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

ii

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.

iii

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

v

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

vi

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

vii

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

viii

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

ix

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

1

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.

2

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

3

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

4

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

5

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.

6

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).

7

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

8

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).

9

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).

10

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).

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

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.

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.

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,

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.,

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

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

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.

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

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

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).

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.

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.

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

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

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

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

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.

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

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

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

33

Mineral Compositions

Results of EMP analys