Multiple Stages of Carbonation and Element Redistribution during Formation of Ultramafic-Hosted Magnesite in Neoproterozoic Ophiolites of the Arabian-Nubian Shield, Egypt

Multiple Stages of Carbonation and Element Redistribution during

Formation of Ultrama

fi

c-Hosted Magnesite in Neoproterozoic

Ophiolites of the Arabian-Nubian Shield, Egypt

Mokhles K. Azer,

Hisham A. Gahlan,

Paul D. Asimow,

Heba S. Mubarak,

and Khaled M. Al-Kahtany

1. Geological Sciences Department, National Research Centre, Cairo, Egypt; 2. Department of Geology and

Geophysics, King Saud University, Riyadh 11451, Saudi Arabia; 3. Geology Department, Assiut University,

Assiut 71516, Egypt; 4. Division of Geological and Planetary Sciences, California Institute

of Technology, Pasadena, California 91125, USA

ABSTRACT

We present a study of the serpentinized peridotites of the Ghadir-Mohagar-Ambaut area, Egypt. They represent the

mantle section of a dismembered ophiolite, tectonically emplaced over a volcanosedimentary succession of island arc

assemblages. The serpentinites are variably metamorphosed from greenschist to lower-amphibolite facies, metasomatized,

and altered, including development of talc-carbonate and quartz-carbonate rocks, especially along shear zones and fault

planes. Nevertheless, some samples contain relics of primary chromian spinel, olivine, and pyroxenes. Relict textures and

whole-rock compositions (Mg#

½

molar

(Mg

1

1

)

p

–

93, with low Al

and CaO contents) both suggest

harzburgite protoliths. The high Mg# and Ni contents of relict olivine and the high Cr# (molar

(Cr

1

Al)) of fresh

chromian spinel cores indicate that the protoliths experienced high degrees of partial melt extraction (

∼

34%

–

39%), well

beyond the limit for exhaustion of clinopyr

oxene from the residue and consistent with

formation in a forearc suprasubduction

zone environment. The serpentinizedultrama

fi

c rocksin the studyareaare divided into massiveserpentinite, serpentinite-

hosted magnesite masses, and magnesite-

fi

lled veins. The carbonation and formation of magnesite ores took place through

two metasomatic stages; the

fi

rst is represented by the magnesite masses and associated with deep-seated metasomatism

and alteration during serpentinization, whereas the second, vein-forming stage postdates serpentinization and occurred

during obduction of the ophiolite. The differences in chemical composition between massive serpentinite and serpentinite-

hostedmagnesite masses suggest leachingofsomeelements and enrichmentofothers duringcarbonation; MgO,Cr, andNi

are depleted, whereas Fe

, CaO, MnO, Nb, Ba, Cu, Pb, Sr, and Zn are enriched in the serpentinite-hosted magnesite

masses, relative to the host massive serpentinite.

Online enhancements:

supplemental tables.

Introduction

Ophiolite sequences provide essential information

about the origin and development of oceanic lith-

osphere in general as well as the speci

fi

chistoryof

formation and closure of oceanic basins in a vari-

ety of tectonic environments (Dilek et al. 2000;

Moores et al. 2000; Pearce 2003; Beccaluva et al. 2004;

Shervais et al. 2004; Azer and Stern 2007; Furnes

et al. 2014). The ophiolites of the Arabian-Nubian

Shield (ANS), in particular, are important for un-

derstanding the formation and assembly of large

tracts of juvenile crust and the events of the pan-

African orogeny. In Egypt, exposures of the north-

western corner of the ANS include Neoproterozoic

ophiolites that represent fragments of oceanic litho-

sphere obducted onto continental crust during the

collision between West and East Gondwana and the

closure of the Mozambique Ocean (e.g., Pallister

et al. 1988; Stern 1994; Zimmer et al. 1995; Stern

et al. 2004; Azer and Stern 2007). Although generally

dismembered and affected by various degrees of al-

teration and metamorphism (e.g., Stern et al. 2004;

Azer and Stern 2007; Khalil et al. 2014; Boskabadi

Manuscript received April 14, 2018; accepted September 13,

2018; electronically published December 7, 2018.

* Author for correspondence; email: mokhles72@yahoo.com,

mk.abdel-malak@nrc.sci.eg.

[The Journal of Geology, 2019 volume 127, p. 81

–

107]

2018 by The University of Chicago.

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et al. 2017), these ophiolites are among the most

distinctive rock units in the basement exposures of

the Eastern Desert of Egypt, especially in the central

and southern sectors (

fi

g. 1).

A conspicuous feature of the ANS ophiolites is

the development of carbonate minerals in their

ultrama

fi

c rocks (e.g., Azer 2013; Boskabadi et al.

2017). Carbonate-altered rocks are spatially asso-

Figure 1.

Regional geological map showing the distribution of late Neoproterozoic ophiolitic rocks in the central and

south Eastern Desert of Egypt (modi

fi

ed after Shackleton 1994; tectonic boundary from Stern and Hedge 1985). The

location of

fi

gure 2 is indicated. Inset map shows the location of the Eastern Desert of Egypt within the Arabian-

Nubian Shield (Stern et al. 2004). A color version of this

fi

gure is available online.

M . K . A Z E R E T A L .

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ciated with magnesite, ta

lc, and gold deposits that

commonly occur along shear zones and fault planes

(e.g., Salem et al. 1997; Ghoneim et al. 1999, 2003;

Botros 2002; Ali-Bik el al. 2012; Azer 2013; Boskabadi

et al. 2017). Thelargevolume of carbonated ultrama

fi

rock implies signi

fi

cant

fl

uxes of CO

-rich

fl

uids,

presumably during amalgamation of the ANS ju-

venilecrust in the lateNeoproterozoic era. Crucially,

thesourceoftheCO

causing the carbonation and the

timing of carbonate alteration are unknown. Further-

more, geochemical redistribution of major and trace

elements between the host rocks and carbonates such

as magnesite has not been well documented. In gen-

eral, then, the processes associated with ultrama

fi

mineral carbonation and the effects of these pro-

cesses

—

in the ANS and elsewhere

—

remain poorly

studied and understood.

This contribution aims to integrate

fi

eld observa-

tions,petrography,whole-rockchemistry,andmineral

composition of serpentinized and carbonated ultra-

fi

c exposures of the dismembered ophiolite in the

Ghadir-Mohagar-Ambaut (GMA) area of the Eastern

Desert of Egypt. The goals of the study include de-

fi

ning the petrogenesis andgeodynamic setting of the

ophiolite and its incorporation into the juvenile crust

of the ANS as well as constraining the nature and

timing of carbonation and alteration of ophiolitic

ultrama

fi

cs in the Eastern Desert of Egypt and the

formation of associated ore deposits.

Field Geology

The Neoproterozoic basement complex of the East-

ern Desert and Sinai constitutes the northwestern

exposure of the ANS, which formed by collision be-

tween East and West Gondwana during the pan-

African orogeny (Stern 1994). A major component of

the ANS, especially in the Eastern Desert of Egypt,

consists of ophiolite sequences representing frag-

ments of oceanic lithosphere obducted onto conti-

nental crust during the amalgamation of the shield.

These ophiolite sequences are typically dismembered

and distributed as blocks in a tectonic mélange (Shack-

letonetal.1980;Riesetal.1983;Church1988;ElGaby

et al. 1988). Sequences may include a lower

“

mantle

”

unitofserpentinizedperidotiteandanupper

“

crustal

”

unit of gabbro (layered and isotropic), sheeted dikes,

andpillowbasalts(Sternetal. 2004;Khalil etal. 2014;

Gahlan et al. 2015; Obeid et al. 2016), though the full

classical sequence is rarely found intact within a

single mélange block.

The study area lies near the southern end of the

area typically designated as the central Eastern Des-

ert of Egypt, about 30 km southwest of Mersa Alam

(

fi

g. 1). It is dominated by outcrops of Neoproterozoic

rocks including dismembered ophiolite, island arc as-

semblages, and intrusive rocks (

fi

g. 2). Where contacts

between the ophiolitic units (especially serpentinites

and related quartz-carbonate rocks) and the island arc

assemblages are exposed, they are thrust contacts with

the island arc assemblages in the footwall. The island

arc footwall rocks form a volcanosedimentary suc-

cession of slates, greywackes, and

fi

ne- to coarse-

grained tuffs, with some lithic fragments of andes-

ite. Both the ophiolite and the island arc sequence

are intruded by younger granodiorites and granites,

with minor tonalite.

The ophiolitic rocks in the study area are one part

of the Ghadir ophiolite sequence, an outstanding

example of the ophiolites of the northernmost ANS

(El Sharkawy and ElBayoumi 1979; El Bayoumi 1983;

Kröner et al. 1992). They are found as allochthonous

blocks within a mélange (e.g., El Sharkawy and El

Bayoumi 1979; Takla et al. 1982; Basta et al. 1986;

Abd El-Rahman et al. 2009) once mapped as geosyn-

clinal metasediments (El Ramly 1972). All the classic

elements of the ophiolite sequence can be found

within the GMA area, including serpentinized peri-

dotites, metagabbros, sheeted dikes, and pillow lavas

(Basta et al. 1986; Khalil and Azer 2007). The ophio-

liticmetagabbrosoccupy the easternpartofthe study

area; they form small hillocks of variable height and

are melanocratic, medium grained, and black. Some

gabbros are sheared and intruded by granodiorite.

The serpentinite outcrop stretching from Gebel

Ghadir to Gebel Ambaut represents one of the largest

and most continuous outcrops of serpentinite in the

Eastern Desert of Egypt. The serpentinite is found

both in a large mass in the middle of the mapped area

and as small olistolith blocks within the ophiolitic

mélange. In the

fi

eld it displays a range of colors,

usually black or greenish black. All the contacts

around the serpentinite masses are tectonic and

marked by strong brecciation and shearing. Serpen-

tinitesarefrequentlytransformedintotalc-carbonate

and quartz-carbonate rocks (listwaenite), particularly

along thrust faults and shear zones. A few chromitite

lenses are encountered within the serpentinite, and

narrow pyroxenite dikes cut across the serpentinites.

Magnesite bodies occur near the contacts between

ultrama

fi

c rocks and siliceous country rocks and

along regional faults cutting the ultrama

fi

c rocks.

Magnesite ores in the GMA area form veins and

stockworks (

fi

g. 3

) as well as massive snow-white

deposits (

fi

g. 3

) within the ultrama

fi

ccountry

rocks. Most magnesite veins are branched (

fi

g. 3

and their contacts with the country rocks are sharp

butirregular.Themagnesiteveinsvaryinwidthfrom

a few centimeters to 1 m. Some fault zones feature

magnesite veins passing gradually upward into mas-

Journal of Geology

MAGNESITE IN NEOPROTEROZOIC OPHIOLITES OF EGYPT

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sive magnesite in the hanging wall. Small magnesite-

fi

lled veinlets and stockworks, a few millimeters to

2 cm wide, in the altered ultrama

fi

c rocks show no

well-de

fi

ned preferred orientation. Some brecciated

serpentinites are present within the magnesite veins

(

fi

g. 3

). Most of the veins occurring outside the main

shear zones have an orientation and coarse texture

suggestive of tension gashes opened with the same

sense of shear as the shear zones.

In places

—

for example, along shear zones

—

the

serpentinites are altered into talc-rich rocks. These

are generally rather uniform in mineralogical com-

position but vary appreciably in fabric and color; the

rocks may be massive, schistose, or gneissic and range

from darkgreenish gray to light gray, yellowish white,

or cream colored. The talc-rich rocks are commonly

soft, but they become progressively harder with in-

creasing iron contents or abundance of carbonate

minerals. They show cavernous weathering and host

small veinlets of crystalline magnesite. A few nodules

and irregular pockets of magnesite and dolomite with

quartz are observed in the talc-rich rocks. The quartz

and carbonate veins generally have a northwest-

southeast trend concordant with the foliation in the

talc-rich bodies. A few secondary veinlets of sutured

quartz grains are observed, interpreted to have de-

veloped along preexisting mylonite zones.

Analytical Techniques

Mineral identi

fi

cation was performed on thin and

polished sections with a polarizing Nikon micro-

Figure 2.

Geological map showing the distribution of late Neoproterozoic rocks in the Ghadir-Mohagar-Ambaut area

(modi

fi

ed after Takla et al. 1982). G.

p

Gebel. A color version of this

fi

gure is available online.

M . K . A Z E R E T A L .

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scope and con

fi

rmed by X-ray diffraction (XRD)

and an electron probe microanalyzer. Powder XRD

analyses were obtained with a Bruker D8 advanced

X-ray diffractometer at the Central Metallurgical

and Development Institute in Cairo, Egypt. Cu ra-

diation and a secondary monochrometer were used

at constant voltage and current of 40 kV and 40 mA,

respectively, with a scanning speed in 2

of 1

/min.

Serpentinite samples with minimal petrographic

evidence of carbonate alteration were selected care-

fully for whole-rock geochemical analysis. A total

of 35 samples (19 serpentinites, 3 chromitites, and

13 magnesite ores) were crushed and powdered in

an agate ball mill. Concentrations of major and

trace elements were determined by X-ray

fl

uores-

cence (XRF; ThermoARL XRF spectrometer) at the

GeoAnalytical Lab, Washington State University.

Each powdered sample was weighed, mixed with

two parts dilithium tetraborate

fl

ux, fused at 1000

in a muf

fl

e furnace, and cooled; the resulting bead

was reground, re-fused, and polished on diamond laps

to provide a smooth,

fl

at analysis surface. Reference

material 650CC from USGS standard rock powder

GSP2 was used as a calibration standard. Detection

limitsfor major oxides andtrace elementsrange from

0.001 to 0.04 wt% and from 0.01 to 0.5 ppm, re-

spectively (full documentation is available online

from the GeoAnalytical Lab; https://environment

.wsu.edu/facilities/geoan

alytical-lab/technical

-notes/). The internal precision of the XRF analyses,

calculated from duplicate samples, is better than

1% for most major elements and better than 5% for

most trace elements (except Ni, Cr, Sc, and V,

which may scatter by up to 10%). Loss on ignition

Figure 3.

Field photographs from the Ghadir-Mohagar-Ambaut area.

, Branched veins and stockwork of magnesite.

, Snow-white deposits of massive magnesite.

, Brecciated serpentinite within a magnesite mass. A color version

of this

fi

gure is available online.

Journal of Geology

MAGNESITE IN NEOPROTEROZOIC OPHIOLITES OF EGYPT

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(LOI) was determined by weight difference after ig-

nition at 1000

C.Three chromitite sampleswere also

analyzed by XRF (Phillips PW 1400 spectrometer) at

the Dipartimento di Scienza della Terra, Università

degli Studi, in Rome.

The chemical compositions of the essential rock-

forming minerals in the serpentinites, chromitites,

and magnesite ores were obtained with a

fi

ve-

spectrometer JEOL JXA-8200 electron microprobe

housed at the Geological and Planetary Sciences

Division Analytical Facility, California Institute of

Technology. Operating conditions were 15-kV ac-

celerating voltage, a 25-nA beam current, a focused

beam (1

m), 20-s on-peak counting times, and 10 s

each on high and low background positions. Natu-

ral and synthetic mineral standards were used for

calibration, and data were reduced with the CITZAF

matrix correction algorithm.

Petrography

Ultrama

fi

c Rocks.

We divide the ultrama

fi

c rocks

in the study area into two main groups: massive

serpentinites and serpentinite-hosted magnesite

rocks. Massive serpentinite consists essentially of

serpentine minerals (80%

–

97%), with minor mag-

nesite, chromite, magnetite, and chlorite. A few

specks of sul

fi

des are observed in some samples. Our

XRD results indicate that the serpentine minerals

are mainly antigorite, with subordinate amounts of

lizardite and chrysotile.

The original textures of the ultrama

fi

c rocks have

been almost completely obliterated by serpentine

replacement; however, the original forms of ortho-

pyroxene and olivine can be recognized by the pres-

ence of bastite and of mesh textures (

fi

g. 4

respectively. Rare relics of olivine and pyroxene are

recorded in a few samples (

fi

g. 4

). Serpentine

minerals vary in color from colorless to pale green

and exhibit pseudomorphic, interpenetrating, and

hourglass textures. Antigorite occurs as

fi

brolamellar

fl

ame-like crystals and interpenetrating bladed

crystals. Chrysotile

fi

lls veins that crosscut the an-

tigorite matrix (

fi

g. 4

), indicating protracted serpen-

tinization. Locally, in the ma

ssive serpentinite, quartz

occurs as brown microcrystalline in

fi

llings (

fi

g. 4

The brown color is due to the presence of micron-sized

inclusions of goethite and hematite.

Chromian spinels in the massive serpentinite form

tabular and elongated crystals with rounded rims.

Primary dark-brown chromian spinel is partially al-

tered along rims and cracks to ferritchromite and

magnetite. In a few thin sections, the altered chromian

spinel rims are surrounded by chlorite aureoles,

including

fl

aky, violet Cr-chlorite (kämmererite)

surrounded in turn by a thin

fi

lm of light-green

chlorite. The petrographic relationship between käm-

mererite and chromian spi

nel suggests that the Cr-

chlorite formed as primary chromian spinel was hy-

drothermally altered to ferritchromite. Secondary

magnetite is common as a result of release of Fe

upon serpentinization of olivine and pyroxene. Mag-

netite commonly forms rims around primary Cr-

spinels and ferritchromite. Fine-grained disseminated

magnetite, locally very abundant, occurs in every

serpentinite section.

The marginal parts of the massive serpentinite

are sheared and display a schistosity formed by

subparallel alignment of serpentine

fl

akes. Sheared

serpentinite samples are brecciated and have mylo-

nitic and cataclastic textures. Fractures within the

sheared serpentinites are

fi

lled with chrysotile, quartz,

and carbonates.

Serpentinite-hosted magnesite rocks are made up

of serpentine minerals (

∼

–

80 vol%) and carbon-

ates (

∼

–

15 vol%), with minor chromite, magne-

tite, and chlorite. A few samples contain rare

fi

brous crystals of tremolite-actinolite. No relics of

primary olivine or pyroxenes are found in this rock

type. The carbonate minerals include sparse cryp-

tocrystalline patches,

fi

ne aggregates of magnesite

(

fi

g. 4

), and rare crystals of dolomite. Texturally,

most carbonates are intergrown with and appear to

be coeval with serpentine minerals. There are also

veinlets of magnesite crosscutting the antigorite

matrix (

fi

g. 4

Chromitite.

Samples of the chromitite pods in-

clude both massive and disseminated chromitites,

which can be clearly distinguished in hand specimen

and thin section. Massive chromitite is composed

mainly(

1

92 vol%) ofcoarse-grained chromian spinel,

with minor interstitial serpentine minerals and rare

olivine. Chromian spinel occurs as equant euhedral

to anhedral crystals. They may be unaltered or show

only slight replacement by ferritchromite along grain

margins and cracks. Very small olivine inclusions are

observed within the chromian spinels, and traces of

intergranular or fracture-

fi

lling Cr-chlorite (kämme-

rerite) can be found. Cumulate, chain structures and

banding textures, typical of magmatic crystallization

(Pal and Mitra 2004), are common.

Disseminated chromitite consists of 40

–

70 vol%

chromian spinel and intergranular minerals, mostly

serpentine minerals, with rare olivine, chlorite, and

opaques. Most chromian spinel grains in dissemi-

nated chromitite are more euhedral and smaller in

size than those in the massive variety. They show

highly porous rims of ferritchromite. The dissemi-

nated chromitites are much more altered than

the massive chromitites, presumably because the

M . K . A Z E R E T A L .

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Figure 4.

Photomicrographs under cross-polarized light.

, Bastite texture after pyroxene.

, Mesh texture after olivine.

, Fresh relics of olivine.

, Fresh relics

of pyroxene.

, Chrysotile-

fi

lled vein crosscutting antigorite matrix.

, Brown, microcrystalline, quartz-

fi

lled vein crosscutting antigorite matrix.

, Crypto-

crystalline sparse crystals of magnesite intergrown with serpentine minerals.

, Magnesite veinlet crosscutting antigorite matrix.

, Cryptocrystalline magnesite

with coarse-grained, dolmoite-

fi

lled vug.

, Cracked chromian spinel crystals within massive magnesite.

, Serpentinite fragment within massive magnesite.

, Quartz veinlet within massive magnesite. Bst

p

bastite; Srp

p

serpentine; Ol

p

olivine; Pyx

p

pyroxene; Ctl

p

chrysotile; Mgs

p

magnesite; Qtz

p

quartz;

Dol

p

dolomite; Spl

p

spinel. A color version of this

fi

gure is available online.

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intergranular silicate phases in the disseminated

variety provide more opportunity for subsolidus

element redistribution.

Magnesite.

Magnesite is found in a range of set-

tings and textures in the GMA serpentinite. It occurs

as sparse crystals and

fi

ne aggregates in massive ser-

pentinite or as veins and irregular, cryptocrystalline

masses (with minor, coarse-grained dolomite and

calcite vug

fi

llings) in the sheared varieties (

fi

g. 4

The massive magnesite may contain chromian spinel

crystals similar to those in the host rock (

fi

g. 4

)as

well as angular to subangular lithic fragments of the

host serpentinite (

fi

g. 4

). Occasional quartz veinlets

arefoundinthemassivemagnesite(

fi

g. 4

Magnesite veins are mineralogically pure and pre-

dominantly cryptocrystalline in texture, with sharp

contacts with the enclosing serpentinites. Rare an-

gularfragmentsofhostserpentiniteareobservednear

the margins of some veins. Outside the main shear

zones there are also better-cr

ystallized, coarse-grained

magnesite veins.

Talc-Rich Rocks.

Talc-rich rocks are

fi

ne grained

and show brownish-yellow to reddish-brown color in

thin section. They are composed essentially of talc

(

1

75 vol%), with subordinate serpentine, carbonate,

and opaques. Talc forms

fi

ne aggregates or platy

grains. Carbonates form irregularly distributed coarse

aggregates. Opaque minerals include magnetite, chro-

mian spinel, and sul

fi

des.

Mineral Chemistry

Major-element data for silicate (serpentine, oliv-

ine, pyroxene, and chlorite), chromian spinel, and

carbonate minerals in serpentinite, chromitite, and

magnesite ores were determined by electron mi-

croprobe. The whole data set of mineral analyses is

given in supplementary tables S1

–

S11, available

online.

Silicate Minerals.

Serpentine Minerals.

Serpen-

tine minerals were analyzed both in the serpentinite

matrix (table S1) and in veins (table S2). The serpen-

tine minerals in both matrix and veins are chemi-

cally homogeneous and compositionally near ideal.

Those in the matrix have higher Mg# (molar

(Mg

1

1

); 0.96

–

0.98) than vein-forming ser-

pentine (0.90

–

0.96). Si content in the vein serpentine

is 1.99

–

2.07 atoms per formula unit (apfu), slightly

lower than that in the matrix serpentine (2.03

–

2.10 apfu). Matrix and vein serpentine both show

low concentrations of TiO

(

≤

0.04 wt%) and Al

(

≤

0.81 wt%).

Olivine.

Rare fresh relict olivines were analyzed

in serpentinites. In chromitite, both interstitial oliv-

ine and discrete inclusions within chromian spinels

were analyzed. Chemical compositions and struc-

tural formulaeofolivinearelistedintableS3. Olivine

in chromitite, whether included in spinel or inter-

stitial,ismoreforsteritic (Fo

½

forsterite

p

–

96)

than the fresh relics in serpentinite (0.90

–

0.93). The

concentrations of TiO

,Cr

, and CaO are low

(

≤

0.05 wt%) in all analyzed olivines. NiO is notably

higher in olivine from chromitite (0.68

–

0.8 wt%)

than in the fresh relics in serpentinite (0.35

–

0.45 wt%),

which plot along the NiO-Fo mantle array (

fi

g. 5

MnO concentration in olivine ranges between 0.07

and 0.22 wt% in serpentinite and is strictly lower

than 0.07 wt% in chromitite.

Pyroxenes.

Both clinopyroxene and orthopyrox-

ene were analyzed in serpentinite (table S4). In

terms of the pyroxene nomenclature of Morimoto

et al. (1988), the orthopyroxene is enstatite (Wo

0.9-2.1

89.5-91.3

7.4-8.4

,whereWo

p

wollastonite, En

p

enstatite,andFs

p

ferrosilite),withMg#

p

–

93,

while the clinopyroxene is mainly diopside (Wo

44.0-47.0

49.6-53.4

2.6-3.4

), with Mg#

p

–

96.

On the CaO-versus-Al

discrimination diagram,

the orthopyroxene analyses plot in the

fi

eld of pri-

mary mantle orthopyroxene from depleted harzburg-

ites (

fi

g. 5

). The low Al

content of orthopyrox-

ene and the high Mg# of clinopyroxene are both

consistent with the

fi

elds of forearc peridotites, con-

sidered to be residues after high degrees of partial melt

extraction (e.g., Hamlin and Bonatti 1980; Bonatti

et al. 1990; Ishii et al. 1992;

fi

g. 5

). At such high

degrees of melting, residual clinopyroxene would have

been exhausted; the observed clinopyroxene there-

fore presumably formed in the subsolidus as the CaO

content of orthopyroxene decreased upon cooling.

Another possibility is tha

t clinopyroxene is of meta-

somatic origin, which is consistent with correlations

described below between normative clinopyroxene

abundance and indicators of alteration.

Chlorite.

Chemical compositions and structural

formulae of disseminated chlorite and chloritic au-

reoles around chromian spinel in serpentinite are given

in table S5. Chlorite in the aureoles is distinguished

into green and violet Cr-bear

ing (kämmererite) varie-

ties. Green chlorite in the aureoles is rich in SiO

(28.6

–

30.1 wt%) and FeO

(18.8

–

19.6 wt%) and de-

pleted in MgO (18.1

–

19.0 wt%) and Al

(17.6

–

18.0 wt%), compared to both disseminated chlorite

and kämmererite. Kämmererite is richer in Cr

(2.6

–

3.7 wt%) than either disseminated chlorite (1.0

–

1.9 wt%) or the green chlorite aureole rims (1.0

–

1.1 wt%). According to the classi

fi

cation scheme of

Hey (1954), disseminated chlorite is mostly ripidolite,

with minor pycnochlorite, whereas in the chloritic

aureoles around chromian spinel, the kämmererite is

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classi

fi

ed as ripidolite and the green chlorite rims are

pycnochlorite.

The boundaries of chlorite stability in the presence

of an aluminous phase can be used to determine the

temperature of chlorite formation (e.g., Cathelineau

and Nieva 1985; Kranidiotis and MacLean 1987;

Hillier and Velde 1991; Bourdelle and Cathelineau

2015; Yavuz et al. 2015). Calculated temperatures

for chlorite formation according to the calibrated geo-

thermometer equation of Kranidiotis and MacLean

(1987), which uses tetrahedra

l Al content adjusted for

the effect of Mg#, are listed in table S5. Given the high

Cr content of associated spinel, the Al activity of the

rock may be lower than that assumed in the cali-

bration, leading to a systematic overestimate of the

temperature. This will not affect the relative se-

quence of temperatures from the different textural

settings and varieties of chlorite in a given rock,

however. The inferred temperatures for formation of

disseminated chlorite (293

–

305

C) and kämmererite

(279

–

301

C) are higher than those for the green

chlorite rims (242

–

261

C), suggesting that the green

chlorite aureole rims may sample a unique, late, hy-

drothermal stage.

Nonsilicate Minerals.

Spinel Group.

The chem-

ical compositions and structural formulae of primary

chromian spinel and its alteration products are given

in table S6 for serpentinite, table S7 for magnesite

masses, and table S8 for chromitite. Primary chromian

spinel shows signi

fi

cant heterogeneity in each sam-

ple. Disseminated chromian spinel in serpentinite

is commonly altered to ferritchromite and Cr-

magnetite; less commonly, Cr-magnetite is found

around chromian spinel in magnesite masses. Al-

teration to ferritchromite and Cr-magnetite concen-

trates FeO

but removes Al

,MgO,andCr

.MnO

concentration in ferritchromite is higher in serpen-

tinite (1.1

–

1.9wt%)thaninchromititelenses(0.8

–

Figure 5.

Chemistry of primary minerals in serpentinite and chromitite.

, Forsterite (Fo) versus NiO content of

olivine; mantle olivine array is from Takahashi et al. (1987).

, CaO versus Al

of orthopyroxene (Opx), compared to

the

fi

elds of Avacha highly depleted harzburgite (Hz) mantle xenoliths (Ishimaru et al. 2007), orthopyroxenes from

alpine-type peridotites (Arai 1980), contact-metamorphosed ultrama

fi

c rocks (Arai 1975; Frost 1975), and regionally

metamorphosed ultrama

fi

c rocks (Evans and Trommsdorff 1974; Trommsdorff et al. 1998).

,Al

versus Cr

orthopyroxenes, with the abyssal peridotite

fi

eld from Johnson et al. (1990) and the forearc peridotite

fi

elds after Ishii

et al. (1992).

,Mg#versusCr

of clinopyroxene (Cpx), compared to the abyssal-peridotite

fi

eld (Hamlin and Bonatti

1980; Johnson et al. 1990; Juteau et al. 1990) and the forearc-peridotite

fi

eld from Ishii et al. (1992). A color version of

this

fi

gure is available online.

Journal of Geology

MAGNESITE IN NEOPROTEROZOIC OPHIOLITES OF EGYPT

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1.0 wt%) or magnesite masses (0.7

–

1.0 wt%) as well

as being higher than in other spinel varieties (Cr-

magnetite has 0.2

–

0.6 wt% MnO in serpentinite and

0.2 wt% in magnesite masses; primary chromian spi-

nel has 0.4

–

0.7 wt% MnO in serpentinite, 0.3

–

0.6 wt%

in magnesite masses, and 0.2

–

0.7 wt% in chromitite).

The variability in spinel compositions is readily

visualizedontheAl-Cr-Fe

1

triangular plot (

fi

g. 6;Fe

1

is estimated from Fe

and stoichiometry). Fresh chro-

mian spinel plots along the Cr-Al join, whereas the

altered phases (ferritchromite and Cr-magnetite) plot

along the Cr-Fe

1

join, re

fl

ecting the loss in Al

and

and the increase in Fe

due to alteration and

metamorphism. Fresh chromian spinel in chromitite

shows continuous zoning from Al- and Cr-rich cores

toward rims enriched in Fe and Cr, whereas chro-

mian spinels in serpentinite and massive magnesite

display an abrupt compositional change from chro-

mian spinel core to ferritchromite and Cr-magnetite

rims (

fi

g.6).Inboth cases, Mgcontinuouslydecreases

from cores toward rims.

Chromian spinel in serpentinite and massive

magnesite is characterized by a moderate Cr#

(molar

(Cr

1

Al)

p

–

70 and 0.63

–

0.70, re-

spectively), whereas chromitite features chromian

spinel with a much higher Cr# (0.77

–

0.89). As indi-

cated in

fi

gure 6, primary chromian spinel has a low

1

p

1

(Fe

1

1

1

Al)

08,whichiscon-

sidered to be a diagnostic feature of primary mantle-

derived spinels (Hattori and Guillot 2007; Bernstein

et al. 2013). In terms of divalent cations, primary

chromian spinel in chromitite has compositions

that are notably more magnesian (Mg#

p

–

74)

than those of disseminated chromian spinels in

serpentinite (Mg#

p

–

62) or massive magne-

site (Mg#

p

–

53). The composition of primary

chromite in the chromitite from the GMA area is

similar to published analyses of chromitite lenses

from other localities throughout the Eastern Desert

of Egypt (Ahmed et al. 2001; Azer and Stern 2007).

Carbonates.

Chemical compositions and struc-

tural formulae of carbonate minerals are given for

serpentinite (table S9), massive magnesite ore (ta-

ble S10), and magnesite veins (table S11). Quality

control of the analyses was veri

fi

ed by ensuring

that CO

was suf

fi

cient to charge balance the ana-

lyzed FeO, MnO, MgO, and CaO to bring the ana-

lytical total to 100

2 wt%; analyses failing this

test were discarded. On the basis of chemical analysis,

the carbonate minerals include magnesite, ferroan

magnesite, and dolomite. Magnesite is the only

carbonate mineral observed in the magnesite veins,

whereas both magnesite and dolomite are recorded

in the magnesite masses. Magnesite from masses

contains lower MgO (46.3

–

48.0 wt%) than that in

the veins (48.3

–

49.2 wt%). Disseminated carbonate

minerals in serpentinite include magnesite, ferroan

magnesite, and dolomite. Disseminated magnesite

in serpentinite has MgO (43.4

–

46.2 wt%) lower

than that in either type of magnesite ore. Ferroan

magnesite, by de

fi

nition, is lower in MgO (39.5

–

41.0 wt%) and richer in FeO

(1.5

–

3.2 wt%) than

magnesite. Dolomite, whether in massive magne-

site or serpentinite, is dominated by MgO and CaO,

with low concentrations of other oxides.

Whole-Rock Geochemistry

Ultrama

fi

c Rocks.

Whole-rock major- and trace-

element data are given in table 1 for nine massive

serpentinite samples and 10 serpentinite-hosted

magnesite samples. Many of the incompatible trace

elements are below the analytical detection limits.

O, K

O, and P

are virtually absent in most

analyses. All the serpentinite samples have high

LOI values due to extensive hydration and carbon-

ation; it is notable that the massive serpentin-

ite samples have lower LOI (12.8

–

14.8 wt%) than

the serpentinite-hosted magnesite samples (15.1

–

16.4 wt%).

The normative mineralogy of each sample was

calculated from the whole-rock major-oxide analyses

(table 2). The norms of the serpentinized peridotites

are dominated by olivine and orthopyroxene, con-

sistent with harzburgite protoliths (

fi

g. 7), whereas

serpentinite-hosted magnesite samples apparently

gained CaO during alteration, re

fl

ected in normative

clinopyroxene and a shift into the lherzolite

fi

eld.

Figure 6.

Ternary Cr-Al-Fe

1

plot with chromian spinels

and their alteration products. A color version of this

fi

ure is available online.

M . K . A Z E R E T A L .

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Table 1.

Major- and Trace-Element Contents in the Massive Serpentinite and Serpentinite-Hosted Magnesite of the Ghadir-Mohagar-Ambaut Area

Rock type

Massive serpentinite

Serpentinite-hosted magnesite

AmS-2 AmS-11 AmS-14 AmS-19 AmS-24 AmS-26 AmS-29 AmS-37 AmS-47 AmS-13 AmS-17 AmS-21 AmS-31 AmS-34 AmS-42 AmS-45 AmS-49 AmS-51 AmS-55

Major oxides (wt%):

SiO

40.13 38.88 38.72 40.18 38.88 40.11 39.33 39.08 39.06 39.26 40.08 39.88 36.38 37.74 38.25 36.72 36.97 37.08 37.69

TiO

.01

.03

.01

.02

.01

.03

.02

.04

.02

.01

.02

.01

.06

.05

.36

.59

.38

.45

.63

.62

.44

.54

.19

.36

.34

.37

.43

.41

.71

.63

1.26

3.23

1.97

FeO

7.04

7.01

6.54

6.88

6.63

6.81

6.89

7.23

6.78

7.56

7.35

7.48

7.67

7.70

8.33

8.48

7.64

8.31

7.78

MnO

.06

.10

.08

.06

.10

.05

.06

.07

.10

.09

.16

.27

.17

.21

.13

.17

.11

.15

.17

MgO

38.25 38.33 38.18 38.54 39.52 38.16 38.61 37.27 39.08 35.76 35.28 34.84 35.09 35.74 34.92 35.81 34.78 33.04 34.62

CaO

.72

.91

.56

.76

.87

.33

.48

1.01

.55

1.34

1.56

1.52

2.41

1.12

1.23

2.10

1.94

1.67

2.65

!

.01

!

.01

!

.01

!

.01

!

.01

!

.01

.02

.10

.03

.06

.00

.33

.04

.01

.12

.05

.13

!

.01

!

.01

!

.01

!

.01

!

.01

!

.01

!

.01

!

.01

!

.01

!

.01

!

.01

.02

!

.01

!

.01

!

.01

!

.01

.02

.01

!

.01

!

.01

.02

!

.01

!

.01

LOI

12.80 14.19 14.78 13.39 13.27 13.38 14.16 14.22 13.78 15.31 15.14 15.32 16.16 16.30 15.54 15.82 16.36 15.72 15.09

Total

99.34 100.02 99.26 100.29 99.69 99.46 99.97 99.45 99.58 99.93 99.93 99.78 98.42 99.58 99.19 99.77 99.04 99.29 99.59

Mg#

91.50 91.55 92.04 91.73 92.19 91.74 91.74 91.08 91.95 90.36 90.49 90.22 90.06 90.19 89.25 89.32 90.02 88.74 89.81

Trace elements (ppm):

2182

2047

2124

2352

2310

2418

2108

2091

2284

1943

1727

2002

2057

2053

1064

1956

1718

948

1845

2526

2748

2845

2332

2191

2467

2516

2216

2960

2127

1889

2201

2036

2156

1785

2065

2207

1924

2283

8.4

6.2

7.8

8.4

8.1

7.5

7.9

6.8

4.6

7.6

2.4

6.3

8.6

9.2

10.1

5.4

7.1

9.2

5.6

31.8

24.2

29.6

33.8

21.3

28.6

26.7

44.3

15.2

22.1

15.6

24.5

21.5

31.2

18.6

20.2

26.7

30.1

16.8

1.3

3.1

2.6

1.4

5.2

2.5

1.8

.05

1.1

1.4

1.2

3.7

4.2

2.8

14.1

3.3

5.5

1.8

6.8

18.1

11.7

29.5

24.6

25.8

13.5

19.2

21.3

31.6

40.1

27.8

3.1

1.5

2.3

2.7

3.6

2.5

4.1

2.8

3.6

2.7

3.3

2.2

1.4

3.6

2.8

4.3

3.6

2.1

1.5

1.3

1.2

1.3

1.2

2.1

1.4

1.2

1.1

.13

.08

.18

.07

.11

.15

.17

.12

.31

.22

1.46

.67

.94

.44

.53

1.11

.45

.66

1.7

2.2

3.3

3.9

4.1

2.7

3.4

2.1

2.6

3.4

2.4

4.1

1.8

2.5

1.4

1.7

3.2

2.2

8.5

6.3

14.2

19.3

18.5

20.1

12.4

6.5

7.8

35.6

22.4

51.5

53.5

13.8

21.3

18.5

31.3

14.9

11.7

38.4

44.59 31.8

47.5

51.8

45.2

46.3

49.8

35.4

55.8

51.3

64.4

76.5

49.2

65.3

74.8

62.3

76.4

59.5

1.1

1.2

2.1

1.8

1.6

2.6

1.9

1.4

1.2

1.7

1.4

Note. LOI

p

loss on ignition.

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Table 2.

Normative Compositions of the Massive Serpentinite and Serpentinite-Hosted Magnesite in the Ghadir-Mohagar-Ambaut Area

Rock type

Massive serpentinite

Serpentinite-hosted magnesite

AmS-

Corundum

... ...

...

.02 ... ...

...

... ... ... ...

...

.12 ...

Orthoclase

. . .

.07 .07 . . .

. . .

.07 . . .

.07 .07 . . .

. . .

.07 .07 .14

Albite

. . .

.2 1

. . .

2.46 .4

.1 1.22 .51 1.29

Anorthite

1.13 1.87 1.17 1.38 1.95 1.9 1.4 1.67 .5

.63 .9

.87 1.39 . . .

2.1 1.99 3.46 9.9 5.55

Acmite

. . .

.78 . . .

. . .

Diopside

2.35 2.65 1.6 2.26 2.32 . . .

1.08 3.18 2.04 5.62 6.45 6.26 10.27 5.24 4.02 8.07 6.4

. . .

7.71

Diopside (Wo) 1.25 1.41 .85 1.2 1.23 . . .

.57 1.69 1.08 2.98 3.42 3.32 5.44 2.78 2.13 4.27 3.39 . . .

4.09

Diopside (En) .99 1.12 .68 .96 .99 . . .

.46 1.34 .87 2.35 2.7 2.61 4.28 2.18 1.66 3.33 2.67 . . .

3.21

Diopside (Fs)

.11 .12 .07 .1

. . .

.05 .15 .09 .29 .33 .33 .55 .28 .23 .46 .34 . . .

.42

Hypersthene 31.74 25.32 28.54 30.73 22.2 34.59 29.37 28.6 26.54 29.79 34.49 34.67 15.32 20.2 29.14 14.96 18.17 23.94 15.11

Hypersthene

(En)

28.69 22.89 25.97 27.86 20.23 31.37 26.63 25.72 24.12 26.55 30.76 30.77 13.59 17.89 25.59 13.15 16.13 20.9 13.36

Hypersthene

(Fs)

3.04 2.43 2.57 2.87 1.97 3.22 2.75 2.88 2.42 3.24 3.73 3.91 1.74 2.3 3.55 1.82 2.04 3.04 1.74

Olivine

62.37 67.72 66.2 63.17 71.11 61.12 65.77 63.9 68.27 60.31 55.18 54.91 70.2 68.93 61.44 71.93 67.89 62.55 67.43

Olivine (Fo) 55.83 60.62 59.69 56.72 64.21 54.9 59.05 56.87 61.46 53.14 48.67 48.16 61.51 60.35 53.29 62.41 59.58 53.89 58.94

Olivine (Fa)

6.54 7.1 6.51 6.45 6.89 6.22 6.72 7.04 6.81 7.17 6.51 6.75 8.68 8.58 8.15 9.53 8.31 8.66 8.49

Magnetite

1.98 2

1.9 1.93 1.88 1.92 1.95 2.07 1.94 2.19 2.14 2.22 2.3 1.9 2.44 2.48 2.27 2.45 2.27

Ilmenite

.03 .03 .06 .03 .04 .04 .02 .02 .07 .04 .08 .05 .02 .03 .04 .02 .14 .11 .11

Apatite

. . .

.03 .03 .03 . . .

. . .

.05 .03 .03 . . .

.03 .03 . . .

.03 .05 . . .

. . .

.03

Color index 98.46 97.72 98.29 98.12 97.54 97.67 98.19 97.77 98.86 97.95 98.35 98.11 98.12 96.29 97.09 97.47 94.87 89.05 92.62

Diff. index

. . .

.07 .07 . . .

. . .

.2 1

.37 .6

.07 2.53 .4

.1 1.29 .58 1.43

Note. Wo

p

wollastonite; En

p

enstatite; Fs

p

ferrosilite; Fo

p

forsterite; Fa

p

fayalite; Diff. index

p

differentiation index.

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Although the compositions of the serpentinite-

hosted magnesite samples are broadly similar to

those of the massive serpentinite samples, we can

recognize signi

fi

cant and consistent depletion in

MgO, Cr, and Ni, accompanied by enrichment in

, CaO, MnO, Nb, Ba, Cu, Pb, Sr, and Zn. The

variations in chemical compositions in serpentinites

and serpentinite-hosted magnesite are summarized

fi

gure 8. Assuming that the serpentinite-hosted

magnesite samples represent the products of alter-

ation of protoliths similar to the associated massive

serpentinite, we infer that these differences indicate

which elements are leached from serpentinite during

alteration and which are brought in by altering

fl

uids.

In the massive serpentinite samples, CaO is restricted

to low concentrations that mostly covary with Al

contents in the sense expected from melt depletion.

This indicates that Ca metasomatism did not signi

fi

cantly affect the massive serp

entinite samples, despite

their proximity to Ca-enriched lithologies such as

serpentinite-hosted magnesite, massive and vein mag-

nesite, talc-carbonate rocks, and quartz-carbonate

rocks.

The whole-rock Mg# of massive serpentinite

samples ranges between 0.91 and 0.92, higher than

that of serpentinite-hosted magnesite samples

(0.89

–

0.91), suggesting that there is depletion of

Mg, addition of Fe, or both during carbonation. On

the other hand, the serpentinites from the GMA area

have, on the whole, a range of Mg# (0.89

–

0.92) con-

sistent with the Mg# of modern oceanic peridotites

(Mg#

89; Bonatti and Michael 1989) and similar

to that of other ophiolites in the Eastern Desert of

Egypt (e.g., Azer et al. 2013; Khalil et al. 2014; Gahlan

et al. 2015; Obeid et al. 2016). It is rare to be able to

compare serpentinized peridotites to fresh peridotites

to assess whether serpentinization itself, though ob-

viously accompanied by mineral-scale redistribution

of Mg and Fe, is accompanied by metasomatic alter-

ation of whole-rock Mg#. In this case, however, we

have measurements of the Mg# of fresh relics of both

pyroxene and olivine, the dominant minerals in the

harzburgite protoliths, enabling an estimate of the

prealteration Mg# of the whole rock as a point of

comparison.

Chromitite Lenses.

Only the massive chromitite

was selected for major- and trace-element analysis

(three samples; table 3). As expected, the chromitite

samples have high Cr

,Fe

,Al

,andMgO

contents.TheyarerichinNi,V, andZnbutlowinRb,

Ba, Sr, Zr, and Nb. The bulk chromitite analyses have

contents of Cr

50 wt%, FeO

20 wt%,

TiO

3 wt%, and MnO

1 wt%. These charac-

teristics are similar to those of alpine-type podiform

chromitites considered to be derived from depleted

upper-mantle sources (Dickey 1976; Jan and Windley

1990).

Magnesite Ores.

Whole-rock geochemical data

for 13 massive and vein magnesite samples are shown

in table 3. The massive and vein types are signi

fi

cantly different from each other. The massive mag-

nesite type is generally higher in SiO

,FeO

,CaO,Cr,

Ni, Sr, Ba, and Pb than the vein-type magnesite. Such

enrichment in major oxides and trace elements not

hosted by pure magnesite is consistent with and can

be attributed to the observed presence of minor ser-

pentine minerals, calcite, dolomite, and magnetite in

the magnesite masses as well as the more FeO- and

CaO-rich compositions of the magnesite mineral

analyses from the massive type samples. Speci

fi

cally,

high Cr concentrations (158

–

276 ppm) in the massive

magnesite re

fl

ect the presence of disseminated chro-

mian spinels. The relatively high Sr (124

–

203 ppm)

and Ba (19

–

25 ppm) contents of the massive magne-

site can be attributed to the presence of Ca-bearing

carbonates and serpentine minerals; substitution of

Ba and Sr into magnesite is quite limited because of

the large ionic-radius mismatch with Mg (Möller

1989; Zachmann and Johannes 1989).

Discussion

The ophiolitic rocks of Egypt have been a subject

of considerable research interest because they offer

important clues for reconstructing the geodynamic

Figure 7.

Ternary plot of olivine (Ol), orthopyroxene

(Opx), and clinopyroxene (Cpx) abundance in the nor-

mative mineralogy of massive serpentinite and

serpentinite-hosted magnesite samples (ultrama

fi

crock

nomenclature from Coleman 1977). A color version of

this

fi

gure is available online.

Journal of Geology

MAGNESITE IN NEOPROTEROZOIC OPHIOLITES OF EGYPT

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Figure 8.

Silica vari-

ation diagrams for

some major and trace

elements, emphasiz-

ing differences in

whole-rock chemical

compositions

be-

tween serpentinite

and

serpentinite-

hosted magnesite. A

color version of this

fi

gure is available

online.

Figure 8.

Silica variation diagrams for some major and trace elements, emphasizing differences in whole-rock chemical compositions between serpentinite

and serpentinite-hosted magnesite. A color version of this

fi

gure is available online.

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Table 3.

Major- and Trace-Element Contents in the Massive Chromitite and Magnesites of the Ghadir-Mohagar-Ambaut Area

Rock type

Massive chromitite

Magnesite

Magnesite veins

Massive magnesite

AMC-1 AMC-5 AMC-8 MV-1 MV-4 MV-9 MV-14 MV-17 MV-21 MV-25 MV-28 MM-4 MM-10 MM-13 MM-17 MM-22

Major oxides (wt%):

SiO

4.89

5.08

5.27

.66

.32

.44

.17

.36

.57

.97

4.64

6.14

3.43

5.96

3.12

TiO

.12

.09

.11

.01

.00

.01

.00

.01

.02

.00

.02

.09

15.79 14.05 15.41

.01

.02

.03

.04

.03

.01

.52

.44

.74

.53

.64

FeO

15.08 17.07 16.04

.16

.05

.08

.09

.04

.18

.16

.22

1.21

1.17

1.01

1.36

1.14

39.77 40.27 39.64 ...

...

MnO

.18

.21

.17

.02

.01

.03

.02

.01

.03

.01

.02

.10

.06

.04

.07

.05

MgO

20.87 19.86 20.52 40.68 42.33 41.08 41.44 42.96 43.55 43.37 42.98 36.89 37.96 40.23 39.46 39.79

CaO

.22

.28

.31 1.94

1.38 1.67

.47

.84

2.08

1.85 2.63

6.78

5.06

2.67

2.89

3.52

O.02

!

.01

!

.01

.02

.01

!

.01

!

.01

.03

.01

.02

!

.01

.03

.02

!

.01

!

.01

!

.01

!

.01

!

.01

.02

.01

.02

.01

!

.01

!

.01

.03

!

.01

!

.01

.02

.01

.02

.01

.03

.24

.14

.04

.02

.40

.05

.04

.05

.02

.14

.05

LOI

2.64

2.37

2.09 55.84 56.01 56.32 56.76 55.73 53.26 54.07 52.78 49.93 48.77 52.02 49.17 51.58

Total

99.59 99.30 99.58 99.37 100.38 99.69 99.33 99.84 99.54 100.47 99.68 100.15 99.70 100.18 99.69 100.02

Trace elements (ppm):

1495

1503

1987

314

441

332

427

205

. . .

158

178

182

198

276

1.2

2.1

1.6

1.5

1.2

1.6

1.1

1.4

1.3

774

834

817

15.6

14.5 17.6 19.2

18.4

20.1

16.8 18.9

20.4

19.7

32.3

20.4

18.9

1.2

1.1

1.8

2.4

1.5

1.4

3.4

1.9

1.5

2.3

3.6

3.4 78.9

39.4 28.8 32.2

42.3

72.5

79.6 82.4 123.7 203.1 131.4 139.4 124.6

1.9

2.3

1.4

2.8

2.6

4.4

1.5

2.6

2.2

1.4

3.4

2.8

3.3

1.9

2.1

6.7

1.8

2.4

3.3

1.4

2.7

3.5

4.6

2.4

1.2

1.3

1.2

2.1

1.6

2.1

13.7

16.8

14.2

1.3

1.7

1.1

1.2

1.4

1.1

1.7

3.6

7.9

2.8

9.4

4.3

1.89 3.45

5.8

3.3

6.5

2.2

3.1

185

246

234

4.6

4.9

6.5

11.2

5.3

6.8

4.9

6.7

4.5

8.8

3.6

7.4

8.7

9.4 10.2 13.6

9.8

11.2

13.5

8.8

23.1

22.3

21.8

22.6

24.52

Note. LOI

p

loss on ignition; ND

p

not determined.

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evolution of the pan-African belt (Kröner et al. 1992;

Shackleton 1994; Zimmer et al. 1995; Stern et al.

2004; Azer and Stern 2007; Ali et al. 2010; Gahlan

et al. 2015; Obeid et al. 2016). Most assessments of

the tectonic setting for Egyptian ophiolites have

focused on the bulk chemistry of lavas and gabbroic

rocks; the abundant serpentinized ultrama

fi

crocks

have rarely been considered for this purpose. How-

ever, de

fi

ning the tectonic setting of Neoproterozoic

ophiolitic rocks on the basis of the bulk chemistry of

the crustal section is challenging because of the ef-

fects of fractional crystallization and alteration (Azer

and Stern 2007). Likewise, the whole-rock chemistry

of highly serpentinized peridotite is of limited use

in tectonic discrimination. Rather, the chemistry

of preserved magmatic minerals

—

particularly oliv-

ine, spinel, and pyroxene

—

fl

ects the magmatic

conditions and tectonic environment of ultrama

fi

protoliths and is often robust against less-than-

complete alteration processes (e.g., Arai 1980, 1992;

Dick and Bullen 1984; Beccaluva et al. 1989; Jan and

Windley 1990; Barnes and Roeder 2001; Proenza

et al. 2004). Hence, we focus on the compositions of

relict magmatic minerals and seek to deduce the tec-

tonic environment and petrogenesis of the serpen-

tinites of the GMA area.

Tectonic Setting of Serpentinized Peridotites.

De-

spite universal agreement that ophiolite sequences

represent fragments of oceanic lithosphere, the dif-

ferences among the detailed tectonic environments

in which oceanic lithosphere can be formed are subtle

and can lead to different interpretations. Thus, al-

though Zimmer et al. (1995) inferred an open-ocean

mid-ocean ridge tectonic setting for the Gerf ophi-

olites in the south Eastern Desert of Egypt, the past

two decades have seen the development of a mount-

ing consensus that all the ophiolitic rocks of Eastern

DesertofEgyptformedinsubduction-relatedsettings

(e.g., El Sayed et al. 1999; Ahmed et al. 2001; Farahat

et al. 2004; Azer and Khalil 2005; Azer and Stern

2007). At a

fi

ner level of detail, while the majority of

recent studies of serpentinites throughout the East-

ern Desert have assigned their origin to forearc

settings (e.g., Azer and Stern 2007; Khalil and Azer

2007; Abd El-Rahman et al. 2009; Azer et al. 2013;

Hamdy et al. 2013; Khedr and Arai 2013; Khalil et al.

2014; Gahlan et al. 2015; Obeid et al. 2016), a few

ophiolitic outcrops continue to be assigned to back-

arc settings on the basis of the transitional geo-

chemical character of their volcanic sequences, be-

tween those of island arcs and mid-ocean ridges (e.g.,

El Sayed et al. 1999; Farahat et al. 2004; Abd El-

Rahman et al. 2009; Basta et al. 2011).

Our data on whole-rock and relict mineral chem-

istry of serpentinites from the GMA area allow us to

examine the question of the tectonic setting of the

Ghadir ophiolite complex by explicit comparison to

corresponding data from ultrama

fi

c rocks of known

geodynamic settings worldwide. Although whole-

rock chemistry is not the most reliable indicator, we

do observe that the GMA serpentinites, like a num-

berofotherperidotiteoutcropsintheANS(e.g.,Stern

et al. 2004; Azer and Stern 2007), plot within the

highly depleted

fi

eld associated with the extreme

melt depletions experienced by forearc peridotites.

Figure 9

, for example, shows that massive serpen-

tinite samples have CaO (

!

0.9 wt%) and Al

(

!

0.6 wt%) contents diagnostic of forearc peridotites

(

fi

g. 9

), and the same is true for other elements, such

as MgO, FeO

, Ni, and Cr.

Perhaps the best petrogenetic indicator for the

tectonic setting of ma

fi

c and ultrama

fi

c igneous

rocks is the chemical composition of primary spinel

(e.g., Dick and Bullen 1984; Jan and Windley 1990;

Arai 1992; Barnes and Roeder 2001; Kamenetsky

et al. 2001; Sobolev and Logvinova 2005; Arif and

Jan 2006). Spinels from mid-ocean ridge and backarc

basinperidotitesgenerally haveCr#

50 (Barnesand

Roeder 2001; Ohara et al. 2002), whereas spinels in

forearc peridotites generally have a higher Cr# (up to

80), and those from boninites typically have a Cr# of

–

90. Keeping in mind that alteration and meta-

morphism can modify spinel compositions and com-

plicate their petrogenetic interpretation (e.g., Zhou

et al. 1996; Barnes and Roeder 2001; Proenza et al.

2008; González-Jiménez et al. 2009), we focus here

only on the petrographically and chemically fresh

chromian spinels in the serpentinite, chromitite,

and magnesite ore from the GMA area. Analyses

from serpentinite and magnesite ores have Fe

1

/Fe

1

ratios and Al

contents consistent with the supra-

subduction zone peridotite

fi

eld (SSZ in

fi

g. 9

). On a

Cr#-Mg# plot (

fi

g. 9

), the chromian spinels from

serpentinite and magnesite ore samples plot above

Cr#

p

60, overlapping the

fi

elds of both modern

forearc peridotites that experienced high degrees of

melt extraction (

∼

34%

–

39%) and other Eastern Des-

ert ophiolites previously assigned to forearc settings

(e.g., Azer and Stern 2007; Azer 2014; Khalil et al.

2014; Gahlan et al. 2015; Obeid et al. 2016).

Fresh chromian spinels in the massive and dis-

seminated chromitite lenses are clearly distinct from

those in the serpentinite and magnesite ore samples.

TheirCr#,Mg#,andTiO

contents all resemble those

of spinels from boninite lavas (

fi

g. 9

). Boninites,

of course, are characteristic of certain stages in the

evolution of the forearc of intraoceanic arcs (e.g.,

Murton 1989; Johnson and Fryer 1990; Bédard 1999;

Beccaluva et al. 2004), and boninitic af

fi

nities have

been used to support a forearc setting in a number of

M . K . A Z E R E T A L .

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other Egyptian ophiolite sequences (e.g., El Sayed

et al. 1999; Abdel Aal et al. 2003; Saleh 2006).

Petrogenesis.

Serpentinized Peridotite.

Exposed

serpentinite sequences in ophiolites have witnessed

a series of processes during their journey from oce-

anic mantle to juvenile crust and

fi

nally to surface

exposure. Their chemistry, mineralogy, and texture

can therefore be used, in principle, to constrain pro-

cesses in all of these environments. However, each

subsequent stage of their com

plex history partly over-

prints evidence of earlier stages, and so their geo-

chemical signatures must be treated with caution to

Figure 9.

Discrimination diagrams for the tectonic setting of ultrama

fi

c rocks on the basis of whole-rock and spinel

chemistry.

, Whole-rock Al

-versus-CaO diagram (Ishii et al. 1992).

,SpinelAl

-versus-Fe

1

/Fe

1

diagram,

showing the

fi

elds of spinel from suprasubduction zone (SSZ) and mid-ocean ridge (MOR) peridotite (after

Kamenetsky et al. 2001).

, Mg#-versus-Cr# diagram for fresh chromian spinels (after Stern et al. 2004), with

fi

eld

boundaries from Dick and Bullen (1984), Bloomer et al. (1995), and Ohara et al. (2002). The

fi

eld for chromian spinels

in serpentinites of the Eastern Desert of Egypt is based on data from Azer and Khalil (2005), Khalil and Azer (2007),

Farahat (2008), Farahat et al. (2011), Azer (2014), Khalil et al. (2014), Gahlan et al. (2015), and Obeid et al. (2016). The

experimental batch melting trend, marked by degree of melting (offset for clarity), is from Hirose and Kawamoto

(1995).

,TiO

-versus-Cr# diagram for fresh chromian spinels (

fi

elds after Dick and Bullen 1984; Arai 1992; Jan and

Windley 1990). MORB

p

MOR basalt. A color version of this

fi

gure is available online.

Journal of Geology

MAGNESITE IN NEOPROTEROZOIC OPHIOLITES OF EGYPT

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