Introduction

The southern portion of the Andean Southern Volcanic Zone (SVZ, Fig. 1; Stern 2004; Stern et al. 2007) is composed of several large composite volcanoes, including Melimoyu, Mentolat, Maca, Cay, and Hudson, and also many (>15; Fig. 1) small monogenetic mafic cones along the Liquiñe-Ofqui Fault System (LOFS; Gutiérrez et al. 2005). Lacustrine sediment cores, from seven lakes formed in this region of Chile (Figs. 1 and 2) after the initiation of the retreat of mountain glaciers, which started by at least ~17,900 calendar years before present (cal years BP) (Miranda et al. 2013), each contain multiple tephra layers derived from an integrated total of >70 explosive eruptions of these volcanoes (Figs. 3 and 4). These eruptions span the period of climate change at the transition from the glacial to the post-glacial warm period and extend through the Holocene.

Fig. 1
figure 1

Map of southernmost portion of the SVZ showing the location of Hudson, Maca, Cay, and Mentolat volcanoes. Also shown are the locations of some of the >15 monogenetic centers along the Liquiñe-Ofqui Fault System (LOFS) and surrounding Hudson (Gutiérrez et al. 2005). The boxed area encloses the lakes from which cores were obtained (Fig. 2; Miranda et al. 2013). Lago Shaman (de Porras et al. 2012) lies further to the northeast

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Fig. 2
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Google Earth image showing the location of the seven lakes from which the cores (Figs. 3 and 4) discussed in the text were taken and the thickness range for the Ho deposits in each lake (Table 2)

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Fig. 3
figure 3

X-ray image of the 8.5 × 1 m sections of core extracted from Lago Unco, >100 km east of Hudson volcano (Figs. 1 and 2). The 74 different tephras in this core appear as white layers due to their higher density compared to the predominantly organic lake sediments in which they are preserved. The large late-glacial Ho tephra is 60-cm thick, extending from core section 8 at 78 cm to section 9 at 38 cm, and contains pumice up to >2 cm in diameter. The inferred age of Ho is constrained in this core by dates in the organic rich sediments overlying and underlying the Ho tephra as between 17,300 and 17,440 cal years BP in age (Table 1; Miranda et al. 2013). These ages are consistent with it underlying a late-glacial tephra sequence S9 dated as <14,910 cal years BP (Miranda et al. 2013)

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Fig. 4
figure 4

X-ray images for six of the other lakes where the Ho tephra has been observed. The large late-glacial Hudson eruption is labeled Ho, with the sequence of nine (S9) closely spaced eruptions and MENo as reference deposits for each core. Each section of each core is approximately 1-m long and 4-cm wide

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Specific source volcanoes for some of the tephra layers in these cores have been identified from their trace-element chemistry. Based on their chemical characteristics (Fig. 5) and their stratigraphic depth in the cores, some of the previously identified and dated large explosive eruptions of Hudson, such as H2 at ~3,920 cal years BP (Naranjo and Stern 1998), and of Mentolat, such as MEN1 at 7,560 cal years BP (Naranjo and Stern 2004; Stern et al. 2013), are represented by tephra in these cores (Fig. 3).

Fig. 5
figure 5

La versus Ba plot showing the fields occupied by the different eruptive products from Hudson (Ho from Tables 3 and 4; H3 in 1991 AD from Kratzmann et al. 2009 and 2010; H1 and H2 from Naranjo and Stern 1998) and Mentolat Volcano (Naranjo and Stern 2004; Stern et al. 2013). Hudson eruptive products are elevated in LIL, HFSE, and REE compared to Mentolat and other southern SVZ tephras

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Hudson (45°54′S; 72°58′W), which is the southernmost volcanic center in the SVZ, has had three large explosive eruptions during the Holocene (Stern 1991, 2008; Scasso et al. 1994; Naranjo and Stern 1998; Kratzmann et al. 2009, 2010). Fifteen sediment cores from the seven lakes each contain an unusually thick late-glacial age tephra layer (Tables 1 and 2; Figs. 3 and 4) resulting from a very large explosive eruption, which based on both its bulk trace-element composition (Fig. 5) and distribution (Fig. 6), was derived from Hudson volcano and is termed Ho (Weller et al. 2013). Here, we characterize some of the physical and chemical features of the late-glacial Ho eruptive products and make an assessment of the total eruptive volume of this and other explosive eruptions of Hudson since the beginning of deglaciation in southern Chile.

Table 1 Age constrains for the Ho tephra deposits from Miranda et al. (2013)
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Table 2 Thickness (in cm) of the Ho and H2 tephra deposits in lake cores from south Chile
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Fig. 6
figure 6

Estimated 50- and 10-cm isopachs of the Ho eruption (solid lines) compared to those for the mid-Holocene H1 eruption (dashed lines; Naranjo and Stern 1998). The distribution of the Ho tephra indicates northeastward dispersion

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Geologic/tectonic setting

Hudson volcano is the southernmost volcano in the Andean Southern Volcanic Zone (SVZ; 33–46°S; Stern 2004; Stern et al. 2007). Volcanism in this section of the Andean cordillera is the result of subduction of the Nazca Plate under the South American continent (Fig. 1). Further to the south is the Austral Volcanic Zone (AVZ; Stern and Kilian 1996), in which volcanism is the result of the subduction of the Antarctic Plate underneath the southern portion of the South American Plate. In between the AVZ and Hudson is a ~350-km gap in the active volcanic arc that lies just south of where the Chile Rise, an active spreading center that separates the Antarctic and Nazca Plates, has been subducted under the South American Plate (Cande and Leslie 1986). Marine geophysical data indicate that the Chile Rise first collided with the Chile Trench near Tierra del Fuego and has migrated north along the continental margin since 15–20 Ma to its current location (Cande and Leslie 1986; Nelson et al. 1994).

Cembrano et al. (1996) suggest that there is a strong connection between the arc-parallel Liquiñe-Ofqui Fault System (LOFS; Fig. 1) and the location of the volcanic centers of the SVZ. The LOFS results from a combination of oblique subduction of the Nazca Plate and the impingement of the Chile Rise against the continent. Most of the southern SVZ volcanoes are located along or to the west of the LOFS. Hudson volcano, in contrast, is situated approximately 30 km east of this fault zone (Naranjo and Stern 1998). Numerous monogenetic mafic cinder cones are also roughly aligned along the LOFS trend (Fig. 1; Gutiérrez et al. 2005), and others surround Hudson as well.

The Quaternary volcanoes of the southern SVZ are located on a portion of the Northern Patagonian Batholith composed of mainly tonalites and granodiorites (Pankhurst et al. 1999; D’Orazio et al. 2003; Gutiérrez et al. 2005). This Cretaceous to late Miocene intrusive complex, which is ~1,000 km in length and extends across the length of the southern SVZ (41°S–52°S), formed by episodic magmatic events over a 125 Myr time interval (Pankhurst et al. 1999).

The magmas produced by the southern SVS volcanoes are generally high-Al2O3 basalts and basaltic andesites (Futa and Stern 1988; López-Escobar et al. 1993; D’Orazio et al. 2003), with only minor andesites, dacites, and rhyolites. Hudson generates magmas which are compositionally distinct from the other southern SVZ volcanic centers such as Mentolat (Fig. 5), Maca, and Cay. Analyzed samples from Hudson range from basalts to trachydacites with high TiO2, FeO, Na2O, and K2O relative to rocks of similar silica content erupted from other volcanic centers in the southern SVZ (Futa and Stern 1988; Stern 1991, 2008; López-Escobar et al. 1993; Naranjo and Stern 1998; D’Orazio et al. 2003; Gutiérrez et al. 2005; Kratzmann et al. 2009, 2010). Hudson’s extrusive rocks are also rich in incompatible trace elements, which include LILE, REE, and HFSE, compared to samples from other southern SVZ and AVZ volcanoes with similar silica content (Stern 1991, 2008; Naranjo and Stern 1998).

Hudson eruptions

The Hudson caldera was first identified as a volcano only after a small Plinian eruption and lahar in 1971 (Tobar 1972; Fuenzalida and Espinosa 1974; Fuenzalida 1976; Cevo 1978; Best 1989, 1992). Hudson erupted as recently as October 2011 and K-Ar dates indicate that this volcano has been active since 1 Ma or before (Orihashi et al. 2004). The most recent large explosive eruption from Hudson (H3 in 1991 AD) occurred on August 8, 1991. It began with an initial phreatomagmatic event (phase 1) which erupted through a 400-m crater and a 4-km-long fissure in the northwestern portion of the caldera (Naranjo 1991; Naranjo et al. 1993; Scasso et al. 1994; Kratzmann et al. 2009, 2010; Wilson et al. 2011, 2012). The paroxysmal phase (phase 2) began on the 12th of August with a Plinian style eruption from an 800-m-wide crater located in the southwestern portion of the caldera. This phase of the eruption continued on and off for 3 days before ceasing on the 15th of August. The tephra fallout from phase 2 was dispersed in an ESE direction covering a narrow, elongated sector of southern Patagonia (Scasso et al. 1994; Wilson et al. 2011, 2012). This eruption produced a volume of at least 4 km3 of material, but the overall volume is likely to have been larger because tephra was also deposited into the sea. Tephra from this eruption ranges in composition from basalt and basaltic-andesite during phase 1 to trachyandesite and trachydacite during phase 2, with a distinct gap between 54 and 60 wt.% SiO2 separating the two phases (Kratzmann et al. 2009, 2010). Furthermore, Kratzmann et al. (2009) observed a trend of progressive decrease in SiO2 content of the eruptive products during the second phase of the eruption. The abrupt transition from basalt/basaltic-andesite (51 to 54 wt.% SiO2) erupted in the first phase to trachyandesite and trachydacite (60 to 65 wt.% SiO2) during the second phase and the progressive decrease in SiO2 during the second phase have been attributed to a combination of magma mixing and fractional crystallization (Kratzmann et al. 2009, 2010).

Soil and sediment sections exposed east and the southeast of the Hudson volcanic center in river and road cuts, as well as in excavated trenches, preserve pyroclastic fallout deposits that were derived from nine explosive Holocene eruptions of Hudson, including two very large ones (Naranjo and Stern 1998). The younger of the two large events (H2) occurred at approximately 3,920 cal years BP, and the older (H1) eruption occurred approximately 7,750 cal years BP (Stern 2008; Prieto et al. 2013). Based on the isopach maps for tephra from these eruptions, Naranjo and Stern (1998) determined that both the H1 and H2 eruptions generated greater volumes of ejected material than did the H3 eruption in 1991 or the 1932 Quizapu eruption (9.5 km3; Hildreth and Drake 1992). Using radiocarbon age constraints and petrochemical data, Stern (1991, 2008) suggested that grey-green tephra layers in southernmost Patagonia, originally termed Tephra II by Auer (1974), are distal deposits of the H1 eruption derived from Hudson. These form >10-cm-thick tephra deposits in Tierra del Fuego, and this eruption may have had a significant impact on the prehistoric people living on this island (Prieto et al. 2013). Based on the proximal and distal maximum isopachs for the H1 tephra, Stern (1991, 2008) estimated that this eruption ejected >18 km3 of material. The chronology established by Naranjo and Stern (1998) shows that the recurrence interval between explosive eruptions of Hudson is irregular, but that the three largest Holocene explosive events (H1, H2, and H3 in 1991 AD) are relatively regularly spaced at 3,870 ± 40 years.

Although Naranjo and Stern (1998) did not observe any early Holocene or late-glacial tephra, Haberle and Lumley (1998) documented two late-glacial tephra deposits in lake sediments on Peninsula Taitao, 150 km southwest of the volcano. Radiocarbon dating of these deposits yielded ages of 14,560 cal years BP for eruption HW1 and 13,890 cal years BP for eruption HW2 (Haberle and Lumley 1998).

Methods

Multiple sediment cores were obtained, using a 5-cm-diameter modified Livingstone piston corer (Wright 1967), from seven small lakes within spatially limited watershed basins selected to minimize the amount of inorganic sediment deposited in the lakes. X-ray images of the cores (Figs. 3 and 4) were taken to allow for better visual identification of the tephra deposits and to provide a means of stratigraphic correlation of the tephra layers between the cores. The white layers in these images are the denser lithologies, often tephra deposits, and the darker layers are less dense organic-rich lacustrine sediments. The stratigraphic position of the Ho tephra layer, combined with several other key marker tephra deposits which were identified in the multiple cores, allowed for correlation of this deposit within the cores. The chronology of the tephra in the cores is controlled by AMS radiocarbon dates of organic material in the sediments (Table 1; Miranda et al. 2013) and chronostratigraphic correlation of tephra layers. Radiocarbon dates were converted to calendar years before present (cal years BP) using the CALIB 5.01 program (Stuiver et al. 1998).

Once identified, the individual tephra deposits were extracted from the cores by removing approximately half of the deposit with a knife. Several of the Ho tephra layers were sufficiently large to sample at discrete 10-cm intervals to assess if any vertical stratification is present within the deposits. The tephra samples were washed to remove any organic matter, and then dried and sieved to remove any coarse fraction material not volcanic in origin. Within the Ho deposits, several different eruptive phases were apparent based on particle color and morphology. Based on these characteristics, the different particle types were hand-picked out of the deposit. Thin sections of these particles were prepared for examination under a petrographic microscope in order to determine tephra morphology and mineral phases present.

Trace-element compositions for 2–10 g of both bulk tephra samples and hand-picked pumice and dense glass fragments were determined using an ELAN D CR ICP-MS. The major element compositions of pumice glass, dense glass, and mineral phases were determined using a Jeol JXA-733 Electron Microprobe. A 5–10-μm beam, 15 KV accelerating potential, and 20-nA probe current were used to analyze glass and mineral grains.

Strontium isotopic ratios were measured using a Finnigan-Mat 261 four-collector static Thermal Ionization Mass Spectrometer. Powdered samples were first dissolved in open containers in HF and HC1O3. Replicate analyses of the SRM-987 standard in this mode yielded a mean 87Sr/86Sr of 0.71025 ± 2 (2σ). Measured 87Sr/86Sr were corrected to SRM-987 = 0.710299 ± 8. Errors of 2σ of the mean refer to the last two digits of the 87Sr/86Sr ratio. Details of the analytical procedure are outlined in Farmer et al. (1991).

Results

General

The length of the sediment record for the seven lakes with Ho tephra is variable, but these lakes each contain sediments that date back to late-glacial time (Miranda et al. 2013). The individual cores all preserve >50 to 74 distinct tephra layers derived from explosive eruptions of volcanoes in the area (Figs. 3 and 4). Based on their trace-element chemistry (Fig. 5) and their stratigraphic depth within the cores, specific source volcanoes for some of the Holocene eruptions have been identified (Figs. 3 and 4). These include the previously identified and dated large explosive H2 eruption of Hudson at ~3,920 cal years BP (Naranjo and Stern 1998) and MEN1 of Mentolat at 7,560 cal years BP (Naranjo and Stern 2004; Stern et al. 2013). Other chemically distinctive tephra from late-glacial eruptions that have not been previously described, such as a sequence of nine thin tephra layers dated as <14,910 cal years BP (S9; Figs. 3 and 4) and tephra from a previously undocumented late-glacial Mentolat eruption (MENo), with chemistry (Fig. 5), mineralogy, and glass color and morphology similar to MEN1, are preserved in all or most of the cores and allow late-glacial strato-chronologic correlations between the cores.

Ho deposits

Fifteen cores from the seven lakes studied contain an unusually thick tephra deposit, which provides evidence of a very large late-glacial explosive eruption. Based on their bulk trace-element composition (Fig. 5), discussed in more detail below, and their spatial distribution (Fig. 6), these deposits were derived from a large late-glacial eruption of the Hudson volcano. We term this eruption Ho (Weller et al. 2013). In three cores from Lago Quijada (Fig. 3), the closest to Hudson volcano (Fig. 2), the Ho layer ranges from 78 to 88 cm in thickness. In ten of the 11 cores from five other lakes, all ≥100 km northeast of the Hudson volcano, the Ho tephra layer ranges between 50 and 70 cm in thickness (Fig. 5; Table 2). Only in one of the cores, from Lago Toro, is Ho less than 50-cm thick (Table 2). Pumice size was not systematically measured in any of the cores, but the bulk samples obtained from the cores for petrochemistry contain Ho pumice grains up to 2 cm in maximum diameter. In a core from the seventh lake, Lago Élida (Figs. 1 and 2), more to the east of Hudson, and southeast of the other lakes, Ho is only 18-cm thick (Fig. 4; Table 2) and all pumice grains are <1 cm in maximum diameter. In another core from Lago Shaman (LS; Fig. 1) more than 200 km to the northeast, a 2-cm-thick late-glacial tephra (tephra V), one among 26 tephra identified in this core (de Porras et al. 2012), has tephra color, morphology, and chemistry (Fig. 5; Table 3) similar to Ho.

Table 3 Trace-element composition (in ppm) of bulk samples of Ho
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Both over-thickening and/or compaction of tephra fall deposits may occur in lake sediments. However, 15 of the cores from the lakes that contain Ho tephra also contain layers of H2 tephra that are similar in thickness to the regionally mapped isopachs for this tephra (Table 2; Naranjo and Stern 1998), suggesting that over-thickening was not a problem for this younger tephra in these small lake basins. Based on this as well as the lack of evidence for mixing of other sediment types within the Ho layers and the consistency of the thickness of the Ho layers both in different cores taken from different locations and depths in individual lakes and among the different lakes, we conclude that the Ho layers in these lakes have also not been significantly thickened.

Tentative 50- and 10-cm tephra isopachs for the Ho eruption are presented in Fig. 6 and compared to the 50-cm and the 10-cm tephra isopachs from the H1 eruption (Naranjo and Stern 1998). The area of these isopachs (Fig. 7) indicates that the late-glacial Ho eruption was larger than any of the other, younger, explosive eruptions of Hudson (H1, H2, or H3 in 1991 AD), and we estimate that its tephra volume exceeds 20 km3.

Fig. 7
figure 7

Thickness (cm) versus (Area km2)1/2 of tephra from the H1, H2 (Naranjo and Stern 1998), and Ho eruptions

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The age of the Ho tephra layer has been inferred using 14C methods to date organic matter located in sedimentary layers just above and below this tephra in the cores (Table 1; Miranda et al. 2013). This provides an age for this eruption of between 17,300 and 17,440 cal years BP. This age range is consistent with its stratigraphic position relative to other dated tephra in these cores (Fig. 3). Late-glacial tephra HW1 (14,560 cal years BP) and HW2 (13,890 cal years BP) documented by Haberle and Lumley (1998) from the Taitao Peninsula to the southwest of the Hudson volcano appear to be too young to be distal Ho tephra, and also occur in the opposite direction of the northeastern dispersal direction of Ho (Fig. 6).

Bulk tephra chemistry

All of the bulk tephra trace-element analytical results (Table 3) are similar to each other and occupy a narrow range of trace-element chemistry characteristic of samples of lavas, pumice, and bulk tephra from other eruptions of Hudson volcano (Fig. 5). They have elevated concentrations of Ti (>7,000 ppm), Rb (>30 ppm), Zr (>200 ppm), Nb (>12 ppm), Y (>30 ppm), Ba (>500 ppm), La (>30 ppm), and Yb (>4 ppm) relative to both more-mafic and felsic eruptions from other volcanoes in the region (Stern 1991, 2008; Naranjo and Stern 1998; Kratzmann et al. 2010).

Tephra components

Clast morphology, mineralogy, and composition

The Ho tephra layers contain a diversity of glass particle morphologies and color (Fig. 8). The different components identified include dark-grey to light-tan glassy vesicle-rich pumice, denser dark black, vesicle-poor glassy fragments with a blocky morphology, and isolated mineral grains. The different components are randomly mixed within the deposit, which does not exhibit any vertical stratification. This may be due in part to the deposition of these layers in lakes rather than as primary fall deposits.

Fig. 8
figure 8

Images of the different components identified in the Ho tephra deposit. a The three different color pumice groups selected from the deposit. The light-tan pumice contains glass with ~66 wt.% SiO2 while the light- and dark-grey pumice contain glass with 59.3 % SiO2 and 57.7 % SiO2, respectively (Table 4). b, d Images of the nonvesicular and vesicular blocky dark glassy material. c, e Back-scattering electron images of the nonvesicular dense black glass with few mineral microlites and the more microlite-rich vesicular dark glassy material, respectively

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The dense blocky black glass fragments themselves have a diversity of morphologies and textures, including fragments which contain few vesicles and few mineral microlites (Fig. 8) as well as fragments with higher abundances of spherical vesicles and olivine, plagioclase, and clinopyroxene microlites, along with minor ilmenite. These blocky dark glassy fragments are the most mafic basaltic-trachyandesite component from the Ho eruption (Table 4), with an average glass composition of 55 wt.% SiO2 (Fig. 9), higher bulk Ti, Mn, and Sr and lower Rb, Ba, Zr, La, and Yb contents, and higher Dy/Yb ratios (Fig. 10), than any pumice grains.

Table 4 Major and trace-element (in ppm) compositions of selected Ho components
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Fig. 9
figure 9

Total alkali content versus silica diagram, modified after Irvine and Baragaar (1971), showing the range of compositions of the Ho tephra glasses. The squares represent the light-tan pumice, the circles represent the light-grey pumice, the triangles represent the dark-grey pumice, the diamonds represent the dark glassy vesicle-poor material in the Ho deposits, and the plusses represent glass from other Hudson eruptive products (Futa and Stern 1988; Naranjo and Stern 1998; Gutiérrez et al. 2005; Kratzmann et al. 2009, 2010)

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Fig. 10
figure 10

Dy/Yb ratios of bulk Ho components versus their glass SiO2 contents (Table 4) compared to the compositions of H1, H2, and H3 glasses plotted by Kratzmann et al. (2010). The dashed line from the average phase 1 basalt erupted in 1991, through the average Ho dense glass basaltic andesite to the Ho trachydacite pumice illustrates a small decrease in Dy/Yb with increasing SiO2, indicative of amphibole fractionation as shown in the inset (Davidson et al. 2007), although amphibole does not occur in any of the Hudson eruptive products

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The small glassy pumice fragments prepared for microprobe analysis contain few or no mineral grains, or these grains were possibly plucked out during the polishing processes. Three separate groups were identified and hand-picked from the tephra deposits on the basis of color: dark grey, lighter grey, and light tan (Fig. 8). The glass in the dark- and lighter-grey pumice groups ranges only between 57.7 and 59.3 wt.% SiO2 (Fig. 9; Table 4) despite their distinct color differences. Although the distinction between the major element chemistries between the glasses in the two grey pumice groups is minor, the bulk-pumice trace-element abundances show clear differences between the two components, with the lighter-grey phase having lower Ti, Mn, and Sr and higher Rb, Ba, and Zr contents, and lower Dy/Yb ratios (Fig. 10), than the dark-grey pumice. The lightest-tan-colored pumice is volumetrically less significant than the darker components, and its glass has an average trachydacite composition of 66 wt.% SiO2 (Fig. 9), with even lower Ti, Mn, and Sr and higher Rb, Ba, Zr, and La contents and lower Dy/Yb ratios (Fig. 10). Thus, nearly the entire range of compositional diversity observed in the H3 Hudson eruption of 1991 AD (Kratzmann et al. 2009) is also present in the Ho tephra, and a similar compositional gap in groundmass glass as seen in the 1991 eruption is also present in the Ho eruption. However, the compositional gap between the dark- and lighter-colored pumice components in Ho glasses, which corresponds to ~7 wt.% between 59 and 66 wt.% SiO2 occurs at a higher silica content than the gap in SiO2 content in the 1991 Hudson eruption, which occurs between 54 and 60 wt.% SiO2.

Figure 9 plots total alkali content versus silica, showing the compositional classification of the various glass components from the Ho eruption. Included in this diagram are tephra from phase 1 and phase 2 of the H3 eruption in 1991 AD (Kratzmann et al. 2009) and other Hudson volcanic products (Futa and Stern 1988; Naranjo and Stern 1998; Gutiérrez et al. 2005). The glasses from the Ho eruption are high K2O and range in composition from basaltic-trachyandesite to trachydacite. Like other Hudson eruptive products, they straddle the curve dividing the alkaline and subalkaline fields described by Irvine and Baragaar (1971).

The light-colored felsic pumice has a bulk 86Sr/87Sr ratio of 0.70445 and the more mafic vesicle-poor dark glassy material has a ratio of 0.70447 (Table 5). These isotopic compositions are very similar and coincide with the strontium isotopic ratios measured on all other bulk tephra and effusive rock samples from previous Hudson eruptions, which overall range in composition from 51 to 66 wt.% SiO2 and 525 to 227 ppm Sr (Fig. 11; Futa and Stern 1988; Stern 1991, 2008; Naranjo and Stern 1998).

Table 5 Sr-isotopes and Sr and SiO2 contents for bulk Ho components and H1, H2, and phase 1 (P1) of H3 in 1991 AD
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Fig. 11
figure 11

87Sr/86Sr ratios versus Sr content for components from the Ho eruption (circles; Table 5), and bulk rock samples from phases 1 and 2 of H3 in 1991 AD (square), H1 (triangles), H2 (diamonds) eruptions (Naranjo and Stern 1998; Stern 1991, 2008), and a basaltic andesite lava sample H-1 from the base of the volcano (Futa and Stern 1988)

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Mineral chemistry

Olivine microlites analyzed from the most mafic dark glassy fragments are generally anhedral to subhedral and commonly occur as aggregates with other microlites such as clinopyroxene, plagioclase, and Fe-Ti oxides. Analyzed olivines occupy a narrow compositional range of Fo75 to Fo67, with olivine microlites from Lago Élida (LÉl) being slightly less MgO rich (Fo70–67) than the olivines in the blocky black glass separates from Lago Mellizas (LM) and Lago Churasco (LC; Table 6; Fig. 12). Clinopyroxene microlites in the dark glassy material also have a restricted range of En48–41, Wo43–35, Fs20–16, and clinopyroxene phenocrysts, which occur as large grains with some attached glass, occupy a similar compositional range En49–46, Wo35–27, and Fs23–19. Clinopyroxenes from LÉl are similar to those analyzed in both LM and LC (Table 6, Fig. 12). Olivine and clinopyroxene microlites observed in the fallout deposits from the H3 1991 AD eruption (Kratzmann et al. 2009), and other Hudson eruptive products (Gutiérrez et al. 2005; Kratzmann et al. 2010) share similar compositions to the mineral phases observed in Ho.

Table 6 Compositions of olivines and pyroxenes in the Ho deposits
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Fig. 12
figure 12

Wo-En-Fs classification diagram for pyroxenes and olivines. Clinopyroxene and olivines are present as microlites in dense black glassy fragments, while the orthopyroxene and some clinopyroxenes occur as large dislodged phenocrysts with some adhering glass. Other Hudson mineral analysis from Gutiérrez et al. (2005) and Kratzmann et al. (2009, 2010)

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No orthopyroxene microlites were observed within the blocky glass fragments or pumice, but orthopyroxene phenocrysts were found as large isolated grains mixed within the bulk tephra deposits. The majority of these crystals are elongate euhedral grains with some adhering glass. They have a restricted compositional range of En61-59, Wo4-3, and Fs38-36 (Table 6; Fig. 12). Orthopyroxene with similar chemistry also occurs in H1 and H2 deposits and was reported as well in trace amounts from the H3 trachydacite tephra (Kratzmann et al. 2009, 2010).

Plagioclase occurs both as microlites contained in the dark blocky mafic glass fragment and also as larger phenocrysts, which like orthopyroxenes have become dislodged from the pumice and tephra in the bulk deposits, but still have glass adhering to their crystal surfaces. Plagioclase microlites within the black glassy fragments are generally euhedral to subhedral and are the dominant microlite observed in this mafic component of the tephra deposit. They have a narrow compositional range of An51–58 (Table 7; Fig. 13) with most of the grains being unzoned. The larger plagioclase phenocrysts occupy a larger compositional range of An41–55 and also have no core-to-rim variation. Included in Fig. 13 are plagioclase compositions from other Hudson volcanic products (Gutiérrez et al. 2005; Kratzmann et al. 2009, 2010), which span a wide range of compositions, with the largest population clustering near the composition of the plagioclase microlites and phenocrysts in the Ho tephra.

Table 7 Electron microprobe analysis of representative plagioclase micolites and phenocrysts from the Ho eruption
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Fig. 13
figure 13

Ab-An-Or compositional triangle for feldspar microlites in dense dark glassy fragments and for large dislodged phenocryst. Other Hudson mineral analysis from Gutiérrez et al. (2005) and Kratzmann et al. (2009, 2010)

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Discussion

Hudson eruptive history

The distribution of the Ho tephra across the area indicates a northeast dispersion pattern of the ash fall with a maximum thickness of 88 cm at a distance of ~100 km from the source vent (Figs. 2 and 6). Based on comparison with all of the three previously documented large Holocene Hudson eruptions H1, H2, and H3 in 1991 AD (Stern 1991, 2008; Scasso et al. 1994; Naranjo and Stern 1998), Ho is larger, with an estimated volume of >20 km3 (Fig. 7) of pyroclastic material. This eruption may have caused or initiated the incremental formation of Hudson’s 10-km-wide summit caldera. The data suggest a sequential temporal decrease in the total volume of each large explosive eruption, from Ho >20 km3, to H1 > 18 km3, to H2 and finally H3 in 1991 AD with only ≥4 km3 (Fig. 14), and supports the suggestion of Watt et al. (2013) that deglaciation resulted in generally increased volcanic activity compared to the present.

Fig. 14
figure 14

Volumes produced by explosive eruptions of different later-glacial and Holocene ages (cal years BP) from volcanoes in the southern SVZ between latitudes 39°S (Llaima) and 46°S (Hudson), modified from Watt et al. (2013) to include the large late-glacial Ho eruption from Hudson (>20 km3). Hudson volcano, with four large explosive eruptions and at least 11 other smaller explosive events, totaling >45 km3 volume over the last 17,500 years, has produced the greatest volume of explosive eruptive products in the SSVZ over this time period

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Adding the total estimated volumes of Ho to that of the previously documented eruptions H1, H2, and H3 in 1991 AD, as well as of at least 11 smaller eruptions (Fig. 14), shows that ≥45 km3 of pyroclastic material has been erupted from Hudson in the last ~17,500 years (Fig. 14). This would make Hudson the most productive volcano in the southern Andes in terms of total volume of pyroclastic material erupted by explosive eruptions since the beginning of deglaciation at approximately 17,900 cal years BP.

Petrogenesis

An interesting distinction between the Ho eruption and the H3 eruption in 1991 AD is the lack of stratification of the Ho deposits with regard to the size or composition of the distinct components in the deposit, even though both eruptions contain bimodal eruptive products. The H3 eruption had a clear temporal distinction between the two chemically different phases as the eruption evolved from basalt and basaltic andesite in phase 1 to trachydacite and trachyandesite in phase 2. The genetic relationship between the two phases has been attributed to a combination of magma mixing and fractional crystallization (Kratzmann et al. 2009), who suggested that fractional crystallization of plagioclase (7.8–18.0 %), clinopyroxene (4.8–11.6 %), orthopyroxene (1.9–6.9 %), magnetite (2.9–6.7 %), and apatite (0.5–1.5 %) from a mafic parental magma similar to the P1 basalt is capable of generating magma with a trachydacite composition similar to the dominant tephra observed in phase 2. Kratzmann et al. (2010) subsequently presented similar fractional crystallization models, involving the anhydrous phases observed in the eruption products, for producing the H1 and H2 trachyandesites and trachydacites from a basalt similar to that erupted during phase 1 of the H3 eruption in 1991 AD. However, they further suggested that trace-element concentrations and ratios, specifically decreasing Dy/Yb ratio trends through the sequence basalts, trachyandesites, and trachydacites erupted in these three eruptions (Fig. 10), could be better explained by amphibole fractionation from a basalt parent at >6-km depth (Davidson et al. 2007), followed by disappearance of amphibole, due to decreasing pressure and increasing temperature, as a result of decompression crystallization, as these magmas rose toward the surface. Ho components exhibit similar Dy/Yb ratio trends (Fig. 10) consistent with this latter cryptic amphibole fractionation model.

The strontium isotope ratios of the mafic dark glassy fragments and the felsic light-tan pumice in the Ho tephra deposits, as well as all the other analyzed extrusive units from the Hudson volcano, which overall range in SiO2 from 51 to 66 wt.%, show almost no variation (Fig. 11; Table 5). The homogenous isotopic composition between the different samples is consistent with the suggestion of Kratzmann et al. (2009, 2010) that the evolution of Hudson magmas are controlled by fractional crystallization processes and that crustal assimilation is not responsible for the generation of the more evolved phases observed in either the H3 eruption in 1991 or any of the other documented Hudson eruptions. It is likely that crystal-liquid fractionation process similar to those suggested by Kratzmann et al. (2009, 2010) for the H1, H2, and H3 eruptions controlled the evolution of the Ho eruptive products, which also range from basaltic-trachyandesite to trachydacite in composition.

The Ho deposits are overall dark in color, with a much greater abundance of mafic dark glass fragments and dark- to light-grey pumice compared to light-tan felsic pumice. The bulk tephra trace-element abundances (Table 3), which are a representative portion of each deposit, can also be used to evaluate the relative proportions of mafic and felsic components (Table 4) in this deposit. Based on the comparison of the bulk-tephra, dense glassy material, and the pumice chemical analysis, the compositions of the bulk-tephra samples (Table 3) are more similar to that of the darker, more mafic dense glass and pumice components, confirming that the eruption was predominantly of relatively mafic basaltic-trachyandesite with only a smaller portion of more felsic trachydacite material. Thus, although Ho at ~17,400 cal years BP was bi-modal, it erupted a much greater proportion of mafic basaltic-trachyandesite compared to more felsic material. H1 at 7,750 cal years BP was essentially trachyandesitic in composition (Naranjo and Stern 1998). H2 at 3,960 cal years BP was more felsic than H1, being composed essentially of trachydacite. H3 in 1991 AD was again bi-modal, but with a much smaller proportion of mafic basalt compared to felsic trachydacite material (Kratzmann et al. 2009). We conclude that the large explosive eruptions of Hudson have evolved on average from more to progressively less mafic through late-glacial to historic times, and their volumes have decreased.