 (4) 0.056 (5) 0.029 (4) 0.067 (6)

Zn 0.0015 (7)

Vac 0.0039 (4)

M site

Mg 0.022 (1) 0.025 (2) 0.039 (1) 0.024 (1) 0.047 (2) 0.039 (2) 0.036 (2) 0.0084 (6) 0.0065 (7)

Al 0.783 (6) 0.841 (5) 0.858 (9) 0.832 (8) 0.678 (6) 0.666 (8) 0.468 (5) 0.498 (4) 0.529 (9)

Ti 0.0068 (9) 0.0089 (5) 0.0073 (5) 0.0049 (2) 0.0058 (2) 0.0053 (2) 0.0043 (2) 0.0047 (5) 0.0042 (5)

Cr 1.103 (7) 1.016 (6) 0.996 (7) 1.040 (8) 1.183 (8) 1.184 (8) 1.428 (7) 1.401 (7) 1.368 (9)

Fe2? 0.035 (2) 0.052 (2) 0.025 (1) 0.041 (1) 0.0009 (2) 0.025 (1) 0.023 (1) 0.031 (1) 0.054 (2)

Fe3? 0.045 (5) 0.050 (3) 0.066 (5) 0.049 (4) 0.077 (6) 0.073 (5) 0.034 (4) 0.050 (5) 0.029 (4)

Ni 0.0022 (7) 0.0032 (7) 0.0035 (7) 0.0043 (7) 0.0039 (7) 0.0033 (8) 0.003 (1) 0.0032 (8) 0.0038 (8)

V 0.003 (1) 0.0041 (7) 0.004 (1) 0.0044 (7) 0.003 (1) 0.0036 (7) 0.0033 (8) 0.003 (1) 0.0035 (8)

Vac 0.0010 (2)

m.a.n.X-ray 57.1 (7) 56.0 (9) 54.9 (1.1) 55.4 (6) 57.5 (8) 57.4 (7) 59.6 (5) 59.1 (2.4) 58.9 (7)

m.a.n.chem 57.0 56.0 54.9 55.0 57.2 57.3 59.5 59.1 58.9

F(X) 0.09 0.15 0.13 0.27 0.45 0.39 0.29 0.25 0.59

X-Ray data

T Site

%Fe2? 76.2 62.6 69.3 61.0 74.4 69.6 62.5 71.6 62.6

%Fe3? 8.0 15.6 8.5 15.9 7.6 8.0 12.1 7.1 16.7

M Site

%Fe2? 7.4 11.1 6.0 10.5 0.2 5.8 5.0 7.6 13.5

%Fe3? 11.4 10.7 16.1 12.6 17.7 16.6 7.4 12.2 7.2

Fe3?/RFe 0.194 0.263 0.247 0.285 0.254 0.246 0.195 0.193 0.239

Mossbauer data

T site

%Fe2? 73.4 67.0 60.9 57.5 53.8 70.0 60.6 66.1 11.8

%Fe3? 7.3 12.4 7.4 13.7 8.3 5.9 7.4 9.1 54.4

M site

%Fe2? 5.0 8.4 18.4 10.2 18.6 3.8 18.4 13.6 9.6

%Fe3? 14.2 12.2 13.6 18.7 19.3 20.3 13.6 11.1 24.2

Contrib Mineral Petrol (2014) 167:958 Page 5 of 17 958

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Table 2 continued

Sample 05-

15N.O.

03-

18N.O.

04-

25N.O.

04-

05N.O.

03-

12N.O.

12-

01N.O.

05-

13N.O.

05-

10N.O.

04-

11cP.O.

Fe3?/RFe 0.175 0.202 0.199 0.270 0.228 0.216 0.171 0.164 0.740

Sample 04-12

P.O.

04-18

P.O.

12-15

P.O.

04-26

P.O.

12-09

P.O.

04-07

Ox.

04-11

Ox.

03-11A

Ox.

12-12

Ox.

MgO 12.0 (4) 14.1 (1) 14.3 (2) 13.4 (2) 13.7 (1) 10.2 (1) 12.7 (1) 12.5 (5) 13.8 (2)

Al2O3 13.9 (3) 14.3 (3) 24.7 (2) 24.8 (2) 12.1 (1) 11.7 (1) 14.0 (1) 12.8 (3) 13.4 (4)

TiO2 0.18 (2) 0.24 (2) 0.35 (3) 0.28 (2) 0.21 (2) 0.17 (2) 0.19 (1) 0.20 (2) 0.19 (2)

Cr2O3 54.5 (6) 54.5 (4) 42.0 (5) 41.5 (5) 57.7 (4) 58.7 (4) 55.8 (5) 57.4 (4) 56.7 (3)

MnO 0.24 (3) 0.23 (3) 0.24 (2) 0.28 (3) 0.21 (2) 0.22 (3) 0.21 (3) 0.22 (3) 0.20 (4)

FeOtot 19.0 (2) 16.6 (2) 18.6 (1) 20.2 (4) 16.1 (2) 17.0 (2) 15.7 (2) 16.2 (3) 15.3 (1)

NiO 0.07 (4) 0.10 (3) 0.12 (3) 0.10 (3) 0.12 (3) 0.08 (3) 0.13 (3) 0.10 (4) 0.13 (3)

V2O3 0.14 (3) 0.09 (3) 0.14 (3) 0.15 (3) 0.14 (3) 0.15 (4) 0.16 (3) 0.12 (3) 0.13 (2)

ZnO 0.05 (4) 0.05 (3) 0.07 (4) 0.05 (3) 0.04 (3) 0.05 (3) 0.04 (4) 0.05 (4) 0.06 (4)

Sum 99.9 100.2 100.3 100.8 100.3 98.3 98.8 99.6 99.9

FeO 15.9 (2) 12.7 (2) 14.3 (1) 15.7 (4) 13.0 (2) 14.3 (2) 14.7 (3) 12.9 (1)

Fe2O3 3.4 4.2 4.8 5.0 3.35 1.5 1.7 2.7

T site

Mg 0.53 (1) 0.634 (5) 0.611 (7) 0.584 (7) 0.637 (5) 0.486 (6) 0.595 (7) 0.58 (2) 0.642 (4)

Al 0.048 (3) 0.016 (2) 0.0184 (9) 0.0078 (7) 0.0014 (3) 0.0037 (5) 0.0020 (2)

Mn 0.0064 (8) 0.0060 (8) 0.0061 (5) 0.0071 (8) 0.0057 (5) 0.0060 (8) 0.0056 (8) 0.0059 (8) 0.006 (1)

Fe2? 0.371 (5) 0.315 (4) 0.301 (4) 0.330 (8) 0.263 (4) 0.081 (7) 0.130 (8) 0.197 (3)

Fe3? 0.041 (9) 0.029 (3) 0.082 (6) 0.060 (5) 0.080 (4) 0.365 (5) 0.268 (8) 0.222 (7) 0.125 (4)

Zn

Vac 0.007 (3) 0.142 (5) 0.05 (3) 0.061 (4) 0.029 (4)

M site

Mg 0.032 (4) 0.027 (1) 0.026 (1) 0.015 (1) 0.0087 (5) 0.0004 (2) 0.0009 (3) 0.0027 (3)

Al 0.486 (9) 0.517 (9) 0.876 (7) 0.862 (6) 0.446 (5) 0.438 (5) 0.513 (6) 0.47 (1) 0.496 (4)

Ti 0.0043 (5) 0.0057 (5) 0.0079 (7) 0.0063 (5) 0.0050 (5) 0.0041 (5) 0.0045 (5) 0.0047 (5) 0.0045 (5)

Cr 1.38 (1) 1.356 (8) 0.999 (8) 0.987 (9) 1.445 (7) 1.480 (7) 1.387 (9) 1.41 (1) 1.409 (6)

Fe2? 0.051 (2) 0.018 (1) 0.058 (2) 0.012 (3)

Fe3? 0.036 (8) 0.070 (5) 0.026 (3) 0.100 (6) 0.066 (3) 0.044 (2) 0.020 (2) 0.067 (4) 0.059 (3)

Ni 0.0025 (8) 0.0029 (7) 0.0024 (7) 0.0031 (8) 0.0021 (8) 0.0034 (7) 0.003 (1) 0.0033 (8)

V 0.0036 (8) 0.0023 (8) 0.0034 (7) 0.0036 (7) 0.0036 (8) 0.004 (1) 0.0040 (7) 0.0030 (8) 0.0033 (5)

Vac 0.021 (5) 0.021 (4) 0.030 (2) 0.05 (4) 0.034 (3) 0.022 (3)

m.a.n.X-ray 60.4 (7) 59.1 (7) 55.9 (4) 56.3 (5) 59.9 (8) 60.4 (2.1) 59.3 (1.1) 60.0 (6) 59.8 (9)

m.a.n.chem 60.4 59.1 55.6 56.1 59.9 61.7 59.7 60.3 59.8

F(x) 0.96 0.20 0.17 0.86 0.39 0.53 0.21 0.50 0.31

X-Ray data

T site

%Fe2? 74.3 72.9 64.5 67.4 64.3 0.0 21.3 31.0 51.7

%Fe3? 8.2 6.7 17.6 12.2 19.6 89.3 70.3 53.0 32.8

M site

%Fe2? 10.2 4.2 12.4 0.0 0.0 0.0 3.1 0.0 0.0

%Fe3? 7.2 16.2 5.6 20.3 16.1 10.7 5.2 16.0 15.5

Fe3?/RFe 0.154 0.229 0.231 0.326 0.357 1.000 0.756 0.690 0.483

Mossbauer data

T site

%Fe2? 11.6 10.8 23.3 66.0 48.4 0.0 13.0 23.0 29.6

958 Page 6 of 17 Contrib Mineral Petrol (2014) 167:958

123

same cr#, but the smaller grains are enriched in fe# relative

to the larger grains. The small grains are unaltered, but the

larger grains have patches of alteration up to about 1 mm

across in which the chromite is irregularly altered to lenses

of ferritchromite about 10 lm wide and 100 lm long.

Samples 03-11, 05-13, 05-10, 12-09 and 12-12 are from

the main abandoned pit. The chromite forms large grains up

to 6 mm in diameter intergrown with olivine and smaller

grains \1 mm in diameter. Olivine is the main interstitial

phase, and olivine inclusions are common in the chromite. In

sample 03-11 (60 % chromite, cr# = 0.75), some grains

show occasional alteration to ferritchromite along fine

fractures which are up to about 5 lm wide and on grain

boundaries in which the cr# increases to 0.82 (Fig. 2c).

Sample 05-10 contains 90 % chromite (cr# = 0.73–0.75)

some of which is altered at grain boundaries to ferritchr-

omite. In these samples, the cr# increases from ca. 0.74 to

0.83 and the fe# from 0.35 to 0.5 in the alteration zones.

Olivine is present as the interstitial phase (Fo96) and in places

is altered to chlorite. Sample 05-13 (90 % chromite, cr#

0.74–0.75) also contains olivine (Fo96–97), and there are

olivine inclusions in the chromite. There is some evidence of

slight alteration of the chromite to ferritchromite along grain

boundaries. Sample 12-09 contains chromite (cr# 0.73–0.77)

with inclusions of clinopyroxene (mg# = 0.96) and calcic

amphibole and interstitial olivine (Fo95–96). The chromite

shows a small amount of oxidation along fractures and at

grain boundaries.

Sample 04-18 contains 75 % chromite (cr# =

0.71–0.72), which is from a banded chromitite–dunite body

cut by serpentinized dunitic veins located to the west of the

main pit. The chromite forms subhedral grains 1 mm in

diameter intergrown with rounded olivine grains up to

2 mm in diameter. The chromite grains contain olivine

inclusions in their cores. Some grains show a small amount

of grain boundary alteration. In the altered regions, the cr#

increases to 0.81 and the fe# from 0.35 to 0.40.

Full petrographic and mineral-chemical descriptions are

given in Rollinson (2008).

Experimental procedures

In this study, we examined eighteen samples of chromitite

from the mantle section of the Oman ophiolite in Wadi

Rajmi, northern Oman. Some of the samples were previ-

ously analyzed by electron probe and Mossbauer spec-

troscopy (Rollinson et al. 2012; Rollinson and Adetunji

2013a). Here, we analyzed the spinels by using X-ray

diffraction and reanalyzed those particular grains with the

electron microprobe. In addition, we refitted out previous

results from the Mossbauer spectroscopy in light of the

new crystal-chemical data obtained, as discussed below.

The data are presented in Tables 1 and 2.

X-ray diffraction

X-ray diffraction data were recorded using an automated

KUMA-KM4 (K-geometry) diffractometer, using MoKaradiation, monochromatized by a flat graphite crystal. Data

were collected, according to the method of Della Giusta

et al. (1996), with up to 55� of 2h in the x-2h scan mode

(scan width 1.8� 2h, with counting times from 20 to 50 s,

depending on the peak standard deviation). Twenty-four

equivalent reflections of (12 8 4) or (8 4 4) peaks

(according to the size of the Cr-spinel), at about 80� or 50�of 2h, respectively, were accurately centered at both sides

of 2h, and the a1 peak barycenter was used for cell

parameter determination. Corrections for absorption were

performed according to the procedures of North et al.

(1968). Structural refinement using the SHELX-97 pro-

gram (Sheldrick 2008) was carried out against Fohkl2 in the

Fd-3 m space group (with the origin at -3 m), since no

evidence of different symmetry appeared. Scattering

Table 2 continued

Sample 04-12P.O. 04-18P.O. 12-15P.O. 04-26P.O. 12-09P.O. 04-07Ox. 04-11Ox. 03-

11AOx.

12-12Ox.

%Fe3? 29.6 2.3 6.6 10.3 24.4 86.1 49.6 43.5 33.2

M site

%Fe2? 17.6 29.0 23.9 2.1 5.5 0.0 11.0 7.6 12.3

%Fe3? 41.2 57.4 46.2 21.5 21.6 14.0 26.4 25.9 24.9

Fe3?/RFe 0.653 0.539 0.464 0.266 0.399 1.000 0.710 0.637 0.518

Mean chemical analyses (10–15 spot analyses for each crystal) and cation distribution in T and M site of the analyzed Cr-spinels on the basis of

four oxygen atoms per formula unit. N.O.: not oxidized; P.O.: partially oxidized; Ox.: oxidized. m.a.n.X-ray and m.a.n.chem: X-ray and chemistry

mean atomic number, respectively. F(x): minimization factor, which takes into account the mean of square differences between calculated and

observed parameters, divided by their standard deviations. Estimated standard deviations are in brackets

Contrib Mineral Petrol (2014) 167:958 Page 7 of 17 958

123

factors were taken from Prince (2004) and Tokonami

(1965). The crystallographic data are presented in Table 1.

Electron microprobe analysis

Ten to fifteen spot analyses were performed on the same

Cr-spinels as were used for X-ray diffraction analysis,

using a CAMECA-CAMEBAX electron microprobe oper-

ating at 15 kV and 15 nA. A 20 s counting time was used

for both peak and total background. Synthetic oxide stan-

dards (MgO, FeO, MnO, ZnO, NiO, Al2O3, Cr2O3 and

TiO2) were used. Raw data were reduced by PAP-type

correction software provided by CAMECA. The mineral-

chemical analyses are reported in Table 2.

Mossbauer spectroscopy

Room-temperature Mossbauer measurements were made

on clean chromite separates with a 25 mCi 57Co source in

an Rh matrix, which was driven at constant acceleration in

a triangular mode. About 50 mg of sample was used, dis-

tributed as 28 mg/cm2 in the sample holder. The spectra

were recorded in 1,024 channels, and the Lorentzian lines

of the folded data were fitted, using the least-square

RECOIL 1.04 computer program developed by Lagarec

and Rancourt (1998). All centroid shift (CS) values are

relative to a-iron. The best fits were obtained by reduced

v2. The uncertainties were calculated using the covariance

matrix. Errors were estimated at about ±0.002 and

0.02 mm/s for isomer shift and quadrupole splitting DEQ,

respectively. A correction for the difference between the

recoil-free fractions for Fe2? and Fe3? at room temperature

was applied (Quintiliani 2005).

In previous studies, we have optimized the

Mossbauer spectra using three doublets. In this study,

the spectra were fitted with four doublets:

Fe2?(T) ? Fe2?(M) ? Fe3?(T) and Fe3?(M) in order to

be consistent with predictions of the site occupancy cal-

culations reported below. The criterion for identification of

Fe-cation sites was based on the relative centroid shift

values, following previous reports by Pal et al. (1994), Rais

et al. (2003) and Rollinson et al. (2012): [dFe3?T] \ [d-Fe3?M] \ [dFe2?T] \ [dFe2?M]. This model was also

found to be appropriate in characterizing natural chromite

samples from ophiolite complexes in the Philippines by

room-temperature Mossbauer spectroscopy (Kuno et al.

2000).

Site occupancy calculations

The distribution of cations between the T and M sites given

in Table 2 was obtained using the method of Carbonin

et al. (1996) and Lavina et al. (2002), in which crystal-

chemical parameters are calculated as a function of the

atomic fractions at the two sites and fitted to the observed

values obtained during the single-crystal X-ray measure-

ments. Site atomic fractions were calculated by minimizing

the function F(x) (Table 2), which takes into account the

mean of the square differences between calculated and

observed parameters from X-ray measurements divided by

their squared standard deviations.

One of the objectives of this study is to better constrain

the Fe3?/RFe ratios in these mantle chromitites, given that

sometimes different solutions to the Mossbauer fitting give

different ratios. In this study, we have the opportunity of

validating the results of Mossbauer spectroscopy with

solutions from the X-ray diffraction data for site occu-

pancies, and vice versa. Thus, we used Mossbauer data for

Fe3? in the site occupancy calculations and then refitted the

Mossbauer data accordingly, iterating this process until a

close agreement was reached between the Mossbauer fit-

ting model (constrained by v2) and the site occupancy

calculations [constrained by F(x)].

Results

Mineral chemistry

The details of compositional variations in the Wadi Rajmi

chromitites were discussed in Rollinson (2008). The prin-

cipal feature is the variation of the cr# [Cr/(Cr ? Al)] with

depth. Chromitites found close to the Moho have lower cr#

Fig. 3 Cr–Fe3?–Al triangular diagram for Oman mantle chromites.

The green circles represent the sample compositions in which Fe3?

was calculated through mineral stoichiometry, and the red spots are

those calculated using Fe3? obtained through Mossbauer and single-

crystal refinement. Fields are from Saumur and Hattori (2013) and

Evans and Frost (1975). Isothermal section at 600 �C for (Fe,Mg)-

Cr2O4-(Fe,Mg)Al2O4-(Fe,Mg)Fe2O4 coexisting with Fo95 olivine is

shown in yellow (Sack and Ghiorso 1991)

958 Page 8 of 17 Contrib Mineral Petrol (2014) 167:958

123

(0.5–0.6) than those found deeper in the mantle section

([0.70). The chromites are magnesian with mg# between

0.686 and 0.611. Small concentrations of the trace elements

Mn, Zn, Ni and V are recorded, but there are no systematic

variations in composition. Inter-crystalline chemical vari-

ations were discussed in Rollinson et al. (2012) and Roll-

inson and Adetunji (2013a). These studies showed a

positive correlation between Ti (apfu) and Fe3? (apfu)

suggesting that some of the variation in Fe3? is magmatic

and relates to variations in the compositions of melts

parental to the chromite.

We have investigated the possible effects of the sub-

solidus alteration of our spinels in two ways. Firstly, we

have plotted our samples on a Cr–Fe3?–Al triangular dia-

gram and show that they plot far away from the ferritchr-

omite field close to the Cr–Al join (Fig. 3). Thus, the bulk

compositions are very different from those of ferritchr-

omite. Secondly, we calculated olivine–chromite Fe–Mg

exchange temperatures using the method of Ballhaus et al.

(1991). Our results show a range of values between 675

and 861 �C, consistent with their position on the Cr–Fe3?–

Al triangular diagram, above the 600 �C solvus for chro-

mite in equilibrium with Fo95 (Sack and Ghiorso, 1991).

However, two of the highly oxidized samples which

coexist with olivine (04-11f and 03-11) show anomalously

high temperatures which indicate that they are not in

equilibrium with olivine. These chromites have lower

Fe2?/Mg ratios than expected which we attribute to the

oxidation of Fe2?–Fe3?. This allows us to constrain the

temperature of oxidation. Since the oxidized samples are

not in equilibrium with olivine, we propose that the oxi-

dation disturbed the olivine–chromite equilibration

(675–861 �C) and so took place at a lower temperature, but

not at a temperature at which ferritchromite forms.

In the samples which are not oxidized, the diffusion of

Mg from chromite to olivine during cooling explains the

very high forsterite content of the olivines. There are weak

positive correlations between the Fe–Mg exchange tem-

peratures and the cr# of the spinel and the Fe3?/RFeMoss

ratio, indicating that there may be a weak compositional

control on the diffusion process (Fig. 4).

X-ray diffraction

In the spinel structure, anions form a near-cubic close-

packed array, parallel to (111) planes, and the cations fill part

of the tetrahedral (T) and octahedral (M) interstices in the

framework. In normal spinels, of which chromite is one,

trivalent cations are found in the M site and divalent cations

in the T site. Oxygen atoms are linked to three octahedral

cations and one tetrahedral cation lying on opposite sides of

the oxygen layer. Variations in the relative effective radii of

cations in the tetrahedral and octahedral sites give rise to the

displacement of the oxygen atom along the cubic diagonal

[111] causing the oxygen layers in the spinel structure to be

slightly puckered. This in turn gives rise to variations in the

oxygen positional parameter (u) which we measure. Previous

studies of oxidation mechanisms in Cr-spinels using crystal

structural parameters and mineral chemistry from a variety

of geological settings (Carbonin et al. 1999; Bosi et al. 2004;

Lenaz et al. 2004a; Derbyshire et al. 2013 among the others)

show that, as a result of oxidation, Fe3? is present on both T

and M sites in the chromite lattice and that the measured

oxygen positional parameter and the cell edge have lower

Fig. 4 Olivine–chromite Fe–Mg exchange temperature (after Ballhaus et al. 1991) versus cr# and Fe3?/RFe

Contrib Mineral Petrol (2014) 167:958 Page 9 of 17 958

123

values than in non-oxidized spinels because of the formation

of a magnetite–magnesioferrite component. For example,

Menegazzo et al. (1997) showed in experimentally oxidized

natural spinels that their u values decreased from about

0.2635 to about 0.2608.

Single crystals of Cr-bearing spinels were analyzed by

X-ray diffraction and their cell edge, a0, and oxygen

positional parameter, u, calculated. Cell edges are in the

range 8.2405 (5) to 8.2912 (3) A, while u values are

between 0.2608 (2) and 0.2629 (1). There is a strong

Fig. 5 a Cr# versus cell edge.

Red spots Oman samples; blue

spots Albania samples (Bosi

et al. 2004). Oxidized spinels

from Oman and Albanian

ophiolites with more than 0.02

atoms per formula units of

vacancies are represented as red

and blue diamonds,

respectively. For comparison

are also represented, yellow

spots, Fe-rich Cr-spinels from

the Rum layered complex

(Lenaz et al. 2011). b Cr#

versus oxygen positional

parameter. Symbols as in

Fig. 4a. c Oxygen positional

parameter, u, versus cell edge,

a0. Symbols as in Fig. 4a.

Yellow diamond oxidized spinel

from terra rossa soil (Carbonin

et al. 1999)

958 Page 10 of 17 Contrib Mineral Petrol (2014) 167:958

123

correlation between the cell edge a0 and the compositional

parameter cr#, and most samples lie on a well-defined

correlation line with an R2 of 0.96 (Fig. 5a). The more Cr-

rich samples have the higher a0 values. However, four

samples with high cr# have lower a0 values than might be

predicted on the basis of their Cr/Al ratios alone and define

a discrete group of samples. This grouping is also apparent

on a u versus cr# (Fig. 5b) and a u versus a0 plot where

there is a main group of samples which define a clear trend

of increasing u with decreasing a0, with an R2 of 0.84

(Fig. 5c) and a smaller group of samples with lower

u values than predicted from their cell edge value.

A similar distribution of u versus a0 and cr# versus a is

seen in samples from the Shetland ophiolite (Derbyshire

et al. 2013) where samples plot along the main u vs a0 and

cr# vs a0 trends illustrated above (Fig. 5c). Samples from

the Albanian ophiolites (Bosi et al. 2004) are also distrib-

uted both on the main chromite trend but also some sam-

ples plot at lower a values. Cr-spinels from the Rum

layered intrusion (Lenaz et al. 2011) plot close to, but do

not overlap, the ophiolite spinel trend for they have a

greater cell edge value and a smaller u value. These sam-

ples have a higher Fe3? content and ferric iron forms a

larger cation than Al and Cr—the tetrahedral bond distance

for Fe3? is between 1.865 and 1.875 A, whereas Al is

between 1.77 and 1.774 A. Similarly, the octahedral bond

distance is 2.025 A for Fe3?, 1.995 A for Cr and between

1.908 and 1.915 A for Al (Shannon 1976; O’Neill and

Navrotsky 1984; Lavina et al. 2002; Lenaz et al. 2004b).

Mossbauer spectroscopy

It is possible to resolve Mossbauer spectra with a variety of

fitting procedures. Rollinson et al. (2012) and Rollinson

and Adetunji (2013a, b) used two different three-doublet

models which considered Fe3? in the M site, and Fe2? in

two different tetrahedral sites and Fe3? in both the T and M

sites and Fe2? also in the M site (see also Rais et al. 2003).

A five-doublet model with three tetrahedral Fe2?, one

octahedral Fe2? and one octahedral Fe3? contribution was

Fig. 6 Mossbauer spectra, each fitted with four doublets, for samples from a highly oxidized zone (04-11), a low oxidized zone and partially

oxidized zones (12-09 and 12-12)

Contrib Mineral Petrol (2014) 167:958 Page 11 of 17 958

123

used by Halenius et al. (2002) for synthetic spinels, and

Lenaz et al. (2004a) successfully adopted a six-doublet

model with four doublets representing the broad Fe2?

contribution (three Fe2? in the T site and one in the M site)

and two doublets representing Fe3? (one in the T and one

in the M site) in natural chromites from Indian komatiites.

In this study, we have adopted a four-doublet fitting

model in order to allow us to compare between our struc-

tural data distributed over four sites and our Mossbauer

data. We have therefore refitted our previously measured

Mossbauer data. This new four-doublet model has the

following range of Mossbauer parameters: Fe2?(T) with

CS of 0.807–0.951 mm/s and quadrupole splittings (QS) of

1.337–1.956 mm/s; Fe2?(M) with CS of 0.937–1.677 mm/

s and QS of 0.346–0.576 mm/s; Fe3?(T) with CS of

0.132–0.517 mm/s and QS of 0.702–1.7426 mm/s; for

Fe3?(B), CS of 0.247–0.407 and QS of 0.410–0.714 mm/s.

Typical spectra are illustrated in Fig. 6.

Site occupancy calculations

Site occupancy calculations allocate the divalent ions Mg,

Mn and Fe2? to the T site, together with a very small

amount of the trivalent ion Al (0.000–0.029 apfu). In

addition, variable amounts of Fe3? (0.03–0.365) are pres-

ent in the T site. Six samples show vacancies in the T site.

The trivalent ions Cr, V and the majority of the Al are

allocated to the M site. In addition, Ti and Ni are allocated

to the M site, together with a small amount of the Mg

(0.000–0.039 apfu) and Fe2? (0.000–0.058 apfu). Some of

the Fe3? is also allocated to the M site, and seven samples

have vacancies on the M site. The total vacancies across

the T and M sites range from 0.005 to 0.172 apfu

(Table 2). Those samples with vacancies have lower oxy-

gen positional parameters (u) than the main data set

(Fig. 7) and show a reduction in the size of the cell edge

(a0) (Fig. 7), properties previously noted by Bosi et al.

(2004) in their study of chromitites from the Albanian

ophiolite. In most samples, the majority of the Fe2? is

located on the T site whereas there is no clear pattern for

the allocation of Fe3? between the T and M sites.

Fe31/RFe ratio calculations

Even if Mossbauer spectroscopy on powder absorbers

requires a relatively large amount of sample material, a

comparison between the Fe3?/Fetot ratio obtained by using

this technique and single-crystal X-ray diffraction has been

commonly used (Carbonin et al. 1996; Bosi et al. 2004;

Lenaz et al. 2004a, 2013; Perinelli et al. 2014) because they

are usually in good accordance. Only recently, Lenaz et al.

(2013) showed for Cr-spinels from kimberlites that there

were large differences in Fe3?/Fetot ratio calculated by these

methodologies and point source Mossbauer, too, due to a

large internal variability in Fe3? compositions, while single-

crystal diffraction and point source Mossbauer gave better

results.

In Fig. 8, we show a comparison between Fe3?/RFe ratios

calculated from mineral stoichiometry and those calculated

from site occupancies using single-crystal X-ray data and

from Mossbauer measurements. Figure 8 shows that there is

good agreement for 12 out the 18 samples between the site

occupancy data and stoichiometric calculations. The site

occupancy data for six samples show that they are more

oxidized (higher Fe3?/RFe ratios) than are predicted by

stoichiometry. These include the four samples 04-07, 04-11f,

03-11 and 12-12 which are structurally anomalous (Fig. 5c)

and which show vacancies on both the T and M sites. These

samples have a relatively high proportion (33–89 %) of their

Fe as Fe3? on the T site and a very small amount of Fe2? on

the M site (0–3 %). They are located in the deep mantle

Fig. 7 Total vacancies versus

oxygen positional parameter,

u. Errors associated with each

technique are lower than 5 %

meaning that error bars are

within the symbols. Symbols as

in Fig. 5

958 Page 12 of 17 Contrib Mineral Petrol (2014) 167:958

123

section close to the boundary with the high-temperature

shear zone in the mantle harzburgites. In addition, two other

samples are oxidized—12-09 from the deep mantle section

and 04-26 from close to the Moho. These samples also show

some vacancies. They have no Fe2? on the M site, but

16–20 % of the Fe present is as Fe3? on the M site. Those

samples with high Fe3? are shown on the Cr–Fe3?–Al dia-

gram in Fig. 3. A comparison between the results of stoi-

chiometric calculations and Mossbauer measurements

shows that 10 out of the 18 samples have Mossbauer -Fe3?/

RFe ratios higher than those calculated from stoichiometry

(Fig. 8) and that the remaining 8 samples have ratios which

define a trend close to but lower than the values predicted

from mineral stoichiometry. The more oxidized samples

include those identified from the X-ray/site occupancy

measurements described above but also another group of

samples which includes 04-18, 04-11c, 04-12 from the deep

mantle and sample 12-15 from close to the Moho.

Our results indicate therefore that Fe3? calculations

based upon the assumption of stoichiometry in these

chrome spinels are inaccurate. Site occupancy calculations

show that six out the 18 samples are more oxidized than

predicted from stoichiometry. Our refined Mossbauer

measurements show virtually no agreement with stoichi-

ometric calculations with many samples more oxidized

than predicted from stoichiometry and a smaller number of

samples less oxidized.

As discussed above, we have sought to optimize the parti-

tioning of Fe into Fe2? and Fe3? using the combined results of

Mossbauer spectroscopy and of X-ray diffraction. Our results,

expressed as Fe3?/RFe ratios, show a good agreement between

the two methods for 14 out of our 18 samples (Fig. 8). This

agreement exists over the range of cr# studied and over the full

range of oxidation ratios observed. However, as noted above,

there is a small group of samples (04-11c, 04-18 and 04-12 from

the deep mantle and 12-15 from Maharra near to the Moho), for

which the results of Mossbauer spectroscopy fitted with four

doublets and of X-ray diffraction do not agree (see Table 2). In

this case, the Mossbauer Fe3?/RFe ratios imply that the samples

are oxidized whereas the X-ray diffraction site occupancy data

do not. If the Mossbauer data are fitted with three doublets, i.e.,

if no Fe2? is allocated in the M site, then the agreement greatly

improves for samples 12-15 and 04-18. However, there still

remains a discrepancy for samples 04-11c where the Fe3?/RFe

ratioMoss = 0.740 and the Fe3?/RFe ratioX-Ray = 0.239 and

04-12 where the Fe3?/RFe ratioMoss = 0.653 and the Fe3?/RFe

ratioX-Ray = 0.154. Both these samples were collected from the

oxidized zone of Wadi Rajmi (see Fig. 1). Sample 04-11 is of

particular interest because our results show that for finer grain

sizes (sample 04-11f = 180–250 lm), the chromite is highly

oxidized and there is good agreement between both the X-Ray

and Mossbauer methods, whereas for coarser grained samples

(sample 04-11c = 250–315 lm), the sample would appear to

be oxidized from the Mossbauer data, but not from the X-ray

data. There are also significant differences from the main

sample suite in the u and a0 parameters measured for these two

samples (Table 1).

Discussion

Our results indicate that on the basis of our precise Fe3?/

RFe ratio measurements, there are three groups of chro-

mitites in the Wadi Rajmi mantle section:

Fig. 8 a a comparison of Fe3?/RFe ratios calculated using single-

crystal X-ray diffraction versus those measured through mineral

stoichiometry; b a comparison of Fe3?/RFe ratios calculated through

Mossbauer spectroscopy versus those measured using mineral stoi-

chiometry; c a comparison of Fe3?/RFe ratios measured using

Mossbauer spectroscopy versus those measured using single-crystal

X-ray diffraction. A total of 5 % error bars are indicated, but usually

the errors are lower than this value (within the symbol). Symbols as in

Fig. 5

Contrib Mineral Petrol (2014) 167:958 Page 13 of 17 958

123

• highly oxidized samples. These have low a0 and

u values for their measured cr#, and which are

anomalous both in the X-ray site occupancy data and

in the Mossbauer data with respect to their Fe3?/RFe

ratios (samples 04-07; 04-11f; 03-11; 12-12);

• partially oxidized samples which are either, anomalous

both in the X-ray site occupancy data and Mossbauer

data, but which are not distinctive with respect to their

lattice parameters (samples 04-26; 12-09), or, those

which appear to be oxidized in the Mossbauer data set

but which are not apparent in the X-ray site occupancy

data set (samples 04-11c; 04-12; 04-18; 12-15);

• samples which are probably not oxidized and thus

represent the original oxidation state of the chromitites

(samples 05-15; 03-18; 04-25; 04-05; 03-12; 12-01;

05-13; 05-10). Fe3?/RFe ratios for these samples are

between 0.164 and 0.285.

Samples in which there is a disparity between the

Mossbauer and X-ray results may represent heterogeneity

within the sample, for the X-ray results are obtained on

single crystals (up to 200 lm in size), whereas the

Mossbauer data are obtained from many grains (ca

50 mg). Electron microprobe studies also show a measure-

able difference in composition with grain size; the smaller

grains have higher Fe/(Fe ? Mg) ratios than larger grains.

The Mossbauer data for 04-11c show very high amount of

Fe3? in T site which should cause a decrease of the oxygen

positional parameter. This is not seen, supporting the

suggestion that the sample is not homogeneously oxidized.

Oxidation mechanisms: crystal-chemical processes

The data for our least oxidized samples show that most of

the Fe is present as Fe2? and is located on the T site. In

contrast in the most oxidized samples, Fe2? on the T site is

largely replaced by Fe3?. A similar oxidation of Fe2? to

Fe3? is also seen on the M site, although the proportion of

total Fe is smaller. We also find vacancies between 0.005

and 0.172 apfu in the most oxidized samples, which

according to our best-fitting procedures are randomly dis-

tributed between T and M sites. These observations are

consistent with previous studies of oxidation mechanisms in

Cr-spinels using crystal structural parameters and mineral

chemistry from a variety of geological settings which show

that, as a result of oxidation, Fe3? is present on both T and

M sites in the chromite lattice (Fig. 6) (Carbonin et al. 1999;

Bosi et al. 2004; Derbyshire et al. 2013; Lenaz et al. 2004a).

This is also consistent with previous thermodynamic mod-

els for the spinel structure which distribute vacancies ran-

domly across both the octahedral and tetrahedral sites

(Mattioli and Wood 1988; Ghiorso and Sack 1991). The

measured oxygen positional parameter and the cell edge in

these oxidized samples have lower values than in non-oxi-

dized spinels because of the formation of a magnetite–

magnesioferrite component (Menegazzo et al. 1997).

The four most oxidized samples, in which there is the

highest number of structural vacancies, are all high-cr#

chromitites from the deeper part of the mantle section in

Wadi Rajmi. They have cr# [ 0.7 and are thought to have

crystallized from melts of boninitic affinity (Rollinson

2008). This study suggests that the high cr# promotes the

formation of structural vacancies during oxidation, possi-

bly because in these samples, the Fe–chromite component

is abundant, so that it is easier to oxidize Fe2? and create

vacancies.

Temperature of oxidation

We argued above that the oxidation process post-dates the

cessation of Fe–Mg diffusion between olivine and chro-

mite, so at temperatures below 675–861 �C. We find no

evidence to indicate that our oxidized chromites are fer-

ritchromite in composition for they do not plot in the field

of ferritchromite on a Fe3?–Cr–Al triangular diagram

(Fig. 3). Further, while some ferritchromite alteration is

visible on electron backscatter images and is typically

developed in very small quantities at grain boundaries and

along fractures, we note that

• the volume of altered chromite (ferritchromite) is very

small relative to the total amount of chromite present.

• separate peaks for the phase ferritchromite are not seen

on the Mossbauer spectra, further confirming its low

volume.

• ferritchromite alteration is not confined to those

samples which show a high Fe3? content but is found

in all samples–oxidized, partially oxidized and not

oxidized.

For these reasons, we conclude that ferritchromite is a

minor phase and is not the major component of the

oxidized samples described here. It is possible that our

samples preserve oxidation–exsolution at the nanoscale,

although we have not detected this in our SEM imaging, in

our X-ray data or in our Mossbauer data.

Thus, the temperature of oxidation is bracketed between

the olivine–chromite Fe–Mg diffusion temperatures

(675–861 �C) and the temperature at which ferritchromite

forms (between 500 and 600 �C, Evans and Frost 1975;

Kimball 1990; Mellini et al. 2005; Gonzalez-Jimenez et al.

2009), so oxidation took place between 600 and 700 �C.

Our site occupancy data suggest that the oxidized

component in chromite has a formula similar to that sug-

gested by Gillot et al. (1981), i.e., ðFe3þ0:9;h0:1ÞðCr1:8

Fe3þ0:2ÞO4. This component makes up about 30 % of the

958 Page 14 of 17 Contrib Mineral Petrol (2014) 167:958

123

spinel phase, although since we only see one phase in the

crystallographic data, we conclude that it is in solid solu-

tion with spinel and magnesiochromite.

The oxidized samples in this study come from the

margin of a shear zone within the mantle harzburgites. The

shear zone comprises a central zone about 2 km wide of

porphyroclastic harzburgite showing significant realign-

ment of olivine grains, flanked by a zone of medium

grained granular harzburgites about 1 km wide (Michi-

bayashi et al. 2006). Our oxidized samples are located on

the northeastern margin of the shear zone just outside the

granular harzburgite domain containing high-temperature

foliations. Grain-scale compositional re-equilibration of

minerals during recovery is reported for chromite (Ghosh

and Konar 2012; Ghosh et al. 2014). Thus, the localized

oxidation of spinel may be linked to deformation, and it is

possible that this oxidation indicates the presence of a

hydrous fluid in association with the shearing.

Robust Fe3?/RFe ratios for Oman mantle chromitites

We have shown above that mineral stoichiometry alone is

incapable of identifying those samples which have expe-

rienced late oxidation. Using the dual approach adopted

here for determining Fe3?/RFe ratios, we have identified

those chromitites which are oxidized. We have removed

these from our data set leaving a suite of samples which are

the least oxidized and which we conclude represent pri-

mary magmatic Fe3?/RFe ratios for the Oman mantle

chromitites. Our X-ray results suggest ratios between 0.193

and 0.285 and our Mossbauer data between 0.164 and

0.270. Whichever methodology is used, the chromitite suite

preserves a range of Fe3?/RFe ratios of 0.09–0.11. We

propose that this variability is real and reflects differences

in the composition of the magmas parental to the chromi-

tites. The X-ray data suggest a weak negative correlation

between Fe3?/RFe ratio and cr# although this is not seen in

the Mossbauer data. If this is real, it suggests that the high-

cr# chromitites of boninitic parentage may have slightly

lower Fe3?/RFe ratios than the lower cr# samples.

The values observed here for Fe3?/RFe ratios in chro-

mite extend from values cited for MORB melts

(0.16 ± 0.1) to those for arc basalts (0.22–0.28) (Fig. 9).

This confirms the findings of previous studies (Rollinson

et al. 2012; Rollinson and Adetunji 2013a, b), but this

refined data set using combined X-ray and Mossbauer

measurements (Fe3?/RFe = 0.164–0.285) suggests lower

Fe3?/RFe ratios than previously measured using the

Mossbauer method alone (0.229–0.350—Rollinson et al.

2012). Thus, we propose that the data presented here for

the least oxidized chromitites are better determined and so

more robust than in our previous studies. Given that most

samples have Fe3?/RFe ratios higher than typically found

in MORB, we suggest that an arc setting is more appro-

priate for these rocks.

Conclusions

• We present a comparison between Fe3?/RFe ratios

measured in chromites from mantle chromitites in the

Oman ophiolite using Mossbauer spectroscopy and

single-crystal X-ray diffraction with ratios calculated

from mineral stoichiometry. Our results show that

mineral stoichiometry does not accurately reflect the

measured Fe3?/RFe ratios.

• There are three groups of samples. The majority

preserve magmatic Fe3?/RFe ratios, whereas a few

are highly oxidized and have high Fe3?/RFe ratios.

There is also a group of partially oxidized samples

• The more oxidized chromites show oxygen positional

parameters among the lowest ever found for chromites.

• Site occupancy calculations using X-ray, electron

microprobe and Mossbauer analyses show that some

chromites are non-stoichiometric and contain vacancies

in their structure randomly distributed between both the

T and M sites. The presence of vacancies lowers the

cell edge of the chromite crystals.

• It is proposed that the oxidation of the magmatic

chromitites took place at temperatures above those

necessary for the formation of ferritchromite but below

about 700 �C, in association with a ductile shear zone

in mantle harzburgites.

• Primary magmatic Fe3?/RFe ratios measured for the

Oman mantle chromitites are between 0.193 and 0.285

Fig. 9 cr# versus Fe3?/RFe ratio plot for Wadi Rajmi chromitites.

Black diamonds are from X-Ray data, red squares from Mossbauer

data. Tie lines are shown between the Mossbauer and X-ray data for

the same sample. Results are plotted relative to the range of Fe3?/RFe

ratios in MORB and arc basalts (after Cottrell and Kelley, 2011 and

Kelley and Cottrell, 2012, respectively)

Contrib Mineral Petrol (2014) 167:958 Page 15 of 17 958

123

(X-ray data) and 0.164–0.270 (Mossbauer data) and

preserve a range of Fe3?/RFe ratios which we propose

is real and reflects differences in the composition of the

magmas parental to the chromitites. The range of

values extends from those MORB melts (0.16 ± 0.1) to

those for arc basalts (0.22–0.28)

• We observe that non-stoichiometric chromites and Cr-

bearing spinels are present in both ophiolites, layered

complexes and mantle xenoliths (Nell and Pollak 1998;

Quintiliani 2005; Quintiliani et al. 2006; Rollinson

et al. 2012; Adetunji et al. 2013; Rollinson and

Adetunji 2013a, b; Perinelli et al. 2014).

Acknowledgments The Italian C.N.R. financed the installation and

maintenance of the microprobe laboratory in Padova. R. Carampin

and L. Tauro are kindly acknowledged for technical support. DL

would like to thank the FRA2009 and PRIN 2010-11 funds. HRR and

JA are supported by research grants from the University of Derby and

by the Nuffield foundation. We thank Chris Ballhaus and B. Ronald

Frost for their helpful comments on an earlier draft of this paper.

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