Article
Reference
Tourmaline as a tracer of late-magmatic to hydrothermal fluid evolution: The world-class San Rafael tin (-copper) deposit, Peru
HARLAUX, Matthieu, et al.
Abstract
The world-class San Rafael tin (-copper) deposit (central Andean tin belt, southeast Peru) is an exceptionally large and rich (>1 million metric tons Sn; grades typically >2% Sn) cassiterite-bearing hydrothermal vein system hosted by a late Oligocene (ca. 24 Ma) peraluminous K-feldspar-megacrystic granitic complex and surrounding Ordovician shales affected by deformation and low-grade metamorphism. The mineralization consists of NWtrending, quartz-cassiterite-sulfide veins and fault-controlled breccia bodies (>1.4 km in vertical and horizontal extension). They show volumetrically important tourmaline alteration that principally formed prior to the main ore stage, similar to other granite-related Sn deposits worldwide. We present here a detailed textural and geochemical study of tourmaline, aiming to trace fluid evolution of the San Rafael magmatic-hydrothermal system that led to the deposition of tin mineralization. Based on previous works and new petrographic observations, three main generations of tourmaline of both magmatic and hydrothermal origin were distinguished and were analyzed in situ for their major, minor, and [...]
HARLAUX, Matthieu, et al. Tourmaline as a tracer of late-magmatic to hydrothermal fluid evolution: The world-class San Rafael tin (-copper) deposit, Peru. Economic Geology, 2020, vol. 115, no. 8, p. 1665-1697
DOI : 10.5382/econgeo.4762
Available at:
http://archive-ouverte.unige.ch/unige:148392
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Tourmaline as a tracer of late-magmatic to hydrothermal fluid evolution:
1
The world-class San Rafael tin (-copper) deposit, Peru 2
3
Matthieu Harlaux1,2*, Kalin Kouzmanov1, Stefano Gialli1, Oscar Laurent3, Andrea Rielli4, 4
Andrea Dini4, Alain Chauvet5, Andrew Menzies6, Miroslav Kalinaj7, and Lluís Fontboté1 5
6 7
1 Department of Earth Sciences, University of Geneva, 1205 Geneva, Switzerland 8
2 Present address: Nevada Bureau of Mines and Geology, University of Nevada, Reno, NV 9
89557-0178, USA 10
3 Institute of Geochemistry and Petrology, ETH Zürich, 8092 Zürich, Switzerland 11
4 Istituto di Geoscienze e Georisorse, CNR, 56124 Pisa, Italy 12
5 Géosciences Montpellier, CNRS-UMR 5243, Université de Montpellier, 34095 Montpellier, 13
France 14
6 Bruker Nano GmbH, Am Studio 2D, 12489 Berlin, Germany 15
7 Minsur S.A., Jr. Lorenzo Bernini 149, San Borja, Lima 27, Peru 16
17
* Corresponding author: mharlaux@unr.edu 18
19
v. 115, no. 8, p. 1665-1697. (Changes introduced at the proof stage not included in this version)
Abstract 20
21
The world-class San Rafael tin (-copper) deposit (central Andean tin belt, southeast Peru) 22
is an exceptionally large and rich (>1 Mt Sn; grades typically >2% Sn) cassiterite-bearing 23
hydrothermal vein system hosted by a late Oligocene (ca. 24 Ma) peraluminous K-feldspar- 24
megacrystic granitic complex and surrounding Ordovician shales affected by deformation and 25
low-grade metamorphism. The mineralization consists of northwest-trending, quartz- 26
cassiterite-sulfide veins and fault-controlled breccia bodies (>1.4 km in vertical and horizontal 27
extension). They show volumetrically important tourmaline alteration that principally formed 28
prior to the main ore stage, similar to other granite-related Sn deposits worldwide. We present 29
here a detailed textural and geochemical study of tourmaline, aiming to trace fluid evolution of 30
the San Rafael magmatic-hydrothermal system that led to the deposition of tin mineralization.
31
Based on previous works and new petrographic observations, three main generations of 32
tourmaline of both magmatic and hydrothermal origin were distinguished and were analyzed in 33
situ for their major, minor, and trace element composition by electron microprobe analyzer 34
(EMPA) and laser ablation–inductively coupled plasma–mass spectrometry (LA-ICP-MS), as 35
well as for their bulk Sr, Nd, and Pb isotopic compositions by multi-collector–inductively 36
coupled plasma–mass spectrometry (MC-ICP-MS). A first late-magmatic tourmaline 37
generation (Tur 1) occurs in peraluminous granitic rocks as nodules and disseminations, which 38
do not show evidence of alteration. This early Tur 1 is texturally and compositionally 39
homogeneous, it has a dravitic composition, with Fe/(Fe+Mg) = 0.36-0.52, close to the schorl- 40
dravite limit, and relatively high contents (10s-100s ppm) of Li, K, Mn, LREE, and Zn. The 41
second generation (Tur 2), the most important volumetrically, is pre-ore, high-temperature 42
(>500°C), hydrothermal tourmaline occurring as phenocryst replacement (Tur 2a) and open- 43
space fillings in veins and breccias (Tur 2b), and microbreccias (Tur 2c) emplaced in the host 44
granites and shales. Pre-ore Tur 2 typically shows oscillatory zoning, possibly reflecting rapid 45
changes in the hydrothermal system, and has a large compositional range that spans the schorl 46
to dravite fields with Fe/(Fe+Mg) = 0.02-0.83. Trace element contents of Tur 2 are similar to 47
Tur 1. Compositional variations within Tur 2 may be explained by the different degree of 48
interaction of the magmatic-hydrothermal fluid with the host rocks (granites and shales), in part 49
due to the effect of replacement vs. open-space filling. The third generation is syn-ore 50
hydrothermal tourmaline (Tur 3). It forms microscopic veinlets and overgrowths, partly cutting 51
previous tourmaline generations, and is locally intergrown with cassiterite, chlorite, quartz, and 52
minor pyrrhotite and arsenopyrite from the main ore assemblage. Syn-ore Tur 3 has schorl- 53
foititic compositions with Fe/(Fe+Mg) = 0.48-0.94 that partly differ from those of late- 54
magmatic Tur 1 and pre-ore hydrothermal Tur 2. Relative to Tur 1 and Tur 2, syn-ore Tur 3 has 55
higher contents of Sr and HREE (10s-100s ppm), and unusually high contents of Sn (up to 56
>1000 ppm). Existence of these three main tourmaline generations, each having specific 57
textural and compositional characteristics, reflects a boron-rich protracted magmatic- 58
hydrothermal system with repeated episodes of hydrofracturing and fluid-assisted reopening 59
generating veins and breccias. Most trace elements in the San Rafael tourmaline do not correlate 60
with Fe/(Fe+Mg) ratios, suggesting that their incorporation was likely controlled by the 61
melt/fluid composition and local fluid-rock interactions. The initial radiogenic Sr and Nd 62
isotopic compositions of the three aforementioned tourmaline generations (0.7160-0.7276 for 63
87Sr/86Sr(i) and 0.5119-0.5124 for 143Nd/144Nd(i)) mostly overlap those of the San Rafael granites 64
(87Sr/86Sr(i) = 0.7131-0.7202 and 143Nd/144Nd(i) = 0.5121-0.5122) and support a dominantly 65
magmatic origin of the hydrothermal fluids. These compositions also overlap the initial Nd 66
isotope values of Bolivian tin porphyries. The initial Pb isotopic compositions of tourmaline 67
show larger variations, with 206Pb/204Pb(i), 207Pb/204Pb(i), and 208Pb/204Pb(i) ratios mostly falling 68
in the range of 18.6-19.3, 15.6-16.0, and 38.6-39.7, respectively. These compositions partly 69
overlap the initial Pb isotopic values of the San Rafael granites (206Pb/204Pb(i) = 18.6-18.8, 70
207Pb/204Pb(i) = 15.6-15.7, and 208Pb/204Pb(i) = 38.9-39.0) and are also similar to those of other 71
Oligocene to Miocene Sn-W ± Cu-Zn-Pb-Ag deposits in southeast Peru. REE patterns of 72
tourmaline are characterized, from Tur 1 to Tur 3, by decreasing (Eu/Eu*)N ratios (from 20 to 73
2) that correlate with increasing Sn contents (from 10s to >1000 ppm). These variations are 74
interpreted to reflect evolution of the hydrothermal system from reducing towards relatively 75
more oxidizing conditions, still in a low-sulfidation environment as indicated by the pyrrhotite- 76
arsenopyrite assemblage. The changing textural and compositional features of Tur 1 to Tur 3 77
reflect the evolution of the San Rafael magmatic-hydrothermal system and support the model 78
of fluid mixing between reduced, Sn-rich magmatic fluids and cooler, oxidizing meteoric 79
waters as the main process that caused cassiterite precipitation.
80 81
Introduction 82
83
Tourmaline is the main borosilicate mineral in granitic rocks and associated magmatic- 84
hydrothermal ore deposits (e.g., Sn-W veins and greisens, pegmatites, Cu-Au breccia pipes, 85
Cu-Mo-(Au) porphyries, IOCG deposits), where it is mostly found as late-magmatic 86
disseminations, hydrothermal veins and breccias, and metasomatic replacements (e.g., London 87
et al., 1996; Slack, 1996; Černý et al., 2005; Slack and Trumbull, 2011). Due to its stability 88
over a wide pressure-temperature range and negligible intra-crystalline diffusion, tourmaline is 89
an especially refractory mineral that preserves the initial signature of physicochemical 90
processes during its crystallization (e.g., Dutrow and Henry, 2011; Marschall and Jiang, 2011;
91
van Hinsberg et al., 2011). Tourmaline is common in granite-related Sn-W deposits and has 92
been used as a proxy for tracing ore-forming processes and fluid origins (e.g., Duchoslav et al., 93
2017; Codeço et al., 2017, 2019; Hong et al., 2017; Launay et al., 2018; Harlaux et al., 2019a).
94
Previous studies have shown that major element compositions of tourmaline reflects 95
dominantly those of the host rocks and mineralizing fluids (e.g., Henry and Guidotti, 1985;
96
Slack et al., 1993; Henry and Dutrow, 1996; van Hinsberg et al., 2011), whereas incorporation 97
of trace elements is controlled by melt/fluid composition, local fluid-rock interactions, and 98
crystal-chemical effects (e.g., van Hinsberg, 2011; Marks et al., 2013; Hazarika et al., 2015;
99
Yang et al., 2015; Kalliomäki et al., 2017).
100
In this study, we address the question of how textural and compositional variations of 101
tourmaline reflect the evolution of late-magmatic to hydrothermal processes leading to the 102
precipitation of ore minerals. Herein, tourmaline was analyzed from the world-class San Rafael 103
Sn (-Cu) deposit (central Andean tin belt, southeast Peru) that represents an exceptionally large 104
and rich, granite-related cassiterite-bearing vein system extensively studied by previous 105
workers (Arenas, 1980; Palma, 1981; Kontak and Clark, 2002; Mlynarczyk et al., 2003;
106
Mlynarczyk, 2005; Mlynarczyk and Williams-Jones, 2006; Wagner et al., 2009; Corthay, 2014;
107
Prado, 2015). The central Andean tin belt is a classic metallogenic province for Sn-W deposits 108
in which widespread tourmaline occurs in veins, stockworks, breccias, and greisens (e.g., 109
Sillitoe et al., 1975; Lehmann et al., 1990, 2000). Tourmalinization is related to the highly 110
evolved character of magmas of the central Andean tin belt as indicated by high boron contents 111
(avg = 225 ppm B) measured in melt inclusions from Bolivian tin porphyries (Dietrich et al., 112
2000; Lehmann et al., 2000; Wittenbrink et al., 2009). The volumetrically important occurrence 113
of tourmaline at San Rafael, principally formed prior to the main ore stage, can be compared to 114
other granite-related Sn-dominated deposits, such as those of the Cornwall province in 115
southwest England (e.g., Drivenes et al., 2015; Duchoslav et al., 2017), as well as to granite- 116
related W-dominated deposits, such as Panasqueira in Portugal (e.g., Codeço et al., 2017, 2019;
117
Launay et al., 2018).
118
Several generations of tourmaline, from late-magmatic to syn-ore hydrothermal, have been 119
previously documented in the San Rafael deposit (Table 1; Kontak and Clark, 2002;
120
Mlynarczyk et al., 2003; Mlynarczyk and Williams-Jones, 2006), making it a particularly 121
appropriate case study to test the use of tourmaline as a tracer of fluid evolution. Based on 122
previous studies on the San Rafael deposit and new petrographic observations, we distinguish 123
for the scope of the present work three main tourmaline generations: (i) late-magmatic Tur 1 124
occurring as nodules and disseminations in peraluminous granites, (ii) pre-ore hydrothermal 125
Tur 2 forming replacements and open-space fillings in veins, breccias, and microbreccias, and 126
(iii) syn-ore hydrothermal Tur 3 occurring as late microscopic veinlets and overgrowths, locally 127
intergrown with cassiterite, chlorite, and quartz from the main ore assemblage. Using a large 128
set of representative samples of these three tourmaline generations, we present here a 129
comprehensive dataset of in situ analysis of major, minor, and trace elements and bulk Sr, Nd, 130
and Pb isotopic compositions. We show that the changing textural and compositional features 131
of Tur 1 to Tur 3 reflect the evolution of the San Rafael magmatic-hydrothermal system, 132
including increasing Sn contents (from 10s to >1000 ppm) and changes from reducing toward 133
relatively more oxidizing conditions. These results are consistent with and support the model 134
of fluid mixing between reduced, Sn-rich magmatic fluids and cooler, oxidizing meteoric 135
waters, as the main process that caused cassiterite precipitation at San Rafael.
136 137
Regional geology of the central Andean tin belt 138
139
The San Rafael Sn (-Cu) deposit (latitude 14°13’58” S, longitude 70°19’18” W) is located 140
within the Cordillera de Carabaya in the Eastern Cordillera of southeast Peru and defines the 141
northern end of the central Andean tin belt (Fig. 1A). This metallogenic province extends ca.
142
1000 km to the southeast as a 30- to 130-km-wide belt from southern Peru through Bolivia to 143
northern Argentina, and hosts hundreds of Sn-W ± Cu-Zn-Pb-Ag vein deposits (e.g., Kelly and 144
Turneaure, 1970; Turneaure, 1971; Grant et al., 1979; Lehmann et al., 1990). These deposits 145
are spatially associated with peraluminous granitoid intrusions and subvolcanic stocks that were 146
emplaced during three major metallogenic periods: (i) Late Devonian - early Carboniferous;
147
(ii) Late Triassic - Early Jurassic, restricted to the northern part of the belt; and (iii) Late 148
Oligocene - Early Miocene, affecting the entire belt (e.g., Grant et al., 1979; Clark et al., 1983, 149
1990; McBride et al., 1983; Rice et al., 2005). The most important episode for deposition of 150
Sn-W ore was the last, between 25 and 12 Ma, including several world-class deposits such as 151
San Rafael, Llallagua, Cerro Rico de Potosi, and Chorolque (e.g., Kelly and Turneaure, 1970;
152
Turneaure, 1971; Sillitoe et al., 1975; Lehmann et al., 1990). The Cordillera de Carabaya in 153
southeast Peru consists of a more than 10-km-thick sequence of lower Paleozoic 154
metasedimentary rocks (San José, Sandia, and Ananea formations) overlying unexposed 155
Precambrian gneissic basement that were all deformed and metamorphosed during the Variscan 156
orogeny (Laubacher, 1974; Clark et al., 1990; Sandeman et al., 1996). The metasedimentary 157
rocks were intruded by Oligocene to Miocene S-type granitic bodies and overlain by volcanic 158
rocks that are part of the Crucero Supergroup (e.g., Kontak et al., 1987; Laubacher et al., 1988;
159
Clark et al., 1990; Kontak et al., 1990; Cheilletz et al., 1992; Sandeman et al., 1997).
160 161
Geology of the San Rafael Sn (-Cu) deposit 162
163
San Rafael currently is one of the largest and highest-grade primary Sn deposits in the 164
world, with total past production of >1 Mt of Sn, and remaining total reserves estimated at 8 165
Mt of ore at 1.74% Sn (Minsur S.A., 2018). According to internal Minsur Reports, between 166
1969, when the production started at around 4900 masl, and 2019, when the underground 167
workings have deepened down to 3600 masl, the San Rafael mine has produced 26,608,702 t 168
ore with 3.7% Sn on average. Production in 2019 amounted 1,159,299 t ore at 1.86% Sn. Until 169
1978, when only the upper parts of the deposit were mined, Cu grades were higher than those 170
of Sn (for example 29,100 t at 4.62% Cu and 1.14% Sn in 1969, and 111,926 t at 1.49% Cu and 171
1.31% Sn in 1978). Subsequently, Cu grades decreased until 1986 (229,784 t at 0.37% Cu and 172
2.84% Sn), last year when Cu production has been reported. Between 1987 and 2012, Sn grades 173
were >3% with a maximum of 6.76% Sn for 356,106 t ore in 1994. At that time, bonanza ore 174
with exceptional grades reaching as high as >20% Sn was mined. Since 2013, Sn grades have 175
decreased from 2.72% in 2013 to 1.86% in 2019, in line with the progressive deepening of the 176
mining operations and the increase of mining efficiency.
177
The mineralization consists of a northwest-trending, quartz-cassiterite-sulfide vein system 178
spatially associated with a late Oligocene granitic complex hosted in Ordovician shales of the 179
Sandia Formation that was affected by contact metamorphism during granite emplacement (Fig.
180
1B; Kontak and Clark, 2002; Mlynarczyk et al., 2003; Gialli et al., 2017). The intrusive complex 181
includes the San Rafael granite that crops out in the southwestern part of the district and the 182
Quenamari granite in the northeast, both being part of the same intrusion at depth, as indicated 183
by drilling and underground workings (Fig. 1C). The San Rafael and Quenamari granitic bodies 184
consist dominantly of peraluminous porphyritic biotite-cordierite-bearing monzogranite, 185
exhibiting several centimeter size phenocrysts of K-feldspar, and were extensively studied by 186
Kontak and Clark (2002). Minor enclaves of biotite microgranite and diorite, showing mingling 187
textures with the host K-feldspar-megacrystic granite, have also been described in the upper 188
part of the stock (Kontak and Clark, 2002; Mlynarczyk et al., 2003). The San Rafael and the 189
Quenamari granitic intrusions are surrounded by porphyritic ring dikes, which are 190
petrologically similar to the central granites and exhibit quenched textures (Fig. 1B; Kontak 191
and Clark, 2002). In the southwestern part of the San Rafael granite, a greisen (quartz + 192
muscovite ± tourmaline assemblage) crops out at an altitude of >4800 masl for a few hundred 193
meters, elongated in a NNW-SSE direction (Fig. 1B). This area of intense greisenization is 194
interpreted to represent the apical part of the granitic pluton, where acidic magmatic fluids 195
accumulated during cooling of the parental magma (Gialli et al., 2019). Clark et al. (2000) and 196
Kontak and Clark (2002) dated the San Rafael granite at 24.6 to 24.7 ± 0.2 Ma (U-Pb zircon 197
and monazite ages). Hydrothermal activity is dated between 24.1 ± 0.1 Ma and 22.0 ± 0.2 Ma 198
(40Ar/39Ar muscovite and adularia plateau ages; Clark et al., 2000; Kontak and Clark, 2002;
199
Gialli et al., 2017). Dikes of lamprophyres, 10 to 50 cm in thickness and having a phlogopite + 200
K-feldspar + plagioclase ± quartz dominated assemblage, cut the San Rafael intrusive complex 201
and partly crop out in the southwestern margin of the Quenamari granite (Fig. 1B-C).
202
Lamprophyre enclaves within the San Rafael K-feldspar-megacrystic granite have also been 203
reported (Kontak and Clark, 2002). Lamprophyric magmas mixing with peraluminous S-type 204
biotite-cordierite magmas (“hybridization”) represent a common feature of the Crucero 205
Supergroup in the Eastern Cordillera, where they emplaced between ca. 26 and 22 Ma (e.g., 206
Carlier et al., 1997; Sandeman et al., 1997; Sandeman and Clark, 2003).
207
The paragenetic sequence of the San Rafael deposit comprises four main stages (Palma, 208
1981; Kontak and Clark, 2002; Mlynarczyk et al., 2003): (i) Stage I corresponds to formation 209
of early barren quartz-tourmaline veins and breccia bodies, which cut the granitic intrusion and 210
surrounding host rocks. This stage was preceded by early episodes of hydrothermal alteration, 211
including incipient potassic and sericitic alteration of magmatic plagioclase and K-feldspar in 212
the granites, and pervasive sodic alteration forming hydrothermal albite haloes around the 213
quartz-tourmaline veins and affecting the granitic groundmass (Kontak and Clark, 2002; Gialli 214
et al., 2019). (ii) Stage II consists of the Sn ore-stage assemblage composed of quartz, 215
cassiterite, chlorite, and minor tourmaline, pyrrhotite, and arsenopyrite, together forming veins 216
and breccias that locally reopened pre-existing quartz-tourmaline veins-breccias. This stage is 217
associated with pervasive chloritization, mainly along veins and breccia bodies, that affected 218
both the granites and the metasedimentary host rocks. (iii) During stage III a sulfide-dominant 219
assemblage, comprising mostly pyrrhotite, chalcopyrite, arsenopyrite, galena, and sphalerite, 220
was deposited. This stage occurs as quartz-sulfide veins and as late infill that cut or reopened 221
pre-existing vein generations. (iv) Late barren quartz-carbonate (calcite, siderite) stage IV 222
veins, containing minor fluorite and adularia, cut all previous veins and breccias.
223
In the present investigation, we studied representative samples from the San Rafael deposit 224
covering a large spectrum of different generations of magmatic and hydrothermal tourmaline.
225
The majority of the studied samples come from the San Rafael mine and were collected from 226
underground workings between 3610 masl and 4475 masl.
227 228
Analytical methods 229
230
Scanning electron microscopy and automated mineralogy (QEMSCAN) 231
232
Mineralogical observations of tourmaline were carried out using transmitted-light 233
microscopy combined with scanning electron microscopy (SEM) using a JEOL JSM7001F 234
SEM equipped with an energy-dispersive X-ray spectrometer (EDS) at the University of 235
Geneva, Switzerland. Back-scattered electron (BSE) images were acquired on carbon-coated 236
polished thin sections using an acceleration voltage of 15 kV, adjusting the image contrast to 237
reveal internal zoning within the tourmaline grains. Scanning electron microscopy was used to 238
study internal textures of tourmaline and to select grains for in situ chemical analyses.
239
Automated mineral analysis and textural imaging were performed using an FEI QEMSCAN 240
Quanta 650F facility at the University of Geneva. The QEMSCAN system is equipped with 241
two Bruker QUANTAX light-element EDS detectors. Analyses were conducted at high 242
vacuum, an accelerating voltage of 25 kV, and with a beam current of 10 nA on carbon-coated 243
polished thin sections. The field image operating mode (Pirrie et al., 2004) was used for 244
analyses. In total, 221 individual fields were measured per sample, with a field size of 1500 x 245
1500 µm and a point spacing of 5 µm. The standard 1000 counts per point were acquired, 246
yielding a limit of detection of approximately 2 wt.% per element for mineral classification.
247
Measurements were performed using the iMeasure v5.3.2 software; the iDiscover v5.3.2 248
software package was used for data processing. Results consist of: (i) spatially resolved and 249
fully quantified mineralogical maps; (ii) BSE images with identical resolution as in the 250
mineralogical maps; and (iii) X-ray elemental distribution maps.
251
252
Electron microprobe analysis 253
254
Major and minor element compositions of tourmaline were determined by electron 255
microprobe analyses (EMPA) using a JEOL JXA-8200 Superprobe microanalyzer equipped 256
with five wavelength-dispersive X-ray spectrometers (WDS) at the University of Geneva.
257
Operating conditions were as follows: acceleration voltage of 15 kV, beam current of 20 nA, 258
and beam diameter of 5 μm. Counting times on element peaks and backgrounds were 16 sec 259
and 8 sec, respectively, for Si, Al, K, Ca, Fe, Mg, Mn, Ti, and 30 sec and 15 sec, respectively, 260
for Na, Cr, F, Cl. The following standards were used for calibration: albite (Si, Al, Na), olivine 261
(Mg), fayalite (Fe), synthetic MnTiO3 (Mn), rutile (Ti), orthoclase (K), wollastonite (Ca), Cr2O3
262
(Cr), topaz (F), and tugtupite (Cl). Limits of detection are approximately: 200 ppm for Fe and 263
F; 170 ppm for Si; 130 ppm for Al, Mg, Mn, and Cr; 100 ppm for Na, K, Ti, and Ca; and 40 264
ppm for Cl. Structural formulae of tourmaline were calculated using the WinTcac software of 265
Yavuz et al. (2014) by normalizing to 15 cations for the Y+Z+T sites and assuming a 266
stoichiometric 3 atoms for B and 4 atoms for OH+F, based on the general formula 267
XY3Z6(T6O18)(BO3)3V3W, where X = Na+, Ca2+, K+, and vacancy (X□); Y = Fe2+, Mg2+, Mn2+, 268
Al3+, Li+, Fe3+, and Cr3+; Z = Al3+, Fe3+, Mg2+, Ti4+, and Cr3+; T = Si4+ and Al3+; B = B3+; V = 269
OH-, O2-; and W = OH-, F-, and O2- (Henry et al., 2011). Chemical compositions of tourmaline 270
are reported in weight per cent (wt.%) oxides and structural formulae are expressed in atoms 271
per formula unit (apfu). EMPA X-ray elemental maps were acquired for F, Ca, Na, Ti, and Cl, 272
using an acceleration voltage of 15 kV, beam current of 20 nA, beam diameter of 5 μm, and 273
dwell times of 320 msec. The presented map is 1024 x 1024 pixels in size with a 5 µm per pixel 274
resolution, corresponding to an investigated area of 5120 x 5120 µm. Total acquisition time 275
was ca. 103 hrs. The program XMapTools 3.1.2 (Lanari et al., 2014) was used for processing 276
QEMSCAN and EMPA maps, including classification and analytical standardization based on 277
EMPA spot analyses along the investigated grains.
278 279
Laser ablation–inductively coupled plasma–mass spectrometry 280
281
Tourmaline trace element analyses were carried out at the ETH Zürich, by laser ablation–
282
inductively coupled plasma–mass spectrometry (LA-ICP-MS) using a RESOlution (Australian 283
Scientific Instruments) 193 nm ArF excimer laser system attached to an Element XR (Thermo 284
Scientific, Germany) sector-field mass spectrometer. Analyses were performed directly on 285
polished thin sections of ca. 30 µm thickness, loaded in a Laurin Technic S-155 dual-volume 286
ablation cell fluxed with carrier gas consisting of ca. 0.5 L·min−1 He (5.0 grade) and sample 287
gas from the ICP-MS consisting of ca. 1 L·min−1 Ar (6.0 grade). We used a laser repetition rate 288
of 5 Hz, spot diameters of 43 or 51 μm, and a laser output energy of ca. 40 to 45 mJ, 289
corresponding to an on-sample energy density of ca. 4 J·cm–2. Three pre-ablation pulses were 290
applied immediately before each analysis for surface cleaning. Signal homogenization was 291
performed using in-house Squid tubing. The ICP-MS was tuned for maximum sensitivity on 292
the high mass range while keeping the production of oxides low (248ThO+/232Th+ <0.25%).
293
Intensities for the 40 following isotopes were acquired using time resolved-peak jumping and 294
triple detector mode: 7Li, 9Be, 27Al, 29Si, 31P, 39K, 43Ca, 45Sc, 49Ti, 51V, 53Cr, 55Mn, 59Co, 60Ni, 295
65Cu, 66Zn, 85Rb, 88Sr, 89Y, 90Zr, 93Nb, 118Sn, 133Cs, 139La, 140Ce, 141Pr, 146Nd, 147Sm, 151Eu, 157Gd, 296
159Tb, 163Dy, 165Ho, 166Er, 169Tm, 172Yb, 175Lu, 181Ta, 208Pb, and 209Bi. Dwell times were set to 297
10 msec, except for 27Al, 29Si, and 49Ti (5 msec), 9Be and 118Sn (20 msec), and all REEs (30 298
msec). With these settings, the total sweep time was 914 msec. Each measurement consisted of 299
70 mass sweeps acquired over ca. 60 sec, split in 30 sec of gas blank measurement followed by 300
30 sec of sample ablation. The raw intensities were processed using the Matlab-based SILLS 301
software (Guillong et al., 2008). LA-ICP-MS spectra were individually checked for possible 302
presence of micro-inclusions whereas spikes unrelated to the analyzed sample were 303
systematically eliminated. Integration windows were defined only for plateau-like signals thus 304
avoiding parts of the sample signal contaminated by the matrix or micro-inclusion. NIST SRM 305
612 glass (Jochum et al., 2011) was used as a calibration reference material (analyzed with a 306
spot diameter of 43 µm), via conventional standard-sample bracketing to correct for sensitivity 307
drift throughout the analytical session. Matrix effects were corrected using as an internal 308
standard the wt.% SiO2 content determined by EMPA for each analytical spot. Repeated 309
analyses of glasses prepared from USGS natural reference materials (GSD-1G and GOR128- 310
G1; analyzed with spot diameter of 43 µm) were processed as unknowns to check accuracy and 311
reproducibility of the analyses. Results show that reproducibility ranges from 3 to 12% (2σ, 312
increasing with decreasing concentration) for all analyzed trace elements, and that the analyses 313
are accurate within this level of analytical uncertainty. Limits of detection (LOD) were 314
calculated using the equation of Pettke et al. (2012) and are reported in Supplementary Table 315
2. We also assessed trace element contamination of the LA-ICP-MS system, which can be 316
significant for Sn (Schlöglova et al., 2017), by repeated analysis of fused silica glass using a 317
laser repetition rate of 5 Hz, spot diameters of 43 and 51 μm, and laser output energy of ca. 168 318
mJ, corresponding to an on-sample energy density of ca. 10 J·cm–2. Results show that 319
contamination is negligible for most trace elements, i.e., <0.5 ppm and even <0.1 ppm for m/z 320
ratios >85. Notable exceptions are 39K (<2 ppm), 43Ca (<40 ppm), and 118Sn (<1 ppm).
321
Nevertheless, the corresponding levels of contamination are within analytical uncertainty (i.e., 322
<10% and mostly <3% relative) for these elements, and therefore can also be considered 323
negligible.
324 325
Multi-collector–inductively coupled plasma–mass spectrometry 326
327
Bulk Sr, Nd, and Pb isotopic analyses of tourmaline were performed at the IGG-CNR in 328
Pisa, Italy. Tourmaline samples were crushed and sieved to 75-350-µm mesh and then 329
processed through a multi-step heavy-liquid procedure using sodium heteropolytungstate 330
solution (LST Fastfloat; density 2.9 g·cm−3) followed by centrifugation at 1000-2000 rpm. The 331
recovered heavy fractions were cleaned with deionized water and then dried in a furnace at 332
40°C overnight. Pure tourmaline concentrates were handpicked (~99% purity) under a 333
binocular microscope and were ground with an agate mortar to <80 µm. Tourmaline separates 334
were powdered and ~100 mg aliquots were taken for Sr, Nd, and Pb isotopic analyses. The 335
aliquots were digested in perfluoroalkoxy alkane (PFA) vials with mixed HF + HNO3 on a hot 336
plate at ~100°C for 5 days until complete sample digestion. The solutions were evaporated to 337
dryness and residues were re-dissolved in HNO3 and HCl 6.6N at ~100°C for 1 day. Strontium, 338
Nd, and Pb were separated using conventional ion-exchange procedures. Isotopic analyses were 339
performed using a Thermo Fisher Neptune Plus multi-collector–inductively coupled plasma–
340
mass spectrometer (MC-ICP-MS) at the IGG-CNR in Pisa. The instrument was equipped with 341
a combined cyclonic and Scott-type quartz-spray chamber, Ni-cones, and a MicroFlow PFA 342
100 µL·min−1 self-aspiring nebulizer. Strontium isotopic analyses were corrected for mass bias 343
fractionation using the 88Sr/86Sr ratio (8.375209) and for mass interference using the ratios 344
83Kr/84 Kr (0.201750), 83Kr/86Kr (0.664740), and 85Rb/87Rb (2.592310). Analytical accuracy 345
and long-term external reproducibility of Sr isotopic analyses were assessed using the reference 346
material NIST SRM 987 that yielded 87Sr/86Sr ratios of 0.710257 ± 0.000019. Instrumental 347
mass fractionation during Nd analyses was corrected using the 146Nd/144Nd ratio (0.7219). Mass 348
interference correction was done using the ratios 147Sm/144Sm(4.838710) and 147Sm/148Sm 349
(1.327400). Analytical accuracy and long-term external reproducibility of Nd isotopic analyses 350
were assessed using the reference material J-Ndi-1 that yielded 143Nd/144Nd ratios of 0.512098 351
± 0.000005. During Pb analyses, mass interference was corrected using the ratio 202Hg/204Hg 352
(4.350370). Samples were spiked with an in-house Tl standard and the ratio 203Tl/205Tl 353
(0.418882) was used to correct for mass bias fractionation. Analytical accuracy and long-term 354
external reproducibility of the Pb isotopic analyses were assessed over the measurement period 355
by analyzing replicates of the international standard SRM 981. The 2σ uncertainties for the 356
206Pb/204Pb, 207Pb/204Pb, and 208Pb/204Pb ratios are 0.02%, 0.03%, and 0.04%, respectively.
357 358
Results 359
360
Tourmaline textures and paragenetic sequence 361
362
Characteristics of the three main generations of tourmaline of both magmatic (Tur 1) and 363
hydrothermal origin (Tur 2 and Tur 3) distinguished in the San Rafael deposit are summarized 364
in the following. This classification is based on new petrographic evidence (crosscutting 365
relationships, replacements, overgrowths) and incorporates the findings of Kontak and Clark 366
(2002) and Mlynarczyk and Williams-Jones (2006). Differences with the classification of these 367
authors (mainly regarding Tur 1 and Tur 2) are shown in Table 1 and are also indicated in the 368
following.
369
Late-magmatic tourmaline (Tur 1): is found in peraluminous granites as quartz-tourmaline 370
nodules (Tur 1a) and disseminated grains (Tur 1b) that represent less than 2% of the rock 371
volume. Tourmaline is intergrown with magmatic quartz and K-feldspar that are devoid of 372
hydrothermal alteration, suggesting that Tur 1 is late-magmatic. Tur 1a nodules are hosted in 373
both the K-feldspar-megacrystic biotite-cordierite monzogranite and the microgranite, and form 374
centimeter-wide, rounded aggregates of intergrown tourmaline and quartz rimmed by a 0.5- to 375
1-cm-thick leucocratic halo composed of quartz and K-feldspar (Fig. 2C). The morphology, 376
distribution, and size of the quartz-tourmaline nodules at the scale of the granitic stock, both 377
vertically and laterally, have not been investigated. At a microscopic scale, Tur 1a forms brown- 378
orange, anhedral to subhedral millimeter size grains having a skeletal radial texture with 379
interstitial 200-1000 µm quartz and K-feldspar grains (Fig. 3A-B). Accessory minerals include 380
fluorapatite, rutile, and ilmenite present as micro-inclusions (10-200 µm) in quartz and 381
tourmaline. Tourmaline Tur 1a shows discrete core-to-rim zonation but lacks oscillatory zoning 382
(Fig. 4A-B). Disseminated tourmaline Tur 1b, described previously only in minor leucogranite 383
plugs and dikes along the southwestern and northeastern margins of the San Rafael granite 384
(Kontak and Clark, 2002; Mlynarczyk and Williams-Jones, 2006), is also observed in 385
microgranite. This tourmaline occurs as millimeter size greenish to orange-brownish, euhedral 386
to sub-euhedral grains intergrown with magmatic quartz, K-feldspar, plagioclase, and accessory 387
fluorapatite (Figs. 2D and 3C), similar to the descriptions of Kontak and Clark (2002) and 388
Mlynarczyk and Williams-Jones (2006). Back-scattered electron images of Tur 1b show a 389
relatively homogeneous internal texture displaying locally a discrete core-to-rim zoning (Fig.
390
4C).
391
Pre-ore hydrothermal tourmaline (Tur 2): occurs as post-magmatic replacements and in 392
veins and breccias in the granites and enclosing shales. Based on new petrographic 393
observations, three sub-generations (Tur 2a, Tur 2b, and Tur 2c) are distinguished. Tur 2a, in 394
part equivalent to the “early post-magmatic tourmaline” of Mlynarczyk and Williams-Jones 395
(2006), locally replaces partly or wholly magmatic minerals (K-feldspar phenocrysts, 396
cordierite, biotite) in the megacrystic monzogranite and microgranite (Fig. 2F). Tur 2a locally 397
forms randomly distributed rosettes that replaced K-feldspar phenocrysts within metasomatized 398
granites, as previously documented by Kontak and Clark (2002). Under the petrographic 399
microscope, Tur 2a is orange-brownish to bluish and typically forms sub-euhedral crystals 400
ranging in size typically from 100 to 500 µm (Fig. 3D-E). In BSE images, tourmaline reveals a 401
complex internal texture with a 50-100 µm-thick core rimmed by thin bands (down to <10 µm) 402
showing oscillatory and sector zoning (Fig. 4D-E). Tur 2b occurs in quartz-bearing veins and 403
clast-supported breccias cemented by quartz. This tourmaline generation represents the most 404
abundant at San Rafael and was previously described as “early hydrothermal tourmaline” by 405
Mlynarczyk and Williams-Jones (2006). Tur 2b veins and breccias are observed at surface up 406
to more than 5000 masl and underground at least down to 3600 masl, thus having a continuous 407
vertical extent of >1.4 km. They cut both the granitic intrusion (Tur 2b(g)) and the shales in the 408
contact metamorphic aureole (Tur 2b(s)). Tourmaline-quartz veins are generally narrow (<3 cm) 409
and systematically surrounded by an alteration halo <1 to 10 cm-wide composed of albite and 410
sericite, with minor apatite and rutile (Fig. 2E). In thin section, tourmaline hosted in the granite 411
(Tur 2b(g)) is orange-brown to green and forms acicular, sub-euhedral crystals ranging from 100 412
µm up to several millimeters in size (Fig. 3F). The tourmaline grains show a typical open-space 413
filling texture with c-axes perpendicular to the host rock contact and growth directions towards 414
the center of the vein. Characteristic is intense oscillatory zoning at the µm-scale as evidenced 415
by BSE images (Fig. 4F) and X-ray elemental mapping (Fig. 5). Prismatic sectors are 416
characterized by higher concentrations of Ca, Ti, and F, and lower Al, whereas oscillatory 417
zoning is marked mainly by variations of the Fe/Mg ratio (Fig. 5). Tur 2b in the veins and 418
breccias is considered broadly synchronous owing to (i) similar color, texture, and zoning 419
patterns; (ii) subsequent re-openings and fillings commonly observed within a single sample;
420
and (iii) absence of systematic crosscutting evidence (Figs. 3G and 4G). Tourmaline from veins 421
hosted in the metamorphic shales (Tur 2b(s)) shows features similar to those in the granite- 422
hosted tourmaline but is typically finer grained (50-200 µm) and is green-blue under plane- 423
polarized transmitted light (Fig. 3H). Tur 2c in tourmaline-quartz microbreccias is 424
volumetrically the most abundant (Fig. 2G). The microbreccias are commonly associated with 425
reopening of pre-existing quartz-tourmaline veins and contain clasts of early tourmaline veins 426
and breccias. Under the microscope, Tur 2c occurs as fine-grained (<10-50 µm) and randomly- 427
oriented brown grains that form a dense aggregate cementing clasts of tourmalinized wall rocks, 428
tourmaline veins, and relict magmatic quartz (Figs. 3I and 4H). The acicular crystals of older 429
quartz-tourmaline veins are in places broken, partially corroded, and replaced by Tur 2c that 430
makes up the microbreccia (Fig. 4I).
431
Syn-ore hydrothermal tourmaline (Tur 3): occurs as microscopic veinlets and overgrowths 432
that are partly cutting previous tourmaline generations (Fig. 2H). Previously identified by 433
Mlynarczyk and Williams-Jones (2006) as “ore-stage tourmaline,” it is mostly blue (Tur 3a) or 434
in places green (Tur 3b). Under the microscope, Tur 3a is pale to dark blue and forms either 50 435
to 200 µm-thick veinlets cutting orange-brown acicular tourmaline grains (Tur 2) or 10 to 100 436
µm-thick overgrowths in optical continuity with the host crystal (Figs. 3J-K and 4H-K). Due to 437
their relatively small size, these fine-scale overprints of Tur 3 onto Tur 2 can be easily 438
overlooked when using transmitted light imagery only. Overprinting of pre-ore hydrothermal 439
quartz-tourmaline (Tur 2) by fine-grained Tur 3 can be recognized and quantified using 440
combined X-ray elemental and QEMSCAN mapping, as Tur 3 has higher Fe concentrations. In 441
the sample shown in Figure 6, more than 15% of the analyzed area corresponds to Tur 3a. Green 442
tourmaline (Tur 3b) and blue tourmaline (Tur 3a) veinlets locally show inter-crosscutting 443
relationships (Fig. 3K), suggesting penecontemporaneous formation. Green tourmaline (Tur 444
3b) forms khaki-green to dark-green acicular, 50 to 200 µm-long grains commonly intergrown 445
with quartz and chlorite in open-space fillings. Locally, green tourmaline occurs as small 446
needle-like crystals intergrown with cassiterite, chlorite, and quartz from the main ore 447
assemblage (Figs. 3L and 4L). This textural relationship indicates that tourmaline continued to 448
form during Sn ore deposition but in volumetrically minor amounts compared to the pre-ore 449
hydrothermal stage.
450 451
Major and trace element composition of tourmaline 452
453
The average major and minor element composition and structural formula of the three 454
tourmaline generations are reported in Table 2 and the full dataset in Supplementary Table 1.
455
Tourmaline from San Rafael shows major element variations, principally for Fe (0.05-2.21 456
apfu), Mg (0.13-2.90 apfu), and Al (5.41-7.20 apfu), as well as for minor elements such as Na 457
(0.30-0.86 apfu), Ca (0.01-0.34 apfu), and Ti (0.01-0.50 apfu). The different tourmaline 458
generations (Tur 1 to Tur 3) have compositions belonging to the schorl-dravite and foitite-Mg- 459
foitite solid solutions (Figs. 7 and 8), which is consistent with the EMPA data reported by 460
Kontak and Clark (2002) and Mlynarczyk and Williams-Jones (2006). Tur 1 belongs to the 461
alkali group (Fig. 7) and has dravitic compositions, close to the schorl-dravite limit, with 462
Fe/(Fe+Mg) = 0.36-0.52 and X□/(X□+Na+K) = 0.08-0.43. Tur 2 is also dominantly within the 463
alkali group and has intermediate compositions corresponding to the schorl-dravite solid 464
solution with Fe/(Fe+Mg) = 0.02-0.83 and X□/(X□+Na+K) = 0.02-0.62. Compared to Tur 2b(g), 465
shale-hosted Tur 2b(s) is distinguished by higher X□/(X□+Na+K) = 0.44-0.62 and Fe/(Fe+Mg) 466
= 0.61-0.80. In contrast, Tur 3 belongs to the alkali to X-vacant group and has schorl to foitite 467
compositions with Fe/(Fe+Mg) = 0.48-0.94 and X□/(X□+Na+K) = 0.11-0.70. The strongest 468
compositional variations relate to the Fe/(Fe+Mg) ratio, reflecting the substitution vector 469
Fe2+Mg2+-1 as indicated by a general linear trend in the Fe vs. Mg diagram (Fig. 8). All 470
tourmaline analyses show values of Fe+Mg <3 apfu and Al >6 apfu, suggesting the presence of 471
excess Al in the Y-site. The positive correlation between X-site vacancies and Al contents 472
indicates a combination of the three substitution vectors (☐Al3+)(Na+Mg2+)-1, (Al3+O2- 473
)(Mg2+OH-)-1, and (☐Al3+2)(Ca2+Mg2+2)-1 (Fig. 9), which also explains the variations in Ca and 474
Na.
475
Trace element contents of tourmaline from San Rafael are summarized in Table 3 and the 476
full dataset is reported in Supplementary Table 2. Tourmaline from San Rafael has generally 477
very low to low concentrations (<0.1 to 10 ppm) of Be, Co, Cu, Rb, Y, Zr, Nb, Ta, Cs, Bi, Pb, 478
and most REE; intermediate concentrations (10s to 100 ppm) of Li, P, Sc, V, Cr, Ni, Zn, Sr, 479
and partly Sn; and high to very high concentrations (100s to >1000 ppm) of K, Ca, Ti, and Mn, 480
along with Sc, V, Cr, and Sn in some cases. Variation diagrams of some trace elements as a 481
function of Fe/(Fe+Mg) ratio are shown in Figure 10, and percentile box and whisker plots are 482
shown in Supplementary Figure 1. The different tourmaline generations have overlapping and 483
highly variable concentrations for some trace elements such as Cr (1-700 ppm), Sc (1-1000 484
ppm), Nb (0.05-35 ppm), Ta (0.01-10 ppm), REE (0.05-50 ppm), and Pb (0.1-14 ppm). Other 485
trace elements show more systematic variations across the three main tourmaline generations.
486
Late-magmatic Tur 1 is characterized by high contents of Li (60-200 ppm), K (150-500 ppm), 487
Ti (1200-12000 ppm), Zn (83-360 ppm), and Mn (65-500 ppm). In contrast, syn-ore 488
hydrothermal Tur 3 is distinguished by higher contents of Be (2-50 ppm), Sr (17-530 ppm), and 489
Sn (105-2240 ppm). Intermediate trace element concentrations between those of Tur 1 and Tur 490
3 characterize pre-ore hydrothermal Tur 2, particularly for Li (3-250 ppm), Be (0.5-65 ppm), K 491
(32-400 ppm), Zn (1-300 ppm), and Sn (3-200 ppm). Compared to Tur 2b(g), shale-hosted Tur 492
2b(s) contains lower Li (3.5-15 ppm) and K (75-112 ppm) contents, and higher V (280-402 493
ppm), Mn (89-121 ppm), and Zn (235-290 ppm) contents (Fig. 10).
494
Positive linear correlations are observed for some trace elements such as Nb vs. Ta (R2 = 495
0.48), K vs. Pb (R2 = 0.52), Mn vs. Zn (R2 = 0.69), V vs. Cr (R2 = 0.58), Sc vs. V (R2 = 0.69), 496
Ti vs. Ca (R2 = 0.42), LREE vs. Ca (R2 = 0.79), and Ca vs. Sr (R2 = 0.52) (Supplementary 497
Figure 2). The observed co-variations likely indicate substitution-controlled mechanisms for 498
their incorporation in the tourmaline structure, principally the X-, Y-, and Z-sites. In multi- 499
element diagrams normalized to upper continental crust (UCC), tourmaline is characterized by 500
relative enrichments in Li, Be, Sn, and Zn between 2 and 1000 times higher than the UCC 501
values, and relative depletions in Sr, K, Ca, Cs, Rb, Y, P, Zr, Bi, Pb, Mn, Co, Ni, and Cu 502
between 2 and 1000 times lower than the UCC values (Fig. 11). The REE content of tourmaline 503
is uniformly low (<0.1 to 50 ppm), between 10 and 1000 times lower than the UCC values, and 504
is characterized by variable LREE/HREE ratios with systematic positive Eu anomalies (Fig.
505
11), which is quantified by the ratio (Eu/Eu*)N = EuN/(SmN x GdN)0.5 where N corresponds to 506
the chondrite-normalization. Tur 1 and Tur 2 are enriched in LREE over HREE, whereas Tur 3 507
has higher contents of HREE relative to LREE. Tourmaline Tur 1 to Tur 3 shows a progressive 508
decrease of the (La/Yb)N (from 1000 to 0.01) and (Eu/Eu*)N (from 20 to 2) ratios, correlating 509
with a progressive increase of the Sn content, from 10s to 100 ppm for Tur 1 and Tur 2, and as 510
high as >1000 ppm for Tur 3 (Fig. 12).
511 512
Multivariate statistical analysis 513
514
Two types of multivariate statistical techniques, principal component analysis (PCA) and 515
discriminant projection analysis (DPA), were applied to the major and trace element dataset of 516
tourmaline in order to understand and quantitatively classify the observed compositional 517
variations.
518
Principal component analysis is a classical multivariate statistical technique particularly 519
useful for treating large geochemical datasets, such as LA-ICP-MS trace element data (e.g., 520
Winderbaum et al., 2012; Belissont et al., 2014; Harlaux et al., 2018, 2019a). The aim of PCA 521
is to provide dimensionality reduction of correlated variables into a reduced set of orthogonal 522
linear combinations, so-called principal components, which maximize the variance and 523
minimize information loss (Izenman, 2008). A classical PCA has been applied to the major and 524
trace element dataset of San Rafael tourmaline (n = 258 spot analyses, separating the three main 525
generations Tur 1 to Tur 3). A total of 18 variables has been selected for the PCA including 526
four major elements (Al, Fe, Mg, Na) quantified by EMPA and 14 trace elements (Li, K, Ti, V, 527
Cr, Sc, Mn, Zn, Sr, Sn, Nb, Ca, LREE, HREE) determined by LA-ICP-MS. Other elements 528
present at very low concentrations (<1 ppm) or below the limits of detection were not included 529
in the analysis. Results of the PCA applied to the log-transformed dataset of major and trace 530
elements in tourmaline are shown in Figure 13A. Data are represented by a two-dimensional 531
projection of the two first principal components (PC1 vs. PC2), which describe the statistical 532
correlations between the investigated variables (n = 18). The elements projected on the PC1 vs.
533
PC2 plane account for 51.2% of element content variability. Four main groups of element 534
correlation clusters are discriminated by the PCA. A first group consisting of K, Nb, Mn, Zn, 535
and Fe, characterizes Tur 1. A second group composed of Li, Ca, Ti, LREE, Sr, and Na, and a 536
third group including Mg, Cr, Sc, and V, explain the variability of Tur 2. A fourth group 537
comprises Sn and HREE and represents mainly Tur 3. These statistical correlation clusters 538
reflect partly the crystal chemical control on the incorporation of major and trace elements in 539
tourmaline. Indeed, elements such as Fe, Mg, Zn, Mn, Cr, Sc, V, Li, and REE are incorporated 540
into the Y-site, whereas Na, Ca, and Sr enter the X-site (Henry et al., 2011; van Hinsberg, 541
2011).
542
Discriminant projection analysis is another multivariate statistical technique that aims to 543
classify a high-dimensional dataset into predefined groups known a priori by calculating a 544
linear discriminant function that maximizes ratios between the groups (Izenman, 2008). This 545
method has been used for geological or forensic investigations of trace element or isotopic 546
analyses of minerals (e.g., Dalpé et al., 2010). We have applied DPA to the same element 547
dataset for San Rafael tourmaline (n = 258 spot analyses, 18 variables) as done for PCA, and 548
defined three groups corresponding to the three generations of tourmaline (Tur 1 to Tur 3).
549
Results of the DPA applied to the log-transformed dataset of major and trace elements in 550
tourmaline are shown in Figure 13B. Data are represented by the two discriminant projections 551
(DP1 vs. DP2), which maximize the separation between the predefined groups. Tur 1 is mainly 552
projected on the DP2 axis and is defined by correlations of K, Mn, Zn, Ti, Ca, Li, Mg, LREE, 553
and V. Tur 2 lies at the center of the DP1 vs. DP2 projection, mainly reflecting correlations of 554
Mg, Al, Cr, Li, and LREE. Tur 3 is defined by the DP2 vector, reflecting correlations of Fe, Sn, 555
HREE, and Sr.
556
557
Sr, Nd, and Pb isotopic compositions of tourmaline 558
559
The Sr, Nd, and Pb radiogenic isotopic compositions of tourmaline are shown in Figure 14 560
and data are reported in Table 4. Results show variable initial ratios of 87Sr/86Sr(i) (0.7160- 561
0.7276) and 143Nd/144Nd(i) (0.5119-0.5124) that plot within the fields of Coastal and Western 562
Cordilleran intrusions and Arequipa-Antafolla metamorphic basement (see references in Fig.
563
14). Most samples of tourmaline have 87Sr/86Sr(i) and 143Nd/144Nd(i) values that overlap those of 564
the San Rafael granites (87Sr/86Sr(i) = 0.7131-0.7202 and 143Nd/144Nd(i) = 0.5121-0.5122;
565
Supplementary Table 3), thus supporting a dominantly magmatic origin for the Sr and Nd, and 566
by extension, for the tourmaline-precipitating hydrothermal fluids. Only one sample of 567
hydrothermal tourmaline from a vein hosted in shale has a relatively high Sr radiogenic 568
composition (sample SRG-21A, 87Sr/86Sr(i) = 0.7276), likely related to fluid-rock interaction 569
with the host rock. The initial Nd isotopic values of tourmaline fall in the compositional range 570
of Bolivian tin porphyries of Llallagua, Chorolque, and Cerro Rico de Potosi (143Nd/144Nd(i) = 571
0.5121-0.5124; Dietrich et al., 2000). The initial Pb isotopic compositions of tourmaline are 572
more variable, with 206Pb/204Pb(i), 207Pb/204Pb(i), and 208Pb/204Pb(i) ratios mostly falling in the 573
range of 18.6-19.3, 15.6-16.0, and 38.6-39.7, respectively. Three samples (PSR-24A and PSR- 574
24A-bis with 206Pb/204Pb(i) = 21.2-21.5, and PSR-12B with 206Pb/204Pb(i) = 22.8) have markedly 575
higher 206Pb/204Pb(i) values and are not shown in Figure 14. These highly radiogenic Pb 576
compositions likely reflect the precipitating fluid that had a high initial 238U/204Pb ratio. Most 577
Pb isotopic compositions of tourmaline plot close to the upper crust and orogen curves of 578
Zartman and Doe (1981) and fall within the compositional ranges of the Arequipa-Antafolla 579
metamorphic basement, the central volcanic zone, and the Coastal and Western Cordilleran 580
intrusions (see references in Fig. 14). These compositions partly overlap the initial Pb isotope 581
values of the San Rafael granites (206Pb/204Pb(i) = 18.6-18.8, 207Pb/204Pb(i) = 15.6-15.7, and 582
208Pb/204Pb(i) = 38.9-39.0; Supplementary Table 3) and are also similar to those of other 583
Oligocene to Miocene Sn-W ± Cu-Zn-Pb-Ag deposits in southeast Peru (206Pb/204Pb(i) = 18.5- 584
25.2, 207Pb/204Pb(i) = 15.6-16.0, and 208Pb/204Pb(i) = 38.6-40.1; Kontak et al., 1990).
585 586
Discussion 587
588
Tourmaline textures as indicator of formation conditions 589
590
The textural features of tourmaline may record particularities of the environment of 591
formation and provide valuable information on physicochemical conditions during its 592
crystallization (e.g., van Hinsberg et al., 2001; Dutrow and Henry, 2016, 2018). The textural 593
variations of tourmaline generations (from Tur 1 to Tur 3) reflect evolving physical and 594
chemical parameters of the San Rafael system during the magmatic-hydrothermal transition.
595
Texturally homogeneous nodules and disseminations of Tur 1 intergrown with the quartz- 596
feldspar granitic groundmass (Figs. 3 and 4) are typical of magmatic tourmaline in evolved 597
peraluminous granites (e.g., London and Manning, 1995; London et al., 1996; Balen and 598
Broska, 2011; Balen and Petrinec, 2011; Drivenes et al., 2015). The formation of tourmaline 599
nodules has been widely discussed in the literature and three main hypotheses have been 600
proposed for their origin: (i) post-magmatic hydrothermal alteration of granitic bodies by 601
externally derived boron-rich fluids (e.g., Rozendaal and Bruwer, 1995); (ii) crystallization 602
from immiscible, hydrous, boron-aluminosilicate melts or boron-rich aqueous fluids that 603
separated from coexisting silicate melt (e.g., Dini et al., 2007; Balen and Broska, 2011;
604
Drivenes et al., 2015; Burianek et al., 2016); and (iii) products of magmatic crystallization on 605
the liquid line of descent of boron-rich granitic melts (e.g., Perugini and Poli, 2007; Balen and 606
Petrinec, 2011; Valentini et al., 2015). The Tur 1 nodules observed in the San Rafael granites 607
are devoid of alteration features (i.e., absence of veins, dissolution-reprecipitation texture, and 608
pervasive alteration), thus arguing against a hydrothermal origin related to post-magmatic 609
alteration. Instead, the rounded shapes of the nodules, presence of leucocratic rims, and limited 610
intergrowths with the granitic groundmass better suggest their physical separation from the 611
silicate melt during crystallization. Experimental studies indicate that tourmaline saturation in 612
strongly peraluminous granitic melts (ASI >1.2) can be reached after extended crystal 613
fractionation when the boron content of the melt attained >2 wt.% B2O3 at 750°C and 2 kbar 614
(Wolf and London, 1997). Other works based on petrographic relations, geochemical data, and 615
compositional phase diagrams suggested that minimum boron content in the range of 0.05-0.3 616
wt.% B2O3 is sufficient to saturate a peraluminous granitic melt in tourmaline between 600 and 617
750ºC (Pesquera et al., 2013). The San Rafael megacrystic granite has a moderately fractionated 618
peraluminous S-type composition (ASI = 1.1-1.5), as indicated by whole-rock and mineral 619
chemistry data (Kontak and Clark, 2002; Mlynarczyk, 2005; Corthay, 2014; Prado, 2015). Its 620
whole-rock boron content (B = 60-160 ppm, avg = 115 ppm; Mlynarczyk, 2005) falls in the 621
same range of peraluminous granites hosting tourmaline nodules elsewhere (B = 10-500 ppm;
622
Dini et al., 2007; Balen and Broska, 2011; Pesquera et al., 2013; Drivenes et al., 2015; Burianek 623
et al., 2016) as well as quartz-hosted melt inclusions from Bolivian tin porphyries (B = 35-640 624
ppm; Dietrich et al., 2000; Lehmann et al., 2000; Wittenbrink et al., 2009). This suggests that 625
the boron content of the San Rafael granitic melt may have been initially much higher owing 626
to the preferential partitioning of boron into a fluid phase relative to the silicate melt during 627
water saturation (e.g., Pichavant, 1981; London et al., 1988; Hervig et al., 2002; Thomas et al., 628
2003; Schatz et al., 2004). In addition, the presence of 0.5- to 1-cm-thick quartz-K-feldspar 629
leucocratic rims surrounding the Tur 1 nodules suggest that tourmaline crystallization was 630
favored by decomposition reactions of biotite and cordierite, possibly at temperatures <750ºC 631
close to solidus (Wolf and London, 1997). Based on the arguments above, we propose that the 632
Tur 1 nodules crystallized from an immiscible, hydrous, boron-aluminosilicate melt that 633
separated from the San Rafael granitic melt during the late-magmatic stage prior or 634
concomitantly to the magmatic-hydrothermal transition.
635
Pre-ore hydrothermal Tur 2 is characterized by intense oscillatory zoning at a micrometer 636
scale (Figs. 3 to 4), which is characterized by fluctuating concentrations of major, minor and 637
trace elements and progressively increasing Fe/(Fe+Mg) ratios along the growth direction 638
(Supplementary Figure 3). Oscillatory zoning in hydrothermal minerals is generally interpreted 639
to reflect (i) periodic changes in the external environment due to varying fluid properties such 640
as chemical composition, pressure, temperature, or fO2 (Holten et al., 1997), (ii) intrinsic self- 641
organization processes during crystal growth controlled by absorption-diffusion reactions at the 642
crystal-fluid interface (L’Heureux and Jamtveit, 2002), or (iii) a combination of both external 643
and internal fluctuating factors (Shore and Fowler, 1996). Quartz- and tourmaline-hosted 644
primary fluid inclusions from hydrothermal veins cutting the San Rafael megacrystic granite 645
yielded trapping temperatures of >500ºC and high salinity (34-62 wt.% NaCl equiv) in 646
lithostatic pressure conditions of 0.8 kbar (Kontak and Clark, 2002; Mlynarczyk et al., 2003;
647
Wagner et al., 2009; Corthay, 2014; Prado, 2015). Under such pressure-temperature conditions, 648
the oscillatory zoning observed in Tur 2 crystals is attributed to rapid changes in the 649
hydrothermal system, possibly caused by evolving physicochemical fluid conditions and/or by 650
rapid crystal growth. This interpretation is supported by (i) the open-space filling textures 651
typical of the Tur 2b crystals that grew perpendicular to the vein selvages, (ii) the increasing 652
Fe/(Fe+Mg) ratios along growth directions, and (iii) evidence for multiple vein-opening events;
653
all together indicating a dynamic fluid environment. Experimental synthesis of tourmaline 654
single crystals from boron-rich hydrothermal solutions (400-750°C, 1 kbar) reported a growth 655
rate of 0.05 mm/day (Setkova et al., 2011), which possibly reproduces the fluid dynamics in 656
natural magmatic-hydrothermal systems. Typical textures of repeated hydrofracturing and 657
fluid-assisted reopening (e.g., Jébrak, 1997) observed in quartz-tourmaline veins and breccias 658
