# New ages, morphometric and geochemical data on recent shoshonitic volcanism of the Puna, Central Volcanic Zone of Andes: San Jerónimo and Negro de Chorrillos volcanoes

Jose-Luis Fernandez-Turiel, Julio Saavedra, F. J. Perez-Torrado, A.

Rodriguez-Gonzalez, Marta Rejas, Hervè Guillou, Meritxell Aulinas

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Jose-Luis Fernandez-Turiel, Julio Saavedra, F. J. Perez-Torrado, A. Rodriguez-Gonzalez, Marta Rejas, et al.. New ages, morphometric and geochemical data on recent shoshonitic volcanism of the Puna, Central Volcanic Zone of Andes: San Jerónimo and Negro de Chorrillos volcanoes. Journal of South American Earth Sciences, 2021, 109, 10.1016/j.jsames.2021.103270. hal-03401862

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# New ages, morphometric and geochemical data on recent shoshonitic volcanism of the Puna, Central Volcanic Zone of Andes: San Jerónimo and Negro de Chorrillos volcanoes

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J.L. Fernandez-Turiel<sup>a, *</sup>, J. Saavedra<sup>b</sup>, F.J. Perez-Torrado<sup>c</sup>, A. Rodriguez-Gonzalez<sup>c</sup>, M. Rejas<sup>a</sup>, H. Guillou<sup>d</sup>, M. Aulinas<sup>e</sup>

## ARTICLE INFO

**Keywords:**

Shoshonite

Phlogopite

Dusty plagioclase

Glass

Central Volcanic Zone

Andes

## ABSTRACT

Shoshonitic volcanic centres are scarce within the back-arc mafic volcanism of the Puna Plateau in the Central Volcanic Zone of the Andes. San Jerónimo and Negro de Chorrillos volcanoes are two well preserved shoshonitic monogenetic edifices that developed mainly Strombolian activity and lava flows, which extended more than 9 and 4 km from the vent, producing a volume DRE (Dense Rock Equivalent) of 0.091 and 0.020 km³, respectively. The new unspiked K-Ar (groundmass) ages of San Jerónimo (144 ± 3 ka) and Negro de Chorrillos (51 ± 2 ka) eruptions presented in this work are both younger than those previously published. Petrographic and compositional features allow distinguishing between potassic and ultrapotassic suites in volcanic products of both eruptions. Phlogopite occurrence and low glass content characterize the ultrapotassic suite. Based on the crustal thickness of this part of the Puna and the experimental constraint, these magmas last equilibrated at Moho depths at around 50 km (the highest depth of spinel lherzolite stability field) for temperatures of 1100-1200 °C. Once in the crust, the K-rich magma underwent a minor interaction with a trachytic-trachydacitic melt, which is evidenced petrographically and geochemically, e.g., mismatch of mineral assemblages (co-occurrence of olivine and quartz), disequilibrium crystal textures (dusty zones in plagioclase, embayments in sanidine), and a wide range of compositional variation in the glasses. Ultrapotassic (i.e., phlogopite-bearing) and potassic lavas in the same eruption are interpreted as pulses of the same magma but with a longer mixing time with the trachytic-trachydacitic melt for the resulting ultrapotassic rocks within an eruptive time frame of hours or days. The singular shoshonitic character of these magmas could be related to their origin along an old terrane boundary. Afterwards, as occurs for other mafic magmas of the Puna, major NW-SE fault systems would favour their ascent to the surface.

## 1. Introduction

The Central Volcanic Zone of the Andes (CVZ; e.g., Thorpe, 1984; Stern, 2004) extends between 14° S in Peru and 27° S in Chile and Argentina (Fig. 1) and records a complex evolution of a volcanic arc and back-arc through hundreds of volcanoes, from constructional (polygenetic and monogenetic) to large collapse calderas (Grosse et al., 2017; Guzmán et al., 2014; Kay and Coira, 2009; Maro et al., 2016b, 2017;

Naranjo et al., 2018; Petrinovic et al., 2017; Tilling, 2009). The preservation of volcanic edifices is usually exceptional due to the prevailing arid or hyperarid conditions since early to middle Miocene times (Stern et al., 2007; Strecker et al., 2007). The CVZ overlaps the Central Andean Plateau or Altiplano-Puna Plateau, a ~4-7 km elevation region with the thickest continental crust on Earth, reaching up to 80 km in some areas (Allmendinger et al., 1997; Beck et al., 2015; Liu et al., 2011; Naranjo et al., 2018; Wörner et al., 2018; Yuan et al., 2002). It is the most

<footnote><sup>a</sup> Geosciences Barcelona, GEO3BCN, CSIC, Barcelona, Spain</footnote>

<footnote><sup>b</sup> IRNASA, CSIC, Salamanca, Spain</footnote>

<footnote><sup>c</sup> Instituto de Estudios Ambientales y Recursos Naturales (i-UNAT), Universidad de Las Palmas de Gran Canaria (ULPGC), Las Palmas de Gran Canaria, Spain</footnote>

<footnote><sup>d</sup> Laboratoire des Sciences Du Climat et L'Environnement, CEA-CNRS-UVSQ, Gif-sur-Yvette, France</footnote>

<footnote><sup>e</sup> Departament de Mineralogia, Petrologia I Geologia Aplicada Facultat de Ciències de La Terra, Universiat de Barcelona, Spain</footnote>

<footnote>* Corresponding author.</footnote>

<footnote>E-mail addresses: jlfernandez@geo3bcn.csic.es (J.L. Fernandez-Turiel), franciscojose.perez@ulpgc.es (F.J. Perez-Torrado), alejandro.rodriguezgonzalez@ulpgc.es (A. Rodriguez-Gonzalez), mrejas@geo3bcn.csic.es (M. Rejas), herve.guillou@lsce.ipsl.fr (H. Guillou), meritxellaulinas@ub.edu (M. Aulinas).</footnote>
extensive orogenic plateau on Earth formed in an oceanic-continental subduction setting (Garzione et al., 2017). This plateau consists of two morphotectonic units, the Altiplano (~14–23° S) and Puna (~23–27° S). In this geotectonic setting, the magmatism is driven by the geodynamic mechanisms of plateau uplift and the processes by which the continental crust overriding plate has been shortened and thickened, whereas the mantle lithosphere that originally underlay it has been thinned or entirely removed (Beck et al., 2015; Garzione et al., 2017; Kay and Coira, 2009). These processes produced variations in the crustal thickness of the plateau. This thickness is around 50 km depth in the studied area (~24–25° S) (Beck et al., 2015; McGlashan et al., 2008; Yuan et al., 2002).

The Andean CVZ's volcanic activity is located in both the frontal arc and the back-arc from Miocene to Quaternary (e.g., Barnes and Ehlers, 2009; Jiménez and López-Velásquez, 2008; Kay and Coira, 2009; Mamani et al., 2010). Polygenetic volcanism produced stratovolcanoes with associated andesitic to dacitic pyroclastic fall deposits and lava flows, and dacitic to rhyolitic domes and pyroclastic flow deposits, mainly in the frontal volcanic arc, although there are also some scarce stratovolcanoes in the back-arc, e.g., Tunupa (Salisbury et al., 2015), Uturuncu (Michelfelder et al., 2013; Sparks et al., 2008), and Tuzgle (Coira and Kay, 1993; Norini et al., 2014). Caldera-sourced silicic deposits occur throughout the Altiplano-Puna plateau, being less abundant and generally younger in the south than in the north (Guzmán et al., 2014; Kay et al., 2010; Schnurr et al., 2007). For example, the Cerro Blanco Volcanic Complex produced 4.2 ka BP ago, one of the most important Holocene eruptions in the world (Fernandez-Turiel et al., 2019). In the back-arc, monogenetic eruptive centres generated andesitic to basaltic andesitic, and minor shoshonitic, pyroclastic fall deposits and lava flows (Acocella, 2014; Báez et al., 2017; Grosse et al., 2017, 2020; Haag et al., 2019; Naranjo et al., 2018; Wörner, 2019) as well as some pyroclastic felsic deposits (Filipovich et al., 2020; Petrinovic et al., 2005; Petrinovic and Colombo Piñol, 2006).

Magmas of both constructional volcanoes and calderas ascend

through and interact with the thick continental crust, mainly associated with N-S, NW-SE and NE-SW regional left-lateral strike-slip fault zones. For example, López-Coranzulí, Calama-Olacapato-El Toro, Archibarca-Galán, and Culampajá-Farallón Negro fault zones (Acocella et al., 2011; Bonali et al., 2012; Lanza et al., 2013; Naranjo et al., 2018; Norini et al., 2013; Petrinovic et al., 2017; Tibaldi et al., 2017; Tibaldi and Bonali, 2018).

The mafic back-arc volcanic centres in the Puna plateau are isolated, loosely clustered or form small volcanic fields (Báez et al., 2017; Garzione et al., 2017; Maro et al., 2016b). The eruptions, mainly of Strombolian style, produced basaltic andesites to andesites with a high-K calc-alkaline signature, and less frequently basaltic trachyandesites and trachyandesites that are at or near shoshonite composition (Déruelle, 1991; Drew et al., 2009; Guzmán et al., 2006; Kay et al., 1994; Knox et al., 1989; Krallmann, 1994; Maro et al., 2017; Murray et al., 2015; Richards et al., 2006; Risse et al., 2013, 2008). Shoshonitic monogenetic volcanoes are distributed in the central Andean back-arc region of Peru (Carlier et al., 2005; Sandeman and Clark, 2004), Bolivia (Hoke and Lamb, 2007; Jiménez and López-Velásquez, 2008; Redwood and Rice, 1997; Salisbury et al., 2015), and Argentina (Déruelle, 1991; Guzmán et al., 2006; Kay et al., 1994; Krallmann, 1994; Schreiber and Schwab, 1991).

The direct partial melting of the primitive mantle cannot produce shoshonitic magmas, so an additional source of K₂O is needed. This enrichment of K₂O and K/Na within the mantle is currently explained by (Förster et al., 2019): 1) a two-stage fusion model of the “vein + wall rock” metasome (Foley, 1992), and 2) a single-stage recycled sediment/crust model (Campbell et al., 2014). The first model is the most suitable for the Andean potassic magmas (e.g., Carlier et al., 2005; Hoke and Lamb, 2007; Murray et al., 2015). In this framework, the potassic and ultrapotassic magmas are the earliest melts to be generated when the metasomatized lithospheric mantle undergoes heating and/or extension (Gibson et al., 1995). The study of shoshonitic volcanoes can provide insights regarding the mantle composition and mantle-crust

<figure><img src="image_4.png" /></figure>

**Fig. 1.** Location of studied San Jerónimo and Negro de Chorrillos and other shoshonitic volcanic centres in the Puna Plateau (triangles). The squares indicate other volcanoes mentioned in the text. The shaded-relief image is derived from ASTER GDEM ver.3 (https://ssl.spacesystems.or.jp/ersdac/GDEM/E/). Inset (a) shows the location in the Andes of the Northern Volcanic Zone (NVZ), the Central Volcanic Zone (CVZ), the Southern Volcanic Zone (SVZ), and the Austral Volcanic Zone (AVZ). Inset (b) displays the cratons and Palaeozoic terranes in southern South America (modified from Franceschinis et al., 2016).
interaction of discrete lithospheric mantle domains at continental subduction zones (Carlier et al., 2005; Jiménez and López-Velásquez, 2008). On the other hand, the use of xenoliths as proxies of the mantle composition is limited by their scarcity in this type of mafic volcanism (e.g., Chapman et al., 2015; Maro et al., 2016a; McLeod et al., 2013), which limits its use to know the mantle composition.

San Jerónimo (SJ), also called in the literature San Gerónimo, and Cerro Negro de Chorrillos (NdC), sometimes called Cerro Chorrillos, are two of the most representative back-arc monogenetic shoshonitic volcanoes in the Puna plateau (Déruelle, 1982, 1991; Ducea et al., 2013; Harmon et al., 1984; Hörmann et al., 1973; Kay et al, 1994, 1999; Knox

et al., 1989; Krallmann, 1994; Murray et al., 2015; Petrinovic, 2008; Schreiber and Schwab, 1991; Tannhäuser, 1906; Urquizo Furlán, 2012) (Fig. 2). Although the studies mentioned above contributed to advancing knowledge of the particular history and petrogenesis of these monogenetic shoshonitic volcanoes, their origin and age have remained controversial. SJ and NdC volcanoes are particularly well suited for a petrological study as they have erupted shoshonitic lavas with a wide variety of mineral contents and textures. Earlier studies have provided representative volcanological (Urquizo Furlán, 2012), petrographic and whole-rock and mineral chemistry data of SJ and NdC (Déruelle, 1982, 1991; Ducea et al., 2013; Harmon et al., 1984; Hörmann et al., 1973; Kay

<figure><img src="image_5.png" /><figcaption>**Fig. 2.** a) Satellite image (from Google Earth, August 3, 2016, CNES/AIRBUS) of study area showing the location of San Jerónimo and Negro de Chorrillos volcanic centres. b-e) Field photographs of San Jerónimo volcano from NW (b); Negro de Chorrillos volcano from NW (c); blocky lava flow of San Jerónimo (d); and blocky lava flow of Negro de Chorrillos (e).</figcaption></figure>
et al, 1994, 1999; Knox et al., 1989; Krallmann, 1994; Murray et al., 2015; Schreiber and Schwab, 1991; Tannhäuser, 1906). In this sense, we expand on these studies and aim to provide a more comprehensive overview of the volcanic history of SJ and NdC. We present new geomorphological, petrological, geochemical and dating constraints to further understand their magmatic plumbing system and the evolution of shoshonitic magmas in the Altiplano-Puna Plateau as in other geodynamic settings.

## 2. Geological setting

The study area is on the Calama-Olacapato-El Toro (COT) fault zone at the transition or limit between the northern and southern Puna (Fig. 1). The COT fault zone transversally crosses, from NW to SE, the volcanic arc and back-arc regions of the Puna plateau, at about 22–25°S. The COT is a NW–SE trending left-lateral transcurrent fault zone with inferred transtensive motions that, in the study area and during the late Pliocene–Quaternary, responds to a kinematic regime of a left-lateral strike-slip with a horizontal ~ E–W σ₁ and ~N–S σ₃ (Lanza et al., 2013). The overall left-lateral transtension of this fault zone created favourable tectonic conditions for the ascent of magma and the associated volcanic activity in the back-arc, developing volcanic centres such as San Jerónimo and Negro de Chorrillos (Fig. 1) (Acocella, 2014; Bonali et al., 2012; Norini et al., 2013; Petrinovic et al., 2017; Riller et al., 2001), which in turn are affected by numerous WNW- to NW-striking subvertical fault scars and associated aligned ponds and springs (Bonali et al., 2012; Lanza et al., 2013; Seggiaro et al., 2016). The magmatic activity along the COT lineament suggests that these lithospheric weakness zones control mantle melting by focusing thermal inputs and fluid ingress in the Puna lithospheric blocks, as was inferred for potassic and ultrapotassic melts in the large lineaments of the Altiplano (Carlier et al., 2005).

The volcanoes of San Jerónimo and Negro de Chorrillos are located 18 and 10 km west of the town of San Antonio de Los Cobres, respectively. The basement of San Jerónimo is formed by Miocene ignimbrites and the one of Negro de Chorrillos by Ordovician sediments, volcanic rocks and granitoids (Seggiaro et al., 2016). The edifice of San Jerónimo (24.2290°S 66.4985°W) consists of a monogenetic scoria cone, which closed a ravine in the south, forming an ephemeral lake (Vilela, 1969), and lava flows reaching 2–3 km to the north and ~9 km to the east from the vent. Negro de Chorrillos is also a monogenetic volcano (24.2715°S 66.4148°W), which in addition to a scoria cone, forms 2–5 km long lava flows to the north and east (Vilela, 1969), generated in three phases with effusive and explosive events (Urquizo Furlán, 2012). The second phase products are distinguished by the presence of phlogopite (Urquizo Furlán, 2012). Some authors have associated pyroclastic flows and ash fallout with this eruption (Seggiaro et al., 2016).

## 3. Methods

A palaeogeomorphological reconstruction of San Jerónimo and Negro de Chorrillos eruptions, based on fieldwork, was carried out to obtain the pre- and post-eruption terrain models following the methods developed by Rodriguez-Gonzalez et al. (2011, 2012). Field studies consisted of examining and identifying the different volcanic units (volcanic cone, lava flows and pyroclastic fall deposits) and the kind of relief below the volcanic deposits. The resulting information was implemented in a Geographical Information System (GIS), together with high-resolution optical images and present-day Digital Elevation Model (DEM) from Google Earth™, to proceed with the morphometric analysis. The GIS software used was TNTgis 2019 (MicroImages, Inc.). The generation of the DEM consisted of three general steps. The elevation point data of two areas containing the SJ and NdC volcanoes in Google Earth™ were extracted in the first step. The elevation point data extraction was performed with an online tool (Terrain Zonum Solution, http://www.zonums.com/gmaps/terrain.php). The coordinates (latitude and longitude

in WGS84 projection) and the elevation (z) in meters for each elevation point sampled were the extraction process's output. These values were then imported into a text file for further process. Once revised for outliers, the second step comprised the generation of a contour map (5 m interval). Finally, the third step consisted of constructing a DEM from the contour lines. A 5 m DEM resolution was appropriate for the accurate quantification of all the morphometric parameters of interest. The second step's contour map (i.e., post-eruption) was also used to generate the pre-eruptive DEM, modifying the contour lines to obtain the pre-eruptive surface based on fieldwork (Rodriguez-Gonzalez et al., 2010). The comparison of pre- and post-eruption surfaces allowed us to

<table><caption>**Table 1**
Morphometric parameters of San Jerónimo and Negro de Chorrillos volcanoes.</caption><thead><tr><th>Volcanic cone</th><th>Units</th><th>San Jerónimo</th><th>Negro de Chorrillos</th></tr></thead><tbody><tr><td>Crater</td><td></td><td></td><td></td></tr><tr><td>Crater rim maximum elevation</td><td>M</td><td>North 4885</td><td>Central 4950</td><td>South 4915</td><td>4425</td></tr><tr><td>Crater rim major axis (acr)</td><td>M</td><td>164</td><td>152</td><td>157</td><td>448</td></tr><tr><td>Crater rim minor axis (bcr)</td><td>M</td><td>144</td><td>83</td><td>117</td><td>350</td></tr><tr><td>Crater rim eccentricity (eer)</td><td></td><td>0.5</td><td>0.7</td><td>0.8</td><td>0.6</td></tr><tr><td>Crater major axis azimuth (theta)</td><td>Degrees</td><td>170</td><td>163</td><td>143</td><td>70</td></tr><tr><td>Crater depth (hcr)</td><td>M</td><td>3</td><td>3</td><td>4</td><td>75</td></tr><tr><td>Cone major axis (aco)</td><td>M</td><td></td><td>1846</td><td></td><td>1188</td></tr><tr><td>Cone minor axis (bco)</td><td>M</td><td></td><td>1340</td><td></td><td>1047</td></tr><tr><td>Cone eccentricity (eco)</td><td></td><td></td><td>0.7</td><td></td><td>0.5</td></tr><tr><td>Cone major axis azimuth (theta)</td><td>degrees</td><td></td><td>366</td><td></td><td>30</td></tr><tr><td>Cone height (h)</td><td>M</td><td></td><td>258</td><td></td><td>162</td></tr><tr><td>Cone volume (V)</td><td>km³</td><td></td><td>0.118</td><td></td><td>0.034</td></tr><tr><td>Cone volume DRE (V<sub>DRE</sub>)</td><td>km³</td><td></td><td>0.030</td><td></td><td>0.009</td></tr><tr><td>Area (A)</td><td>km²</td><td></td><td>1.904</td><td></td><td>0.699</td></tr><tr><td>Cone slope (beta)</td><td></td><td></td><td></td><td></td><td></td></tr><tr><td>Median</td><td>degrees</td><td></td><td>20</td><td></td><td>20</td></tr><tr><td>Mean</td><td>degrees</td><td></td><td>21</td><td></td><td>19</td></tr><tr><td>Minimum</td><td>degrees</td><td></td><td>0</td><td></td><td>0</td></tr><tr><td>Maximum</td><td>degrees</td><td></td><td>45</td><td></td><td>45</td></tr><tr><td>Standard deviation</td><td>degrees</td><td></td><td>8</td><td></td><td>8</td></tr><tr><td>Basement slope (pre-eruption)</td><td></td><td></td><td></td><td></td><td></td></tr><tr><td>Median</td><td>degrees</td><td></td><td>8</td><td></td><td>16</td></tr><tr><td>Mean</td><td>degrees</td><td></td><td>9</td><td></td><td>15</td></tr><tr><td>Minimum</td><td>degrees</td><td></td><td>0</td><td></td><td>0</td></tr><tr><td>Maximum</td><td>degrees</td><td></td><td>39</td><td></td><td>40</td></tr><tr><td>Standard deviation</td><td>degrees</td><td></td><td>6</td><td></td><td>6</td></tr><tr><td>Lava flow</td><td></td><td></td><td></td><td></td><td></td></tr><tr><td>Length (L)</td><td>m</td><td></td><td>9390</td><td></td><td>4520</td></tr><tr><td>Bottom width (wb)</td><td></td><td></td><td></td><td></td><td></td></tr><tr><td>Median</td><td>m</td><td></td><td>444</td><td></td><td>767</td></tr><tr><td>Mean</td><td>m</td><td></td><td>913</td><td></td><td>1034</td></tr><tr><td>Minimum</td><td>m</td><td></td><td>26</td><td></td><td>61</td></tr><tr><td>Maximum</td><td>m</td><td></td><td>3287</td><td></td><td>2437</td></tr><tr><td>Standard deviation</td><td>m</td><td></td><td>960</td><td></td><td>868</td></tr><tr><td>Area (A)</td><td>km²</td><td></td><td>8.801</td><td></td><td>4.631</td></tr><tr><td>Height (h) (thickness)</td><td></td><td></td><td></td><td></td><td></td></tr><tr><td>No. Profiles for calculation</td><td></td><td></td><td></td><td></td><td></td></tr><tr><td>Median</td><td>m</td><td></td><td>11</td><td></td><td>8</td></tr><tr><td>Mean</td><td>m</td><td></td><td>12</td><td></td><td>11</td></tr><tr><td>Minimum</td><td>m</td><td></td><td>3</td><td></td><td>2</td></tr><tr><td>Maximum</td><td>m</td><td></td><td>22</td><td></td><td>33</td></tr><tr><td>Standard deviation</td><td>m</td><td></td><td>5</td><td></td><td>8</td></tr><tr><td>Volume (V)</td><td>km³</td><td></td><td>0.081</td><td></td><td>0.015</td></tr><tr><td>Volume DRE (V<sub>DRE</sub>)</td><td>km³</td><td></td><td>0.061</td><td></td><td>0.011</td></tr><tr><td>Lava flow basis slope</td><td></td><td></td><td></td><td></td><td></td></tr><tr><td>Median</td><td>degrees</td><td></td><td>8</td><td></td><td>8</td></tr><tr><td>Mean</td><td>degrees</td><td></td><td>11</td><td></td><td>8</td></tr><tr><td>Minimum</td><td>degrees</td><td></td><td>0</td><td></td><td>0</td></tr><tr><td>Maximum</td><td>degrees</td><td></td><td>31</td><td></td><td>20</td></tr><tr><td>Standard deviation</td><td>degrees</td><td></td><td>7</td><td></td><td>6</td></tr></tbody></table>
estimate cone and lava flow volumes and morphometric parameters (Table 1). Surface modification by the pyroclastic sheet-like fall deposits was almost imperceptible because of their low magnitude, and hence their volumes were not estimated. These cone and lava flow bulk volume estimates were transformed into Dense Rock Equivalent (DRE) volumes considering a porous volume fraction of 75% for pyroclasts in the cones (Mangan and Cashman, 1996) and 25% for basaltic a'a' lava flows (Wolfe, 1988).

Three representative sites from the SJ lava flows and other three
from the NdC lava flows were sampled based on the suitability for K-Ar

dating as well as for petrological and geochemical studies (Fig. 3). The
identification number IGSN of samples was registered by SESAR: SJ-1
has IGSN IEJLF001T, SJ-2 is IEJLF001U, SJ-5 is IEJLF001V, NdC-3 is
IEJLF001W, NdC-6 is IEJLF001X, and NdC-7 is IEJLF001Y. We also
sampled the Alumbrera volcano (Fig. 1), near the town of Antofagasta de
la Sierra, a recent mafic eruption of the Puna only for dating to compare
the ages, i. e., CB3001-D registered with IGSN IEJLF001Z.

Petrography and modal mineralogy were performed by conventional
optical petrographic microscopy. For modal analysis, 1000 points per
sample were counted. All whole-rock samples were crushed and

<figure><img src="image_6.png" /><figcaption>**Fig. 3.** Deposits and structures of San Jerónimo (a) and Negro de Chorrillos (b) volcanoes. The eruptive phases are according to Urquizo Furlán (2012). The insets show topographic profiles of the palaeosurface before the eruption and the present-day topography.</figcaption></figure>
powdered using standard jaw crusher and agate milling equipment. Major and some trace elements (Ba, Cr, Ni, Rb, Sr, V, Y, Zn and Zr) were determined by X-ray fluorescence (XRF) spectrometry at the Scientific and Technological Centres of the University of Barcelona (CCiTUB), using a Philips PW2400 sequential X-Ray spectrometer with a Rh tube, by fusion of 2 g of powdered sample with lithium tetraborate to make glass beads and then analysed by XRF for major elements. Pressed powder pellets with 5 g of sample were used for trace element XRF analysis. Additional trace elements (Nb, Ta, Cs, Hf, Pb, Th, U and the rare Earth elements or REE) were analysed by High Resolution-Inductively Coupled Plasma-Mass Spectrometry (HR-ICP-MS, Element XR, Thermo Scientific) at the Laboratory of Geochemistry labGEOTOP (GEO3BCN, CSIC, Barcelona). A split of 0.1 g of each sample was digested with HNO₃:HClO₄:HF (2.5:2.5:5 ml, v/v), doubly evaporated to incipient dryness with the addition of HNO₃, and finally made up to 100 ml with 1% (v/v) HNO₃; before to HR-ICP-MS measures, a known amount of indium was added as the monitor. The precision and accuracy of analytical determinations were monitored using reference materials of the Geological Survey of Japan (rhyolite JR-3 and basalt JB-3).

The compositions of olivine, pyroxenes, phlogopite, feldspars, spinel and other minor minerals as well as of glasses were analysed in situ at the CCiTUB using an electron microprobe analyser (EMPA) Cameca SX50. An acceleration voltage of 20 kV and a current beam of 15–20 nA were used during all the analyses except for sodium. A current beam of 6 nA and an incident electron beam of 5–10 µm diameter were used to minimise the sodium loss. The instrument was calibrated using primary natural and synthetic standards. The quality of mineral analyses was reviewed, verifying that the structural formula gave a correct stoichiometry.

Treatment of whole-rock and mineral chemistry data was performed with GCDkit version 6.0 (Janousek et al., 2006, 2016). Mg number, Mg# = (100 x Mg/[Mg + Fe²⁺]), was calculated assuming 10% of the iron present in the melt as Fe³⁺ and the remaining 90% as Fe²⁺ (Gill, 2010). Europium anomalies (Eu/Eu*) were calculated as Eu/Eu* = (Eu)cn/[(Sm)cn x (Gd)cn]⁰.⁵; cn, normalized to chondrite from Boynton (1984). Values of εNd were used with published Nd isotope ratios and (¹⁴³Nd/¹⁴⁴Nd)CHUR = 0.512638.

Whole-rock major and trace element concentrations of this study (Fernandez-Turiel et al., 2020) were compiled in a larger dataset together with previously reported geochemical data for SJ and NdC lava flows (Déruelle, 1991; Hörmann et al., 1973; Kay et al., 1994; Krallmann, 1994; Murray et al., 2015; Schreiber and Schwab, 1991), and with other Puna shoshonitic centres: Los Gemelos and El Saladillo (Guzmán et al., 2006; Hausen, 1937; Hörmann et al., 1973; Kay et al., 1994; Krallmann, 1994), Llanaleri (Krallmann, 1994) and Cerro Barro Negro (Maro et al., 2016b) (Fig. 1). This shoshonitic dataset consists of 54 samples. As reference, the whole-rock major and trace element concentrations and Pb, Sr, Nd isotope data of the Neogene mafic volcanic rocks of northern Puna (Kay et al., 2010; Krallmann, 1994; Maro et al., 2016b) and southern Puna (Báez et al., 2017; Drew et al., 2009; Risse et al., 2013) were also compiled.

Magnetostratigraphy was used in combination with isotopic dating and geomorphological observations to provide the most reliable time framework for volcanic deposits' geological reconstruction. The preliminary definition of polarity sequences was carried out in the field using a portable fluxgate magnetometer. After these paleomagnetic works, two samples of lava flows were selected for age determinations: SJ-2 from San Jerónimo volcano and NdC-3 from Negro de Chorrillos volcano. Samples were K-Ar dated at the Laboratoire des Sciences du Climat et de l'Environnement (LSCE) at Gif-sur-Yvette (France), using the unspiked K-Ar method (Charbit et al., 1998). Groundmass splits from fresh samples were prepared (Guillou et al., 1996), and age calculations were based on the decay and abundance constants from Steiger and Jäger (1977). The samples were crushed and sieved to 0.250–0.125 mm size fraction and ultrasonically washed in HNO₃ to remove phases formed by alteration. Both K and Ar measurements were performed on

the groundmass, which is considered representative of the phase that crystallized during the lavas' solidification. Phenocrysts and xenocrysts, which may carry excess ⁴⁰Ar, were removed using heavy liquids of appropriate densities, magnetic separations and final inspection under binocular picking microscope. Potassium was analysed by ICP-MS. Argon was extracted from 1 to 2.5 g groundmass samples by radio-frequency heating induction in a high vacuum glass line and purified with titanium sponge and Zr-Al getters. Argon was analysed using a 180°, 6 cm radius, 620 V accelerating potential mass spectrometer working in a semi-static mode. The isotopic composition of Ar and Ar content were determined using an unspiked K-Ar technique (Charbit et al., 1998; Guillou et al., 1996).

# 4. Results

## 4.1. Morphometry

Table 1 shows the values of the most significant morphometric parameters, including areas and volumes, of cones and lava flows of SJ and NdC. SJ has a multiple cone with three closed craters roughly aligned NNW-SSE (Figs. 2–3), reaching a maximum altitude of 4950 m a.s.l. At the central crater. It covers a surface of about 1.9 km² and is 258 m high. The median slope of the cone flanks is 20, with a maximum of 45°. It developed on a palaeotopography where the median slope was 9°, although reaching 39° in some places (Fig. 3a). NdC has a horseshoe-shaped cone with a crater open towards the NE in the main direction of the basement's prevalent slope (Figs. 2–3). The maximum altitude is 4425 m a.s.l. The surface of this cone is 0.7 km², and it is 162 m high. The median of the NdC cone is also 20, with a maximum of 45° (Fig. 3b).

The lava flows of SJ and NdC consist of blocky-type lava and, exceptionally, a'a'-type lava with typical scoriaceous and very fragmented surface (Fig. 2d-e). The lava flows of SJ cover an area of 8.8 km² and those of NdC 4.6 km². The maximum lengths reached by these flows were 9.4 and 4.5 km from SJ and NdC vents, respectively. The thickness of the lava flows varies between 1 and 5 m, although sometimes several lava lobes overlap, reaching up to 20 and 30 m thickness in SJ and NdC, respectively.

Pyroclastic sheet-like fall deposits have not been considered in volume estimates due to their small occurrence. According to the methodology, the comparison between the generated pre- and post-eruption DEMs allow estimating the cone and lava flow bulk volumes, as represented schematically in Fig. 3 (insets). In this way, the bulk volumes estimated for cone and lava flows of SJ are 0.118 and 0.081 km³, respectively (Table 1). The corresponding Dense Rock Equivalent (DRE) volumes for SJ are 0.030 and 0.061 km³. For the NdC volcano, the bulk volume estimates are 0.034 and 0.015 km³ for cone and lava flows, and 0.009 and 0.011 km³ expressed as DRE volumes. Thus, the total volume DRE (Dense Rock Equivalent) was 0.091 and 0.020 km³ for SJ and NdC volcanoes, respectively. It is noted that the lava flow DRE volumes are greater than the cone DRE volumes in both cases, being the lava-to-cone ratio of 2.0 in SJ and 1.2 in NdC.

## 4.2. Petrography

Both SJ and NdC lavas have a general hipocrystalline, porphyritic-aphanitic, vesicular texture with macrocrysts (0.2 and up to 7 mm) of olivine, clino- and orthopyroxene, quartz, feldspars and, in some cases, phlogopite in a very fine-grained, commonly glassy, groundmass with microcrysts (<0.2 mm) of feldspars (plagioclase and sanidine), clino-pyroxene, olivine, opaque minerals and apatite. Xenoliths were not observed in the field.

The occurrence of phlogopite allows separating the samples into two groups. We have observed this mineral in samples of the second phase of NdC eruption (NdC-3 and NdC-6) but not in the sample of the third phase (NdC-7). Phlogopite was not observed in SJ's longest lava flow (samples SJ-1, SJ-2, SJ-5; Table 2). Samples without phlogopite present
<table><caption>**Table 2**
Petrography: modal analysis (% in volume).</caption><thead><tr><th rowspan="2">Sample</th><th colspan="4">Potassic series</th><th colspan="2">Ultrapotassic series</th></tr><tr><th>SJ-1</th><th>SJ-2</th><th>SJ-5</th><th>NdC-7</th><th>NdC-3</th><th>NdC-6</th></tr></thead><tbody><tr><td><em>Phenocrysts</em></td><td></td><td></td><td></td><td></td><td></td><td></td></tr><tr><td>Olivine</td><td>5.9</td><td>6.4</td><td>10.7</td><td>6.8</td><td>8.5</td><td>7.8</td></tr><tr><td>Pyroxene</td><td>4.2</td><td>2.9</td><td>2.4</td><td>3.4</td><td>0.9</td><td>2.0</td></tr><tr><td>Cpx (corona)</td><td>0.6</td><td>1.1</td><td>1.4</td><td>0.4</td><td>0.4</td><td>1.6</td></tr><tr><td>Phlogopite</td><td>-</td><td>-</td><td>-</td><td>-</td><td>6.9</td><td>3.8</td></tr><tr><td>Apatite</td><td>tr</td><td>tr</td><td>tr</td><td>tr</td><td>0.1</td><td>0.1</td></tr><tr><td><em>Xenocrysts</em></td><td></td><td></td><td></td><td></td><td></td><td></td></tr><tr><td>Quartz</td><td>1.4</td><td>1.2</td><td>2.3</td><td>0.8</td><td>0.5</td><td>2.2</td></tr><tr><td>Feldspar</td><td>0.1</td><td>0.1</td><td>0.7</td><td>0.1</td><td>0.6</td><td>0.9</td></tr><tr><td><em>Matrix</em></td><td></td><td></td><td></td><td></td><td></td><td></td></tr><tr><td>Feldspar</td><td>23.1</td><td>42.1</td><td>25.6</td><td>42.9</td><td>37.5</td><td>35.2</td></tr><tr><td>Apatite</td><td>1.1</td><td>1.3</td><td>1.0</td><td>2.0</td><td>3.1</td><td>2.3</td></tr><tr><td>Cpx</td><td>4.7</td><td>10.9</td><td>13.3</td><td>12.7</td><td>25.8</td><td>17.5</td></tr><tr><td>Opaque minerals</td><td>7.2</td><td>6.6</td><td>5.8</td><td>11.7</td><td>14.5</td><td>8.1</td></tr><tr><td>Glass</td><td>37.1</td><td>14.2</td><td>32.1</td><td>14.7</td><td>0.4</td><td>2.1</td></tr><tr><td>Vesicles</td><td>14.6</td><td>13.2</td><td>4.7</td><td>4.5</td><td>0.8</td><td>16.4</td></tr><tr><td>Total</td><td>100</td><td>100</td><td>100</td><td>100</td><td>100</td><td>100</td></tr></tbody></table>

porphyric (11–17 vol % macrocrysts, ranging ~200 µm to 5 mm in size), vesicular (4–15 vol % vesicles), and sometimes glomeroporphyric textures with pyroxenes enclosing spinel inclusions (Fig. 4a). They have a glassy to fine-grained, plagioclase-dominated microlite groundmass, showing a fluidal texture (more evidenced in SJ-2 and NdC-7), with macrocrysts of olivine (sometimes with reaction rims), feldspars, clinopyroxene and rare orthopyroxene. Plagioclase macrocrysts are often corroded and show a sieve texture defined by a core, a dusty or fritted zone, and a rim (Fig. 4b). The dusty texture consists of a diffuse zone that mantles the crystal’s interior, as defined by Tsuchiyama (1985). It varies in thickness, entering into the pre-existing crystal from the margin inward and sometimes extending throughout its core. An overgrowth of clear plagioclase invariably mantles the dusty zone. Sanidine is very scarce and occurs as macrocrysts of ~200 µm to 3 mm in size. Euhedral to subhedral olivine macrocrysts and microcrysts often show corrosion gulfs (Fig. 4c). Subhedral-euhedral pyroxene macrocrysts show variable sizes (0.2–6 mm). Olivine and pyroxene macrocrysts include tiny opaque minerals.

Phlogopite-bearing lavas are texturally similar to the previous ones, although the macrocryst content is higher (≈18 vol %), being olivine and phlogopite the main mineral phases. Phlogopite mainly occurs as macrocrysts (3.8–6.9 vol %, Table 2), with evident opacitic rims due to disequilibrium processes (Fig. 4d), and minor microcrysts. Sanidine occurs as corroded macrocrysts of ~200 µm to 3 mm in size (Fig. 4e) but is also very common in the groundmass as well as clinopyroxenes. Glass is much scarcer than in the previous group (0.4–2.1 vs. 14.2–37.1 vol %). There are also remnants of much-altered minerals, in part perhaps amphiboles.

One of the most striking features of the SJ and NdC lavas is the presence in the same sample of olivine and quartz macrocrysts. These peculiar quartz crystals, always in small quantities (0.5–2.3 vol %) (Table 2), do not show evidence of notable deformations (such as mosaic or wavy extinctions, intense fractures, etc.), but reaction coronas are common (Fig. 4f). The coronas consist of acicular clinopyroxene and glass rimming the quartz crystals. Where corrosion is scarce and the corona undeveloped, angular, automorphic or subautomorphic forms are observed.

### 4.3. Whole-rock geochemistry

The SJ and NdC samples of this work range from 53.19 to 57.44 wt % SiO₂, 6.31–7.72 wt % MgO, 3.56–5.98 wt % K₂O, 2.49–3.02 wt % Na₂O, and 1.02–2.12 wt % TiO₂. The volatile contents determined in the whole-rocks are 0.09–0.90 wt % LOI, and magnesium number (Mg#) ranges between 67 and 70, except for sample SH-7 with a Mg# of 60. All

these samples, on anhydrous basis, are classified as basaltic trachyandesite and trachyandesite in the total alkalis-silica (TAS) diagram (Le Bas et al., 1986) (Fig. 5a). While SJ samples are mainly in the sub-alkaline field, the NdC samples are mostly in the alkaline field (Fig. 5a) (Irvine and Baragar, 1971). In the SiO₂ vs. K₂O diagram, the samples plot in the shoshonite series field (Fig. 5b) (Peccerillo and Taylor, 1976). According to the K₂O and Na₂O contents, most of the samples are potassic (1 ≤ K₂O/Na₂O ≤ 2 wt %), except for some lavas belonging to NdC volcanic edifice (i.e., NdC-3 and NdC-6), which are ultrapotassic (K₂O/Na₂O ≥ 2 wt %, K₂O > 3 wt %, and MgO > 3 wt %) (Fig. 5c) (Foley et al., 1987; Gill, 2010). The two samples that are ultrapotassic are the two phlogopite-bearing samples. From the perspective of geotectonics, the studied rocks in the Yb–Ce plot reflect a continental magmatic arc setting (Gill et al., 2004) (Fig. 5d).

In a primordial mantle normalized diagram, all samples display enrichment in large-ion lithophile elements (LILE, K, Rb, Sr, Cs, Ba and light REE) with pronounced negative anomalies in some high-field strength elements (HFSE, i.e., Nb, Ta and Ti) and a positive anomaly of Pb and Li (Fig. 6a). Patterns of potassic and ultrapotassic lavas are similar in the general shape, but the last exhibit, in addition to the enrichment of K, slightly higher contents of Rb, Ba, Zr, and Hf.

REE patterns are similar for all lavas, with a steep slope pattern for light REE (LREE) ((La/Dy)ₙ, 12–14) that becomes practically flat in the Ho–Lu segment ((Ho/Lu)ₙ, 1.0–1.1), whereas the Eu anomaly is slightly marked (Eu/Eu*, 0.81 to 0.97) (Fig. 6b). In detail, ultrapotassic lavas of the NdC are slightly more enriched in all REE and, in the potassic lavas, sample NdC-7 overlaps considerably with the SJ lavas for La–Dy segment and with the NdC ultrapotassic lavas from Ho to Lu.

Harker diagrams show good negative correlations of Fe₂O₃t (Fig. 7a), MgO and CaO (Fig. 7b) with SiO₂. However, Al₂O₃, Na₂O, K₂O, TiO₂ and P₂O₅ plot poorly defined trends (not shown). Trace elements do not present simple trends (e.g., SiO₂ against Rb and Ba, Fig. 7c–d).

### 4.4. Mineral and glass compositions

Mineral and glass compositions are reported in an open dataset (Fernandez-Turiel et al., 2020). This section exposes the results obtained for olivine, pyroxene, phlogopite, feldspars, glass, and opaque minerals.

#### 4.4.1. Olivine

The olivine composition from SJ and NdC lava flows ranges from Fo₈₉ to Fo₇₁ (mol % forsterite component) (Fig. 8a–c). Two separate groups can be distinguished. The most forsteritic olivine (Fo₈₂-8₉) is found in the macrocryst cores of both potassic and ultrapotassic lavas. Forsterite content is lower (Fo₇₁-8₁) in the rims and some macrocryst cores of potassic lavas.

Nickel content shows a broadly positive correlation with forsterite content in olivine cores at Fo₈₂-Fo₈₉ (potassic series, 0.01–0.43 wt % NiO, n = 25; ultrapotassic series, 0.09–0.68 wt % NiO, n = 10) (Fig. 8a). This group is also characterized by lower MnO contents (<0.25 wt %) and slightly lower CaO contents (0.07–0.16 wt % CaO), compared to the olivine rims (potassic series, 0.32–0.45 wt % MnO, 0.13–0.29 wt % CaO, n = 6; ultrapotassic series, 0.36 wt % MnO; 0.16 wt% CaO, n = 1) (Fig. 8b–c).

Most olivine crystals are not in equilibrium with the whole rock (Fig. 8d). The most primitive olivines are cores in both the potassic and ultrapotassic series and, in some cases, show a Mg# exceeding the equilibrium line. The compositions in equilibrium (n = 14) were used to estimate temperatures by using the olivine-melt thermometer (eqn. 21; WFFeo = 0.9; spreadsheet Olivine_and_glass_03_14_18. xls) (Putirka, 2008). For melt composition, we used the corresponding whole-rock composition. Mineral-pairs were checked to ensure they satisfy the equilibrium condition K₉ (Mg–Fe)ol-liq = 0.30 ± 0.03 (Roeder and Emslie, 1970), and yielded temperatures of 1209–1303 °C for 7 kbar (~21 km) in anhydrous conditions. The olivines of potassic lavas gave 1209–1252 °C (n = 10), whereas those of the phlogopite-bearing
<figure><img src="image_7.png" /><figcaption>**Fig. 4.** Photomicrographs (plane-polarized light, PPL; cross-polarized light, CPL) of representative petrographic features of San Jerónimo and Negro de Chorrillos lava flows. (a) Orthopyroxenes (opx) forming a glomeroporphyric texture with inclusions of spinels (sample SJ-5, PPL). (b) Plagioclase with sieve texture defined by a core, a dusty or fritted zone, and a rim (sample SJ-5, CPL). (c) Olivine phenocryst with corrosion gulfs (sample SJ-5, PPL). (d) Phlogopite phenocrysts embayed and rimmed by opaque minerals (NdC-3, PPL). (e) Sanidine xenocryst with anhedral appearance due to the strongly rounded edges of the crystal (NdC-3, CPL). (f) Quartz xenocryst with the corona of glass and clinopyroxene (cpx) (sample SJ-5, PPL).</figcaption></figure>

ultrapotassic compositions reach 1283–1303 °C (n = 4).

### 4.4.2. Pyroxenes

Clinopyroxene macrocrysts, microcrysts, and coronas in SJ and NdC lava flows are classified (Morimoto, 1988) as augite-diopside and, only in SJ, as clinoenstatite. Orthopyroxene macrocrysts are enstatite and were only observed in the SJ potassic series (sample SH-5), forming cumulates. Clinopyroxene macrocrysts (Wo₃₃.₄₇ En₄₂.₅₃ Fs₉.₁₄) and microcrysts (Wo₄₀.₄₅ En₄₁.₄₉ Fs₁₀.₁₆) of the potassic series define a variation trend from 0.42 to 1.76 wt % TiO₂ and 1.63 to 5.22 wt % Al₂O₃. Despite the textural difference and petrogenetic significance, the

composition of coronas (Wo₃₈.₄₅ En₄₁.₅₂ Fs₇.₁₅), from 0.14 to 2.04 wt % TiO₂ and 0.05 to 5.60 wt % Al₂O₃, overlaps those of the macro and microcrysts. Orthopyroxene macrocrysts in this potassic series show high ferromagnesian contents (Wo₃.₉ En₇₆.₈₃ Fs₁₄.₁₈) and lower values and narrower ranges for TiO₂ (0.19–0.35 wt %) and Al₂O₃ (1.06–2.52 wt %). In the ultrapotassic series, clinopyroxene macrocrysts (Wo₃₉.₄₅ En₄₁.₅₂ Fs₈.₁₅), microcrysts (Wo₄₂.₄₄ En₄₆.₅₀ Fs₈.₁₀) and coronas (Wo₄₂.₄₄ En₄₇.₅₀ Fs₈.₁₁) overlap the compositions observed in the potassic series, although the variation ranges are narrower for TiO₂ (0.41–1.69 wt %) and Al₂O₃ (0.84–4.07 wt %). Clinopyroxene macrocryst and microcryst compositions are mostly outside the range of crystal-melt equilibrium
<figure><img src="image_8.png" /><figcaption>**Fig. 5.** Composition on anhydrous basis of San Jerónimo and Negro de Chorrillos samples using data of this work and from other studies (Déruelle, 1991; Hörmann et al., 1973; Kay et al., 1994; Krallmann, 1994; Murray et al., 2015; Schreiber and Schwab, 1991), and comparison with the shoshonitic centres of Los Gemelos and El Saladillo (Guzmán et al., 2006; Hausen, 1937; Hörmann et al., 1973; Kay et al., 1994; Krallmann, 1994), Llanaleri (Krallmann, 1994), and Cerro Barro Negro (Maro et al., 2016b). (a) Classification in the Total Alkali-Silica (TAS) diagram (Le Bas et al., 1986) with a dashed line separating the subalkaline from the alkaline field (Irvine and Baragar, 1971). (b) Classification in the $SiO_2$ vs. $K_2O$ plot (Peccerillo and Taylor, 1976) (Irvine and Baragar, 1971). (c) $Na_2O-K_2O$ classification of potassic and ultrapotassic rocks ($MgO > 3 wt \%$) (Foley et al., 1987; Gill, 2010). (d) Ce-Yb plot showing the studied samples compared with shoshonites fields of several regions with alkaline volcanism worldwide (Gill et al., 2004). The reference of the symbols for all plots is in (c). The compositional fields of the Neogene mafic volcanic rocks of northern Puna (Maro et al., 2016b) and southern Puna (Báez et al., 2017; Drew et al., 2009; Risse et al., 2013) are also indicated as reference.</figcaption></figure>

compositions for potassic and ultrapotassic samples, using an iron-magnesium distribution coefficient of 0.28 ± 0.08 for clinopyroxene (Putirka, 2008) (Fig. 9a). Corona compositions are mostly outside of the equilibrium field, but some are also inside (particularly in sample SJ-1). The P-T conditions for clinopyroxene-liquid equilibrium are sensitive to both temperature and pressure and calculated using an iterative simultaneous solution method (WFFeo = 0.9; anhydrous conditions; spreadsheet Clinopyroxene-P-T-9-26-16. xls) (Putirka, 2008). This geothermobarometer yielded pressure estimates between 7.1 kbar and 0.0 kbar and temperature between 1180 and 1100 °C (Fig. 9b). Pyroxene coronas in the ultrapotassic samples yielded 6-2 kbar, whereas coronas in the potassic samples are in the entire range of calculated pressures.

### 4.4.3. Phlogopite

Macrocysts and microcrysts of phlogopite occur only in the ultrapotassic series (Fig. 4d). The composition ($Fe^{2+}/(Fe^{2+}+Mg) < 0.33$; $Al^{IV} = 2-3$ apfu; n = 8)) classify them as phlogopite (Deer et al., 2013).

### 4.4.4. Feldspars

The ternary Ab-An-Or diagram for feldspars is shown in Fig. 10. In

the potassic series, macrocrysts (core and rim) and microcrysts are mostly within plagioclase solid-solution series, whereas macrocryst dusty zones show compositions between anorthoclase and sanidine solid-solution series close to hotter isotherms (Fuhrman and Lindsley, 1988) (Fig. 10). The cores of these macrocrysts are quite sodic ($An_{21-52}Ab_{44-70}Or_{3-25}$), while dusty zones are more potassic ($An_{11-63}Ab_{32-52}Or_{5-50}$), and their overgrowth rims are significantly more calcic ($An_{45-70}Ab_{28-48}Or_{2-7}$). Microcryst compositions ($An_{44-71}Ab_{27-47}Or_{2-14}$) overlap the compositions of the last rims.

Conspicuous by its absence is the lack of dusty zones on feldspar macrocrysts of the ultrapotassic series. The core and rim compositions of these macrocrysts and the microcrysts are more frequent within the sanidine and anorthoclase solid-solution series than in the plagioclase solid-solution series. The core compositions of these macrocrysts show a quite potassic trend ($An_{0.33}Ab_{21-52}Or_{22-79}$), largely overlapping the compositions of the macrocryst rims ($An_{5-26}Ab_{28-36}Or_{38-67}$) and microcrysts ($An_{6-45}Ab_{32-48}Or_{11-62}$).

### 4.4.5. Glass

Glass is more common in the potassic (14.2–37.1 vol %) than in the
<figure><img src="image_9.png" /><figcaption>**Fig. 6.** (a) Spidergram showing incompatible element enrichment patterns relative to the primitive mantle (Sun and McDonough, 1989). (b) Chondrite-normalized (Boynton, 1984) REE patterns. The grey field comprises the compositions of San Jerónimo and Negro de Chorrillos samples using data of this work and from other studies (Déruelle, 1991; Kay et al., 1994; Krallmann, 1994; Murray et al., 2015; Schreiber and Schwab, 1991), the shoshonitic centres of Los Gemelos and El Saladillo (Guzmán et al., 2006; Hausen, 1937; Hörmann et al., 1973; Kay et al., 1994; Krallmann, 1994), Llanaleri (Krallmann, 1994), and Cerro Barro Negro (Maro et al., 2016b).</figcaption></figure>

<figure><img src="image_10.png" /><figcaption>**Fig. 7.** Selected Harker diagrams of San Jerónimo and Negro de Chorrillos using data of this work and from other studies (Déruelle, 1991; Kay et al., 1994; Krallmann, 1994; Murray et al., 2015; Schreiber and Schwab, 1991), the shoshonitic centres of Los Gemelos and El Saladillo (Guzmán et al., 2006; Hausen, 1937; Hörmann et al., 1973; Kay et al., 1994; Krallmann, 1994), Llanaleri (Krallmann, 1994), and Cerro Barro Negro (Maro et al., 2016b). The compositional fields of the Neogene mafic volcanic rocks of northern Puna (Maro et al., 2016b) and southern Puna (Báez et al., 2017; Drew et al., 2009; Risse et al., 2013) are also indicated as reference.</figcaption></figure>

ultrapotassic (0.4-2.1 vol %, series (Table 2)). The glass compositions define two groups in the SJ and NdC lava flows: glasses in the groundmass and glasses in coronas together with clinopyroxene surrounding

quartz crystals. There are glass compositions of the potassic series in both groups but only from coronas in the ultrapotassic series, since microanalysis of the interstitial glass in the latter was not possible
<figure><img src="image_11.png" /><figcaption>**Fig. 8.** Forsterite (mol % forsterite component) versus (a) NiO (wt %), (b) CaO (wt %), and (c) MnO (wt %) in olivines from the studied eruptive potassic and ultrapotassic series. Note the change of trend at ~Fo<sub>82</sub>, a reflection of the bimodality of olivine compositions. Olivine cores display a broadly positive correlation of NiO with forsterite content and a negative correlation with MnO, constituting an olivine-dominated crystallization trend. Instead, olivine compositions from rims show lower forsterite (Fo < 82) and NiO (<0.15 wt %). (d) Rhodes diagram (whole-rock Mg# vs. olivine Mg#) for San Jerónimo and Negro de Chorrillos lava flows. The yellow field represents the range of crystal-liquid equilibrium compositions for K<sub>D</sub> (Mg–Fe)<sup>ol-liq</sup> = 0.30 ± 0.03 (Roeder and Emslie, 1970).</figcaption></figure>

<figure><img src="image_12.png" /><figcaption>**Fig. 9.** Mg# in whole-rock vs Mg# in mineral for clinopyroxene of San Jerónimo and Negro de Chorrillos lava flows (a); the yellow field represents the range of crystal-melt equilibrium compositions using an iron-magnesium distribution coefficient of 0.28 ± 0.08 for clinopyroxene (Putirka, 2008). The P-T estimates (b) were calculated by iterative simultaneous solution based on the clinopyroxene-melt equilibrium (Putirka, 2008).</figcaption></figure>

because of their practically fully crystallized groundmass (Fig. 11). The potassic series's interstitial glasses range mainly from basaltic–trachyandesite to the trachyte-trachydacite field in a TAS diagram, reaching even the rhyolitic field (Fig. 11). Instead, the group of corona glasses ranges from trachyte-trachydacite to rhyolitic TAS fields, being all corona glasses of the ultrapotassic series of rhyolitic composition.

### 4.4.6. Opaque minerals

Opaque minerals in the SJ and NdC lava flows are included in olivine, clinopyroxene and orthopyroxene macrocrysts in the potassic series and only in olivine in the ultrapotassic series, as well as micro-crysts in the groundmass in both cases. Macrocryst inclusions are mainly Cr-spinel, whereas titanomagnetite occurs as microcrysts in the groundmass. Only in one case, ilmenite was found in the groundmass (SH-1) and, in another case, rutile in apatite (SH-2). In the potassic
<figure><img src="image_13.png" /><figcaption>Fig. 10. Compositions of feldspars in the Ab-Or-An diagram. Isotherms for 1 kbar from Fuhrman and Lindsley (1988).</figcaption></figure>

<figure><img src="image_14.png" /><figcaption>Fig. 11. Total alkalis vs. SiO₂ ( TAS ) classification diagram ( Le Bas et al., 1986) for the interstitial glass in the groundmass and the glasses associated with coronas surrounding quartz crystals. Data are normalized to 100%. The shoshonitic melt, trachytic-trachydacitic melt and quartz end members are reported to explain the glass composition variation. See details in the discussion section.</figcaption></figure>

series, Cr-spinel crystals included in olivine (n = 39) show 1.02–10.95
wt % TiO2 and chromium numbers from 0.56 to 0.73 (Cr# = Cr/(Cr +
Al), mol); those included in clinopyroxene (n = 7) have 1.62–7.09 wt %
TiO2 and Cr# from 0.55 to 0.71, and the composition of one crystal
included in orthopyroxene is 2.24 wt % TiO2 and Cr# of 0.58 (Fig. 12).
In the matrix, the composition of titanomagnetites ranges from 3.31 to
17.36 wt % TiO2. The ilmenite in the matrix has 35.29 wt % TiO2 (n = 1)
and the rutile included in apatite 89.15 wt % TiO2 (n = 1). In the
ultrapotassic series, Cr-spinel crystals included in olivine (n = 12) show
1.45–1.92 wt % TiO2 and Cr# from 0.74 to 0.80 (Fig. 12). The compo-
sition of titanomagnetites ranges from 8.44 to 12.98 wt % TiO2 (n = 5) in
the matrix.

4.5. *Geochronology*

New K–Ar ages were determined on unaltered groundmass from fresh samples from SJ and NdC lavas. The age determined for sample SJ-2 of San Jerónimo is 144 ± 3 ka (±2σ error) (Table 3). The sample NdC-3 of Negro de Chorrillos provided an age of 51 ± 2 ka (±2σ error) (Table 3).

To help constrain these eruptions temporally within the younger
mafic volcanism of the Puna, we also determined the age of the Alum-
brera eruption, considered one of the younger volcanoes of the region
due to its well-preserved landforms (scoria cone and lava flows). The
sample CB3001-D of a lava flow of the Alumbrera volcanic centre pro-
vided an age of 37 ± 2 ka (±2σ error) (Table 3).
<figure><img src="image_15.png" /><figcaption>**Fig. 12.** Spinel compositions in the potassic and ultrapotassic series of the San Jerónimo and Negro de Chorrillos lava flows. (a) Spinel (Al + Mg)-Cr-(Fe + Ti) ternary diagram; arrow denotes the Fe-Ti variation trend of spinel compositions, along which Fe³⁺, Fe²⁺/(Mg + Fe²⁺) and TiO₂ increase at constant Cr/(Cr + Al) (Barnes and Roeder, 2001). (b) FeO–TiO₂–Fe₂O₃ ternary diagram for spinels. (c) Relationship between the content of forsterite in olivine and Cr# of spinels enclosed in this mineral; OSMA, olivine-spinel mantle array (Arai, 1994).</figcaption></figure>

<caption>**Table 3**</caption>

K–Ar ages of samples from the shoshonitic volcanoes of San Jerónimo and Negro de Chorrillos and the basaltic andesite Alumbrera volcano, one of younger mafic eruptions of the Puna. Age calculations are based on the decay and abundance constants from Steiger and Jager (1977). Errors on the age include the 1% uncertainty on the ⁴⁰Ar* content of the calibration standards.

<table><thead><tr><th>SAMPLE<br>Experiment</th><th>M. P.</th><th>K<br>(wt%) ± 1σ</th><th>Mass molten<br>(g)</th><th>⁴⁰Ar*(%)</th><th>⁴⁰Ar*10<sup>-13</sup><br>(mol/g) ± 1σ</th><th>Weighted mean<br>⁴⁰Ar*10<sup>-13</sup><br>(mol/g) ± 1σ</th><th>Age<br>ka ± 2σ</th></tr></thead><tbody><tr><td><strong>CB3001-D</strong></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td></tr><tr><td><strong>IGSN IEJLF001Z</strong></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td></tr><tr><td>8674</td><td>(+)</td><td>2.225 ± 0.022</td><td>2.00294</td><td>0.784</td><td>0.139 ± 0.058</td><td>0.143 ± 0.049</td><td>37 ± 2</td></tr><tr><td>8698</td><td></td><td></td><td>1.94134</td><td>1.855</td><td>0.151 ± 0.089</td><td></td><td></td></tr><tr><td><strong>NdC-3</strong></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td></tr><tr><td><strong>IGSN IEJLF001W</strong></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td></tr><tr><td>8586</td><td>(+)</td><td>5.105 ± 0.151</td><td>1.01915</td><td>1.041</td><td>4.310 ± 0.167</td><td>4.486 ± 0.113</td><td>51 ± 2</td></tr><tr><td>8602</td><td></td><td></td><td>1.02974</td><td>1.855</td><td>4.633 ± 0.154</td><td></td><td></td></tr><tr><td><strong>SJ-2</strong></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td></tr><tr><td><strong>IGSN IEJLF001U</strong></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td></tr><tr><td>8585</td><td>(+)</td><td>3.188 ± 0.032</td><td>1.04809</td><td>2.894</td><td>7.964 ± 0.198</td><td>7.946 ± 0.109</td><td>144 ± 3</td></tr><tr><td>8601</td><td></td><td></td><td>1.00288</td><td>2.891</td><td>7.939 ± 0.130</td><td></td><td></td></tr></tbody></table>

M. P. Magnetic polarity, (+) normal polarity.

# 5. Discussion

## 5.1. Morphometry

SJ and NdC eruptions produced Strombolian scoria cones and lava flows, the most common expression of the monogenetic volcanism in the Puna (Grosse et al., 2020; Haag et al., 2019). Tephra fallout deposits and pyroclastic flows are rare in this kind of volcanism. However, up to 10 cm thick ash-fall deposits identified in fluvial terraces were associated with the SJ eruption (Kay et al., 2008), and pyroclastic flows and ash fallout were related to the NdC eruption (Seggiaro et al., 2016).

Some morphometric parameters (cone height, h; cone major axis, a<sub>co</sub>, crater major axis, a<sub>cr</sub>; flank slope) and some geometrical aspect ratios (h/a<sub>co</sub> and a<sub>cr</sub>/a<sub>co</sub>) are commonly used to compare cinder cones. When compared to the Puna cinder cones (Haag et al., 2019), the cone heights of SJ (258 m) and NdC (162 m) are larger than the mean value (40 m). Their cone major axes (1846 and 1188 m for SJ and NdC) also exceed the mean (400 m). On the other hand, an ideal cinder cone is defined as having h/a<sub>co</sub> = 0.18, a<sub>cr</sub>/a<sub>co</sub> = 0.40 and a flank slope of ~31° (Fornaciai et al., 2012). Both SJ and NdC h/a<sub>co</sub> ratios are 0.14, whereas the a<sub>cr</sub>/a<sub>co</sub> ratio is 0.08 in SJ and 0.38 in NdC, and the slope median is 20° for the two edifices. The SJ edifice's strong deviation from an ideal
cone is because the SJ edifice is a multiple cone with three main craters (Figs. 2–3).

Lava flows of SJ and NdC have a fresh appearance showing well-preserved structures (levees, lobes, etc.). The maximum lengths reached by these flows (9.4 and 4.5 km for SJ and NdC, respectively) are in the range of 1–14 km observed in the southern Puna's mafic centres (Grosse et al., 2020). The maximum thicknesses of SJ and NdC (20 and 30 m, respectively) are in the high part of the average thicknesses (2–28 m) of these centres' lava flows. The southern Puna's lava flows cover areas between 0.6 and 44 km², whereas lava Flows of SJ and NdC covered 8.8 and 4.6 km², respectively.

The bulk and DRE volumes of cones are 0.118 and 0.030 km³ for SJ and 0.034 and 0.009 km³ for NdC. Compared with cone bulk volumes of the southern Puna ranging between 0.004 and 0.17 km³ (mean = 0.05 km³) (Grosse et al., 2020), the SJ cone is among the larger, whereas the NdC cone is close to the average. The bulk and DRE volumes of lava flows are 0.081 and 0.061 km³ for SJ, and 0.015 and 0.011 km³ for NdC. Compared with the southern Puna's lava flows, where bulk volumes are between 0.004 and 1.5 km³ (mean = 0.2 km³) (Grosse et al., 2020), SJ and NdC lava flow bulk volumes are among the lower values. Finally, the SJ and NdC whole centres (edifice plus lava flows) cover an area of 10.7 and 5.3 km², respectively. These values are between 0.1 and 44 km², the range observed for mafic centres in the southern Puna (Grosse et al., 2020).

The morphometric differences of monogenetic landforms result from eruption dynamics, pre-eruptive topography, and erosion processes (Di Traglia et al., 2014; Kereszturi et al., 2012; Rodriguez-Gonzalez et al., 2011, 2012). In the case of SJ and NdC, there are contrasts in the pre-eruptive surface slope with median values of 8 and 16°, respectively, and the larger time of exposition to erosion (144 ± 3 ka for SJ against 51 ± 2 ka for NdC). However, the main factor determining the SJ and NdC cones' distinctive morphometries was the eruption dynamics. The three craters of the SJ edifice indicate a complex eruptive history with shifts in vent location, whereas a breaching event characterises the NdC edifice morphology.

In summary, SJ and NdC eruptions produced scoria cones, lava flows and practically imperceptible pyroclastic fall deposits resulting from a Strombolian-type eruption with a Volcanic Explosivity Index (VEI) of 2 (Newhall and Self, 1982).

## 5.2. Petrography

Three of the most remarkable petrographic features of the studied samples are the presence of phlogopite in some lava flows, the mismatch of mineral assemblages (co-occurrence of olivine and quartz, the latter considered as xenocryst), and the common occurrence of disequilibrium textures. We have not observed phlogopite in SJ's longest lava flow (samples SJ-1, SJ-2, SJ-5; Table 2), although this mineral was previously referred at SJ volcano (Déruelle, 1991). Concerning the NdC eruption, Urquizo Furlán (2012) identified three eruptive phases, and we have observed phlogopite in the second phase (samples NdC-3, NdC-6; Table 2). Sample NdC 7 corresponds to a lava flow of the third phase without phlogopite. Thus, both eruptions had phases that erupted products with and without phlogopite.

Common examples of disequilibrium crystal textures are the sieve and dusty textures in plagioclase macrocrysts, where pervasive resorption may affect the entire crystal leading from a microscopic sieve texture (Stewart and Pearce, 2004) to a dusty texture consisting in a nebula of nearly sub-microscopic glass droplets (Tsuchiyama, 1985). The dusty texture (Fig. 4b) may be related to magma mixing, where preexisting sodic plagioclase is suddenly exposed to hotter, more mafic conditions (e.g., with more Ca), thereby reacting and dissolving to produce dusty rims (Hibbard, 1981; Kawamoto, 1992; Richer et al., 2004; Tsuchiyama, 1985; Viccaro et al., 2010). The thickness of the reaction rim is proportional to the duration of heating (Tsuchiyama, 1985).

Other disequilibrium crystal textures are due to peripheral dissolution and may range from subtle rounding of crystal edges to strongly rounded or embayed crystals. This process affected feldspars (Fig. 4e), olivine (Fig. 4c) and pyroxenes. Replacement reactions because of disequilibrium result in rims of opaque minerals around phlogopite phenocrysts (Fig. 4d) and the occurrence of coronas of clinopyroxene (augite) and glass around quartz xenocrysts (Fig. 4f). The latter feature is relatively common in the Puna back-arc mafic rocks (Coira and Kay, 1993; Maro et al., 2017; Risse et al., 2013) and evidences the quartz's reaction with mafic melt (Sato, 1975).

## 5.3. Whole-rock geochemistry

The sample compositions presented in this work overlap the previously reported data for SJ and NdC in the TAS diagram (Fig. 5a), classifying all considered samples as basaltic trachyandesites and trachyandesites. The shoshonitic assignment of SJ and NdC lava flows (Fig. 5b) was described in previous works (Déruelle, 1982, 1991; Ducea et al., 2013; Knox et al., 1989; Murray et al., 2015; Schreiber and Schwab, 1991). Considering the grouping in potassic and ultrapotassic rocks (Fig. 5c) the total dataset contains five ultrapotassic samples, all from NdC (Déruelle, 1991; Kay et al., 1994; Murray et al., 2015). The occurrence of phlogopite phenocrysts in these samples is confirmed in three cases, two in this work (Table 2) and one from Déruelle (1991; N16W). Their projection in the Yb-Ce plot indicates a continental magmatic arc setting, close to the Cascades shoshonites from the western USA (Gill et al., 2004) (Fig. 5d).

The shoshonites are distinct from the high-K calc-alkaline volcanic mafic rocks of the northern and southern Puna in their higher K and LILE concentrations, except for Cs (Fig. 6a). The negative anomalies of Nb, Ta and Ti also occur in the mafic rocks of the Puna. This pattern is common in arc magmas, and it is widely believed that it is the geochemical signature of peridotites metasomatized by LILE-LREE-bearing hydrous fluids in the mantle wedge above a subduction zone (Baier et al., 2008; Gill, 2010). In this setting, the negative Nb and Ta anomalies must be related to the very low solubility of Nb and Ta in aqueous fluids and the high solubility of these elements in aluminous clinopyroxene, rutile, and potentially some other phases, under relatively low pressures (10–20 kbar) in the shallow part of the subduction zone (Baier et al., 2008). Concerning the positive anomalies, Pb provides evidence for crustal assimilation. On the other hand, the enrichment in Li in Neogene volcanic rocks of CVZ is due to the combination of Palaeozoic basement assimilation with the mantle-derived signature (Meixner et al., 2020).

SJ and NdC samples show low Ba/Rb (5.6–11.4) and high Rb/Sr (0.1–0.2) ratios, typical ranges for lavas derived from phlogopite-bearing sources because Rb and Ba are compatible in phlogopite, whilst Rb, Sr and Ba are moderately compatible in amphibole (Wei et al., 2017). The enriched LREE patterns may be related to a thick continental crust (Winter 2010), while the flat heavy REE (HREE) tail suggests a garnet-free source, implying relatively shallow depths of magma formation in the upper mantle (Davidson et al., 2007a; Winter 2010). The lack of a significant Eu anomaly (Eu/Eu*, 0.81 to 0.97) suggests that plagioclase is not a major phase involved in fractional crystallization processes.

## 5.4. Mineral and glass compositions

### 5.4.1. Olivine

The compositional variations of NiO, CaO and MnO (Fig. 8a–c) as a function of forsterite indicate that the olivine macrocrysts crystallized from a compositionally related suite of magmas and are not xenocrysts from the mantle or other sources. With rising of forsterite, the slight fall in CaO agrees with experimental works explaining that both the forsterite content of the olivine and the magma CaO content are major controls on CaO in olivine (Libourel, 1999; Thompson and Gibson, 2000). Furthermore, CaO and forsterite contents are higher than those of
typical mantle xenocrysts (Liu et al., 2011; Straub et al., 2011; Thompson and Gibson, 2000). Besides, in both the potassic and ultra-potassic series, some olivine cores show a Mg# exceeding the equilibrium line, suggesting that they are antecysts (Davidson et al., 2007b; Larrea et al., 2014) (Fig. 8d). Instead, the olivines with Mg# under the equilibrium line evidence fractionation processes in a closed system. These features support the hypothesis that all analysed olivines are magmatic and not from mantle sources. Finally, the evolution trend defined by lower values of forsterite (Fo < 82, Fig. 8a–c) is consistent with a lower temperature and pressure crystallization, i.e., the rims crystallized at lower temperature and pressure compared to the cores. In these conditions, plagioclase and clinopyroxene co-crystallized with olivine. Crystallization of plagioclase at high crustal levels increases the CaO/Al₂O₃ ratio and alkalis (Na₂O + K₂O) content in the melt favouring the Ca incorporation in olivine rims (Libourel, 1999).

### 5.4.2. Pyroxenes

In both the potassic and ultrapotassic lavas, the compositions of clinopyroxene macrocrysts and microcrysts are mainly outside the crystal-melt equilibrium compositional range (Putirka, 2008) (Fig. 9a). This supports that these clinopyroxenes are the product of fractionation processes. The corona compositions are also mostly outside of the equilibrium field, but their texture demonstrates quartz xenocrysts' reaction with mafic melt (Sato, 1975).

The wide range obtained for pressure (up to 7.1 kbar, up to ~21.3 km) and the narrow range of temperature (1180–1100 °C) suggest a fast ascent rate (Fig. 9b). While the pyroxene coronas in the ultrapotassic lavas formed between 6 and 2 kbar, the coronas of the potassic lavas grew between 7.1 kbar and the surface. Clinopyroxene temperatures are ~30–~200 °C lower than those obtained for olivine (1209–1303 °C). This feature is interpreted as a consequence of clinopyroxene crystallizing after olivine, although no petrographic evidence supports this.

### 5.4.3. Phlogopite

The wide compositional variation in TiO₂ (2.89–8.85 wt %) and FeOt (5.14–9.53 wt %) in macrocrysts of phlogopite allows us to interpret them as phenocrysts, which are liquidus and near-liquidus phases, excluding their origin as mantle xenocrysts (Fritschle et al., 2013). The high Mg# (0.77–0.89) of phlogopites and their host rocks (0.70) indicate a deep origin, i.e., high P-T conditions in the range of 20–50 km depth for temperatures of 1100–1200 °C (Condamine and Médard, 2014; Gill, 2010). In this context, our magmas would be generated at the highest depth of spinel lherzolite's stability field (40–50 km). On its ascent and up to a depth of about 20 km, the phlogopite would form.

### 5.4.4. Feldspars

Dusty feldspar macrocrysts (Fig. 4b) are suggested to have formed by resorption of relatively albitic plagioclase xenocrysts due to the influx of the more calcic shoshonitic magma. Subsequently, more An-rich plagioclase crystallized from the hybrid host to form the thin high-calcium outermost rim. The occurrence of K-feldspar in phlogopite-bearing rocks results from the equilibria that include phlogopite, either with olivine or with pyroxene, where the potassium of phlogopite will remain as sanidine (Deer et al., 2013).

### 5.4.5. Glass

The wide range of compositional variation in glasses (Fig. 11) also evidences the mixing of magmas. We observe two compositional trends of glasses. The trend of groundmass glasses may result from both hybridization of melts and fractionation. Hybridization is represented by samples of the potassic series, with a high glass content (14–37%). Similar trends were observed in groundmass glasses of the Grizzly Lake complex of Yellowstone Volcano and were interpreted as the hybridization of mafic and felsic magmas (Morgavi et al., 2016). Instead, fractionation leads to rocks with low contents of more primitive interstitial glass (0–2%), i.e., samples of the ultrapotassic series. The second

compositional trend corresponds to glasses of coronas around quartz xenocrysts (Fig. 11). In the potassic series, there was a large quantity of melt to react and the result was a wide range of compositions. However, in the ultrapotassic series, the melt was scarce and more evolved, resulting in a silica-rich glass. The two before mentioned compositional trends of glasses converge in the trachyte-trachydacite field of the TAS diagram (Fig. 11), i.e., the composition of the felsic end-member involved in the mixing process. This composition is not uncommon in the region. Ignimbrites of the Puna have a composition similar to the postulated trachyte-trachydacite composition of these felsic melts (Murray et al., 2015).

### 5.4.6. Opaque minerals

The compositions of Cr-spinels and titanomagnetites in a ternary plot of (Al + Mg)–Cr-(Fe + Ti) (Fig. 12a) show that the potassic and ultrapotassic series follow the Fe–Ti variation trend of spinel compositions, along which Fe³⁺, Fe²⁺/(Mg + Fe²⁺) and TiO₂ increase at constant Cr/(Cr + Al) (Barnes and Roeder, 2001). However, the Cr content is slightly higher in the spinels enclosed in the ultrapotassic series's olivines (Fig. 12b–c). On the other hand, the compositions of Cr-spinels enclosed in olivine (i.e., forsterite content in olivine vs. Cr# of spinels) plot outside the olivine-spinel mantle array (OSMA) (Fig. 12c) (Arai, 1994).

### 5.5. Geochronology

The new groundmass K–Ar ages of San Jerónimo (144 ± 3 ka) and Negro de Chorrillos (51 ± 2 ka) confirm that NdC is younger than SJ, but both ages are significantly younger than those published before. San Jerónimo volcano was previously dated at 0.78 ± 0.10 Ma (Aquater, 1979; in Coira and Paris, 1981; Linares and Gonzalez, 1990), and Negro de Chorrillos three times with resulting ages of 0.45 ± 0.10 Ma (Linares and Gonzalez, 1990; the original reference is missing in the reference list of this catalogue of ages), 0.20 ± 0.15 Ma (Schwab and Lippolt, 1974), and 0.20 ± 0.08 Ma (Aquater, 1980; in Coira and Paris, 1981; Linares and Gonzalez, 1990). The new ages are younger because the previous ages were obtained from whole-rock while the groundmass provided the new ones. Potassium from older feldspar xenocrysts can cause older whole-rock ages.

The new ages of San Jerónimo and Negro de Chorrillos volcanoes are not enough to decipher the association of pyroclastic density currents (Chorrillos section) and also of reworked ash deposits interbedded in alluvial fans (Corte Blanco section) that have been associated with Negro de Chorrillos volcano (Seggiaro et al., 2016). Lava flows of San Jerónimo eroded the former primary deposits. This fact leads Seggiaro et al. (2016) (p. 475) to deduce that the San Jerónimo volcano had eruptive pulses before and after the Negro de Chorrillos eruption. Dating results confirm the relative chronology between the two eruptions and leave open the discussion about the origin of both deposits. Tentatively, we advance the hypothesis that pyroclastic flows could be related to an unknown eruption before San Jerónimo and Negro de Chorrillos eruptions, while the reworked basaltic ashes east of San Antonio de los Cobres village must be associated with an eruption of greater explosivity than those of SJ and NdC. In the area studied, if the discussed section is of Upper Quaternary age, the most probable candidate is the Tuzgle volcano (Fig. 1), which is located 30 km NW of this section.

Considering the shoshonitic eruptive centres of the Puna, the Cerro Barro Negro is possibly the oldest event recognized because its products are above ignimbrites of 6.4–6.6 Ma (Maro and Caffe, 2013, 2016). Among the young shoshonitic volcanoes, San Jerónimo is the oldest (144 ± 3 ka), followed by Negro de Chorrillos (51 ± 2 ka), and Los Gemelos volcanoes that were formed at around 35 ka, as inferred from the age of lacustrine sediments overlying the lava flows (Guzmán et al., 2006). The age of El Saladillo is unknown.

Although there are still many undated mafic volcanic centres in the Puna, existing data suggest that this activity was restricted in the northern Puna to the Late Miocene-Pliocene interval, with a probable
peak in the Late Miocene (Maro and Caffe, 2016). Instead, the peaks of mafic activity were at 5–4, 3–1 and <1 Ma in the southern Puna (Grosse et al., 2020; Risse et al., 2008; Schoenbohm and Carrapa, 2015), where this monogenetic volcanism is particularly abundant (Grosse et al., 2020). This temporal displacement of the back-arc mafic volcanism from north to south is interpreted as a result of variations in the geodynamic setting, changing slab dips and delamination of the thickened lower crust and lithosphere (Kay and Coira, 2009; Maro and Caffe, 2016).

The excellent preservation of geological formations and structures is a noteworthy characteristic of the Puna due to its aridity (Aulinas et al., 2015). This feature gives a fresh appearance to many recent mafic volcanic edifices suggesting a Late Quaternary age, including Holocene (Báez et al., 2017; Déruelle, 1991). In this framework are the young shoshonitic volcanoes, but also the mafic centres along the western margin of the Arizaro Salar with ages of 0.08–0.22 Ma (Schoenbohm and Carrapa, 2015), the Peinado region with ages of 0.151–0.215 Ma (Grosse et al., 2020), and the region north of Antofagasta de la Sierra with ages of 0.03 and 0.34 Ma (Risse et al., 2008). The Alumbrera Volcano (37 ± 2 ka), located south of this town, is one of the younger basaltic andesite eruptions recorded. In summary, the mafic volcanism of the Puna shows a low recurrence at the Human time scale (Báez et al., 2017; Grosse et al., 2020; Maro et al., 2016b; Risse et al., 2013, 2008).

## 5.6. Petrogenesis

The long-held hypothesis on the origin of back-arc mafic magmatism in the Puna, including the shoshonitic volcanism, suggest that it is the result of partial melting in the mantle as a consequence of continental lithospheric delamination or foundering (Beck et al., 2015; Coira et al., 1993; de Silva and Kay, 2018; Heit et al., 2014; Kay and Mahlburg Kay, 1993; Kay et al., 1994; Maro et al., 2016b; Murray et al., 2015). This model is supported by seismic tomography images of the mantle wedge (Bianchi et al., 2013; Calixto et al., 2013; Delph et al., 2017; Ibarra et al., 2019; Liang et al., 2014; Mulcahy et al., 2014; Schurr et al., 2006) and by thermomechanical modelling (Sobolev and Babeyko, 2005). It should be noted that only a very minor part of this magmatism is shoshonitic (Maro et al., 2017).

The primitive mantle-normalized incompatible element ratio diagram shows high enrichment in large-ion lithophile elements (LILE, e.g., Rb, Ba, Th, and K), with remarkable depletion in high field-strength elements (HFSE, e.g., Nb, P, Ti) in the studied shoshonites (Fig. 6a). This pattern is similar for the high-K calc-alkaline mafic rocks of the Puna, showing the latter higher K and LILE concentrations, except for Cs. The metasomatism of peridotites can explain this pattern by LILE-LREE-bearing hydrous fluids in the mantle wedge above a subduction zone (Baier et al., 2008; Gill, 2010). On the other hand, flat HREE-Y pattern (normalized to primitive mantle) evidences a spinel lherzolite as the mantle source in this scenario, i.e., a source without garnet (Bédard, 2005; Gill, 2010; Green et al., 2000; LaTourrette et al., 1995; Sun and Liang, 2013; Zhang et al., 2019) (Fig. 6b). This model disagrees with the pyroxenitic source (±garnet-bearing) interpreted previously (Ducea et al., 2013; Murray et al., 2015) and precludes other models based on garnet peridotite (Coira and Kay, 1993; Kay et al., 1994; Knox et al., 1989). Spinel lherzolite is stable at a lesser depth in the mantle than garnet lherzolite, in the range of 20–50 km depth for temperatures of 1100–1200 °C (Condamine and Médard, 2014; Gill, 2010). Based on the crustal thickness of this part of the Puna and the experimental constraint, these magmas last equilibrated at Moho depths at around 50 km (the highest depth of spinel lherzolite stability field). Besides, considering concentrations of 0.05 wt % K₂O and 0.07 wt % TiO₂ for continental spinel lherzolite average (Maaløe and Aoki, 1977), the K₂O/TiO₂ ratio normalized to source composition for shoshonitic samples ranges between 2.23 and 6.81 (n = 49), which indicates a percentage of partial melting lower than 5% (Hirschmann et al., 1999). This threshold of 5% is only exceeded for some samples of shoshonitic mafic centres of the northern and southern Puna.

There is no simple way to explain the transition from primary melts generated in the mantle wedge to shoshonitic magmas. One way can be through the interaction of mantle-derived magma with felsic lower crustal rocks to produce evolved shoshonitic compositions (Richards et al., 2006). Another possibility could be the transition from calc-alkaline to shoshonitic magma compositions due to a general increase of magmatic temperatures and a decrease of H₂O activity (aH₂O) in response to extensional tectonics and related increase of heat flow and declining influence of slab-derived fluids (Beermann et al., 2017). Experimental works point out the feasibility to obtain trachytic melts from K-rich basaltic andesites, the latter formed by melting of a metasomatized phlogopite-bearing mantle in the spinel stability field (Condamine and Médard, 2014). A mantle source consisting of mica-bearing peridotite was already proposed by Déruelle (1991). It should be noted that there is no evidence of mantle phlogopite, spinel, or any other phase of this origin in the samples studied. In any case, the mantle-derived magma evolved and was modified by later processes within the crust, something on which all models of mafic magmatism of the Puna agree (Coira et al., 1993; Déruelle, 1991; Guzmán et al., 2006; Knox et al., 1989; Maro et al., 2017; Risse et al., 2013).

Geochemical data of SJ and NdC shoshonites suggest the development of fractional crystallization processes and the interaction of the shoshonitic melts with others more evolved during the ascent to the surface. Fractional crystallization trends are evidenced by Harker diagrams showing negative correlations of FeO_t (Fig. 7a), MgO, CaO (Fig. 7b), and TiO₂ with SiO₂. However, Al₂O₃, Na₂O, K₂O (Fig. 5b) and P₂O₅ plot in poorly defined trends possibly due to mixing with a more felsic magma. The variations in trace elements also suggest a combination of fractional crystallization and magma mixing to explain the observed variations. For example, Rb and Ba trends show complex variations with SiO₂, the most evolved terms coinciding with the phlogopite-bearing lavas (Fig. 7c–d). Thus, despite the mafic lavas being affected by mixing with more felsic melts, bulk rock geochemistry supports that the lavas with phlogopite represent the most evolved terms reached by fractional crystallization within the mafic melts. This fractional crystallization is coherent with Rb and Ba's partition coefficients for phlogopite (LaTourrette et al., 1995). Similar behaviours are observed for P₂O₅ and Zr variations against SiO₂, in these cases associated with the crystallization of apatite and zircon, respectively.

Fig. 13 presents published Pb, Sr and Nd isotope data of shoshonitic rocks of the Puna (Déruelle, 1991; Guzmán et al., 2006; Kay et al., 1994; Maro et al., 2017; Murray et al., 2015), shown relative to other mafic volcanic rocks of the northern and southern Puna (Drew et al., 2009; Kay et al., 2010; Krallmann, 1994; Maro et al., 2016b; Murray et al., 2015; Risse et al., 2013). The ²⁰⁶Pb/²⁰⁴Pb, ²⁰⁷Pb/²⁰⁴Pb, ²⁰⁸Pb/²⁰⁴Pb ratios of shoshonites range from 18.42 to 18.75, 15.40–15.70, and 38.45–39.20 (n = 9), the ⁸⁷Sr/⁸⁶Sr ratios are between 0.70633 and 0.70827 (n = 17), and their ¹⁴³Nd/¹⁴⁴Nd values are 0.51223–0.51255 (n = 13) and εNd −1.72 to −7.96. These isotopic compositions overlap or are close to the values of other mafic centres of the Puna (Fig. 13).

The Pb isotopic ratios are enriched relative to the NHRL (Northern Hemisphere Reference Line), pointing out a crustal component's contribution. However, they do not vary with silica or other element concentrations (e.g., Sr, Nd), which is interpreted as mixing end-members having the same isotopic signature for Pb. Several studies have outlined domains within the central Andean orocline (13–28°S) based on radiogenic Pb (Aitcheson et al., 1995; Mamani et al., 2008, 2010; Wörner et al., 2018). These domains would correspond to terranes that were amalgamated to the western Gondwana margin (Ramos, 2008). The Pb isotope ratios of shoshonitic rocks are slightly lower than those of the southern's Puna mafic volcanic rocks (Fig. 13a), suggesting another terrane's involvement. We speculate that Pb ratios of shoshonites reflect the Pampia terrane, whereas the values of mafic volcanoes from the southern Puna are closer to those of the Antofalla domain/-terrane (Mamani et al., 2010) (inset b in Fig. 1).

The shoshonitic sample set shows narrower ranges for Sr and Nd
<figure><img src="image_16.png" /><figcaption>**Fig. 13.** Sr-Nd-Pb diagrams for shoshonitic volcanic centres of the Puna (data from Déruelle, 1991; Guzmán et al., 2006; Kay et al., 1994; Krallmann, 1994; Maro et al., 2017; Murray et al., 2015). The compositional fields of the Neogene mafic volcanic rocks of northern Puna (Kay et al., 2010; Krallmann, 1994; Maro et al., 2016b) and southern Puna (Drew et al., 2009; Risse et al., 2013) are also indicated as reference.</figcaption></figure>

isotope ratios compared to northern and southern Puna mafic volcanoes, although they mostly overlap with these distributions (Fig. 13b–f). The general positive trend of $^{87}$Sr/$^{86}$Sr isotope ratio against silica of all Puna mafic rocks compiled reveals a typical behaviour of the assimilation and

fractional crystallization (AFC) process (DePaolo, 1981) (Fig. 13c), which increases the $^{87}$Sr/$^{86}$Sr ratio values by crustal assimilation of older material with higher $^{87}$Sr/$^{86}$Sr compositions during differentiation (González-Maurel et al., 2019). The general variation of $^{87}$Sr/$^{86}$Sr vs. Sr
(Fig. 13d) is also compatible with the AFC process (González-Maurel et al., 2019; Hildreth and Moorbath, 1988). On the other hand, the assimilation during turbulent ascent (ATA) process (Huppert et al., 1985) has been proposed to explain the evolution of these mafic magmas (Guzmán et al., 2006; Maro et al., 2016b). In this process, the least-evolved magmas may be the most contaminated, and it is documented by reverse trends in SiO₂ and Sr vs. ⁸⁷Sr/⁸⁶Sr diagrams (González-Maurel et al., 2019). These trends are not dominantly depicted in Fig. 13c-d.

Crustal contamination by magma mixing in San Jerónimo and Negro de Chorrillos shoshonitic magmas is corroborated by physico-chemical disequilibria in crystals and glass. Accordingly, both volcanic centres must be modelled as an open system in which external melts were introduced in the shoshonitic magma. A wide variety of microscopic disequilibrium textures provides unequivocal evidence of the interaction between melts and the magmatic system's inability to reach equilibrium conditions (Streck, 2008). Thus, we find a mismatch of mineral assemblages (e.g., the co-occurrence of olivine and quartz), crystal textures (e.g., dusty or resorbed textured macrocrysts, replacement reactions), and a wide range of compositional variation in the glass. This last feature suggests the contamination by trachytic-trachydacitic melts.

These interactions with pre-existing trachyte-trachydacite crustal melts were relatively limited, e.g., little to no inherited minerals (quartz, feldspars) and a slight change in bulk rock geochemistry. This type of magmatic evolution is typically ascribed to polybaric differentiation within a trans crystal column of magmatic mush (Cashman et al., 2017; Cashman and Edmonds, 2019), and the magmas must be removed from the AFC site and erupted immediately to preserve correlations between indices of differentiation and isotopic ratios (Davidson and de Silva, 1995). According to recent numerical simulations, these mixing processes are produced in a short time, in the order of hours or days (Montagna et al., 2015; Morgavi et al., 2019; Perugini et al., 2010). The wide range obtained for pressure (up to 7.1 kbar, up to ~21.3 km) and the narrow range of temperature (1180–1100 °C) also supports a fast ascent rate (Fig. 9b).

These considerations are common to the origin of shoshonitic and high-K calc-alkaline magmas in the Puna, and they do not resolve how to obtain a shoshonitic magma in specific and rare locations. We suggest that an additional geodynamic factor could be involved. The shoshonitic centres are approximately located in the intersection of the N-S suture between the terranes of Antofalla and Pampia with the major NW-SE fault systems of López-Coranzúl (Cerro Barro Negro volcano) and Calama-Olacapato-El Toro (San Jerónimo, Negro de Chorrillos, El Saladillo and Los Gemelos volcanoes) (Fig. 1). The Antofalla-Pampia suture is a significant weakness zone developed in Paleozoic times along the southwestern margin of Gondwana (Augustsson et al., 2015; Beck and Zandt, 2002; Heredia et al., 2018; Omarini et al., 1999; Ortiz et al., 2019; Ramos, 2010, 2018). Although its precise location on its entire length is still a matter of debate, this suture is essentially sub-parallel to the modern Andean chain in the central Andes. This hypothesis would represent a mechanism by which a preferential terrane boundary distribution could develop a small volume of potassic melts (precursors of the shoshonitic rocks), while the major NW-SE fault systems would favour their ascent to the surface.

Finally, potassic and ultrapotassic (i.e., phlogopite-bearing) magmas in the same eruption are interpreted as pulses of the same magma feeding the volcano. Petrographic and chemical differences are possibly due to changes in the interaction time with the felsic magma. A longer interaction during mixing, within the eruptive time window of hours or days, allowed a practically fully crystallized groundmass in the ultrapotassic rocks, reducing the quantity of glass and forming phlogopite.

## 6. Conclusions

San Jerónimo and Negro de Chorrillos shoshonitic volcanoes are small monogenetic edifices mainly built by effusive activity, resulting

from a Strombolian-type eruption with a Volcanic Explosivity Index (VEI) of 2 (Newhall and Self, 1982). Unspiked K-Ar ages (groundmass) yield 144 ± 3 ka for San Jerónimo and 51 ± 2 ka for Negro de Chorrillos eruptions.

Volcanic products of potassic and ultrapotassic (i.e., phlogopite-bearing) compositions occur in both eruptions. The main petrographic differences are the occurrence of phlogopite and the low glass content in the ultrapotassic rocks. Geochemically, the lava flows of SJ and NdC exhibit the same spinel lherzolite mantle source composition, which is stable in the range of 20–50 km depth for temperatures of 1100–1200 °C. According to the crustal thickness of this part of the Puna, our magmas would be generated at the highest depth of lherzolite’s stability field with spinel (40–50 km). The phlogopite would form on its ascent and up to a depth of about 20 km.

Once in the crust, the K-rich magma interacted with a trachytic-trachydacitic melt, on the order of hours or days, producing minor contamination of the mafic magma. This mixing is evidenced petrographically and geochemically, e.g., mismatch of mineral assemblages (co-occurrence of olivine and quartz), disequilibrium crystal textures (dusty zones in plagioclase, embayments in feldspars), and a wide range of compositional variation in the glasses. The presence of potassic and ultrapotassic rocks in the same eruption is interpreted as pulses of the same magma but with differences in interaction time during magma mixing. A longer interaction allowed a practically fully crystallized groundmass in the ultrapotassic rocks.

The scarcity and location of shoshonitic centres, compared with other mafic volcanic centres of the northern and southern Puna, suggest that a preferential N-S terrane boundary distribution could develop a small volume of potassic melts (precursors of the shoshonitic rocks), while the major NW-SE fault systems would favour their ascent to the surface. The eruptive histories of SJ and NdC may give insights on the past and future behaviour of similar volcanoes in other back-arc regions of large plateaus as the Altiplano in the CVZ or the Tibetan plateau.

### CRediT author statement

J.L. Fernandez-Turiel: Conceptualization, Investigation, Writing – review & editing, Investigation, data curation. J. Saavedra: Conceptualization, Methodology, Investigation, Writing – original draft preparation. F.J. Perez-Torrado: Conceptualization, Methodology, Investigation. A. Rodriguez-Gonzalez: Conceptualization, Investigation, Visualization. M. Rejas: Resources (geochemical analyses). H. Guillou: Resources (geochronological analysis). M. Aulinas: Investigation, formal analysis.

### Declaration of competing interest

The authors declare that they have no known competing financial interests or personal relationships that could have appeared to influence the work reported in this paper.

### Acknowledgements

Financial support was provided by Project QUECA (MINECO, ref. CGL 2011-23307). This study was carried out in the Research Consolidated Groups GEOVOL (Canary Islands Government, ULPGC) and GEOPAM (Generalitat de Catalunya, 2017 SGR 1494). We acknowledge the assistance in the analytical work of labGEOTOP Geochemistry Laboratory (infrastructure co-funded by ERDF-EU Ref. CSIC08-4E-001) of GEO3BCN, CSIC, and EPMA Laboratories of CCiTUB (X. Llovet), and D. Fernández, her support on reference management. We thank the collaboration of P. E. Urquizo Furlán. We are grateful to the editor P. Samaniego for the effective handling of the manuscript and P. Grosse and S. Guzmán for providing insightful comments on a preliminary version of this manuscript.
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