the magmatic-hydrothermal transition in peralkaline ......the lct group comprises peraluminous...
TRANSCRIPT
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The Magmatic-Hydrothermal Transition in Peralkaline Rhyolite Magma at Terceira, Azores
by
Caitlin Beland
A thesis submitted in conformity with the requirements for the degree of Masters of Applied Science
Department of Earth Sciences University of Toronto
© Copyright by Caitlin Beland 2014
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The Magmatic-Hydrothermal Transition in Peralkaline Rhyolite
Magma at Terceira, Azores
Caitlin Beland
Masters of Applied Science
Department of Earth Sciences
University of Toronto
2014
Abstract
The geochemistry of quartz-hosted melt (MI) and fluid inclusions (FI) in quartz syenite from
Terceira, Azores was investigated to provide insight into late-stage evolution of peralkaline melts
and the behaviour of high field strength (HFSE) and rare-earth elements (REE) at the magmatic-
hydrothermal transition. Crystalline and hydrous MI analyzed by laser ablation-inductively-
coupled plasma mass-spectrometry (LA-ICP-MS) show extreme magmatic enrichment of HFSE
and REE. Sanidine crystallization resulted in enrichment of the melt in HFSE, REE and volatiles.
Halite-saturated FI analyzed by LA-ICP-MS show lower total REE abundances than melts, and a
general enrichment in HREE. Comparison of REE distribution patterns of MI and miarolitic
zircon and monazite suggest late-stage melt evolution by monazite, then zircon and pyrochlore
fractionation. Microthermometry of FI suggests maximum trapping conditions of 675°C, 120
MPa. The residual evolved to very volatile-rich compositions and initially exsolved a hydrosaline
melt that was diluted to lower salinities by aqueous-fluid exsolution on cooling.
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Acknowledgments
Special thanks to Professor Jim Mungall for patience, assistance and guidance. Thanks to
Professor Jake Hanley at Saint Mary’s University for assistance and guidance with data
processing and many helpful discussions. Also, many thanks to Dr. Colin Bray for guidance with
microthermometry, and Dr. Duane Smythe for general guidance with analytical techniques. Last
but not least, infinite thanks to my father especially for teaching me the values of perseverance
and determination, my family and Simon Urbain for their continued support.
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Table of Contents
Acknowledgements .......................................................................................................................iii
Table of Contents ...........................................................................................................................iv
List of Tables ................................................................................................................................vii
List of Figures ..............................................................................................................................viii
List of Appendices .........................................................................................................................ix
Chapter One: Introduction .............................................................................................................1
Origin of peralkaline melts ................................................................................................2
Solubility of incompatible lithophile elements, volatiles and their magmatic
enrichment in peralkaline melts .....................................................................................5
Solubility of the HFSE in peralkaline melt ............................................................5
Solubility of volatiles in peralkaline melt ..............................................................6
Influence of volatiles on solubility of the HFSE in peralkaline melt .....................7
Magmatic enrichment of the HFSE in peralkaline melt .........................................8
Solubility of the HFSE in aqueous fluid and their hydrothermal transport ........................8
Solubility of the HFSE in aqueous fluid .................................................................8
The role of aqueous fluids in the genesis of HFSE deposits .................................10
Partitioning behaviour of the HFSE between silicate melt and aqueous fluid .................11
Chapter Two: Research paper prepared for the Journal of Petrology ...........................................13
Abstract .............................................................................................................................13
Introduction .......................................................................................................................14
Geology of Terceira...............................................................................................18
Materials and Methods ......................................................................................................21
Sample collection and petrography .......................................................................21
Microthermometry ................................................................................................22
LA-ICPMS ............................................................................................................22
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Electron Microprobe .............................................................................................26
Petrographic Observations ................................................................................................27
Quartz syenite .......................................................................................................27
Mineral compositions ............................................................................................33
Quartz-hosted inclusions .......................................................................................37
Fluid inclusions .........................................................................................38
Melt inclusions ..........................................................................................40
Hydrous melt inclusions ...........................................................................41
Fluid Inclusion Microthermometry ...................................................................................43
Compositions of Melt and Fluid Inclusions ......................................................................47
REE .......................................................................................................................47
Element ratio variations ........................................................................................50
Discussion .........................................................................................................................52
Phase assemblages ................................................................................................52
Late-stage mineralogy of Terceira quartz syenite and inferred
compositional features of the residual liquid ............................................52
Miaskite or agpaite? ..................................................................................53
Intensive parameters .............................................................................................57
Oxygen fugacity........................................................................................57
Pressure and temperature of entrapment of melt and fluid inclusions.......60
Temperature of quartz crystallization .......................................................64
Compositions of minerals, melts and fluids ..........................................................65
Boundary layer effects and post-entrapment crystallization .....................65
Miarolitic monazite and zircon compositions ...........................................66
Melt inclusion compositions .....................................................................67
Fluid inclusion compositions ....................................................................70
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FI-MI pairs and distribution coefficients ..................................................72
Conclusions ......................................................................................................................80
References ........................................................................................................................82
Chapter Three: Conclusions and future work ..............................................................................93
References …...………………...………………………………………………………………..96
Appendix A: Microthermometry Results ....................................................................................111
Appendix B: LA-ICPMS Data of Fluid and Melt Inclusion Analyses .......................................117
Appendix C: Abandoned/Failed Components of Research ........................................................127
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List of Tables
Table 1: Comparison of analytical results of Mungall & Martin (1996) and the present study on
sample P16 ....................................................................................................................................25
Table 2: EMP operating conditions ..............................................................................................26
a: monazite analyses .........................................................................................................26
b: zircon analyses ..............................................................................................................27
Table 3: A list of HFSE-bearing minerals found in miarolitic cavities ........................................30
Table 4: Ti abundances in different zones (seen in CL) of selected quartz grains and calculated
temperatures of crystallization at various pressures .....................................................................32
Table 5: EMP results for monazite analyses .................................................................................33
Table 6: EMP results of zircon analyses .......................................................................................34
Table 7. Summary of vapour disappearance and halite dissolution temperatures, salinity and
estimate of minimum trapping pressure for FI homogenizing by halite dissolution ....................45
Table 8: Pairs of melt and fluid compositions and calculated distribution coefficients ..............75
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List of Figures
Figure 1.1: Schematic phase diagram of NaAlSi3O8-SiO2-H2O at moderate pressure ..................4
Figure 2.1: Examples of time-resolved spectra for LA-ICPMS analysis of inclusions ................24
Figure 2.2: Paragenetic sequence diagram for Terceira quartz syenite ........................................29
Figure 2.3: Back-scattered electron (BSE) images of various miaroles in quartz syenite ............31
Figure 2.4: BSE images of botryoidal zircon ................................................................................36
Figure 2.5: Chondrite-normalized REE distribution patterns for monazite and zircon ................37
Figure 2.6: Representative images of various types of quartz-hosted fluid inclusions ................39
Figure 2.7: Representative photomicrographs of various quartz-hosted melt inclusion types in
quartz syenite ................................................................................................................................41
Figure 2.8: Representative photos of transitional (HMI) inclusions .............................................42
Figure 2.9: Histogram of measured homogenization temperatures for all fluid inclusions ..........44
Figure 2.10: Histogram of calculated salinities for all fluid inclusions ........................................44
Figure 2.11: Box-whisker plots for Type I and II fluid inclusions ...............................................46
Figure 2.12: Chondrite-normalized plots of REE abundances of MI and HMI ............................48
Figure 2.13: Chondrite normalized plot of REE abundances for FI .............................................49
Figure 2.14: Element ratio plots for Terceira inclusions ..............................................................51
Figure 2.15: Qualitative μCaO versus μNa2O plot .......................................................................57
Figure 2.16: Homogenization temperature-salinity plot of all FI .................................................62
Figure 2.17: Pressure-temperature diagram of trapping conditions of FI .....................................63
Figure 2.18: HFSE+REE distribution plots for co-entrapped melts and fluids ............................73
Figure 2.19: Fluid-melt distribution coefficient as a function of agpaitic index of the melt .......77
Figure 2.20: Calculated fluid-melt distribution coefficients (Df/m
) for La, Y and Yb as a function
of ASI ...........................................................................................................................................78
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List of Appendices
Appendix A: Microthermometry Results ....................................................................................111
Appendix B: LA-ICPMS Data of Fluid and Melt Inclusion Analyses .......................................117
Appendix C: Abandoned/Failed Components of Research ........................................................127
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Chapter 1
Introduction
Peralkaline igneous rocks are important hosts for economically exploitable deposits of high field
strength elements (HFSE), including rare earth elements (REE) and in particular the heavy-REE
(HREE). For example, of the 41 known HFSE occurrences of economic interest in North
America, which are hosted in or sourced from 10 different geologic environments, seven
deposits occur in peralkaline granites, syenites or nepheline syenites (Mariano & Mariano,
2012).
There are two main geochemical trends and deposit types of incompatible lithophile elements
associated with extremely fractionated magmas; LCT and NYF (Cerny, 1992). LCT is so named
for enrichment in Li, Cs and Ta>Nb, as well as Be, B, P, Mn, Ga, Rb, Sn and Hf (Cerny, 1992;
London, 2008). The LCT group comprises peraluminous granitic pegmatites derived from I- or
S-type peraluminous granite plutons (Cerny, 1992; London, 2008), generally in orogenic tectonic
settings (Martin & De Vito, 2005). When this enrichment pattern originates from I-type granites
it is produced by anatexis of intermediate igneous source materials, which releases abundant Li
(and likely Ta) to the partial melt (London, 2008). The S-type LCT granite suite is produced by
anatexis of pelitic metasediments that are inherently rich in Li, Al and P followed by extreme
fractionation (London, 2008). The NYF group is named for its enrichment in Nb>Ta, Y, and F,
as well as the HREE and the HFSE (Zr>Hf) (Cerny, 1992; London, 2008). This group comprises
small, hypabyssal plutons that are peralkaline, silica-undersaturated or – oversaturated in
composition (Cerny, 1992). NYF-type deposits occur in anorogenic tectonic settings (Martin &
De Vito, 2005) and are ultimately mantle-derived (London, 2008). The origin of the NYF
geochemical signature remains poorly understood (London, 2008). HREE enrichment common
in NYF-deposit types may be associated with partial melts of source rocks with abundant biotite
and amphibole (LREE-selective phases) as restite (London, 2008). Known occurrences of
strongly peralkaline, NYF-type deposits include Ilimaussaq, Greenland, the Lovozero and
Khibina complexes, Russia, Tamazeght, Morocco, Strange Lake, Thor Lake, Kipawa and Mont
Saint Hilaire, Canada.
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Origin of Peralkaline Melts
There is general agreement that peralkaline melts are produced by differentiation of low degree
partial melts of the mantle. Low degrees of partial melting appear to be critical for generation of
peralkaline melts (Schwab & Johnston, 2001). For example, Schiano et al. (1998) analyzed
olivine-hosted glass inclusions in spinel lherzolite and found that the inclusions’ compositions
represented a cogenetic set of melts formed by partial melting of spinel lherzolite with melt
fractions (F) ranging from 0.2 – 5 %. Initial melts (F=0.2%) are highly silicic (64 wt % SiO2) and
alkali-rich (11 wt % alkali-oxides) (Schiano et al., 1998). With further melting, however, the
proportions of FeO, CaO, MgO and Cr2O3 in the melt increase, with concomitant decrease in
SiO2, Na2O, K2O, and Al2O (Schiano et al., 1998). From these observations, it is apparent that
low degrees of partial melting are necessary to generate peralkaline melt from a mantle source
rock. Another apparent requisite is a metasomatized mantle source rock. Subduction of spilitized
basaltic crust releases an alkali-rich fluid with high Na/K and this fluid can metasomatize mantle
rocks to produce sodic amphiboles and/or increase the aegirine or jadeite component in pyroxene
(Markl et al., 2010). Later anatexis of this metasomatized mantle rock can yield sodic or persodic
melts (Markl et al.., 2010) if melting is limited to small fractions. Consequently, many
researchers call for metasomatism of the mantle below peralkaline igneous provinces prior to
melt generation (Eby, 1985; Montero et al., 1998; Smith et al., 1999; Bea et al., 2001;
Goodenough et al., 2002; Muzio et al., 2002; Jahn et al., 2009; Kohler et al., 2009; Ozgenc &
Ilbeyli, 2009).
Fractional crystallization of a transitional or alkali basaltic parental magma, derived from partial
melting of metasomatized mantle, is the essentially undisputed model for the genesis of
peralkaline, silica-undersaturated melts (Sorensen, 1997; Marks et al., 2003; Man0n et al., 2006;
Schonenberger & Markl, 2008; Schilling et al., 2011). However there remains contention over
the genesis of peralkaline silica-saturated rocks (Bonin, 2007; Di Carlo et al., 2010). For these
rocks, many researchers invoke a petrogenetic model of fractional crystallization from an alkali
basalt parent combined with crustal assimilation to reach silica-saturation (Marks et al., 2003;
Kozlovsky et al., 2007; Martin 2006; Markl et al., 2010). Nonetheless, a number of experimental
and geochemical modelling studies have shown that peralkaline, silica-saturated melts can be
generated in exactly the same fashion as their silica-undersaturated counterparts; by protracted
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fractional crystallization of a mantle-derived transitionally alkaline basalt (Mungall, 1993;
Mungall & Martin, 1995; Montero et al., 1998; Muzio et al., 2002; Nekvasil et al., 2004; Bonin,
2007; Macdonald et al., 2008; Jahn et al., 2009; Ozgenc & Ilbeyli, 2009; Shellnutt et al., 2009;
White et al., 2009; Di Carlo et al., 2010; Frost &Frost, 2010; Ronga et al., 2010; Macdonald,
2012). Specifically, geochemical modelling of basalt-pantellerite suites has indicated that
pantellerite melt can be produced from basalt by 90-95 % fractional crystallization, with removal
of amphibole playing an essential role in generating peralkaline residua from metaluminous
intermediate magmas (Mungall & Martin, 1995, White et al., 2009; Ronga et al., 2010).
Evidence from experiments and modelling therefore demonstrate that silica-saturation in
peralkaline rocks can be achieved by fractionation of transitionally alkaline basalt without any
input of crustal material.
From the above discussion, it is clear that peralkaline silica-saturated compositions are the end
result of extreme fractionation. It follows that peralkaline igneous rock–hosted HFSE deposits
are an even more extreme case. Why then are such extremely fractionated magmas relatively rare
in nature? This is because common alkalic, tholeiitic and calc-alkaline basalts lose geochemical
degrees of freedom by removal of calcic plagioclase and ferromagnesian minerals as they evolve
towards either of the silica-saturated or silica-undersaturated eutectics in the system SiO2-
NaAlSiO4-KAlSiO4, at which the melt completely solidifies over a very short range in
temperature (Mungall, 2014). However, the eutectic temperature is influenced by composition
and compositions with [Na+K]/Al >>1, or
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and aqueous fluid shrinks, the crest of the critical curve moves to lower temperature and the
invariant point moves to more water-rich compositions, until at approximately 3.5 GPa, when the
solvus shrinks to a point (critical point) that joins the eutectic (Mungall, 2014). At pressures
above 3.5 GPa the invariant point no longer exists, so the melt can evolve to very hydrous
compositions without freezing until it becomes an aqueous fluid rich in dissolved silicate
components (Mungall, 2014).
Fig. 1. Schematic phase diagram of NaAlSi3O8-SiO2-H2O at moderate pressure, reproduced
after Boettcher & Wyllie (1969).
Supercritical behaviour in the system NaAlSi3O8-SiO2-H2O only occurs at very high pressures.
Such behaviour has been documented for similar systems, for example experiments of Paillat et
al., (1992) in the system NaAlSi3O8-H2O determined the position of the critical end point at
670°C, 1.5 GPa, and Bureau & Keppler (1999) performed experiments in the system NaAlSi2O6-
H2O and found complete miscibility between jadeite melt and aqueous fluid at 800°C, 1.5 GPa.
The critical pressure however can be lowered by the addition of fluxes such as B, F, or excess
Na2O, as indicated by the experimental results of Sowerby & Keppler (2002) in the system
albite+H2O. Sowerby & Keppler (2002) showed a depression of the critical end point to
pressures as low as 0.4 GPa (shallow crustal values). Smirnov et al. (2012) performed
experiments in the system Na2O-SiO2-H2O and found supercritical behaviour when initial Na2O
was above 2 wt %, with the critical end point at 600°C, 0.15 GPa. However, as initially
proposed by Tuttle and Bowen (1958), supercritical behaviour has been documented at very low
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temperatures and pressures in the system Na2Si2O5-NaAlSi3O8-H2O (Mustart, 1972). Mustart
(1972) determined the position of the second critical end point in this system to be at 322°C and
90 MPa. Fractional crystallization of strongly peralkaline compositions can therefore proceed to
arbitrarily small melt fractions without ever reaching a eutectic and becoming vapour saturated
(Mungall, 2014). Though not directly observed (i.e., experimentally), supercritical behaviour in
strongly peralkaline systems with compositions analogous to natural pantellerites has been
inferred on the basis of textural or melt and fluid inclusion evidence by a number of researchers
(Mungall & Martin, 1996; Webster & Rebbert, 1998; Thomas et al., 2000; Thomas et al., 2006;
Marks et al., 2003; Andersen et al., 2010).
Solubility of incompatible lithophile elements and volatiles and their magmatic enrichment in peralkaline melts
Solubility of the HFSE in peralkaline melt
High fields strength elements are highly soluble in peralkaline silicate melts, attaining
concentrations in the weight percent range at extremely high peralkalinity. It has been
experimentally shown that the solubility of baddeleyite increases with increasing peralkalinity in
silicate melts (Marr et al., 1998). Similarly, experimental work has demonstrated increased
solubility of zircon (Watson, 1979; Watson & Harrison, 1983; Linnen & Keppler, 2002), hafnon
(Linnen & Keppler, 2002) and manganotantalite in granitic melts with increasing peralkalinity
(Van Lichterfelde et al., 2010). The solubility of Hf is positively and linearly correlated with
excess Na2O (Davis et al., 2003). Increased solubility of columbite with increasing peralkalinity
of the melt has been demonstrated experimentally (Linnen & Keppler, 1997; Fiege et al., 2011).
Solubility of monazite is also strongly compositionally dependent and is positively correlated
with melt peralkalinity (Montel, 1985; Krenn et al., 2012). These components can reach
concentrations up to a few wt % in peralkaline melts. For example, experimental results of
Linnen & Keppler (2002) indicated a maximum solubility of about 3.75 wt % baddeleyite and
almost 7 wt % hafnon in a melt with (Na+K)/Al of 1.6.
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Solubility of volatiles in peralkaline melts
In addition to the HFSE, peralkaline melts are capable of dissolving large concentrations of
volatile and fluxing components such as H2O, F and Cl. Dingwell et al. (1997) demonstrated
increasing H2O solubility with increasing (Na+K)/Al of the melt, and that H2O solubility at
shallow crustal levels doubles with the addition of excess Na2O equivalent to levels observed in
peralkaline provinces.
Fluorine solubility is also dependent on peralkalinity, as shown by the positive correlation
between fluorite solubility and peralkalinity of granitic melts (Scaillet & Macdonald, 2004;
Gabitov et al., 2005). Furthermore, Wilding et al. (1993) studied quartz-hosted melt inclusions in
peralkaline rhyolite and observed correlation of high H2O contents with high F concentrations
and peralkalinity. Early experimental studies suggest that increased Cl solubility is linked with
increasing (Na+K)/Al of the melt. The work of Metrich & Rutherford (1992) for example,
showed that although total iron oxide concentration (FeO*) is an important control on Cl
solubility in SiO2 rich melts, peralkalinity is an overriding factor. Furthermore, Signorelli &
Carroll (2000) conducted Cl solubility experiments on phonolitic melts ranging from
peraluminous to peralkaline, coexisting with aqueous fluid and Cl-rich brine, and observed the
highest Cl concentrations in peralkaline compositions. These authors also demonstrated a
negative correlation between Cl solubility and pressure (Signorelli & Carroll, 2000). In contrast,
more recent work has shown that although (Na+K)/Al has a significant influence on the
solubility of Cl in phonolitic compositions, it is less important for more silica-rich compositions
(Signorelli & Carroll, 2002). Instead, the degree of depolymerisation of the melt is more
influential on Cl solubility in trachytic or rhyolitic compositions, and solubility of Cl increases
with increasing proportions of non-bridging oxygen (Signorelli & Carroll, 2002). However,
increasing peralkalinity serves to depolymerise the melt, so the positive correlation of Cl
solubility and peralkalinity still stands, and is explained by the depolymerising effects of excess
alkalis. This is consistent with the findings of various experimental investigations indicating
enhanced Cl solubility with increasing F concentrations, which serve to depolymerise the melt
(Webster 1997a, 1997b; Webster & Rebbert, 1998).
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Influence of volatiles on solubility of the HFSE in peralkaline melts
Extensive evidence provided in the literature for high solubility of both the HFSE and volatile
components in peralkaline silicate melts leads to the question of whether one influences the
other; however available investigations of this problem are limited. Linnen (2005) investigated
the effect of [H2O]melt on the solubility of columbite, tantalite, wolframite, rutile, zircon and
hafnon in peralkaline melts and found that it has no influence whatsoever. In contrast, monazite
solubility has been shown to increase with increasing H2O content of the melt (Rapp & Watson,
1986), suggesting different controls of REE and HFSE solubility in silicate melts. Fiege et al.
(2011) found that F concentration has no effect on the solubility of columbite or tantalite, and
suggested instead that the solubility of these phases is related to melt structure, and enhanced by
increased amounts of network modifying cations (minimum solubility at (Na + K)/Al = 1). The
findings of Fiege et al. (2011) indicate that Nb and Ta do not form complexes with F- in silicate
melts. Similarly, the solubility of REE-phosphates is independent of F in water-saturated
haplogranitic systems (Keppler, 1993). In marked contrast, experiments by Keppler (1993)
indicated strongly enhanced solubility of manganocolumbite and manganotantalite in the
presence of F. Recent experiments have also shown that the Nb content of loparite increases
when the mineral crystallizes in the presence of F-bearing fluid (Suk et al., 2013). Additionally,
solubility of Ti and Zr in hydrous haplogranitic melts in equilibrium with rutile and zircon show
positive linear and quadratic correlations with F content, respectively (Keppler, 1993). Keppler
(1993) suggested that complexation with non-bridging oxygen (made available by fluorine-
induced depolymerisation) or direct complexation with F is the solubility mechanism for these
elements. In a recent study of Zr complexation in hydrous silicate melts and aqueous fluids at
high pressure, Louvel et al. (2013) found the dominant species to be polymeric Zr-O-Si/Na, in
both F-bearing and F-free haplogranitic melts containing 15.5-33 wt % dissolved H2O, but state
that their results regarding the extent of Zr-F complexation in such compositions are
inconclusive.
In summary, current experimental evidence on the influence of volatile components on the
solubility of HFSE in peralkaline silicate melts via complexation or depolymerization is unclear.
Considering the hypothesis of Fiege et al. (2011) that Nb and Ta solubility is a function of melt
depolymerisation, it is unclear why H2O has no effect on the solubility of these elements since
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these H2O serves also to depolymerise the melt. Moreover, natural systems that are highly
enriched in the HFSE are very commonly also highly enriched in F. It is evident that solution
mechanisms for the REE and HFSE are exceptionally complex and likely involve interplay
between volatile content and peralkalinity of the melt. Fiege et al. (2011) also suggest that low
fO2 of the melt can enhance HFSE solubility. This hypothesis definitely warrants further
investigation as many peralkaline, igneous HFSE ore deposits have been shown to form under
reducing conditions (Marks et al., 2003; Ryabchikov & Kogarko, 2006; Salvi & Williams-Jones,
2006; Schonenberger & Markl, 2008; Markl et al., 2010).
Magmatic enrichment of the HFSE in peralkaline melts
Magmatic enrichment of HFSE to near exploitable levels in silicate melts residual to protracted
fractional crystallization has been indicated by a number of studies (Kovalenko et al., 1995;
Chabiron et al., 2001; Schmitt et al., 2002; Thomas et al., 2006; Andreeva & Kovalenko, 2011;
Kynicky et al., 2011; Papoutsa & Pe Piper, 2013; Sun et al., 2013). Enrichment in REE, U and
Th at the Kvanefjeld deposit at Ilímaussaq is considered purely magmatic in origin and occurred
as a result of closed-system, protracted and uninterrupted crystallization of peralkaline nepheline
syenitic magma, pre-enriched (through fractional crystallization of an alkali basaltic parent) in
HFSE (Sorensen et al., 2006; Sorensen et al., 2011). The colossal Nb+REE deposits at
Lovozero, Kola Peninsula, Russia are also believed to be purely magmatic in origin (Kogarko et
al., 2002). Replacement of loparite by complex Nb and REE minerals at Lovozero is attributed to
reaction with the residual melt (not hydrothermal fluids) (Kogarko et al., 2002).
Solubility of the HFSE in aqueous fluid and their
hydrothermal transport
Solubility of the HFSE in aqueous fluid
Hydrothermal processes can concentrate HFSE through remobilization, transport and re-
deposition. ZrO2 solubility increases with decreasing temperature in HF-bearing (0.1 m) fluids;
sufficient amounts of Zr can be transported to account for concentrations observed in natural, F-
rich hydrothermal systems (Migdisov et al., 2011). This concentration of HF in aqueous fluid is
comparable to greisenizing fluids with up to 1.2 m HF (Shapovalov & Setkova, 2012).
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Furthermore, with similar concentrations of HF in hydrothermal systems, significant quantities
of REE can be dissolved (Migdisov et al., 2009). Recent numerical modelling of fluid interaction
with HFSE/REE-bearing rocks of the Strange Lake Complex by Gysi & Williams-Jones (2013)
suggests that the HFSE/REE can be hydrothermally transported at HF concentrations as low as
0.4 m. It was also shown that the LREEs are more soluble in hydrothermal fluids than the HREE
(Migdisov et al., 2008; Migdisov et al., 2009), and that REE-chloride complexes predominate
over those with fluoride in aqueous fluids at high temperature (200°C) (Migdisov & Williams-
Jones, 2002, 2007). Due to the common association of hydrothermal fluorite with HFSE
deposits, fluoride has long been considered as the most likely transporting ligand of these
elements (Wood, 1990; Haas et al., 1995). However, the aforementioned studies demonstrate
that this is improbable, and Williams-Jones et al. (2012) suggest instead that F- acts as a binding
ligand promoting precipitation and that the bulk of REE transportation in hydrothermal fluids is
achieved through complexation with chloride. Correspondingly, Mayanovic et al. (2007)
determined that Gd-Cl complexation is significant at 300°C, and that at higher temperatures Gd
is associated with partially hydrated chloride complexes. Hydroxide can also be an important
complexing agent for the REE in hydrothermal solutions, as indicated by the experiments at
shallow crustal pressures and 300°C of Pourtier et al. (2010).
Evidence for hydrothermal mobilization and transport of HFSE has also been documented in
natural systems; for example, Salvi et al. (2000) and Sun et al. (2013) reported the occurrence of
HFSE daughter phases in fluid inclusions, and Gilbert & Williams-Jones (2008) documented
vapour transport of REE as evidenced by REE-enriched (relative to natrocarbonatite)
encrustations surrounding volcanic vents at Oldoinyo Lengai, Tanzania. Fluorite-REE
mineralization in an NYF-type pegmatite in the Pikes Peak granite, Colorado, (Gagnon et al.,
2004) and in quartz-syenite in New Mexico, USA (Williams-Jones et al., 2000) and the
mineralization at the colossal Bayan Obo REE deposit in China is considered hydrothermal in
origin, in the sense that an orthomagmatic fluid remobilized and redeposited the REE in a
carbonatite (Lai et al., 2012; Lai & Yang, 2013). Though not directly observed, many authors
infer hydrothermal mobilization and transport of the HFSE when comparing mineral or whole
rock compositions in hydrothermally altered rock to its fresh counterpart and finding enrichment
in the HFSE (Jiang et al., 2003).
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The role of aqueous fluids in the genesis of HFSE deposits
Most peralkaline igneous HFSE deposits are believed to have been generated by a combination
of magmatic and hydrothermal processes. At the Thor Lake rare metal deposit, NWT, for
example, the main rare metal ore mineral is zircon of both magmatic and hydrothermal origin
(Sheard et al., 2012). Repeated injections of aegirine-nepheline syenite, fractional crystallization
and convection produced cumulate layers of magmatic eudialyte, then zircon (Sheard et al.,
2012). Late-stage exsolution of orthomagmatic aqueous fluid from the syenite then remobilized
Zr and the HREE to form pseudomorphic zircon after eudialyte, secondary colloform zircon rims
on pre-existing grains and REE-fluorcarbonates (Sheard et al., 2012). The LREE were also
remobilized, but deposited more distally as a result of later mixing with an externally derived Ca-
bearing fluid (Sheard et al., 2012).
Similarly, the proposed deposit model for Strange Lake, Labrador, involves magmatic
enrichment in HFSE via fractional crystallization of peralkaline granite followed by exsolution
of a highly saline, F-rich orthomagmatic fluid (Salvi & Williams-Jones, 2006). Interaction of this
fluid with peralkaline granite allowed leaching of HFSE from the primary mineral assemblage,
complexation of the HFSE with the ligands F- and Cl
- and their subsequent transport as
complexes (Salvi & Williams-Jones, 2006). The orthomagmatic aqueous fluid then mixed with
Ca-bearing meteoric water, resulting in precipitation of fluorite and a subsequent decrease in the
activity of F in the system, which in turn destabilized HFSE complexes and led to deposition of
HFSE-bearing minerals (Salvi & Williams-Jones, 2006). Primary HFSE minerals were replaced
by Ca-bearing equivalents (Salvi & Williams-Jones, 2006). It must be noted however that the
above interpretation of the origin of mineralization at Strange Lake runs counter to recent work
indicating the unlikelihood of transport of the HFSE by F (Williams-Jones et al., 2012). In light
of recent work, this model may be refined to transportation of the HFSE by Cl, and binding with
F to promote precipitation. Though they are not enriched in HFSE at exploitable concentrations,
similar models have been proposed for the origin of mineralization at Mont Saint Hilaire,
Quebec (Schilling et al., 2011), Tamazeght, Morocco (Salvi et al., 2000; Schilling et al., 2009),
Brockman, western Australia (Ramsden et al., 1993), and the Loch Loyal Syenite Complex,
northern Scotland (Walters et al., 2013).
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11
Partitioning behavior of the HFSE between silicate melt and
aqueous fluid
Investigations of HFSE and REE partitioning between fluid and melt as a function of fluid and
melt composition, temperature and pressure, are scarce in the literature. All available
experimental studies examined metaluminous to peraluminous melt compositions (Webster et
al., 1989; Bai & Koster van Groos, 1999; Reed et al., 2000; Bureau et al., 2003), except for
Borchert et al. (2010), who also considered peralkaline compositions. There have also been two
partitioning studies completed on natural samples of metaluminous rocks containing cogenetic
fluid and melt inclusions (Audetat & Pettke, 2003; Zajacz, 2008). Partitioning behaviour of the
HFSE is strongly dependent on both fluid and melt composition (Borchert et al., 2010), and for
this reason only peralkaline melt compositions will be discussed here.
Borchert et al. (2010) determined partition coefficients Dif/m
(Dif/m
= Cifluid
/Cimelt
) for Ba, La, Y,
and Yb, as a function of melt composition, represented by the Aluminum Saturation Index (ASI:
molar Al2O3/[Na2O + K2O + CaO]), salinity of the fluid (wt % NaCl), pressure and temperature,
by performing both quench experiments and hydrothermal diamond anvil cell experiments
coupled with synchrotron radiation X-ray fluorescence microanalysis of K-lines. The authors
chose Ba as representative for the large ion lithophile elements (LILE), La for the REE, Y for the
MREE or HFSE, and Yb to represent the HREE (Borchert et al., 2010). Their results showed that
all four elements partition into the melt for compositions within a range of ASI from 0.76 to 1.32
(Borchert et al., 2010). With increasing salinity of the aqueous fluid phase, Dif/m
does increase,
but remains well below 1 (Borchert et al., 2010). Experiments involving Cl-bearing fluids
revealed a melt-compositional dependence on the partitioning behaviour of the REE and Y,
where DREE,Yf/m
decreases with decreasing ASI (ie increasing peralkalinity) (Borchert et al.,
2010). Borchert et al., (2010) explain this finding in part as a function of the degree of
depolymerisation of the melt; peralkaline silicate melts contain more NBO-coordinated sites than
haplogranitic melts and can therefore incorporate more REE. Furthermore, when peralkaline
silicate melt is in the presence of a Cl-bearing aqueous fluid, the alkalis strongly partition into
the fluid, causing the ASI to increase initially (Borchert et al., 2010). Consequently, pH of the
fluid increases as the incorporation of alkalis lowers the activity of Cl-, causing H
+ to partition
into the melt and producing OH- in the fluid (Borchert et al., 2010). This behaviour is evidenced
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12
by the pH of peralkaline quench products of 9.5 (Borchert et al., 2010). The opposite behaviour
was observed in peraluminous compositions whose higher DREE,Yf/m
can be explained by a higher
activity of Cl- in the fluid, leading to more efficient chloride complexation of the REE (Borchert
et al., 2010). One very intriguing piece of evidence from this study is that in situ analyses show
nearly identical concentrations of La and Y in both Cl-bearing and Cl-free fluids, implying
complexation of the REE by dissolved silicate components (Borchert et al., 2010). The key
finding of Borchert et al. (2010) however is that exsolution of H2O±(Na,K)Cl fluids, even of
high salinity, from peralkaline melts does not deplete the melt in REE. This is to some degree
contradictory to observations of hydrothermal enrichment of the REE (and other HFSE) by
internally derived fluids of various peralkaline HFSE ore deposits (e.g., Thor Lake, Strange
Lake, Canada; see above). However, hydrothermal fluorite and F-bearing phases are abundant in
such deposits, whereas the systems investigated by Borchert et al. (2010) are F-free and therefore
may not be totally representative of natural peralkaline melt-fluid systems. Their findings do
suggest though that F- may indeed be a more important ligand for REE transportation in
hydrothermal fluids than suggested by Williams-Jones et al. (2012) (see above). Furthermore,
the high mobility of aqueous fluids might permit the spatially focussed deposition of
hydrothermal REE mineralization from fluids even if the fluids are not enriched in REE
compared to their parental silicate melts.
It should be anticipated that as strongly peralkaline melts approach the conditions of the second
critical endpoint, their properties increasingly approach those of the immiscible fluid phase, with
the result that partition coefficients for all elements will approach unity in the most peralkaline
compositions because at the second critical endpoint the two phases have converged to an
identical composition. Much remains to be learned about HFSE and REE partitioning in highly
peralkaline fluid-melt systems.
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13
Chapter 2
Research Paper Prepared for the Journal of Petrology
Abstract
The geochemistry of melt and fluid inclusions trapped in quartz that occurs as partial fillings of
miarolitic cavities in quartz syenite from Pico Alto, at Terceira, Azores was investigated in order
to provide further insight into late-stage evolution of peralkaline melts and the behaviour of high
field strength (HFSE) and rare earth elements (REE) in a system at the magmatic-hydrothermal
transition. Crystalline and hydrous melt inclusions analyzed by laser ablation-inductively
coupled plasma-mass spectrometry (LA-ICP-MS) show extreme magmatic enrichment of the
HFSE and REE, with up to 8.5, 3.4 and 4.4 wt % Zr, Nb and REE+Y. Fractionation was
associated with the crystallization of sanidine, resulting in progressive enrichment of the melt in
the HFSE, REE and volatiles and an increase in the agpaitic index to values as great as 6. Halite-
saturated fluid inclusions analyzed by LA-ICP-MS show lower total REE abundances than melts,
and a general enrichment in the HREE. Comparison of REE distribution patterns of melt
inclusions and miarolitic zircon and monazite suggest late-stage melt evolution by monazite,
then zircon and minor pyrochlore fractionation. Microthermometry of aqueous fluid inclusions
suggests maximum trapping conditions of 675°C, 120 MPa, similar to the greatest recorded
temperature of crystallization of quartz as determined by Ti geothermometry. The residual melt
of Terceira quartz syenite evolved to very volatile-rich compositions and initially exsolved a
hydrosaline melt (71.6 wt % NaCl eq) that was diluted to lower salinities by further exsolution of
aqueous fluid upon cooling. The appearance of quartz and aqueous fluid occurred at
approximately the same stage in the evolution of the system, at a temperature very close to the
pantellerite melt solidus. Preliminary calculated melt-fluid distribution coefficients suggest that
exsolution of hydrosaline melt does not significantly alter the HFSE and REE content of
peralkaline hydrous silicate melt. Had the residual melts represented as melt inclusions within
quartz been separately emplaced outside the they would be comparable in grade and composition
to ore in known HFSE deposits.
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14
Introduction
Peralkaline igneous rocks are important hosts for economically exploitable deposits of high field
strength elements (HFSE), including rare earth elements (REE) and in particular the heavy-REE
(HREE). For example, of the 41 known HFSE occurrences of economic interest in North
America, which are hosted in or sourced from 10 different geologic environments, seven
deposits occur in peralkaline granites, syenites or nepheline syenites (Mariano & Mariano,
2012). Because they lack a distinctive geophysical signature suitable for their detection by
remote sensing, a complete understanding of the genesis of such deposits is crucial for
identifying future exploration targets. Many aspects of the petrogenesis of HFSE deposits related
to peralkaline igneous rocks are still poorly understood. In particular, it remains unclear whether
such deposits are purely magmatic or if hydrothermal processes are integral to their formation.
Also enigmatic is the common enrichment in HREE over the light REE (LREE) in these deposits
despite the generally LREE-enriched nature of their host plutons (Boily & Williams-Jones, 1994;
Schmitt et al., 2002; Marks et al., 2003, Salvi & Williams-Jones, 2005; London, 2008; Sheard et
al., 2012).
High field strength elements are highly soluble in peralkaline silicate melts. It has been
experimentally shown that the solubility of baddeleyite increases with increasing peralkalinity in
silicate melts (Marr et al., 1998). Similarly, experimental work has demonstrated increased
solubility of zircon (Watson & Harrison, 1983), hafnon (Linnen & Keppler, 2002) and
manganotantalite in granitic melts with increasing peralkalinity (Van Lichterfelde et al., 2010).
Prolonged fractional crystallization of such a melt further concentrates the HFSE. Magmatic
enrichment of HFSE to near exploitable levels in residual silicate melts has been indicated by a
number of studies (Kovalenko et al., 1995; Chabiron et al., 2001; Schmitt et al., 2002; Thomas et
al., 2006; Andreeva & Kovalenko, 2011; Kynicky et al., 2011). It has also been well documented
that hydrothermal processes can concentrate HFSE through remobilization, transport and re-
deposition (Boily & Williams-Jones, 1994; Salvi & Williams-Jones, 1990, 1996). ZrO2 solubility
increases with decreasing temperature in HF-bearing (0.1 m) fluids (Migdisov et al., 2011).
Sufficient amounts of Zr can be transported to account for concentrations observed in natural, F-
rich hydrothermal systems where greisenizing fluids may contain up to 1.2 m HF (Shapovalov
& Setkova, 2012). Furthermore, with similar concentrations of HF in hydrothermal systems,
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15
significant quantities of REE can be dissolved (Migdisov et al., 2009). Recent numerical
modelling of fluid interaction with HFSE/REE-bearing rocks of the Strange Lake Complex by
Gysi & Williams-Jones (2013) suggests that the HFSE/REE can be hydrothermally transported at
HF concentrations as low as 0.4 m. It was also shown that the LREEs are more soluble in
hydrothermal fluids than the HREE (Migdisov et al., 2008; Migdisov et al., 2009), and the
predominance of REE-chloride complexes over those with fluoride in aqueous fluids at high
temperature (200°C) (Migdisov & Williams-Jones, 2002, 2007). Due to the common association
of hydrothermal fluorite with HFSE deposits, fluoride has long been considered as the most
likely transporting ligand of these elements (Wood, 1990; Haas et al., 1995). However, the
preceding studies demonstrate that this is improbable, and Williams-Jones et al. (2012) suggest
instead that F- acts as a binding ligand promoting precipitation and that the bulk of REE
transportation in hydrothermal fluids is achieved through complexation with chloride. Evidence
for hydrothermal mobilization and transport of HFSE has also been documented in natural
systems. For example, Salvi et al. (2000) report the occurrence of HFSE daughter phases in fluid
inclusions, and the mineralization at the colossal Bayan Obo REE deposit in China is considered
hydrothermal in origin (though intimately tied to the magmatic evolution of the carbonatite) (Lai
et al., 2012; Lai & Yang, 2013).
Enrichment in REE, U and Th at the Kvanefjeld deposit at Ilímaussaq is considered purely
magmatic in origin and occurred as a result of closed-system, protracted and uninterrupted
crystallization of peralkaline nepheline syenitic magma, pre-enriched (through fractional
crystallization of an alkali basaltic parent) in HFSE (Sorensen et al., 2006; Sorensen et al.,
2011). The giant Nb+REE deposits at Lovozero, Kola Peninsula, Russia are also believed to be
purely magmatic in origin (Kogarko et al., 2002). Replacement of loparite cumulates by complex
Nb and REE minerals at Lovozero is attributed to reaction with the residual melt rather than
hydrothermal fluids (Kogarko et al., 2002).
Most peralkaline igneous HFSE deposits are believed to have been generated by a combination
of magmatic and hydrothermal processes. At the Thor Lake rare metal deposit, NWT, for
example, the main rare metal ore mineral is zircon of both magmatic and hydrothermal origin
(Sheard et al., 2012). Repeated injections of aegirine-nepheline syenite, fractional crystallization
and convection produced cumulate layers of magmatic eudialyte, then zircon (Sheard et al.,
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16
2012). Late-stage exsolution of orthomagmatic aqueous fluid from the syenite then remobilized
Zr and the HREE to form pseudomorphic zircon after eudialyte, secondary, colloform zircon
rims on pre-existing grains and other REE-fluorcarbonates (Sheard et al., 2012). The LREE were
also remobilized, but deposited more distally as a result of later mixing with an externally
derived, Ca-bearing fluid (Sheard et al., 2012). Similarly, the proposed deposit model for Strange
Lake, Labrador, involves magmatic enrichment in HFSE via fractional crystallization of
peralkaline granite followed by exsolution of a highly saline, F-rich orthomagmatic fluid (Salvi
& Williams-Jones, 2006). Interaction of this fluid with peralkaline granite allowed leaching of
HFSE from the primary mineral assemblage, complexation of the HFSE with the ligands F- and
Cl- and their subsequent transport as complexes (Salvi & Williams-Jones, 2006). The
orthomagmatic aqueous fluid then mixed with Ca-bearing meteoric water, resulting in
precipitation of fluorite and a subsequent decrease in the activity of F in the system, which in
turn destabilized HFSE complexes and led to deposition of HFSE-bearing minerals (Salvi &
Williams-Jones, 2006). Primary HFSE minerals were replaced by Ca-bearing equivalents (Salvi
& Williams-Jones, 2006). It must be noted however that the above interpretation of the origin of
mineralization at Strange Lake runs counter to recent work indicating the unlikelihood of
transport of the HFSE by F (Williams-Jones et al., 2012). In light of recent work, this model may
be refined to transportation of the HFSE by Cl, and binding with F to promote precipitation.
Though they are not enriched in HFSE at exploitable concentrations, similar models have been
proposed for the origin of mineralization at Mont Saint Hilaire, Quebec (Schilling et al., 2011),
Tamazeght, Morocco (Salvi et al., 2000; Schilling et al., 2009), Brockman, western Australia
(Ramsden et al., 1993), and the Loch Loyal Syenite Complex, northern Scotland (Walters et al.,
2013).
Mungall & Martin (1996) examined glassy pantellerite and comagmatic holocrystalline syenite
from Pico Alto at Terceira, Azores and found essentially identical lithophile element abundances
in both rocks, indicating evolution to equal degrees of HFSE enrichment and supporting the view
that the syenites are holocrystalline equivalents of the pantellerites. However, the syenites are
depleted in U, Y, Hf and REE relative to pantellerite, demonstrating that some HFSE were
mobile during late-stage crystallization (Mungall & Martin, 1996). Comparison of unaltered
granite with hydrothermally altered granite at Strange Lake, Labrador reveals a complementary
pattern wherein these same mobile elements are enriched in the ore zone by hydrothermal
-
17
processes (Mungall & Martin, 1996). Mungall & Martin (1996) advanced a genetic model for
HFSE deposits involving fractional crystallization of peralkaline melt followed by exsolution of
a saline orthomagmatic fluid with high concentrations of dissolved HFSE (and possibly silicate
minerals), with eventual deposition of HFSE minerals promoted by a decrease in temperature or
by mixing with an externally derived fluid.
Cann (1967) examined two syenite xenoliths with different mineralogy, carried to surface in
trachytic pumice at Agua de Pao, Sao Miguel, Azores and considered them to be the slowly
cooled equivalents of the most evolved lavas on the volcanic island. Despite a very similar major
element composition, the trace element composition of these two blocks was very different
(Cann, 1967). Cann (1967) concluded that the two types of syenite xenoliths were related by
fractional crystallization, and proposed crystallization at a eutectoid point for the major element
composition of the solids to match that of the liquid. Widom et al. (1993) revisited these same
samples and noted that their composition was similar to that of the most differentiated trachytes
at Agua de Pao, but with some notable differences including a higher concentration of SiO2 and
the incompatible elements and depletion in H2O, Rb, Sr, Pb and U relative to trachyte. These
authors concluded that the xenoliths represent the bulk liquid composition at Agua de Pau, but
that they must have been pervasively altered by a quartz-saturated aqueous fluid, consistent with
some samples having pore spaces completely filled with quartz (Widom et al., 1993). When the
aqueous fluid eventually percolated away, it would have depleted the xenoliths in alkalis and U
(Widom et al., 1993). Ridolfi et al. (2003) also studied syenite xenoliths from Agua de Pao and
similarly concluded that the quartz-saturated syenites represent the plutonic equivalents of
evolved trachytes. Trachyte at Agua de Pao is strongly depleted in Ba, Eu and Sr and enriched in
incompatible elements (Ridolfi et al., 2003). In trace element variation diagrams, quartz-
saturated syenites overlap with the evolved trachyte, though some samples of the plutonic ejecta
are more highly enriched in Zr, Nb, and Th than their volcanic equivalent (Ridolfi et al., 2003).
Similar work has been done on peralkaline xenoliths from Ascension Island and Tenerife,
Canary Islands. Roedder & Coombs (1967) examined peralkaline granite xenoliths entrained in
trachyte at Ascension Island and concluded that the xenoliths represented the equivalent of
trachyte that was cooled along the inner walls of the volcanic conduit, subsequently released by
later eruptions. Webster & Rebbert (2001) completed a melt inclusion study on quartz-hosted
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18
glass inclusions in the same xenoliths from Ascension and found the inclusion compositions to
be very similar to those of whole rock data for both xenoliths and trachyte. However, melt
inclusions are distinctly enriched in Na2O, F, Cl and H2O relative to granite or trachyte (Webster
& Rebbert, 2001). Additionally, Wolff & Toney (1993) analyzed interstitial glass in a nepheline-
syenite xenolith entrained in phonolitic pumice (which they consider the volcanic equivalent of
the xenoliths) at Tenerife, Canary Islands. Interstitial glass in nepheline-syenite is enriched in Zr
by one order of magnitude relative to the most Zr-rich pumice (Wolff & Toney, 1993). Other
incompatible elements should be accordingly enriched in the glass relative to pumice, but REE,
Y and Th were below detection limits in the late-stage liquid (Wolff & Toney, 1993). Wolff &
Toney (1993) attribute the difference in trace element composition between glass, xenolith and
pumice to the occurrence in nepheline-syenite of titanite and loparite; minerals that are
significant hosts for the REE, Y, Nb and Th, and whose crystallization depletes the residual
liquid in these elements. Also relevant to the present study is the work of Ferguson (1978), who
compared the mineralogy of a nepheline-syenite xenolith to that of its entraining ignimbrite. This
author documented needles of titanian-aegirine as well as apatite, låvenite, eucolite (eudialyte
group mineral) and sodalite projecting from the walls of cavities in the nepheline-syenite and
concluded that these phases crystallized from brine or vapour (Ferguson, 1978). Ignimbrite at
Tenerife does not contain these HFSE-bearing minerals and Ferguson (1978) suggested that the
difference in mineralogy between plutonic and volcanic equivalent rocks to be controlled by
volatile content and pressure.
In this contribution, we examine the compositions of melt and fluid inclusions trapped in quartz
that occurs as partial fillings of miarolitic cavities in quartz syenite from Pico Alto, at Terceira,
Azores, and indicate extreme magmatic enrichment of HFSE as well as preferential extraction of
the HREE into the exsolving aqueous fluid.
Geology of Terceira
Terceira is one of nine volcanic islands that comprise the Azores archipelago in the North
Atlantic (37 - 40°N), approximately 1300 km from the Portuguese mainland. The Azores
straddle the mid-Atlantic Ridge (MAR) and lie at the triple junction of the American, Eurasian
and African plates (Madureira et al., 2011). Terceira lies east of the MAR on the Terceira Rift, a
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19
very slowly spreading plate boundary (estimated ~4 mm/year) separating the Eurasian and
African plates (Madureira et al., 2011). The Azores are commonly divided into three groups:
Western, Central and Eastern islands (Self, 1973). Terceira is considered a member of the
Central group, along with Faial, Pico, Sao Jorge and Graciosa (Self, 1973).
Terceira is the third largest island in the Azores, measuring 28 by 18 km and covering an area of
401 km2 (Self, 1973). It is composed of 4 overlapping stratovolcanoes. From east to west these
are Cincos Picos, Pico Alto, Guilherme Moniz and Santa Barbara (Self, 1973; Mungall &
Martin, 1995). Calvert et al. (2006) considered the Pico Alto volcanic centre to be a portion of
the active flank of the otherwise inactive Guilherme Moniz volcano immediately to the south.
Trachytic pyroclastics of the Lajes Ignimbrite, which were erupted from the Pico Alto caldera ca
21 ka, cover the eastern two thirds of the island (Self, 1973). The volcanic centers decrease in
age from east to west, though adjacent volcanoes overlap in age (Calvert et al., 2006). Basaltic
vents are widely distributed across Terceira, yet are heavily concentrated along a rift running
from a point between Guilherme Moniz and Pico Alto in the centre of the island to Santa Barbara
in the west (Mungall & Martin, 1995; Calvert et al., 2006; Madureira et al., 2011). Cincos Picos
and Guilherme Moniz are extinct, whereas the Pico Alto Volcanic Center (north flank of
Guilherme Moniz) and Santa Barbara remain active (Calvert et al., 2006).
Two distinctly different magmatic trends are observed in the recent volcanic suites. The Santa
Barbara suite includes off-rift basalt through to comendite, whereas pantelleritic lavas erupted at
Pico Alto represent the felsic termination of an evolutionary sequence from rift basalt to
pantellerite (Mungall & Martin, 1995). Mungall & Martin (1995) suggested that the differences
between these trends stemmed from a number of factors including their derivation from distinct
mantle sources, evolution at different pressures, and more complete degassing of volatiles from
the Santa Barbara trachytes compared with those from Pico Alto. Oxygen fugacity must have
been low throughout the evolution of the Pico Alto suite to produce such Fe enrichment and a
pantelleritic trend rather than a comenditic one (Mungall & Martin, 1995). Furthermore, Mungall
& Martin (1995) determined by petrogenetic modelling that the Pico Alto rift basalt evolved at
low pressure and underwent little differentiation before erupting, whereas at Santa Barbara the
off-rift basalts underwent significant fractionation at high pressure prior to eruption.
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20
The Pico Alto magmatic center is a complex of pantellerite domes and flows occupying and
overflowing an older caldera collapse structure related to the eruption of the Lajes ignimbrite
(Mungall & Martin, 1996). The evolutionary trend in the Pico Alto suite follows through from
olivine-augite basalt, hawaiite, mugearite, benmoreite, trachyte to pantellerite (Self, 1973).
Mungall (1993) modelled the liquid line of descent for the suite and found that Si and Al remain
at relatively constant concentrations while Ti is initially enriched as plagioclase, lesser
clinopyroxene and olivine are removed from the parental transitionally alkaline basalt until Fe-
Ti-oxides appear on the liquidus resulting in a decrease in Ti and increase in Si. The Pico Alto
suite can be continuously represented by this liquid line of descent up to the trachytes by
approximately 75% fractional crystallization (Mungall, 1993). Peralkaline compositions are
produced in the model by early crystallization of amphibole from benmoreite (and a
corresponding decrease in Al) at higher pressure to form mildly peralkaline trachyte (Mungall,
1993). Similarly, experiments of Nekvasil et al (2004) produced peralkaline residual liquids by
kaersutite fractionation. Removal of an anhydrous mineral assemblage matching observed
phenocryst phases from early trachyte successfully models the most evolved pantellerite at Pico
Alto (Mungall, 1993). The trachytic magma fractionates only a small proportion of ilmenite
(0.56 – 1.21 %), resulting in enrichment of Fe and Mn (Mungall, 1993). The most evolved
peralkaline melt composition observed at Pico Alto is in an olivine-hosted melt inclusion
containing 69.0 wt % SiO2, 12.6 wt % Fe2O3 and 5.5 wt % Al2O3, with an agpaitic index of ≥ 2.4
(Mungall, 1993). Mungall & Martin (1995) showed through their least-squares mixing model of
fractional crystallization that the most evolved melt inclusion composition at Pico Alto could be
produced from the least evolved melt inclusion composition by 74% fractional crystallization of
ilmenite, fayalite, sodic hedenbergite and sanidine. These authors also documented magmatic
enrichment of high field strength elements (HFSE) over 30-fold by fractional crystallization in
the Pico Alto suite (Mungall & Martin, 1995).
The Lajes ignimbrite carries xenoliths of quartz syenite that are considered comagmatic and
essentially equivalent in composition with the most evolved pantellerite lavas (Mungall &
Martin, 1996). Pico Alto pantellerite typically consists of a microlite-rich, glassy matrix
containing several modal percent of phenocrysts of sanidine, aegirine-augite, fayalite, ilmenite,
apatite, and rare pyrite, aenigmatite and amphibole (Mungall & Martin, 1996). Quartz syenite
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21
xenoliths are fully crystalline and composed mainly (65 – 95 modal %) of alkali feldspar, with
minor aegirine-augite, fayalite and albite, and trace apatite and ilmenite. Feldspar laths form an
interlocking framework that encloses a significant network of miarolitic cavities (≤ 10 vol %)
that commonly contain or are filled by euhedral, zoned quartz. Aside from quartz, a variety of
minerals decorate and project into cavities, including numerous HFSE-bearing phases. These
high field strength element (HFSE)-rich or rare earth element (REE)-rich minerals were
interpreted by Mungall & Martin (1996) to have precipitated from a fluid as indicated by their
habits and their projection from pore walls. Secondary alteration of the quartz syenites includes
oxidation and replacement of ferromagnesian minerals and replacement of sanidine by perthite.
Materials and Methods
Sample Collection and Petrography
Fourteen quartz-syenite xenoliths were collected from the beach at Baia de Calderinha, Terceira,
Azores in June 2012. At this location, xenoliths are readily weathered out of thick, friable
ignimbrite deposits. One sample (CB1205), was collected from a subterranean basaltic lava tube,
now collapsed and inaccessible, by workers at the Gruta do Natal, Terceira, Azores. To facilitate
cutting, given the friability of the highly porous quartz syenite samples, the rocks were vacuum
impregnated with epoxy. Once impregnated, samples were cut into a first set of duplicate
polished thin sections (30 μm) and doubly polished thin sections (100 μm). Additional doubly
polished thin sections were later cut for particularly interesting samples. All thin section
preparation was done at Queen’s University, Kingston, Ontario.
Thin sections were examined by transmitted light microscopy to select quartz grains hosting
fluid and/or melt inclusions. Grains hosting both types of inclusions were preferred, but those
with only one type were also included. Each grain selected for study was photographed using an
Olympus BX51 transmitted/reflected light microscope equipped with an Olympus Q-color 3
RTV camera. Using a JEOL-6610LV SEM operating at 20 kV and 8.3 nA, with an Oxford
Instruments x-ray detector with an area of 20 mm2 and manufacturer-provided software, a full
back scattered electron (BSE) image archive of each individual grain was compiled. Energy
dispersive spectroscopy (EDS) was used to identify mineral phases in those polyphase melt
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22
inclusions which were observed to lie at the surface and visible in BSE, as well as rare metal
mineral phases in cavities. Additionally, a Gatan MiniCL cathodoluminescence imaging system
was used to identify growth zones in quartz grains.
Microthermometry
Once targeted and imaged, quartz grains were cut in water from thin sections using a small hand-
held diamond saw. To dissolve epoxy from pore spaces and to remove the selected quartz chips
from the glass thin section substrates they were then soaked for hours to days in either methylene
chloride or acetone (depending on the adhesive used for vacuum impregnation).
Microthermometry of fluid inclusions was performed on a microscope-mounted Linkham
THMS600 heating-cooling stage, calibrated using synthetic CO2 and aqueous fluid inclusions
with known phase transitions at -56.6°C, 0.0°C and 374.1°C. Measurements were reproducible to
± 1.5 °C (1 sigma standard deviation) for temperatures below zero, and to ± 7.7°C for higher
temperatures. Salinities, estimates of minimum trapping temperature (TT) and associated
isochores for all inclusions were calculated using the program SoWat, which employs the
equation of state formulae of Driesner (2007) and Driesner & Heinrich (2007).
LA-ICPMS
Element concentrations in fluid and melt inclusions were determined by laser ablation ICPMS at
the University of Toronto. The analytical system uses a Newwave UP-213 Nd-YAG laser
operating at 213 nm coupled to a VG-PQExcell ICP-MS with He gas flushing the ablation cell.
Software supplied by the manufacturer was used to acquire time-resolved count data. NIST 610
standard glass was used as the external calibration standard. Analyses were performed in blocks
of 20, with the first and last two done on the NIST 610 standard. The following element masses
were analyzed (dwell times in ms) in the inclusions: 7Li (30),
11B (30),
23Na (10),
25Mg (10),
27Al
(10), 29
Si (10), 39
K (10), 44
Ca (10), 49
Ti (10), 51
V (10), 53
Cr (10), 55
Mn (10), 57
Fe (10), 59
Co (10),
62Ni (10),
65Cu (30),
, 66Zn (10),
75As (10),
77Ar Cl (10),
82Se (10),
83Kr (10),
85Rb (10),
88Sr (10),
89Y(10),
90Zr (10),
93Nb (10),
95Mo (30),
107Ag (10),
118Sn (10),
121Sb (10),
133Cs (50),
137Ba (10),
139La (10),
140Ce (10),
141Pr (10),
146Nd (10),
147Sm (10),
151Eu (10),
157Gd (10),
159Tb (10),
163Dy
(10), 165
Ho (20), 166
Er (20), 169
Tm (20), 173
Yb (20), 175
Lu (20), 178
Hf (30), 181
Ta (30), 182
W (10),
208Pb (10),
232Th (10),
238U (10).
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23
Melt and fluid inclusion element concentrations were quantified from raw signals using the
software SILLS (Guillong et al., 2008). This involved deconvolution for mixed melt or fluid
inclusion+host and host-only signals after calculation of background-corrected count rates for
each isotope, and quantification of inclusion and host composition. An example time-resolved
spectrum for LA-ICPMS analysis of a silicate melt inclusion is given in Figure 2.1a.
For melt inclusions, internal standardization was performed using the average K2O and SiO2
content of melt inclusions determined independently by Mungall & Martin (1996). This method
was determined by Gray et al., (2011) to yield accurate melt inclusion analyses for trace
elements in Si-rich inclusions hosted in quartz, based on comparison of LA-ICPMS data and data
obtained independently by EMP. Inter-element ratios determined for individual melt inclusions
by this method are precisely constrained, but the absolute concentrations may be in error by
several wt% relative, and nominal major element totals of the melt inclusions vary between
approximately 80 and 120 wt%. To check the validity of this approach, LA-ICPMS analyses of a
glass (sample P16), reported by Mungall & Martin (1996), were obtained and quantified with
generally good agreement with published data for the majority of trace elements and major
elements with significant exceptions being Al2O3 and Sr (underreported by Mungall & Martin,
1996) and Ba and U (over-reported by Mungall & Martin, 1996). Results of Mungall & Martin
(1996) are given in Table 1, with oxides reported in wt % and trace elements in ppm. Since melt
inclusions are polyphase (i.e., they contain several component minerals that grew during cooling
after entrapment), complete inclusions buried in the host phase had to be analyzed to ensure that
the bulk analyses did not exclude phases that had been polished away during sample preparation.
Estimated relative uncertainties for melt inclusion analyses are generally better than 5% with the
exception of a few elements such as Ba and Sm that show uncertainties as high as 25% relative.
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24
Fig. 2.1. Example of time-resolved spectrum for LA-ICPMS analysis of (a) a silicate melt
inclusion and (b) a halite-saturated fluid inclusion.
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25
Table 1. Comparison of Analytical Results of Mungall & Martin (1996) and the Present Study on
Sample P16
Mungall & Martin,
1996
Present Study
SiO2 66.78 66.78
TiO2 0.53 0.44
Al2O3 10.55 13.37
Fe2O3 2.57 -
FeO 5.8 6.1
MnO 0.34 0.27
MgO 0.08 0.15
CaO 0.51 0.48
Na2O 7.3 7.8
K2O 4.45 5.24
P2O5 0.04 -
Ba 201 69
Nb 304 326
Zr 1742 1523
Y 152 136
Sr 2 5
Rb 187 198
Th 29 26
U 19 9
La 163 183
Ce 296.7 368.1
Pr 33.6 39.5
Nd 120.9 144.8
Sm 22.7 28.8
Eu 2.8 3.6
Gd 21 26
Tb 3.2 4.3
Dy 21.5 26.5
Ho 4.1 5.2
Er 12.1 15.1
Tm 1.7 2.1
Yb 10.5 14.1
Lu 1.6 2.0
Hf 23.4 36.2
Ta 15.5 18.8
For fluid inclusions, internal standardization was performed using Na concentrations assumed to
be equal to the average bulk salinity (wt % NaCl eq.) determined from microthermometry.
Owing to large ranges in salinity for melt inclusion assemblages, fluid inclusion trace element
concentrations should be considered to be semi-quantitative only and carry relative uncertainties
(corresponding to uncertainties in bulk salinity) of 20-30%. Despite the large uncertainty in
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26
absolute concentration, it should be noted that the relative concentrations of elements in each
analysis are much better constrained, so that element ratios can be compared with some
confidence. An example time-resolved spectrum of LA-ICPMS analysis of a halite-saturated
fluid inclusion is given in Figure 2.1b.
Electron Microprobe
Selected monazite and zircon grains in cavities were analyzed by a Cameca SX-50/51 (DCI 1300
DLL) electron microprobe (EMP) at the University of Toronto, equipped with 3 tunable
wavelength dispersive spectrometers. Operating conditions were 40° takeoff angle and a beam
energy of 20 keV. A beam size of 1 μm was used. The same times were used for on and off-peak
counting. Other operating conditions for monazite and zircon EMP analyses are given in Tables
2a,b, respectively. All standardization was performed under the exact same conditions as
analysis of unknowns. Oxygen was calculated by stoichiometry.
Table 2a: EMP operating conditions for monazite analyses
Analyzer
Crystal
Counting
Time (s)
Off-peak
Correction
Method
Standards
Used
P (ka) PET 20 Linear CePO4REE/6
Y (la) PET 20 Average Y2O3sx2
La (la) LiF 20 Linear LaPO4REE/6
Ce (la) PET 20 Average CePO4REE/6
Pr (lb) LiF 40 Linear PrPO4REE/6
Nd (la) LiF 40 High Only NdPO4REE/6
Sm (lb) LiF 40 Linear SmPO4REE/6
Ho (la) LiF 40 Average HoPO4REE/6
Th (ma) PET 40 High Only ThSiO4sx2
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27
Table 2b: EMP operating conditions for zircon analyses
Analyzer
Crystal
Counting
Time (s)
Off Peak
Counting
Time (s)
Off-peak
Correction
Method
Standards Used
F (ka) TAP 10 10 Average CaF2sx2
Al (ka) TAP 40 40 Average pxTiAlsx1
Si (ka) TAP 20 20 Slope (Hi) ZrSiO4sx2
P (ka) PET 20 30 Linear YPO4_REESX6/6
Ti (ka) LiF 20 30 Linear TiO2sx2
Fe (ka) LiF 20 20 Linear pyropKsx1
Y (la) PET 30 100 Average YPO4_REESX6/6
Zr (la) PET 20 20 Low Only ZrSiO4sx2
Nb (la) PET 100 100 High Only NaNbO3sx1
Hf (la) LiF 60 60 Average HfSiO4sx2
Ce (la) LiF 100 40 Linear CePO4REE/6
Nd (la) LiF 150 60 Linear NdPO4REE/6
Gd (la) LiF 200 80 Linear GdPO4REE/6
Dy (la) LiF 100 40 Linear DyPO4REE/6
Er (la) LiF 100 40 Average ErPO4REE/6
Yb (la) LiF 200 80 Linear YbPO4REE/5
Petrographic Observations
Quartz syenite
The paragenetic sequence for Pico Alto quartz syenite xenoliths has been established by
petrography employing transmitted and reflected light microscopy as well as detailed
observation by SEM (Fig 2.2). Four stages of crystallization are recognized in the textural
evolution of the rocks. Earliest crystallizing magmatic phases, which appear as phenocrysts in
the pantellerite lava and are also present in the xenoliths, are fayalite, ilmenite, sanidine,
aegirine-augite, aenigmatite, apatite and rare pyrite. Aenigmatite commonly surrounds fayalite.
Aegirine-augite grains are often patchy in BSE, showing brighter domains where the
composition nearer to end-member aegirine. The xenoliths display an essentially diabasic texture
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28
dominated by an interlocking framework of former sanidine crystals enclosing volumetrically
minor amounts of the other phases. Continuing growth of sanidine after grain-grain
impingement (referred to here as main stage crystallization) led to the partial or complete
overgrowth of other early minerals within the much larger feldspar crystals during this main
growth stage. The liquid residual to the main stage of solidification occupied a network of
polyhedral pores interstitial to the sanidine crystal framework. This residual liquid was highly
enriched in volatile components and HFSE, leading to the deposition of a variety of minerals
within the pore space at the late stage of solidification, beginning with pyrochlore, followed by
quartz, then other HFSE-bearing minerals including most commonly zircon, end-member
aegirine, monazite, fersmite and britholite, and more rarely aeschynite group minerals,
baddeleyite, bastnäesite, chevnikite-(Ce), elpidite, eudialyte, kainosite, låvenite, nacareniobsite-
(Ce), nafertisite, narsarsukite, parisite-(Ce), REE-oxyfluoride, samarskite-(Y), synchisite,
vlasovite and xenotime. Earlier zircon is euhedral whereas later grains facing into cavities are
botryoidal. A list of the HFSE-bearing minerals and their formulae is given in Table 3 (in most
cases the rare minerals were identified solely based on their compositions). During this late-stage
interstitial crystallization, limpid overgrowths of a subsolvus assemblage of near end-member
albite and K-rich sanidine were formed facing into the cavities, locally enclosing rare minerals
that had previously nucleated on pore walls (Fig 2.3b). Spectacular textures are present in the
miaroles (Fig. 2.3), for example, botryoidal zircon and aeschynite, radiating masses of fayalite
and fibrous aegirine, pyrochlore and bastnäesite. In one instance, the edges of botryoidal zircon
are decorated with fine grained britholite. Colloform quartz is commonly the latest phase (later
alteration aside) in miaroles, decorating earlier deposited minerals (Fig. 2.3). Titanium
abundances in different CL zones of selected quartz grains were measured by LA-ICPMS and
range from 5.54 to 256.92 ppm. Full results are given in Table 4. A final stage in the textural
evolution, referred to here as the alteration and veining stage, led to subsolidus re-equilibration
of most primary magmatic phases to a new mineral assemblage, both pervasively and along
distinctly recognizable fractures, apparently simultaneously with the final growth of new phases
within the interstitial pore spaces. Pervasive low temperature alteration in the quartz syenites is
seen as the replacement of early sanidine by patchy perthitic intergrowths and the breakdown of
ferromagnesian minerals to hydrous minerals clinoptilolite-(K) and tuperssuatsaite, possible
cronstedtite or greenalite, ferro-richterite, potassic-magnesiohastingsite, and iron
oxides/hydroxides goethite, ferrihydrite, hematite, and other unidentified iron oxide/hydroxide
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29
phases. Early ilmenite commonly contains oriented lamellae of hematite. Not only confined to
the cavities, late-stage HFSE-bearing minerals are also commonly seen filling fractures.
Fig. 2.2. Bar diagram showing the general paragenetic relations of minerals and
evolutionary stages of crystallization in the quartz-syenite xenoliths. Dashed bars indicate
that the mineral is not observed in all samples. Circles indicate minerals that make up the
phenocryst assemblage in pantellerite. *Limonite used as a general term for unidentified
iron oxides, hydroxides.
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30
Table 3 . HFSE-bearing minerals found in miarolitic cavities
Mineral Formula Abbreviation
Used
Aenigmatite Na2Fe2+
5Ti(Si6O18)O2 aen
Aeschynite-(Ce) (Nd,Ce,Ca)(Ti,Nb)2(O,OH)6
Aeschynite-(Y) (Y,Ca,Fe,Th)(Ti,Nb)2(O,OH)6
Baddeleyite ZrO2
Bastnaesite (La,Ce,Nd,Y)(CO3)F
Britholite (Ce,Y,Ca)5(SiO4,PO4)3(OH,F)
Chevkinite-(Ce) (Ce,La,Ca,Th)4(Fe2+,Mg)(Fe2+,Ti,Fe3+)2(Ti,Fe
3+)2(Si2O7)2O8
Dalyite K2ZrSi6O15
Elpidite Na2ZrSi6O15 ∙ 3H2O
Eudialyte Na15Ca6(Fe2+,Mn2+)3Zr3(Si25O73)(O,OH,H2O)3(OH,Cl)2 eud
Fersmite (Ca,Ce,Na)(Nb,Ta,Ti)2(O,OH,F)6 fer
Ilmenite FeTiO3 ilm
Kainosite Ca2(Ce,Y)2Si4O12CO3∙H2O
Låvenite (Na,Ca)2(Mn2+,Fe2+)(Zr,Ti)(Si2O7)(O,OH,F)2
Monazite LREEPO4 mnz
Nacareniobsite-(Ce) Na3Ca3(Ce,La,Nd)Nb(Si2O7)OF3 nac
Nafertisite Na3(Fe2+,Fe3+)6(Ti2Si12O34)(O,OH)7∙2H2O
Narsarsukite Na4(Ti,Fe)2(Si8O20)(O,OH,F)2
Parisite-(Ce) Ca(Ce,La)2(CO3)3F2
REE-oxyfluoride REEOF
Pyrochlore (Na,Ca)2Nb2O6(OH,F)
Samarskite-(Y) Y(Fe2+,Fe3+,U,Th,Ca)2(Nb,Ta)2O8
Synchisite Ca(Ce,Nd,Y)(CO3)2F
Vlasovite Na2ZrSi4O11
Xenotime YPO4
Zircon ZrSiO4 zrn
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31
Fig. 2.3. Back-scattered electron (BSE) images of various miaroles in quartz syenite.
Mineral abbreviations from Kretz (1983) where possible. (a) Limpid sanidine overgrowths
on perthite are the earliest phase in this miarole, partially enclosing nacareniobsite-(Ce),
botryoidal zircon and fersmite. Later, finer grained monazite and botryoidal zircon are
decorated by colloform quartz. A ferromagnesian mineral has been altered to unidentified
iron oxide/hydroxides (lim). (b) Limpid albite overgrowth on perthite is seen enclosing
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32
zircon and monazite; fibrous, end-member aegirine occupies the majority of the pore, and
late colloform quartz decorates earlier minerals. (c) Early, euhedral aegirine precipitated
early in the cavity and was subsequently decorated by botryoidal zircon and fine grained,
euhedral fersmite. Colloform quartz commonly decorates zircon, though the reverse
pattern is also seen here. (d) Enlargement of c), delineated by the red box..Euhedral
fersmite, botryoidal zircon and alkali feldspar are decorated by later colloform quartz. A
ferromagnesian mineral has been altered to unidentified iron oxides/hydroxides (lim). (e)
Fibrous aegirine projecting into pore followed by sanidine, ilmenite, aenigmatite, eudialyte
and quartz. (f) Aegirine projecting into pore and coated with colloform quartz.
Table 4. Ti abundances in different zones (seen in CL) of selected quartz grains and calculated
temperatures of crystallization at various pressures P=134 bars P=1202 bars P=2000 bars
Analysis Ti
49
(ppm) T (°C) T (°C) T (°C)
12060-9-12 45.36 463.23 463.25 463.27
12060-9-22 13.20 381.98 382.00 382.01
12060-9-32 52.45 474.14 474.16 474.18
12060-9-42 29.56 432.86 432.89 432.90
12091-3h-12 51.78 473.15 473.17 473.19
12091-3h-22 166.44 573.81 573.84 573.85
12091-3h-32 84.30 512.11 512.13 512.15
12150-4a-12 5.54 334.83 334.84 334.86
12150-4a-22 89.77 517.44 517.46 517.48
12150-4a-32 19.30 405.02 405.04 405.05
12150-4b-12 7.86 353.01 353.03 353.04
12150-4b-22 5.80 337.12 337.13 337.15
12150-4b-32 15.80 392.71 392.73 392.75
12094-5b-12 256.92 618.52 618.55 618.57
12094-5b-22 164.35 572.58 572.61 572.62
12094-5c-12 192.96 588.53 588.55 588.57
12094-5c-32 41.45 456.62 456.64 456.65
12094-3b-12 78.81 506.48 506.51 506.53
12094-3b-22 70.66 497.54 497.57 497.59
12160-6d-12 22.05 413.49 413.51 413.53
12160-6d-22 129.59 550.01 550.04 550.06
12160-6d-32 8.70 358.48 358.50 358.51
12160-6d-42 19.53 405.78 405.80 405.82
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33
Mineral Compositions
Results of EMP analys