The Western Coastal Range Domain comprises the metamorphic rocks that crop out mostly in the western flank of the Coast Range (Figure 2). Two units with different lithological and structural characteristics have been defined by Duhart (1999). The Estaquilla Metabasites correspond to nematoblastic mafic schists with N-MORB geochemical signature (Crignola et al., 1997) and Cretaceous to early Cenozoic U-Pb zircon crystallization ages (Hervé et al., 2016) metamorphosed at greenschist facies. The Río Llico Micaschists are composed of fine interlayers of metasandstones and metapelites with greenschist facies mineral associations of quartz-white mica-albite and quartz-albite-chlorite interleaved with minor NW-SE trending nematoblastic to granonematoblastic greenschist bodies with an actinolite-epidote-chlorite-quartz-albite assemblage (Duhart, 1999). Hervé et al. (2016) obtained a ca. 330 Ma zircon U-Pb maximum depositional age based on a micaschist of this unit (Figure 2A).

The Eastern Coastal Range Domain is composed of the Zarao Metavolcanic Rocks (Vildoso, 2017) and the neighboring garnet-bearing micaschist (Palape, 2020), which we define as the Cajonmo Garnet Micaschist. The metavolcanic rock unit is constituted by the Zarao metarhyolite and the spatially associated mafic schists (Figure 2). The crystallization age of the igneous protolith of the Zarao metarhyolite is constrained by 396.7 ± 1.3 Ma U-Pb age (Duhart et al., 2001) and 393 ± 3 Ma U-Pb SHRIMP age (Hervé et al., 2018) and corresponds to the oldest unit in the area. The metarhyolite has been interpreted to be either formed in an ocean island environment (Quezada et al., 2018) or a continental forearc rift setting (Rapela et al., 2021). In the Zarao area (Figure 1B), the aeromagnetic anomaly has a NE-SW strike, which differs with the NW-SE regional pattern of magnetic anomalies (Godoy and Kato, 1990).

Located at the east of the Coast Range, buried under the Cenozoic sedimentary cover, the Llanquihue Basement Complex (Duhart et al., 2001) was recognized in the exploration boreholes (Elgueta et al., 2000). It is mostly composed of micaschists with Devonian detrital zircon U-Pb maximum depositional age (Quezada, 2015; Hervé et al., 2016; 2018). Duhart et al. (2001) report the existence of an undeformed hornblende tonalite with ca. 359 Ma (Ar/Ar in hornblende) crosscutting micaschists at the bottom of the Rahue-2 borehole (Figure 1B). Geophysical surveys show the existence of an NNE-SSW-trending dense block at depths (Tašárová, 2007; Yañez and Rivera, 2019; Maksymowicz et al., 2022), which is interpreted as the subsurface extension of the Llanquihue Basement Complex (Figure 1B). The northern limit of this basement complex is located near Lago Ranco where the Trafún Metamorphic Complex crops out (Figure 1B), and the latter shares a significative Devonian component in the detrital zircon age population (Hervé et al., 2016). The Trafún Metamorphic Complex is intruded by the late Carboniferous coastal batolith, where Deckart et al. (2014) reported a U-Pb SHRIMP age of 306 ± 1 Ma.

The metamorphic fabric of the Río Llico Micaschists is constituted by two foliations, Sw1 and Sw2. In the western sector of this unit (Figure 2), the foliation Sw1 dips 70°NW and is crenulated by foliation Sw2 that dips 40°W (Figures 3A,B). The L2 lineation formed by the intersection of the Sw1 and Sw2 foliations dips 42° to the SW. Parasitic Z type microfolds related to the Sw2 axial plane crenulation (Figure 3C) are facing to the SW. The sigmoidal shape of the microfolds indicates top-up-to-the-northeast (i.e., reverse-sense) shear (Figure 3B), as shown by Bell and Johnson (1992) for determination of the shear sense of crenulation development in a fold limb. The mineral stretching lineation associated with the Sw2 dips 20° to the SW. In addition, the Sw1 foliation at microscale is anastomosed and oblique to the Sw2 foliation (Figure 3D).

In the eastern sector of this unit (Figure 2), the Sw1 foliation dips 40° to SE, while the Sw2 foliation dips 20°E. The structure associated with the Sw2 foliation of this unit is part of a kilometric scale northeast-vergent duplex thrust sheets. The contact between the Río Llico Micaschists and Pinuno Metasandstones is inferred to be related to a later west-vergent thrust (Figure 2C).

Three foliations (Se1, Se2, and Se3) are recognized in the Cajonmo Garnet Micaschist unit, while only the Se2 structures are present in the Zarao Metavolcanic Rock unit. The Se1 structures in the Cajonmo Garnet Micaschist are rootless isoclinal folds of centimetric and millimetric wavelength (Figures 4A,B, 8B). The Se2 foliation is formed by an asymmetric foliation that dips 60° to the W (Figures 4A,B) and folds the previous Se1 foliation in parasitic isoclinal S-type folds that face to the W. At microscale, the foliation Se2 is anastomosed and envelopes a few syn tectonic garnet porphyroblasts with asymmetric sigma tails (Figure 4D); however, most garnet crystals have post-tectonic textural features (Figure 4C). In addition, the sigmoidal shape of the Se2 micas and quartz bands indicate top-down-to-the-west shear sense representing a reversed normal limb of a west-vergent Fe2 isoclinal fold (Figures 2C–4B–4B). The Se2 foliation is folded by cylindrical Fe3 tight folds. A displacement between the pelitic and psammopelitic layers was recognized in their hinge, produced by the flexural slip. These cylindrical folds develop a 40°E dipping Se3 axial plane crenulation cleavage representing an inverse limb of a west-vergent fold (Figures 2C–4B–4B). The contact between the Cajonmo Garnet Micaschists and Zarao Metavolcanic Rocks lies in the west flank of the Fe3 synform. While this contact may be tectonic or interlayed, further fieldwork is required to define it. The fold axes Fe3 are conformed by the intersection of the Se2 and Se3 foliation planes and dip 15° to the south (Figure 4A). A detailed analysis of the (micro)structure and metamorphic mineral growth is given in Section 4.4. The deformation of the Se3 foliation is coherent with the ductile thrusting responsible of sitting the Cajonmo Garnet Micaschist together with the Zarao Metavolcanic Rocks above the poorly deformed Pinuno Metasandstones (Figure 2C).

The structures of the Pinuno Metasandstones are well exposed at the Las Cuyas locality (Figure 2A). They are poorly deformed and exhibit a 30° SE-dipping Se2 cleavage. The preserved sedimentary structures indicate a normal position of the beds, and bedding (S0) dips 50° to the SE (Figures 3G,H). The angular relationship between the Se2 cleavage and the bedding S0 is observed at microscale as well (Figures 3I,J). Hence, the exposed outcrops of the Pinuno Metasandstones represent the normal limb of the NW-vergent kilometric structure (Figure 2C).

The SiO₂ content of the metarhyolite is ∼66 wt% and in the metabasaltic rocks ranges between ∼49 and 54 wt%; among the latter, the higher contents (∼52–54 wt%) correspond to samples FO1696 and CM20608. The metabasalts have FeO_t (FeOt = 0.9 * Fe₂O_3T) values between 6.9 and 10.08 wt%, MgO between 3.85 and 11.41 wt%, CaO between 9.17 and 11.99 wt%, and K₂O+ Na₂O values between 2.91 and 5.11 wt%. In terms of major element composition, sample FO1695 is the less fractionated and samples FO1696–CM20608 are the most evolved.

The metarhyolite is enriched in Zr (1,322 ppm) and has a Nb/Y ratio of 0.51, and its protolith is classified as an alkaline rhyolite (Figure 5A; Winchester and Floyd, 1977; modified by Pearce, 1996). The chondrite-normalized REE pattern (Figure 5C) shows a steep pattern for the LREE (La_N/Sm_N= 5.29), a nearly flat pattern for the HREE (Sm_N/Yb_N= 1.19), and a pronounced Eu negative anomaly (Eu/Eu*= 0.26). The N-MORB-normalized multi-element pattern (Figure 5D) is characterized by a negative Nb-Ta anomaly, a small Zr positive anomaly, and a marked Ti negative anomaly.

Based on the Ti/Zr and Nb/Y ratios, the protoliths of the metamafic rocks correspond to subalkaline basalts (Figure 5A; Winchester and Floyd, 1977; modified by Pearce, 1996). Samples FO1696 and CM20808 have Nb/Y values of 0.6–0.65 and yield compositions close to the alkali basalt field. The Zr/Y and Zr/Nb ratios (Figure 5B) point to three types of basaltic protoliths like, normal mid-ocean ridge basalts (N-MORBs), enriched mid-ocean ridge basalts (E-MORBs), and plume-type MORB (P-MORB), the latter transitional between E-MORB and alkaline oceanic within plate basalts (OIBs). The chondrite-normalized REE pattern shows significant differences among the metabasaltic rocks. The metabasalts with E-MORB and P-MORB like signatures are enriched in LREE (La_N/Sm_N= 1.74 and ∼2.5, respectively), while the N-MORB like is depleted in LREE (La_N/Sm_N= 0.54). On the other hand, the HREE pattern of N-MORB and E-MORB like is nearly flat (Sm_N/Yb_N= 0.89 and 1.09, respectively) when compared to the more step pattern in the P-MORB like metabasalts (Sm_N/Yb_N ∼ 1.95). A marked Eu positive anomaly is displayed by the N-MORB like (Eu/Eu*=1.36) and a subtle one by the E-MORB (Eu/Eu*=1.07). Furthermore, the P-MORB-like metabasalt has a small negative Eu anomaly (Eu/Eu* ∼0.91). The N-MORB like shows a Nb negative anomaly, while the others display coupled Nb-Ta negative anomalies. The P-MORB like basalt has a small negative anomaly of Ti, and the E-MORB and N-MORB like basalts have small negative anomalies of Hf-Zr (Figure 5D).

The ⁸⁷Sr/⁸⁶Sr_t isotope composition shows values between 0.7055 and 0.7067 (Figure 7A). These high values suggest that post-magmatic processes (hydrothermal/metamorphic events) have modified the original signature of the igneous protolith, particularly in the metabasaltic rocks, perhaps involving the participation of sea water. The εNd_t values have a wider range; the lowermost values correspond to the metarhyolite and the P-MORB (-3.04 to -1.64), while the metabasalts with signatures like E-MORB and N-MORB have positive values of 3.28 and 7.6, respectively (Figure 7A). The ²⁰⁷Pb/²⁰⁴Pb_t isotope composition has a narrow range of values between 15.57 and 15.64, as well as the ²⁰⁶Pb/²⁰⁴Pb_t composition with values between 17.79 and 18.11 (Figure 7B). A wider range of values is shown by the ²⁰⁸Pb/²⁰⁴Pb_t; the metarhyolite and P-MORB have values between 38.69 and 38.97, while the E-MORB and N-MORB have values of 38.01 and 37.15, respectively (Figure 7C).

The sample FO1710 contains the mineral assemblage biotite + white mica + garnet + quartz + plagioclase + ilmenite ± epidote ± apatite ± chlorite (mostly retrograde). Garnet is up to 1 mm in diameter. The crystallization–deformation relationships and mineral paragenesis, shown in Figure 8, suggest the existence of two metamorphic episodes affecting the Cajonmo Garnet Micaschists. The essential metamorphic minerals formed in the De1 deformation event are white mica + ilmenite, constituting the folded foliation Se1. The metamorphic minerals formed in the Se2 comprise biotite + white mica + ilmenite aligned in the crenulation cleavage Se2 domain spaced by quartz and plagioclase microlithons (Figure 8B). According to the petrographic observations in Section 4.1.2.2, most garnets, chlorites, and epidotes are post-tectonic to the Se2 foliation. The Syn-Se2 garnet porphyroblasts exhibit an internal foliation (S1) formed by the inclusion of epidote, quartz, apatite, biotite, white mica, and ilmenite.

The chemical composition of garnet was analyzed along two profiles (Figures 9A–C). The Syn-Se2 garnet #1 is relatively large (500 μm). It is idiomorphic and shows straight borders, and the Se2 foliation is defined by biotite, white mica, and ilmenite wrapping around. Garnet #2 is smaller and post-tectonic. In the central part, garnet #2 contains a linear trail of inclusions, while close to the edge’s inclusions are randomly oriented. The chemical analyses indicate that garnet is essentially almandine. Garnet #1, unlike garnet #2, exhibits a noticeable chemical zonation. Mn decreases from core to rim (X_Sps = 0.14–0.20), whereas Fe shows the opposite behavior (X_Alm = 0.67–0.73). Ca and Mg remain almost constant from the core to rim (X_Grs = 0.05–0.07; X_Prp = 0.06–0.08). Garnet #2 shows no chemical zonation; noteworthy, its composition essentially conforms to the rim of garnet #1. It should be noted that at the very edges of garnet #2, there is a decrease in X_Alm from 0.733 to 0.724 and X_Prp from 0.065 to 0.061, and an increase in X_Sps from 0.146 to 0.150 and X_Grs from 0.053 to 0.063 (Figure 9E).

The white mica and biotite analyzed are mainly in the foliation Se2 and occur as fine-grained crystals with lepidoblastic texture in mica-rich bands. On the basis of 11 oxygens for the structural formula, the white mica composition is essentially muscovite that is characterized by a considerable composition range: Si_iv = 3.06–3.18 per formula unit (pfu), Al_iv = 0.82–0.94, and Al_vi = 1.76–1.88 pfu. The compositional variation is mainly due to the Tschermak exchange (Si_iv + (Mg, Fe²⁺)_vi = Al_iv + Al_vi) (Figure 9F). A slight deviation can be explained by the presence of minor amounts of ferric iron. Biotite occurs as fine-grained crystals in mica-rich bands. The biotite composition shows only little variation: Si_iv = 2.63–2.69 per formula unit (pfu) unit based on 11 oxygens, Al_iv = 1,30–1.36, Al_vi = 0.30–0.37 pfu, Ti_vi = 0.08–0.10 pfu, and X_Fe = Fe/(Fe+Mg) = 0.63–0.65. Plagioclase is scarce and occurs preferentially as microlithons in the granoblastic texture displaying an oligoclase to albite composition of X_An = 0.06–0.24. Ilmenite contains about 5% of the pyrophanite component. Retrograde chlorite replacing biotite shows composition with X_Fe = 0.57–0.60 (Figure 9G). The epidote is scarce and scatters on the granoblastic and lepidoblastic bands with a pistacite composition of X_Ps = Fe/(Fe+Al) = 0.24.

For P–T evaluation, P–T and T–O₂ pseudosections were calculated using the thermodynamic Gibbs energy minimizing software Perple_X (downloaded 23.02.2022; Connolly, 1990; 2005) in the chemical system Mn-Na-Ti-Ca-K-Fe-Mg-Al-SiO-H-O, for a range from 2 to 12 kbar and 300 to 650°C. The thermodynamic database used was TC-DS633 by Holland & Powell (1998). The following solution models as provided by the Perple_X were applied: Gt(W) for garnet, Bi(W) for biotite, Chl(W) for chlorite, Mica(W) for white mica, Ilm(DS6) for ilmenite, feldspar, and Ep(HP11) for epidote.

For the calculation, the analyzed composition of the studied rock FO1710 from the Cajonmo Garnet Micaschists has been slightly corrected to fit the 11-component system. Furthermore, the CaO content was reduced according to the bulk-rock phosphorus content, assuming that these elements are bound to apatite (Calderón et al., 2017). Accordingly, the CaO content is reduced from 0.93 to 0.77 wt% (Table 2). To estimate the portions of FeO and Fe₂O₃, a T-O₂ pseudosection at 5 kbar was calculated (not shown). Owing to the presence of ilmenite and the absence of both hematite and magnetite, the portion of ferric iron is limited to values clearly lower than 10% of the total iron, in the temperature range under present consideration. A further constraint is given by the composition of epidote. Accordingly, for the calculation of P–T pseudosections, the amount of Fe³⁺ is set to 3% of total Fe (0.16 wt% Fe₂O₃). The water content is set sufficiently high to ensure the water-excess condition in the considered P–T span.

Figure 10 shows major aspects of the calculated P–T pseudosection for sample FO1710. The P–T conditions of the formation of garnet can be derived from the intersection of the X_Sps and X_Prp isopleth of this mineral (Figures 11A,B–H). The X_Prp isopleth indicates a temperature of 540°C, independent of pressure. A pressure of about 7 kbar is then derived from the intersection with the X_Sps isopleth. This P–T condition is valid for the core and rim of the large garnet as well as for the small garnet, which all have the same Ca and Mg contents. The increase of Mn from the core to rim in garnet #1 may therefore reflect a Rayleigh fractionation of Mn during growth of garnet at constant P–T of 7 kbar/540°C (Figure 11A). Iron then behaves complementary to Mn. Noteworthy, the derived P–T of garnet growth is close to the low-T stability limit of the garnet. The onset of garnet growth there corroborates the interpretation of fractionation, causing the zonation of this mineral, supported by the X_Sps = 0.20 garnet core isopleth that does not intersect with Mg and Ca garnet core isopleths, showing non-equilibrium conditions. In addition, the derived P–T is compatible with the presence of biotite, plagioclase with oligoclase composition (Figure 11D), chlorite (Figure 11F), epidote (Figure 11G), and ilmenite. Because the presence of platy ilmenite aligned in the Se2 foliation, the P–T of 7 kbar/540°C may be indicative for the Se2–Fe2 ductile structure as a whole.

The calculated white mica composition at 7 kbar/540°C is characterized by a Si content of 3.06 Si pfu (Figure 11E). This corresponds to a low Si cluster of the analyzed mica grains being in equilibrium with the garnet. Higher Si contents up to 3.18 Si pfu in white mica must therefore have been crystallized at either high pressure or lower temperature or a combination of both. Previous higher P conditions are not very probable because no indication for Ca rich garnet cores has been found (Figure 11C). A more reliable interpretation is that high-Si white mica represents an earlier, subgreenschist facies generation that has not been erased during prograde evolution (Figure 11E).

Such conditions in the range of 5–7 kbar/300°C would imply that titanite was stable instead of ilmenite. Obviously, titanite was completely converted to ilmenite (+garnet) during temperature increase (Figure 10). The late-stage metamorphic evolution is characterized by the formation of thin Mn richer rims in the garnet border pointing to corrosion still in the stability field of the garnet. Later, the replacement of biotite and garnet by chlorite occurred. According to the derived data, the most probable interpretation is to assume a P–T path characterized by roughly isobaric heating starting at subgreenschist facies conditions up to 7 kbar/540°C (Figure 11H).

Based on the petrological and geochemical data presented in Section 4.1 and 4.3, the protoliths of the metabasaltic rocks are subalkaline basalts with N-MORB, E-MORB, and P-MORB geochemical affinities (Figures 5B,C). A supra-subduction zone setting of formation of the basaltic protolith is supported by the Nb-Ta negative anomalies of the P-MORB and E-MORB and by the Nb negative anomaly of the N-MORB type metabasalts (Figure 5D). The input of the subdction is also observable in the Th/Yb vs. Nb/Yb diagram, where the P-MORB plot above the MORB-OIB array (Figure 6A, Pearce, 2008). Similar geochemical abundances (Figures 5, 6) are shown by the metabasalt from the Huilma-2 borehole (sample HU2-T5) which is located in the eastern slope of the Coastal Cordillera, next to the garnet-bearing schist of Huilma-1 borehole (>460–400 Ma U-Pb zircon maximum depositional age; Hervé et al., 2018). These features suggest a wider geographical extension of the Zarao Metavolcanic Rocks than what was previously considered by Palape (2020).

The depth of melting for basaltic magma formation can be addressed using the TiO₂/Yb vs. Nb/Yb diagram proposed by Pearce (2008), and it indicates shallow melting for the Zarao metabasalts (Figure 6B). High Sm/Yb_N ratios are related to the influence of garnet on the petrogenesis of basalts (e.g., Kay and Kay, 1991; Kay and Mpodozis, 2001). Willbold and Stracke (2006) compiled geochemical data from ocean island representing the major OIB families and obtained a Sm/Yb_N mean value of 4.79 ± 0.56 among them. The ∼1.95 Sm/Yb_N ratio of the P-MORB metabasalts is lower than the average OIB, in agreement with shallower mantle melting in the spinel stability field. The E-MORB and N-MORB have lower Sm/Yb_N ratios of 1.09 and 0.89, respectively. The Ce/Y parameter has been used as a proxy to estimate the Moho depth in arc lavas, and thus the crustal thickness, with lower values related to thinner crust (Mantle and Collins, 2008). The Ce/Y ratios of the Zarao metabasalts vary from 1.42 to 1.48 in the P-MORB, to 0.63 in the E-MORB and toward 0.23 in the N-MORB (Table 1), which supports the formation of these rocks in an environment of thin crust, but of variable thickness (between ∼25 and 10 km thick following the work of Mantle and Collins, 2008).

The Sr-Nd-Pb isotope composition of the Zarao Metavolcanic Rocks shows a narrow composition in the ²⁰⁷Pb/²⁰⁴Pb_t vs. ²⁰⁶Pb/²⁰⁴Pb_t space (Figure 7B) and is similar to the Bulk Silicate Earth (BSE) reservoir of Zindler and Hart (1986), not shown. However, a significative variation in the εNd_t vs. ⁸⁷Sr/⁸⁶Sr_t and ²⁰⁷Pb/²⁰⁴Pb_t vs. ²⁰⁸Pb/²⁰⁴Pb_t space is observed among the three types of basaltic protoliths (Figures 7A–C), reflecting a progressive change in the mantle source from an Enriched Mantle 1 (EM1) type toward a Depleted MORB Mantle (DMM) type.

The heterogeneity observed in the geochemistry of the Zarao Metavolcanic Rocks probably reflects different mantle sources and slab fluids involved in the magma generation, which is not uncommon on backarc volcanism environments (e.g., Ducea et al., 2015). Furthermore, we consider that the systematic variations of mantle sources, subduction input, and depth of melting observed in the metabasaltic protoliths may have a diachronic origin. The P-MORBs are derived from an EM-1 like mantle source, have a more marked slab input, and were formed in a relatively thicker crust. The N-MORBs were formed by mantle melting of the DMM-like source, with minor slab influence, and formed in a thin crust setting. Finally, the E-MORB metabasalt is a transitional composition between P-MORB and N-MORB.

Despite the high LOI values (5.89 wt%) obtained in the Zarao metarhyolite sample, the geochemical abundance of the less mobile elements is similar to those reported by Rapela et al. (2021), and therefore, the obtained results are useful for the geochemical evaluation (Figure 5). The Zarao metarhyolite has an alkaline geochemical signature as deduced from the MALI value of 10.22 (MALI calculated in wt% as Na₂O+ K₂O–CaO; Frost et al., 2001), in agreement with the Nb/Y ratio of 0.51 (Winchester and Floyd, 1977; modified by Pearce, 1996). These geochemical ratios like those of quartz-saturated magmas agree with the preservation of quartz microphenocrysts of the igneous protolith. The ferroan character of the metarhyolite (Fe*= 0.96, Fe* calculated as Fe*= FeO_t/(FeO_t/MgO)) is typical of A-type magmas (Frost et al., 2001). Quezada et al. (2018) distinguished the OIB-like (A₁-type) and arc-like (A₂-type) mixed source of the Zarao metarhyolite based on the Y-Nb-Ga geochemical ratios proposed by Eby (1992) to subdivide A-type granitoids.

The marked negative Eu anomaly (Figure 5C) indicates that the fractionation of feldspar played a major role in its petrogenesis. On the other hand, the flat HREE pattern (Sm_N/Yb_N= 1.19) is indicating that the fractionation of garnet and amphibole during its petrogenesis was not significative (Davidson et al., 2012). Negative anomalies of Nb-Ta and Ti in the NMORB-normalized spider diagram (Figure 5D) together with an impoverishment in the content of high field strength elements (HFSEs) are characteristics like those observed by magmas formed in subduction-related environments or derived from the subcontinental lithospheric mantle (more details below).

The Sr-Nd-Pb isotope composition of the Zarao metarhyolite is similar to the P-MORB metabasalts (Figure 7). This unit shows the Sr-isotope composition (⁸⁷Sr/⁸⁶Sr_t = 0.7061, Figure 7A) lower that the one obtained for the P-MORB, and probably closer to the original composition of the igneous protolith of the metarhyolite. The ²⁰⁷Pb/²⁰⁴Pb_t and ²⁰⁸Pb/²⁰⁴Pb_t composition is like the EM1-type mantle source (Figure 7C). Thus, the isotope data strongly suggest a cogenetic link between the P-MORB metabasalts and the metarhyolites, implying that the enriched isotope composition of the metarhyolite is related to the magma source. Rapela et al. (2021) reported the zircon εHf_t and δ¹⁸O isotope compositions of the Zarao metarhyolite showing negative εHf_t values (−5.5 to −1.2), which agree with the result here presented in whole rock εNd_t data, and δ¹⁸O values between and below mantle array (3.9–5.1), the latter supporting nearly the absence of contamination/assimilation of the upper continental crust in its petrogenesis.

While tholeiitic and calc-alkaline magmas can coexist on the same supra-subduction zone settings, a ∼2 wt% H₂O content has been proposed as the limit between both magmatic trends (Zimmer et al., 2010). While lower water contents in magmas (<2 wt%) are expected in the tholeiitic series and are characterized by the suppression of early crystallization of magnetite, water-bearing silicates, such as amphibole and biotite, and plagioclase, the opposite occurs within the calc-alkaline series (e.g., Gill, 1981; Grove and Baker, 1984; Grove et al., 2003). The extreme fractionation of a tholeiitic magma following the crystallization of olivine + plagioclase + pyroxene + Fe-Ti oxides could develop the formation of alkaline rhyolites (e.g., Spulber and Rutherford, 1983; Geist et al., 1995; Whitaker et al., 2008), and it seems to be a dominant process in the generation of bimodal magmatism in continental rift-related settings and zones of mantle upwelling (Bachmann and Bergazantz, 2008; Ayelew and Ishiwatari, 2010). The large Eu and Ti negative anomalies displayed by the Zarao metarhyolite agree with the major fractionation of feldspar and Fe-Ti oxides, respectively. According to Spulber and Rutherford (1983), during the fractionation of a tholeiitic magma, plagioclase fractionation is far more relevant under low-pressure conditions, while amphibole fractionation plays a major role at pressures >2 kbar, despite the oxygen fugacity of the magma. Then, it is possible that the magmatic differentiation toward rhyolitic compositions occurred in a magmatic chamber located at shallow crustal depths. Furthermore, the tholeiitic differentiation trend promotes the early crystallization of ¹⁸O-enriched mineral phases such as feldspar, instead of ¹⁸O-depleted minerals like amphibole or magnetite; thus, extreme fractional crystallization could lead to the formation of an ¹⁸O-depleted magma (Taylor, 1968; Gao et al., 2018). While this is a plausible mechanism to explain the low δ¹⁸O values of the Zarao metarhyolite, the assimilation of the high-temperature altered mafic crust cannot be discarded (Troch et al., 2020).

The age of the basaltic protolith of the Zarao Metavolcanic Rocks is shaded by late Carboniferous to Permian tectono-thermal events that affected the metamorphic units in the Coast Range. Duhart et al. (2001) reported K-Ar ages of ca. 297 Ma and 269 Ma for this unit (Figure 2B), and thus, some inferences need to be done to unravel the relative ages among the Zarao metabasalts. As stated in the previous section, the origin of the Zarao metarhyolite is associated with the extreme fractional crystallization of a mafic magma such as the P-MORB metabasalts. Considering that the Zarao metarhyolite was formed from the differentiation of basaltic magmas, it requires that the basalt sources were not solidified yet at the time of magma differentiation, and therefore, it is interpreted that the P-MORB metabasalts and the Zarao metarhyolite are roughly contemporaneous. In this scenario, the geochemical variations observed in the Zarao metabasalt are diagnostic of progressive lithospheric thinning over a supra-subduction zone setting. This tectonic context agrees with the Middle Devonian tectonic evolution proposed by Rapela et al. (2021) for the segment.

Basaltic magma geochemistry can be used as a tectonic setting tracer due to its low fractionation and low degree of crustal contamination (e.g., Pearce, 2014; Xia and Li, 2019). In the Y-La-Nb tectonic discrimination diagram (Cabanis and Lecolle, 1989), the Zarao metabasalts plot mostly in the continental tholeiites and oceanic backarc fields (Figure 6C). These compositional fields are intermediate between the orogenic, arc-related basalt domains (with a low La/Nb ratio), and the anorogenic N-MORB/E-MORB/OIB domains (with a high La/Nb ratio). The progression from a continental extensional setting (P-MORB type) toward an oceanic backarc setting (N-MORB type) is in concordance with the crustal thickness changes deduced from the proxies previously used (e.g., Ce/Y, Sm/Yb_N, Section 5.1.1). Thus, it is proposed that the Zarao metabasalts represent the northward prolongation of the Chaitenia marginal basin, the latter previously described in the Reñihue Fjord area (Figure 1A), where a similar geochemical diversity of metabasaltic-types (Figures 5, 6) having a Middle Devonian age crops out (Hervé et al., 2018; Rapela et al., 2021).

Rapela et al. (2021) noted the similar zircon εHf_t values between the calc-alkaline plutons of the San Martin arc (ages between ca. 405 and 395 Ma) and the Zarao metarhyolite (ca. 393 Ma) and suggested that both were formed over the same lithospheric block. The EM-1 like mantle source of the P-MORB metabasalts, combined with their depletion of high field strength elements (HFSEs) such as Nb, Ta, and Ti, suggests that they may have formed by low-pressure partial melting of the metasomatized subcontinental lithospheric mantle (SCLM) (Salters and Shimizu, 1988; O’Reilly and Griffin, 2012). The partial melting was favored by the rollback of the down going slab and adiabatic decompression due to primitive (DMM-type) mantle upwelling, which lead to enhanced extension in the upper plate and the detachment of the forearc away from the waning San Martin arc (Rapela et al., 2021).

As the locus of arc magmatism switched toward the detached forearc, namely, the ‘Chaitenia island arc’, a spreading center with a typical DMM-like mantle source was progressively stablished in the backarc basin, where the E-MORB and N-MORB Zarao metabasalts were formed. Backarc basin basalts have a mixed tholeiitic geochemical signature between typical N-MORB and arc magmas (Stern, 2010); however, the subduction input in backarc basin basalts will be more subtle as we move away from the volcanic arc front, i.e., toward the more mature portions of the backarc system (e.g., Xia and Li, 2019). This is the case proposed for the Zarao metabasalts, with the E-MORB being formed in early stages of backarc spreading (Nb-Ta anomaly plus a mantle source transitional between EM-1 and DMM), and the N-MORB metabasalts formed during more mature stages (an Nb anomaly and a DMM source). However, we acknowledge that these tectonic processes are complex and other interpretations may be suitable as well.

Coeval with the initial stages of forearc extension and the formation of the Zarao metarhyolite, the Caleta Gonzalo diorite was formed near Reñihue (392 ± 3 Ma, Figure 1A). It corresponds to the older calc-alkaline intrusive of the Chaitenia oceanic arc (Rapela et al., 2021). The REE abundances of this diorite are characterized by a “spoon-shaped” REE pattern indicative of amphibole fractionation (Figure 5) (Davidson et al., 2012). The zircon O-Hf isotope composition is marked by low δ¹⁸O values (+3.8 to +4.5) like those shown by plagiogranites intruding supra-subduction zone ophiolites (Grimes et al., 2013), and highly primitive εHf_t values (14.8–16.7) indicating derivation from a depleted mantle source (Rapela et al., 2021), as it was observed in the N-MORB Zarao metabasalts. The north-south mantle source variations between the Zarao area (enriched mantle) and the Reñihue area (depleted mantle) may be indicating diachronism in the opening of the backarc basin; while the Zarao area was at an embryonic stage of continental thinning, the Reñihue area seafloor spreading in the backarc was already active at ca. 393 Ma. Further north, at the La Cabaña area (Figure 1A), ultramafic rocks have been interpreted to be formed in an oceanic backarc setting (González-Jiménez et al., 2014). This unit is related to schists having youngest detrital zircon age peaks between ca. 352 and 341 Ma (Romero et al., 2018). Speculatively, the spatial and temporal relationships observed along the strike of the Chaitenia arc/backarc system might be indicative of backarc basin opening from south to north.