JOURNAL OF PETROLOGY
VOLUME 49
NUMBER 6
PAGES 1097^1131
2008
doi:10.1093/petrology/egn019
The Alkaline^Peralkaline Tamazeght Complex,
High Atlas Mountains, Morocco: Mineral
Chemistry and Petrological Constraints for
Derivation from a Compositionally
Heterogeneous Mantle Source
MICHAEL A. W. MARKS1*, JULIAN SCHILLING1,
IAN M. COULSON1,2, THOMAS WENZEL1 AND GREGOR MARKL1
INSTITUT FU«R GEOWISSENSCHAFTEN, AB MINERALOGIE UND GEODYNAMIK, EBERHARD-KARLS-UNIVERSITA«T,
1
WILHELMSTRASSE 56, D-72074 TU«BINGEN, GERMANY
2
SOLID EARTH STUDIES LABORATORY (SESL), DEPARTMENT OF GEOLOGY, UNIVERSITY OF REGINA, REGINA,
SASKATCHEWAN, S4S 0A2, CANADA
RECEIVED SEPTEMBER 4, 2007; ACCEPTED MARCH 26, 2008
ADVANCE ACCESS PUBLICATION APRIL 25, 2008
The EoceneTamazeght complex, High Atlas Mountains, Morocco is
a multiphase alkaline to peralkaline intrusive complex. A large variety of rock types (including pyroxenites, glimmerites, gabbroic to
monzonitic rocks, feldspathoidal syenites, carbonatites and various
dyke rocks) documents a progression from ultramafic to felsic magmatism. This study focuses on the silicate plutonic members and the
genetic relationships between the various lithologies. Based on
detailed petrographic and mineral chemical data we show that the
various units crystallized under markedly different oxygen fugacity
and silica activity conditions and demonstrate how these parameters
influence both the phase assemblage and the detailed chemical evolution of the fractionating phases. Nepheline, olivine^clinopyroxene
and hornblende^plagioclase thermometry indicate equilibration
temperatures 8008C for all major rock types. Highly oxidized
conditions (close to the hematite^magnetite buffer) are characteristic of the garnet-rich pyroxenites, ultrapotassic glimmerites and
associated olivine-shonkinites. The parental magmas to these rocks
evolved from low initial aSiO2 values of 01 to values of 05^08
during nepheline and alkali feldspar saturation. In contrast, the
monzonitic rocks evolved from initially high aSiO2 values (up to
075) down to about 01 at intermediate values of oxygen fugacity
*Corresponding author. E-mail: [email protected]
(FMQ ¼ þ2^5 to 1, where FMQ is the fayalite^magnetite^
quartz buffer). For nepheline syenites and malignites, more reduced
conditions (FMQ ¼ 2) and intermediate aSiO2 values (between
025 and 05) dominate. We conclude that fractional crystallization
is not a likely mechanism to explain the large variety of lithologies
present in theTamazeght complex. It is more probable that successive
melting of a compositionally heterogeneous mantle source region gave
rise to several melt batches with distinct chemical and physicochemical characteristics. Low-degree melts from a K-phase-bearing
mantle domain resulted in the formation of ultrapotassic glimmerites,
whereas garnet-rich pyroxenites and olivine-shonkinites may have
originated from hybrid melts and partly from a pyroxene-dominated
source. Less alkaline lithologies such as monzonites potentially reflect
larger degrees of melting and the increased importance of a basaltic
component, whereas nepheline syenites and malignites may be
explained by lower degrees of melting and a more alkaline character
for the parental melt of these rocks.
KEY WORDS: Tamazeght; Morocco; alkaline magmatism; source
heterogeneity;Ti-bearing andradite
The Author 2008. Published by Oxford University Press. All
rights reserved. For Permissions, please e-mail: journals.permissions@
oxfordjournals.org
JOURNAL OF PETROLOGY
VOLUME 49
I N T RO D U C T I O N
Alkaline to peralkaline igneous rocks represent a volumetrically small, but mineralogically highly variable, group
typically located within intracontinental extensional settings. Chemically, these rocks are characterized by high
contents of alkalis and incompatible elements, particularly
the high field strength elements (HFSE; such as Ti, Zr, Hf
and Nb). The residual fluids of such rock associations are
known to give rise to a number of exotic mineral associations in pegmatites and hydrothermal veins (e.g. Salvi &
Williams-Jones, 1990; Khomyakov, 1995; Chakhmouradian
& Mitchell, 2002), some of them being of economic interest
(Srensen, 1992).
The exceptional geochemical character of alkaline to
peralkaline igneous rocks is reflected by an unusual phase
assemblage and by the chemical composition of these
phases; otherwise less-common minerals can appear as
major constituents. For example, Ti-bearing andradite is
commonly found in ultramafic alkaline lithologies (e.g.
Coulson et al., 1999; Vuorinen et al., 2005) and eudialytegroup minerals (Na^Ca-zircono- and titanosilicates) are
typical of highly evolved agpaitic nepheline syenites
(e.g. Srensen, 1997; Mitchell & Liferovich, 2006). It has
been shown that the evolution of intensive parameters
(e.g. fO2, aSiO2) during the crystallization of such rock
types significantly influences the chemical composition of
the phases present (e.g. Jones & Peckett, 1980; Coulson,
2003; Marks & Markl, 2003; Mann et al., 2006).
The association of ultramafic pyroxenites, leucocratic ijolites, and highly evolved nepheline syenites carbonatites is a common feature of alkaline plutonic
complexes world-wide (e.g. Harmer, 1999; Dunworth &
Bell, 2001; Vuorinen et al., 2005). Petrological and geochemical studies have revealed two principal genetic
relationships in such complexes: either (1) closed-system
fractionation of a common parental magma produces the
various lithologies (e.g. Beccaluva et al., 1992; Markl et al.,
2001; Marks et al., 2004; Halama et al., 2005) or (2) the
various lithologies represent crystallization of magmas
derived from different sources, or are related to each other
by combined assimilation^fractionation^mixing processes
(e.g. Kramm & Kogarko, 1994; Morikiyo et al., 2000;
Arzamastsev et al., 2006).
The Tamazeght complex, which is the focus of this study,
comprises numerous intrusive phases that document a progression from ultramafic to felsic alkaline to peralkaline
rock types. A wide range of lithologies is present, including
pyroxenites, glimmerites, gabbroic to monzonitic rocks,
and predominating feldspathoidal syenites. Additionally,
several carbonatitic diatremes and dyke rocks of lamprophyric, carbonatitic, phonolitic and foiditic composition
occur throughout the complex and its sedimentary cover
(Agchmi, 1984; Bouabdli et al., 1988; Mourtada et al., 1997;
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JUNE 2008
Neukirchen & Markl, in preparation). These are not, however, the focus of this work.
The large variety of rock types present in the Tamazeght
complex questions the possibility that these rocks were
derived from one parental magma by fractional crystallization alone; thus important questions concerning the
origin and the genetic relationships between the different
lithologies remain to be answered. Until now, there has
been no systematic study of the mineral chemical and petrological evolution of the Tamazeght rocks. In this study
we investigate in detail the chemical evolution of the fractionating phases to derive crystallization conditions for the
various rock types in terms of oxygen fugacity (fO2) and
silica activity (aSiO2). We further show how these parameters influence the chemical evolution of the mineral
phases present and how such investigations are useful in
deciphering the role of chemically different source components for such multiphase intrusive complexes.
GEOLOGIC A L S ET T I NG A N D
P R EV IO U S WOR K
The Tamazeght complex (also known as Tamazert complex) is located in the Moroccan High Atlas Mountains,
about 20 km south of the city of Midelt (Fig. 1). Here, in
the northern range of the High Atlas, NE^SW-striking
dome and trough structures are the dominant structural
features. Jurassic to Cretaceous marine sediments were
deposited in intra-continental pull-apart basins that are
related to the opening of the Atlantic Ocean (Laville, 1981;
Laville & Harmand, 1982). The Tamazeght complex is the
largest of several alkaline intrusions associated with these
graben structures. It intrudes Liassic marbles and crops
out as an elongated body (16 km NE^SW and 5 km
NW^SE) following the trend of the graben. Along the
medial axis of the complex, Mesozoic marbles form the
roof of the intrusion (Agard, 1960).
Numerous intrusive phases in the Tamazeght complex
document a progression from ultramafic to felsic magmatism. Kchit (1990) proposed the following chronology
based on structural features, cross-cutting relationships
and enclaves within the various units: (1) ultramafic rocks
(pyroxenites & glimmerites); (2) shonkinites; (3) monzogabbroic rocks; (4) foid-monzosyenites; (5) malignites and
associated pegmatites; (6) a range of textural varieties of
nepheline syenites; (7) monzonitic rocks; (8) several carbonatite diatremes associated with carbonatitic and phonolitic
to foiditic dyke rocks; (9) a lamprophyric dyke swarm.
All the intrusive units show vertical or near-vertical
planar internal structures. Close to their margins these
tend to be oriented parallel to the contacts, which are also
sub-vertical. This led Kchit (1990) to conclude that the
Tamazeght intrusive units represent irregular pipe-shaped
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MARKS et al.
TAMAZEGHT COMPLEX, MOROCCO
4°40′
4°35′
ANOUGAL
Rabat
Fès
Casablanca
Meknès
Midelt
Tamazeght
Marrakesh
M
o
or
cc
o
32°35′
ultramafic group
monzogabbro
N
monzonite
foid-monzosyenite
shonkinite
malignite
foyaitic nepheline syenite
porphyritic nepheline syenite
granular nepheline syenite
carbonatite
subvolcanic breccia
various dyke rocks
32°30′
0
4°40′
1
2 km
Country rocks:
Jurassic gabbro
Mesozoic carbonatites
Fig. 1. Geological sketch map of the Tamazeght complex, Morocco [modified after Kchit (1990)]. The small village of Anougal in the
northeastern corner is the gateway to the complex.
bodies, which cross-cut each other. These magmatic ‘pipes’
were interpreted to represent magma in-fills of crustal fractures created by the same SW^NE sinistral shearing that
characterizes the post-Cretaceous Atlas folding (Laville &
Harmand, 1982). The presence of roof pendants, numerous
pegmatites and contact metamorphic minerals within the
surrounding marbles suggests that these magmatic bodies
intruded to shallow depths of 53 km (Salvi et al., 2000).
Radiometric ages of 44 4 Ma (Rb/Sr) and 42 3 Ma
(K/Ar) (Tisserant et al., 1976) have been determined for
some of the monzonites. Nephelinitic dyke rocks, however,
have an age of 35 Ma (Klein & Harmand, 1985). This relatively large time gap led Khadem Allah et al. (1998) to
question the genetic relationship between the various
intrusive phases. Nevertheless, based on geochemical data,
Bouabdli et al. (1988) and Kchit (1990) assumed that all the
rock units originated by fractional crystallization of
a common parental magma of nephelinitic or monchiquitic composition. This parental magma was considered
to have originated by low-degree partial melting of a
carbonated amphibole-lherzolite mantle source. The carbonatites were thought to have formed through liquid
immiscibility (Bouabdli et al., 1988).
The most recent studies of the Tamazeght complex
focused on the fenitizing effects of carbonatitic fluids
(Bouabdli & Liotard, 1999; Neukirchen & Markl, in preparation), on the influence of sedimentary carbonate rocks
on the evolution of the peralkaline to agpaitic pegmatites
of the complex (Khadem Allah et al., 1998), on the
hydrothermal mobilization of HFSE within some of the
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VOLUME 49
P E T RO G R A P H Y
In this section, we describe the phase assemblages observed
in the various lithologies and the micro-textural characteristics of the investigated samples. Figure 2 gives an overview of the mineral assemblages present.
The ultramafic group
pyroxenites
glimmerites
monzogabbros
monzonites
foid-monzosyenites
olivine-shonkinites
amphibole-shonkinites
porphyritic nepheline syenites
granular nepheline syenites
foyaitic nepheline syenites
miaskitic malignites
agpaitic malignites
pegmatites and veins
Fig. 2. Summary of the mineral assemblages observed in the various Tamazeght rocks.
1100
nepheline / sodalite
alkali feldspar
plagioclase
calcite
amphibole
Two types of ultramafic rocks can be distinguished:
(1) pyroxenites consisting of variable amounts of clinopyroxene, nepheline and garnet; (2) glimmerites, which
are dominated by biotite.
Pyroxenites (TMZ23b, 23c and 25) are dominated by
euhedral clinopyroxene (Cpx), nepheline and euhedral
to subhedral garnet (Fig. 3a and b). Minor phases are
apatite, calcite, mica, magnetite, titanite and pyrite
garnet
Fe-Ti oxides
pyroxene
olivine
Field relations between the various rock units were
described in great detail by Kchit (1990) and Al-Haderi
et al. (1998), constraining the chronological order of emplacement noted above. However, our own field work in
the Tamazeght complex has shown that the distinction
between the rock types in the field is not as obvious as
indicated on the geological map of Kchit (1990) and the
labeled rock types are not exclusively restricted to the
areas indicated. For example, according to Kchit (1990)
monzogabbro occurs as only one distinct body in the
northeastern part of the complex (Fig. 1). However, the
detailed petrographic investigation of our own samples
(4150 specimens) has revealed that, for example, monzogabbroic rocks also occur within the foid monzosyenitic
unit and vice versa. Obviously, the ratio between plagioclase and alkali feldpar is highly variable within single
rock units on a small scale and clear intrusive contacts
between them are only rarely visible in the field. If present,
these are in many cases gradual without sharp contacts,
and thus provide evidence for only a rather short time
gap between the emplacement of the various rock units.
Thus, here we treat monzogabbros, monzonites and foid
eudialyte
F I E L D R E L AT I O N S
JUNE 2008
monzosyenites as a single group; namely, the monzonitic
group. All of these rock types are characterized by the
occurrence of both feldspar types and we classify them
based on their plagioclase:alkali feldspar ratio, irrespective
of where they have been collected in the field. Similarly,
the nepheline syenitic group shows heterogeneities in
terms of grain size, mineral assemblage and macroscopic
textures. In all, this indicates that most of these rocks
were probably emplaced within a rather short time interval as a crystal mush, possibly also allowing for mixing
and mingling among them.
mica
nepheline syenites (Salvi et al., 2000, 2001) and the compositional variation of clinopyroxene in some of the nepheline
syenites (Khadem Allah et al., 1996).
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titanite / zircon
JOURNAL OF PETROLOGY
MARKS et al.
TAMAZEGHT COMPLEX, MOROCCO
(occasionally with inclusions of pyrrhotite). Magnetite is
locally transformed to hematite (Fig. 3c). Kchit (1990)
also reported the rare occurrence of olivine. Amphibole
and feldspar are absent in these rocks. With increasing
nepheline content, some of the rocks are classified as
mafic foidolites (melteigites and ijolites). Cumulus minerals are colourless to pale green clinopyroxene (showing
discontinuous zonation patterns; Fig. 3d), nepheline, oscillatory zoned garnet and magnetite. Nepheline and garnet
have a prolonged crystallization interval and are also
present as intercumulus phases (Fig. 3b). Along the rims,
nepheline is in places altered to cancrinite and/or sodalite.
In one sample (TMZ23b) a zone several centimetres wide
consisting of euhedral calcite, analcime and a late clinopyroxene generation is observed to cross-cut garnet-rich
pyroxenite. Pyroxene in this vein is bright green and very
fine-grained (generally 5100 mm) and occurs as radiating
clusters.
In glimmerites (TMZ20 and 22), poikilitic biotite is
the dominant phase, at up to 65 modal %. Early minerals
Fig. 3. Typical microtextures observed in ultramafic rocks from the Tamazeght complex. (a) Typical cumulate texture showing euhedral
garnet (grt) with nepheline inclusion (ne) and interstitial calcite (cal) in garnet-rich pyroxenite TMZ23b. (b) Interstitial nepheline occurs
together with euhedral clinopyroxene (cpx) and garnet (TMZ25). (c) BSE image of a magnetite (mag) grain partly transformed to hematite
(hem) (TMZ23b). (d) In all pyroxenites, clinopyroxene shows discontinuous zonation, displaying three distinct pyroxene compositions
(TMZ25). (e) In glimmerites early clinopyroxene-I is transformed to a mixture of fine-grained biotite and magnetite (TMZ20).
(f) Throughout the glimmerites, ocelli-like textures, consisting of granular clinopyroxene-II in the outer parts and of interstitial calcite and
magnetite, are observed (TMZ22).
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are clinopyroxene-I, garnet and minor perovskite, magnetite (occasionally with cores of chromite) and
apatite. Feldspar, nepheline and amphibole are lacking.
Compared with the pyroxenites, garnet is not euhedral
but occurs in subhedral to anhedral granular aggregates.
Clinopyroxene-I is locally rimmed or even replaced by
a mixture of fine-grained mica and magnetite (Fig. 3e);
early perovskite is always overgrown by titanite. Compositionally zoned ocelli-like textures, which occur
throughout the rocks, consist of granular and colourless
clinopyroxene-II in the outer parts and of interstitial calcite and/or magnetite in their cores (Fig. 3f). Commonly,
magnetite is replaced by pyrite and hematite. Other
opaque minerals include sphalerite and chalcopyrite. In
places, a third pyroxene generation (Cpx-III) crystallized
interstitially with respect to clinopyroxene-II.
The monzonitic group
This group of rocks is characterized by the occurrence of
both plagioclase and alkali feldspar in addition to foid
minerals (nepheline and minor sodalite and cancrinite).
Based on their relative modal abundance, it is subdivided
into foid (-bearing) monzogabbros, (foid-bearing) monzonites, (foid-bearing) syenites and foid-monzosyenites.
Monzogabbros are generally rich in euhedral to subhedral grey to pale green clinopyroxene and magnetite with
minor amounts of ilmenite. However, both amphibolerich and biotite-rich varieties exist. In amphibole-rich varieties (TMZ159), minor pyroxene is commonly overgrown
by reddish brown amphibole, and in places, small rounded
relics of clinopyroxene can be seen within euhedral amphibole (Fig. 4a). Both minerals contain subhedral inclusions
of magnetite, and titanite occurs as subhedral crystals and
as narrow (5200 mm wide) rims overgrowing earlier
ilmenite (Fig. 4b). In biotite-rich varieties (TMZ320), subhedral biotite is commonly associated with pyroxene, but is
never seen to overgrow or to resorb it, unlike amphibole
in amphibole-rich varieties. Also, biotite hosts inclusions
of subhedral magnetite and needles of apatite. In these
varieties, titanite is much more abundant and occurs exclusively as large (up to 2 mm) subhedral to euhedral crystals,
occasionally with rounded inclusions of ilmenite (Fig. 4c).
Monzonites are porphyritic with euhedral phenocrysts
of plagioclase and alkali feldspar set in a medium-grained
matrix of clinopyroxene, magnetite, ilmenite, amphibole,
biotite, titanite and feldspars. The grain size ranges from
several centimetres to 51mm. Accessory minerals are
apatite and zircon. Subhedral to euhedral pyroxene
co-crystallized with Fe^Ti oxides and titanite. In common
with the monzogabbros, these rocks initially crystallized
magnetite and ilmenite, the latter of which almost exclusively occurs as partly resorbed inclusions within titanite.
Biotite is commonly corroded and shows a rim of finegrained magnetite (Fig. 4d). Subordinate amphibole is
NUMBER 6
JUNE 2008
subhedral and some of the amphibole cores host tiny
patches of exsolved Fe^Ti oxides.
In foid-monzosyenites the relative modal amounts of
pyroxene and amphibole are highly variable. Generally,
euhedral pyroxene is pale grey to green in the core and has
distinct bright green to yellow^green outer parts that show
patchy heterogeneities. Euhedral titanite and subhedral
magnetite (now coarsely exsolved to ilmenite and magnetite) appear to have co-crystallized with clinopyroxene
(Fig. 4e) and both occur as inclusions in amphibole. No primary ilmenite was found in these rocks. Additionally, most
amphibole shows the above-mentioned exsolution textures
and in some samples (TMZ157 andTMZ312), a late pyroxene
population overgrows earlier amphibole (Fig. 4f). Biotiterich varieties typically are amphibole-free (TMZ318), but
in samples with both phases (TMZ313), biotite predominates, occurring as rounded inclusions in amphibole or
as a complex intergrowth.
The foid syenitic group
Following the IUGS nomenclature, these rocks are
subdivided into shonkinites, nepheline syenites and
malignites, based on the proportion of mafic minerals
(Le Maitre, 2002). Mafic minerals include clinopyroxene, amphibole, perovskite^titanite, apatite, olivine,
biotite, magnetite, eudialyte, zircon, garnet,
calcite, fluorite. Primary felsic minerals include alkali
feldspar, nepheline and sodalite. Pure albite occurs as
a late magmatic phase.
Shonkinites have a colour index 460 (Le Maitre, 2002)
and are subdivided into olivine-bearing and amphibolerich varieties. Olivine-shonkinites (TMZ12 and 130) are
characterized by large (1^5 mm) phenocrysts of olivine,
pale grey clinopyroxene-I and magnetite-I set in a finegrained groundmass of greenish clinopyroxene-II, garnet,
apatite, alkali feldspar and nepheline. The last mineral is
in most cases strongly altered to calcite, analcime, cancrinite and sodalite. Olivine phenocrysts have rounded grain
boundaries and are partly rimmed by a fine-grained
mixture of magnetite-II, amphibole and biotite (Fig. 5a);
clinopyroxene-I is commonly overgrown or invaded by
clinopyroxene-II (Fig. 5b). If present, perovskite is
rimmed by titanite (Fig. 5c). However, titanite also occurs
as rims around magnetite-I and, occasionally, also as
euhedral grains. Garnet is present as a minor phase in
one sample (TMZ12).
Amphibole-shonkinites (TMZ68, 139 and 308) are
coarse-grained and do not show any noticeable phenocrysts. Here, euhedral amphibole (up to 5 mm) strongly
dominates over clinopyroxene and is commonly associated
with subhedral biotite (Fig. 5d). Minor euhedral clinopyroxene is finer grained (generally 52 mm) than amphibole
and grey to pale green in colour. Occasionally, it shows
irregular and patchy heterogeneities, where the outer
regions of the crystals are more greenish and the inner
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MARKS et al.
TAMAZEGHT COMPLEX, MOROCCO
Fig. 4. Typical micro-textures observed in rocks of the monzonitic group. (a) In amphibole-rich monzogabbro TMZ159 small rounded relics of
clinopyroxene occur as inclusions within reddish brown euhedral amphibole (amph). (b) In the same sample primary Fe^Ti oxide (ox) grains
are rimmed by thin seams of titanite (ttn). (c) In biotite-rich monzogabbro TMZ320 subhedral titanite commonly has rounded relics of Fe^Ti
oxides and euhedral apatite needles as inclusions. (d) In monzonites, biotite is corroded and exhibits a rim of fine-grained Fe^Ti oxide (TMZ2).
(e) A typical texture in foid-monzosyenite TMZ219 showing euhedral titanite and subhedral Fe^Ti oxides coexisting with pyroxene.
(f) In amphibole-bearing foid-monosyenite TMZ157, a late pyroxene population overgrows earlier amphibole.
regions are more greyish in colour. Olivine and perovskite
are absent and titanite is always euhedral. Magnetite and
minor ilmenite occur either as subhedral grains associated
with clinopyroxene, amphibole and titanite (TMZ139
and 308) or as subhedral to anhedral rounded grains as
inclusions in these three minerals (TMZ68). Petrographically this group shows similarities to some of the monzonitic rocks.
Nepheline syenites have a colour index 530 (Le Maitre,
2002) and, based on their general texture, a number of
varieties can be distinguished.
Foyaitic nepheline syenites (TMZ165, 221 and 223) are
generally coarse-grained and consist of a framework of
large alkali feldspar laths (up to 5 mm) associated with
euhedral to subhedral nepheline and minor sodalite, the
latter of which is strongly altered. Locally, interstitial
albite also occurs. Euhedral clinopyroxene with grey cores
and distinct yellow^greenish rims forms larger aggregates
and is commonly associated with euhedral titanite, apatite,
magnetite and biotite (Fig. 6a). In samples, which were
collected in the vicinity of carbonatite dykes, biotite is
commonly intergrown with clinopyroxene and appears
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Fig. 5. Typical textures observed in shonkinitic rocks from the Tamazeght complex. (a) Rounded olivine phenocryst rimmed by a fine-grained
mixture of Fe^Ti oxides, amphibole and biotite (TMZ130). (b) Phenocryst of clinopyroxene-I overgrown by clinopyroxene-II, which also occurs
as subhedral to euhedral grains in the groundmass (TMZ12). (c) Early perovskite (prv) rimmed by anhedral titanite (TMZ12; BSE image).
(d) Amphibole-shonkinites do not show any noticeable phenocrysts and are more coarse-grained than olivine-shonkinites. In these rocks, subhedral to euhedral amphibole (amph) strongly dominates over clinopyroxene and is commonly associated with titanite and apatite (TMZ68).
to replace it (Fig. 6b). Primary subhedral amphibole in
these rocks is rare. If present, it shows fine-grained exsolution textures in the core region (TMZ221).
In granular nepheline syenites (TMZ74, 94, 95 and 126)
subhedral clinopyroxene shows grey cores with distinct
greenish coloured rims (Fig. 6c) and is associated with
euhedral titanite and magnetite. Euhedral to subhedral
amphibole is brown to dark green in colour and shows
similar exsolution textures to those in the foyaitic varieties.
Locally, it is overgrown by green fine-grained clinopyroxene (Fig. 6d). Clinopyroxene, amphibole and titanite contain tiny needle-shaped inclusions of apatite.
Porphyritic nepheline syenites (TMZ311) consist of
centimetre-sized euhedral alkali feldspar phenocrysts with
euhedral to subhedral nepheline, clinopyroxene, amphibole, titanite and rounded magnetite filling the space
between them. Clinopyroxene is generally pale green,
showing no distinct greenish rim but a patchy inhomogeneity. As in the foyaitic and granular varieties, amphibole is subhedral to euhedral and also shows characteristic
exsolution in the core regions (Fig. 6e).
Malignites have a colour index of 30^60 (Le Maitre,
2002) and are generally amphibole- and biotite-free.
Occasionally, more leucocratic varieties exist. However, to
distinguish this rock type from the other foid syenites,
we call them malignites throughout this work. Based on the
presence of eudialyte or lafivenite [simplified formula
(Na,Ca)2(Mn2þ,Fe2þ)(Zr,Ti,Nb)Si2O7(O,OH,F)], they
are subdivided into miaskitic and agpaitic varieties.
Euhedral apatite and titanite occur as inclusions in clinopyroxene, nepheline or alkali feldspar. Mostly euhedral green
to yellow^green clinopyroxene (up to 5 mm in size) shows
irregular heterogeneities throughout most samples (Fig. 6f).
Nepheline and alkali feldspar are both subhedral in habit,
and sodalite occurs as an interstitial phase. In some spatially
restricted areas eudialyte (Fig. 6g) or lafivenite were formed
during the late-magmatic stage, accompanied by felty
clinopyroxene-II and small albite laths (Fig. 6h). Late-stage
hydrothermal processes are documented by the formation
of symplectitic cancrinite^sodalite seams around precursor
nepheline. Within the malignites, a number of pegmatites
and hydrothermal veins are recognized. The pegmatites
have been intensively studied by Khadem Allah et al.
(1998) and, thus, we investigated only one aegirine-rich
pegmatite sample (TMZ247) for this study. It consists of
euhedral centimetre- to decimetre-sized yellow^green
1104
MARKS et al.
TAMAZEGHT COMPLEX, MOROCCO
Fig. 6. Typical microtextures observed in nepheline syenites of the Tamazeght complex. (a) In foyaitic nepheline syenites TMZ165 and 221,
euhedral clinopyroxene (cpx) with pale grey cores and distinct green rims is associated with euhedral titanite (ttn), apatite, Fe^Ti oxide (ox)
and biotite (bt). (b) In foyaitic nepheline syenite sample TMZ223, which was collected close to a carbonatitic dyke, pyroxene is replaced
by biotite (bt). (c) Discontinuously zoned pyroxene crystal in granular nepheline syenite TMZ94. (d) Twinned amphibole crystal in granular
nepheline syenite TMZ74 overgrown by pale green pyroxene. (e) In porphyritic nepheline syenite (TMZ311) amphibole shows characteristic
exsolution of tiny Fe^Ti oxide needles in the core. This texture is observed in most rock types of the Tamazeght complex. (f) In malignites,
pyroxene is subhedral to euhedral and shows irregular small-scale heterogeneities (TMZ233; BSE image). (g) Agpaitic malignite (TMZ176)
characterized by the coexistence of alkali feldspar, pyroxene, nepheline and eudialyte (eud); ab, albite. (h) In most malignitic samples a late
generation of pure albite occurs along grain boundaries of earlier alkali feldspar (TMZ295).
1105
JOURNAL OF PETROLOGY
VOLUME 49
sector-zoned clinopyroxene. Locally, hematite occurs interstitially between the pyroxene and as rounded inclusions
within pyroxene crystals. Minor minerals are alkali feldspar
and eudialyte. Hydrothermal veins (TMZ177, 229, 231 and
234) are generally several centimetres wide. They consist
of alkali feldspar laths several centimetres in size and
yellow^green pyroxene needles of similar length. Accessory
minerals are zircon or eudialyte, catapleite, magnetite and
Nb- and Mn-rich ilmenite.
NUMBER 6
JUNE 2008
Table 1: Representative electron microprobe analyses of
olivine from the olivine-shonkonite, Tamazeght Complex,
Morocco
Sample:
TMZ12
TMZ12
TMZ130
TMZ130
3874
SiO2
4032
3985
4022
Al2O3
003
003
005
002
FeO
963
1254
1244
2129
MINER A L COMPOSITIONS
Analytical techniques
MnO
016
020
022
113
MgO
4928
4629
4684
3873
The major and minor element compositions of the
constituent minerals were determined using a JEOL 8900
electron microprobe in wavelength-dispersion mode at
the Institut fu«r Geowissenschaften, Universita«t Tu«bingen
(Germany). For silicate minerals, we used a beam current
of 15 nA and an acceleration voltage of 15 kV; for Fe^Ti
oxides, we used 20 nA and 20 kV. The peak counting time
was 16 s for major elements and 30^60 s for minor elements.
Background counting times were half of the peak counting
times. The peak overlap between the Fe Lb and F Ka was
corrected for. To avoid Na migration under the electron
beam, analyses of feldspar, nepheline and sodalite were
performed with a defocused beam of 10 mm diameter.
In cases where Fe^Ti oxides showed fine-grained exsolution textures they were analysed with a defocused beam of
20^40 mm diameter. For calibration, both natural minerals
and synthetic phases were used as standards. Processing
of the raw data was carried out with the internal frZ
correction method of JEOL (Armstrong, 1991). Analytical
uncertainties are below 1% relative for major elements
and around 15^20% for minor elements.
The bulk composition of coarsely exsolved Fe^Ti oxide
grains was reconstructed by combining image processing
(NIH Image software) of back-scattered electron (BSE)
images of the exsolved mineral grains with point analyses
of exsolved ilmenite and magnetite. The bulk composition
was then recalculated using the area proportions of both
exsolved phases and using molar volumes of 4452 and
3170 cm3/mol for magnetite and ilmenite, respectively.
Generally, this procedure was applied to 3^5 grains in
each investigated sample.
NiO
037
029
010
014
CaO
021
026
030
035
10000
9946
10017
10040
100
Olivine
Within two samples of olivine-shonkinite, the compositional variation of olivine is small (Fo90^87 in TMZ12
and Fo88^75 in TMZ130; Table 1). Most of this variation is
related to normal growth zonation with decreasing XMg
[Mg/(Mg þ Fe2þ)] from core to rim but essentially
unzoned olivine is also present (Fig. 7). However, the olivine from the two samples differs significantly in terms of
its minor element composition. In TMZ12, the olivine is
relatively high in NiO (up to 04 wt %) but low in CaO
Total
Formula based on 4 oxygen atoms
Si
099
100
100
Al
000
000
000
000
Fe2þ
020
026
026
046
Mn
000
000
000
002
Mg
180
173
173
149
Ni
001
001
000
000
Ca
001
001
001
001
Sum
301
301
300
298
mol % end-members
Fo
90
87
87
76
Fa
10
13
13
24
(503 wt %) and MnO (5025 wt %), whereas in sample
TMZ130, the opposite is the case (NiO 5016 wt %; CaO
504 wt %, MnO 513 wt %).
Clinopyroxene
As is typical for alkaline intrusive complexes, clinopyroxene shows a wide range of compositions. For a detailed
chemical classification, 10 end-members were computed,
assuming stoichiometry (six oxygen atoms and four
cations). Details of the applied calculation scheme are
given in the Appendix. In addition to the Quadcomponents (enstatite [En], ferrosilite [Fs], diopside [Di]
and hedenbergite [Hed]), we include the Na-bearing components aegirine [Aeg], (Ti, Zr)-aegirine [Ti-Aeg], and
jadeite [ Jd]. In most analyses, the calculated Fe3þ content
exceeds the Na content, implying the presence of a ferriTschermak component [Fe-Ts]. The AlIV-bearing components Ca-Tschermak [Ca-Ts] and Ti-Tschermak [Ti-Ts]
are also considered. The variation of these components in
the various lithologies is illustrated in Figs 8 and 9 and
representative clinopyroxene analyses are given in Table 2.
The end-members diopside, hedenbergite and aegirine
are the most important ones to describe the compositional
1106
MARKS et al.
TAMAZEGHT COMPLEX, MOROCCO
TMZ 130
100
TMZ 12
100
Fo
80
80
60
60
40
40
20
20
Fa
Fa
0
0
1.5
1.5
1
MnO
wt. %
wt. %
1
MnO
0.5
0.5
0
0
0.4
0.4
0.3
0.3
NiO
CaO
0.2
0.2
0.1
0.1
CaO
NiO
0
rim
wt. %
wt. %
mol %
mol %
Fo
0
core
rim
≈
rim
core
rim
≈
Fig. 7. Zoning profiles across olivine grains from olivine-shonkinites TMZ130 (left column) and TMZ12 (right column). Olivine in sample
TMZ130 shows considerable chemical zonation, with high amounts of Fo and CaO in the core and high Fa and MnO at the rim. For NiO,
however, no obvious systematic trend is observed. In contrast, olivine from sample TMZ12 is generally unzoned and shows only minor and
unsystematic within-grain heterogeneities. Also, NiO contents are considerably higher compared with olivine from TMZ130, despite a very
similar Fo content (see text for further discussion).
evolution of clinopyroxene from the Tamazeght suite. In all
lithologies, similar Di-rich pyroxene compositions are
found and these evolve towards more Hed- and Aeg-rich
compositions with progressive differentiation. However,
the various rock types show variable relative amounts of
Hed enrichment while evolving towards Aeg-rich compositions, resulting in rather flat evolutionary trends for rocks
of the monzonitic and nepheline syenitic group and comparatively steep trends for shonkinites. Ultramafic rocks
show an intermediate trend. Also, the overall variation of
clinopyroxene compositions observed in one rock type
is highly variable, with clinopyroxene from nepheline
syenites showing by far the largest chemical variation.
Within one sample, the whole trend from Di-rich via intermediate towards Aeg-rich compositions can be traced
(Fig. 8). These differences are of major importance and
will be discussed in detail below.
Intermediate (aegirine^augite) pyroxene compositions
in shonkinites and in some malignites do not follow welldefined trends as is found for most other rock types. These
broad compositional fields can be correlated with irregular
heterogeneities as is evident from BSE images (Fig. 6f).
The minor components Fe-Ts, Ca-Ts and Ti-Ts are generally 510 mol %. Typically, they are lower in ultramafic
rocks than in the other rock types. However, within the
ultramafic rocks, Fe-Ts is relatively enriched in the two
1107
JOURNAL OF PETROLOGY
VOLUME 49
NUMBER 6
JUNE 2008
Aeg
Aeg
pyroxenites
glimmerites
inner zone
cpx-I
cpx-II
cpx-III
intermediate zone
outer zone
late pyroxenes (TMZ 23b)
Di
Hed
Di
Hed
Aeg
Aeg
monzonitic group
shonkinites
monzogabbros
cpx in amph-shonkinites
monzonites
cpx-I in ol-shonkinites
foid-monzosyenites
cpx-II in ol-shonkinites
cpx overgrowing amphibole
Di
Di
Hed
Aeg
Aeg
malignites and
late-stage rocks
nepheline
syenites
miascitic malignites
porphyritic rocks
agpaitic malignites
foyaitic rocks
hydrothermal veins
granular rocks
Di
Hed
pegmatite
Di
Hed
Hed
Fig. 8. Aegirine^diopside^hedenbergite pyroxene triangle illustrating the observed variation of clinopyroxene composition within the various
rock types.
inner zones of discontinuously zoned clinopyroxene from
pyroxenites (Figs 3d and 9). In glimmerites, partly
resorbed phenocrysts of clinopyroxene-I (Fig. 3e) are relatively rich in all three Tschermak components (Fig. 9).
Significant amounts of the Ti-Aeg end-member are generally restricted to Aeg-rich clinopyroxene compositions,
where it may be up to 30^40 mol%.
Zoning profiles for clinopyroxene are different in the
various rock groups. In pyroxenites, clinopyroxene shows
discontinuous zonation, displaying three distinct pyroxene
compositions (Figs 3d and 9). Pyroxene in monzonitic rock
types shows continuous zoning patterns, starting with
Di-rich compositions in the cores and evolving towards
more Hed- and Aeg-bearing compositions at the rims,
coinciding with an increase in Tschermak components.
In granular syenite, clinopyroxene is discontinuously
zoned (Fig. 6c). Within the Di-rich core a sudden increase
and decrease of Fe-Ts and Ti-Ts is observed. The rim
1108
MARKS et al.
TAMAZEGHT COMPLEX, MOROCCO
pyroxenites
glimmerites
monzogabbros
monzonites
foid-monzosyenites
olivine-shonkinites
amphibole-shonkinites
porphyritic nepheline syenites
granular nepheline syenites
foyaitic nepheline syenites
malignites
pegmatites and veins
0
5
10
mole % CaAl[AlSiO6]
0
5
10
15
mole % CaTi[Al2O6]
0
5
10 0 10 20 30 40 50
mole % CaFe[FeSiO6]
mole % NaTi0.5Fe0.5[Si2O6]
Fig. 9. Variation in minor components for the Tamazeght clinopyroxenes.
evolves towards high amounts of the Aeg component but is
low again in Tschermak components. This evolutionary
trend is accompanied by a continuous decrease in Ca-Ts
from core to rim. According to their heterogeneous
appearance (Fig. 6f), zoning profiles for clinopyroxene
from malignites reveal that, despite a rough trend of
increasing Aeg-component from core to rim, the evolution
in XMg and Tschermak components is not strictly
systematic.
Fe^Ti oxides
Fe^Ti oxides show considerable variation in terms of phase
assemblage, composition and exsolution textures among
the various rock types. Figure 10 illustrates this variation
and Table 3 provides representative analyses.
In pyroxenites, primary, almost Ti-free magnetite
(Mag99^100Usp0^1) is commonly replaced by hematite
(Ilm0^2Hem98^100Pyr0^1). In glimmerites, euhedral
opaque phases in the mica-rich matrix are mostly Ti-poor
magnetite (Mag79^93Usp4^8Spl1^13) and in rare cases, these
show distinct Cr-rich cores (Mag14^17Usp3^5Spl78^81), with
an XCr value [Cr/(Cr þAl)] of 075 and an XMg value
of 02. The composition of magnetite from ocelli-like
textures overlaps with the range observed in matrix magnetite (Mag90^95Usp5^10Spl0^1).
In monzogabbros and monzonites, homogeneous and
Ti-poor magnetite (Mag97^100Usp0^2Spl0^1) and ilmenite
(Ilm80^83Hem10^12Pyr7^10) were observed; the latter is
much less abundant and occurs almost exclusively as
rounded inclusions in titanite or is overgrown by the
latter. In foid-monzosyenites, Ti-bearing magnetite
(Mag67^83Usp15^27Spl2^5) shows coarse sandwich-type
exsolution textures, but primary ilmenite is absent.
In olivine-shonkinites, ilmenite is absent but the magnetite composition is relatively variable (Mag55^90Usp6^40
Spl4^7). In amphibole-shonkinites both Ti-bearing
magnetite (Mag88^98Usp1^11Spl0^4) and ilmenite (Ilm84^87
Hem8^11Pyr4^7) are present. In nepheline syenites, the
composition of magnetite (Mag76^99Usp1^23Spl0^2) shows
no systematic variation between the textural varieties.
In miaskitic malignites, magnetite (Mag83^99Usp1^17
Spl0^1) is present; this is, however, lacking in the agpaitic
varieties. Late-stage hydrothermal veins contain either
Mn-rich ilmenite (Ilm46^49Hem1^2Pyr50^53 ; TMZ234) or
Ti-poor magnetite (Mag86^98Usp2^13Spl0^1 ; TMZ229).
Generally, V2O3 contents are higher in ilmenite (up to
about 3 wt %) than in magnetite (506 wt %) but no systematic differences between the various rock types were
observed. The highest Cr2O3 contents were found in the
cores of spinel grains from glimmerites (up to 386 wt %);
however, the vast majority of the magnetite in the ultramafic rocks contains 53 wt % Cr2O3 . In shonkinites, magnetite contains up to 15 wt % Cr2O3, but in all other rock
types, Cr2O3 contents in magnetite and ilmenite are
much lower (504 and 501wt %, respectively). ZnO contents are generally below 51wt %, with no obvious difference between magnetite, hematite or ilmenite nor with
any systematic evolution within the complex. Only in the
Cr-rich spinel of the glimmerites were increased ZnO
contents (up to 32 wt %) detected. For all analyses, ZrO2
contents are below detection limit, as are Nb2O5 contents,
except for the Mn-rich ilmenite from hydrothermal vein
TMZ234, where up to 16 wt % of Nb2O5 was detected.
Garnet
Garnet in the Tamazeght rocks is generally rich in Ca,
Fe3þ and Ti, covering the compositional range between
Ti-bearing andradite and schorlomite (Ca3Ti4þ2[Si3^x
(Fe3þ,Al,Fe2þ)x]O12). Representative compositions are
reported in Table 4. The nomenclature concerning schorlomite is somewhat controversial [see Chakhmouradian &
McCammon (2005) for a recent discussion]. In the absence
1109
Table 2: Representative electron microprobe analyses of clinopyroxene from theTamazeght Complex, Morocco
Rock
Pyroxenite
Glimmerite
Monzogabbro
Monzonite
Foid-monzosyenite
TMZ159
TMZ2
TMZ321
Olivine-shonkinite
type:
Sample:
TMZ25
TMZ25
TMZ25
TMZ23b
TMZ22
TMZ22
TMZ22
inner zone
intermediate
outer zone
late cpx
cpx-I
cpx-II
cpx-III
TMZ320
TMZ313
TMZ312
TMZ157
TMZ219
TMZ130
TMZ12
cpx-I
cpx-II
5271
5178
5236
5210
4700
5425
5189
4856
5012
5143
4779
4944
4920
4587
4844
4528
5041
TiO2
032
040
011
476
306
022
385
195
070
052
064
138
222
083
251
336
017
ZrO2
000
000
001
066
004
003
073
002
000
002
017
010
002
074
003
001
005
Al2O3
117
111
041
077
532
005
032
388
230
154
383
386
400
425
512
691
053
b.d.
b.d.
b.d.
b.d.
b.d.
Cr2O3
b.d.
b.d.
b.d.
Fe2O3
295
612
842
2299
419
076
2713
528
645
482
678
557
312
965
363
601
1407
FeO
339
251
359
478
167
220
221
220
268
350
1063
423
408
901
316
137
1039
MnO
051
065
104
018
005
020
016
021
045
061
097
046
020
104
012
014
082
MgO
1450
1348
1096
042
1415
1725
029
1338
1272
1377
652
1211
1362
557
1380
1289
353
CaO
2411
2353
2051
022
2413
2409
058
2284
2278
2362
2052
2272
2267
1987
2328
2326
1582
Na2O
057
106
264
1310
024
023
1343
093
120
065
175
108
069
224
054
062
K2O
000
000
000
007
000
000
000
000
000
001
006
001
002
006
000
10023
10064
10005
10005
9985
9928
10059
9925
9940
10048
9965
10096
9984
9913
10061
9985
10068
Total
489
000
JOURNAL OF PETROLOGY
SiO2
Formula based on 4 cations and 6 oxygen atoms
197
199
175
199
198
182
189
191
186
184
183
181
179
171
196
001
000
014
009
001
011
005
002
001
002
004
006
002
007
009
000
Zr
000
000
000
001
000
000
001
000
000
000
000
000
000
001
000
000
000
Al
005
005
002
003
023
000
001
017
010
007
018
017
018
020
022
030
002
Cr
000
000
000
000
000
000
000
000
000
000
000
000
000
000
000
000
000
Fe3þ
008
017
024
066
012
002
078
015
018
013
020
016
009
029
010
017
041
Fe2þ
010
008
011
015
005
007
007
007
008
011
035
013
013
030
010
004
034
Mn
002
002
003
001
000
001
001
001
001
002
003
001
001
003
000
000
003
Mg
080
074
061
002
078
093
002
074
071
076
037
067
075
033
076
072
020
Ca
095
093
082
001
096
095
002
092
092
094
086
090
090
084
092
093
066
Na
004
008
020
098
002
002
099
007
009
005
013
008
005
017
004
004
038
K
000
000
000
000
000
000
000
000
000
000
000
000
000
000
000
000
000
Sum
400
400
400
400
400
400
400
400
400
400
400
400
400
400
400
400
400
mol % end-members
Aeg
4
8
20
67
2
2
81
7
9
5
13
8
5
17
4
5
Jd
0
0
0
2
0
0
0
0
0
0
0
0
0
0
0
0
38
0
Ti-Aeg
0
0
0
29
0
0
19
0
0
0
0
0
0
0
0
0
0
Fe-Ts
2
5
2
0
5
0
0
4
5
4
3
4
2
6
3
6
2
Ti-Ts
1
1
0
0
9
0
0
6
2
1
2
4
6
4
7
9
1
Al-Ts
2
1
1
0
3
0
0
3
3
2
6
4
2
6
4
6
0
Di
79
75
62
0
74
87
0
72
71
75
37
65
68
33
69
68
21
Hed
12
10
15
0
5
7
0
7
10
12
37
14
12
34
9
4
38
En
0
0
0
0
2
4
0
1
0
1
1
1
4
0
4
2
0
Fs
0
0
0
2
0
0
0
0
0
0
1
0
1
0
0
0
0
(continued)
JUNE 2008
192
001
NUMBER 6
195
Ti
VOLUME 49
1110
Si
Table 2: Continued
Rock-type:
Amphibole-shonkinite
Porph.
Foyaitic nepheline syenite
Granular nepheline syenite
Miaskitic malignite
TMZ94
TMZ94
TMZ94
TMZ82
TMZ82
Agpaitic malignite
Vein
Pegmatite
TMZ229
TMZ247
TMZ247
neph syenite
Sample:
TMZ68
TMZ311
TMZ165
SiO2
TiO2
TMZ165
TMZ223
TMZ178
TMZ298
5091
4720
4985
5030
5032
5293
4378
4859
4847
4476
148
273
068
155
029
044
388
167
170
381
ZrO2
000
000
011
002
033
048
010
000
004
Al2O3
278
546
284
309
132
130
787
556
Cr2O3
043
049
b.d.
b.d.
b.d.
b.d.
b.d.
Fe2O3
223
522
624
407
1363
2885
FeO
259
242
722
427
1030
043
MnO
007
021
050
026
131
MgO
1558
1208
956
1288
CaO
2345
2349
2139
Na2O
037
100
172
TMZ238
TMZ238
5182
5145
5099
5219
5055
5221
5015
5215
026
096
029
128
098
036
084
040
008
020
078
024
016
003
061
031
020
564
774
088
112
076
107
107
083
105
127
045
033
b.d.
b.d.
b.d.
b.d.
b.d.
b.d.
b.d.
b.d.
b.d.
627
402
477
521
1546
2864
1696
2364
2859
2937
1914
2765
117
151
120
261
567
083
1033
357
006
043
905
158
066
009
010
012
013
103
069
107
094
102
124
101
078
276
108
1227
1456
1442
1199
527
013
088
156
089
033
083
059
2289
1393
284
2338
2228
2250
2305
1333
266
1094
578
549
272
950
350
102
564
1236
061
075
079
070
627
1243
740
1053
1110
1237
801
1184
000
000
004
000
002
000
000
000
000
003
002
002
002
004
006
000
001
001
Total
9990
10029
10016
10036
9984
10137
9941
9949
9998
10011
10021
9971
9988
10076
9984
10047
9990
9997
Formula based on 4 cations and 6 oxygen atoms
1111
Si
187
176
189
188
197
199
165
181
179
167
198
198
200
199
195
200
197
200
Ti
004
008
002
004
001
001
011
005
005
011
001
003
001
004
003
001
002
001
Zr
000
000
000
000
001
001
000
000
000
000
000
001
000
000
000
001
001
000
Al
012
024
013
014
006
006
035
024
025
034
004
005
004
005
005
004
005
006
Cr
001
001
000
000
000
000
000
001
001
000
000
000
000
000
000
000
000
000
Fe3þ
006
015
018
011
040
082
018
011
013
015
044
083
050
068
083
084
056
080
Fe2þ
008
008
023
013
034
001
004
005
004
008
018
003
034
011
000
001
030
005
Mn
000
001
002
001
004
002
000
000
000
000
003
002
004
003
003
004
003
003
Mg
086
067
054
071
016
006
069
080
079
067
030
001
005
009
005
002
005
003
Ca
093
093
086
091
058
012
094
088
088
092
055
011
046
024
023
011
040
014
Na
003
007
013
007
043
090
004
005
006
005
046
093
056
078
083
092
061
088
K
000
000
000
000
000
000
000
000
000
000
000
000
000
000
000
000
000
000
Sum
400
400
400
400
400
400
400
400
400
400
400
400
400
400
400
400
400
400
mol % end-members
Aeg
3
7
13
7
41
83
4
5
6
5
45
87
51
68
88
88
57
Jd
0
0
0
0
3
5
0
0
0
0
2
3
4
4
0
3
2
6
Ti-Aeg
0
0
0
0
0
4
0
0
0
0
0
7
3
7
0
4
3
3
Fe-Ts
2
4
3
2
0
0
7
3
4
5
0
0
0
0
0
0
0
0
Ti-Ts
4
8
2
4
1
0
11
5
5
11
1
1
0
1
3
0
2
0
Al-Ts
1
4
4
2
0
0
6
7
7
6
0
0
0
0
0
0
0
0
Di
81
78
69
54
70
16
5
67
71
69
63
30
0
5
8
5
1
5
3
Hed
8
8
25
14
39
3
4
4
4
8
22
2
37
12
4
4
31
7
En
4
0
0
1
0
0
1
5
5
2
0
0
0
0
0
0
0
0
Fs
0
0
0
0
0
0
0
0
0
0
0
0
0
0
0
0
0
0
TAMAZEGHT COMPLEX, MOROCCO
K2O
MARKS et al.
TMZ68
JOURNAL OF PETROLOGY
VOLUME 49
NUMBER 6
TiO2
(rutile)
TiO2
(rutile)
FeTiO3
(ilmenite)
Fe2TiO4
(ulvöspinel)
FeO
pyroxenites
magnetite
replacement hematite
FeTiO3
(ilmenite)
glimmerites
spl (ocelli)
spl (matrix)
Fe3O4
(magnetite)
Fe2TiO4
(ulvöspinel)
Fe2O3
(hematite)
FeO
Fe2TiO4
(ulvöspinel)
FeO
foid-monzo syenites
magnetite
Fe2O3
(hematite)
TiO2
(rutile)
ol-shonkinites
magnetite
FeTiO3
(ilmenite)
amph-shonkinites
magnetite
ilmenite
Fe3O4
(magnetite)
monzogabbros & monzosyenites
magnetite
ilmenite
Fe3O4
(magnetite)
TiO2
(rutile)
FeTiO3
(ilmenite)
JUNE 2008
nepheline syenites & malignites
magnetite
Fe2TiO4
(ulvöspinel)
Fe2O3
(hematite)
FeO
Fe3O4
(magnetite)
Fe2O3
(hematite)
Fig. 10. Composition of Fe^Ti oxides in the various Tamazeght lithologies.
of crystallographic and spectroscopic data, we do not
attempt to constrain the distribution of Ti, Fe and Al
between the cation sites.
In the ultramafic rocks, TiO2 contents range between
399 and 1429 wt %. Si shows significant deviation from
the ideal stoichiometry (296^249 p.f.u.) and Fe3þ varies
considerably (112^157 p.f.u.). In olivine-shonkinites,
TiO2 (196 and 1884 wt %), Fe3þ (099^166 p.f.u.) and
Si (212^30 p.f.u.) display strong variance. In nepheline
syenites, TiO2 and MgO contents are comparatively lower
(313^438 wt % and 022^028 wt %, respectively),
whereas Al2O3 (314^412 wt %), MnO (089^154 wt %)
and ZrO2 (up to 12 wt %) are significantly higher than
in other lithologies (Table 4). This may reflect the more
evolved character of these rocks.
Many workers advocate a simple homovalent substitution of Ti , Si, based on the negative correlation between
these two elements, to account for the apparent deficit on
the Z-site. However, for the Tamazeght garnets the strong
negative correlation betweenTi and Si does not exactly fall
on the ideal 1:1 correlation and even when Si is ideal
(at 3 a.p.f.u.) 02 Ti p.f.u. is present (Fig. 11). This indicates that Ti and/or Si are involved in other substitutions.
In an attempt to further identify important substitutions
occurring within the Tamazeght garnets, we employed
principal component analysis (PCA), a statistical method
that has proven useful in petrological studies (Jime¤nezMilla¤n et al., 1994; Ragland et al., 1997) to our data, using
XLSTAT 2007.6 (Addinsoft). The method extracts a set of
principal components, which allows us to explain the
observed variability in compositions. We deduce that the
most substitution schemes are Ti4þFe3þFe3þ1Si1,
Ti4þMgFe3þ2 and/or Ti4þFe2þFe3þ2, which are responsible for about 80% of the observed variability of the data.
Amphibole
The majority of amphibole analyses show a good 1:1
correlation between Ti and (Mg, Fe, Mn) (Fig. 12) indicating that the substitution Ti4þO22Mg1(OH)2
1112
Table 3: Representative electron microprobe analyses of Fe^Ti oxides from theTamazeght Complex. Morocco
Rock-type: Pyroxenite
Sample:
TMZ25
Glimmerite
TMZ23b
TMZ25
TMZ23b TMZ20
Monzogabbro
TMZ23
TMZ23
TMZ20
TMZ20 TMZ320
magnetite magnetite hematite hematite magnetite magnetite magnetite magnetite spinel
Monzonite
TMZ159
TMZ320 TMZ2
Foid-monzosyenite
TMZ321
TMZ2
TMZ219
TMZ157
TMZ312
magnetite magnetite ilmenite magnetite magnetite ilmenite magnetite magnetite magnetite
Nb2O5
b.d.
b.d.
b.d.
b.d.
b.d.
b.d.
b.d.
b.d.
b.d.
b.d.
b.d.
b.d.
b.d.
b.d.
b.d.
b.d.
b.d.
SiO2
004
003
009
008
000
002
000
000
010
003
002
001
001
005
002
001
001
000
TiO2
004
057
036
028
291
339
169
265
190
069
022
4615
034
017
4522
663
484
1069
Al2O3
015
004
003
008
003
000
000
003
1484
016
009
003
018
026
003
058
048
039
V2O3
011
035
022
015
042
042
044
037
007
030
032
262
031
037
255
071
056
093
Cr2O3
006
004
004
007
048
014
215
857
3813
007
005
004
007
007
005
107
115
104
Fe2O3
6903
6775
9854
9928
6316
6269
6341
5545
1123
6703
6807
1148
6738
6681
1030
5367
5681
4698
FeO
3141
3183
000
022
3158
3112
2993
2880
2324
3098
3089
3438
3065
3066
3598
3544
3426
3849
MnO
000
002
043
011
174
246
209
379
537
069
037
283
052
012
436
095
053
140
MgO
002
002
000
000
036
050
038
056
325
005
007
239
005
009
010
054
035
096
ZnO
000
000
000
002
025
024
020
047
315
005
001
003
008
002
015
016
015
017
10086
10065
9971
10029
10093
10098
10029
10069
10128
10005
10011
9996
9959
9862
9876
9976
9914
10105
1113
Nb
000
000
000
000
000
000
000
000
000
000
000
000
000
000
000
000
000
000
Si
000
000
000
000
000
000
000
000
000
000
000
000
000
000
000
000
000
000
Ti
000
002
001
001
008
010
005
008
005
002
001
085
001
000
087
019
014
030
Al
001
000
000
000
000
000
000
000
059
001
000
000
001
001
000
003
002
002
V
000
002
001
000
002
002
002
002
000
001
001
008
001
002
008
003
003
004
Cr
000
000
000
000
001
000
006
026
102
000
000
000
000
000
000
003
003
003
Fe3þ
199
195
197
199
180
178
182
157
029
195
198
021
196
196
020
153
164
131
Fe2þ
100
101
000
000
100
098
095
091
066
099
099
071
099
100
076
113
110
121
Mn
000
000
001
000
006
008
007
012
015
002
001
006
002
000
009
003
002
004
Mg
000
000
000
000
002
003
002
003
016
000
000
009
000
001
000
003
002
005
Zn
000
000
000
000
001
001
001
001
008
000
000
000
000
000
000
000
000
000
Sum
300
300
200
200
300
300
300
300
300
300
300
200
300
300
200
300
300
300
mol % end-members for magnetite
Mag
100
98
91
90
92
79
14
98
99
98
99
78
83
68
Usp
0
2
8
10
5
8
5
2
1
1
0
19
14
30
Sp
0
0
1
0
3
13
81
0
0
1
1
3
3
2
mol % end-members for ilmenite
Ilm
Hem
Pyr
0
0
81
99
100
12
80
10
1
0
7
10
(continued)
TAMAZEGHT COMPLEX, MOROCCO
Formula based on 3 (2) cations and 4 (3) oxygen atoms for magnetite (ilmenite)
MARKS et al.
Total
b.d.
Table 3: Continued
Rock-type:
Sample:
Olivine-shonkinite
Amphibole-shonkinite
Nepheline syenite
Miaskitic malignite
Vein
TMZ12
TMZ130
TMZ130
TMZ68
TMZ139
TMZ68
TMZ139
TMZ165
TMZ94
TMZ311
TMZ82
TMZ82
TMZ288
TMZ234
TMZ234
TMZ229
TMZ229
magnetite
magnetite
magnetite
magnetite
magnetite
ilmenite
ilmenite
magnetite
magnetite
magnetite
magnetite
magnetite
magnetite
ilmenite
ilmenite
magnetite
magnetite
b.d.
b.d.
b.d.
b.d.
b.d.
b.d.
b.d.
b.d.
b.d.
b.d.
b.d.
b.d.
b.d.
124
163
005
000
SiO2
007
007
002
002
003
003
000
007
003
001
003
001
003
002
000
000
007
TiO2
1268
255
740
377
025
4717
4684
028
345
803
588
313
036
4982
4970
062
463
Al2O3
324
118
297
026
014
002
001
010
029
026
015
021
017
000
000
005
017
V2O3
022
014
022
034
042
267
262
022
045
074
016
021
024
213
215
017
015
Cr2O3
032
041
028
007
059
006
005
008
006
016
003
004
004
003
005
007
004
Fe2O3
4158
6334
5095
6090
6628
850
970
6766
6215
5262
5754
6228
6799
142
164
6757
6004
FeO
3973
3239
3513
3441
3061
3792
3748
3085
3201
3653
3302
3192
3094
2167
2065
3142
3296
MnO
107
077
142
009
032
200
297
047
217
102
337
181
038
2281
2368
027
245
MgO
190
064
108
009
002
137
092
002
011
067
001
002
003
002
002
001
008
ZnO
009
006
018
002
005
007
005
003
045
018
036
021
032
000
000
000
019
10090
10155
9966
9998
9871
9980
10064
9979
10116
10024
10056
9983
10050
9792
9788
10019
10077
Total
1114
000
000
000
000
000
000
000
000
000
000
000
000
004
004
000
000
000
000
000
000
000
000
000
000
000
000
000
000
000
000
000
000
000
Ti
036
007
021
011
001
088
087
001
010
023
017
009
001
094
095
002
013
Al
014
005
013
001
001
000
000
000
001
001
001
001
001
000
000
000
001
V
001
001
001
002
002
008
008
001
002
003
001
001
001
006
006
001
001
Cr
001
001
001
000
002
000
000
000
000
000
000
000
000
000
000
000
000
Fe3þ
114
178
143
175
193
016
018
196
177
149
164
180
196
003
003
195
171
Fe2þ
121
101
110
110
100
079
078
099
101
116
105
102
099
045
043
101
105
Mn
003
002
005
000
001
004
006
002
007
003
011
006
001
048
049
001
008
Mg
010
004
006
001
000
005
003
000
001
004
000
000
000
000
000
000
000
Zn
000
000
000
000
000
000
000
000
001
001
001
001
001
000
000
000
001
Sum
300
300
300
300
300
200
200
300
300
300
300
300
300
200
200
300
300
mol % end-members for magnetite
Mag
57
90
72
88
98
99
89
76
83
90
99
98
87
Usp
36
7
21
11
1
1
10
23
17
9
1
2
13
Sp
7
3
7
1
1
0
1
1
0
1
0
0
0
mol % end-members for ilmenite
Ilm
87
84
48
Hem
9
9
1
46
2
Pyr
4
7
51
52
JUNE 2008
000
Si
NUMBER 6
Nb
VOLUME 49
Formula based on 3 (2) cations and 4 (3) oxygen atoms for magnetite (ilmenite)
JOURNAL OF PETROLOGY
Nb2O5
MARKS et al.
TAMAZEGHT COMPLEX, MOROCCO
Table 4: Representative electron microprobe analyses of garnet from theTamazeght Complex, Morocco
Rock-type:
Glimmerite
Sample:
Pyroxenite
TMZ20
TMZ22
SiO2
3390
3329
TiO2
691
777
ZrO2
007
Al2O3
Fe2O3
TMZ22
Olivine-shonkinite
TMZ25
TMZ12
TMZ12
TMZ12
Granular
Foyaitic nepheline
neph. syenite
syenite
TMZ126
TMZ221
TMZ23c
TMZ23b
TMZ221
3277
3413
3323
3222
2476
3360
3556
3462
3437
3390
795
639
670
1036
1884
1028
300
345
420
418
008
010
020
021
004
100
031
005
082
109
110
128
106
112
105
101
187
165
079
098
324
350
358
2135
2112
2207
2183
2314
1817
1739
1665
2536
2352
2168
2198
FeO
228
272
169
229
199
230
266
430
090
155
177
159
MnO
042
057
063
057
060
046
035
030
015
089
154
152
MgO
066
067
066
061
062
127
146
033
025
027
027
026
CaO
3317
3262
3279
3291
3245
3322
3187
3304
3353
3235
3170
3144
Na2O
009
012
019
012
013
006
015
045
017
013
023
022
Total
10013
10002
9997
10009
10007
9997
10013
10005
9996
10084
10036
9979
Formula based on 8 cations and 12 oxygen atoms
Si
284
281
277
286
279
269
212
281
298
287
287
284
Ti
043
049
050
040
042
065
121
065
019
022
026
026
Zr
000
000
000
001
001
000
004
001
000
003
004
005
Al
013
010
011
010
010
018
017
008
010
032
034
035
Fe3þ
134
133
140
138
147
114
112
105
160
147
136
139
Fe2þ
016
019
012
016
014
016
019
030
006
011
012
011
Mn
003
004
004
004
004
003
003
002
001
006
011
011
Mg
008
008
008
008
008
016
019
004
003
003
003
003
Ca
297
294
295
295
293
297
292
296
301
287
283
282
Na
002
002
003
002
002
001
002
007
003
002
004
004
Sum
800
800
800
800
800
800
800
800
800
800
800
800
1.5
1
ide
Ti p.f.u.
al 1
0.5
:1 c
orr
ela
tion
glimmerites
pyroxenites
olivine-shonkinites
nepheline syenites
0
2
2.5
Si p.f.u.
3
Fig. 11. Correlation of Ti vs Si for garnet from the various Tamazeght
rocks. (See text for further discussion.)
(e.g. Oberti et al.,1992) plays an important role in the incorporation of Ti in the Tamazeght amphiboles; consequently,
the amphibole formula unit was calculated on the basis of
(23 þ Ti) oxygen atoms and 16 cations. Table 5 gives some
representative amphibole analyses. Following the nomenclature scheme of Leake et al. (1997), all analyses are calcic
amphiboles of hastingsitic (AlVI Fe3þ) and kaersutitic
(Ti 05 p.f.u.) composition. The latter is restricted to monzogabbros, some monzonites and to shonkinites (the least
evolved rock types of the respective lithological groups).
Figure 13 illustrates the variation in XK [K/(Na þ K)], AlVI
p.f.u., F p.f.u., XMg, XFe3þ [Fe3þ/(Fe2þ þ Fe3þ)], andTi p.f.u.
observed throughout the complex. The last three variables
show a systematic evolution within the two lithological rock
groups, each parameter decreasing with progressive evolution. The variation in octahedrally coordinated aluminium
(AlVI) and XK appears to be unsystematic. However, within
the monzonitic group the absolute range of XK seems to
decrease from monzogabbros via monzonites towards foidmonzosyenites, whereas the maximum XK value in question
increases slightly. In terms of halogens, chlorine content is
always low (5004 p.f.u.) and fluorine contents are variable,
with foid-monzosyenites and foyaitic nepheline syenites
showing comparatively high F contents of 5053 and
5072 p.f.u., respectively.
1115
JOURNAL OF PETROLOGY
monzogabbros
monzonites
foid-monzosyenites
3.8
(Mg, Fe2+,Mn) p.f.u.
VOLUME 49
shonkinites
porphyritic nepheline syenites
granular nepheline syenites
foyaitic nepheline syenites
3.6
3.4
3.2
3
1:1
2.8
0
0.2
0.4
0.6
Ti p.f.u.
0.8
1
Fig. 12. Correlation of (Mg þ Fe2þ þ Mn) p.f.u. vs Ti p.f.u. in amphibole from various Tamazeght rocks. The good 1:1 correlation indicates
that the substitution Ti4þO22Mg1(OH)2 (e.g. Oberti et al., 1992)
plays an important role in the incorporation of Ti in the Tamazeght
amphiboles.
Biotite
Biotite occurs in all samples of the ultramafic and the
monzonitic groups. Within the foid syenitic group, only
shonkinites and foyaitic nepheline syenites contain biotite.
Similar to the amphiboles, the Tamazeght biotites are characterized by low Si contents and 8 ^ Si þ Al deficits of up to
034 p.f.u., which indicates the presence of tetrahedrally
coordinated Fe3þ or Ti4þ (e.g. Dunworth & Wilson, 1998;
Mann et al., 2006). Figure 14 illustrates the variation in
XMg, Ti and F p.f.u. observed throughout the complex
and Table 6 gives some representative analyses.
XMg values are highest in biotites from the ultramafic
rocks (up to 096) and decrease towards the more evolved
rock types, reaching their lowest values (503) in some of
the foyaitic nepheline syenites. In olivine-shonkinites, however, two types of biotite occur, groundmass biotite and
biotite growing at the expense of olivine, with the latter
having exceptionally high XMg values (around 09), reflecting the XMg value of the precursor olivine.
Ti contents show considerable variation, being lowest in
ultramafic rocks (5035 Ti p.f.u.), in biotite replacing pyroxene in foyaitic nepheline syenites (Fig. 6b) and in the
biotite from the olivine-shonkinites that overgrows olivine
(5023 Ti p.f.u.). All other biotites have elevated Ti contents, with the highest Ti contents found in monzogabbros
(up to 116 Ti p.f.u.) and shonkinites (up to 078 Ti p.f.u.).
It should be noted that these two rock types also contain
the most Ti-rich amphiboles. Such high Ti contents could
potentially explain the 8 ^ (Si þ Al) deficits on the tetrahedral site. The positive Ti^Al correlation and the negative
Ti^Si and Ti^(Mg,Fe,Mn) correlations (Fig. 15) imply the
importance of the coupled substitution MgSi2Ti4þ1Al2,
which was proposed by Wagner et al. (1987) and
NUMBER 6
JUNE 2008
Mann et al. (2006) for biotite from alkaline rocks of
the Katzenbuckel volcano, Germany. While replacing divalent cations by Ti4þ on octahedral sites, charge-balance
might also be reached by the substitution mechanism
MgK2Ti4þ1Al2, which creates vacancies on the X site
(Deer et al., 1992). According to the applied formula calculation (normalization to 22 oxygens), up to 15% of the
X site may be vacant. In Fig. 15a and b, low-Ti biotites
from glimmerites and from foyaitic nepheline syenite
TMZ223 deviate from the trend shown by biotites from
all other samples. In Fig. 15c, only the latter plot off the
trend. This feature also coincides with elevated F contents
in these samples and may indicate that their chemistry is
governed by other substitution mechanisms, potentially
implying a different origin for these micas.
In most biotites, chlorine contents are5003 p.f.u, except
for monzogabbros and monzonites, where slightly higher
Cl contents of up to 007 p.f.u were found. Fluorine contents are highly variable (from 5001 to 41p.f.u.; Fig. 14).
Generally, F is negatively correlated with Ti content and
reaches high values in ultramafic rocks, in biotite around
olivine from shonkinites, and in biotite in evolved foyaitic
nepheline syenites. F contents in biotite from the monzonites do not fit this relationship, but this might be
explained by simple alteration of mica in these rocks
(see above and Fig. 4d).
Feldspar
Ca-bearing plagioclase is restricted to rocks of the monzonitic group where individual grains are strongly zoned
with decreasing mol % anorthite (An) and increasing
mol % albite (Ab) from core to rim; orthoclase (Or) is
generally low. Overall, plagioclase composition varies
between An68Ab31Or1 and An22Ab74Or4 (Fig. 16; Table 7).
The most anorthite-rich compositions are found in samples
of monzogabbro, whereas the most anorthite-rich plagioclase in monzonites and foid-monzosyenites is very similar
(An52 and An44, respectively).
Alkali feldspar is in most cases exsolved into pure albite
and orthoclase. These textures are partly rather coarse
and/or heterogeneous, and this feature makes it difficult to
reconstruct a primary magmatic composition. However,
in many samples (except for foyaitic nepheline syenites)
some alkali feldspar grains (or at least parts of them) show
no signs of exsolution. Unexsolved alkali feldspar in rocks
of the monzonitic group varies in composition between
Ab48Or48An4 and Ab20Or78An2 (Fig.16;Table 7), and exhibits no systematic evolution from monzogabbros to monzonites to foid-monzosyenites. In porphyritic and granular
nepheline syenites, alkali feldpar composition varies
between Ab70Or26An4 and Ab20Or80An0, and in malignites,
as well as in hydrothermal veins, An-free and relatively
Or-rich (Ab29Or71^Ab12Or88) alkali feldspar is found.
Interstitial albite in foyaitic nepheline syenites as well as
1116
MARKS et al.
TAMAZEGHT COMPLEX, MOROCCO
Table 5: Representative electron microprobe analyses of amphibole from theTamazeght Complex, Morocco:
Rock-type: Monzogabbro
Monzonite
Foid-monzosyenite
Olivine-
Amph-shonkinite
shonkinite
Porph. neph. Granular
Foyaitic neph.
syenite
neph. syenite
syenite
Sample:
TMZ159 TMZ159 TMZ321 TMZ321 TMZ157 TMZ219 TMZ313 TMZ130
TMZ68
TMZ139
TMZ311
TMZ74 TMZ95 TMZ221
SiO2
3940
3993
3878
3762
3906
4036
3948
3956
3979
3836
3982
3901
3971
4112
TiO2
594
737
418
152
243
304
330
680
348
325
267
343
174
116
ZrO2
004
006
003
014
013
006
008
009
006
005
015
002
011
020
Al2O3
1157
1203
1234
1232
1179
1162
1142
1110
1162
1207
1174
1188
1132
857
FeO
1291
1047
1348
2052
2103
1655
1736
1226
1665
1812
2033
1547
1756
2157
MnO
021
010
036
095
094
081
065
029
041
037
059
061
072
173
MgO
1077
1201
1081
685
733
1032
961
1127
999
830
813
991
945
790
CaO
1145
1187
1141
1047
1110
1113
1173
1154
1127
1084
1096
1088
1066
886
Na2O
275
277
252
273
267
295
274
288
276
250
302
285
283
427
K2O
184
158
167
185
186
188
193
16
193
209
178
185
208
180
Cl
016
005
007
008
003
004
004
009
007
011
004
007
004
001
F
011
015
012
000
059
108
022
018
014
012
000
030
039
145
9715
9839
9577
9505
9896
9984
9856
9766
9817
9618
9923
9627
9661
9864
Total
Formula based on 16 cations and (23 þ Ti) oxygen atoms
Si
603
600
600
600
603
607
600
602
605
602
605
604
615
635
Ti
068
083
049
018
028
034
038
078
040
038
031
040
020
013
Zr
000
000
000
001
001
000
001
001
000
000
001
000
001
001
Al
209
213
225
231
215
206
205
199
208
223
210
217
207
156
Fe3þ
102
098
083
089
094
100
111
111
099
091
100
097
088
133
Fe2þ
064
034
091
184
178
108
109
045
113
147
158
103
139
145
Mn
003
001
005
013
012
010
008
004
005
005
008
008
009
023
Mg
246
269
249
163
169
232
218
256
227
194
184
229
218
182
Ca
188
191
189
179
184
179
191
188
184
182
179
180
177
147
Na
082
081
076
084
080
086
081
085
081
076
089
086
085
128
K
036
030
033
038
037
036
037
031
037
042
035
037
041
035
Cl
004
001
002
002
001
001
001
002
002
003
001
002
001
000
F
005
007
006
000
029
051
011
009
007
006
000
014
019
071
1600
1600
1600
1600
1600
1600
1600
1600
1600
1600
1600
1600
1600
1600
Sum
late-stage albite laths in some of the malignites are An-free
and contain generally52 mol% orthoclase.
Foid minerals
Nepheline
The variation of nepheline composition is illustrated in
Fig. 17; representative nepheline compositions are given
in Table 8. In pyroxenites, nepheline composition varies
between Ne60Ks25Qtz15 and Ne72Ks25Qtz4. The relatively
Qtz-rich and Ne-poor compositions are typically found
in the cores of euhedral nepheline grains, which occur
as inclusions in garnet, whereas the Qtz-poor and
Ne-rich compositions are typical of interstitial nepheline
grains.
In rocks of the monzonitic and nepheline syenitic group,
nepheline varies in composition between Ne67Ks12Qtz21
and Ne72Ks18Qtz10 with no systematic differences between
the various rock types. The evolution from Qtz-rich and
Ks-poor to relatively Qtz-poor and Ks-rich compositions
is in contrast to the compositional evolution of nepheline
from the pyroxenites and has been described as being typical of post-magmatic re-equilibration (Powell, 1978).
A similar compositional variation is observed within
miaskitic and agpaitic malignites (Ne70Ks12Qtz18^Ne75
Ks22Qtz3), with a tendency for Ks-rich compositions to be
more frequent in agpaitic malignites. Nepheline compositions in a hydrothermal vein overlap with the Qtz-poor
compositions of the malignites.
1117
JOURNAL OF PETROLOGY
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NUMBER 6
JUNE 2008
monzogabbros
monzonites
foid-monzosyenites
shonkinites
porphyritic nepheline syenites
granular nepheline syenites
foyaitic nepheline syenites
0.1
0.2
0.3
0.4
0
0.5
0.1
K / (Na+K)
0.2
0.3
AlVI
p.f.u.
0.4
0.6
0.4
0.5
0
0.2
0.4
0.6
0.8
1
0.8
1
F p.f.u.
monzogabbros
monzonites
foid-monzosyenites
shonkinites
porphyritic nepheline syenites
granular nepheline syenites
foyaitic nepheline syenites
0
0.2
0.4
0.6
0.8
1
0
0.2
Mg / (Fe2++Mg)
0.8
1
0
Fe3+ / (Fe2++Fe3+)
0.2
0.4
0.6
Ti p.f.u.
Fig. 13. Diagram illustrating the compositional variation in amphibole from various Tamazeght rocks.
glimmerites
pyroxenite
monzogabbros
monzonites
foid-monzosyenites
olivine-shonkinites
amphibole-shonkinites
foyaitic nepheline syenites
0.2
around
olivine
0.4 0.6 0.8
Mg / (Fe2+ + Mg)
around
olivine
around
olivine
1
0
0.5
1
F p.f.u.
1.5 0 0.2 0.4 0.6 0.8 1 1.2
Ti p.f.u.
Fig. 14. Diagram illustrating the compositional variation in biotite from various Tamazeght rocks.
CaO and Fe2O3 contents may be up to 15 wt % and
16 wt %, respectively. The highest Fe contents are present
in pyroxenites and malignites, and the lowest contents were
observed in monzonites and nepheline syenites.
observed. Minor elements include Fe (505 wt % Fe2O3),
K (501wt % K2O) and Ca (504 wt % CaO).
Sodalite
DISCUSSION
Evidence from the mineral chemical
variations for a heterogeneous
magma source
Sodalite-group minerals occur as euhedral and interstitial
phases in most samples. Sodalite is not found in either
the ultramafic or shonkinitic lithologies. Compositional
differences between primary sodalite and sodalite associated with cancrinite in reaction textures are not obvious.
The chemical composition of both types is close to endmember sodalite, with Cl between 155 and 189 a.p.f.u.
and SO3 ranging from 001 to 04 a.p.f.u. (Table 9).
A weak negative correlation between S and Cl is
Clinopyroxene occurs in all lithologies and is therefore
most suited to track the physico-chemical evolution of the
Tamazeght magmas. The evolution from diopside-rich
pyroxene compositions towards end-member aegirine is
typical of alkaline complexes worldwide. The major difference between various complexes is the amount of Fe2þ
1118
MARKS et al.
TAMAZEGHT COMPLEX, MOROCCO
Table 6: Representative electron microprobe analyses of biotite from theTamazeght Complex, Morocco
Rock-type: Glimmerite
Pyroxenite
Monzo- Monzonite Foid-monzosyenite
Olivine-shonkinite
Amphibole-shonkinite
gabbro
Foyaitic nepheline
syenite
Sample:
TMZ20 TMZ22 TMZ25 TMZ23c TMZ320 TMZ2
TMZ313
TMZ318
TMZ130 TMZ130 TMZ68
TMZ139
TMZ165
TMZ221
SiO2
4039
3862
3879
4125
3589
3657
3575
3587
3588
4141
3829
3484
3803
3816
TiO2
211
352
194
074
507
585
597
550
484
160
364
437
178
233
Al2O3
1292
1432
1183
1135
1385
1375
1421
1400
1401
1152
1204
1397
1341
1218
FeO
613
487
1617
977
1506
1493
1732
1757
1900
538
1706
2037
1632
1956
MnO
027
012
088
140
040
029
052
039
026
010
041
033
086
137
MgO
2423
2359
1689
2121
1437
1491
1286
1309
1167
2437
984
1090
1445
1205
CaO
001
001
005
001
003
003
003
001
001
038
668
001
001
002
Na2O
007
010
012
010
072
048
047
039
035
031
197
054
011
032
K2O
1105
1065
1015
1075
963
978
967
968
959
1034
881
807
914
977
Cl
000
000
000
000
026
020
003
004
007
005
006
011
001
001
F
152
112
122
223
041
197
040
018
041
185
008
022
117
250
9871
9691
9804
9881
9568
9875
9723
9672
9608
9731
9888
9433
9529
9827
Total
Formula based on 22 oxygen atoms
Si
572
551
579
599
544
544
537
541
549
592
572
545
578
585
Ti
022
038
022
008
058
065
067
062
056
017
041
051
020
027
Al
216
241
208
194
247
241
252
249
253
194
212
258
240
220
Fe
073
058
202
119
191
186
218
222
243
064
213
267
208
251
Mn
003
001
011
017
005
004
007
005
003
001
005
004
011
018
Mg
511
502
376
459
325
331
288
294
266
520
219
254
328
276
Ca
000
000
001
000
001
000
000
000
000
006
107
000
000
000
Na
002
003
003
003
021
014
014
011
010
009
057
016
003
009
K
200
194
193
199
186
186
185
186
187
189
168
174
177
191
Cl
000
000
000
000
007
005
001
001
002
001
001
003
000
000
F
068
051
058
102
020
093
019
008
020
084
004
011
056
121
1599
1589
1594
1597
1578
1570
1569
1571
1567
1592
1593
1569
1563
1578
Sum
enrichment relative to Na and Fe3þ enrichment during
their evolution (Fig. 18). In that sense, two extreme evolutionary paths have been documented: from diopside to
aegirine without significant Fe2þ enrichment [e.g. Murun,
Siberia (Mitchell & Vladykin, 1996) and Katzenbuckel,
SW Germany (Mann et al., 2006)] and from diopside
via hedenbergite, and thus strong Fe2þ enrichment, prior
to evolution towards aegirine-rich compositions [from the
Il|¤ maussaq Complex, South Greenland (Larsen, 1976;
Marks & Markl, 2001; Markl et al., 2001)]. In addition to
these two extremes, intermediate paths have been documented in many studies (e.g. Mitchell & Platt, 1982;
Korobeinikov & Laajoki, 1994; Coulson, 2003; Vuorinen
et al., 2005). The most obvious factor influencing the extent
of Fe2þ enrichment in clinopyroxene during differentiation
is the oxidation state of the magma (e.g. Larsen, 1976).
However, the presence of coexisting mafic minerals (e.g.
amphibole, biotite, garnet) has also been shown to play
a role (e.g. Chakhmouradian & Mitchell, 2002; Vuorinen
et al., 2005) and it seems likely that the Na/Ca ratio of the
melt or fluid from which the pyroxene crystallizes also
influences the evolutionary path. For the two extreme
trends, quantitative data on oxygen fugacitiy (fO2) are
available for the Katzenbuckel (SW Germany) and the
Il|¤ maussaq suite (South Greenland), and indeed, for these
two suites relatively oxidized (FMQ ¼ þ1 to þ2, where
FMQ is the fayalite^magnetite^quartz buffer) and extremely reduced crystallization conditions (FMQ ¼ 2 to
4), respectively, were determined (Marks & Markl, 2001;
Markl et al., 2001; Mann et al., 2006). For intermediate
suites (e.g. North Qo“roq, South Greenland; Alno«, Sweden;
Fig. 18), no quantitative estimates have been reported, but
it seems likely that, in terms of fO2, these suites formed
under conditions somewhere around the FMQ buffer.
Although in all Tamazeght units similar diopside-rich
compositions are observed, the amounts of Na and Fe3þ
1119
JOURNAL OF PETROLOGY
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NUMBER 6
(a) 1.5
JUNE 2008
An
glimmerites
replacing pyroxene in TMZ 223
biotite from all other samples
Ti p.f.u.
1
plagioclase in monzonitic rocks
alkali feldspar in monzonitic rocks
0.5
alkali feldspar in all other rocks
900°C
750°C
0
1.5
Ab
2
2.5
3
Al p.f.u.
(b) 1.5
Ti p.f.u.
1
0.5
0
5
5.5
6
Si p.f.u.
(c) 1.5
Ti p.f.u.
1
0.5
0
4
5
(Mg, Fe, Mn) p.f.u.
6
Fig. 15. Correlation diagrams illustrating the compositional variation in biotite from various Tamazeght rocks. A positive correlation
between Ti and Al (a) and negative correlations for Ti vs Si (b) and
Ti vs (Mg,Fe,Mn) (c) imply the importance of the coupled substitution MgSi2Ti4þ1Al2.
increase during Fe2þ enrichment (compare the slopes of the
clinopyroxene evolutionary path) and the lengths of the evolutionary paths differ between the various rock types (Fig. 8).
By far the steepest slope is observed in the evolution trend
Or
Fig. 16. Feldspar composition in the An^Ab^Or triangle plotted
in comparison with the temperature-dependent feldspar solvus of
Fuhrman & Lindsley (1988) at 1 kbar.
for early clinopyroxene in shonkinites whereas a comparatively flat path is tracked by clinopyroxene from the monzonitic group and from nepheline syenites. Qualitatively, these
differences might indicate differences inthe oxidation state of
the parental magma, with shonkinites probably crystallizing
under more oxidized conditions compared with both rocks
of the monzonitic group and nepheline syenites. This observation is in accordance with the composition of coexisting
Fe^Ti oxides in the respective rocks. Some of the shonkinites
contain Ti-poor magnetite, whereas nepheline syenites and
monzonites contain eitherTi-enriched magnetite or magnetite and ilmenite (Fig.10).The trend shown by clinopyroxene
from the ultramafic rocks is intermediate, although in these
rocksTi-free magnetite and hematite occur.
In shonkinites and in some malignites, intermediate
pyroxene compositions (aegirine^augite) were found,
which do not follow a well-defined path, but are remarkably variable in composition and plot along a broad band
within the central part of the Di^Hed^Aeg triangle
(Fig. 8). This is in contrast to the well-defined clinopyroxene trend observed in foyaitic nepheline syenites, which
similarly evolve via intermediate to aegirine-rich compositions but follow a tight path. Such tightly defined
evolutionary paths most closely resemble the chemical evolution of clinopyroxene during its primary crystallization
history, which is directly linked to the physico-chemical
evolution in the crystallizing melt. In contrast, the poorly
defined compositional fields of late clinopyroxene from
shonkinites and from miaskitic malignites (Fig. 8) and
their heterogeneous micro-textural appearance (Fig. 6f)
imply that these compositions reflect different extents of
diffusional re-equilibration with a fluid phase during subsolidus conditions.
Detailed clinopyroxene zoning profiles reveal that the
relative proportions of Al-Ts, Fe-Ts and Ti-Ts change
systematically during evolution and that this systematic
evolution is different in rocks of the monzonitic group
1120
MARKS et al.
TAMAZEGHT COMPLEX, MOROCCO
Table 7: Representative electron microprobe analyses of feldspar from theTamazeght Complex, Morocco
Rock-type:
Monzogabbro
Monzonite
Foid-monzo-
Granular nepheline syenite
Miaskitic malignite
Syenite
Sample:
TMZ320
TMZ320
TMZ320
TMZ2
TMZ2
SiO2
5107
Al2O3
2869
5407
6698
5463
2793
1925
2739
Fe2O3
038
040
026
BaO
007
012
TMZ94
6609
6357
6653
1839
2308
1987
029
018
026
044
038
020
048
053
000
089
000
CaO
1453
1041
070
861
045
Vein
malignite
TMZ159
SrO
Agpaitic
TMZ82
TMZ83
TMZ295
6543
6578
6539
6550
6802
1821
1873
1857
1818
1868
016
017
011
024
044
b.d.
071
019
028
b.d.
032
462
085
TMZ94
b.d.
008
b.d.
007
b.d.
b.d.
b.d.
b.d.
b.d.
TMZ229
016
b.d.
b.d.
005
b.d.
1132
Na2O
365
514
521
569
227
892
791
225
342
159
135
K2O
026
029
789
049
1327
078
455
1361
1247
1424
1466
022
Total
9913
9889
10073
9837
10085
10123
10090
9994
10086
10004
10018
9840
302
Formula based on 8 oxygen atoms
Si
237
248
299
252
300
279
295
300
299
300
301
Al
157
151
101
149
098
119
104
099
100
100
098
098
Fe3þ
001
001
001
001
001
001
001
001
001
000
002
001
Ba
000
000
001
001
000
000
001
000
000
000
000
000
Sr
001
001
000
002
000
000
001
000
000
000
000
000
Ca
071
051
003
043
002
022
004
000
000
000
000
000
Na
033
046
045
051
020
076
068
020
030
014
012
097
K
002
002
045
003
077
004
026
080
072
083
086
001
Sum
502
500
495
502
498
501
500
500
502
497
499
499
mol % end-members
Ab
31
46
48
53
20
74
70
20
30
15
12
Or
1
2
48
3
78
4
26
80
70
85
88
1
An
68
52
4
44
2
22
4
0
0
0
0
0
compared with that in the nepheline syenites. In monzonitic rocks, all three Tschermak components increase from
core to the rim of crystals, whereas the opposite is the case
in nepheline syenites. The schematic equilibrium
CaAl½AlSiO6 ðin clinopyroxeneÞ þ SiO2
, Ca½Al2 Si2 O8 ðin plagioclaseÞ
ð1Þ
implies that Al-Ts in clinopyroxene and the anorthite component in plagioclase buffer silica activity in a crystallizing
melt. Consequently, the presence of both clinopyroxene
and plagioclase in monzonitic rocks indicates that silica
activity was higher in these rocks compared with foid syenites, which lack plagioclase (note that the amount of Al-Ts
component in clinopyroxene of monzonitic rocks is very
similar to that in most foid syenitic rocks). In fact, qtzbearing monzosyenites have been reported to occur rarely
in Tamazeght (Kchit, 1990).
The observed variations in Ti-Ts between the rock units
(Fig. 9) show that, generally, clinopyroxene in the ultramafic rocks is significantly lower in Ti-Ts than in all other
99
rock types (neglecting some of the most evolved nepheline
syenites). The most obvious difference between the ultramafic and the other rocks is the presence or absence of
Ti-bearing andradite. Although no simple schematic equilibrium between Ti-bearing andradite and the Ti-Ts molecule can be expressed, it seems likely that Ti-bearing
andradite acts as a sink for Ti, and thus coexisting clinopyroxene (Fig. 9) and biotite (Fig. 14) are comparatively
starved of Ti. The exceptionally high Ti contents in
early clinopyroxene-I phenocrysts in glimmerites do not
contradict this observation, as micro-textures show that
clinopyroxene-I is not in equilibrium with garnet and
mica (Fig. 3e); they may simply have crystallized before
garnet appeared on the liquidus.
The composition of clinopyroxene-I in glimmerites is
distinct from that in most other rocks from the complex.
In addition to being diopside-dominated (which does not,
however, make clinopyroxene unique for the Tamazeght
suite), clinopyroxene-I in glimmerites shows comparatively
high proportions of all three Tschermak components and is
extremely low in Na (Figs 8 and 9). We interpret these data
1121
JOURNAL OF PETROLOGY
SiO
2
VOLUME 49
80
pyroxenites
1068
°
775°
7
00°
90
500°
Ne
90
80
70
60
Ks
SiO
2
wt%
80
monzonitic group &
nepheline syenites
1068
°
775°
700°
90
500°
Ne
90
80
70
60
Ks
SiO
2
wt%
miaskitic malignites
agpaitic malignites
hydrothermal veins
80
1068
°
775°
700°
90
500°
Ne
90
80
70
60
Ks
wt%
Fig. 17. Nepheline compositions in the Ne^Ks^SiO2 triangle (on a
wt% basis) for the various Tamazeght rocks. The isotherms are from
Hamilton (1961); œ, the Morozewicz nepheline composition; i, the
Buerger nepheline composition.
as evidence that clinopyroxene-I from glimmerites crystallized from a melt source chemically distinct from the parental melt of the other rocks. Furthermore, given the high
abundance of biotite in these rocks, this parental magma
must have been exceptionally rich in potassium. Bouabdli
& Liotard (1992) reported major and trace element
data for the Tamazeght glimmerites and suggested that a
kimberlitic magma was a likely parent to these rocks.
However, Tamazeght glimmerites differ from typical
kimberlites in the lack of Mg-rich ilmenite (instead,
a Cr-bearing but relatively Mg-poor spinel phase is present), the atypical Na-poor and Ti-rich clinopyroxene
NUMBER 6
JUNE 2008
compositions (see above) and the occurrence of Ti-rich
and Cr-poor garnet. In any case, the presence of calcite
and large amounts of phlogopite implies a potassium-rich
and carbonated mantle source; such a source rock has
already been proposed for the lamprophyre dyke swarm
that cross-cuts the Tamazeght complex rocks (Bouabdli
et al., 1988). In all, it seems likely that a carbonated
amphibole-lherzolite was the source rock for the generation of the lamprophyric dykes, carbonatites and the
Tamazeght glimmerites.
In the remaining rock units, very similar diopside-rich
core compositions are observed. However, the various
rock types document dissimilar evolutionary paths resulting in different phase assemblages and different phase
compositionsça fact that is hard to reconcile with the
assumption of a homogeneous parental melt source for all
rock types. We thus argue that the various rock units in
the Tamazeght complex possibly resulted from successive
(or progressive) melting of a chemically and mineralogically heterogeneous mantle source. The generated melt
batches were very similar in their XMg value, but in terms
of their physico-chemical characteristics they were obviously distinct from each other. In turn, these differences
in intensive parameters (fO2, aSiO2, aH2O) resulted in the
stabilization of different phase assemblages (e.g. ilmenite
or magnetite, Ti-andradite or titanite, amphibole or pyroxene, presence or absence of plagioclase) and these influenced the continuing chemical evolution of the melts from
which they crystallized. The influence of plagioclase fractionation on the chemical evolution of the remaining melt
and the composition of later crystallizing phases can be
seen in the monzonitic rocks. Plagioclase crystallization
(K/Na ratio 1) increases the K/Na ratio of the melt.
Amphibole in these rocks shows increasing K/(Na þ K)
ratios from core to rim and the minimum K/(Na þ K)
ratio of amphibole in the various monzonitic members
increases with evolution from monzogabbros via monzonites to foid-monzosyenites. However, such a systematic
evolution is not seen in the plagioclase-free rocks (Fig. 13).
Olivines from two samples of olivine-shonkinite have
similar high XMg values of around 09 in their cores.
Together with their relatively high Ni contents, this indicates that olivine in these rocks crystallized from a nearprimary mantle melt. However, Ni, Ca and Mn contents
are very different in these two samples (see above; Fig. 7).
The concentration of such elements in olivine of a fixed
XMg value is relatively independent of parameters such as,
for example, oxygen fugacity, being mainly dependent on
the composition of the crystallizing melt (Snyder &
Carmichael, 1992). Thus, the observed heterogeneities in
olivine from this rock type show that melt source heterogeneities may occur on a relatively small scale and these may
later be documented not only in different lithologies but
also in slight chemical variations of phases within a single
1122
MARKS et al.
TAMAZEGHT COMPLEX, MOROCCO
Table 8: Representative electron microprobe analyses of nepheline from theTamazeght Complex, Morocco
Rock-type: Pyroxenite
Monzo-
Monzonite Granular neph. Foyaitic neph. Porph. neph Miascitic malignite
gabbro
syenite
syenite
syenite
TMZ95
TMZ165
TMZ311
Sample:
TMZ23 TMZ25 TMZ320 TMZ321
SiO2
4185
4266
4333
4479
4360
4411
Al2O3
3403
3491
3467
3345
3460
3369
Fe2O3
127
081
046
062
044
CaO
001
002
077
149
Na2O
1464
1232
1614
1575
Agpaitic malignite
Vein
TMZ240
TMZ299
TMZ234
TMZ231
TMZ83
4460
4306
4334
4302
4581
4541
3374
3524
3350
3251
3130
3230
054
040
082
080
101
122
092
082
074
140
002
002
000
002
002
1594
1580
1533
1244
1645
1620
1651
1675
K2O
767
845
616
464
591
562
434
852
648
633
506
465
Total
9947
9917
10153
10074
10131
10050
9981
10010
10059
9906
9992
10005
864
Formula based 32 oxygen atoms
Si
815
826
821
847
826
840
847
879
831
837
875
Al
781
797
774
745
772
756
755
705
757
746
704
724
Fe3þ
019
012
006
009
006
008
006
015
012
015
018
013
Ca
001
000
000
000
000
016
030
017
015
028
000
000
Na
565
568
565
552
562
593
577
585
583
565
616
611
K
192
198
194
190
209
149
112
143
137
105
106
158
2373
2401
2360
2343
2375
2362
2327
2342
2335
2367
2319
2370
Sum
mol % end-members
Ne
68
56
72
68
71
69
67
70
74
73
70
72
Ks
23
25
18
13
17
16
12
12
19
19
14
13
Qtz
9
19
10
19
12
14
21
18
7
8
16
15
rock type. However, other possibilities, such as mixing of
different magma batches having distinct trace element
compositions, cannot be excluded.
Quantitative constraints on the evolution
of intrinsic parameters
At a given depth of intrusion, the parameters mainly governing the evolution of theTamazeght magmas areT, fO2, aSiO2
and aH2O . The Al-in-hornblende barometer (e.g. Schmidt,
1992) commonly provides the only means of constraining
the emplacement depth of plutonic complexes, such as the
monzonitic group of the Tamazeght complex. However,
Anderson & Smith (1995) showed that this barometer is
significantly affected by T and fO2. Given this, the known
restrictions of the application (Schmidt, 1992), the unusual
Ti-rich composition of amphiboles and the strong zonation
of plagioclase in the monzonitic rocks, the results need to be
treated with extreme caution. However, a combination of the
Al-in-hornblende barometer and amphibole^plagioclase
thermometry (Blundy & Holland,1990; Holland & Blundy,
1994) yields pressure estimates between 01 and 23 kbar
(uncertainty of 06 kbar; Anderson & Smith, 1995) and
equilibration temperatures between 790 and 8608C (uncertainty of 408C; Holland & Blundy, 1994). These estimates
seem reasonable, as they are in accordance with estimated
conditions in upper crustal alkaline magma chambers
elsewhere (e.g. Larsen & Srensen, 1987; Potter et al., 2004).
The presence of numerous pegmatites, roof pendants and
contact-metamorphosed sediments (Salvi et al., 2000) indicates a shallow depth of intrusion. Consequently, we apply
a pressure of 1kbar in subsequent calculations.
In addition to amphibole^plagioclase thermometry,
further constraints on minimum liquidus temperatures
can be made by plotting the feldspar compositions on
the temperature-dependent feldspar solvus of Fuhrman &
Lindsley (1988) and by nepheline thermometry (after
Hamilton, 1961). The results of the latter represent minimum liquidus temperatures, as a result of the known lateto postmagmatic equilibration of nepheline resulting in
Si loss and hence lower estimates of temperature (Powell,
1978). Near-solidus temperatures for olivine-shonkinites
can be calculated with the QUILF program (Frost &
Lindsley, 1992; Lindsley & Frost, 1992; Andersen et al.,
1993) from the assemblage olivine^clinopyroxene based on
the Fe^Mg-exchange equilibrium between these two
phases. For these calculations, average olivine core compositions and the most Fe-rich pyroxene core compositions
were used, to minimize the possible effects of later diffusive
re-equilibration, during which pyroxene tends to become
enriched in Mg (Markl et al., 1998; Marks & Markl, 2001).
1123
JOURNAL OF PETROLOGY
VOLUME 49
NUMBER 6
JUNE 2008
Table 9: Representative electron microprobe analyses of sodalite from theTamazeght Complex, Morocco
Rock-type:
Foid-monzosyenite
Granular neph. syenite
Foyaitic neph.
Miaskitic malignite
Agpaitic malignite
TMZ221
TMZ231
TMZ238
TMZ295
TMZ299
syenite
Sample:
TMZ157
TMZ74
TMZ95
SiO2
3610
3661
3623
3846
3759
3797
3776
3730
Al2O3
3175
3254
3245
3187
3091
3175
3123
3102
Fe2O3
021
006
007
026
042
018
020
031
CaO
013
036
012
000
002
001
003
006
Na2O
2394
2379
2415
2426
2515
2545
2474
2546
K2O
005
007
008
007
007
004
005
001
Cl
670
684
683
608
656
640
588
577
SO3
019
005
006
097
028
067
170
203
Total
9906
10032
9999
10197
10100
10247
10159
10195
605
Formula based 21 oxygen atoms
Si
588
591
586
606
608
604
607
Al
616
612
619
592
589
595
591
593
Fe3þ
001
003
001
003
005
002
002
004
Ca
006
002
002
000
000
000
000
001
Na
741
759
757
741
789
784
770
801
K
001
001
002
001
001
001
001
000
1953
1968
1966
1944
1993
1986
1972
2004
Sum
Cl
186
186
187
162
180
172
160
159
SO3
001
002
001
011
003
008
021
025
Sum
187
188
188
174
183
180
181
183
In addition to reaction (1) above, various phase equilibria allow us to constrain the T^fO2^aSiO2 evolution of
the different rock types:
CaTiO3 þ SiO2 , CaTiSiO5
ð2Þ
ZrO2 þ SiO2 , ZrSiO4
2 Fe3 O4 þ 3 SiO2 , 3 Fe2 SiO4 þ O2
ð3Þ
ð4Þ
NaAlSiO4 þ 2 SiO2 , NaAlSi3 O8
3 Ca3 Fe2 ðSiO4 Þ3 þ 2 Fe3 O4 þ 9 SiO2
, 9 CaFeSi2 O6 þ 4 O2
ð5Þ
3 CaFeSi2 O6 þ 3 FeTiO3 þ O2
, 3 CaTiSiO5 þ 2 Fe3 O4 þ 3 SiO2 :
ð6Þ
ð7Þ
Phase diagrams were calculated using the GEOCALC
software of Berman et al. (1987) and Liebermann &
Petrakakis (1990) with the database of Berman (1988).
Thermodynamic data for titanite and perovskite were taken
from Robie & Hemingway (1995). End-member component
activities were calculated using the solution model of
Fuhrman & Lindsley (1988) for feldspar, the models of
Wood (1979) and Green et al. (2007) for clinopyroxene,
and a mixing-on-site model for nepheline. The activity of
andradite was calculated after Cosca et al. (1986), and for
Fe^Ti oxides either unit activities or, if necessary, the
solution models implemented in QUILF were used.
Titanite, perovskite, baddeleyite and zircon were treated
as pure phases. Unit activity of SiO2 was referred to the
standard state of the relevant pure SiO2 phase at P andT.
Estimation of equilibration temperatures
Two-feldspar thermometry using the temperaturedependent feldspar solvus of Fuhrman & Lindsley (1988)
was applied to the monzonitic rocks and resulted in
minimum liquidus temperatures between 750 and 9008C
(Fig. 16).
Applying nepheline thermometry, maximum temperatures for the pyroxenites reach 10008C. For monzonitic
rocks and nepheline syenites, slightly lower but still high
temperatures well above 8008C are indicated, as is the
case for malignites. Nepheline from one of the hydrothermal veins yields temperatures of about 400^5008C. It is
interesting to note that the evolution of nepheline compositions is different in pyroxenites compared with the
other rock types (Fig. 17). In pyroxenites, the Ne content
increases with decreasing SiO2 component, whereas in the
other rock types, Ne content decreases, which was
1124
TAMAZEGHT COMPLEX, MOROCCO
1000
T > 800°C
(nepheline thermometry)
700
∆FMQ = −2 to −4
D
600
0.01
F
0.05
0.1
aSiO2
G
3 O8
1000
T > 950°C
(ol-cpx thermometry)
olivine-shonkinites
8
aAb = 0.5 – 1
aNe = 0.25 – 0.45
600
0.01
monzonitic group
0.2 0.3 0.5 0.75 1
aAn = 0.27 – 0.50 aAn = 0.67 – 0.50
aTs = 0.06 – 0.08 aTs = 0.03
900
790 – 860°C
(hbl-plag thermometry)
800
CaAl[
8
NaAlS
i3 O
AlSiO
700
Variations of aSiO2
0.1
aSiO2
2
temperature (°C)
1000
0.05
6 ] + SiO
2
CaTiO
3 + SiO
CaTiS
2
iO
5
Ca[A
l2 Si
2 O8 ]
ZrO
2 + Si
O
ZrSi
2
O
4
NaAlS
iO + 2
4
SiO
considered to indicate sub-solidus re-equilibration by
Powell (1978). The nepheline trend observed in pyroxenites
may be interpreted to reflect the primary evolution trend
of nepheline in these rocks, evolving towards Ne-rich compositions during differentiation.
QUILF calculations for olivine-shonkinites resulted in
equilibrium temperatures between 950 and 9808C, which
represent near-solidus conditions.
600
0.01
aAb = 0.5 – 0.7
aNe = 0.3 – 0.45
0.05
0.1
aSiO2
0.2 0.3 0.5 0.75 1
1000
aAb = 0.35 – 0.6
aNe = 0.3 – 0.4
nepheline syenites
Zr
3 O8
lSi
O
2 +
600
NaA
CaTiO
700
lSiO
3 + SiO
2
CaTiS
iO
Zr
S
5
i
Si
O O2
800
4 +2S
iO
2
900
temperature (°C)
The presence or absence of perovskite provides an important constraint on silica activity [equilibrium (2)].
Perovskite occurs only in some of the ultramafic rocks and
the olivine-shonkinites, where it always exhibits rounded
grain boundaries and rims of titanite (Fig. 5c). Titanite
itself occurs (although rarely) as euhedral grains in equilibrium with andradite. In all other rock types, perovskite
is absent. This implies that silica activity in these rocks was
initially significantly lower than in the other rock types
and the preserved textures indicate an increase of aSiO2
during the evolution of these rocks. For high temperatures
above 8008C (as indicated by nepheline thermometry),
the transformation of perovskite to titanite takes place at
aSiO2 values of 01. An upper limit of aSiO2 is given by
the absence of alkali feldspar according to equilibrium (5),
which results in aSiO2 values between 05 and 075 (Fig. 19).
A very similar evolution is observed in olivine-shonkinites
(Fig. 19) and an upper limit of aSiO2 of about 08 is estimated for these rocks.
NaAlSiO
5
700
i3 O
800
4 + 2 SiO
2
900
O3 +
SiO
2
CaTi
SiO
temperature (°C)
Fig. 18. Clinopyroxene evolution trends from various alkaline suites.
A, Katzenbuckel, SW Germany (Mann et al., 2006); B, Murun, Russia
(Mitchell & Vladykin, 1996); C, Lovozero, Russia (Korobeinikov &
Laajoki, 1994); D, Alno«, Sweden (Vuorinen et al., 2005); E, Coldwell
nepheline syenites, Canada (Mitchell & Platt, 1982); F, North Qo“roq,
South Greenland (Coulson, 2003); G, Il|¤ maussaq, South Greenland
(Larsen, 1976; Marks & Markl, 2001; Markl et al., 2001). Quantitative
data on oxygen fugacitiy (given as FMQ units) available for suites
A and G imply that the chemical evolution of clinopyroxene might be
useful as a qualitative indicator of oxygen fugacity (see text for further
discussion).
NaAlS
Hed
CaTi
Di
0.2 0.3 0.5 0.75 1
NaA
E
4
C
aAb = 0.5 – 1
aNe = 0.25 – 0.30
2
800
NaAlS
i
B
A
900
NaAlSiO
4 + 2 SiO
temperature (°C)
∆FMQ = +1 to +2
ultramafic rocks
@ 1 kbar
5
Aeg
CaTi
O3 +
SiO
2
CaTi
SiO
MARKS et al.
500
≈
400
0.01
0.05
0.1
aSiO2
veins:
aAb = 0.8 – 0.9
aNe = 0.48 – 0.53
0.2 0.3 0.5 0.75 1
Fig. 19. T^aSiO2 diagrams (calculated for 1 kbar) illustrating the
evolution of these parameters in various Tamazeght rocks.
1125
−10
−15
andr
adite
+m
hede
@600˚C
aAndr = 0.4 – 0.7
aHed = 0.05 – 0.1
amag = 1
FMQ
0.05
0.2 0.3 0.5 0.75 1
perovskite
titanite
olivine-shonkinites
@900°C
−10
log fO2
0.1
aSiO2
nepheline
−25
0.01
−5
albite
gite
nber
−20
HM
ad
andr
etite
magn
−15
mag
ite +
hede
e
netit
gite
nber
te
fayali
−20
aAndr = 0.3 – 0.5
aHed = 0.05 – 0.1
amag = 0.5 – 1
aFa = 0.05 – 0.1
FMQ
−25
amphibole-shonkinites
& monzonitic rocks
@800°C
−10
0.2 0.3 0.5 0.75 1
HM
−15
te
agneti
e+m
FMQ
−5
0.01
0.05
nepheline syenites
& malignites
@800°C
−10
aHed = 0.05 – 0.2 aHed = 0.0 – 0.3
amag = 0.5–1
amag = 1
0.1
aSiO2
perovskite
−25
enite aIlm = 0.7 – 0.8 aIlm = 0.7 – 0.8
e + ilm
bergit
heden
0.2 0.3 0.5 0.75 1
zircon
−20
baddeleyite
titanit
titanite
log fO2
albite, anorthite
0.1
aSiO2
baddeleyite
zircon
nepheline, Ca-Al cpx
−5
0.05
perovskite
titanite
0.01
−15
titanit
−20
FMQ
−25
0.01
e+m
agneti
0.05
te
enite
e + ilm
bergit
heden
0.1
aSiO2
albite
nepheline
log fO2
HM
Significance of clinopyroxene^garnet^Fe^Ti oxide^titanite
textures: further constraints on fO2 and aSiO2 evolution
In the ultramafic rocks, oxygen fugacity is buffered by
reaction (6) and was initially 2^5 log units above the
FMQ buffer (Fig. 20). The replacement of magnetite by
hematite indicates that fO2 rose subsequently to values
around the hematite^magnetite buffer. The intersection of
reaction (5) with the hematite^magnetite buffer (indicated
by a grey dot in Fig. 20) is temperature-dependent and
@1000°C
@800°C
tite
agne
albite
log fO2
HM
nepheline
ultramafic rocks
@800°C
@ 800°C
−5
In monzonitic rocks, silica activity is constrained by
equilibria (1), (2) and (4). Decreasing activity of anorthite
(in plagioclase) and increasing activity of the Tschermak
component (in clinopyroxene) displaces reaction (1) to
lower values of aSiO2 (Fig. 19). Using the most An-rich plagioclase composition of the monzonitic group and core
compositions of clinopyroxene, the calculated initial aSiO2
ranges between 05 for foid-monzosyenites and 075 for
monzogabbros (at temperatures of 790^8608C as calculated above); similar aSiO2 values between 04 and 07 are
calculated based on equilibrium (5). Additionally, a lower
limit of aSiO2 of 025 for the early crystallization stage
is given by the occurrence of zircon. Higher Tschermak
components in the rims of clinopyroxene and decreasing
An contents in plagioclase imply that aSiO2 dropped significantly during differentiation, which is in contrast to the
evolutionary trend determined for the ultramafic rocks.
Combined with the presence of titanite, a lower limit of
aSiO2 of about 01 can be determined. The even lower
aSiO2 indicated by the most An-poor plagioclase compositions can be explained by the fact that these compositions
were no longer in equilibrium with clinopyroxene.
For nepheline syenites and malignites, equilibrium (5)
was used to constrain aSiO2 (Fig. 19). Calculated initial
aSiO2 ranges between 025 and 05 and was, therefore,
initially lower than in the monzonitic rocks (assuming
T ¼ 800^9008C as indicated by nepheline thermometry).
This is in accordance with the absence of plagioclase in
the nepheline syenites, but similar amounts of Tschermak’s
component in clinopyroxene.
Constraints on the crystallization conditions of the hydrothermal veins are given by equilibrium (5) and the presence
of zircon in some of these veins [equilibrium (3)]. At
temperatures around 5008C (as indicated by nepheline
thermometry), aSiO2 was around 01^02 (Fig. 19). In many
of the investigated rocks, late-stage to hydrothermal alteration features are documented. Nepheline and sodalite are
altered to cancrinite, whereas in malignites late-stage
pure albite is observed. Based on mineral textures, the formation of agpaitic rocks, which crystallize eudialyte-group
minerals, catapleite, lafivenite and other Na^Zr-silicates, is
also seen to occur at late-magmatic stages (Schilling et al.,
2007). However, a detailed account of the late-stage to hydrothermal processes observed in theTamazeght rocks is not the
subject of this study and will be discussed in detail elsewhere.
JUNE 2008
titanite @ 1000°C
NUMBER 6
@ 600°C
VOLUME 49
perovskite
JOURNAL OF PETROLOGY
aIlm = 0.7 – 1
aHed = 0.1 – 0.45
amag = 0.5 – 0.95
0.2 0.3 0.5 0.75 1
Fig. 20. fO2^aSiO2 diagrams (calculated for 1 kbar) illustrating the
evolution of these parameters in various Tamazeght rocks. The bold
lines labelled HM and FMQ represent the position of the hematite^
magnetite and fayalite^magnetite^quartz buffers at unit activities.
1126
MARKS et al.
TAMAZEGHT COMPLEX, MOROCCO
takes place at aSiO2 values of between 055 (at 6008C)
and 08 (at 10008C). Although no estimate on the temperature of this transformation is possible, it is implied
that during this oxidation process aSiO2 simultaneously
increased to higher values. The relative scarcity of magnetite and titanite in garnet-rich pyroxenites compared with
all other rock types may indicate that Ti-rich garnet influences the stability of these phases. However, a detailed
quantitative treatment of the relevant phase relations is
not possible, as thermodynamic data for Ti-rich garnet are
lacking.
In olivine-shonkinites oxygen fugacity is buffered by
reactions (4) and (6) and similarly oxidized conditions
(FMQ ¼ þ25 to þ4) to the ultramafic rocks are indicated (Fig. 20). When comparing pyroxenites with olivineshonkinites, it seems surprising that despite the similarly
oxidized crystallization conditions, their clinopyroxene
evolutionary paths are distinct from each other (Fig. 8). In
fact, clinopyroxene-I from garnet-poor olivine-shonkinite
has higher Fe3þ/(Fe2þ þ Fe3þ) ratios of 085^093 than clinopyroxene from garnet-rich pyroxenites (04^07). Additionally, Fe3þ/(Fe2þ þ Fe3þ) values in the latter increase
from the inner to the outer zone; Figs 3d and 8; Table 2).
The observed irregular zonation patterns in clinopyroxene
from pyroxenites (Fig. 3d) do not exclude the possibility
of redistribution of several cations as a result of secondary
re-equilibration and, therefore, the achievement of equilibrium cannot ultimately be assumed. In contrast, garnet in
both rock types has similar Fe3þ/(Fe2þ þ Fe3þ) ratios of
08^095, which is the same range as found for clinopyroxene from olivine-shonkinites.
In olivine-shonkinites, the Fe3þ/(Fe2þ þ Fe3þ) ratios for
both minerals are similarly high and very similar to the
Fe3þ/(Fe2þ þ Fe3þ) ratio in garnet from pyroxenites, and it
therefore seems likely that in olivine-shonkinites garnet
and clinopyroxene co-crystallized or at least reflect the
same evolutionary stage of the melt they crystallized from.
In the remaining rock types, titanite textures show
some interesting variation. In monzogabbros, monzonites
and amphibole-shonkinites, either most titanite occurs
as rims around ilmenite, or subhedral titanite contains
abundant inclusions of rounded relics of Fe^Ti
oxides (Fig. 4b and c). Despite the fact that these rocks
contain much more magnetite than ilmenite, these relics
are in almost all cases ilmenite and the above-mentioned
titanite rims almost exclusively occur around ilmenite.
Primary magnetite grains do not seem to be affected
by this reaction. In the more evolved rock types, which
do not contain primary ilmenite (foid-monzosyenites and
nepheline syenites), titanite is always euhedral and seems
to have co-precipitated with clinopyroxene and magnetite
(Fig. 4e). Additionally, it occurs much more commonly as
euhedral inclusions in clinopyroxene and amphibole in
these rocks. These textures imply that both oxygen fugacity
and silica activity in these rocks were buffered by the
schematic equilibrium (7).
Wones (1989), Xirouchakis et al. (2001a, 2001b) and
Ryabchikov & Kogarko (2006) demonstrated that the
stability of titanite is controlled by T, fO2, aSiO2 and the
composition of the coexisting oxides and Fe^Mg silicates.
If reaction (7) is calculated for rocks of the monzonitic
group, which contain both magnetite and ilmenite (monzogabbros and monzonites), comparatively less oxidized
conditions between 05 and 25 log units above the FMQ
buffer are indicated (Fig. 20). Foid-monzosyenites, however,
lack ilmenite, and the magnetite has a higher ulvo«spinel
content. Nevertheless, we calculated equilibrium (7) for
these rocks, using the full range of magnetite and ilmenite
compositions observed in the monzonitic group. In this
case, the range of estimated fO2 expands towards relatively
reduced conditions up to 1 log unit below the FMQ buffer.
Using a similar approach for nepheline syenites and malignites, and taking into account the observed compositional
variations of the phases (including pure ilmenite as a possible lower limit for oxygen fugacity), fO2 values around
and significantly below the FMQ buffer (FMQ ¼ þ1
to ^2) are estimated (Fig. 20).
S U M M A RY A N D C O N C L U S I O N S
Our work on the various lithologies of the Tamazeght
complex demonstrates that the combination of detailed
petrographic studies with careful interpretation of mineral
chemical variations not only reveals details of their petrological evolution, but also can be used to constrain the role
of compositionally distinct mantle domains in their origin.
If the relevant phase assemblages indicative for intensive
parameters are identified, quantification of the important
phase equilibria is generally straightforward, if reliable
thermodynamic data for the phases of interest are available. For all rocks, high temperatures between 750 and
10008C for initial crystallization conditions are demonstrated. However, in terms of aSiO2 and fO2, the principal
rock groups crystallized and evolved under markedly
different conditions.
The most oxidized conditions were determined for the
ultramafic rocks (FMQ up to þ5) and olivine-shonkinites
(FMQ ¼ þ25 to þ 4). Both groups evolved from low
initial aSiO2 values (possibly as low as 01) to higher values,
reaching nepheline saturation in the ultramafic rocks
(aSiO2 around 05) and alkali feldspar saturation in the
olivine-shonkinites (aSiO2 ¼ 05^08). In terms of their
crystallization conditions and phase assemblages, the
olivine-shonkinites share some similarities with pyroxenites,
although the modal abundances for the phases (garnet,
clinopyroxene, olivine, nepheline, feldspar) are very different. We conclude that these two lithologies might have a
similar parental magma and could be linked to each by
crystal^liquid differentiation processes.
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JOURNAL OF PETROLOGY
VOLUME 49
For amphibole-shonkinites and monzonitic rocks, intermediate fO2 conditions are calculated (FMQ ¼ þ25
to 1). Their evolution with respect to aSiO2 is in the opposite sense to that indicated for the ultramafic rocks and
olivine-shonkinites. The fractionation of plagioclase and
clinopyroxene resulted in a decrease in aSiO2 from around
075 in the early stages to about 01, still during magmatic
conditions.
For nepheline syenites and malignites, relatively low
aSiO2 values of between 025 and 05 were calculated.
Although the values for fO2 (FMQ ¼ 2) have to be
taken as rough estimates, they indicate rather reduced
conditions of formation. The formation of hydrothermal
veins occurred at temperatures around 5008C and low
aSiO2 values between 01 and 02.
This study shows that the conditions of crystallization in
alkaline plutonic rocks influence both the crystallizing
mineral assemblage and the detailed chemical evolution
of the phases during differentiation and cooling. Both
in terms of fO2 and aSiO2, we found very different crystallization conditions for the various lithologies. In a general
sense, high fO2 favours the crystallization of garnet.
At intermediate fO2 titanite and magnetite are the preferred phases, whereas relatively low fO2 will lead to an
enhanced stability of clinopyroxene and ilmenite. The evolution of aSiO2 during magmatic differentiation also shows
contrasting trends. In the most primitive lithologies, low
initial aSiO2 prevents the crystallization of alkali feldspar
and plagioclase. In these rocks, aSiO2 increases during
differentiation. In turn, in plagioclase- and alkali feldsparbearing rocks, aSiO2 is buffered by the co-crystallization
of Al-Tschermak-bearing clinopyroxene and nepheline,
respectively, and indicates decreasing aSiO2 with progressive differentiation for some of the lithologies.
From the perspective of the origin of the large lithological variation found in the Tamazeght complex, in contrast
to the findings of Kchit (1990) and Bouabdli et al. (1988) we
suspect that the various rock types probably originated
from distinct melt batches derived from a heterogeneous
mantle source (heterogeneity caused by earlier metasomatic enrichment processes) or were produced from a
stratified mantle source. However, crystal fractionation
and accumulation processes may also play a role for some
of the rocks. Models for mantle metasomatism include
cryptic and patent mantle metasomatism (e.g. Wilshire &
Shervais, 1975), and the vein-plus-wall-rock mantle model
of Foley (1992) and others. The main differences between
the various models are the proposed metasomatizing
agents (melt vs fluid phase), the spatial effects of this metasomatism (locally vs universal) and the resulting mineralogical and geochemical changes [formation of hydrous
phases vs enrichment in incompatible elements without
other obvious changes; see review by Wilshire (1987)].
Regardless of the process, a later melting event in such
NUMBER 6
JUNE 2008
a modified mantle source region will initially produce
highly alkaline melts, which are strongly enriched in
incompatible elements if the degree of melting is low
enough. The higher the degree of melting, the less alkaline
and more basalt-like the resultant melts will be. In this
sense, the various lithologies in the Tamazeght complex
might be interpreted either as representing variable
degrees of melting of a cryptically metasomatized mantle
domain or, if the vein and wall-rock model of Foley (1992)
is applied, as reflecting melts of hydrous vein material,
pristine wall-rocks and hybrid mixtures between them.
The data presented here will serve as a basis for further
geochemical and geochronological work, which is needed
to resolve the origin of the Tamazeght rocks in detail.
AC K N O W L E D G E M E N T S
We acknowledge the support of Francois Fontan, Pierre
Monchoux and Stefano Salvi (Toulouse, France), who
gave us interesting insights into their earlier work on
Tamazeght and encouraged us to investigate this intrusive
complex in more detail. Ali Bajja (Marrakesh, Morocco)
is thanked for his co-operation, which facilitated the
field trip in May 2006. We also thank the citizens of
the small Berber village of Anougal for their hospitality,
and Boujemaa Boudaoud (Azrou, Morocco) for being our
guide in the High Atlas Mountains and for supplying
important infrastructure in the field. Florian Neukirchen
is thanked for his assistance in the field. We also thank
Sebastian Staude for his help with reflected light microscopy and Sylvia Mettasch for her interest in this work
and for careful and detailed petrographic work on some
of the samples. Funding for this work by the Deutsche
Forschungsgemeinschaft (grant Ma 2135/11-1) and the
Natural Sciences and Engineering Research Council
of Canada (IMC: Discovery grant funds) is gratefully
acknowledged. The constructive comments of M. Wilson,
R. Mitchell, A. Chakhmouradian and one anonymous
reviewer are greatly appreciated.
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A P P E N D I X 1: C A L C U L AT I O N O F C L I N O P Y ROX E N E E N D - M E M B E R S
I N T H E 10 - C O M P O N E N T S Y S T E M D i ^ H e d ^ E n ^ Fs ^ A e g ^ J d ^ T i - A e g ^ Fe Ts ^ T i -Ts ^ A l -Ts
End-member
Formula
Calculation
Conditions or restrictions
Aegirine (Aeg)
NaFe[Si2O6]
¼ Fe3þ 100
if Fe3þ Na
¼ Na 100
if Fe3þ4Na
¼ 05 Fe3þrest(1) 100
if Fe3þrest40
Ferri-Tschermak (Fe-Ts)
CaFe[FeSiO6]
Jadeite (Jd)
NaAl[Si2O6]
[if AlVI40 and (Na – Aeg)40]
VI
Ti-Aegirine (Ti-Aeg)
Ti-Tschermak (Ti-Ts)
¼ Al 100
if AlVI (Na – Aeg)
¼ (Na – Aeg) 100
if AlVI4(Na – Aeg)
¼ (Na – Aeg – Jd) 100
if (Na – Aeg – Jd) 2 (Ti þ Zr)
¼ 2 (Ti þ Zr) 100
if (Na – Aeg – Jd)42 (Ti þ Zr)
NaTi05(R2þ)(2)05[Si2O6]
[if (Na – Aeg – Jd)40 and (Ti þ Zr)40]
[if (Ti þ Zr)rest40 and AlIV40]
CaTi[AlAlO6]
¼ (Ti þ Zr)rest
(3)
100
IV
Al-Tschermak (Al-Ts)
¼ 05 Al 100
if (Ti þ Zr)rest4(05 AlIV) and Ca
¼ Ca 100
if Ca (Ti þ Zr)rest and 05 AlIV
(if Ca – Fe-Ts – Ti-Ts40 and AlIV – 2 Ti-Ts40 and AlVI – Jd40)
CaAl[AlSiO6]
¼ (Al
VI
– Jd) 100
if (AlVI – Jd) (AlIV – 2 Ti-Ts) and (Ca – Fe-Ts – Ti-Ts)
¼ (Al
IV
– 2 Ti-Ts) 100
if (AlIV – 2 Ti-Ts) (AlVI – Jd) and (Ca – Fe-Ts – Ti-Ts)
¼ (Ca – Fe-Ts – Ti-Ts) 100
Diopside (Di)
¼
(4)
(1 –
R2þrest(6) (1
XFe(5)) 100
– XFe) 100
¼
Mg2[Si2O6]
if R2þrest Carest
if R2þrest5Carest
(if R2þrest40 and Carest40)
CaFe[Si2O6]
if R2þrest Carest
¼ Carest XFe 100
Enstatite (En)
if (Ca – Fe-Ts – Ti-Ts) (AlVI – Jd) and (AlIV – 2 Ti-Ts)
(if R2þrest40 and Carest40)
CaMg[Si2O6]
¼ Carest
Hedenbergite (Hed)
if (Ti þ Zr)rest (05 AlIV) and Ca
R2þrest XFe 100
if R2þrest5Carest
¼ 05 (R2þrest – Di – Hed)
(if R2þrest – Di – Hed40)
(1 – XFe) 100
Ferrosilite (Fs)
Fe2[Si2O6]
¼ 05 (R2þrest – Di – Hed)
(if R2þrest – Di – Hed40)
XFe 100
(1)
Fe3þrest ¼ Fe3þ – Aeg; (2) R2þ ¼ Fe2þ þ Mg2þ þ Mn2þ; (3) (Ti þ Zr)rest ¼ [(Ti þ Zr) – (05 Ti-Aeg)];
Ti-Ts – Al-Ts; (5) XFe ¼ (Fe2þ þ Mn2þ)/(Fe2þ þ Mn2þ þ Mg2þ); (6) R2þrest ¼ R2þ – 05 Ti-Aeg.
1131
(4)
Carest ¼ Ca – Fe-Ts –