Field trip guide: Neogene evolution of the central Andean Puna plateau and southern Central Volcanic Zone
Published:January 01, 2008
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Suzanne Mahlburg Kay, Beatriz Coira, Constantino Mpodozis, 2008. "Field trip guide: Neogene evolution of the central Andean Puna plateau and southern Central Volcanic Zone", Field Trip Guides to the Backbone of the Americas in the Southern and Central Andes: Ridge Collision, Shallow Subduction, and Plateau Uplift, Suzanne Mahlburg Kay, Víctor A. Ramos
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This seven-day field trip is designed to examine the distinctive magmatic, structural, and sedimentological features of the late Oligocene to Recent evolution of the southern part of the high central Andean Puna plateau and the southern Central Volcanic Zone magmatic arc. The stops for Days 1–5 between 23° and 27°S latitude in Argentina emphasize the distinctive magmatic and structural features of the Puna region, which comprises the southern half of the central Andean Puna–Altiplano plateau. Differences between the northern and southern Puna are highlighted. Among the features to be observed are (1) giant Miocene ignimbrites of the northern Puna, (2) distinctive normal and strike-slip faults and associated shoshonitic lavas of the central Puna, (3) the intraplate and calc-alkaline lavas of the southern Puna, (4) the silicic calderas of the southern Puna, and (5) internally drained salar basins. The stops for Days 6 and 7 between 26.5° and 27.5°S latitude in Chile present a view of the Miocene to Recent frontal arc region on the western side of the plateau. The stops particularly highlight the late Miocene to Pliocene displacement of the magmatic arc front from the Maricunga Belt on the western edge of the plateau to its present position in the Central Volcanic Zone. Evidence for the timing of plateau uplift, changes in the angle of the underlying subduction zone, delamination of the underlying continental crust and mantle lithosphere, and forearc subduction erosion are examined throughout the course of the trip.
DISCLAIMER: This is not a geologic field guide to be used in the traditional sense of a detailed road log, but rather a survey of the southern Puna region that can be done with the use of Google Earth. Latitude and longitude coordinates are given for all stops in WGS84 (world geodetic system) coordinates that can be used with Google Earth (downloadable from the Web) or any other georeferenced imagery. Driving instructions are given where access is possible by paved or high-quality unpaved roads that are present on road maps that are generally available in the region. Due to the lack of information on available road maps, changing driving conditions, and the need for caution in accessing the sites, precise instructions are not given for others. A significant number of these stops are on primitive, unmaintained roads or tracks that require a serious four-wheel drive vehicle, an experienced off-road driver, and supporting equipment (winch, spare tires, jack, etc.) to navigate safely. Many of the stops are in the Atacama Desert, where there is almost no water. The majority of the stops are at high elevations—most are between 3500 and 4500 m (~11,500–14,750 ft); all are over 2000 m (6500 ft). Attention needs to be paid to the potential for altitude sickness (called the Puna in Argentina). There are no cell phone towers, no service stations, or towing facilities in most of the region to bail you out! The availability of fuel can be questionable at times. Contact a knowledgeable guide to the region before attempting to use this field guide in the more isolated parts of the area.
Introduction to The Field Area
Overview of the Central Andean Puna-Altiplano Plateau
The central Andes are known for a number of special features whose origins have been long debated. The most prominent geographic feature is the extensive Puna-Altiplano plateau (Fig. O.1), which after the Tibetan plateau is the world's highest (average elevation 3700 m) and largest (700-km–long and 200-km–wide) plateau. Unlike the Tibetan plateau, which has little volcanic cover, the Puna-Altiplano is covered by an extensive array of Neogene volcanic centers that extend in chains to the eastern side of the plateau. The plateau is underlain by thick continental crust (>50 km) with seismic data indicating thicknesses as much as 65–80 km in the central plateau (Yuan et al., 2002; McGlashan et al., 2008). The highest regions correspond with sites of Neogene volcanic activity. Uplift of the plateau is considered to have largely occurred in the Neogene with the principal cause being crustal thickening in response to crustal shortening along with limited magmatic addition (e.g., Allmendinger et al., 1997). Delamination of the lower continental crust and lithosphere and the resulting thermal conditions have contributed to the high elevation (e.g., Kay and Kay 1993; Kay et al., 1994a). General overviews on the plateau have been presented in Allmendinger et al. (1997), Kay et al. (1999, 2004), Beck and Zandt (2002), Oncken et al. (2006), and references in those papers.
The central Andes are characterized by a dominantly compressional Neogene stress regime not related to continental collision. Contractional deformational belts of Neogene age border the eastern side of the plateau (Fig. O.1; see Allmendinger et al., 1997). From north to south, they include the Subandean and Eastern Cordilleran fold-thrust belts, the Santa Barbara belt where shortening is accommodated by inversion of Cretaceous normal faults, and the high-angle, reverse-faulted Sierras Pampeanas . The amount of shortening is variable, and exact amounts are widely debated (e.g., Kley and Monaldi, 1998; Kley et al., 1999).
The subducting Nazca plate beneath the central Andean plateau is Oligocene to Eocene in age. A feature of this slab is its relatively shallow angle (~30°) compared to other regions around the circum-Pacific where subduction angles are rarely less than 45°. To the north and south of the plateau lie shallower subducting segments under the volcanically quiescent Peruvian and Chilean flat-slab regions (Fig. O.1; Isacks 1988; Cahill and Isacks, 1992). The southward transition to the Chilean flat slab is relatively smooth, whereas that with the Peruvian flat-slab segment is abrupt. Flattening is expressed as a bench that develops in the seismic zone between 90 and 135 km (Fig. O.1; Cahill and Isacks, 1992). Gephart (1994) emphasized the degree of bilateral symmetry in the shape of the seismic zone and topography of the land surface from 33°S to 5°S latitude. Isacks (1988) argued that the modern geometry of the subducting Nazca plate and uplift of the central Andean plateau are the result of the oroclinal bend in the South American plate overriding the subducting Nazca plate. Models for the evolution of the plateau call for a “collision” between the shallowly dipping Nazca plate and the overriding South American plate. Recent papers have emphasized the importance of the westward drift of South America (Sobolev and Babeyko, 2005; Oncken et al., 2006) as originally proposed by Silver et al. (1998).
The central Andean plateau can be divided into the Puna (~22° to 27°S) and the Altiplano (~22° to 15°S) as discussed by Turner (1970). Some important differences are shown in Figures O.2 and O.3. As summarized by Whitman et al. (1996), Allmendinger et al. (1997), and Yuan et al. (2002), the Puna differs from the Altiplano in that: (a) the basement is generally younger and has a larger component of Paleozoic magmatic rocks; (b) small discontinuous, diachronous basins replace the large Altiplano basin; (c) widespread Miocene to Quaternary volcanic rocks erupted in broad northwest-trending chains separated by regions of basement outcrop (Fig. O.2); (d) average elevations are higher (Figs. O.2 and O.3); and (e) seismic, topographic, and gravity data support a thinner lithosphere and crust (Figs. O.3–O.6). Geophysical characteristics of the northern Puna and Altiplano are better known due to the seismic studies of ANCORP (2003) and the Arizona group (e.g., Beck and Zandt, 2002). The ANCORP seismic profile near 22°S is shown in Figure O.5. A seismic array was installed in the southern Puna by a Cornell-Missouri-Potsdam group in 2007.
The Puna Plateau
The modern Puna plateau shows some important north to south differences as initially suggested by Alonso et al. (1984a).
Among the contrasts are a higher average topography (Isacks, 1988) and a thinner lithosphere in the south. Evidence for the thinner lithosphere comes from asthenospheric shear wave attenuation, lower seismic Q, and effective elastic thickness (Whitman et al., 1992, 1996). Evidence for a thinning lithosphere at the transition between the northern and southern Puna comes from seismic tomography (Schurr et al., 2003, 2006) and Quaternary shoshonitic lavas erupted along faults (Coira and Kay, 1993; Kay et al., 1994a). Gravity data support a region of distinctively thin lithosphere and crust near 25°S (Tassara et al., 2006) and are in accord with seismic data indicating crustal thicknesses near 42–49 km (Fig. O.4). Crustal thicknesses from 22°S to 25°S range from ~50 to 68 km; thicknesses are unavailable farther south (Fig. O.4). The southern Puna also differs from the northern Puna in having normal and strike-slip faults (e.g., Marrett et al., 1994) associated with <7 Ma mafic volcanic rocks (e.g., Kay et al., 1994a, 1999; Risse et al., 2008) and glassy andesitic to dacitic lava flows (e.g., Kay et al., 1999, 2005).
Contrasts in the Magmatic and Structural History of the Northern and Southern Puna
The northern and southern Puna are characterized by distinctive late Oligocene to Recent magmatic, structural, and sedimentological histories that are highlighted below and are the focus of the stops on Days 1–5. The magmatic front west of the southern most Puna in Chile is the focus of Days 6 and 7. The main late Oligocene to Recent magmatic features of the northern and southern Puna are shown in Figures O.6 to O.8 and related to general tectonic models for transects in the northern (~21° to 24°S) and southern (~26° to 28°S) Puna in Figures O.9 and O.10. More complete discussions can be found in Coira et al. (1993), Allmendinger et al. (1997), Kay et al. (1999), Oncken et al. (2006), and references therein. Additional information on the magmatic arc front in the southernmost Puna can be found in Mpodozis et al. (1995, 1996) and Kay et al. (1994a, 2006).
Northern Puna History (Fig. O.6)
The late Oligocene to middle Miocene (28–17–16 Ma) history of the northern Puna is characterized by a general gap in magmatic activity, contractional deformation, and foreland sedimentary basin development from the arc to the backarc of the modern plateau. This situation changes in the middle Miocene at 15–12 Ma as small generally andesitic porphyritic stocks and domes erupted in the far backarc (e.g., Caffe et al., 2002). In the late Miocene, the picture changed dramatically as large ignimbrite centers erupted across the plateau (de Silva, 1989; Coira et al., 1993; Kay et al., 1999) and contractional deformation ceased on the plateau (e.g., Gubbels et al., 1993). The eruptions include those of the ~150 km3 Granada center at ca. 10 Ma (Caffe et al., 2007), the ~1400 km3 Vilama center at ca. 8.5 Ma (Soler et al., 2007), the ~650 km3 Panizos center at ca. 6.7 Ma (Ort, 1993; Ort et al., 1996), and the ~650 km3 Coranzulí center at ca. 6.7–6.4 Ma (Seggiaro and Aniel, 1989; Seggiaro, 1994) (volumes in dense rock equivalents [DRE]). After ca. 10 Ma, the main locus of foreland basin formation and thrusting shifted to the Subandean Belt to the east where thrusting began at ca. 9–8.5 Ma in the west and at ca. 6.9 Ma to the east (Echavarria et al., 2003). Major surface uplift at this time based on paleobotany (Gregory-Wodzicki, 2000) is supported by oxygen isotopic analyses of carbonates (Garzione et al., 2006). Large ignimbrite eruptions continued in the latest Miocene to Pliocene with the centers concentrated farther to the west. The eruptions include those of the >1600 km3 La Pacana caldera at ca. 4 Ma (e.g., Lindsay et al., 2001) and the ~1500 km3 Puripicar ignimbrite erupted at ca. 4.2 Ma (Barquero-Molino, 2003; de Silva and Gosnold, 2007). By the latest Pliocene, nearly all magmatic activity was concentrated near the modern Central Volcanic Zone arc front. Thrusting continued in the Subandean Belt with out-of-sequence thrusting occurring after 4.5 Ma (Echavarria et al., 2003).
From 27 to 19 Ma, widespread ignimbrite and dacitic dome complexes were erupting in the frontal arc west of the Puna. Backarc activity was limited to mafic volcanism just east of the arc near 26°S (Segerstrom basalt; Kay et al., 1999) and small dacitic to rhyodacitic complexes in the Puna (Coira et al., 1993). This picture changed between ca. 20 and 16 Ma when small stratovolcanoes and minor ignimbrites began erupting in the backarc and contractional deformation took place east of the Maricunga Belt (Mpodozis and Clavero, 2002). By the middle Miocene, 16–12 Ma andesitic stratovolcanoes were erupting in the Maricunga Belt arc, and volcanism initiated at ca. 15–14 Ma at the long-lived backarc stratovolcanic complexes like Beltran, Antofalla, and Tebenquicho (e.g., Coira et al., 1993; Kraemer et al., 1999; Richards et al., 2006). Ignimbrites erupted in the transitional region between the northern and southern Puna near 24°S (Petrinovic, 1999; Petrinovic et al., 1999). The magmatic style changed at 11–7 Ma as volcanism in the Maricunga Belt arc became concentrated in the Copiapó dacitic ignimbrite-dome complex (Mpodozis et al., 1995; Kay et al., 1994b) and local 11–10 Ma ignimbrites in the backarc preceded andesitic lavas and dome complex formation in the long-lived Puna stratovolcanic complexes. Contractional deformation continued in the southern Puna, and the Pampean ranges east of the Puna began to uplift at ca. 9–8 Ma. By ca. 6–5 Ma, volcanism had terminated in the Maricunga Belt arc as frontal arc activity shifted eastward toward the Central Volcanic Zone arc (Mpodozis et al., 1996). A major change occurred in the Puna at ca. 6.7 Ma as mafic lavas erupted along normal and strike-slip faults (e.g., Kay et al., 1999) and the Cerro Galán ignimbrite eruption began (Sparks et al., 1985). By ca. 4 Ma, ignimbrites were erupting in both the arc (Laguna Verde, Amarga, Vallecito; Mpodozis et al., 1996; Siebel et al., 2001) and backarc (Real Grande ignimbrite at Cerro Galan; Sparks et al., 1985). The arc front was stabilized in the Central Volcanic Zone (Fig. O.8.; Mpodozis et al., 1996) by 3–2 Ma and the ~1000 km3 Cerro Galán ignimbrite erupted at ca. 2.2 Ma (Sparks et al., 1985). During this time, intraplate-like mafic lavas erupted over the modern gap in intermediate depth slab seismicity (intraplate region in Figs. O.1B and O.8) and arclike magmas erupted to the north and south (Kay et al., 1994a, 1999). Since 2 Ma, silicic magmas have erupted from the Cerro Blanco caldera, and Puna mafic volcanism has continued (Siebel et al., 2001; Kay et al., 2006; Risse et al., 2008).
Models for the Neogene Magmatic-Tectonic Evolution of the Northern and Southern Puna
Northern Puna: Shallowing and Steepening of the Subduction Zone (Fig. O.9)
The model for the evolution of the northern Puna in Figure O.9 involves a transition from a very shallow, late Oligocene to early Miocene subduction zone to a moderately steep subduction zone (Coira et al., 1993; Kay et al., 1999). Periodic delamination of the lower crust and lithosphere could have accompanied slab steepening (e.g., Beck and Zandt, 2002).
Stage 1 at 26–14 Ma: The virtual lull in volcanism, widespread contractional deformation, and basin formation across the arc and backarc are consistent with a shallowly dipping subduction zone like that under the modern Chilean flat-slab region. A similar Oligocene flat slab has been suggested by James and Sacks (1999) to explain geophysical and geological observations under the southern Altiplano.
Stage 2 at 14–6 Ma: Steepening of the subduction zone after ca. 14 Ma is consistent with small backarc dacitic eruptions preceding widespread, voluminous late Miocene ignimbrite eruptions. The transfer of contractional deformation into the Subandean Belt and important uplift of the plateau after 10 Ma fit with the model of Isacks (1988) in which upper crustal shortening is compensated by ductile lower crust under the plateau at the time of uplift. Accumulation and fractionation of lower crustal melts at the brittle-ductile transition near 20 km depth is in accord with magma chambers inferred near that depth from geophysical data by Zandt et al. (2003) and on the ANCORP (2003) profile (Fig. O.5). The horizontal compressional failure of melt-weakened crust can explain the transfer of these magmas to the shallow crustal chambers from which they are inferred to erupt (e.g., Lindsay et al., 2001). The major ignimbrite eruptions are also potentially linked to the crustal failure that produced the Subandean thrusts. Delamination of the crust and mantle lithosphere under the northern Puna and southern Altiplano (Yuan et al., 2002; Beck and Zandt, 2002; Garzione et al., 2006) could have enhanced crustal melting by intrusion of mantle melts as argued by Kay and Kay (1993) for Cerro Galán. However, in detail, this delamination needs to differ from that under the southern Puna because the mafic lavas and mixed extensional and/or strike slip and/or contractional fault system found in the southern Puna are essentially absent.
Stage 3 at 6–3.8 Ma: Further steepening of the subducting slab is consistent with eruption of giant ignimbrites near the modern arc front and continuing contraction in the Subandean Belt.
Stage 4 at <3.8 Ma: In the most recent stage, andesitic and dacitic magmatism has been essentially confined to the arc front region by the steeper subduction zone, and shortening has continued in the Subandean Belt. The eruption of small backarc shoshonitic lavas (Kay et al., 1994b) and seismic images (e.g., Beck and Zandt, 2002; Schurr et al., 2006) are in agreement with removal of pieces of lithosphere beneath the Eastern Cordillera.
Southern Puna: Shallowing to Steepening Subduction and Crust and Mantle Delamination (Fig. O.10)
The model for the southern Puna in Figure O.10 involves moderate middle Miocene shallowing of the subduction zone, late Miocene steepening, and late Miocene to Pliocene lower crustal and lithospheric delamination coincident with frontal arc migration. The delamination model is modified from Kay et al. (1994a, 1999) and the entire model is presented in Kay et al., 2005.
Stage 1 at ca. 26–20 Ma: The concentration of latest Oligocene and early Miocene andesitic to dacitic volcanism in the arc is consistent with a relatively steeply dipping slab (Fig. O.8). A noncontractional setting with local extension fits with eruption of voluminous dacitic-rhyolitic domes and ignimbrites in the arc and the Segerstrom mafic lavas in the backarc.
Stage 2 at 19–16 Ma: A change to a compressional stress regime at this time is consistent with contractional deformation in the backarc in Chile (e.g., Mpodozis and Clavero, 2002) and in the Puna to the east (e.g., Kraemer et al., 1999). The initiation of Miocene contraction along much of the Andean chain (e.g., Kay and Mpodozis, 2002; Oncken et al., 2006) at this time has been suggested to be due to the accelerated westward drift of South America over the Nazca plate (Kay and Copeland, 2006; Oncken et al., 2006) in the manner suggested by Silver et al. (1998).
Stage 3 at 15–8 Ma: Shallowing of the subduction zone during this period can explain the eastward broadening of the magmatic arc, the arc-like geochemistry (high field strength element depletion; Kay et al., 2005) of the backarc volcanic rocks and backarc contractional deformation. A period of shallowing fits with subduction of the Juan Fernandez ridge on the Nazca plate playing a role in the shallowing because the reconstruction of Yañéz et al. (2001) shows the ridge approaching and subducting under the region from ca. 14–8 Ma. Shallowing needs to have been less pronounced than under the northern Puna because no temporal magmatic gap is observed.
Stage 4 at 7–4 Ma: A major change in magmatic and deformational style at ca. 7 Ma is consistent with steepening of the subducting slab and delamination of the lower crust and underlying mantle lithosphere. Steepening of the slab is again consistent with the Yañéz et al. (2001) model that shows the Juan Fernandez Ridge subducting farther to the south by this time. A regional change in the deformational style to a complex mix of normal, strike-slip, and contractional faults (Marrett et al., 1994), the eruption of mafic lavas along faults, and uplift of the plateau (Alonso et al., 2006) fit with delamination of gravitationally unstable lower crust into the mantle wedge at this time (Kay and Kay, 1993; Kay et al., 1994a, 1999). The resulting influx of asthenospheric mantle into the thickening mantle wedge can explain mantle melting leading to basaltic volcanism and the widespread crustal melting producing the early Cerro Galan ignimbrite eruptions. This is also the time of the extinction of the Maricunga Belt arc and widespread volcanic activity between the Maricunga Belt, and the future Central Volcanic Zone arc front (Figs. O.7 and O.8; Mpodozis et al., 1995, 1996; Kay et al., 1994b; Kay and Mpodozis, 2002). Migration of the arc front is consistent with a major pulse of forearc subduction erosion (Kay and Mpodozis, 2002; Kay, 2006).
Stage 5 after 3 Ma: By 3 Ma, a large amount of crustal melting following delamination resulted in the major Cerro Galán ignimbrite eruption at ca. 2.2 Ma. A thick mantle wedge fits with eruption of intraplate-like lavas and the geophysical evidence for a thin mantle lithosphere (Sn attenuation; see Whitman et al., 1992, 1996).
Field Trip Log Part 1: Days 1 to 5—The Argentine Puna
Beatriz Coira and Suzanne Mahlburg Kay
The objectives of Days 1–5 are to examine the latest Oligocene to Recent history of the Puna plateau and to briefly view the Paleozoic basement. The route of the trip, which generally progresses from north to south, is shown in Figure 1.1. Day 1 features the Eastern Cordillera and the northern Puna; Day 2 features the transitional zone between the northern and southern Puna; Days 3, 4, and the first part of Day 5 feature the southern Puna, and the last part of Day 5 features the transition to the northern Sierras Pampeanas. A general overview of these regions is presented in the Introduction to the field area.
Day 1—Northern Puna
The objective of Day 1 is to examine a transect from the Eastern Cordillera onto the northern Puna. The route starts in the city of Jujuy, goes to the north and then to the west ascending through the Eastern Cordillera fold-and-thrust belt onto the Puna (see Figs. 1.2 to 1.4).
Stop 1-1: Arroyo del Medio Alluvial Cone
Directions: Take Argentine Route 9 north from the city of Jujuy in the direction of Humahuaca for ~36 km. Stop along road near 23°56′50″S; 65°27′30″W; 2100 m asl.
The large alluvial cone seen here is principally composed of the series of debris flows that dammed the valley of the Rió Grande in 1945. The narrowing of the valley at this point is a result of these debris flows. The large size of the flows reflects both the active seismicity of the region and the coincidence of this sector of the river valley with the line of maximum storm intensity during the rainy season (January and February). The voluminous Holocene to Recent clastic sediments along the Quebrada de Humahuaca to the north further reflects the crustal seismicity that demonstrates the ongoing tectonic activity in the region. A seismic array deployed by the PANDA group (Cahill et al., 1992) along the eastern edge of the Puna, the Eastern Cordillera, and into the Santa Barbara structural belt elucidated the locations of the crustal earthquakes associated with contractional structures along the eastern deformational front. A peak of seismicity at a depth of 20–25 km and fault plane solutions for these earthquakes are in accord with a mid-crustal detachment on which significant east-west late Cenozoic foreland shortening has occurred (Cahill et al., 1992).
Stop 1-2: View of Typical Eastern Cordillera Structures Just East of Purmamarca
Directions: Continue north on paved Route 9 and turn left (west) on paved Route 52 toward the village of Purmamarca. Stop is at a pull-out on the right side of the road ~3 km after the turn and before entering Purmamarca. Stop is at 23°44′50″S; 65°29′39″W; 2313 m asl.
The Eastern Cordillera is a thrust belt with a typical piggy-back structural style that shows both west and east vergence. Balanced cross sections indicate ~60 km of shortening across the Eastern Cordillera. Cenozoic deformation and uplift of the Eastern Cordillera began at ca. 40 Ma and were largely over by 10 Ma (Müller et al., 2002; Ege et al., 2007). Some deformation is still occurring as shown by earthquake activity and evidence for very recent thrust and strike-slip deformation. An array of west- and east-verging thrusts involves late Precambrian, Lower Paleozoic, and Cretaceous to Paleogene stratigraphic units (Figs. 1.3 and 1.5). At this stop, imbricated slices of Cambrian and Cretaceous-Paleocene rocks can be seen thrust over the sedimentary sequences of the Precambrian–Lower Cambrian Puncoviscana Formation. The complex structures in this region are interpreted as having formed from a combination of west-verging compressional thrusts in the Late Ordovician (Ocloyic Event), extensional rift faults in the Cretaceous to Paleogene, and east-verging contractional structures in the Miocene to Pliocene. The multi-episodic deformation resulted in a wide variation in vergence and partial inversion of extensional structures leaving younger over older fault relationships.
This is a tourist spot for lunch and a great place to buy local merchandise typical of the Puna plateau region. Among the attractions is the Santa Rosa de Purmamarca chapel, which is a Jesuit church that was constructed between 1648 and 1779. Paintings of the seventeenth century Cuzco school exhibited in the chapel depict the life of Santa Rosa de Lima. Other paintings are entitled La Piedad and La Inmaculada. At the chapel entrance stands a five-century–old algarrobo tree. The town preserves a Spanish colonial style with a strong indigenous and Argentine overprint. A native market place is located in the main square.
Stop 1-3: View of Alluvial Deposits Looking East from the Cuesta de Lipan
Directions: Continue west on Route 52 up the eastern side of the plateau. Stop at the Cuesta de Lipan near 23°40′19″S; 65°36′13″W; 3390 m asl. View is to the east.
The outstanding sequence of highly colored Pliocene-Pleistocene deposits of the Purmamarca Formation that formed under periglacial conditions can be viewed from this stop. The deposits are strongly dissected by the Purmamarca River marking the ongoing uplift of the Eastern Cordillera. A monadnock, which is an isolated remnant of a former erosional cycle, can be seen in a tributary valley.
Stop 1-4: View of Cretaceous Rocks Affected by Cenozoic Deformation at El Angosto
Directions: Continue west on Route 52 climbing the eastern side of the plateau, passing the summit, and start descending the switchbacks to the west. Stop is at El Angosto near 23°42′W; 65°41′W; 3891 m asl.
Structures associated with Cretaceous rifting events have had a profound influence on the style of Cenozoic thrusting and folding in the Eastern Cordillera. Many north- to north-northeast–trending normal faults that were active during the Cretaceous are oriented approximately perpendicular to the Cenozoic contractional direction. These former normal faults have been reactivated as reverse faults. At this stop, Andean age Miocene-Pliocene thrusts can be seen to cut Cretaceous-Paleocene extensional faults that affected the synrift Cretaceous Pirgua Subgroup. The post-rift deposits of the Paleogene Balbuena and Santa Barbara Subgroups are also involved in the thrusts. Thrusts that put sheets of Precambrian-Cambrian basement over younger units are visible at this stop.
Stop 1-5: Salar de Salinas Grandes
Directions: Continue west on Route 52 to Salinas Grandes. Stop at the tourist stop in the middle of the salar near 23°35′47″S; 65°52′58″W; 3429 m asl.
Salars are a distinctive feature of the Puna. They are basically intermontane basins that have principally developed over the past 7–8 Ma as large volumes of continental evaporites accumulated during the uplift of the Puna. Evaporate deposition is controlled by geothermal activity related to volcanism, internal drainage, and climate. The Salinas Grandes salar at this stop hosts one of the major halite deposits of the Puna. Production is 10,000–20,000 tons per year depending on the weather. Borate deposits, which are mainly from ulexite [NaCaB5O6 (OH)6-5H2O], are exploited in the Tres Morros and Niño Muerto sectors of the salar.
The Salinas Grandes is a symmetrical basin bounded by inverted Cretaceous high-angle normal faults. An east-southeast–trending seismic reflection profile between the Salinas Grandes and Laguna de Guayatayoc from Coutand et al. (2001) in Figure 1.6 shows three main units. Unit 1—Paleozoic basement units and the Early Cretaceous Tusaquillas pluton; Unit 2—Cambrian to Early Ordovician quartzites subjected to folding and faulting prior to Andean deformation (Gangui, 1998); and Unit 3—Mesozoic and Cenozoic strata separated from Units 1 and 2 by a major angular unconformity (Gangui, 1998). The lower units show evidence for Cretaceous normal faults that were inverted during Andean compression (Acevedo and Bianucci, 1987). The Salinas Grandes depocenter lies between two major thrusts that were active during Cenozoic sedimentation.
Stop 1-6: Ordovician Basement (Cobres Plutonic Complex: Churcal and Las Burras Granitoids) and Distal Miocene Ignimbrite Flows
Directions: Continue west on Highway 52, crossing the Salinas Grandes, and stop on the north side of the road after entering the low ranges to the west. Stop is at ~23°25′21″S; 66°11′18″W; 3525 m asl.
The deformed magmatic and sedimentary rocks seen here and along the field trip route to the west (Fig. 1.2) are representative of Ordovician sequences, which form an important component of the Puna basement. The exposures in the northern Puna are dominated by low- to medium-grade metamorphic rocks, syntectonic rhyolites, and rare granitoids. The exposures in the southern Puna are dominantly composed of higher grade metamorphic rocks and abundant granitoids. The dominance of Ordovician basement outcrops in the Puna contrasts with the situation in the Eastern Cordillera, where the principal exposures are Precambrian to Cambrian sedimentary sequences.
The outcrops at Stop 6 (Figs. 1.2 and 1.7) are part of the syntectonic lower Ordovician Cobres Plutonic Complex (Kirschbaum et al., 2006), which is composed of granodiorite and monzogranite plutons. The plutons record a tectonomagmatic event at 476 Ma (Lork and Bahlburg, 1993; Haschke et al., 2005), which is the time of collision of the Laurentian-derived Precordillera (Cuyania) terrane with the Gondwana continent to the south (Coira et al., 1999). An initial low-temperature deformational event produced the intense folding and cleavage in the exposures in this region. Subsequent medium-temperature deformational features can be associated with pluton emplacement and metamorphism. The last event in this tectonic cycle was the intrusion of the post-tectonic Las Burras Granite (428 ± 17 Ma, Zappettini, 1989).
The Ordovician basement and valleys along the field trip route are covered by the distal flows of the Las Termas ignimbrite (6.45 Ma; Seggiaro, 1994) that erupted from the Coranzulí center ignimbrite center (Fig. 1.8).
Stop 1-7: Las Termas Ignimbrite North of the Village of Barrancas
Directions: Turn right (north) on a prominent secondary dirt road near Stop 1-6 to the village of Barrancas. Stop 1-7 is 15 km northeast of Stop 1-6 at the well-exposed ignimbrite exposures near 23°19′50″S; 66° 05′28″W; 3624 m asl. After the stop, return to Route 52.
The Las Termas ignimbrite (6.45 ± 0.15 Ma; Seggiaro, 1994) at this stop is the longest and youngest of the four ignimbrite sheets that erupted from the late Miocene Coranzulí caldera (Figs. 1.7 and 1.8). The Coranzulí caldera, which is some 35 km to the north (Fig. 1.7), is better seen at the next stop. The Las Termas ignimbrite is well exposed all along the Barrancas valley, where its distribution was controlled by preexisting relief. The Las Termas ignimbrite is a moderately to highly welded, crystalrich (plagioclase, sanidine, and biotite) dacite with a moderate amount of pumice. The flow unconformably overlies deformed sedimentary deposits of the middle Miocene Pastos Chicos Formation and puts a minimum age on the last stage of contractional deformation in this region of the Puna.
Stop 1-8: View of Coranzulí Caldera–Resurgent Center and Ignimbrite Flows
Directions: Continue west on Route 52, stop near 23°26′26″S; 66°17′29″W; 3852 m asl.
The Coranzulí Caldera complex, seen to the north at this stop, is located at the intersection of major northeast- and north-trending regional faults (Fig. 1.8) that controlled the relative movement of the Rinconada and Cochinoca structural blocks. These faults, together with the northwest-striking Ramallo fault, define a transtensional system, which controlled the development of the collapse caldera (Seggiaro, 1994). On a regional scale, the northeast-trending faults define part of the Lipez-Coranzulí lineament that traverses the Puna. Volcanic centers aligned along this lineament form one of the distinctive transverse Puna volcanic chains. According to Riller et al. (2001), onset of activity on these faults at ca. 10 Ma marks an important temporal transition from a deformational regime controlled by dominantly vertical thickening to one controlled by orogen-parallel stretching as critical crustal thicknesses were achieved in the Puna.
The Coranzulí Caldera complex erupted four discrete ignimbrite sheets between 6.8 and 6.5 Ma (Seggiaro et al., 1987). These ignimbrites unconformably overlie deformed Miocene sedimentary sequences that fill the Pastos Chicos sedimentary-tectonic depression. The combined volume of the four ignimbrites is estimated to be over 650 km3 (Seggiaro et al., 1987). All four have rhyodacitic compositions, moderate amounts of pumice, high crystal contents, and moderate welding. The proximal and distal facies are similar. Ash-fall deposits are scarce. Ground surge deposits can be seen in some regions. The eruption of the 6.5 Ma Las Termas ignimbrite can be related to the collapse caldera of the Coranzuli center, which has a diameter of 5 km. Eruption was followed by a magmatic resurgence that is represented by the late Miocene Coranzulí lava dome. The dome fills and partially obscures the caldera depression.
DAY 2—TRANSITION ZONE BETWEEN THE NORTHERN AND SOUTHERN PUNA
The objective of Day 2 is to examine the Neogene evolution of the transition zone between the northern and southern Puna. The route will cross the El Toro lineament, which is the boundary between the southern and northern Puna (Alonso et al., 1984b). South of 24°S, late Neogene strike-slip and extensional faults, small discontinuous and diachronous sedimentary basins, and mafic lavas appear.
Stop 2-1: Miocene Pastos Chicos Formation in Front of Pastos Chicos Hostería on Edge of Village of Susques
Directions: 23°25′16″S; 66°22′58″W; 3660 m asl. After the stop, take a secondary dirt road just east of the Pastos Chicos Hostería to the south in the direction of the villages of Huancar and Sey (see map in introduction to Days 1–5).
This stop provides a view of clastic sedimentary sequences in the Pastos Chicos Formation that lie unconformably above deformed lower Ordovician sedimentary basement rocks. A tuff in the Pastos Chicos Formation has been dated at 9.5 ± 0.3 Ma (Schwab, 1973). At this stop, the Pastos Chicos Formation is unconformably covered by the undeformed 6.4 Ma Las Termas ignimbrite from the Coranzulí center. South of Susques and the Rio Sijes, the Pastos Chicos Formation unconformably overlies deformed sediments in the Miocene Trinchera Formation (Schwab, 1973) that were deposited unconformably on Ordovician and Cretaceous units. The Trinchera Formation consists of clastic and calcareous beds intercalated with tuffs and ignimbrites that have yielded a K/Ar age of 10.8 ± 0.3 Ma (Schwab and Lippolt, 1976). On the flank of the Salar de Olaroz-Chachari to the west (see Figs. 1.7 and 2.1), Ordovician sedimentary rocks are thrust over conglomerates in the Pastos Chicos Formation. To the north, undeformed ca. 10 Ma Granada and ca. 8–7 Ma Vilama ignimbrites (Fig. 1.8) overlie deformed Miocene clastic and pyroclastic rocks (Soler et al., 2007). The regional relations show that late Miocene shortening was going on in the Susques region before the eruption of the Las Termas ignimbrite at 6.4 Ma and that deformation had ceased by 9–10 Ma. Similar field relations have been used to argue that contractional deformation had ceased in the northern Puna and southern Altiplano and shifted into the eastern foreland by ca. 10–9 Ma (Gubbels et al., 1993).
Optional Stop near Village of Huancar: Lower Ordovician Sedimentary and Volcanic Rocks under Deformed Late Miocene Sediments of Pastos Chicos Basin
The deformed Lower Ordovician turbidities, synsedimentary dacitic lavas and hyaloclastites, and sparse mafic lavas that form the basement under the Miocene sequences (Coira, 1996; Coira et al., 1999) are well exposed near the village of Huancar. As elsewhere in the northern Puna, the Ordovician sequence is characterized by low-grade metamorphism. The overlying late Miocene clastic, evaporitic, and pyroclastic sequences of the Miocene Trinchera and Pastos Chicos Formations record the evolution of the Pastos Chicos intermontane basin between 10.8 and 8.9 Ma (Schwab and Lippolt, 1976).
Stop 2-2: Eruption Sequence and Origin of Tuzgle Volcano
From this stop, the Tuzgle volcanic complex can be seen to the south. Tuzgle volcano is on the northern side of the El Toro lineament, which is a first-order, northwest-trending transtensional fault system that crosses the plateau (see Fig. 2.1). The small San Gerónimo and Negro de Chorrillos shoshonitic volcanoes seen at Stop 2-7 are along the southern side of the lineament (Figs. 2.4 and 2.5). On a regional scale, the Tuzgle volcano is located near the eastern edge of the Puna plateau, some 275 km east of the main front of the Central Volcanic Zone and some 200 km above the Wadati-Benioff zone. The underlying subducting Nazca plate marks the northern end of the transition zone where the slab begins to shallow into the Chilean flat slab south of 28°S (Cahill and Isacks, 1992; Figs. O.1 and O.2).
The Tuzgle eruptive sequence in Figure 2.2 is discussed in Coira and Kay (1993). Volcanic activity (see Fig. 2.2) began at 0.5 ± 0.2 Ma (Aquater, 1980) with the eruption of ~0.5 km3 of rhyodacitic ignimbrite that flowed to the north. The ignimbrite was followed by the 0.3 ± 0.1 Ma Old Complex dacitic lava dome complex that has a total volume of ~3.5 km3. Next, the Old Complex was partially covered by the pre-platform andesitic lava flows that erupted from a new crater. This new crater was in turn partially filled by the Platform mafic andesite flows. These early-stage Tuzgle volcanic rocks were then cut by northwest- and east-trending faults that controlled the emplacement of the Post-platform and Young Tuzgle mafic andesite lava flows (Coira and Paris, 1981). The total volume of the Pre-platform and Young lavas is ~0.5 km3.
The Tuzgle region is an active geothermal area (Coira, 1995) with the Tuzgle and Agua Caliente hot springs having temperatures of 40 °C to 56 °C and the Planta Mina hot springs having a temperature of 21 °C. Geothermal reservoir temperatures are 132 °C to 142 °C. Geochemical studies indicate mixing of deep and shallow water.
Petrologic and geochemical data show that the Tuzgle volcanic rocks are notably diverse and require variable thermal conditions and diverse mantle and crustal components to explain their origin (Coira and Kay, 1993). One general group includes the Tuzgle ignimbrite, Old Complex, and pre-platform units, which all have intraplate-like high field strength element (HFSE) signatures and rare earth element (REE) patterns with La/Yb <30. A second group includes the post-platform and Young andesitic flows, which all have arc-like HFSE anomalies and very steep REE patterns with La/Yb >35. A third group in the region includes the San Gerónimo and Chorrillos shoshonites to be seen at Stop 2-7. All of these magmas have isotopic and trace element evidence for upper crustal contamination. Major element and 87Sr/86Sr ratio similarities between plagioclase (~AN33) xenocrysts in plagioclase-phenocryst–free mafic lavas and ignimbrite phenocrysts suggest that the plagioclase and quartz xenocrysts in the mafic lavas were introduced by mixing of variable composition mafic magmas with the rhyodacitic like those which produced the ignimbrites magmas (Kay et al., 1994a).
Figure 2.3 from Coira and Kay (1993) shows a model to account for these diverse volcanic rocks. The model involves (1) melting in the asthenospheric mantle above the subducting slab; (2) variable contamination of the melts by delaminated slab-modified lithosphere in the asthenosphere and by in situ continental lithosphere and crust; and (3) magma accumulation, fractionation, and mixing of diverse mafic magmas with the rhyodacite that produced the ignimbrite at décollement depths near 20 km before eruption. The most shoshonitic magmas require the lowest percentage mantle melts and the most slab-modified lithospheric contamination.
The characteristics of the Tuzgle magmas are consistent with the underlying mantle being in a transitional setting from a thick continental lithosphere in the incipient stages of delamination in the Altiplano to the north to a thinner post-delamination mantle wedge producing calc-alkaline and intraplate magmas in the southern Puna (Kay et al., 1994a, 1999). Hoke et al. (1994) and Myers et al. (1998) argued for incipient delamination under the Altiplano based on mantle 3He/4He anomalies and seismic data, respectively. In the Tuzgle region, Schurr et al. (2003) use a tomographic model (Fig. 2.3) to argue for melting or fluids ascending from an earthquake cluster at 200 km depth into the crust below Tuzgle. Schurr et al. (2006) interpret a more recent tomographic image below Tuzgle as showing blocks of continental lithosphere in the process of being delaminated, melting in the lower crust, and melt accumulation at ~20 km in accord with the model of Coira and Kay (1993) in Figure 2.3.
Stop 2-3: Exposure of the Tuzgle Ignimbrite
Directions: Continue south on secondary dirt road. Stop is 2 km south of the village of Sey at 23°57′28″S; 66°29′35″W; 4040 m asl.
The Tuzgle ignimbrite is well seen on the west side of the road at this stop. This ignimbrite, which was the first magma to erupt from the Tuzgle center, flowed to the north over Miocene sedimentary and volcanic rocks and the Ordovician volcanic and sedimentary sequences of the Faja Eruptiva Oriental Formation. The flow can be seen to be moderately crystal rich and to have a low degree of welding. The basal part of the ignimbrite contains sparse lithic fragments that are mainly derived from the underlying Ordovician basement. Pumice fragments constitute 10%–20% of the middle to upper sections of the ignimbrite.
Stop 2-4: Exposure of the Tuzgle Pre-Platform Andesite Flows
Directions: Continue south on dirt road; stop at 24°05′13″S; 66°30′47″W; 4461 m asl.
The blocky lava of the Tuzgle pre-platform andesite unit is well seen at this stop. The phenocrysts in the lava include clinopyroxene, orthopyroxene, and amphibole. Plagioclase phenocrysts are absent, and large plagioclase (andesine) and quartz xenocrysts are abundant. The disequilibrium nature of the xenocrysts is shown by sieve textures and labradorite rims on the andesine xenocrysts, and by embayed brown glass rims and clinopyroxene aggregates around quartz xenocrysts.
Stop 2-5: Tuzgle Volcano and Middle Miocene Ignimbrite Flows
Directions: Continue south on dirt road; stop at 24°07′01″S; 66°27′45″W; 4355 m asl.
View to the north of Tuzgle volcano in the Pastos Chicos– Aguas Calientes depression. The depression is bounded by Ordovician magmatic and sedimentary sequences on the east and by the Miocene clastic-pyroclastic facies of the Pastos Chicos Formation and the deformed sediments of the underlying Trinchera Formation on the west. The Tuzgle Young andesite lava flow is well seen from here (see also satellite image in Fig. 2.2). The flow has a blocky to aa structure, pressure ridges developed transverse to the flow direction, and a craggy front.
The outcrop to the west is a middle Miocene dacitic ignimbrite flow with a K/Ar age of 15.2 ± 0.5 Ma (Aquater, 1980) that erupted over Ordovician Faja Eruptiva granodioritic porphyries (Fig. 2.4). This moderately welded ignimbrite is crystal rich and has a moderate lithic content. The Miocene volcanic rocks of this age that extend as far east as 65°50′W at this latitude (Hongn and Seggiaro, 2001) are argued by Coira et al. (1993) and Kay et al. (1999) to record the initial steepening of the Miocene flat slab in the northern Puna.
Stop 2-6: Abra del Charcos and Structure of the Eastern Cordillera
Directions: Continue south; stop at 24°09′01″S; 66°24′44″W; 4350 m asl.
Looking to the southeast from this stop, low-grade metasediments of the Precambrian–lower Cambrian Puncoviscana Formation and quartzites of the middle-upper Cambrian Meson Group can be seen thrust over synrift facies of the Cretaceous Pirgua Subgroup. Just to the west of this fault, Pirgua Subgroup sandstones and conglomerates can be seen in high-angle fault contact with Ordovician Faja Eruptiva Oriental porphyries. To the east of the fault, Ordovician sediments are thrust over Precambrian and Cambrian sequences in a younger over older westvergent thrust that records the inversion of the normal faults of the Cretaceous rift structures. These relations are shown on the geologic map in Figure 2.4. This sequence of structures records the complex history of Late Ordovician (Ocloyic) west-vergent contractional deformation, Cretaceous-Paleogene rifting, and Miocene-Pliocene contractional deformation in this region.
Outcrops along the road at this point are Ordovician Faja Eruptiva Oriental granodioritic porphyries, which are distinctive for their large feldspar megacrysts.
Stop 2-7: View of the San Jerónimo and Negro de Chorrillos Shoshonites along the Calama-Olacapato-Toro Fault Zone
Directions: Continue south on Route 74. Stop is at 24°12′37″S; 66°24′24″W; 4118 m asl. This is a complex interchange; follow the road to Santa Rosa de los Pastos Grandes (Route 129).
The monogenic San Jerónimo and Negro de Chorrillos shoshonitic volcanoes seen to the south and southeast from this stop erupted along northeast-trending, strike-slip faults (Fig. 2.5). K/Ar ages are 0.78 ± 01 Ma for San Geronimo and 0.2 ± 0.08 Ma for Negro de Chorrillos (Aquater, 1980). The northwest-trending Chorrillos fault cutting the northern part of the Negro Chorrillos center and the Incahule fault (Petrinovic, 1999) south of the San Jerónimo center define the Calama–Olacapato–El Toro fault zone (megatraza del Toro) in this part of the plateau. The latest movement on the Chorrillos fault was characterized by oblique normal–left-lateral motion (Marrett et al., 1994). The shoshonitic centers are composed of block and aa-type lavas, bombs, scoria, and ash-fall deposits. Lavas from San Jerónimo flowed up to 10 km from the vent; those from Negro de Chorrillos flowed some 4 km down the valley. Ash-fall deposits up to 10 cm thick from the San Jerónimo center are observed in Quaternary fluvial terraces. Déruelle (1991) suggested the shoshonites formed from melts of mica-bearing peridotites that were substantially contaminated in the crust. Knox et al. (1989) and Coira and Kay (1993) argued that the shoshonitic magmas formed as small-percentage melts of a mantle wedge contaminated by subduction-modified lithosphere followed by some degree of crustal contamination.
Kinematic studies on the faults in the Quebrada del Toro by Marrett and Strecker (2000) reveal a complex regional structural history. This history involves northwest-southeast contraction on northeast-trending faults that began in the Miocene and ended after 0.98 Ma, and horizontal northeast-southwest contraction on northwest-trending faults that had initiated by 4.17 Ma and is still active today. Both fault regimes were active between 4.17 Ma and 0.98 Ma. The present kinematic regime along the northwest-striking Calama-Olacapato-Toro volcanic belt is controlled by active northeast-striking strike-slip and northwest-trending thrust faults (Fig. 2.5). The two fault sets accommodate the differential shortening between major north-striking thrust fault systems.
Stop 2-8: Tajamar Valley near Eastern Topographic Rim of Aguas Calientes Caldera
Directions: Stop along Route 129 at 24°14′30″S; 66°26′54″W; 4228 m asl.
Three distinctive volcanic units can be seen at this stop. The first is the lower Ordovician Faja Eruptiva granodiorite porphyry with large feldspar megacrysts that is typical of the basement in this region. The second is the “blocky” and intermediate aa-type shoshonitic lava flows from the San Jerónimo center. The third is the highly welded Tajamar ignimbrite (Coira and Paris, 1981), which has a mean K/Ar age of 10.6 Ma, and is the intracaldera facies related to the Aguas Caliente caldera (Petrinovic et al., 1999) to be seen at the next stop.
Stop 2-9: View of Rim of Aguas Calientes Caldera
Directions: Continue on Route 129; stop near km 11 at 24°17′16″S; 66°27′07″W; 4473 m asl.
From this point, the eastern rim of the Aguas Calientes caldera is visible to the east, and the 10.6 Ma Tajamar ignimbrite is seen to overlie the hydrothermally altered Verde and Aguas Calientes ignimbrites on the caldera rim. The older Aguas Calientes ignimbrite has been dated at 17.5 ± 0.5–16.8 ± 0.5 Ma (Petrinovic et al., 1999). The hydrothermal alteration of the older ignimbrites is attributed to the intrusion of subvolcanic dacitic domes. This alteration is responsible for the epithermal Ag-Pb-Zn mineralization exploited in the La Poma mining district near Stop 2-8 (Fig. 2.4).
Petrinovic (1999) argues that the collapse of the Agua Calientes caldera was controlled by left-lateral strike-slip faults related to the northwest-southeast–striking Calama-Olacapato-Toro fault system. If so, this fault system must have been active at 11–10 Ma. The initial collapse was asymmetric and produced the distinctive rim that is seen only on the eastern side of the caldera. The continuous collapse and opening of rim vents facilitated the eruption of the dacitic Chorrillos, Tajamar, and Abra de Gallo ignimbrites at 10.8 Ma to 10 Ma (Petrinovic et al., 1999). This caldera collapse was followed by a period of resurgence that led to tumescence of the Tajamar ignimbrite intracaldera deposits. Sb-Au epithermal mineralization in the Incachule mining district is associated with the final stages of the volcanic complex.
Stop 2-10: Abra del Gallo Ignimbrite
Directions: Stop along road at 24°20′20″S; 66°29′28″W; 4691 m asl in the Abra del Gallo.
View of the Abra del Gallo ignimbrite, which is the extracaldera ignimbrite facies associated with the last stage of collapse of the Agua Calientes caldera (Petrinovic, 1999). The east-southeast rim of the caldera and related proximal deposits can be seen from this stop. Ash-flow deposits that are up to 170 m thick comprise the ignimbrite plateau to the east and southeast of the rim. In contrast to the Tajamar ignimbrite (intracaldera facies), the Abra del Gallo ignimbrite (extracaldera facies) seen from here is crystal rich, has a low degrees of welding, and lacks vapor phase or hydrothermal alteration.
Stop 2-11: Quevar Stratovolcano and Serranía de Barreal
Directions: Stop is along road at 24°27′59″S; 66°39′25″W; 3941 m asl.
Quevar stratovolcano is a large late Miocene composite center that overlies Ordovician basement. It is the largest volcanic edifice in the west-northwest–trending Calama-Olacapato-El Toro volcanic chain. The center has been dated and studied by Goddard et al. (1999). At this stop, the andesitic lava flows from the Azufre center (K/Ar age of 7.53 ± 0.02 Ma), and the ca. 5.5 Ma Cerro Gordo dacitic dome can be seen. The lavas cover part of the extensive 10.3 ± 0.3 Ma rhyodacitic-dacitic Olacapato and Pastos Grandes ignimbrites related to the Aguas Calientes caldera. These ignimbrite sheets can be up to 400 m thick under the northern and southeast flanks of the Quevar center.
Volcanic activity at Quevar began with the emplacement of the nonexplosive rhyolitic lava domes to the west of Cerro Azufre at 8.6 ± 0.03 Ma. These domes were then covered in large part by major andesitic-dacitic lava flows and minor block and ash deposits that erupted from five vents between 8.3 and 7.53 Ma. The highest part of Cerro Quevar (6180 m) is a dacitic dome dated at 5.58 ± 0.03 Ma. This dome along with the ca. 5.5 Ma Cerro Gordo dacitic dome mark the last magmatism directly related to the Quevar complex. Important hydrothermal alteration and epithermal deposits occur along northeast- and west-northwest–trending faults in the complex. The largest alteration zone is on the western flank of the volcano in the Quebrada Incahuasi, where a Pb-Ag mine and a precious metal epithermal deposit prospect are located.
Looking eastward from this stop, the Tajamar ignimbrite can be seen to cover deformed continental Eocene-Miocene clastic sedimentary sequences. Because the Tajamar ignimbrite is un deformed, the subhorizontal west-northwest–directed shortening and subvertical extension that affected these sediments must have occurred before 10 Ma (Marrett et al., 1994).
Stop 2-12: Pastos Grandes Basin
Directions: Stop along road at 24°41′49″S; 66°41′37″W; 3880 m asl.
The Pastos Grandes basin is located between two major transverse Puna volcanic chains—the Calama–Olacapato–El Toro chain to the north and the Ratones-Archibarca-Galán chain to the south. The excellent exposures of the Paleogene to Neogene sediments and Miocene borate deposits in the Pastos Grandes basin near this stop provide a record of the Cenozoic evolution of the Puna. Two principal sedimentary cycles are recorded (Alonso et al., 1991; Alonso, 1999a). The first cycle coincides with deposition of the Paleogene Gestes Formation (Teg; Fig. 2.4) red beds, which contain a subtropical continental fauna and indicate an exoreic fluvial environment with warm and humid climatic condition. These sediments accumulated in broken foreland basins. The second cycle coincides with the deposition of the Pozuelos and Sijes Formations (Tmp and Tms; Fig. 2.4), which formed in endorheic basins (closed with no outflow to rivers) under semiarid to arid conditions. The basins have evaporitic and alluvial deposits (wet salar) along with significant volcanic and geothermal deposits. These early Miocene to late Pliocene (or early Pleistocene) sediments were deposited in compressional retroarc basins in which volcanic centers created barriers (Jordan and Alonso, 1987; Kraemer et al., 1999; Voss, 2002). The presence of Oligocene?-Miocene evaporites shows that the onset of hyperarid conditions was associated with establishment of internal drainage (Alonso et al., 1991; Vandervoort et al., 1995).
Strata representing these sedimentary cycles are well exposed near Stop 2-12. The Eocene Geste formation (Turner, 1961; Teg in Fig. 2.4) is a more than 2-km-thick sequence of alluvial and fluvial deposits. These deposits can be seen on the eastern side of the Salar Pozuelos to the west of this stop where they are in angular unconformity with lower Paleozoic metamorphic basement. The exposure is an eastward-dipping homoclinal section of mammal fossil-bearing red and purple conglomerates, sandstones, siltstones, and mudstones. The Geste sequence gives way to the east to the strongly folded clastic and evaporitic (halite, gypsum, and borate) sedimentary beds of the Pozuelos Formation (Turner, 1961; Tmp in Fig. 2.4). These strata are well exposed in the homoclinal structure along the southeastern margin of the Salar de Pastos Grandes. They have yielded a K/Ar age of 7.6 ± 1.1 Ma (Alonso et al., 1991). Above and to the east of the Pozuelos Formation is a thick sequence of siltstones, clay stones, and tuffs in the Sijes Formation (Turner, 1961; Tms in Fig. 2.4). The tuffs have K/Ar ages of 6.8–4.0 Ma (Alonso et al., 1991). North- and north-northeast–trending folds in the Sijes Formation show that active shortening was still occurring at 4 Ma. The Sijes formation hosts an approximately north-northeast–oriented belt of discontinuous borate deposits that are mined in hydroboracite deposits in the Monte Amarillo and Monte Azul mines and in colemanite deposits at Esperanza, Sol de Mañana, and Santa Rosa (see Fig. 2.8). Concordantly above the Sijes Formation is a conglomeratic and psammitic sequence known as the Siguel Formation (Alonso and Gutierrez, 1986; Tps on Fig. 2.4). An intercalated tuff in this sequence yielded a 40Ar/39Ar age of 2.9 ± 0.04 Ma (Vandervoort et al., 1995). Deformation in the Siguel Formation records a period of tectonic reactivation.
Terraces made of clastic and evaporitic deposits interbedded with pyroclastic rocks record the presence of a Pleistocene salar within and on the southern margin of the Salar de Pastos Grandes (Blanca Lila Formation; Alonso and Menegatti, 1990). The terraces have been dated at 1.5 Ma (see Alonso et al., 2006).
Stop 2-13: The Salar de Ratones Depression and the Ratones Volcano
Directions: 25°02′42″S; 66°47′5″W; 3870 m asl.
Cerro Ratones, which is prominently seen to the south from this stop, is a partially eroded andesitic stratovolcano. This volcanic center is part of the transversal Puna volcanic chain that includes the Archibarca volcano and the Cerro Galán ignimbrite complex (e.g., Alonso et al., 1984b). A K/Ar age of 30 ± 3 Ma (Linares and González, 1990) on an andesitic flow on Cerro Ratones suggests the center is one of the oldest Cenozoic volcanoes in the region. The 40Ar/39Ar age of 7 Ma age in Vandervoort et al. (1995) on a younger flow indicates a later stage of activity.
Basement orthogneises and subordinate amphibolites and migmatites on the eastern slope of the Salar de Ratones and in the Salar de Diablillos have yielded a Sm-Nd mineral isochron age of 509 Ma (Lucassen et al., 2000) and a U-Pb age of 508 Ma (Becchio et al., 1999). A Silurian thermal overprint near 440 Ma has been reported by Lucassen et al., (2000).
To the west, the metamorphic basement along the western margin of the Salar de Ratones depression consists of quartzaluminosilicate–bearing schist and granitic orthogneiss with subordinate pegmatite, amphibolite, and migmatite (Viramonte et al., 1976). An Ordovician thrust puts these high-grade units over Ordovician sedimentary rocks (Hongn, 1994; Mon and Hongn, 1991). These Paleozoic rocks are unconformably overlain by the Eocene Geste Formation, which is in turn under a tilted and deformed middle Miocene clastic and pyroclastic sequence with north-northeast–trending fold axes. The Miocene deformation is consistent with subhorizontal WNW-ESE shortening and sub-vertical extension (Marrett et al., 1994).
Stop 2-14: Salar del Hombre Muerto Depression (View to the South)
Directions: Stop is along road at 25°11′17″S; 66°58′04″W; 4041 m asl.
The Salar del Hombre Muerto contains borate deposits interbedded in its sedimentary fill. Borate-bearing salars like these are principally along the eastern border of the southern Puna and in the northern Puna. Ulexite, which is the dominant boron-bearing mineral, occurs in “potatoes” and “bars.” In some salars, borax (or “tincal”) occurs as disseminated (Diablillos and Centenario) or small (Turi Lari, Cauchari, and Rincon) deposits. The formation of Tertiary and Quaternary borate deposits in the Puna can be attributed to three factors (Alonso, 1999a, 1999b): (1) active hot springs with borate-bearing water, (2) closed basins, and (3) an arid to semiarid climate.
Borate deposits in the Salar del Hombre Muerto are concentrated on the island in the salar called the Farallon Catal, and on the Tincalayu and Hombre Muerto peninsulas. The Farallon Catal has a 5000-m-thick sedimentary sequence that is mapped in the Catal Formation (Tmc on map in Fig. 2.4). This formation is composed of red sandstones and pelites at the base, green pelites with gypsum and travertine in the middle, and ignimbrites and reworked tuffs and coarse conglomerates at the top. K/Ar dates limit the age at the base to be 15.0 ± 2.4 Ma and that at the top to be 7.2 ± 1.4 Ma (Alonso et al., 1991). Evaporites near the middle of the section are pure gypsum interbedded with gypsiferous clay (Alonso and Gutierrez, 1986). A basaltic andesite flow overlying the sequence has a 40Ar/39Ar age of 0.8 ± 0.1 Ma (Risse et al., 2008). The folded sedimentary sequence on the Tincalayu penin sula in the northwestern part of the salar is mapped in the Sijes Formation (Tms) in Figure 2.4. This sequence has halite near the base, borax and gypsum in the middle, and clastic strata at the top. A tuff in the borax deposits has an age of 5.86 ± 0.14 Ma. Basaltic andesitic lavas overlying these strata have ages of 2.43 ± 0.7 Ma (Risse et al., 2008) and 0.75 ± 0.3 Ma (Alonso et al., 1984b). The Tincalayu mine contains the largest borax deposit in the Southern Hemisphere.
Stop 2-15: Southern Puna Mafic Andesitic Flow Southwest of the Salar de Hombre Muerto
Directions: Stop is along side of dirt road that is route 43 near 25°27′30″S; 67°11′41″W; 4115 m asl.
The southern Puna is characterized by late Neogene strikeslip and extensional faults, small discontinuous diachronous basins, and mafic, glassy lava flows. Basaltic andesitic to andesitic flows around the Salar de Hombre Muerto like the one at this stop erupted from monogenetic to polygenetic centers along north-northeast–, north- and northeast-oriented faults (Figs. 2.6 and 3.1). They have ages of near 4.6, 2.4, and 0.8 Ma (Risse et al., 2008). The mafic andesite lava (~54% SiO2) at this stop is typical in having high MgO, Cr (300 ppm), and Ni (100 ppm) contents and containing acidic plagioclase and quartz xenocrysts. These features fit with the rapid rise of mantle-derived basaltic magmas (~80%) that mixed with lower crustal melts (~20%) accumulated at mid-crustal depths (Kay et al., 1999, 2005). The quartz and feldspar xenocrysts can be explained as the phenocrysts in the silicic magmas. Mid-crustal mixing of this type is supported by seismic anomalies interpreted to be magma accumulations at mid-crustal depths (e.g., Chmielowski et al., 1999; ANCORP, 2003; Fig. O.5).
Days 3 to 5—Southern Puna
The objective of the next three days is to examine characteristic features of the southern Puna and the differences with the northern Puna. On the afternoon of Day 5, the route descends the southeastern margin of the plateau and crosses into the northern Sierras Pampeanas. A four-wheel drive vehicle is required for many of the stops on Days 3–5.
Day 3—Oligocene to Quaternary Evolution of The Southern Puna
The objective of Day 3 is to examine the characteristic structural and magmatic features of the southern Puna and the differences in the Miocene to Recent history with the northern Puna (see Overview). A main structural difference is the normal and strike-slip faults along which mafic magmas erupted. Another difference is that the distribution of volcanic centers con tributed to a distinctive sedimentation pattern in a broken foreland basin. Magmatic differences with the northern Puna include (Figs. O.6–O.8): (1) lack of a Neogene magmatic gap, (2) large 15–8 Ma andesitic to dacitic backarc stratovolcanoes and dome complexes, (3) post–7 Ma mafic lavas and glassy andesites and dacites along faults, (4) small- to moderate-sized ca. 6.6–4 Ma ignimbrites followed by the 1000 km3 Cerro Galán eruption at ca. 2.2 Ma, and (5) young mafic lavas with intraplate chemical affinities. These differences are interpreted to reflect moderate shallowing of the subducting slab as the Juan Fernandez ridge subducted beneath the region at ca. 14–7 Ma, steepening of the slab after 7 Ma, and delamination of eclogitic lower crust and mantle lithosphere. Limited crustal shortening east of the plateau is consistent with crustal flow contributing to crustal thickening of the southern Puna during the ca. 10–4 Ma peak of ignimbrite eruption in the northern Puna and shallowing of the Chilean flat slab to the south.
Stop 3-1: Outflow Sheet of the Cerro Galán Ignimbrite ~30 km West of the Caldera Rim
Directions: Take Route 43 (gravel road) north from Antofagasta de la Sierra toward Salar de Hombre Muerto. Stop is along Route 43 at 25°54′58″S; 67°21′49″W; 3681 m asl.
The outflow facies of the large-volume dacitic Cerro Galán ignimbrite that erupted at ca. 2 Ma are well seen at this stop. The flow at this locality is moderately welded and exhibits spectacular columnar jointing that reflects the geometry of the paleophreatic surface. The ignimbrite overlies a strongly folded Ordovician sedimentary sequence.
Stop 3-2: View of Cerro Galán Caldera from the Northwest
Directions: Stop is along Route 43 at 25°46′38″S; 67°16′22″W; 4310 m asl.
This stop affords a panoramic view of the Cerro Galán ignimbrite sheet that erupted in a major explosive episode at ca. 2 Ma from the giant Cerro Galán caldera (35 × 20 km in diameter) to the east. The ignimbrite has an estimated volume of 1000 km3 (Sparks et al., 1985). The volume of the outflow facies is estimated at 280 km3. The outflow sheet is from 30 to 200 m thick and extends up to 100 km from the caldera rim in all directions. The distal flows visible at this stop cover folded Ordovician sedimentary rocks. Looking to the south, early Pliocene dacitic to rhyodacitic lavas from volcanic centers like Merihuaca (4.86 Ma) can be seen to be partially covered by the Cerro Galán flows.
To the east-southeast, the dome structure of the resurgent center in the caldera is seen to rise to an altitude near 6300 m. The Cumbres de Luracatao peaks seen to the north (see Thematic Mapper image in Fig. 2.6) expose the lower Paleozoic orthogneisses and granitoids basement that is under the Cerro Galán caldera.
Stops 3-3a and 3-3b: Views of Beltran Stratovolcanic Complex
Directions: Stops are along Route 43 at (a) 25°38′56″S; 67°13′58″W; 4351 m asl; and (b) 25°34′41″S; 67°13′50″W; 4351 m asl.
The Beltran stratovolcano seen from Stop 3-3 is one of the large, long-lived composite volcanoes in the southern Puna. The center is dominantly composed of voluminous dacitic to andesitic lava flows and was active from 14.1 to 7.7 Ma (Kraemer et al., 1999). This center, like the Tebenquicho and Antofalla stratovolcanic complexes in the distance to the northwest and west on the margin of the Salar de Antofalla margin (Figs. 3.1 to 3.3), are part of the northwest-trending transcurrent Archibarca–Galán Puna volcanic chain.
Stop 3-4: Laguna Caro Depression and Glassy Late Miocene Andesites
Directions: Turn west from Route 43 onto a dirt track heading toward Laguna Caro. Stop on the south side of Laguna Cara at 25°34′42″S; 67°17′36″W; 4041 m asl.
The Beltran seen at the last stop flanks the southern side of the Laguna Caro depression. The eroded glassy andesitic flows visible at this stop and seen elsewhere around the depression are a distinctive volcanic feature of the southern Puna. These flows are characterized by a glassy character, plagioclase (oligoclase) and quartz xenocrysts, and high Mg numbers and Cr (~100–200 ppm) and Ni (~25–100 ppm) contents for their SiO2 (59%–64%) contents. Moderately steep REE patterns (La/Yb >25) are consistent with the magmas equilibrating with residual garnet-bearing lower crust. The glassy character suggests that the flows erupted as hot, degassed magmas that ascended rapidly along faults. A glassy andesite on the northeastern margin of the Laguna Caro depression has a whole-rock 40Ar/39Ar age of 4.6 ± 0.5 Ma (Risse et al., 2008).
Stop 3-5: View of Salar de Antofalla, Antofalla, and Tebenquicho Stratovolcanoes
Directions: Continue west on dirt track to 25°34′58″S; 67°30′54″W; 3928 m asl. Stop at top of ridge where road descends into the Salar de Antofalla. A map of the area is in Figure 3.4.
The Antofalla stratovolcano on the western margin of the Salar de Antofalla to the west is a voluminous and long-lived 6400 m volcanic complex with a diameter of 35 km. The center erupted in three main stages. The first stage produced: (1) basaltic andesitic lava cones like Cerro de la Aguada (12.9 ± 0.8 Ma; Kraemer et al., 1999), Cajero (13.0 ± 0.5 Ma; Schnurr, 2001) and Antofalla (13 Ma; Coira and Pezzutti, 1976); (2) dacitic lava domes like Cerro Lila (10–9 Ma; Coira and Pezzutti, 1976); (3) dacitic to rhyolitic ignimbrites (10.9–9.6 Ma; Kraemer et al., 1999); (4) andesitic to dacitic lavas (9–8 Ma; Coira and Pezzutti, 1976); and (5) minor basaltic andesitic flows erupted along WNW-ESE and north-south structures. In the next stage between 7 and 4 Ma, basaltic andesites and minor basalt flows erupted on the margins of the complex and from flank vents on the older centers. Basaltic, basaltic andesite, and andesitic flows also erupted from structurally controlled vents on the eastern and southern margins of the Salar de Antofalla. The last stage produced sparse basaltic andesitic lavas and small rhyolitic lava domes like those at Las Cuevas and Cerro Botijuela (2.5–2.3 Ma; Schnurr, 2001; Richards et al., 2006). Pleistocene (1.5–1.2 and 0.1–0.5) andesitic to basaltic scoria cones and lava flows (Marrett et al., 1994; Kay et al., 1997) also erupted on the margins of the Salar de Antofalla. The Tebenquicho stratovolcanic complex to the north is dominated by andesitic to dacitic lava flows that erupted between 14 and 6 Ma (Kraemer et al. 1999).
The Antofalla and Tebenquicho volcanic complexes unconformably overlie thick sequences of Cenozoic sediments that record the tectonic-sedimentary evolution of the region. Figure 3.5 from Carrapa et al. (2005) provides a summary in the Salar de Antofalla region. Sedimentation began in the latest Eocene with deposition of the clastic foreland basin deposits in the late Eocene to Oligocene Quiñoas Formation that were derived from the west as a result of the Incaic deformation (Jordan and Alonso, 1987; Kraemer et al., 1999; Voss, 2002). An arid climate was established at this time. Late Oligocene thick-skinned contractional deformation (D1 in Figure 3.5; Adelmann and Görler, 1998) triggered syntectonic deposition of the coarse-grained late Oligocene to early Miocene Chacras Formation alluvial fan deposits (Kraemer et al., 1999). These fan sediments were largely derived from the Sierra de Calalaste region farther south (Carrapa et al., 2005), which will be visited on Day 4. Uplift associated with Oligocene deformation led to reorganization of the depositional systems in the Salar de Antofalla area. Renewed east-west to WNW-ESE shortening in the early Miocene (D2 in Fig. 3.5, ca. 20–17 Ma) reactivated the Paleogene west-vergent fault system (Adelmann and Görler, 1998). During this time, the Salar de Antofalla region was separated into small intra-arc depocenters in which the Potrero Grande Formation alluvial fan and fluvial sediments accumulated (Kraemer et al., 1999; Voss, 2002). The Miocene west-vergent thrusts affected Lower Paleozoic, Permian, and Tertiary rocks. Younger Miocene shortening (D3 in Fig. 3.5) produced east- and west-vergent basement thrusts that tilted the Potrero Grande Formation alluvial fans and generated further deposition. The middle Miocene to Pliocene Juncalito Formation of Kraemer et al. (1999) was deposited at this time. The thick evaporate deposits in these sequences show that the Salar de Antofalla basin was internally drained by the late Miocene. A mixed Pliocene stress regime (D4 in Fig. 3.5) produced both contractional and local strike-slip deformation. The present narrow and elongate shape of the Salar de Antofalla basin reflects this deformation as well as Quaternary erosional processes. Unlike other southern Puna salars, evaporates in the Salar de Antofalla are largely halite rather than borate deposits.
Seismological studies near 25.5°S by the ANCORP group provide information on the crust beneath the Salar de Antofalla (Heit, 2005). A high-velocity anomaly coincides with a portion of the central Andean gravity high that Götze and Krause (2002) associated with Lower Paleozoic ultrabasic rocks in the basement. The sides of the Antofalla depression are flanked by low-velocity anomalies. The anomaly to the west extends under the main Central Volcanic Zone arc. The anomaly to the east at ~67°W is along strike with the Cerro Galán caldera. The shape and extent of the narrow northeast-trending Salar de Antofalla and the topographic difference between the salar surface (3400 m) and the flanks (~4000 m) appears to reflect a deeper structure bounded by these velocity anomalies. In November of 1973, a magnitude 5.8 earthquake occurred at a depth of 6 km beneath the region (Chinn and Isacks, 1983). The kinematics of the focal mechanism indicate north-south–dipping, 30° extension, and east-west horizontal shortening. One of the nodal planes strikes parallel to the major axis of the Salar de Antofalla.
Stop 3-6: Faults and Mafic Lavas in the Vega de los Colorados
Examples of the strike-slip faults in the southern Puna can be observed in the Vega de los Colorados. These faults have been studied by Marrett et al. (1994) and Kraemer et al. (1999). The main branch of the regional-scale Acazoque fault and a branching synthetic strike-slip fault are well seen. The Acazoque fault strikes NNE-SSW and dips southeast. The fault is marked over much of its length by a conspicuous scarp along which basaltic andesite flows have erupted. Right lateral motion has produced extensional pull-apart deformation.
Outcrops in the Vega de los Colorados region include Ordovician low-grade metamorphic sedimentary rocks that are unconformably overlain by clastic sediments and mafic lava flows. The age of the sedimentary sequence is constrained by a 40Ar/39Ar age of 26.3 ± 1.6 Ma on an interbedded tuff (Vandervoort et al., 1995). The sequence is affected by folds with NNE-SSW–trending axes that are consistent with subhorizontal WNW-ESE shortening and subvertical extension. One of the basaltic andesitic lava flows has a 40Ar/39Ar plateau age of 4.6 ± 0.2 Ma (Risse et al., 2008; isochron age is 4.9 ± 0.2 Ma). The mafic flows and scoria cone in the Vegas de los Colorados are cut by a fault that strikes 342° and dips 45° northeast. This fault is interpreted as a Tertiary reverse fault reactivated as a normal fault (Marrett et al., 1994).
Stop 3-7: Nacimientos—Basaltic Andesites and Their Tectonic Control
Directions: Take dirt road from Vega de los Colorados south in the direction of Nacimientos. Stop is at 25°52′13.9″; 67°26′9.3″; ~3750 m asl.
Basaltic andesite lavas with olivine and clinopyroxene pheno crysts at this stop are typical of flows erupted from monogenetic and simple polygenetic cones in the Nacimientos volcanic field of the southern Puna (Figs. 3.1 to 3.3). This flow yielded a 40Ar/39Ar age of 2.8 ± 0.2 Ma (Risse et al., 2008). The source of the flow is a small monogenic center located on a north-northeast–striking fault system. The fault shows right-lateral movement, which produced an extensional pull-apart setting that controlled the location of the volcanic vents.
Stop 3-8: View of Basaltic Andesites and Andesites Centers
Directions: Continue on dirt road. Stop is at 25°53′03″S; 67°27′045″W; 3735 m asl.
This stop provides a general view of the numerous small basaltic to mafic andesitic centers west of the Cerro Galán caldera. These centers were emplaced along north-south– to north-northeast–striking fault systems. The mafic lavas show a temporal change in chemical character from more arc-like (e.g., La/Ta ratios = 36–55) at ca. 6.6 Ma to more intraplate-like (e.g., La/Ta = 20–35) affinities after 3 Ma (Kay et al., 1994a, 1999). This change is interpreted as being related to the loss of a backarc-subducted component in the thicker mantle wedge that resulted from steepening of the subduction zone and delamination of the crust and mantle.
Day 4—Southern Puna Silicic Calderas
The main objective of Day 4 is to observe the Cerro Galán and Cerro Blanco silicic caldera complexes and their regional settings. The centers are shown in the map in Figure 4.1.
Stop 4-1: Early Pliocene Cerro Merihuaca Lava Flows, ca. 2 Ma Galán Ignimbrite, and a View of Mafic Cinder Cones and Lava Flows to the West
Directions: Go north on Route 43 from Antofagasta de la Sierra. Turn east at the sign for the road to Real Grande, ~6 km north of the village of Antofagasta de la Sierra; turn is near 26°02′30S; 67°23′041″W; 3441 m asl. Stop is at 25°59′07″S; 67°17′27″W; 3984 m asl.
Dacitic lava flows from the Merihuaca volcanic center (4.86 ± 0.19 Ma; Sparks et al., 1985) are well seen at this stop (Figs. 4.2 and 4.3). They are surrounded by the Cerro Galán ignimbrite. The surface of the ignimbrite viewed from this stop exhibits a distinctive wavy form that reflects its eruption on a smooth slope dropping into the Punilla River depression to the west. Numerous late Miocene to Pleistocene basaltic andesitic, basaltic trachyandesitic, and andesitic monogenic and polygenic centers can be seen to the west (Figs. 4.1 and 4.2).
Stop 4-2: Wavy Upper Surface of the Galán Ignimbrite
Directions: Continue on dirt track; stop is at 25°58′56″S; 67°17′08″W; 4021 m asl.
This stop provides a closeup view of the wavy upper surface of the ca. 2 Ma Cerro Galán ignimbrite. This ignimbrite is the most extensive and youngest flow unit from the Cerro Galán caldera complex with flows reaching distances of up to 100 km from the rim. At this point we are ~18.5 km west of the rim (see Fig. 4.2). The ignimbrite here is dacitic in composition (~68% SiO2), crystal rich (plagioclase, sanidine, quartz, and biotite), has small amounts of pumice (1–5 cm in diameter), and contains sparse lithic fragments.
Stop 4-3: Real Grande Valley: Cerro Galán Ignimbrite Sequence near Camp 1 Locality of Sparks et al. (1985)
Directions: Continue on dirt track. Stop is near end of track at 25°59′23.6″S; 67°13′57.4″W; 4212 m asl. View is that in Fig. 4.4A.
A complete section of the Cerro Galán ignimbritic sequence is exposed in the Real Grande Valley at this stop. A labeled photo looking to the north is shown in Figure 4.4A. The locality is some 12 km west of the caldera rim and is near section CG2 studied by Sparks et al. (1985; Fig. 4.5). The base of the profile is composed of the Blanco ignimbrite, which is overlain by the 6.4−5.1 Ma Merihuaca ignimbrite. The Merihuaca ignimbrite is separated into lower, middle, and upper members at this point. The next unit is the 5.14–4.8 Ma Real Grande ignimbrite. The Merihuaca and Real Grande ignimbrites units are characterized by basal plinian deposits, numerous individual flows, and proximal co-ignimbrite lag deposits. Overall, they are moderate to rich in pumice and have a high lithic and a moderate to low crystal content (plagioclase, quartz, biotite, and absent to minor hornblende). Variable compositions are indicated by banded pumices that occur at the top of some units. A dense rock equivalent (DRE) volume of >500 km3 was estimated for the Merihuaca and Real Grande ignimbrites by Sparks et al. (1985). At the top of the section is the ca. 2 Ma Cerro Galán ignimbrite that is separated from the Real Grande ignimbrite by an important erosional unconformity. As at previous stops, the Galán ignimbrite is a crystal-rich, pumicepoor, and lithic-poor unit. In contrast to the underlying ignimbrites, the Cerro Galán ignimbrite lacks basal plinian deposits, proximal lag breccias, and compositional zoning. Sparks et al. (1985) proposed that the Cerro Galán ignimbrite eruption resulted from a catastrophic foundering of a cauldron block into a magma chamber leading to caldera collapse. Post-caldera eruptions took place along the northern segment of the fault on the eastern side of the caldera and on the western margin. Resurgent doming centered along the eastern fault is shown by radial tilting of the ignimbrite and lake deposit accumulations.
Stop 4-4: Merihuaca Ignimbrite Section in the Real Grande Valley
Directions: From near Stop 4-3, head north-northeast on a minor dirt track and continue cross country toward the section viewed at Stop 4-3. Do not do this without a good four-wheel drive vehicle and an experienced driver. Stop is at 25°58′46.7″S; 67°13′08.6″W; 4305 m asl. After the stop, retrace the trip route to the village of Antofagasta de la Sierra.
This stop provides a closeup view of part of the Cerro Galán ignimbrite sequence in column CG2 in Figure 4.5. The closest outcrop is in the lower member of the Merihuaca member (Ml in Fig. 4.5). This ignimbrite is a rhyodacite (68% SiO2) with 10%–15% crystals (biotite, quartz and plagioclase). The lower layer of the unit is a massive pumice flow with a slight reverse grading in pumice clasts and up to 20% of lithic clasts (lava fragments and metamorphic rocks). The overlying layer shows reverse grading in pumice and internal bedding defined by variations in pumice size. Four types of pumice occur: (1) a moderately compact gray to white crystal-rich type, (2) a white laminar type, (3) a finely vesiculated ochre type, and (4) an almost aphanitic type. The top layer is a massive pumice flow deposit that is moderately welded and has an intercalated, fine ash-fall deposit. A walk to the north provides a view of the upper units.
Salar de Incahuasi and Cerro Blanco Caldera
The second half of Day 4 includes a transit through the Salar de Incahuasi to the Cerro Blanco volcanic complex to look at the structural relations (Stop 4-5) and late Miocene to Pleistocene volcanic rocks of the southernmost Puna (Stops 4-6 to 4-9). The late Miocene Rosada ignimbrite and the Quaternary volcanic rocks associated with the Cerro Blanco caldera will be visited. The region south and west of the Cerro Blanco caldera in the Cor dillera de Buenaventura will be seen on Day 6. The stops are shown on the geologic maps in Figures 3.3 and 4.6 and on the Thematic Mapper image in Figure 4.7. The stops require a four-wheel drive vehicle and can only be reached when conditions are appropriate for driving on the surface of the salar. The road is a track that is poorly marked and hard to follow in some segments.
Stop 4-5: Neo-Paleozoic Structures Reactivated by Tertiary Andean Tectonism
Directions: Head west from the village of Antofagasta de la Sierra on a road to the Salar de Incahuasi. Ask for directions in the village. Stop 4-5 is at the north end of the salar at 26°6′35″S; 67°35′10″W; 3540 m asl.
View of the Sierra Phyllo Colorado, which exposes red Permian conglomerates and sandstones (Patquia and de la Cuesta Formations; Fernández Seveso et al., 1991) and Ordovician turbidites (Fig. 4.6). A NNW-SSE–striking normal fault in the northern part of the exposure puts the Ordovician sequence in fault contact with the Permian sequence. These Permian deposits show an abrupt decrease in thickness west of the Sierra Phyllo Colorado where they are exposed in an upthrown block of an east-dipping normal fault. Both the Ordovician and Permian units are covered in low-angle unconformity by Paleogene (Eocene) Geste Formation reddish to brown sandstones, conglomerates, and pelites. The Geste Formation was seen in the Pastos Grandes Basin on Day 2. The Paleogene and Permian are also related by a reverse fault. This structural picture has been interpreted to result from Andean contractional inversion of extensional structures formed in the Permian basin (Seggiaro et al., 2002).
Stop 4-6: Rosada Ignimbrite
Directions: Stop is at 26°26′39.6″S; 67°41′21.0″W; 3254 m asl.
The Rosada ignimbrite (8.1 ± 0.5 Ma, Kay et al., 2006; 6.3 ± 0.2 Ma, Kraemer et al., 1999) seen at this stop flowed from the south through the Incahuasi salar depression and partially covered Ordovician sedimentary sequences on the eastern flank of the Sierra de Calalaste (Figs. 4.6 and 4.7). Subsequently, the Rosada ignimbrite was eroded and partially covered by the Pleistocene Blanco ignimbrite that flowed through the low part of the center of the depression. The Rosada ignimbrite is best exposed on the border of the Blanco ignimbrite along the margins of the Incahuasi depression. The Rosada ignimbrite is a highly to moderately welded rhyodacitic ash flow (66.8%–68.7% SiO2) that is crystal rich, has moderate to large amounts of pumice (15%–35%), and is poor (<5%) in lithic fragments (dacites and Ordovician rocks). The pumice is white, crystal rich, and variably flattened depending on the degree of welding. Some very flattened pumice fragments are up to 15 cm in length.
Stop 4-7: Blanca (White) Ignimbrite
Directions: Stop is at 26°31′06.7″S; 67°42′15.4″W; 3466 m asl.
The Blanca ignimbrite (White ignimbrite) exposed here has a 40Ar/39Ar age of 0.2 ± 0.1 Ma (Siebel et al., 2001) and is related to the Cerro Blanco caldera complex (Figs. 4.6 and 4.7). The complex consists of two nested calderas embedded in a major structure with a diameter of ~15 km. The main ash flows form two units. One is the Campo de la Piedra Pomez unit that flowed northeastward along the Carachipampa valley, where it has been dated at of 0.55 ± 0.1 Ma (Seggiaro et al., 2002). The other is the Blanca ignimbrite seen here. This ignimbrite flowed to the north for more than 25 km through the Incahuasi depression. The flow has a rhyolitic composition (70%–74.4% SiO2) and is massive, crystal poor (quartz, plagio clase, and biotite), pumice rich, and poorly welded. Different facies of the ignimbrite vary from co-lag proximal to distal pumice-rich deposits.
The margins of the Cerro Blanco caldera complex are formed by distributed co-ignimbrite breccias that contain older ignimbrite and lava blocks. Several intracaldera lava domes record an episode of magmatic resurgence at 0.15 Ma (Siebel et al., 2001) postdating the Blanca ignimbrite. Pritchard and Simons (2002) reported a thermal anomaly (~100 °C) in the central part of the Cerro Blanco caldera, and detected a subsidence of 2.5 cm for a period from 1992 to 2000 using radar interferometry. Their interferogram is shown as an inset in Figure 4.7. Evidence for subsidence has been absent since 2000.
Stop 4-8: Cueros de Purilla Obsidian Lava Dome
Directions: Stop at 26°34′3.6″; 67°44′58.2″; 3895 m asl.
The Cueros de Purulla is a rhyolitic (72.3% SiO2) obsidian lava dome (Fig. 4.7) with a 40Ar/39Ar date of 0.4 ± 0.1 Ma (Siebel et al., 2001). The location of the dome is controlled by extensional and strike-slip faults. Dome formation can be attributed to the release of batches of degasified magmas at the time of the eruption of the Blanca ignimbrite and the formation of the Cerro Blanco caldera.
Stop 4-9: Domes at the Margin of the Cerro Blanco Caldera
Directions: Stop at 26°40′37.8″; 67°45′6.0″; 4020 m asl.
Dacitic lava domes at this stop on the northern margin of the Cerro Blanco caldera complex erupted in episodes of degasification of the caldera system. They are phenocryst-poor vitrophyric lavas (plagioclase, biotite, and quartz). One has a K/Ar age of 1.3 ± 0.4 Ma (Kay et al., 2006) and is considered to constitute the easternmost dome in the east-west–trending Cordon de Buenaventura dome complex.
Day 5—Southern Puna Young Mafic Lavas and Faults and The Transition Into The Northern Sierras Pampeanas
Stop 5-1: Quaternary Basaltic Andesite Flows of the Laguna Volcano
Directions: Head south from Antofagasta de la Sierra on Route 43 (being improved and paved in late 2007). Stop at 26°07′55″S; 67°25′26″W; 3368 m asl.
The Laguna volcano is a prominent monogenetic cinder cone (Figs. 4.7 and 5.1) with a lava flow dated by 40Ar/39Ar at 0.34 ± 0.06 Ma (Risse et al., 2008). The flows have an aa morphology with distinctive jagged flow fronts. In some places, layers of fragmented clinkers alternate with massive lavas. The lavas are characterized by olivine and plagioclase phenocrysts and a basaltic andesitic composition (53% SiO2). Their trace elements show an intraplate-like signature (La/Ta ~25; Ba/Ta ~386) and a moderately steep REE pattern (La/Yb = 17) that is interpreted to reflect an increase in mantle wedge volume above the subducting Nazca plate in response to lithospheric delamination and steepening of the slab (Kay et al., 2005; Overview).
Stop 5-2: Late Miocene Glassy Lavas under the Laguna Flows
Directions: Stop along Route 43 at 26°08′23″S; 67°25′14″W; 3343 m asl.
The glassy andesitic lava flow (~62% SiO2) at this stop, which has a 40Ar/39Ar age of 7.34 ± 0.04 Ma (Risse et al., 2008) is partially covered by the Quaternary Laguna lava flow. The flow shows the characteristic flaggy parting typical of many latest Miocene and Pliocene southern Puna lavas The trace element chemistry contrasts sharply with that of the Laguna flow in having an arc-like signature (strong HFSE element depletion; La/Ta = 46.9). This signature along with a steep REE pattern (La/Yb = 28) showing equilibration with garnet typifies many southern Puna lavas of this age. These characteristics are interpreted to reflect a thick underlying crust and the introduction of arc components into the magma source in association with late Miocene shallowing of the subducting slab under the southern Puna region (Kay et al., 2005).
Stop 5-3: View of the Mafic Jote and Alumbrera Volcanoes
Directions: Stop along Route 43 at 26°12′0″S; 67°24′14″W; 3300 m asl.
The Jote volcanoes are a cluster of monogenic basaltic andesitic centers composed of cinder cones and lava flows. Their distribution is well seen in the satellite image in Figure 4.7. The centers are distributed along a north-northeast–striking fault system that has a component of horizontal extension. An olivine phyric basalt (51.5% SiO2) from one of these flows yielded a 40Ar/39Ar age of 3.2 ± 0.02 Ma (Risse et al., 2008). The flows are chemically similar to the Laguna flows (Kay et al., 1994a) indicating that generally similar magmatic conditions have persisted in the region for the past 3 million years.
Stop 5-4: View of Carachipampa Volcano, Campo de Piedra Pumice, and Cerro Blanco Caldera
Directions: Stop is along Route 43 at 26°28′49″S; 67°20′15″W; 3160 m asl.
The prominent Carachipampa monogenetic mafic volcanic center is visible to the west from this stop. The Campo de Piedra Pomez, which is the pumice-rich flow erupted from the Cerro Blanco caldera at 0.55 Ma (Seggiaro and Hongn, 1999), is visible to the southwest (Figs. 4.7 and 5.1). The Carachipampa volcanic center is composed of mafic scoria cones and radiating lava flows. The lavas are clinopyroxene-bearing basalts (52% SiO2) with an intraplate-like chemistry (Kay et al., 1999, 2006) that yielded a 40Ar/39Ar age of 0.75 ± 0.08 Ma (Risse et al., 2008).
Stop 5-5: Pasto Ventura Monogenic Volcanoes and Quaternary Faults
Directions: Entrance to Pasto Ventura region is along old Route 43 at 26°41′56″S; 67°10′40″W; 3792 m asl. A four-wheel drive vehicle is required. Follow track to south and then turn at 26°44′03″S; 67°11′03″W; 3623 m and drive up gentle slope to west. Stop is at 26°44′24″S; 67°13′08″W; 3675 m asl.
As shown in Figures 5.2 and 5.3, this is another region where young southern Puna mafic lava flows are clearly related to normal and strike-slip faults. The locality has been studied by Allmendinger et al. (1989) and Marrett et al. (1994). In the region, Paleozoic metamorphic basement rocks can be seen to be thrust over tightly folded continental clastic sequences that are overturned beneath west-northwest to northwest-dipping and north-northeast to northeast- striking thrust faults. These thrust faults are considered to be Miocene in age. Along strike, these faults and possible new faults offset young basaltic andesitic cinder cones and lavas and alluvial surfaces like those seen at this stop. Kinematic indicators along the reactivated fault provide evidence for oblique normal-right lateral motion with moderately northeast and east-northeast–plunging shortening and subhorizontal north-northwest to north-south extension (Allmendinger et al., 1989). A small eroded basaltic andesitic cone displaced by a normal fault to the north of this stop has a 40Ar/39Ar age of 1.3 ± 0.6 Ma (Fig. 5.2; Marrett et al., 1994).
The young basaltic to mafic andesitic lavas (52%–56% SiO2) in this region contrast with the young mafic lavas erupted at the Laguna, Jote, and Carachipampa centers in having arclike high field strength element depletions (La/Ta = 32–55; Ba/Ta = 430–670; Kay et al., 1994a).
Stop 5-6: Late Miocene Stratigraphy and Mio-Pliocene Deformation near Pasto Ventura
Directions: Continue south along rough track to 26°45′23″S; 67°13′29″W; 3684 m asl. After stop, retrace route back to main road (Route 43).
The late Miocene sequence at this point consists of a red bed sequence containing an ash-fall tuff dated by 40Ar/39Ar at 10.04 ± 0.05 Ma (Marrett et al., 1994). The sequence is affected by overturned folds that are cut by west-northwest to northwest-dipping thrust faults. The geometry of the structures indicates subhorizontal west-northwest to northwest-trending shortening and subvertical extension during the late Miocene to Pliocene (Marrett et al., 1994).
Stop 5-7: View to South of the Vicuña Pampa Volcanic Complex
Directions: Stop along Route 43 near 26°44′29.2″S; 67°04′15.7″W; 3897 m asl.
The Vicuña Pampa Volcanic Complex seen to the south-southeast is a succession of ignimbrites, lavas, and minor dikes of basaltic to basaltic andesitic composition associated with rare plinian deposits. The complex is constructed on Precambrian–Lower Paleozoic metamorphic basement. Viramonte and Petrinovic (1999) proposed that this volcanic complex developed as a collapse caldera with a circular geometry and a diameter of 15 km. The elements of the caldera are shown on the Thematic Mapper image in Figure 5.2. An andesitic lava flow associated with the center at Peñas Frias has a K/Ar age of 12.1 ± 0.6 Ma (Kay and Coira, unpublished). A N40°W-striking fault system cuts the complex. A fault displacing the rim is reported to have a left-lateral strike-slip sense with ~3 km of displacement (Allmendinger et al., 1989).
Stop 5-8: Rio El Bolsón—Phases and Styles of Late Cenozoic Deformation
Directions: Stop along Route 43 at ~27°01′28″S; 66°45′43″W; 2400 m asl.
Two distinct phases and styles of late Cenozoic deformation have been described in the northern Sierras Pampeanas near the boundary with the southern Puna by Allmendinger (1986, 1989). The older phase is characterized by contraction exhibited by folding and dip-slip thrust faults. Northeast-trending thrust faults that dip 20°–55° to the northwest bound structural blocks and put basement rocks over folded Tertiary sequences. Northwest horizontal shortening and vertical extension directions are prevalent (Allmendinger, 1986). En echelon faults terminate in plunging folds like those that are typical in transfer zones of foreland thin-skinned thrust belts. The fault segments in the Sierra Pampeanas are shorter than those fault segments in the Puna (average 20–30 km in length). The timing of contractional deformation is constrained by ages in the Corral Quemado sequence (Fig. 5.4) that indicate that contractional deformation began after ca. 11 Ma and continued until ca. 3 Ma (Allmendinger et al., 1989). The second and younger style of late Cenozoic deformation in the northern Sierras Pampeanas is latest Pliocene-Quaternary in age and is associated with a thrust kinematic regime typified by nearly east-west shortening and verti cal extension. This geometry is seen along the Rio El Bolsón where the Punaschotter conglomerates (Fig. 5.4) unconformably overlie folded sequences in the Andalhuala Formation that are cut by northeast-striking thrust faults. Allmendinger et al. (1989) contrast this style with the contemporaneous strike-slip and normal faults on the plateau to the west, which were seen at Stop 5-5 at Pasto Ventura and other southern Puna stops. The southern Puna extensional faults often have a generally northerly strike and dip 0° to 47°N; the strike-slip faults have a nearly east-west shortening direction like the thrust faults in the northern Sierras Pampeanas.
At this stop, Andalhuala Formation medium- to coarsegrained sandstones, conglomerates, and siltstones are affected by folds with northeast-trending axes that are typical of the first deformation stage. Elsewhere in the region, northeast-striking faults thrust Paleozoic basement over these units. The Miocene age of the Andalhuala Formation is based on 40Ar/39Ar dates of 6.68–7.14 Ma on interbedded tuffs (Allmendinger, 1986).
Field Guide Part 2: Arc Region in Chile
Suzanne Mahlburg Kay and Constantino Mpodozis
The objectives of Days 6 and 7 are to examine the Neogene to Recent history of the Miocene to Recent frontal volcanic arc region in Chile. During the late Oligocene to late Miocene (26–6 Ma), the frontal arc was located along the Maricunga Belt and then migrated eastward to the site of the Pleistocene to Recent Central Volcanic Zone arc. The southernmost Central Volcanic Zone Ojos del Salado region is visited on Day 6 and the Maricunga Belt on Day 7.
Day 6—Southernmost Central Volcanic Zone, Neogene Arc Migration, And Quaternary To Holocene Activity
Directions: Take paved Route 45 from Fiambalá through Chaschuil to the Argentine customs and immigration station at Las Grutas just east of Paso San Francisco and the Chilean border. The distance from Fiambalá to Paso San Francisco is ~205 km.
Introduction to the Volcanic Centers of the Ojos del Salado Region
The Ojos del Salado volcanic region is located near the southern termination of the modern Central Volcanic Zone near 27°S latitude (Figs. 6.1 and 6.2). This region is home to the Central Volcanic Zone arc and the late Oligocene to Miocene volcanic centers that erupted in the backarc of the Maricunga Belt arc, which is over 40 km to the west (Fig. 6.1). Between ca. 9 and 6 Ma, Maricunga Belt volcanic activity dramatically decreased and then ceased as volcanism increased significantly in the back arc. By 4 Ma, the main frontal volcanic arc was essentially in the Ojos del Salado region. The distribution, age, and geochemistry of the Ojos del Salado region volcanic rocks reflect complex magmatic-tectonic interactions associated with arc migration, crustal thickening, and uplift (Figs. 6.1 and O.10; Mpodozis et al., 1996; Kay et al., 1997, 1999, 2006). Considering the modern subduction geometry and assuming that the magmatic front of the Maricunga Belt was located ~100 km above the subducting Nazca slab, a significant portion of the forearc crust and frontal arc lithosphere should have been removed from below this region after 8 Ma (Kay and Mpodozis, 2000, 2002; Kay et al., 2006). The objectives of Day 6 are to see the Central Volcanic Zone volcanic front, view the Miocene backarc volcanic centers, and discuss the processes associated with eastward arc migration.
Most of the volcanic activity in the Ojos de Salado region can be summarized in the three episodes listed below (see Mpodozis et al., 1996). Centers mentioned are labeled on Figures 6.1 to 6.4; K/Ar ages are plotted in Figures 6.5 to 6.7. Discussions of the geology, physical Volcanology, and geochemistry of the centers can be found in González Ferrán et al. (1985), Baker et al. (1987), Mpodozis et al. (1996), Gardeweg et al. (1997, 1998, 2000), Kay and Mpodozis (2000, 2002), Kay et al. (1999, 2006), and Clavero et al. (2000).
Late Miocene (9–5 Ma) Volcanism behind the Dying Maricunga Belt Arc
An older group includes andesites and dacites erupted from isolated 9–7 Ma stratovolcanoes. They include a lava from the Laguna Verde volcano (8.9 ± 0.4 Ma), flow from the lower part of the Cerro Los Patos volcano in Argentina (7.6 ± 0.6 Ma), and 7.8–7.6 Ma lavas in the Cordon Foerster. A younger group includes an andesite along the Robertson Escarpment (6.5 ± 0.7 Ma), 5.9–5.1 Ma high-K andesites and dacites around the Wheelwright caldera, flow banded rhyolites associated with the Fuenzalida dome (5.8 ± 0.9 Ma), and dacites associated with the La Barda Complex (4.9 ± 0.4 Ma).
Pliocene (5–1.8 Ma) Volcanism at the Emerging Central Volcanic Zone Front
After volcanism ended in the Maricunga Belt, the front migrated eastward, and large-scale activity began in the Ojos del Salado region. From 4 to 1.4 Ma, diverse centers erupted moderate volumes of basaltic andesite to rhyolitic lavas. Older units include silicic andesitic flows at Los Patos volcano (4.5 ± 0.9 Ma), silicic andesitic and dacitic lavas and domes southwest of Tres Cruces (4.0–4.5 Ma; Gardeweg et al., 1997), glassy andesitic lavas from the Rodrigo cone (4.4 Ma), and andesites from the Peñas Blancas (4.2 Ma) and Ermitaño volcanoes (3.6–3.7 Ma) and a center north of the Cerro Laguna Verde volcano (3.9 ± 0.9 Ma). The Laguna Verde rhyodacitic ignimbrite erupted at 3.8–4.4 Ma. Younger eruptions include the thick 3.5–3.4 Ma dacitic flows that are the precursors to the Ojos del Salado complex (Fig. 6.6) and pyroxene andesite flows at Cerro El Muertito (3.2 ± 0.3 Ma) and Volcán del Inca (3.3 ± 0.3 Ma). Similar age silicic rocks erupted around the Nevado Tres Cruces (Fig. 6.7).
Pleistocene to Holocene Centers at the Central Volcanic Zone Front (<1.5 Ma)
The most striking centers of the Ojos del Salado region are the Quaternary volcanic edifices that form a northeast- to east-northeast–trending belt that extends for more than 50 km (Fig. 6.4). Most of these centers are well over 6000 m high. Major andesitic to dacitic stratovolcanic complexes near the border include Incahuasi (6610 m), San Francisco, Cerro Condor, and Cerro Peinado. Centers farther east include El Fraile, El Muerto (6420 m), Ojos del Salado (6880 m), Cerro Solo (6190 m), and Nevado de Tres Cruces (6330 m), which are dominantly dacitic dome complexes and lava domes. Centers with K/Ar and 40Ar/39Ar ages less than 1.5 Ma constitute more than half of the erupted rocks in the Ojos del Salado region. The dominance of andesites and presence of basaltic lavas (Incahuasi and San Francisco) in the Paso San Francisco region and dacitic complexes elsewhere can be related to regional differences with contractional stresses dominating in the west and extensional stresses along faults being more important in the east.
Stop 6-1: Las Grutas—View of the Latest Pliocene to Quaternary (Holocene?) San Francisco and Incahausi Volcanoes at the Southern End of the Central Volcanic Zone
Directions: Stop along Route 45 in Argentina just west of immigration and customs buildings and east of Paso San Francisco (pass at 4725 m) at 26°54′23″S; 68°23′55″W; 4424 m asl.
The centers of the Ojos del Salado region near Paso San Francisco (Incahuasi, San Francisco, and Falso Azufre) are typical Andean stratovolcanoes (typically 58%–61% SiO2) that dominantly erupted pyroxene and hornblende andesites. The Incahuasi and San Francisco volcanoes viewed from this stop also have monogenetic parasitic cones composed of olivine- and clinopyroxene-bearing basalts (~53% SiO2). Kay et al. (1999) interpreted these mafic lavas as products of mantle-derived magmas that were contaminated by siliceous magmas like those from the nearby Pliocene Laguna Verde ignimbrite (Figs. 6.2 and 6.4). Contamination occurred as the magmas ascended rapidly along deep-seated faults.
The east-northeast–trending Cordillera de San Buenaventura, just to the north of Las Grutas (Fig. 6.3) includes an east-northeast–trending chain of latest Pliocene to Pleistocene andesitic to dacitic domes with ages ranging from 2.3 to 0.4 Ma (Kay et al., 2006). The east-northeast–trending faults along which these domes were emplaced reflect complexities related to major structural blocks. Baldwin and Marrett (2004) have argued that the faulting and the domes in the Cordillera de San Buenaventura are related to a releasing bend connecting regional-scale, north-northeast–trending, right-lateral strike-slip faults.
Stop 6-2: Late Miocene Mulas Muertas Volcanic Center
Directions: Head west through Paso San Francisco crossing the border into Chile; stop along Chilean National Route 31 near 26°54′54″S; 67°07′59″W; 4019 m asl.
The Mulas Muertas volcanic complex seen directly to the west (Figs. 6.4 and 6.5) from this stop is the largest of the backarc centers that erupted in the Ojos del Salado region as the late Miocene Maricunga volcanic arc was expiring to the west. The oldest part of the complex is a hornblende-biotite-sanidine dacite dome dated at 5.6 ± 0.4 Ma. The main eruption of the center produced two-pyroxene andesites and hornblende-bearing andesites and dacites with K/Ar ages of near 5.1–5 Ma. Other late Miocene backarc volcanic rocks in the region include dacitic lavas dated at 6.5 ± 0.7 Ma that along the Robertson escarpment to the west of the San Francisco volcano and andesitic to dacitic rocks with ages from 5.9 to 5.3 Ma near the Wheelwright volcano (see Fig. 6.5).
This stop also affords a nice view of the Incahuasi volcano up the valley to the south-southeast and a view of the El Fraile center directly to the south (see Figs. 6.4 and 6.5). Laguna Verde is the lake to the west.
Stop 6-3: Laguna Verde and Pliocene Laguna Verde Ignimbrite
Directions: Stop along Chilean Route 3 near 26°53′25″S; 68°29′0″W; 4375 m asl.
The Laguna Verde ignimbrite (Fig. 6.4) is a thick and widespread, pumice-rich, dacitic to rhyolitic ignimbrite. The upper levels are formed by a biotite-quartz-sanidine–bearing welded rhyolitic tuff (>74% SiO2) with K/Ar ages ranging from 4.4 to 3.8 Ma (Fig. 6.5; Mpodozis et al., 1996). The source is uncertain. Since the Laguna Verde ignimbrite is similar in both composition and age to the Laguna Amarga and Vallecito ignimbrites (Siebel et al., 2001; Kay et al., 2006) just to the northeast (Fig. 6.3), the source has been suggested to be related to the Laguna Amarga caldera described by Seggiaro et al. (2002). The Laguna Verde ignimbrite erupted as the volcanic front was moving to the present site in the southernmost Central Volcanic Zone (Mpodozis et al., 1996; Kay et al., 1997).
Stop 6-4: View of Pliocene to Holocene Ojos del Salado Complex
Directions: Stop along Chilean Route 31 near 26°54′55″S; 68°35′10″W; 4466 m asl.
The Ojos del Salado peak with a height of 6880 m is well seen through the valley to the south from this stop. The Ojos del Salado complex is a large constructional volcanic center built by numerous superposed small volcanic edifices. The center, which has been mapped and described by Gardeweg et al. (1997, 1998; see Fig. 6.6), can be seen to be extensively covered by air fall pumice deposits that originated from an explosive eruption from the Nevado de Tres Cruces to the west. The pumice deposits are thickest on the lower slopes. The lowermost part of the Ojos del Salado complex consists of thick dacitic flows that have yielded K/Ar ages of ca. 3.4–3.5 Ma (Mpodozis et al., 1996). These flows are similar in age to low-silica pyroxene andesite lavas dated at 3.2 ± 0.3 Ma at Cerro El Muertito (Figs. 6.4 and 6.5). The main part of the Ojos del Salado complex is composed of short, thick, dacitic lava flows and domes clustered in an area 13 km long by 12 km wide. The lower part dates to the early Pleistocene as indicated by a K/Ar age of ca. 1.53 Ma from a long, glassy, two-pyroxene andesite flow in the lower part. The upper part is composed of short, thick, youthful-looking glaciated dacitic lava flows. Those on the western flank form a north-northeast–trending line of small-volume, two-pyroxene, hornblende-biotite dacite cones. A summit complex is developed over northward-extending flows that have yielded a K/Ar age of 1.08 ± 0.09 Ma. The summit complex consists of two coalesced, east-west–aligned edifices whose diameters are less than 2.5 km. The hornblende-biotite dacites in these edifices yielded a 40Ar/39Ar date of 0.34 ± 0.19 Ma (Gardeweg et al. 1997, 1998). The active crater (1.3 × 0.5 km), which shows signs of fumarolic activity, is located between the two edifices.
The peaks seen to the south on the west side of the valley constitute the Cordon Foerster. The main center has a late Miocene K/Ar age of 7.5 ± 0.9 Ma (Fig. 6.5). The center on the east side of the valley is the Mulas Muertas volcano that was viewed from another perspective at the previous stop. The view here is of the hornblende-clinopyroxene–bearing andesites and dacites of the main eruptive phase with K/Ar ages of 5.0 ± 0.3 Ma and 5.1 ± 0.3 Ma.
Stop 6-5: Southern Rim of the Wheelwright Caldera
Directions: Stop along Chilean Route 31 near 26°54′40″S; 68°37′15″W; 4510 m asl. The outcrop along the road at this point is the Pliocene Laguna Verde ignimbrite.
The enigmatic Wheelwright caldera seen to the north of the road at this stop is defined by a ~19-km–wide, subcircular, topographic rim that surrounds a depression. The rim and depression are well seen in the satellite image in Figure 6.4. A problem with interpreting this structure is that no major ignimbrites are associated with the caldera-like depression. The main Wheelwright structure formed between 5.2 and 4.2 Ma (Baker et al., 1987; Clavero et al., 2000) just after the Maricunga arc shut off and as the new arc front was starting to form in the position of the modern southern Central Volcanic Zone.
A series of Pliocene stratovolcanoes sit along the rim of the Wheelwright caldera (see Fig. 6.5). These include the large Peñas Blancas volcano (4.2 ± 0.6 Ma) on the southwest side (seen to the northwest from this stop), the Ermitaño volcano (3.7 ± 0.6 Ma and 3.6 ± 0.6 Ma) on the southeast side, and a small unnamed center north (3.9 ± 0.9 Ma) of the Laguna Verde volcano (Fig. 6.5). The late Miocene (5.9 ± 0.9 Ma) Pico Wheelwright center sits just west of the Peñas Blancas center. The Volcán del Inca on the south side of the Peñas Blancas center (viewed immediately to the north from this stop) is an olivine-pyroxene basaltic andesite parasitic cone with a K/Ar age of 3.3 ± 0.3 Ma.
Stop 6-6: Early Miocene Segerstrom Basalt, Middle Miocene Cerros Amarillos Volcanic Center, and the Pliocene Rodrigo Volcano
Directions: Stop along Chilean Route 31 near 26°56′50″S; 68°45′58″W; 4500 m asl.
Geological features at this stop include the southwestern rim of the Wheelwright caldera to the northeast (discussed at Stop 6-5), Oligocene to Miocene sedimentary and mafic volcanic sequences of the Cordon Segerstrom to the north, altered middle Miocene Cerros Amarillos lavas to the east-northeast, and the early Pliocene Rodrigo Volcano to the south (Figs. 6.4 and 6.5).
The view to the north into the Cordón Segerstrom shows the basaltic and basaltic andesitic lavas called the Segerstrom basalt (González-Ferrán et al., 1985; Baker et al., 1987). These lavas cover the late Oligocene–early Miocene vol canic and sedi mentary Claudio Gay sequence (Mpodozis and Clavero, 2002), which is well exposed in the Cordillera Claudia Gay to the west (Stop 6-8). These sediments and volcanic rocks were emplaced in the backarc of the late Oligocene to Miocene volcanic arc, which is preserved in the Maricunga Belt some 35 km to the west (Day 7). The Segerstrom lavas are generally olivine- and plagioclase-bearing mafic flows (48%–57% SiO2) that dip gently eastward and have K-Ar ages of 25–24 Ma (Mpodozis et al., 1996; Kay et al., 1999). Their tholeiitic (FeO/MgO = 1.0–2.4) affinities, moderate (2%–3%) K2O contents, and low La/Ta (24–26) and Ba/La (16–19) ratios are consistent with eruption in a mildly extensional backarc. Their REE patterns (La/Yb = 9–14) are consistent with the parent lavas being small degree melts of mantle peridotite. Decreasing εNd (+4.3 to +1.5) and increasing 87Sr/86Sr (0.7039–0.7033) values show that the parent magmas were contaminated in the crust. The Segerstrom belt is one of several occurrences of Oligo cene and early Miocene backarc mafic lavas in the central Andes. Others are the Máquinas basalt in the Chilean flat slab (30°S; Kay et al., 1991) and the Chiar Kkollu basalts in the Bolivian Altiplano (Davidson and de Silva, 1992).
Isotopic differences between the Segerstrom basalts and the Central Volcanic Zone Incahuasi and San Francisco mafic flows (87Sr/86Sr ~0.7055; εNd ~+2) at Stop 6-1 reflect a change from a less enriched early Miocene backarc mantle to a more enriched Quaternary mantle wedge. This change is consistent with incorporation of crustal components into the mantle wedge by intense forearc subduction erosion at the time of arc migration (Kay and Mpodozis, 2000; Kay, 2006).
The Cerros Amarillos lavas seen to the northeast are one of the oldest manifestations of arc-like andesitic to dacitic volcanism in the backarc of the Maricunga Belt arc. A dacite sample from a dome in this group yielded a K/Ar age of 13.9 ± 0.7 Ma (Mpodozis et al., 1996). This age fits with the middle Miocene expansion of arc volcanism into the backarc of the southern Puna as indicated by the abundant Miocene volcanic rocks in that region.
The two-pyroxene andesite lavas (61% SiO2) from the Pliocene Rodrigo volcano to the south are distinctive for their glassy texture and chemical similarities to the southern Puna glassy lavas. The Rodrigo lavas have a K/Ar age of 4.4 ± 0.6 Ma (Mpodozis et al., 1996; Fig. 6.5). Chemical similarities of these lavas with the southern Puna lavas include pronounced arc-like HFSE depletions (La/Ta = 70–80) and steep REE patterns (La/Yb ~50). Volcanic rocks like these are largely restricted to the latest Miocene and Pliocene and reflect the special magmatic conditions at the time of arc migration (see Fig. 6.1; Kay and Mpodozis, 2002).
Stop 6-7: Nevado de Tres Cruces Complex and Campo de Piedra Pomez
Directions: Near 27°03′38″S; 68°55′00″W; 4350 m asl.
The Nevado de Tres Cruces Massif seen to the east and south at this stop is formed by three large dacitic cones—Norte in the north (6206 m), Centro in the center (6629 m), and Sur in the south (6748 m), aligned along a NNW-trending regional fault system. The southern cone is considered to be the source of the nearly 100 m thick dacitic Tres Cruces ignimbrite (Pampa Blanca in Fig. 6.7, conspicuous white ignimbrite on TM scene in Fig. 6.4) with a 40Ar/39Ar age of 0.067 Ma (Gardeweg et al., 2000). The lavas, domes, and ignimbrites of the Macizo Tres Cruces volcanic complex (Fig. 6.7) are younger than 1.5 Ma. The youngest 40Ar/39Ar age (0.28 ± 0.022 Ma) is from the southern cone. The dacitic pumice of the Campo de Piedra Pomez near the road at this stop has a 40Ar/39Ar age of 1.56 ± 0.15 Ma. Older Pliocene dacitic flows around the Nevado de Tres Cruces Massif include the La Barra flow (2.3 ± 0.3 Ma) to the west (Fig. 6.5), the Lemp (2.8 ± 0.3 Ma), and Cristi (2.5 ± 1.3 Ma) volcanoes to the north, and the Cerro El Plateado flow to the east (Gardeweg et al., 1997, 2000).
Stop 6-8: Rio Lamas—Middle Miocene Syntectonic Strata and Interbedded Ignimbrite Sheets in Cordillera Claudio Gay
Directions: Watch for turnoff to Rio Lamas Waterfall from Route 31 near 27°04′54″S; 68°55′ (there are several tracks). Stop is along road, northeast of waterfall at 27°04′54″S; 68°55′59″W; 4280 m asl. After this stop, take Route 31 west. The route heads north along the Salar de Maricunga in the pre-Andean depression. At the north end of the salar, take the main road north to Chilean Customs and Immigration, and then continue north to La Ola. From La Ola, the road passes the south side of the Salar de Pedernales and heads to Montandón. West of Montandón, take the paved road north to Portal del Inca and continue to the town of El Salvador, where there are restaurants and hotels. The distance from the turnoff from Route 31 to the paved road to Porta del Inca is ~100 km. The distance along paved road to El Salvador is ~32 km.
This stop at the southern termination of the Cordillera Claudio Gay provides a view of the synorogenic gravels and interbedded ignimbrites of the middle Miocene Rio Lamas sequence. Regionally, the Cordillera Claudio Gay (Figs. 6.1, 6.4, and 6.5) is an almost 150-km–long, north-south range with peaks up to 4800 m. At the latitude of Stop 6-8, the Cordillera Claudio Gay forms the eastern boundary of the modern pre-Andean depression that hosts the active, internally drained Salar de Maricunga and Salar de Pedernales basins. The western boundary of the pre-Andean depression is the Maricunga Belt, some 20 km to the west. The Maricunga Belt and the Salar de Maricunga will be well seen as the route of the trip heads west descending the western slope of the Cordillera de Claudia Gay.
The Cordillera de Claudia Gay exposes a more than 1.5-km–thick Eocene to middle Miocene volcanic and sedimentary sequence that overlies Late Paleozoic acidic volcanic and granitoid basement (Mpodozis and Clavero, 2002). The oldest Tertiary unit is the Claudio Gay sequence, which is mainly composed of coarse dacitic pyroclastic deposits and domes. Epiclastic sediments and saline-lake carbonates accumulated in small intramontane basins from the coeval Rio Juncalito sequence (Mpodozis and Clavero, 2002). The Claudio Gay sequence is capped by the late Oligocene–early Miocene backarc Segerstrom basalt (Stop 6-6). At ca. 20–21 Ma, this depositional regime changed dramatically as east-verging, high-angle reverse faults uplifted the Late Paleozoic basement of the Cordillera Claudia Gay. Intense volcanism followed as large 20–19 Ma dome complexes with extensive block and ash-flow aprons erupted along the northern Cordillera Claudio Gay. These domes are covered by the widespread 18–19 Ma dacitic Vega Helada ignimbrite that has been correlated with the huge Rio Frio ignimbrite (Cornejo and Mpodozis, 1996). These ignimbrites are likely linked with the collapse that created the giant Aguilar caldera ~79 km to the north (e.g., Mpodozis and Clavero, 2002). In the middle Miocene, the Cordillera de Claudia Gay was affected by the last compressional deformation in the region of the modern pre-Andean depression. Evidence for this deformation comes from the alluvial gravels interbedded with distal ignimbrites (K/Ar age of 15–16 Ma) in the Rio Lamas sequence. These gravels show progressive unconformities and intraformational folds indicative of synsedimentary deformation (Gardeweg et al., 1997; Mpodozis and Clavero, 2002).
Day 7—The Maricunga Belt: The Miocene Arc Front
The objective of Day 7 is to examine the Tertiary vol canic units of the Maricunga Belt where the active arc front was located from the late Oligocene to the late Miocene (26–6 Ma, Fig. 7.1). Stop 7-1 in the El Salvador mining district provides a view of pre-Maricunga Belt Cenozoic magmatism in the Chilean Precordillera just east of the Maricunga Belt (Figs. 7.1 and 7.2). Stop 7-2 examines the Miocene Atacama gravels associated with the uplift of the Puna plateau. Stops 7-3 to 7-6 provide representative views of the early Miocene, middle Miocene, and late Miocene magmatic styles of the Maricunga Belt.
Stop 7-1: Paleocene El Salvador Caldera and El Salvador Porphyry Copper Deposit
Directions: Take paved road south from El Salvador and stop along road ~3.5 km south near 26°17′34″S; 69°34′27″W; 2490 m asl.
The area around the El Salvador mining district provides an excellent opportunity to examine the earliest Cenozoic stages of pre-Maricunga Belt magmatism and associated mineralization along the western Andean slope in the Chilean Precordillera.
Cenozoic volcanism in this region, as elsewhere in the Central depression and Precordillera of northern Chile began in the Paleocene when a series of stratovolcanoes, collapse calderas, and rhyolitic dome complexes erupted over deformed Jurassic to Late Cretaceous volcanic and sedimentary rocks (Mpodozis and Ramos, 1990; Cornejo et al., 1993a, 1993b, 1997, 1999; Arriagada et al., 2006). One of the largest of these centers was the 61–57 Ma El Salvador caldera seen from this stop. This center erupted high-K, calc-alkaline trachybasalts and trachyandesites and sanidine-bearing rhyolitic lavas, domes, and tuffs in response to the collapse of a shallow-level, zoned magma chamber. The lavas and ignimbrites have within-plate trace element signatures consistent with eruption in a period of very slow or arrested Farallon plate subduction. REE patterns show the magmas equilibrated with residual amphibole in a moderately thick crust (Cornejo et al., 1993b, 1999; Cornejo and Matthews, 2001).
During a time of highly oblique, fast convergence between the Farallon and South American plates in the Eocene–early Oligo cene (45–35 Ma) (Pardo Casas and Molnar, 1987), the Incaic deformation affected large parts of the Chilean Precordillera and formed the more than 1000-km–long Domeyko fault system (Tomlinson et al., 1994; Cornejo et al., 1997; Maksaev and Zentilli, 1999). The major fault in the El Salvador region is the left-lateral strike-slip Sierra del Castillo fault (Fig. 7.2), which will be crossed on the route to Stop 7-2. This fault separates Mesozoic volcanic rocks on the west from Jurassic–early Cretaceous backarc basin marine sediments in the Tarapacá backarc basin on the east side (Mpodozis and Ramos, 1990; Cornejo et al., 1993a, 1998). The spectacular thin-skinned folds and thrusts developed in the dominantly Mesozoic carbonate sequences of the Potrerillos fold-thrust belt are beautifully exposed along the deep canyons of the Quebrada Asientos and Río de la Sal. This thrust belt is seen on the route to Stop 7-2 and as the road continues eastward through Montandón.
Ages on syn- and post-tectonic intrusions along the Domeyko fault system show that the Incaic deformation had begun in the El Salvador region by 42 Ma, was occurring at 36 Ma, and was over by 32 Ma (Tomlinson et al., 1994; Cornejo et al., 1993a). This deformation produced important thickening and exhumation and uplift of the crust as shown by apatite fission-track ages of 40–38 Ma (Maksaev and Zentilli, 1999; Nalpas et al., 2005). Extreme transpressive conditions trapped most of the magma in the crust. Small subvolcanic stocks and porphyries emplaced along major fault strands are hosts to giant Eocene-Oligocene porphyry copper deposits like those at Chuquicamata and La Escondida to the north.
The El Salvador porphyry copper, which is often considered the type example of Chilean porphyry copper deposit (Cornejo et al., 1997), is hosted in the large Indio Muerto rhyolitic dome (58 ± 2 Ma, sensitive high-resolution ion microprobe [SHRIMP] U-Pb age; Cornejo et al., 1997, 1999) along the rim of the Paleocene El Salvador caldera. The chemistry of the El Salvador porphyry fits with mineralization being associated with fluids expelled from water-rich magmas as tectonic thickening led to destabilization of amphibole-bearing lower crust and formed an eclogitic crustal keel in the Oligocene (Cornejo et al., 1999; Cornejo and Matthews, 2001).
Stop 7-2: Atacama Gravels, San Andreas Ignimbrite, and Uplift of the Andean Plateau
Directions: Continue south and east along the paved and major gravel roads toward the village of Montandón. Stop ~5.6–6 km east of the junction of the paved and gravel road on the road to Montandón near 26°21′45″S; 69°27′34″W; 2490 m asl.
The voluminous Miocene Atacama gravels (Mortimer, 1973), which occur along the western slope of the Andes, can be seen from this stop. These unconsolidated gravels, whose thicknesses exceed 400 m in places, filled drainage networks and covered the pre-middle Miocene topography. Those at this stop filled the paleovalley of the Quebrada Asientos; similar sequences filled the paleovalley of the Rió de la Sal to the north (Fig. 7.2). The gravels were shed westward from the rising Puna Plateau to the east from ca. 18–11 Ma (Cornejo et al., 1993a; Nalpas et al., 2005) as the hyperarid climate developed that has prevailed in the Atacama Desert since the Miocene (Alpers and Brimhall, 1988; Rech et al., 2006; Nalpas et al., 2005). 40Ar/39Ar ages on interbedded ashes in the gravels range from 15.3 to 10.5 Ma (Owen et al., 2003).
A regional pediment surface (Atacama Pediplain of Mortimer, 1973) developed as gravel accumulation ceased. Dating of cobbles on this surface using cosmogenic nuclides (10Be, 26Al, and 21Ne) yield exposure ages of 9 Ma (Nishiizumi et al., 2005). Near El Salvador, the pediment is covered by the widespread, upper Miocene San Andrés ignimbrite with biotite K/Ar ages ranging from 10.2 to 9.3 Ma (Clark et al., 1967; Mortimer, 1973; Cornejo et al., 1993a). The San Andrés ignimbrite is contemporaneous with Copiapó complex ignimbrites in the Maricunga Belt that will be seen at Stop 7-6. A remnant of the San Andrés ignimbrite on the pediment surface northeast of the abandoned Potrerillos airport can be seen from this stop.
Introduction to Stops 7-3 to 7-6: Oligocene to Miocene Maricunga Belt Volcanic Arc Front
The objective of Stops 7-3 to 7-6 is to examine the late Oligocene to late Miocene magmatic and tectonic history of the Maricunga Belt arc. The volcanism in the Maricunga Belt coincides with the shallowing of the subducting Nazca slab under the Chilean flat slab to the south (see Kay et al., 1991) and the uplift of the Puna plateau to the east. The magmatic history reflects crustal thickening as the subducting plate changed shape and modified the overlying crust and lithospheric mantle. Overviews of the Maricunga Belt can be found in Mpodozis et al. (1995) and Kay et al. (1994b, 1999). Other discussions of the Maricunga Belt region and individual volcanic centers are in Mercado (1982), González-Ferrán et al. (1985), Baker et al. (1987), Davidson and Mpodozis (1991), Walker et al. (1991), Mpodozis et al. (1991), Vila and Sillitoe (1991), Vila et al. (1991), Moscoso et al. (1993), Cornejo et al. (1993a, 1998), McKee et al. (1994), Sillitoe et al. (1991), Tomlinson et al. (1999), and Kay and Mpodozis (2000, 2001, 2002).
The petrology and geochemistry of the Maricunga Belt arc rocks are discussed by Kay et al. (1994b, 1999), Mpodozis et al. (1995), and Kay and Mpodozis (2000, 2002). The volcanic rocks are mostly medium- to high-K calc-alkaline andesites and dacites (56%–68% SiO2) whose geochemistry reflects a tem poral trend from low-pressure, pyroxene-dominated to higher pressure, garnet-bearing eclogitic residual mineral assemblages in equilibrium with the magmas. Increasingly steeper heavy REE patterns show a growing role for garnet, and decreasing negative Eu anomalies and increasing Sr and Na2O concentrations show a decreasing role for plagioclase. The trends are interpreted to reflect the temporal evolution of these magmas in relation to a cooling mantle wedge and a thickening crust. Extreme chemical characteristics in late Miocene to Pliocene lavas reflect magma evolution in a very thick crust (>65 km) and introduction of continental crust into the mantle wedge by forearc subduction erosion as the arc front was displaced eastward.
The magmatic and tectonic history of the Maricunga Belt can be summarized in five stages:
Late Oligocene to early Miocene (26 to 21–20 Ma). During the first stage, andesitic to dacitic volcanic complexes erupted over a moderately dipping subduction zone through about a 40-km–thick crust. Centers include the large andesitic to dacitic Cerro Bravos–Barros Negros stratovolcano and dome complexes in the north (Stop 7-3) and widespread, small multiple dacitic dome complexes in the south. Many of the dome complexes and associated tuff and pyroclastic breccia rings are hydrothermally altered. Epithermal, high-sulfidation, gold-silver and porphyry-style gold mineralization occurred during this stage at La Coipa, Esperanza, Pantanillo, Refugio, and La Pepa (Vila and Sillitoe, 1991; Mpodozis et al., 1995; Muntean and Einaudi, 2001).
Early Miocene (21–17 Ma). The second stage was characterized by a virtual volcanic lull during a period of compressional deformation and crustal thickening. Evidence for compressional deformation was seen in the Cordillera Claudio Gay at Stop 6-8 (Mpodozis and Clavero, 2002).
Middle Miocene (17–12–11 Ma). The third stage was marked by the construction of voluminous, andesitic to dacitic stratovolcanic complexes along the length of the Maricunga Belt. From north to south, these centers include the Doña Inés, Ojos de Maricunga (Stop 7-4), Santa Rosa, Pastillos-Pastillitos (Stop 7-5), La Laguna, and Cadillal-Yeguas Heladas volcanoes. Ignimbrites were emplaced during the initial stages at some centers. Most centers are little eroded and preserve much of their original form. The Atacama gravels seen at Stop 7-2 were deposited in the Precordillera and Central depression to the west at this time. The third stage ended with the emplacement of structurally controlled, shallow-level, quartz-dioritic stocks hosting gold and copper mineralization (Marte, Lobo, and Aldebarán-Casale gold porphyries; Vila and Sillitoe, 1991).
Late Miocene (11–10–7 Ma): A radical change in the distribution of volcanic centers occurred during the fourth stage as most of the volcanic activity north of 27.5°S became concentrated in the silicic andesitic to dacitic Copiapó volcanic complex (Stop 7-6).
Late Miocene to Pliocene (7 Ma to 5 Ma): The last stage is marked by the eruption of small-volume bimodal centers at the southern end of the Maricunga Belt. These eruptions include ignimbrites associated with the fault-controlled Jotabeche rhyodacitic caldera and glassy mafic andesitic to andesitic Pircas Negras lavas emplaced along faults (Mpodozis et al., 1991, 1995; Kay et al., 1991, 1994b). The chemistry (very high La/Yb ratios and Na2O and Sr contents) of these magmas are uncommon in the central Andes and have been associated with a very thick crust and contamination of the mantle wedge by crustal material removed from the forearc by subduction erosion (e.g., Kay and Mpodozis, 2002).
Stop 7-3: Late Oligocene–Early Miocene Cerros Bravos–Barros Negros Volcanic Complexes
Directions: From Stop 7-2 continue on major gravel road through Montandón to La Ola. At La Ola, take major road to south and stop near 26°40′40″S; 69°03′44″W; 3758 m asl.
The overlapping Cerros Bravos–Barros Negros volcanoes seen to the west from this stop (Figs. 7.1 and 7.2) were emplaced in the earliest volcanic episode of the Maricunga Belt. They erupted over a moderately dipping subduction zone through a crust that was likely near 40 km thick. These centers mark the frontal arc west of the backarc Segerstrom basalt seen at Stop 6-6. The Cerros Bravos–Barros Negros centers are located along sinistral northwest-trending strike-slip faults that are part of the Eocene Potrerillos–Maricunga system that parallels the southern side of the Paleozoic Pedernales batholith (Cornejo et al., 1993a; Mpodozis et al., 1995; Tomlinson et al., 1999). K/Ar ages from these centers range from 25 ± 1 Ma to 21.7 ± 1 Ma (Moscoso et al., 1993; Cornejo et al. 1993a; McKee et al., 1994; Kay et al., 1994b; Mpodozis et al., 1995; Tomlinson et al., 1999).
Activity at the Cerros Bravos–Barros Negros centers began with the eruption of small-volume rhyodacitic ignimbrites (eMi in Fig. 7.2) that were subsequently covered by main stage pyroxene- and hornblende-andesitic lava and block and ash deposits (eMa in Fig. 7.2). Crystal-rich hornblende and biotite-bearing dacitic domes (eMd in Fig. 7.2) were then emplaced in the cores of the stratovolcanoes. A series of subcircular domes that host the Esperanza epithermal gold and silver deposit (Vila and Sillitoe, 1991; Moscoso et al., 1993) were emplaced in an 8-km–long belt along a reactivated WNW-trending fault zone (Cornejo et al. 1993a) on the northeastern flank of Cerros Bravos . The domes at Esperanza are surrounded by plinian rhyo litic lapilli tuffs (eMt in Fig. 7.2) with K/Ar ages ranging from 24 to 20 Ma. Alteration ages range from 23.2 ± 1.4 to 19.3 ± 0.7 Ma (Sillitoe et al. 1991; Moscoso et al. 1993). The northeasternmost domes are intruded by unaltered dacitic porphyry dikes with biotite K/Ar ages of 22.5 ± 1.2 and 22.4 ± 0.7 Ma (Cornejo et al., 1993a; Kay et al., 1994b).
A few kilometers south of Cerro Bravos, a multistage dome complex was emplaced at La Coipa. The La Coipa complex is located where west-verging, north-trending thrusts (Vegas la Junta) intersect the sinistral northwest-trending Quebrada Indagua fault. The La Coipa dome cluster consists of five 300- to 400-m–wide, crystal-rich, dacitic to rhyodacitic domes. K/Ar ages from unaltered early Miocene domes range from 24.6 ± 0.9 Ma to 22.9 ± 0.8 Ma (Zentilli 1974; Moscoso et al., 1993); ages of hydrothermally altered domes range from 20.2 ± 1.2 to 17.3 ± 1.0 Ma (Sillitoe et al., 1991; Moscoso et al., 1993; Mpodozis et al., 1995). The domes are surrounded by an extensive coeval blanket of intensely altered, coarse pyroclastic breccias, and poorly welded lapilli tuffs with biotite K/Ar ages of 24.7 ± 0.7 and 24.0 ± 1.0 Ma. These rocks host the epithermal silver-gold mineralization at the La Coipa, Can-Can, and Purén mines and prospects (e.g., Cornejo et al., 1993a; Mpodozis et al., 1995). Volcanism in the La Coipa region ended with the emplacement of middle Miocene domes (mMd in Fig. 7.2; Mpodozis et al., 1995).
Stop 7-4: Early Middle Miocene Volcanism—Ojos de Maricunga Stratovolcano
Directions: Continue south, passing the Chilean Customs, and take a major gravel road to the southeast toward Laguna Santa Rosa; stop near 27°00′36″S; 69°03′56″W; 3785 m asl.
The largest group of middle Miocene volcanic centers in the Maricunga Belt is the cluster of stratovolcanoes to the west and south of the Salar de Maricunga. The northernmost of these centers is seen to the west from this stop (Figs. 7.1 and 7.2). This is the well-preserved Ojos de Maricunga volcano (4985 m) with a basal diameter of 15 km and a central crater filled by a dacite dome dated at 15.8 ± 0.9 Ma (whole-rock K/Ar; Mpodozis et al., 1995). The slopes of the volcano are covered by unconsolidated hornblende andesite block and ash deposits (mMl on Fig. 7.2) that have yielded K/Ar ages from 16.2 ± 0.6 to 15.1 ± 0.7 Ma (Zentilli, 1974; Mpodozis et al., 1995) and a 40Ar/39Ar age of 14.5 ± 0.1 Ma (McKee et al., 1994). The block and ash deposits overlie two ignimbrites of uncertain origin (mMt on Fig. 7.2). The older is a welded red tuff (60%–62% SiO2) that is up to 100 m thick and has a whole-rock K/Ar age of 15.8 ± 0.8 Ma. The younger, which is exposed on the southwestern slope, is a slightly welded, pumice-rich biotite-bearing ignimbrite with biotite K/Ar ages of 14.3 ± 1.6 Ma and 13.7 ± 2.6 Ma (see Zentilli, 1974).
Other middle Miocene volcanic centers to the south are principally made of hornblende- and pyroxene-bearing andesite. They include the Santa Rosa, Cerro Lagunillas, Pastillitos, and Cerro Las Cluecas volcanoes (Figs. 7.4 and 7.5). Blocks from a coarse blanket of reworked pyroclastic block and ash deposits on the slopes of the Santa Rosa cone have K/Ar ages of 15.4 ± 0.55 Ma (hornblende, McKee et al., 1994) and 13.8 ± 0.6 Ma (whole rock, González-Ferrán et al., 1985). Whole-rock K/Ar ages from a Cerro Lagunillas lava, a block from the Pastillitos center, and a Cerro Las Cluecas lava range from 16.2 ± 0.6 to 15.9 ± 1.4 Ma (Mpodozis et al., 1995).
Stop 7-5: Earliest Late Miocene Volcanism and the Gold Porphyries: Pastillos Volcano and Mina Marte
Directions: Continue south, passing the road to Laguna Santa Rosa and Tambo and the Ex Mina Marte. South of the Ex Mina Marte, stop on the road near 27°12′04″S; 69°00′43″W; 4000 m asl.
Just to the east of the Ojos de Maricunga centers are the Villalobos and Pastillos volcanoes. Their K/Ar ages are summarized by Mpodozis et al. (1995) (Fig. 7.3). The Villalobos center is a deeply eroded stratovolcano with ages of 16.1 ± 2.1 and 14.8 ± 0.8 Ma (K/Ar) and 13.3 ± 0.5 Ma (40Ar/39Ar, McKee et al., 1994). The andesitic to dacitic Pastillos volcano is superimposed on the Villalobos volcano. Flows on the western flank of the Pastillos volcano have ages from 12.0 ± 1.8 to 12.9 ± 0.51 Ma. The Pastillos lavas and flows west of Cerro Las Cluecas (K/Ar ages of 13.1 ± 1.6–12.4 ± 1.4 Ma) are the youngest in the Ojos de Maricunga region.
The Pastillos-Villalobos center is intruded by shallow level, hornblende-bearing, andesitic and dacitic to quartz diorite porphyries. Those hosting gold-rich quartz stockwork alternation in the Marte, Lobo, Escondido, and Valy prospects along the eastern flank of the Villalobos center are called gold porphyries (Sillitoe, 1994; Mpodozis et al., 1995). The main targets are in the underlying Oligocene tuffs and the core of Pastillos volcano. The gold porphyries were emplaced along a west- northwest–trending fault system that projects from the Valle Ancho to the east into the northwest-trending sinistral Eocene fault system of the Porterillos–La Coipa region to the west (Tomlinson et al., 1994). An unusual association of porphyry-style mineralization and epithermal advanced argillic alteration (telescoping) in the Marte-Lobo district was attributed to rapid erosion linked to volcanic edifice collapse as the porphyry intruded by Sillitoe (1994). A problem is that physical evidence for collapse at Marte is unclear (Mpodozis et al., 1995). The time of mineralization is based on alunite K/Ar ages of 13.3 ± 0.42 Ma (Sillitoe et al., 1991) and 12.0 ± 0.69 Ma (Zentilli et al., 1991) in acid-sulfate altered rocks and a K/Ar whole-rock age of 11.4 ± 0.5 Ma on the Valy porphyry (see Mpodozis et al., 1995).
The first Maricunga Belt deposit called a porphyry gold deposit was at the Marte prospect (Vila et al., 1991). This horn-blende-biotite diorite porphyry stock was emplaced and mineralized at ca. 13–14 Ma. The stock is in a 2-km–diameter subcircular depression that is considered to be the small eroded summit crater of the Pastillos stratovolcano (Sillitoe et al., 1991; Mpodozis et al., 1995). It consists of two mineralized intrusives and a younger weakly porphyritic microdiorite. The mineralization is in a stockwork of quartz veinlets surrounded by chlorite-sericite-clay alteration, accompanied by disseminated pyrite and hematite that likely formed by hypogene martitization (replacement of magnetite by hematite). Isolated remnants of hydro thermal biotite and alkali feldspar indicate an overprint of potassic alteration. The deposit contains ~66 tons of gold. An open-pit, heap-leaching operation began in 1989, but closed shortly after, due to metallurgical problems.
Stop 7-6: Late Miocene Copiapó Volcanic Complex
About 0.6 km south of Stop 7-5, turn west on a track south of a lagoon; continue to near 27°13′17″S; 69°04′56″W; 4130 m asl. This road is best traversed with four-wheel drive.
Late Miocene volcanic activity in the Maricunga Belt from 11 to 7 Ma was largely restricted to the Copiapó volcanic complex at the intersection of the northwest-trending Valle Ancho–Potrerillos fault system with the Maricunga Belt. The silicic andesitic to dacitic pyroclastic flows, domes, and lavas that make up the Copiapó complex cover an area of more than 200 km2 The Copiapó complex is shown on the satellite image in Figure 7.4 and the map in Figure 7.5 from Mpodozis et al. (1995). Volcanic activity took place in the two main stages discussed below. Ages are summarized in Mpodozis et al. (1995) and Kay et al. (1994b).
Stage 1 (11–10 Ma)
Volcanic activity in the Copiapó complex began with the construction of the Azufre stratovolcano, north of the Laguna Negro Francisco. The main edifice is made of hornblende dacite flows and block and ash-flow deposits that are intruded by late-stage, fine-grained dacitic porphyries. Extensive hydrothermal alteration accompanied emplacement of the por phyries. The northern side of the Azufre center is covered by flows from the Azufre Norte center. Whole-rock K/Ar ages are 11.2 ± 0.9 and 9.2 ± 0.9 Ma for the Azufre cone, 11.9 ± 0.7 and 10.5 ± 0.5 Ma for the Azufre Norte flows, and 10.8 ± 0.5 Ma for the Azufre porphyries. At about the same time, large explosive eruptions likely linked to caldera formation produced the well-exposed pyroclastic deposits north and west of the main Copiapó center. These 100- to 300-m–thick deposits are visible from this stop. A basal sequence consists of small-volume ignimbritic flows and block and ash deposits alternating with thin plinian ash falls and proximal detrital deposits, lahars, and fluvial sandstones. They are overlain by the light-gray Copiapó I ignimbrite, which is a small-volume, weakly to moderately welded unit with abundant pumice and lithic fragments. The younger and overlying pink to reddish-brown Copiapó II ignimbrite is a welded columnar-jointed cooling unit with concentrations of large pumice fragments at the top. Both ignimbrites are hornblende-bearing silicic andesite and/or dacite. Ignimbrite whole-rock K/Ar ages range from 11.0 ± 2.2 Ma to 9.6 ± 0.8 Ma.
To the south and east, the Copiapó ignimbrites interfinger with a large coalesced dome and lava-dome complex that is in turn intruded by two fine-grained porphyritic stocks with acid-sulfate hydrothermal alteration (Azufrera de Copiapó). Four whole-rock K/Ar ages from the dome complex range from 11.7 ± 0.6 to 9.8 ± 0.7 Ma. A porphyritic stock in the Azufrera has a K/Ar age of 10.8 ± 0.5 Ma.
Stage 2 (8–7 Ma)
The shield-like Copiapó dome complex produced in Stage 1 forms the base for the younger stage 2 Copiapó cone. The central edifice of the younger cone is composed of hornblende and biotite-bearing dacite lavas and lava domes. Four whole-rock K/Ar ages range from 8.4 ± 0.4 Ma to 6.9 ± 1.1 Ma. A small parasitic dome at Cerro San Román, north of Copiapó has a K/Ar age of 6.9 ± 1.6 Ma and a 40Ar/39Ar hornblende age of 6.6 ± 0.3 Ma (Mulja, 1986).
Directions: End of field trip. To reach the city of Copiapó, return to the gravel road heading north. Continue through Ex Mina Marte and turn west on the road to Tambo and Laguna Santa Rosa. Join Route 31 at La Puerta and continue through La Puerta and Puquois to intersect the paved road to the city of Copiapó. The distance from the turnoff to Laguna Santa Rosa to Copiapó is ~190 km. An alternative and better maintained route is to continue north from Ex Mina Marte to the intersection of Route 31 and follow this road west to Copiapó.
We would like to thank the participants of the Geological Society of America–Asociación Geológica Argentina “Backbone of the Americas” field trip 405 that was held 9–15 April 2006 for their enthusiastic participation and their comments that improved the presentation of this guide. We would particularly like to thank Pablo Caffe, Magdalena Koukharsky, and Alejandro Pérez for their assistance in leading the trip and their invaluable logistical support. Alan Glazner, Robert Kay, and Víctor Ramos provided valuable comments that improved the presentation of the guide. This field trip guide is dedicated to the memory of William Hiss, who died in Antofagasta de la Sierra during the field trip in 2006.