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American Mineralogist, Volume 93, pages 315–326, 2008
Wagnerite in a cordierite-gedrite gneiss: Witness of long-term fluid-rock interaction in the
continental crust (Ile d’Yeu, Armorican Massif, France)
Pavel Pitra,1,* Philippe Boulvais,1 Vladimir Antonoff,1,† and Hervé Diot2
Géosciences Rennes, UMR CNRS 6118, Université Rennes 1, 35042 Rennes, France
Pôle Sciences et Technologie, Université de La Rochelle, 17042 La Rochelle, France
1
2
Abstract
We describe the first occurrence in the Variscan Belt of Western Europe of the relatively rare phosphate wagnerite, ideally Mg2PO4F. It occurs in albite-rich, cordierite-gedrite-bearing gneisses on the
island of Ile d’Yeu, southern Armorican Massif, France. These gneisses are associated with a network
of shear zones that crosscut granitoid orthogneisses of calc-alkaline affinity. Wagnerite is zoned and
displays a rimward decrease of Fe/(Fe + Mg) from 0.16 to 0.08 and a concomitant increase in F. The
F content ranges 0.46–1.05 apfu, but critically depends on the choice of the analytical standard. Based
on phase diagrams calculated with THERMOCALC, we infer that the wagnerite-bearing orthoamphibole + cordierite + biotite + chlorite paragenesis equilibrated at ca. 550 °C, and pressures lower
than 4 kbar. The presence of staurolite relics requires similar temperatures, but pressures higher than
4 kbar, implying an evolution dominated by decompression. On the basis of whole-rock chemistry
and stable isotopes, we suggest that superimposed periods of metasomatic alteration throughout the
metamorphic history led to the prograde stabilization of the cordierite-gedrite gneiss at the expense
of the orthogneiss. This alteration involved aqueous fluids in isotopic equilibrium with local rocks
and caused significant loss of Ca, K, and Si, and gain of Mg and Na. We argue that the Na-enrichment
is the most significant difference between wagnerite-bearing and wagnerite-free Mg-rich, Ca-poor
rocks on Ile d’Yeu. This emphasizes the possible importance of Na metasomatism for the formation
of wagnerite. In light of comparisons with other wagnerite occurrences, we conclude that a long-term
fluid-rock interaction, typically associated with shear-zones, may be the rule rather than the exception
for the formation of wagnerite in metamorphic rocks unaffected by anatexis.
Keywords: Wagnerite, cordierite-gedrite gneiss, fluid-rock interaction, metamorphism, phase
diagrams, P-T and P-X pseudosections
Introduction
Wagnerite, a relatively rare constituent of metamorphic and
igneous rocks, is the Mg- and F-dominant member of the triploidite group, (Mg,Fe,Mn)2(PO4)(OH,F). It has been reported from
pegmatites (e.g., Staněk 1965), anatectic veins in a ultrahightemperature complex (e.g., Grew 1981; Grew et al. 2006), lowtemperature carbonate veins (Hegemann and Steinmetz 1927),
and from a wide range of metamorphic conditions: diagenetic
environments (Braitsch 1960a, 1960b), amphibolite-facies rocks
(Sheridan et al. 1976; Irouschek-Zumthor and Armbruster 1985;
Leroux and Ercit 1992), very high-pressure rocks (Chopin and
Sobolev 1995; Brunet et al. 1998), and low- to high-pressure
granulites (Novák and Povondra 1984; Vry and Cartwright
1994; Simmat and Rickers 2000; Ouzegane et al. 2003; Ren
et al. 2003). Thus, there appears to be no restriction on the P-T
stability range of wagnerite. In contrast, the OH-analog of wagnerite (β-Mg2PO4OH) has been synthesized only at pressures
above ca. 8 kbar, suggesting that wagnerite is stabilized toward
* E-mail: pavel.pitra@univ-rennes1.fr
† Present address: Institut National de la Recherche Scientifique,
490 rue de la Couronne, Québec G1K 9A9, Canada.
0003-004X/08/0203–315$05.00/DOI: 10.2138/am.2008.2597
315
lower pressures by the incorporation of F (Brunet et al. 1998).
The chemical composition of wagnerite could therefore be of
geobarometric interest.
Most authors concur that wagnerite occurrences are limited
to rocks that have high Mg and low Ca contents. The relatively
rare cordierite + orthoamphibole-bearing gneisses, which typically result from metamorphism under low- to medium-pressure
and moderate- to high-temperature conditions, display such
uncommon bulk compositions (Ca- and K-poor, Mg- and
Al-rich). These unusual chemical characteristics are ascribed
either to particular protoliths (evaporitic sediments, products of
pre-metamorphic weathering, or hydrothermal alteration) or to
metasomatic alteration during metamorphism (e.g., Spear and
Schumacher 1982). Many cordierite-orthoamphibole rocks are
associated with massive sulfide deposits and are interpreted
as volcanic rocks hydrothermally altered by seawater prior to
metamorphism (e.g., Schumacher 1988; Smith et al. 1992; Pan
and Fleet 1995; Witt 1999; Peck and Valley 2000; Roberts et
al. 2003). In other, less common cases, cordierite + orthoamphibole formed at the expense of granulite-facies parageneses
during fluid-assisted retrogression (Guiraud et al. 1996; Owen
and Greenough 2000) that may be localized along shear zones
(Dasgupta et al. 1999).
316
Pitra et al.: Wagnerite in a cordierite-gedrite gneiss
In this paper, we describe the first occurrence of wagnerite
in the west European Variscan belt. Wagnerite is hosted in
cordierite-gedrite gneisses associated with granodioritic orthogneisses. Using petrological and geochemical arguments, we show
that wagnerite crystallized together with the cordierite-gedrite
assemblage during decompression under amphibolite-facies
conditions. We argue that wagnerite and the cordierite-gedrite
gneiss formed at the expense of the orthogneiss by virtue of
a long-term synmetamorphic fluid-rock interaction. Our primary aim was to understand the geological context of this new
wagnerite occurrence and its possible implications for other
wagnerite-bearing localities.
Mineral analyses have been performed with a Cameca SX50 electron microprobe (Microsonde Ouest, Brest, France) operating in the wavelength-dispersive
mode. Data reduction was made using the ϕ-ρ-Z program (PAP). No allowance
was made during data reduction for the “excess oxygen” calculated by stoichiometry in the F-bearing phases, which may be a reason for the relatively high totals
obtained for F-rich wagnerite.
diameter of 1 µm and 5 µm. Whatever the conditions, the analyzed spots showed
no “beam browning” and the measured FKα intensity remained reasonably stable
through the 600 s. Similarly, point analyses were repeated 5–6 times on the same
spot. Again, the analyzed spots showed no “beam browning” and the measured
amount of F remained stable within 2σ of the first value (Fig. 1). Consequently,
we felt that there was no need to calculate a hypothetical “zero time” value of the
F content and that the point analyses yielded reproducible results.
Both apatite (Brunet et al. 1998; Ren et al. 2003, 2005) and topaz (Leroux and
Ercit 1992; Fialin and Chopin 2006) have been used as F-standards for microprobe
analyses of wagnerite. Point analyses of the same spots under different analytical
conditions and using the same standard yielded closely similar results, but the
results obtained using topaz systematically yielded values up to two times higher
than those using apatite (e.g., Fig. 1), and consequently we were not able to obtain
accurate F contents of wagnerite. The F concentrations in wagnerite are probably
underestimated when apatite is used since our apatite crystal is probably not in
the ideal crystallographic orientation, but overestimated when topaz is used; this
standard gave concentrations higher than the stoichiometric maximum possible
in wagnerite. Grew et al. (2007) also reported a discrepancy between F contents
obtained from topaz and apatite standards. Clearly, the choice of the analytical
standard is critical for accurate analysis of F and the analytical results depend more
on this choice than on the analytical conditions. It is possible that the only practical
approach to analyzing wagnerite for F would be to prepare a wagnerite F standard
in which F content would have been determined by another method.
Constituents other than fluorine
Formula recalculation of F-bearing minerals
Operating conditions for standard spot analyses for all minerals and all analyzed
elements (with the exception of F) were 15 keV accelerating voltage, 20 nA beam
current, 10 s counting time on the peak and a beam diameter of 1 µm. Standards
were natural albite (Na), orthoclase (K), corundum (Al), wollastonite (Ca, Si),
forsterite (Mg), hematite (Fe), apatite (P) and synthetic ZnS (Zn), MnTiO3 (Mn,
Ti), and Cr2O3 (Cr).
Direct application of the procedure explained in Deer et al. (1992, p. 680–681)
to EMP analyses of F-bearing minerals, where H2O is not directly analyzed, leads to
incorrect formulae. Taking wagnerite as an example, our procedure is to calculate
formulae on the basis of 4.5 oxygen-equivalents (or 9 positive charges). Starting
from the standard microprobe output (wt% oxides plus wt% F), the atomic proportions of all constituents, including F, were multiplied by 4.5/x, where x is the sum
of atomic proportions of oxygen from each analyzed oxide. However, the negative
charges of F must not be included in calculating x. The amount of the OH groups
is calculated by subtracting F apfu from the theoretical (OH + F) content. Consequently, the O = F correction plays no role in the structural formula calculation.
Fluorine
Electron microprobe analysis of minerals for F is difficult (e.g., Ottolini et
al. 2000) and there is no consensus on a favorable analysis strategy. The issue is
related primarily to the choice of the analytical standards, but analytical conditions
are also critical (e.g., Stormer et al. 1993; Ottolini et al. 2000; Fialin and Chopin
2006). Consequently, we have tested several sets of operating conditions combined
with two standards.
Fluorine was analyzed using a TAP crystal. Either Thomas Range (Utah) topaz
or Durango apatite were used as standards. The reliability of both standards was
tested using timed sequential analyses, in particular because of the problems related
to the crystallographic orientation (Stormer et al. 1993). Topaz yielded a nearly
constant FKα intensity over 600 s at 15 keV, 20 nA, and 1 µm beam diameter. At
6 keV, 30 nA, and 5 µm beam diameter, the FKα intensity was nearly constant
during the first 360 s, then slightly increased. For apatite, at 15 keV, 20 nA, and 1
µm beam diameter, integrated FKα intensity increased over the first 120 s (from
about 230 to about 280 counts for 6 s interval), and then regularly decreased and
approached a constant value (of about 200 counts for 6 s interval) after the next
240 s. At 6 keV, 30 nA, and 5 µm beam diameter, the FKα intensity was nearly
constant for the first 120 s and then increased slightly (from about 300 to about
350 counts for 6 s interval) over the next 360 s, approaching a constant value. This
behavior suggests that the apatite standard was not analyzed on the favored (100)
plane. However, the FKα intensity variation is much lower than that reported by
Stormer et al. (1993) for analyses on (001) even under the less favorable conditions (15 keV, 20 nA, 1 µm), and the problem nearly disappeared when apatite
was analyzed at 6 keV, 30 nA, and 5 µm beam diameter.
Fluorine in silicate minerals was analyzed using the two standards, but only at
the standard conditions described above (15 keV, 20 nA, 10 s, 1 µm in diameter).
Fluorine in wagnerite was first analyzed also using these standard conditions.
Analyses were then repeated on close, but not identical locations, using lower
voltage and a larger beam diameter (6 keV, 30 nA, 15 s, 5 µm). X-ray lines and
background offsets were carefully selected to minimize interferences from higherorder lines of heavier elements, in particular Fe and Mg (e.g., Raudsepp 1995).
The absence of interference from Fe and Mg on the analysis of F was confirmed
by the lack of correlation between the number of counts for Fe or Mg peak plotted against the counts on F background. Finally, a timed sequence and sequential
replicate analyses were collected on several locations to check the evolution of the
F signal (cf. Fialin and Chopin 2006). The intensity of FKα was measured over 600
s (100 count intervals of 6 s) at 15 keV, 20 nA and at 6 keV, 30 nA, with a beam
Geological setting
Ile d’Yeu is a small island off the coast of the department of
Vendée, France, in the southern part of the Armorican Massif
(Fig. 2), which represents the internal zones in this part of the
Variscan belt. Most of the island comprises granitoid orthog-
9
apparent F (wt%)
Electron microprobe analytical procedure
8
F-std = topaz
7
6
F-std = apatite
5
1
2
3
4
5
6
analysis no.
Figure 1. Apparent F content in sequentially replicated microprobe
analyses at the same spot on wagnerite using topaz and apatite as FPitra
et al.
- Fig. 16 kV, 30 nA, 5 µm beam diameter, and
standards.
Analytical
conditions:
15 s counting time on the peak per analysis.
317
PITRA ET AL.: WAGNERITE IN A CORDIERITE-GEDRITE GNEISS
neisses crosscut by an array of NW-SE-trending ductile thrust
shear zones with southward vergence. The most important shear
zone is located in the southern part of the island. The orthogneiss
is moderately deformed with foliation directions roughly parallel
to the shear zones (Fig. 2). The shear zones are highlighted by
alignments of “micaschist” layers (Mathieu 1938, 1945), which
locally display interesting mineral parageneses with staurolite,
cordierite, corundum, kyanite, sillimanite, and andalusite. The
paragenetic sequence has been interpreted in terms of a Barrovian
clockwise P-T path (up to ≈600–700 °C, 5 kbar) followed by
decompression and cooling (Semelin and Marchand 1984).
Sassier et al. (2006) argued that these “micaschists” are
actually products of metasomatic alteration of the orthogneiss
in amphibolite-facies shear zones that developed during the
compressive Variscan deformation. The orientation, vergence,
petrological, and geochemical characteristics of the shear zones
suggest that at least some of them formed early in the prograde
part of the metamorphic evolution and remained active until the
temperature peak, contemporaneous with the development of
the pervasive foliation in the orthogneisses. Associated intense
and long-term fluid circulation locally transformed the granitoid
orthogneisses into biotite-rich schists and kyanite-bearing lithologies (Sassier et al. 2006).
The Variscan age of the shear zones is supported by tectonic
correlations and by 40Ar/39Ar data on biotite (cooling age of 300–
304 Ma; G. Ruffet, unpublished data), whereas the emplacement
of the orthogneiss protolith is dated at 530 ± 8 Ma (2σ; in situ U/Pb
zircon LA-MC-ICPMS dating; C. Guerrot, unpublished data).
field relations, PetroGraPhy, and mineral
chemistry
The wagnerite-bearing samples were collected in the center of
a layer of spectacular cordierite-gedrite gneiss (Fig. 3), reported
here for the first time. The gneiss is found along the southern
coast of the island near the “Vieux Château” and the “Les Vieilles” beach, in association with the major shear zone (Fig. 2).
At “Les Vieilles,” only boulders were accessible and none were
found in situ. At the “Vieux Château,” the cordierite-gedrite
gneiss forms a layer several meters thick in the hanging wall of
the major shear zone within the orthogneiss. The other dimensions of the layer are difficult to estimate since it crops out in a
weathered cliff only ca. 5 m high that is hidden by vegetation on
one side and disappears under the sea on the other.
The orthogneiss is medium-grained (around 1–5 mm) and contains plagioclase, quartz, K-feldspar, biotite, accessory apatite and
zircon, and locally minor garnet. It displays a pervasive foliation
well defined by the preferred orientation of biotite and flattening of feldspar and quartz crystals. Plagioclase is the dominant
mineral and forms subhedral coarsely anti-perthitic crystals, up
to 5 mm in size. Quartz grains vary in size, reaching 5 mm, and
are characterized by lobate boundaries, which are attributed to
recrystallization by grain-boundary migration. Chessboard pattern extinction reveals the presence of both prismatic and basal
subgrain boundaries, typical of high-temperature deformation.
Myrmekite reveals the presence of subordinate K-feldspar. Biotite
flakes are locally partly chloritized, up to 3 mm long and display a
well-defined preferred orientation. Primary muscovite is absent.
Armorican
Massif
granitoid orthogneiss
major shear zone
shear zone
49°N
0°
Brest
Rennes
48°N
Port Joinville
47°N
Ile d’Yeu
5°W
3°W
N
Vieux
Château
sample location
high / low tide coast line
Les Vieilles
2 km
fiGure 2. Simplified geological map of Ile d’Yeu (modified after Sassier et al. 2006). Inset shows the location in the Variscan Armorican
Pitra et al. - Fig. 2
Massif, western France.
318
Pitra et al.: Wagnerite in a cordierite-gedrite gneiss
Figure 3. Field photograph of the cordierite-gedrite gneiss. Dark
acicular crystals of gedrite, without a clear preferred orientation are
distributed in the rock, form veins and large intergrowths with cordierite
(round clusters). The diameter of the coin is 19 mm.
Apatite forms subhedral stubby prisms up to 0.4 mm long.
Progressive lateral transition from the orthogneiss into the
cordierite-gedrite gneiss was documented at the outcrop scale. The
beginning of the transition is marked by the local appearance of
anhedral crystals of red garnet (up to 5 mm) and blue cordierite
(up to 10 mm) within the orthogneiss. Further toward the interior
of the layer, cordierite forms large anhedral bluish spots (up to 5
cm), in general intergrown with sheaves of dark acicular gedrite
and surrounded by a feldspar-rich aureole, up to 2 cm thick, devoid
of dark minerals. These blebs are commonly connected to gedritefilled veins, about 5–20 mm thick (Fig. 3). Parts of the rock still
preserve the texture of the neighboring orthogneiss.
In the central part of the layer, the rock is a relatively homogeneous, coarse-grained plagioclase-rich leucocratic gneiss
containing large oval nests of cordierite (up to 5 cm) and
numerous sheaves of gedrite, which display a weak preferred
orientation, but lack signs of intracrystalline strain. Orange to
rusty-brown crystals of wagnerite are irregularly distributed in
the matrix of this gneiss. Relative proportions of the minerals
vary from one thin-section to another due to large grain size
and local aggregations of plagioclase, gedrite, and cordierite.
The thin section VY11 is representative of most of the textural
and mineralogical variability observed in the studied samples
and contains plagioclase (≈45%), gedrite (≈26%), cordierite
(≈23%), biotite (≈2%), chlorite (≈2%), staurolite (<1%), and
locally quartz. The proportions were measured by point counting
on the scale of the thin-section. Accessory phases include wagnerite, ilmenite, pyrrhotite, rutile, zircon, monazite, and white
mica. Apatite is absent in the matrix and only appears along thin
cracks in wagnerite. On the thin-section scale, the minerals lack
any clear preferred orientation.
Representative mineral analyses are given in Table 1. The abbreviations used are and = andalusite; bi = biotite; cd = cordierite;
chl = chlorite; cor = corundum; ctd = chloritoid; ilm = ilmenite;
ky = kyanite; mu = muscovite; oa = orthoamphibole; pl = plagioclase; po = pyrrhotite; q = quartz; sil = sillimanite; st = staurolite;
wag = wagnerite; apfu = atoms per formula unit.
Plagioclase. Plagioclase forms stubby subhedral prisms up
to 5 mm long, whereas fine-grained plagioclase is granoblastic.
Grains are unzoned albite (XAn = 0.05, XOr < 0.01) in all the
analyzed samples.
Gedrite. Gedrite forms long euhedral prisms and needles
(up to 15 × 2 mm), generally gathered in sheaves or clusters
and contains numerous tiny inclusions of biotite, chlorite and
ilmenite. Formulae were calculated from stoichiometry to minimize Fe3+ content and maximize the amount of Na assigned to
the A site (Robinson et al. 1971; Spear 1980); the resulting Fe3+
contents were consistently zero. XFe ranges 0.45–0.49, Si ranges
6.06–6.37 apfu and octahedral Al (calculated as atomic Al + Si
– 8) is between 1.12 and 1.36 apfu. The Na content is between
0.46 and 0.66 apfu, whereas the K content is negligible (<0.01
wt%). Fluorine content is highly variable and reaches 0.23 apfu.
The amount of Cl is negligible (<0.03 wt%).
Cordierite. Cordierite forms large anhedral crystals (up to
10 mm) or granoblastic aggregates that include biotite, chlorite,
ilmenite and locally quartz. It is frequently found in association
with gedrite and wagnerite and systematically rims staurolite
(Fig. 4). It has a homogeneous chemical composition: XFe =
0.22–0.26; Na = 0.05–0.07 apfu.
Biotite. Biotite occurs as subhedral to euhedral crystals
of variable size (0.2–1.5 mm) that typically lack any signs of
replacement by chlorite. It is found in contact with all minerals except staurolite. Values of XFe range from 0.34 to 0.39, Ti
ranges from 0.09 to 0.13 apfu, octahedral Al (atomic Al + Si – 4)
ranges from 0.23 to 0.31. The analyzed F content depends on the
analytical standard: topaz and apatite gave 0.11–0.32 apfu and
0.02–0.06 apfu, respectively.
Chlorite. Chlorite is found in two distinct textures. (1) Subhedral to euhedral individual flakes or fan-like sprays (0.2–0.5 mm)
are commonly included in cordierite and plagioclase together
with gedrite and fresh biotite (Fig. 4b). Chlorite systematically
lacks inclusions of tiny rutile needles (sagenite), characteristic of
chlorite formed from alteration of biotite; these chlorite crystals
are observed in textural equilibrium with all the minerals, including staurolite. Values of XFe range from 0.30 to 0.35, octahedral
Al (atomic Al + Si – 4) ranges 1.33–1.44; chlorite inclusions
in wagnerite have a higher XFe (0.35) and a lower octahedral
Al (1.19–1.22). (2) Only exceptionally, tiny chlorite crystals in
association with white mica are present along cracks or at the
rims of cordierite, gedrite, and biotite, and are interpreted to have
grown at the expense of these minerals.
Wagnerite. Wagnerite forms subhedral to anhedral crystals
of variable size, up to 3 mm (Fig. 4). It encloses gedrite, biotite,
and chlorite, and is commonly rimmed by cordierite. Exceptionally, tiny inclusions of quartz are found in the core of large
crystals. Wagnerite crystals display distinct chemical zoning,
characterized by a rimward decrease in XFe from 0.16 to 0.08 and
a concomitant increase in F (Fig. 5). Analyses that used apatite
as a standard gave 0.46–0.70 F apfu, whereas analyses with
topaz gave 0.76–1.05 F apfu, i.e., the maximum values exceed
the amount allowed by stoichiometry. The amount of Ti ranges
between 0.01 and 0.03 apfu (0.60–1.16 wt%). The Mn content
is low, not exceeding 0.22 wt% (0.00 apfu). The amount of Cl
does not exceed 0.02 wt%.
Staurolite. Anhedral relics of staurolite (1–2 mm) are irregu-
Pitra et al.: Wagnerite in a cordierite-gedrite gneiss
319
Table 1. Representative electron microprobe mineral compositions
Sample
VY11
VY11
VY11
VY11
VY11
VY11
VY11
VY11
VY11 VY11
VY11
VY11
VY11
VY11
008
008b
008cT 008cA
301
301b
109
013
013b
204
h008
109
110
009
Anal. no.
wag
wag
wag
wag
wag
wag
oa
bi
bi
cd
chl
st
st
pl
Mineral
core
core
core
core
rim
rim
core
rim
Position
Anal. cond.*
T1
A1
T2
A2
T1
A1
T1
T1
A1
T1
1
1
1
1
SiO2
0.10
0.06
0.03
0.06
41.54
38.05 37.27 49.36
26.40
27.89
26.88
66.99
0.78
0.84
0.73
0.87
0.65
0.63
0.19
2.32
2.25
0.00
0.10
0.89
0.88
0.00
TiO2
0.00
0.05
0.01
0.09
18.20
16.86 16.74 33.44
22.27
53.71
54.82
20.56
Al2O3
39.58
39.16
39.84
39.51
42.69
41.85
13.02
14.42 14.77
9.83
20.32
3.04
2.40
0.00
MgO
12.25
12.49
12.41
12.70
9.74
9.45
22.14
14.72 14.59
6.10
18.05
14.17
14.77
0.16
FeO
0.07
0.05
0.12
0.22
0.09
0.00
0.04
0.01
0.08
0.00
0.09
0.03
MnO
0.17
0.00
0.02
ZnO
0.05
0.10
0.14
0.11
0.04
0.00
0.00
0.00
0.00
0.00
0.00
1.04
CaO
43.05
42.79
43.09
43.04
42.97
43.16
0.11
0.00
0.00
0.01
0.00
0.05
0.00
0.24
P2O5
0.00
0.00
0.01
0.00
1.94
0.84
0.78
0.32
0.00
0.02
0.00
11.24
Na2O
0.00
0.01
0.00
0.00
0.00
7.96
8.30
0.00
0.00
0.00
0.00
0.06
K2O
0.96
2.76
0.96
3.03
0.00
2.01
1.86
3.55
3.94
11.79
H2O(calc.)
F
9.35
5.48
9.36
4.98
11.87
7.22
0.37
1.09
0.19
0.00
Cl
0.00
0.00
0.00
0.00
0.00
Total
106.19 103.79 106.39 104.13 108.23 104.81
99.50
99.81 98.86 99.07
99.01
99.94
99.84
100.35
3.94
2.31
3.94
2.10
5.00
3.04
0.15
0.46
0.08
0.00
O = F
Total cor. 102.25 101.48 102.45 102.03 103.23 101.77
99.35
99.35 98.78 99.07
O basis
4.5
4.5
4.5
4.5
4.5
4.5
23
11
11
18
14
48
48
8
Si
0.003
0.002
0.001
0.002
6.136
2.806 2.773 4.993
2.683
7.935
7.678
2.927
0.016
0.018
0.015
0.018
0.013
0.013
0.021
0.129 0.126 0.000
0.007
0.190
0.189
0.000
Ti
0.000
0.002
0.000
0.003
3.169
1.465 1.468 3.987
2.667 18.012 18.459
1.059
Al
1.640
1.631
1.649
1.637
1.745
1.721
2.865
1.585 1.639 1.482
3.079
1.290
1.022
0.000
Mg
0.285
0.292
0.288
0.295
0.223
0.218
2.735
0.908 0.908 0.516
1.534
3.372
3.529
0.006
Fe2+
0.002
0.001
0.003
0.005
0.011
0.000 0.003 0.001
0.007
0.000
0.022
0.001
Mn
0.037
0.000
0.001
Zn
0.001
0.003
0.004
0.003
0.006
0.000 0.000 0.000
0.000
0.000
0.000
0.049
Ca
1.013
1.012
1.013
1.013
0.998
1.008
0.014
0.000 0.000 0.001
0.000
0.011
0.001
0.009
P
0.000
0.000
0.001
0.000
0.555
0.120 0.112 0.063
0.000
0.010
0.000
0.953
Na
0.000
0.000
0.000
0.000
0.001
0.749 0.788 0.000
0.000
0.000
0.000
0.003
K
0.178
0.516
0.178
0.562
0.000
0.370
1.829
1.746 1.956
8.000
OH
F
0.822
0.484
0.822
0.438
1.029
0.630
0.171
0.254 0.044 0.000
Cl
0.000
0.000
0.000
0.000
0.000
Total
3.960
3.961
3.965
3.963
4.017
3.973 17.514
9.762 9.817 11.043 17.977 30.857 30.900
5.008
XFe
0.148
0.152
0.149
0.153
0.113
0.112
0.488
0.364 0.357 0.258
0.333
0.723
0.775 Xan
0.049
XF
0.822
0.484
0.822
0.438
1.000
0.630
0.085
0.127 0.022
Xab
0.948
Xor
0.003
Notes: All Fe as FeO.
* Standards for F: T = topaz, A = apatite. Analytical conditions: 1–15 kV, 20 nA, 1 µm, 10 s. 2–6 kV, 30 nA, 5 µm, 15 s for F; other elements as 1. Structural formulae
are based on the number of oxygen-equivalents indicated as “O basis.” Blank = not analyzed. XFe = atomic Fe/(Fe + Mg); XF = atomic F/(F + Cl + OH).
larly dispersed in the rock and everywhere surrounded by a rim
of cordierite (Fig. 4c). Exceptionally, tiny subhedral staurolite
grains are enclosed in plagioclase. Staurolite is found in textural
equilibrium with chlorite and ilmenite. Values of XFe increase
rimwards from 0.72 to 0.78. Manganese and zinc are low (<0.03
and <0.04 apfu, respectively), Ti ranges 0.18–0.23 apfu.
Ilmenite. Ilmenite forms subhedral laths up to 0.7 mm long.
It contains up to 3% of the hematite component and up to 3% of
geikielite. The amount of pyrophanite is negligible.
Quartz. Quartz is absent from the matrix of most samples,
but is locally found as tiny rounded inclusions in plagioclase,
cordierite, and wagnerite. In some samples, quartz forms aggregates of relatively large lobate grains (up to 2 mm) that display
both prismatic and basal subgrain boundaries.
Based on textural relations, we infer that the main equilibrium
assemblage is orthoamphibole (gedrite) + cordierite + biotite +
chlorite, which replaces an earlier assemblage containing staurolite, probably in equilibrium with chlorite. Despite the macroscopic
weak preferred orientation of gedrite, microscopic textures suggest mostly static recrystallization, postdating the development
of the foliation in the orthogneiss. Inclusions of biotite, chlorite,
and gedrite and straight contacts with cordierite suggest that the
crystallization of wagnerite is related to the (quartz-free) orthoamphibole (gedrite) + cordierite + biotite + chlorite assemblage.
However, rare inclusions of quartz in wagnerite cores suggest that
its crystallization may have started earlier.
Modeling of mineral equilibria
To infer the P-T conditions and relate the observed mineral
assemblages to a specific part of the metamorphic history of Ile
d’Yeu, phase diagrams were calculated with THERMOCALC
v. 3.25 (Powell and Holland 1988) and the internally consistent
thermodynamic data set of Holland and Powell (1998, November
2003 update). Taking into account the mineral chemistry, phase
relations should be treated in the model system NCKFMASH.
However, because Na cannot be taken into account in the present
solid-solution model for orthoamphibole, plagioclase is effectively the only Ca- and Na-bearing phase in the system and there
is no advantage of including it in the calculations. Consequently,
phase relations were calculated in the model system KFMASH.
320
Pitra et al.: Wagnerite in a cordierite-gedrite gneiss
0.16
a)
0.15
Fe / (Fe+Mg)
0.14
0.13
0.12
0.11
0.10
0.09
0.08
0.45
F-std = apatite
0.50
0.55
0.60
0.65
0.70
F (apfu)
0.16
b)
0.15
Fe / (Fe+Mg)
0.14
0.13
0.12
0.11
0.10
0.09
0.08
0.75
F-std = topaz
0.80
0.85
0.90
0.95
1.00
1.05
F (apfu)
Figure 5. Wagnerite is characterized by a rimward decrease in XFe
and a concomitant increase in F. However, absolute values of F depend
al. - Fig.
5 (a) F-standard = apatite. (b)
on the choice Pitra
of the et
analytical
standard.
F-standard = topaz.
Figure 4. Photomicrographs (plane-polarized light) of the cordieritegedrite gneiss. (a) Large wagnerite crystal, partly including gedrite,
surrounded by albite and cordierite, with inclusions of biotite and chlorite.
(b) Cordierite enclosing crystals of biotite, chlorite, gedrite, ilmenite,
pyrrhotite, and wagnerite. Biotite and chlorite coexist without replacing
one another. (c) Anhedral relict of staurolite surrounded by cordierite.
Small anhedral crystal of wagnerite partly including biotite in a matrix
bearing oa + cd + bi + chl.
Mixing models for most solid solutions were taken from Holland
and Powell (1998). Orthoamphibole was modeled according
to Xu et al. (1994) and Guiraud et al. (1996). In addition, the
free energy of gedrite in orthoamphibole (Iged,oa) was increased
by 10 kJ/mol (M. Guiraud, personal communication) using the
DQF approach, to better fit the composition of natural orthoamphiboles. For P-T pseudosection calculations, we used a local
bulk composition, derived from the modal proportions and the
chemical compositions of minerals rather than the whole-rock
composition obtained by chemical analysis because of the large
grain-size of the rock. This approach is particularly appropriate
when trying to model observed mineral assemblages and textures
within a given compositional domain of heterogeneous rocks.
Modal proportions were determined by point counting, assuming
equilibration at the scale of the thin-section. Pure H2O-fluid was
considered to be present in excess.
The result is shown in Figure 6. Orthoamphibole is stable
at temperatures above 500 °C and cordierite is stable at temperatures higher than 500 °C and pressures lower than 6 kbar.
Staurolite occupies two distinct regions: staurolite + chlorite is
stable in the range 530–600 °C at pressures higher than 4 kbar,
whereas staurolite + biotite is stable above 640 °C and 7 kbar.
The assemblage orthoamphibole + cordierite + biotite + chlorite
is stable between 500 and 600 °C at pressures lower than 4 kbar.
Pitra et al.: Wagnerite in a cordierite-gedrite gneiss
tion from which we infer that the cordierite-gedrite gneiss was
derived from the orthogneiss through fluid-rock interaction. This
genetic relation is a posteriori confirmed by the good correlation
defined by the elements that are classically considered immobile during fluid-rock interaction at mid-crustal conditions (Al,
Ti, P, Th, Nb, Ta, Zr, Hf, Y, Ni). These elements define a line,
which corresponds to an isocon (Grant 1986). The slope of the
isocon allows for the calculation of the mass variation during
alteration. The slope value (a = 1.35) points to about 26% loss of
mass, which corresponds to a major decrease in volume during
alteration, provided that the rock densities were not significantly
affected by metasomatism (Sassier et al. 2006 reported only
1–3% difference in density between unaltered orthogneiss and
metasomatized, strongly hydrated “micaschist” from other shear
zones). Elements that plot below the isocon were lost during
alteration. From Figure 8, we infer that leaching of SiO2, K2O,
and CaO account for most of the mass lost. An interesting feature
is the nearly complete leaching of K2O, Rb, Ba, and Pb, which
probably reflects the destabilization of K-feldspar at some stage
of the evolution. On the other hand, the amounts of MgO and
Na2O have significantly increased. The enrichment of MgO was
described by Sassier et al. (2006) in other shear zones of the
island and was attributed to high-temperature interaction. Enrichment of Na2O distinguishes the wagnerite-bearing gneisses from
most other shear zones on Ile d’Yeu (see below). Regarding REE,
the slight enrichment in the cordierite-gedrite gneiss relative to its
protolith, especially in LREE, is likely related to the stabilization
of an accessory phase like monazite, which displays a special
affinity for LREE over HREE. The immobility of P (and leaching
of Ca) during the metasomatic process would have allowed the
growth of monazite, which then would have acted as a trap for
LREE from the fluid in an open system.
The inferred stability of staurolite with chlorite implies conditions above 4 kbar and between 530 and 600 °C for this early
assemblage. In consequence, the observed replacement of the
relic staurolite + chlorite-bearing assemblage by the wagneritebearing matrix assemblage orthoamphibole + cordierite + biotite
+ chlorite suggests a metamorphic evolution dominated by
decompression at temperatures around 550 °C.
Whole-rock chemistry
KFMASH (+ H2O)
hl mu
2
ky chl
mu q
4
3
2
and chl mu q
chl mu q
and bi chl q
oa
1
hl
9
cd
hl mu
i
il b
s
oa
st and
chl mu q
cd chl
mu q
7
c
bi
l
ch
q
or
bi c
o
oa sil bi cor
oa sil
bi chl
sil l
cd ch
a
o bi 5
d sil
oa c
r
bi co
8
6
oa cd
bi chl
cor
oa cd
bi chl
q
oa cd bi cor
VY 11
SiO2 Al2O3 MgO FeO K2O
50.27 13.36 20.70 15.41 0.26
cd bi chl q
500
y
ak
4
y bi
oa st bi
chl q
5
oa ky
bi chl
oa k
6
3
chl q
q
mu
chl
st ky
P (kbar)
ctd
ky
7
st chl
mu q
oa st c
1
ctd chl
mu q
chl m
uq
8
oa st ky bi
q
oa ky c
9
Figure 6. Pressure-temperature pseudosection
for the cordierite-gedrite gneiss (sample VY11). Thick
unnumbered lines correspond to Al2SiO5 univariant
equilibria. Small filled circles are KFMASH invariant
points and numbered lines are the following univariant
equilibria: 1 – ctd + ky = st + chl + q, 2 – st + chl +
q = oa + and/sil/ky (degenerate = continues across
the two invariants and ends at the and + chl + mu + q
field), 3 – oa + chl + mu = bi + ky + q, 4 – st + chl ±
ky = oa + cor, 5 – chl + sil = oa + cd + cor, 6 – chl +
and/sil + q = oa + cd, 7 – oa + chl + mu = st + bi + q,
8 – oa + chl + mu = bi + and/sil + q, 9 – chl + mu +
and + q = cd + bi. White = divariant fields (5 phases);
light gray = trivariant fields (4 phases); dark gray =
quadrivariant fields (3 phases). Some divariant fields
are so narrow that they appear as lines (e.g., oa + st +
ky + bi + cor), whereas other fields are not labeled for
the sake of clarity; their assemblages can be deduced
from assemblages in adjacent fields.
q
Samples were finely ground in an agate mortar before analysis.
Whole-rock chemical compositions were obtained by ICP-AES
(major elements) and ICP-MS (minor elements) at the SARM
laboratory, CRPG-CNRS, Nancy, France. Results are reported in
Table 2. Analytical uncertainties range from ±1% (SiO2, Al2O3) to
±10% for major and trace elements, depending on concentration
level. Only Cu and Mo were found to be below detection; the limits
for these elements are 3 and 0.5 ppm, respectively.
The orthogneiss has the composition of a slightly peraluminous granodiorite (A/CNK = 1.12), similar to the rocks analyzed
by Sassier et al. (2006). Niobium and tantalum are depleted
relative to other incompatible elements (spidergram not shown
here), which is consistent with a calc-alkaline signature. The
REE pattern (Fig. 7) with LaN about 100, an enrichment of
LREE over HREE (LaN/LuN = 6), a flat HREE pattern and a large
negative Eu anomaly (Eu/Eu* = 0.47), is typical of granodioritic
compositions.
The cordierite-gedrite gneiss is richer in REE than the orthogneiss but has a similar pattern (Fig. 7): LaN about 200, LaN/LuN =
9.5, Eu/Eu* = 0.41, and a flat HREE pattern. A derivation of the
cordierite-gedrite gneiss from the orthogneiss is thus suggested
by their respective REE patterns. Comparison of the two rocks
in an isocon diagram (Fig. 8) is allowed by the field observa10
550
Pitra et al. - Fig. 6
600
650
321
T (°C)
700
322
Pitra et al.: Wagnerite in a cordierite-gedrite gneiss
Sample
VY11G
VY11OC
VY11G
VY11OC
Rock
ortho
oa-cd
ortho
oa-cd
SiO2
67.39
56.52
Ba
682.5
26.4
Al2O3
14.63
19.20
Pb
19.6
4.2
Fe2O3
4.89
8.07
Sr
106.6
48.8
MnO
0.06
0.06
Rb
104.7
9.4
MgO
1.22
5.40
Cs
4.4
1.1
CaO
2.10
0.56
U
3.5
6.1
Na2O
3.28
6.95
Th
11.8
17.0
K2O
3.52
0.37
Ta
1.2
1.5
TiO2
0.66
0.91
Nb
12.2
17.0
P2O5
0.20
0.31
Zr
250.9
387.7
LOI
0.85
1.38
Hf
6.8
9.4
Total
98.79
99.73
Y
37.7
55.0
Zn
66.4
20.9
La
33.4
68.1
Ga
19.8
23.0
Ce
70.6
148.9
Ge
1.5
1.1
Pr
8.6
17.8
Ni
11.2
16.7
Nd
34.1
68.3
Cr
35.0
37.8
Sm
7.5
14.3
Co
6.6
13.0
Eu
1.1
1.9
Cu
19.7
bdl
Gd
7.1
13.7
V
41.5
41.2
Tb
1.1
2.2
Mo
bdl
1.5
Dy
6.9
11.9
Sn
4.0
4.8
Ho
1.3
2.0
W
0.6
0.9
Er
3.9
5.5
Tm
0.6
0.8
A/CNK
1.123
1.494
Yb
3.8
4.9
Ca/P
13.26
2.34
Lu
0.6
0.7
δ18O‰
10.4
7.6
Notes: Oxides in wt%, other elements in ppm. Ca/P is atomic ratio; A/CNK =
mol Al2O3/(CaO + Na2O + K2O), calculated without any correction. bdl = below
detection limit.
Sample/ REE chondrite
1000
VY11OC - cd-oa gneiss
100
Table 3. Oxygen isotope compositions (‰ vs. SMOW) of whole rocks
(WR) and minerals separates from the cordierite-gedrite
gneiss (VY11OC) and the host orthogneiss (VY11G)
Sample
VY 11G
VY 11OC
WR
10.4
7.6
q
11.3
9.5
bi
6.4
5.1
cd
7.4
5.8
ged
18
4.9‰/470 °C
4.4‰/520 °C
∆ Oq-bi/Tapp
18
∆ Oq-cd/Tapp
2.1‰/710 °C
3.7‰/600 °C
∆18Oq-ged/Tapp
Note: Tapp = apparent temperatures of equilibration calculated using quartzmineral fractionation factors of Zheng (1993a, 1993b).
100
gains
80
70
0.2LREE
HREE
40
5Na
st
n
co
Si
10Mg
50
8Hf
e
um
ol
tv
n
a
5Th
La
60
Y
3Ni
losses
0.1Zr
30
20
10
0
0
3Fe
Nb
10Ta
10Ti
10Ca
20P
Pb
10
20
30
0.5Sr
10K
40
0.5Rb
50
60
0.1Ba
70
y = 1.35x
R2 = 0.99
80
90
100
protolith (orthogneiss VY11G)
10
VY11G - orthogneiss
1
5Al
90
co
n
Table 2. Whole-rock compositions of the wagnerite-bearing cordierite-gedrite gneiss (VY11OC) and the adjacent orthogneiss
(VY11G)
iso
The oxygen isotope compositions (Table 3) of whole rocks
and mineral separates were determined at the stable isotope
laboratory of the University Rennes 1. Minerals were separated
by careful hand picking under the microscope and crushed in a
boron carbide mortar. Powders were reacted using BrF5 following the method of Clayton and Mayeda (1963). The liberated
O2 was then converted to CO2 by reaction with hot graphite.
Isotopic compositions were measured on CO2 using a VG SIRA
10 triple collector mass spectrometer. During the analytical session, measurements of NBS 28 quartz standard gave δ18O = 9.34
± 0.09 (1σ, n = 14). Analyses were normalized to NBS 28 (δ18O
= 9.60‰) by adding 0.26‰ to measured values.
The δ18O value of the orthogneiss (δ18O = 10.4‰) lies in the
range defined by other orthogneisses from Ile d’Yeu (Sassier et al.
2006) and is consistent with a calc-alkaline signature. Quartz and
biotite separates provided values that compare well with minerals from other orthogneisses and are consistent with relatively
high-temperature isotopic equilibration: an apparent temperature
of equilibration of 470 °C (Table 3) has been calculated using
the fractionation factor of Zheng (1993b). The δ18O value of the
cordierite-gedrite gneiss (δ18O = 7.6‰) is lower than the values
of the other shear zones of the island. However, if the fluid, with
which the orthogneiss reacted to form the cordierite-gedrite
gneiss, had been in isotopic equilibrium with the local crustal
rocks, as in the case of the shear zones studied by Sassier et al.
(2006), the theoretical δ18O value of the cordierite-gedrite gneiss
should match the measured value. This theoretical value can be
product (cd-oa gneiss VY11OC)
Stable isotopes
La Ce Pr Nd Pm Sm Eu Gd Tb Dy Ho Er Tm Yb Lu
Figure 7. Chondrite-normalized (Boynton 1984) REE patterns of the
orthogneiss and the cordierite-gedrite gneiss. Plotted using the program
GCDKit (Janoušek
et al.
Pitra et
al.2006).
- Fig. 7
Figure 8. Isocon diagram (Grant 1986) showing variations in
element concentration between the orthogneiss (protolith) and the
cordierite-gedrite gneiss (product). Major elements are plotted as wt%
oxide and trace elements as ppm (Table 2). Scaling factors are indicated.
Pitra
et al.
- Fig.
8 immobile elements (black diamonds).
Straight line
is the
isocon
fit to
Slope and correlation coefficient (R2) of the isocon are indicated in
the lower right corner of the diagram. Squares and triangles indicate
elements that increased and decreased in amount, respectively, during
metasomatism.
323
Pitra et al.: Wagnerite in a cordierite-gedrite gneiss
Discussion
Metamorphic evolution
P-T pseudosections only represent the mineralogy of one specific rock and it is therefore legitimate to inquire to what extent
the conclusions drawn from such a diagram (Fig. 6) are dependent
on the choice of this composition. This problem is even more
evident in metasomatic environments that by definition imply
variations of bulk chemistry. The composition used in Figure 6
corresponds to a final state, whereas the transformation of the
orthogneiss might have been progressive. To check the influence
of the bulk chemistry and in particular the possibility that the
staurolite relics do not represent different P-T conditions, but only
reflect the progressive evolution of the effective bulk composition under constant pressure and temperature, we have calculated
an isothermal P-X pseudosection (Fig. 9). This diagram displays
the mineralogical variation as a function of pressure and chemical
composition of the rock at 560 °C. The bulk composition varies
between that of an unaltered Ile d’Yeu orthogneiss and that of the
cordierite-gedrite gneiss. The geometrical progression from the
left to the right of the diagram models a progressive metasomatic
evolution related to increasing fluid-rock interaction. To avoid
local compositional heterogeneity, we chose as the starting composition an average of several orthogneiss samples rather than
the sample (VY11G) of the orthogneiss immediately adjacent
to the cordierite-gedrite gneiss. However, the composition of
sample VY11G gives a similar diagram and leads to the identical conclusion: the staurolite-bearing parageneses and cordierite
+ orthoamphibole parageneses develop under distinct pressure
conditions, respectively, above and below 4 kbar, whatever the
composition of the rock.
Consequently, despite the metasomatic history of the rock,
the observation of staurolite relics within the orthoamphibole +
cordierite + biotite + chlorite assemblage implies a P-T evolution
dominated by decompression at temperatures around 550 °C.
This evolution passes close to the aluminum silicate invariant
point and is fluid-assisted. The abundance of fluids facilitates
phase transitions and such an evolution is consistent with the
observation of replacement of kyanite by andalusite and locally,
by sillimanite (V. Antonoff, unpublished data).
Furthermore, staurolite only appears on the right-hand side
of the diagram, implying that an important part of the total
7
6
5
P (kbar)
calculated using: (1) the mean δ18O value of quartz separates
from orthogneisses (δ18Oq = 9.8‰, Sassier et al. 2006); (2) the
modal composition given above; and (3) the fractionation factors of Zheng (1993a, 1993b) at 550 °C, the metamorphic peak
temperature. The value calculated this way (7.5‰) is close to
the one actually measured (7.6‰). Moreover, the δ18O values
of the quartz and biotite separates in both the cordierite-gedrite
gneiss and the shear zones of Sassier et al. (2006) are similar.
This result supports the conclusion that the metasomatizing
fluid was in isotopic equilibrium with the Ile d’Yeu rocks at
metamorphic temperatures. In addition, quartz-mineral pairs in
the cordierite-gedrite gneiss gave isotopic temperatures (Table
3) consistent with the peak conditions estimated for the area,
≈600–700 °C (Semelin and Marchand 1984), thereby ruling out
the involvement of fluids on the retrograde path.
4
KFMASH (+ H2O)
@ T = 560°C
oa
bi
mu
q
2
1
oa bi chl
mu q
bi
mu cd bi
q and q
oa
chl
mu
q oa st chl
mu q 2
8
d
an q
oa mu oa and
bi
bi q
3 and
chl
mu
q
oa and bi chl q
iq
nd b
da
oa c
6
oa cd bi chl q
oa
cd
bi
oa cd
bi q
cd bi
mu q cd
bi
q
st chl
mu q
oa
cd
bi
chl
cd bi ksp q
0
0.1
0.2
orthogneiss
0.3
0.4
SiO2 Al2O3 MgO FeO K2O
81.16 6.19 3.98 6.17 2.51
0.5
0.6
0.7
0.8
0.9
1
shear-zone: cd-oa gneiss
SiO2 Al2O3 MgO FeO K2O
50.27 13.36 20.70 15.41 0.26
Figure
pseudosection. Numbering of
Pitra 9.
et Pressure-composition
al. - Fig. 9
univariant equilibria and color of fields are the same as in Figure 6.
metasomatic transformation of the rock likely occurred during
or prior to the crystallization of staurolite, possibly during the
prograde evolution of the orthogneiss.
Nature and timing of chemical alteration
Field relations and whole-rock chemistry (Fig. 8) suggest
that the wagnerite-bearing cordierite-gedrite gneiss resulted
from the metasomatic alteration of the neighboring granodioritic
orthogneiss. Sassier et al. (2006) presented structural, petrological, and geochemical evidence for the synmetamorphic origin of
the other metasomatic shear zones on Ile d’Yeu. Fluid flow was
active over a significant proportion of the prograde metamorphic
evolution, i.e., from the low-temperature stages until the temperature peak. Sassier et al. (2006) cited both low-temperature strain
localization and high-temperature syntectonic crystallization
within the shear zones as evidence for this conclusion. These
authors also noted that there is good geometrical consistency
between shear zone pattern and regional pervasive fabric, which
would not be expected if deformation affected a preexisting net
of alteration zones; this geometrical consistency argues against
a late magmatic or near-surface origin of the alteration zones.
Our observations support the conclusions reached by Sassier
et al. (2006).
A particular feature of the alteration is the strong SiO2 leaching, which induced large mass and volume losses (Fig. 8). At
mid- to lower-crustal conditions, fluids most often flow upward,
along a down-temperature gradient. This generally leads to SiO2
enrichment in infiltrated rocks (Dipple and Ferry 1992) because
silica solubility decreases with decreasing temperature. Sassier
et al. (2006) described such silica enrichment in shear zones that
display smooth strain gradients, which is the structural sign of
initiation at high temperature. On the other hand, shear zones with
very sharp strain gradients display silica depletion (and volume
loss). Sassier et al. (2006) interpreted these characteristics to
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Pitra et al.: Wagnerite in a cordierite-gedrite gneiss
result from the flow of fluids along an up-temperature gradient, early in the deformation history, at a time when isotherms
were reversed due to thrusting (cf. Selverstone et al. 1991). It
is very likely that, in our case, the strong silica depletion is also
due to such an alteration. A retrograde origin for the alteration
is precluded by the observations that cordierite and biotite are
relatively free of secondary phases and oxygen isotopes preserve amphibolite-facies temperatures. Silica depletion in the
wagnerite-bearing gneiss is thus related to an early alteration, at
some stage during the prograde metamorphic evolution.
Another peculiar feature is the significant Na2O enrichment
of the wagnerite-bearing gneiss. This Na-metasomatism was not
reported by Sassier et al. (2006) in the other shear zones of the
island. Sodium metasomatism is a very common phenomenon
that occurs in a variety of contexts (e.g., Perez and Boles 2005).
It can be related to the per descensum alteration of crustal rocks
by the flow of surface-derived fluids in extensional settings
(McLelland et al. 2002) or to the flow of metamorphic fluids in
ductile shear zones (Rubenach and Lewthwaite 2002). Alteration
results from changes in fluid pressure and/or temperature related
to movement rather than from chemical disequilibrium between
fluid and rocks. In many cases, desilication and albitization are
associated (Cathelineau 1986; Boulvais et al. 2007). It is therefore
tempting to relate the silica loss and the sodium enrichment observed in the metasomatic gneiss to a single event. This alteration
episode must have been confined to the shear zone on the south
side of the island, as it is not recognized in most rocks of the
island, and appears to have occurred early in the metamorphic
evolution when fluids were channeled in faults or narrow shear
zones. Part of the Ca and K loss evident in the cordierite-gedrite
gneiss may be as well related to the same event, by albitization
of plagioclase and K-feldspar, respectively.
At high temperature, fluid flow was still active, as suggested
by the occurrence of gedrite-filled veins and the coarse grain size
in the cordierite-gedrite gneiss (Fig. 3). Enrichment of MgO also
might be related to this alteration episode, as it was in the shear
zones studied by Sassier et al. (2006).
The wagnerite-bearing cordierite-gedrite gneiss thus results
from successive episodes of alteration. It is unlikely that externally
derived fluids were involved during these episodes because the
fluid that interacted and equilibrated with the cordierite-gedrite
gneiss also appears to be in equilibrium with the host orthogneisses
(rock-dominated fluid system). Nevertheless, one cannot rule out
the possibility that the second, high-temperature flow erased the
isotopic effects of the first, low-temperature one, even if this lowtemperature event consisted of the invasion of low δ18O marine
or meteoric water. Second, many elements display an immobile
behavior during metasomatism. The isocon is indeed rather well
defined in the Grant diagram (Fig. 8) and there is no fractionation
of element pairs like Ga/Al, Zr/Hf, or Y/Ho. This argues against
the influx of fluids with a high metasomatic potential able to
induce non-charge-and-radius-controlled fractionation. The
simplest way to interpret the formation of the wagnerite-bearing
cordierite-gedrite gneiss is then to invoke the involvement of
aqueous fluids moving along an up- and then a down-temperature
gradient, at different stages of the metamorphic history. This model
supports the long-lasting, syn-metamorphic character of fluid-rock
interactions at Ile d’Yeu. However, it is not possible to distinguish
whether these interactions resulted from a continuous fluid influx
or a succession of discrete events.
Compositional controls on wagnerite formation
Although it appears problematical to obtain a reliable electron
microprobe estimate of the F content in wagnerite, the rimward
decrease in XFe and the concomitant increase in XF observed in the
wagnerite crystals are significant as they are independent of the
choice of the standard (Fig. 5). This relation could be attributed
to the F-Fe avoidance principle (Ekström 1972; Valley et al.
1982). Ren et al. (2003) suggest that, despite the existence of the
synthetic end-member zwieselite [Fe(PO4)F], this phenomenon
may operate in triplite-group minerals. However, they argue that
it cannot be extrapolated to the ordered polytypes of wagnerite,
due to the crystal structure. If this is the case, the observed relation would be purely fortuitous and the evolution of XFe and XF
of wagnerite would be related to independent processes. Several
hypotheses may be proposed.
The XFe ratio of wagnerite is controlled by the Fe-Mg exchange between wagnerite and other Fe-Mg minerals under
changing P-T-X conditions. However, the minerals coexisting
with wagnerite do not display significant zoning. Thus, the only
explanation would be that Fe-Mg diffusion in wagnerite is much
slower than in silicate minerals and wagnerite is the only mineral
where compositional zoning is preserved.
The increase in the amount of F either reflects the evolution
of P-T conditions and in particular the pressure decrease, or it
echoes the progressive evolution of the fluid toward more F-rich
compositions. The first hypothesis is attractive. Brunet et al.
(1998) reported experiments showing that β-Mg2PO4OH, the
OH-analog of wagnerite, is stable at pressures above about 8
kbar and suggested that F stabilizes wagnerite toward low pressures. Accordingly, OH-rich wagnerite would be expected at high
pressures and the observed rimward increase in F would neatly
fit with the calculated pressure decrease. The observed inclusion
of quartz in wagnerite core tends to strengthen this scenario. It
suggests that wagnerite crystallization may have started in the
stability field of the early, higher-pressure paragenesis staurolite
+ chlorite + muscovite + quartz before equilibrating with orthoamphibole + cordierite + biotite + chlorite. However, because of
the high variance of the parageneses, compositions of the phases
in equilibrium (and specifically the XF ratio of wagnerite) depend
not only on P-T conditions, but also on the bulk chemistry of the
system, and in particular, on the composition of the fluid, even
though wagnerite is not the only (F-OH)-bearing mineral present.
Although the hypothesis of the pressure control of wagnerite composition is appealing, the influence of the progressively evolving
fluid chemistry cannot be excluded in the present case.
Irouschek-Zumthor and Armbruster (1985) demonstrated
that an Mg-rich, Ca-poor bulk composition is not a sufficient
condition for wagnerite formation. The rocks from the metasomatic shear zones described by Sassier et al. (2006) are Mg-rich
and Ca-poor, but lack wagnerite and contain apatite instead.
Apatite is the stable phosphate in most geological conditions.
Consequently, the formation of wagnerite (or other less common
phosphates such as lazulite, bearthite, arrojadite, sarcopside, and
chopinite) instead of or in addition to apatite is possible only in
rocks that possess an effective Ca/P ratio lower than that of apatite
Pitra et al.: Wagnerite in a cordierite-gedrite gneiss
– 5/3 (=1.67, Irouschek-Zumthor and Armbruster 1985; Brunet
and Chopin 1995). In the analyzed wagnerite-bearing sample
(VY11OC), bulk Ca/P ratio is of 2.34, similar to the ratios observed in the shear zones of Sassier et al. (2006). However, Ca/P
ratio drops to 0.57 or less when the amount of Ca sequestered in
plagioclase is deducted. Indeed, plagioclase (An5) is very abundant in the wagnerite-bearing samples, whereas the shear zones
described by Sassier et al. (2006) are virtually plagioclase-free
due to their low Na content. A similar situation is reported from
the Lepontine Alps (Irouschek-Zumthor and Armbruster 1985),
where wagnerite-bearing rocks have a similar and even higher
bulk Ca/P than rocks lacking wagnerite. However, wagneritebearing samples are systematically rich in Na2O (5.5–6.0 wt%)
and contain abundant albite (An6–10), which results in a significant
decrease in the available Ca relative to P. Likewise, on Santa Fe
Mountain (Sheridan et al. 1976), wagnerite occurs in sillimaniteplagioclase gneisses that contain 46–75% plagioclase (An6–8).
These observations suggest that Na enrichment is conducive to
the formation of wagnerite because it leads to stabilization of
abundant albite, which acts as a sink for Ca. In summary, it is
the availability of Ca and P, rather than bulk Ca/P ratio, which
determine whether wagnerite will form.
The role of fluids in wagnerite formation
Before the work of Sassier et al. (2006) and this study, Ile
d’Yeu was interpreted as a sequence of interlayered metasedimentary and metamagmatic rocks (Mathieu 1938, 1945; Semelin
and Marchand 1984). We have provided evidence that the development of wagnerite-bearing cordierite-gedrite gneisses on
Ile d’Yeu is related to long-term fluid-orthogneiss interaction
involving low temperature sodium metasomatism followed by
a higher-temperature flow event. Intriguing similarities lead us
to suggest that a similar scenario might apply to other wagnerite
occurrences in metamorphic rocks around the world, although
many of the host rocks were interpreted as being derived from
precursors having an unusual composition.
The Mg-rich, Ca-poor coesite-bearing pyrope-phengite
quartzites (“whiteschists”) from Dora Maira, western Alps, occur as layers, from a few centimeters to several tens of meters
thick, within granitoid orthogneisses. They contain wagnerite
inclusions in pyrope megablasts, accompanied by other uncommon phosphates such as bearthite (Chopin et al. 1993; Chopin
and Sobolev 1995). Compagnoni and Hirajima (2001) proposed
long-term fluid-rock interaction within shear zones for the origin of these rocks. In this first example, wagnerite formation
might well be a consequence of fluid circulation. Interestingly,
in another massif of the western Alps, Monte Rosa, bearthite
and lazulite, albeit without wagnerite, were described in similar
Mg-rich, Ca-poor whiteschist ascribed to fluid-assisted metasomatism of adjacent granites (e.g., Chopin et al. 1991; Pawlig and
Baumgartner 2001).
At the classic localities of east-central Front Range, Colorado
(Sheridan et al. 1976), wagnerite occurs at several places within
thin lenses of a sillimanite-plagioclase gneiss. This gneiss is a
local lithologic variant of a thin, regionally persistent layer of
a rutile-bearing sillimanite-quartz gneiss within quartz-feldspar
gneisses (Marsh and Sheridan 1976; Sheridan et al. 1976). This
layer may well represent a regional-scale shear zone, where
325
fluid-rock interaction led to the development of Mg-rich, Ca-poor
peraluminous gneisses from the quartz-feldspar-dominated protolith. Interestingly, cordierite-gedrite gneisses occur in the same
band (Heimann et al. 2006). In the Simano Nappe, Lepontine
Alps (Switzerland), wagnerite occurs within a 2.5 m thick layer
of “Mg-rich metapelitic rock” that “cuts through monotonous
granitic gneisses” (Irouschek-Zumthor and Armbruster 1985).
Again, this layer likely corresponds to a shear zone that would
have channeled fluid flow. Finally, a metasomatic origin has been
inferred for wagnerite-bearing Mg- and Al-rich rocks from the
Reynolds Range, Australia (Vry and Cartwright 1994). Thus,
wagnerite stabilization and fluid flow seem associated once more.
The situation is different in the case of wagnerite occurrences in
anatectic environments (e.g., Ren et al. 2005; Grew et al. 2007),
where the favorable bulk chemistry may be achieved by virtue
of migmatitic differentiation or interaction of rocks with P-rich
melts. In the light of these comparisons, we believe that intense
and long-lived fluid-rock interaction, typically associated with
shear-zones, may be the rule rather than the exception for the
formation of wagnerite in non-migmatitic metamorphic rocks.
Acknowledgments
We are grateful to Eric Beneteau from the LEBIM (University of Angers) base
at Ile d’Yeu for his kind assistance during the fieldwork. Marcel Bohn helped with
the microprobe analyses. We benefited from enlightening discussions with Ramón
Capdevila, Serge Fourcade, David Dolejš and Denis Gapais. We are thankful to
Christian Chopin, Jennifer Thomson and Ed Grew for thorough and constructive
reviews and discussions. Additionally, we also acknowledge Ed Grew’s helpful
editorial handling of the manuscript.
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Manuscript received February 5, 2007
Manuscript accepted September 6, 2007
Manuscript handled by Edward Grew