Implications for Lithospheric Composition and Evolution with an Emphasis on Lithospheric Foundering

Maﬁ c volcanism on the Puna Plateau | RESEARCH

Maﬁ c volcanism on the Puna Plateau, NW Argentina: Implications for lithospheric composition and evolution with an emphasis on lithospheric foundering

Scott T. Drew1,*, Mihai N. Ducea2, and Lindsay M. Schoenbohm1 1SCHOOL OF EARTH SCIENCES, OHIO STATE UNIVERSITY, 125 SOUTH OVAL MALL, 275 MENDENHALL LABORATORY, COLUMBUS, OHIO 43210, USA 2DEPARTMENT OF GEOSCIENCES, GOULD-SIMPSON BUILDING #77, 1040 E 4TH STREET, UNIVERSITY OF ARIZONA, TUCSON, ARIZONA 85721, USA

ABSTRACT

Lithospheric foundering has drawn increasing attention as an important contributor to continental plateau formation, especially as a driver for increased elevation, extension, and mafi c magmatism. This contribution focuses on the mafi c magmatism that led to the creation of monogenetic volcanoes throughout the Puna Plateau of NW Argentina. Lavas from these volcanoes provide a means to evaluate the recent petrotectonic development of the plateau and, in combination with basement intrusive rocks, determine the isotopic composition and long-term evolution of the lithosphere beneath the central Andean back-arc domain. Mafi c samples have trace-element concentrations and isotopic values typical of an enriched magma source region. We propose that the mafi c magmas originated from an aged, metasomatized subcontinental lithospheric mantle. The lavas have isotopic values nearly identical to those of Early Ordovician Famatinian gabbro and granodiorite. We suggest the most primitive Puna lavas and Famatinian magmas originated from the same subcontinental lithospheric mantle. This implies that at least a thin portion of the subcontinental lithospheric mantle has remained intact beneath NW Argentina for the past ~485 Ma. A comparison to coastal Jurassic igneous rocks and mantle xenoliths from the nearby Salta rift system suggests that the subcontinental lithospheric mantle is chemically decoupled from the depleted mantle to the west and east. This has been the case for hundreds of millions of years despite long-term tectono-magmatic activity along the proto-Andean and Andean margin and within the continental interior. Our data almost certainly rule out large delaminating bodies, suggesting instead partial or piecemeal removal of the lithosphere beneath the Puna Plateau.

LITHOSPHERE; v. 1; no. 5; p. 305–318; Data Repository item 2009236. doi: 10.1130/L54.1

INTRODUCTION a more complex picture of the evolution of the subcrustal lithosphere in a long-lived subduction-zone setting. Research into the formation of continental plateaus has increasingly The young (ca. 7 Ma to present), mostly mafi c magmatism that led to focused on the role of lithospheric loss in generating uplift, extension, the creation of numerous, small-volume monogenetic and simple poly- and magmatism, particularly in northern Tibet (England, 1993; Molnar genetic volcanoes throughout the Puna Plateau is of particular interest to et al., 1993; Haines et al., 2003), the central Andean Plateau (Garzione this study. Lavas from these volcanoes provide a means to evaluate the et al., 2006; Ehlers and Poulsen, 2009), and in eastern Anatolia (Göğüs petrotectonic development of the region and, in combination with base- and Pysklywec, 2008a). Early research in the southern part of the Andean ment intrusive rocks exposed throughout the plateau, determine the chem- Plateau, the Puna of NW Argentina, was critical in establishing the links ical composition and long-term evolution of the lithosphere beneath the among high elevation, thinned lithosphere, extensional faulting, and Central Andes. Specifi cally, mafi c volcanic rocks that are demonstrably small-volume, mafi c volcanism and lithospheric loss (Kay and Kay, 1993; derived from the upper mantle and have not undergone interaction with Kay et al., 1994). More recently, attention has refocused on this problem the overlying crust are of great importance in deciphering recent modifi - because of stable isotope paleoaltimetry, which suggests the rapid rise cations of the uppermost mantle in response to various tectonic processes of the northern Andean Plateau, the Altiplano, and links this increase in such as lithospheric foundering (e.g., Kay et al., 1994). elevation to convective removal of mantle lithosphere (Ghosh et al., 2006; Here, we present new whole-rock major- and trace-element data and Garzione et al., 2006). However, evidence in the Puna Plateau has not been Sr-Nd-Pb isotope values for lavas sampled from the Antofagasta and Pasto signifi cantly reevaluated for nearly two decades. This contribution focuses Ventura Basins in the southern Puna Plateau. We compare these isotopic on geochemical data from the mostly mafi c, small-volume, late Miocene, data to Sr and Nd isotope values from mafi c and intermediate intrusive Pliocene, and Quaternary lava fl ows that were originally instrumental in rocks of the Early Ordovician Famatinian arc (Otamendi et al., 2009), shaping ideas about lithospheric loss. Additional geochronological data demonstrating their striking similarity. This important fi nding suggests have been produced in this region in the intervening decades (e.g., Risse et that parts of the subcontinental lithospheric mantle and lower crust have al., 2008), but new geochemical data from mafi c rocks are scarce. Our new remained intact and nearly invariant in composition near an active conti- geochemical data and interpretations challenge previous analysis, painting nental margin for ~485 Ma. This also implies that recent foundering of lower crustal and mantle lithospheric material as proposed by Kay and *Present address: Brown University, Geological Sciences, 324 Brook Street, Kay (1993) and Kay et al. (1994) may have only thinned the subcontinen- Box 1846, Providence, RI 02912, USA; [email protected]. tal lithospheric mantle behind the active volcanic arc rather than removing

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it completely. Finally, it suggests that the “subduction signature” found intervening, high-elevation basins (Allmendinger et al., 1997; Coutand in recent back-arc central Andean volcanism may not only be the result et al., 2001). Crustal thickness beneath the Puna Plateau is on average of Neogene subduction and subduction erosion, but rather a composite 50–55 km, as opposed to 75 km beneath the Altiplano Plateau (Yuan et al., signal inherent to the subcontinental lithospheric mantle from long-lived 2000). The crust is thought to be mostly intermediate to felsic in composi- subduction beneath central South America. A comparison of our data to tion (Beck and Zandt, 2002; Tassara et al., 2006). The total thickness of coastal Jurassic volcanic and plutonic data supports this conclusion: the lithosphere is only 60 km, suggesting a relatively thin (5–10 km) mantle isotopic dissimilarity between our data and both the Jurassic magmatism lithosphere (Tassara et al., 2006). This value is considerably less than and Cretaceous mantle xenolith–bearing basalts from the Salta rift sys- expected considering the amount of shortening (>150 km) evidenced in tem shows that the back-arc region of the central Andean volcanic zone is the upper crust (Baby et al., 1997; Kley and Monaldi, 1998). underlain by a subcontinental lithospheric mantle keel that is an isolated, Miocene-Pliocene deformation on the Puna Plateau is characterized relatively homogeneous chemical domain intact since the Early Ordovi- by thrust faults that accommodate northwest-southeast shortening and cian, despite long-term periodic tectono-magmatic activity and postulated vertical extension (Allmendinger et al., 1989; Marrett et al., 1994). A recent lithospheric loss. compressional tectonic regime continues today around the margins of the plateau, where there is clear evidence for high-magnitude shortening in LATE CENOZOIC TECTONO-MAGMATIC DEVELOPMENT OF THE the Subandean zone in Peru and Bolivia (Hindle et al., 2005; McQuarrie et PUNA PLATEAU al., 2005) and in the Sierra Pampeanas of northwest Argentina (Allmend- inger, 1986; Coutand et al., 2001). However, at high elevations, an overall The Altiplano-Puna Plateau is the second largest modern plateau northeast-southwest to north-south extensional regime now exists, with on Earth and the largest formed in the absence of continental collision evidence for approximately N-S extension along normal and strike-slip (Fig. 1). The plateau is primarily the result of extensive crustal shortening faults since at least Quaternary time (Allmendinger et al., 1989; Marrett and thickening of the overriding plate as Nazca subducts beneath South et al., 1994). These extensional faults may refl ect the increase in gravi- America (Isacks, 1988; Allmendinger et al., 1997, and references therein). tational potential energy accompanying detachment of foundering litho- Magmatism may have contributed to crustal thickening (Beck and Zandt, sphere, but they could also refl ect an incipient gravitational extensional 2002). Loss of the lithospheric root, whether through delamination or con- collapse of the plateau as a whole (Schoenbohm and Strecker, 2009). vective removal, may have contributed to the modern elevation, structural Arc-related volcanism in the Andes has generated a frontal, linear development, and magmatism of the plateau as well (Kay et al., 1994; (orogen-parallel) arc in the Western Cordillera that takes the form of large Garzione et al., 2006). The geography of the southern segment of the stratovolcano complexes composed primarily of andesite to dacite lavas, plateau, the Puna Plateau, is characterized by fault-bounded ranges and domes, and pyroclastic fl ows (de Silva and Francis, 1991; Allmendinger et

75°W 70°W 65°W 67°30’W kilometers 0 5 10 15 20 field season 2007 Scoria cone 15°S Flow Normal fault

15°S Strike-slip fault 26 0’S ° ° ’ 26 0’S 26 A l t i p W Figure 1. (Left) A digital topog- l a raphy map of the central Andes e n s (SRTM 90 data). Black box shows t o Antofagasta e location of the Antofagasta and ~80 mm/a (NUVEL-1A) 20°S r Pasto Ventura Basins. Nazca– n

South America convergence vec- 20°S 6870 m C

Nazca plate o tor is from NUVEL-1A data (Mar- r rett and Strecker, 2000). (Right) d Hillshade map (SRTM 90 data) i

l ° 30’S l 30’S of the Antofagasta and Pasto e ° 26 r Ventura Basins. Included are

a P

P sampling locales, volcanoes, and u

u related lava ﬂ ows and fault traces. n n •

3,000 m 25°S

a a 25°S

Kilometers Pasto Ventura 0 100 200 300 400 67°30’W 75°W 70°W 65°W

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al., 1997, and references therein), and a more diffuse domain of large calc- share common petrographic and mineralogical features (Table 1). They alkaline ignimbrite centers known as the Altiplano-Puna volcanic complex have low phenocryst counts (mostly <10%), and many exhibit a glassy (de Silva, 1989). The Altiplano-Puna volcanic complex is the world’s larg- texture. Olivine is the most common and largest phenocryst. Clinopyrox- est continental volcanic fi eld, covering an area of at least 50,000 km2 in the ene is also present in many samples. Clinopyroxene is typically subor- southern Altiplano and northern Puna Plateaus. In the Puna Plateau, the dinate to olivine, smaller in size, and often clustered. Both olivine and Altiplano-Puna volcanic complex is marked by widespread andesite and clinopyroxene are typically euhedral. Olivine and clinopyroxene pheno- dacite eruptions that occurred from 16 to 12 Ma. Proliferation of magma- crysts often have glass and Fe-Ti–oxide inclusions. Elongate plagioclase tism was temporally linked to uplift in the southern Puna Plateau (Coira et phenocrysts are present in some samples. The matrix consists of olivine, al., 1993). Large-volume ignimbrite eruptions began during late Miocene clinopyroxene, plagioclase, and glass. Two of our samples, P07-ANT02 time (12–5 Ma) and continued until Pliocene time (Coira et al., 1993). and P07-J201, are more felsic (“evolved”), suggesting that their precursor They were likely the result of extensive crustal melting caused by mantle- magmas must have interacted with the local crustal basement. They have derived magmas entering the continental crust (Francis et al., 1989). In very low phenocryst counts (<5%) and are glassy. Orthopyroxene as well general, arc-related volcanism coincides with the timing of compression as apatite and Fe-Ti oxides are present in the evolved samples. in the region (Allmendinger et al., 1997). Risse et al. (2008) presented a compilation of existing ages for Puna Late Miocene to recent volcanism is typifi ed by andesite to dacite volcanism and added 22 new 40Ar/39Ar ages to the database. Ages range stratovolcanic-dome complexes and relatively small-volume, mafi c vol- from 7.3 Ma to samples that give zero ages. The older samples are less canoes in the Puna Plateau. Mafi c volcanism takes the form of mono- mafi c and possibly relate to the arc volcanism in the area. The majority of genetic and simple polygenetic cinder cones, lava cones, and lava fl ows. mafi c volcanics were erupted subsequent to 3.5 Ma. Three of our samples In the southern and central Puna, most of these lavas were erupted after are from lava fl ows dated by Risse et al. (2008): P07-ANT01 (0.3 Ma), 3.5 Ma, but some date back to 7.3 Ma (Risse et al., 2008). Lava fl ows are P07-CH01 (0.8 Ma), and P07-J401 (4.5 Ma). Some of our other samples basalt to andesite and are thought to be dominantly mantle derived (Kay can be assigned relative ages based on edifi ce degradation and crosscut- et al., 1994). Many of the volcanic centers in the southern Puna Plateau ting relationships (see Table 1). No reliable age data exist for lavas from appear to relate to local strike-slip and dip-slip, generally N-S–trending the Pasto Ventura region. Elemental concentrations and isotopic ratios faults (Marrett and Emerman, 1992; Marrett et al., 1994) (Fig. 1). In gen- do not show variation that correlates with age except that older volca- eral, individual volcanoes and related lava fl ows cover an area less than nism (>7 Ma) in the Puna tends to be more felsic in composition then the a few km2. The lava fl ows are usually no more than 3–4 m thick. Lavas younger, generally mafi c lavas. are dominantly calc-alkaline with minor shoshonitic composition in areas of thickest lithosphere. Kay et al. (1994) argued based on trace-element Major and Trace Elements compositions that some younger, mafi c fl ows have intraplate-like (ocean- island basalt [OIB]) geochemical signatures. They point to a spatial cor- Whole-rock samples were analyzed by X-ray fl uorescence (XRF) relation between the generally OIB geochemical signature and location of and inductively coupled plasma–mass spectrometry (ICP-MS) at the possible lithospheric foundering beneath the Puna Plateau. Geoanalytical Laboratory in the School of Earth and Environmental Sciences at Washington State University for major and trace elements GEOCHEMISTRY (Table 2) (Knaack, 2003; Johnson et al., 1999). The whole-rock samples were essentially liquids because of their low phenocryst count. Seventeen lava samples were collected from the southern Puna Plateau Justifi cation for whole-rock analysis includes low phenocryst count, along a N-S transect in April and May of 2007 (Fig. 1). We collected some sample homogeneity, and lack of alteration. The samples show a range samples from lava fl ows previously analyzed by Kay et al. (1994). We also of compositions, including basalt, trachybasalt, basaltic andesite, basal- provide analyses of never before sampled lavas in the Antofagasta Basin tic trachyandesite, andesite, and dacite (Fig. 2). A plot of MgO versus

and Pasto Ventura, the southernmost region of the plateau. Maﬁ c samples SiO2 shows a strong correlation (Fig. 3). Other major element trends

TABLE 1. LOCATION, ROCK TYPE, AGE, AND PHENOCRYST DESCRIPTION OF OUR 17 SAMPLES

Sample Location Latitude (°S) Longitude (°W) Rock Age % Phenocrysts Phenocryst type type (Ma) (%) P07AL-01 Antofagasta 26 08.354 67 22.936 B-TA <0.3 10 Plag (70), Ol (30) P07AL-02 Antofagasta 26 08.365 67 22.835 B-TA <0.3 10 Plag (60), Ol (40) P07AL-03 Antofagasta 26 11.777 67 24.217 B-TA <0.3 40 Plag (80), Ol (20) P07ANT-01 Antofagasta 26 07.786 67 25.582 B-TA 0.3 5 Ol (100) P07ANT-02 Antofagasta 26 11.801 67 24.226 D >0.3 2 Opx (60), Plag (20), Ap (20) P07J1-01 Jote 26 17.569 67 25.245 TB 8 Ol (60), Cpx (40) P07J2-01 Jote 26 16.706 67 20.514 A 2 Opx (70), Plag (15), Ap (10), Qtz (5) P07J3-01 Jote 26 20.185 67 22.709 B <4.5 5 Ol (90), Cpx (10) P07J4-01 Jote 26 19.764 67 23.387 BA 4.5 5 Ol (95), Plag (5) P07CH-01 Carachi Pampa 26 26.993 67 25.933 TB 0.8 15 Plag (50), Ol (45), Cpx (5) P07PV1-01 Pasto Ventura 26 42.271 67 14.923 TB 5 Ol (90), Cpx (10) P07PV1-02 Pasto Ventura 26 42.134 67 14.853 TB 10 Ol (50), Cpx (50) P07PV2-01 Pasto Ventura 26 43.906 67 12.655 B-TA 15 Ol (70), Cpx (20), Plag (10) P07PV2-02 Pasto Ventura 26 43.740 67 12.298 TB 5 Ol (80), Cpx (20) P07SUR1-01 South Pasto Ventura 26 53.974 67 16.497 B 8 Ol (75), Cpx (25) P07SUR2-01 South Pasto Ventura 26 53.064 67 15.387 B 10 Ol (65), Cpx (35) P07SUR3-01 South Pasto Ventura 26 51.895 67 15.289 B-TA 5 Ol (80), Cpx (20) Note: Rock-type abbreviations are: basalt (B), trachybasalt (TB), basaltic andesite (BA), basaltic-trachyandesite (B-TA), andesite (A), and dacite (D). Radiometric ages for samples P07ANT-01, P07J4-01, and P07CH-01 are from Risse et al. (2008). Relative ages are from field observations. Ol—olivine; Cpx—clinopyroxene; Ap—apatite; Plag—plagioclase; qtz—quartz.

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TABLE 2. MAJOR-OXIDE, TRACE-ELEMENT, RARE EARTH ELEMENT, AND RADIOGENIC ISOTOPIC COMPOSITIONS OF OUR 17 SAMPLES

Sample: P07-AL01 P07- AL02 P07-AL03 P07-ANT01 P07-ANT02 P07-J101 P07-J201 P07-J301 P07-J401

SiO2 (wt%) 53.06 52.65 53.03 52.17 64.22 51.00 62.85 51.69 52.25

TiO2 1.91 2.09 2.10 1.44 0.83 1.40 0.95 1.32 1.45

Al2O3 16.45 16.36 16.53 15.69 15.80 14.86 15.32 15.27 15.03 FeO* 8.09 8.07 8.10 7.14 4.53 7.69 4.78 7.77 7.49 MnO 0.14 0.14 0.13 0.13 0.07 0.14 0.07 0.14 0.12 MgO 6.43 6.15 6.17 6.58 2.47 8.75 2.81 9.06 6.81 CaO 7.20 6.96 6.90 7.78 3.91 8.56 4.53 8.39 8.55

Na2O 3.68 3.63 3.72 3.31 3.62 3.24 3.58 3.10 3.28

K2O 2.47 2.50 2.51 2.16 3.16 1.99 3.24 1.78 1.75

P2O5 0.57 0.50 0.50 0.35 0.23 0.42 0.24 0.34 0.45 LOI –0.04 –0.06 –0.25 1.38 –0.06 0.76 0.53 0.19 0.92 Ni (ppm) 99 85 85 107 23 167 29 187 111 Cr 161 130 133 212 82 385 81 434 271 Sc 21 20 19 21 11 24 10 25 20 V 189 181 180 187 110 198 127 196 211 Ga 18 20 19 17 22 17 21 16 18 Cu 45 46 42 48 28 45 22 55 35 Zn 95 94 94 86 92 77 87 79 96 La 49.53 40.81 40.46 36.23 45.93 40.74 46.97 42.13 39.50 Ce 97.38 82.55 81.95 73.20 89.91 79.15 92.81 82.48 80.46 Pr 11.46 10.00 9.94 8.75 10.36 9.26 10.68 9.72 9.91 Nd 43.29 39.13 38.92 33.27 37.63 35.62 38.75 36.65 39.25 Sm 8.24 7.89 7.87 6.49 6.83 6.72 6.91 6.73 7.69 Eu 2.32 2.25 2.23 1.80 1.47 1.91 1.48 1.82 2.00 Gd 6.67 6.52 6.43 5.38 4.99 5.63 5.13 5.46 6.26 Tb 0.96 0.94 0.95 0.80 0.66 0.83 0.71 0.80 0.89 Dy 5.26 5.26 5.26 4.50 3.40 4.82 3.68 4.57 4.85 Ho 0.96 0.98 0.98 0.85 0.62 0.89 0.65 0.88 0.87 Er 2.47 2.47 2.42 2.22 1.54 2.35 1.62 2.34 2.19 Tm 0.33 0.33 0.33 0.31 0.21 0.33 0.22 0.31 0.29 Yb 1.99 1.98 1.99 1.86 1.28 2.00 1.31 1.96 1.77 Lu 0.30 0.30 0.30 0.28 0.19 0.31 0.20 0.30 0.27 Ba 578 512 508 519 649 517 595 493 450 Th 7.62 6.55 6.37 7.19 16.84 6.48 19.78 6.62 5.63 Nb 39.84 36.46 36.37 25.07 10.00 31.31 12.57 23.08 16.73 Y 24.52 24.55 24.61 21.36 15.85 23.32 16.76 22.51 22.48 Hf 5.72 5.75 5.73 4.99 5.80 4.26 5.91 4.10 4.72 Ta 2.49 2.46 2.45 1.65 0.74 2.08 1.02 1.56 1.17 U 1.75 1.53 1.48 1.46 3.13 2.01 3.41 1.60 1.23 Pb 9.52 8.82 8.57 10.23 18.07 8.93 18.16 9.03 7.65 Rb 62.1 59.6 58.5 63.6 132.5 57.8 130.9 50.8 47.2 Cs 1.86 1.94 1.79 1.95 4.85 2.69 3.89 1.88 1.48 Sr 740 654 671 577 426 714 399 721 716 Zr 238 241 240 196 201 168 207 158 175 87Sr/86Sr 0.706252 0.705943 0.706494 0.711532 0.705595 0.709549 0.706509 0.707039 143Nd/144Nd 0.512497 0.512525 0.512403 0.512257 0.512603 0.512261 0.512490 0.512396 206Pb/204Pb 21.67635 18.80116 19.31945 18.95672 19.02703 33.09525 19.03653 18.92106 207Pb/204Pb 18.04610 15.65331 16.06665 15.69004 15.66314 27.41939 15.66453 15.66419 208Pb/204Pb 44.78762 38.83096 39.86414 39.18847 38.95627 68.50141 39.02849 39.03886 (continued)

compared to MgO include increasing FeO and CaO and decreasing our data set normalized to normal mid-ocean-ridge basalt (NMORB)

Na2O with increasing MgO (Fig. 4), which support a model of contami- (Sun and McDonough, 1989). In this case, normalization to NMORB nation of primitive magmas by means of fractional crystallization and/ allows for identiﬁ cation of elemental enrichment and depletion by or assimilation of crustal material to produce the more evolved compo- subduction processes of a typical, depleted upper mantle composition.

sitions. Two distinct sample groups (high and low) are seen on the Na2O Figure 6 shows the rare earth element (REE) composition of the three versus MgO plot, which reﬂ ect the “trachy” composition of some of most primitive samples as well as the most evolved sample compared to

our samples (see Fig. 2). TiO2 and Al2O3 show no correlation with MgO NMORB and normalized to depleted MORB mantle (DMM; Workman

(Fig. 4). The lack of correlation between TiO2 and MgO is primarily the and Hart, 2005). Important trends in these plots include: enrichment of result of six mafi c samples that contain a relatively high concentration the lighter elements compared to depleted upper mantle compositions; of Ti at a given MgO value. These six samples also have higher Nb and fractionation between fl uid and/or melt mobile large ion lithophile ele- Ta concentrations compared to our other samples but are still depleted ments (LILEs) and immobile high fi eld strength elements (HFSEs); in Nb-Ta relative to Th and La (Fig. 5). The variable Ti-Nb-Ta nega- fractionation between light rare earth elements (LREEs) and heavy rare tive anomaly is discussed later herein in more detail. K/Ti versus MgO earth elements (HREEs); and negative Ti-Nb-Ta anomalies. shows a fairly strong correlation (Fig. 4), especially if the Ti anomaly The fl uid and melt mobile elements (Cs, Rb, Ba, Th, La, Ce, Pb) are is taken into consideration. This correlation suggests that crustal mate- enriched compared to mantle values and the HFSEs and HREEs. These rial is being assimilated during differentiation of basalt toward more elements have only a weak correlation with wt% MgO (R value for Rb evolved compositions. versus MgO = 0.88, Cs = 0.82, Th = 0.79) or none at all (Ba = 0.42, Ce Figure 5 shows the overall trace-element composition of the three = 0.27, La = 0.22), suggesting that their enrichment is a signature of the most primitive samples (P07-J101, P07-J301, and P07-SUR101) from magma source region rather than a local crustal infl uence. Thus, there

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TABLE 2. MAJOR-OXIDE, TRACE-ELEMENT, RARE EARTH ELEMENT, AND RADIOGENIC ISOTOPIC COMPOSITIONS OF OUR 17 SAMPLES (continued)

Sample: P07-CH101 P07-PV101 P07-PV102 P07-PV201 P07-PV202 P07-SUR101 P07-SUR201 P07-SUR301

SiO2 (wt%) 51.55 50.79 51.66 54.45 51.69 47.53 48.77 54.62

TiO2 1.45 1.22 1.08 1.05 1.05 1.28 1.17 0.95

Al2O3 15.99 15.09 14.08 14.96 14.65 15.20 14.47 15.55 FeO* 7.92 7.91 6.89 6.18 6.48 9.18 7.50 6.52 MnO 0.14 0.15 0.13 0.11 0.11 0.16 0.14 0.13 MgO 7.03 8.72 7.49 5.84 6.38 10.30 7.84 6.48 CaO 8.35 8.92 9.39 8.03 8.59 10.06 10.38 7.45

Na2O 3.69 3.45 3.26 3.68 3.39 2.81 3.15 3.38

K2O 1.69 1.82 1.91 2.20 1.92 1.41 1.64 2.38

P2O5 0.39 0.48 0.46 0.48 0.39 0.45 0.41 0.28 LOI 0.67 0.08 1.82 1.87 2.62 0.55 2.58 0.92 Ni (ppm) 91 144 104 74 102 163 110 89 Cr 250 346 328 204 283 505 354 278 Sc 23 25 22 17 19 29 22 21 V 183 214 185 164 158 248 198 170 Ga 19 18 18 19 18 20 18 19 Cu 47 46 54 33 39 56 45 34 Zn 87 83 80 81 83 91 84 72 La 37.84 52.27 54.47 55.45 45.14 36.65 38.49 33.62 Ce 74.80 99.74 99.78 103.07 84.67 75.17 76.26 65.40 Pr 8.84 11.66 11.48 11.83 9.94 9.41 9.18 7.86 Nd 33.95 44.54 43.02 44.13 37.99 38.04 36.12 30.54 Sm 6.53 8.21 7.59 7.75 6.96 7.58 7.05 6.11 Eu 1.87 2.19 2.05 2.06 1.89 2.05 1.90 1.65 Gd 5.54 6.42 5.86 5.88 5.38 6.13 5.73 5.04 Tb 0.82 0.91 0.81 0.79 0.74 0.87 0.82 0.73 Dy 4.74 5.01 4.37 4.34 3.98 4.80 4.58 4.08 Ho 0.90 0.93 0.80 0.79 0.74 0.88 0.85 0.75 Er 2.33 2.37 2.07 2.03 1.82 2.23 2.14 1.92 Tm 0.32 0.33 0.28 0.28 0.25 0.31 0.29 0.26 Yb 2.01 2.03 1.71 1.70 1.54 1.84 1.78 1.60 Lu 0.30 0.30 0.26 0.25 0.23 0.28 0.27 0.24 Ba 484 608 617 697 641 484 470 761 Th 5.86 8.92 9.32 10.42 8.40 5.34 6.71 8.57 Nb 24.47 17.73 18.29 20.09 15.25 14.14 17.08 15.73 Y 22.93 24.04 20.84 20.51 18.91 22.37 21.69 19.32 Hf 4.36 4.76 4.67 5.46 5.34 5.25 5.20 4.50 Ta 1.50 1.05 1.08 1.27 0.94 0.76 1.00 1.09 U 1.27 1.83 2.60 3.40 2.34 0.94 1.49 2.44 Pb 8.69 8.50 10.14 13.66 12.37 5.73 7.96 13.24 Rb 43.8 47.4 50.8 64.6 50.7 28.8 41.2 68.0 Cs 1.14 1.01 1.78 2.77 1.90 0.51 0.98 2.31 Sr 680 998 853 930 810 785 753 702 Zr 177 179 177 212 210 201 201 159 87Sr/86Sr 0.705529 0.705621 0.706520 0.705534 0.706647 0.706330 143Nd/144Nd 0.512552 0.512569 0.512521 0.512600 0.512535 0.512461 206Pb/204Pb 18.88831 18.98622 18.81676 18.83085 21.77646 18.87297 207Pb/204Pb 15.65461 15.64703 15.65606 15.63798 18.05348 15.65660 208Pb/204Pb 38.92797 39.07449 38.94734 38.86455 44.99540 39.10891 Note: LOI—loss on ignition.

8 12

Trachyandesite 10 7 Basaltic- trachyandesite 8 6 Trachy-

O (wt%) basalt 2 6

O+K Dacite MgO (wt%) MgO

2 5

Na Andesite 4 Basaltic andesite 4 Basalt 2 R = .92

3 0 45 50 55 60 65 70 45 50 55 60 65

SiO2 (wt%) SiO2 (wt%)

Figure 2. Total alkali versus silica (TAS) diagram for our Puna Plateau Figure 3. MgO (wt%) versus SiO2 (wt%) for our Puna Plateau samples. samples. Samples are basalt, trachybasalt, basaltic andesite, basaltic-trachyandesite, andesite, and dacite. Fields are after Le Maitre et al. (1989).

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10 3.8

9 3.6

8 3.4

7 3.2 O (wt%) O 2 FeO* (wt%) Na 6 3

5 2.8 R = .84 R = .69

4 2.6 024681012 024681012 MgO (wt%) MgO (wt%)

11 17

10 16.5

16 8

7 15.5 (wt%) 3 O 2 CaO (wt%) 6 Al 15

5 14.5 4 R = .31 R = .88

3 14 024681012 024681012 MgO (wt%) MgO (wt%)

2.2 3.5

2 3

1.8 2.5

1.6 2 (wt%) 2 K/Ti 1.4 TiO 1.5 1.2

1 1 R = .20 R = .79

0.8 0.5 024681012 024681012 MgO (wt%) MgO (wt%) Figure 4. Major oxides versus MgO (wt%) for our Puna Plateau samples. K/Ti versus MgO is provided as a check for crustal contamination. Line of best-ﬁ t and R values are given for reference.

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1000 is a clear subduction signature in the mafi c lavas, which is exemplifi ed by the similarity between LILE concentrations and GLOSS (Plank and Langmuir, 1998) seen in Figure 5. GLOSS is a global average subducted sediment composition and is a useful tool for determining if fl uid and/ 100 or melt from the sediment have contributed to the composition of the magma source region, typically by causing enrichment in LILEs and LREEs. Certain HFSEs (Nb and Ta) do not correlate with MgO. There is signifi cant variation in depletion of these elements in rocks of simi-

10 lar SiO2 and MgO concentrations. Therefore, the Ti-Nb-Ta signature is also considered to be a characteristic of the magma source region. The HREEs are slightly enriched compared to DMM, possibly due to frac- Sample/NMORB tional crystallization, but are depleted relative to NMORB. This deple- NMORB 1 tion suggests long-term melt extraction from the mantle source region. Further, the evolved samples show lower middle (M) REE and HREE compositions than the basalts, indicating a strong crustal component.

Sr-Nd-Pb 0.1 Cs Ba U Ta La Pb Sr Sm Hf Ti Tb Y Er Yb Rb Th Nb K Ce Pr Nd Zr Eu Gd Dy Ho Tm Lu We analyzed fourteen of our samples for Sr, Nd, and Pb isotopes Element (Table 2). The isotopic ratios of 87Sr/86Sr and 143Nd/144Nd were mea- Figure 5. Normal mid-ocean-ridge basalt (NMORB)–normalized sured by thermal ionization mass spectrometry (TIMS) on whole-rock (Sun and McDonough, 1989) trace-element diagram showing samples. Fresh whole-rock samples were crushed to a ﬁ ne powder in a the composition of samples P07-J101, P07-J301, and P07-SUR101. shatterbox. Rock powders were put in large Savillex vials and dissolved These samples are the three most primitive basalts from our in mixtures of hot concentrated HF-HNO3 or, alternatively, mixtures of sample set and are representative of the subcontinental litho- cold concentrated HF-HClO . Strontium and the bulk of the REEs were spheric mantle beneath the Puna Plateau. Global average sub- 4 separated in cation columns containing AG50W-X4 resin, using 1 N to ducted sediment (GLOSS) compositions (Plank and Langmuir, 1998) for large ion lithophile elements (LILEs) and Nb-Ta are plot- 4 N HCl. Separation of Sm and Nd was achieved in anion columns con- ted as diamonds for a subduction signature reference. taining LN Spec resin, using 0.1 N to 0.5 N HCl. Strontium samples

were loaded onto single Ta fi laments with Ta2O5 slurry in order to keep the Sr cut loaded onto the middle third of the fi lament. Neodymium samples were loaded onto single Re fi laments using resin beads. 1000 Mass spectrometric analyses were carried out at the University of Arizona on an automated VG Sector multicollector instrument fi tted with adjustable 1011 Ω Faraday collectors and a Daly photomultiplier (Ducea and Saleeby, 1998). Typical runs consisted of acquisition of 100 isotopic ratios. Fifteen analyses of standard Sr987 during the course of this study yielded mean ratios of: 87Sr/86Sr = 0.710255 ± 7 and 84Sr/86Sr = 100 0.056316 ± 12. Fifteen measurements of the La Jolla Nd standard were performed during the course of this study. The standard runs yielded the following isotopic ratios: 142Nd/144Nd = 1.14184 ± 2, 143Nd/144Nd = 5.118530 ± 11, 145Nd/144Nd = 0.348390 ± 2, and 150Nd/144Nd = 0.23638 SCLM magma

Sample/DMM ± 2. The Sr isotopic ratios of standards and samples were normalized 10 N-MORB to 86Sr/88Sr = 0.1194, whereas the Nd isotopic ratios were normalized to 146Nd/144Nd = 0.7219. The estimated analytical ±2σ uncertainties for samples analyzed in this study are: 87Sr/86Sr = 0.0014%, and 143Nd/144Nd crustal = 0.0012%. Procedural blanks averaged from ﬁ ve determinations were: magma Rb 10 pg; Sr 150 pg; Sm 2.7 pg; and Nd 5.5 pg. Washes from the cation column separation were used for separating 1 La Ce Pr Nd Sm Eu Gd Tb Dy Ho Er Yb Lu Pb in Sr–Spec resin (Eichrom, Darien, Illinois) columns using protocol developed at the University of Arizona. Samples are being loaded in Element 8 M HNO3 in the Sr–Spec columns. Pb elution was achieved via 8 M Figure 6. Rare earth element (REE) diagram normalized to HCl. Pb isotope analysis was conducted on a GV Instruments multicol- depleted mid-ocean-ridge basalt (MORB) mantle (DMM; Work- lector inductively coupled plasma–mass spectrometer (MC-ICP-MS) at man and Hart, 2005) showing the composition of samples the University of Arizona (Thibodeau et al., 2007). Samples were intro- P07-J101, P07-J301, and P07-SUR101, the same as in Figure 3. duced into the instrument by free aspiration with a low-ﬂ ow concentric P07-ANT02 is plotted as a reference for evolved Puna magmas nebulizer into a water-cooled chamber. A blank, consisting of 2% HNO , that are derived from the continental crust. We use normal (N) 3 was run before each sample. Before analysis, all samples were spiked MORB (Sun and McDonough, 1989) to represent the approxi- mate REE composition of an uncontaminated magma derived with a Tl solution to achieve a Pb/Tl ratio of ~10. Throughout the experi- from the asthenosphere beneath the subcontinental litho- ment, the standard National Bureau of Standards (NBS)-981 was run to spheric mantle (SCLM). monitor the stability of the instrument.

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All results were Hg-corrected and empirically normalized to Tl by cal evidence for crustal interaction/magma mixing in our most primitive using an exponential law correction. To correct for machine and inter- samples. Sr and Nd isotopic compositions from our basalts are therefore laboratory bias, all results were normalized to values reported by Galer interpreted to be representative isotopic values for the magma source and Abouchami (2004) for the NBS-981 standard (206Pb/204Pb = 16.9405, region. The 206Pb/204Pb, 207Pb/204Pb, and 208Pb/204Pb values have ranges of 207Pb/204Pb = 15.4963, and 208Pb/204Pb = 36.7219). Internal error reﬂ ects 18.8012–19.0365, 15.6380–15.6900, and 38.8310–39.1885, respectively the reproducibility of the measurements on individual samples, whereas (Fig. 8). These values, including those for our basalts, are relatively radio- external errors are derived from long-term reproducibility of NBS-981 Pb genic compared to values for a depleted mantle magma source region. standard and result in part from the mass bias effects within the instru- Four samples are not included in our Pb isotope diagrams due to failure to ment. In all cases, external error exceeds the internal errors and is reported collect the correct Pb cut in the Sr–Spec columns. here. External errors associated with each Pb isotopic ratio are as follows: 206Pb/204Pb = 0.028%, 207Pb/204Pb = 0.028%, and 208Pb/204Pb = 0.031%. We MAGMA PETROGENESIS AND EVOLUTION OF THE CENTRAL did not perform any age correction to the measured isotopic ratios, given ANDEAN LITHOSPHERE the young age of the volcanic rocks. The samples show two clusters in 87Sr/86Sr versus 143Nd/144Nd space Small-volume lavas have erupted on the Puna Plateau over approxi- (Fig. 7). One grouping consists of the two evolved lavas at high 87Sr/86Sr mately the past 7 Ma (e.g., Kay et al., 1994; Risse et al., 2008). Our sam- and low 143Nd/144Nd. The other grouping has a range of 87Sr/86Sr values ples span the majority of this time (Table 1). Major-element chemistry between 0.70553 and 0.70704 and 143Nd/144Nd values between 0.51240 shows that 15 of our samples are basalt, trachybasalt, basaltic andesite, and 87 86 143 144 and 0.51260. The Sr/ Sr and Nd/ Nd values for the samples display basaltic trachyandesites, some with <50.0 wt% SiO2 and >8.0 wt% MgO, only a weak correlation with wt% MgO. Further, there is no mineralogi- which are indicative of a mantle-derived melt that has experienced little to

no crustal interaction. One sample, P07-SUR101 (47.52 wt% SiO2, 10.30 wt% MgO, 505 ppm Cr, and 163 ppm Ni), is the most primitive basalt yet reported for the central Andean Plateau region in NW Argentina. Unlike 0.5127 some of the basaltic andesite, this and other basalt samples do not have crustal xenoliths or quartz/feldspar xenocrysts. Further, P07-SUR101 con- 0.5126 100 (LM) tains only ~8% phenocrysts, which are mostly olivine (Table 1). Removal 80 of phases reduces the MgO content by ~2%, leaving a mantle-type MgO 80 0.5125 concentration of >8 wt%. The basalts experienced minimal interaction with the overlying crust during ascent because they show no chemical or petrographic evidence for assimilation of felsic rock. Therefore, the ele- 0.5124 Nd 40 40 mental and isotopic composition of our basalt and trachybasalt samples 144 is useful for evaluating the chemical characteristics of the magma source Nd/ 0.5123 region beneath the Puna Plateau. Two of our samples are a silicic andesite 143 20 20 and a dacite with low wt% MgO. They likely have an extensive crustal 0.5122 component and thus provide insight into the composition of possible crustal assimilant. For breadth of study, we have included all previous iso- CC1 topic analyses from the Puna Plateau in Figures 7, 8, 9, and 10. 0.5121 Importantly, LILEs and LREEs in basalt samples show enrichment CC2 compared to NMORB and enrichment relative to HFSEs and HREEs in 0.5120 the same sample (Figs. 5 and 6). This is a widely documented phenom- 0.704 0.708 0.712 0.716 0.720 0.724 0.728 enon in convergent-plate settings, and it suggests that Puna back-arc mag- 87Sr/86Sr mas originated from a magma source region that was enriched during a Figure 7. 87Sr/86Sr versus 143Nd/144Nd plot comparing Puna Plateau lava com- time when the mantle currently beneath the Puna Plateau was part of the positions (fi lled circles are this study; empty circles are compiled from the mantle wedge fl uidization zone. Spatial and temporal variations in Th, literature) to that of Famatinian gabbro (empty squares) and granodiorite Ba, La, and other LILE and LREE concentrations indicate a complicated (squares with crosses). Compiled data are from Caffe et al. (2002), de Silva mixture of fl uid and possibly melt addition to the magma source region. (1991), de Silva et al. (1994), Deruelle (1982), Francis et al. (1989), Guzmán Samples also show variable HFSE negative anomalies (Fig. 5). This et al. (2006), Knox et al. (1989), Kraemer (1999), Kraemer et al. (1999), Kay et al. (1994), Lindsay et al. (2001), Mattioli et al. (2006), Ort et al. (1996), Petri- occurs most prominently with Nb and Ta. Most samples show large nega- novic et al. (2005), Schnurr et al. (2007), and Siebel et al. (2001). See supple- tive Nb-Ta anomalies compared to other incompatible elements, which mental material for more information.1 The lithospheric mantle (LM) source are common in convergent-margin magmas. However, there is variation in region is represented by our most primitive sample, P07-SUR101. CC1 is the degree of the Nb-Ta negative anomalies and concentration at similar crustal end member 1, which is a high-86Sr and low-143Nd/144Nd Famatinian SiO2 and MgO wt%. For example, the Nb concentration range in basalt granodiorite. This end member may be located in the lower or midcrust. and trachybasalt samples from our data set is 14.1–31.3 ppm. There is no CC2 is crustal end member 2, which is a hybrid composition consisting of the least radiogenic 87Sr/86Sr and 143Nd/144Nd ratios from the Famatinian correlation between Nb-Ta and MgO. Thus, we consider Nb and Ta con- granodiorite data. This end member may be located in the lowermost crust centrations and related anomalies to be inherited from the magma source as a more mafi c layer that gained its lower Nd isotopic composition from region, not the overlying continental crust. Lavas with higher Ti-Nb-Ta subduction-related magmatism prior to Puna magmatism. concentrations, and thus smaller negative anomalies relative to neighbor- ing elements on a trace-element diagram, are scattered throughout the 1GSA Data Repository Item 2009236, Table DR1, is available at www.geosociety. Puna Plateau. They also do not show a temporal pattern, making varia- org/pubs/ft2009.htm, or on request from [email protected], Documents Sec- tions in recent melt removal an unlikely explanation. It is possible they retary, GSA, P.O. Box 9140, Boulder, CO 80301-9140, USA. were derived from source regions with increased amounts of amphibole,

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23.00 40.50 A B 22.00 40.00

21.00 39.50

CC Pb Pb 204 20.00 204 39.00 Pb/ Pb/ 206 208 CC LM 19.00 38.50 LM MORB MORB 18.00 38.00

17.00 37.50 0.512 0.5122 0.5124 0.5126 0.5128 0.513 0.5132 0.5134 17.5 18 18.5 19 19.5

143Nd/144Nd 206Pb/204Pb Figure 8. (A) 206Pb/204Pb versus 143Nd/144Nd comparing Puna lava compositions to that of Jurassic volcanic and plutonic rocks (Lucassen et al., 2006) and Cretaceous mantle xenoliths from the Salta rift system (Lucassen et al., 2005). (B) 208Pb/204Pb versus 206Pb/204Pb showing the same sample comparison. Symbols are the same as Figure 7 with the addition of coastal Jurassic volcanic and plutonic rocks (empty triangle and empty upside-down triangle, respectively) and the mantle xenoliths (empty crosses). Both plots show the same decoupling between Puna magmas and the Jurassic volcanic and plutonic samples. It is important to note that the Puna Plateau and Jurassic continental margin data sets represent the full spectrum of magma compositions, ranging from basalt to highly evolved samples. MORB—mid-ocean-ridge basalt; LM—lithospheric mantle.

0.5132

0.718 A B 0.5130

0.714 0.5128 Nd Sr 0.5126 144 86 0.710 Sr/ Nd/ 87 143 0.5124

0.706 0.5122

0.702 0.5120 024681012 024681012

MgO (wt%) MgO (wt%) Figure 9. (A) 87Sr/86Sr versus MgO (wt%) comparing Puna lava compositions to that of Jurassic volcanic and plutonic rocks (Lucassen et al., 2006). (B) 143Nd/144Nd versus MgO (wt%) showing the same sample comparison. Symbols are the same as in Figure 8. Both plots show a signiﬁ cant difference in isotopic composition between Jurassic igneous rocks and late Cenozoic Puna volcanics at similar MgO (wt%).

titanite, illmenite, and/or rutile. However, REE concentrations and pat- where (e.g., lithosphere-derived “western Great Basin” basalts discussed terns for basalts do not show supporting evidence for this hypothesis in Kempton et al., 1991; for review of physical and chemical properties (Fig. 6), such as the expected negative Dy anomalies if amphibole were of continental mantle, see Carlson et al., 2005) (Fig. 7). This observa- responsible for variation in the HFSE anomalies. Variable magma source tion includes our basalts and published data from mafi c samples (e.g., region trace-element characteristics are common for spatially extensive Kay et al., 1994), which have 87Sr/86Sr values greater than 0.70550 and monogenetic volcanic fi elds, and clarifi cation of the Ti-Nb-Ta anomaly in 143Nd/144Nd less than 0.51265. As pointed out by Kempton et al. (1991) the Puna Plateau data set warrants further study. and many others, it is impossible for magma from the depleted mantle Puna lavas have Sr and Nd isotopic compositions that are typical to obtain these Sr and Nd isotope values through crustal assimilation or for enriched continental mantle lithosphere in South America and else- assimilation and fractional crystallization (AFC) processes and maintain

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0.714 0.5132 A B 0.712 0.5130

0.710 0.5128 Nd Sr 144 86 0.708 0.5126 Sr/ Nd/ 87 143 0.706 0.5124

0.704 0.5122

0.702 0.5120 -71 -70 -69 -68 -67 -66 -71 -70 -69 -68 -67 -66

Longitude Longitude

20.0 40.5 C D 40.0

19.5

39.5 Pb Pb 204 19.0 204 39.0 Pb/ Pb/ 208 206

38.5 18.5

38.0

18.0 37.5 -71 -70 -69 -68 -67 -66 -71 -70 -69 -68 -67 -66

Longitude Longitude Figure 10. Plots of (A) 143Nd/144Nd, (B) 87Sr/86Sr, (C) 206Pb/204Pb, and (D) 208Pb/204Pb versus longitude for Puna Plateau volcanic samples (our samples and those from the literature) and coastal Jurassic extrusive and intrusive igneous rocks from Lucassen et al. (2006). Symbols are

the same as in Figure 8. Only samples with <57 wt% SiO2 were included in order to ﬁ lter out those with strong crustal isotopic signatures. All plots show decoupling between the coastal rocks and Puna Plateau volcanics, which suggests different mantle source region compositions.

a basalt major-element composition, particularly such high wt% MgO tion on the margin of the Gondwana foreland. It is generally agreed that and Ni and Cr concentrations. Further, Pb isotope values for Puna basalts the major phase of magmatism took place during the Early Ordovician are also radiogenic (Fig. 8) compared to what is expected for magmas (ca. 490–480 Ma) in both the south (Pankhurst et al., 1998, 2000) and the derived from a depleted mantle. Based on both the trace-element and north (Pankhurst et al., 2000; Kleine et al., 2004). Later intrusions of per- isotopic evidence, we suggest that mafi c Puna magmas originated from aluminous granitoids occurred until the mid-Ordovician. Preserved rocks an aged subcontinental lithospheric mantle that had also been previously are dominantly calc-alkaline and include gabbro, gabbronorite, norite, metasomatized. tonalite, monzonite, granodiorite, and granite, some with mafi c enclaves. Plutonic rocks of the Ordovician Famatinian arc are exposed through- We compare Sr and Nd isotopic data of Puna volcanics to the most out northwest Argentina. Geochemistry, geochronology, and examination recently acquired data for the Famatinian arc (Stair, 2008; Otamendi et al., of fi eld relations over the past 20 years have shed some light on the evolu- 2009) in Figure 7. Mafi c (gabbro and gabbronorite) Famatinian samples tion of this magmatic arc (Ramos et al., 1986; Dalla Salda et al., 1992; have initial Sr and Nd isotopic values nearly identical to mafi c lavas from Rapela et al., 1992; DeBari, 1994; Kleine et al., 2004; Pankhurst et al., the southern Puna Plateau. The limited Sr and Nd isotopic data in the lit- 1998; Pankhurst et al., 2000). In short, the arc was formed by plate subduc- erature, such as that from Kleine et al. (2004) for samples from the central

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and northern Puna Plateau, support this conclusion. Therefore, we suggest isotopic composition to MORB (Figs. 8 and 9) than an aged subcontinen- that Famatinian gabbro and Puna basalt were generated from the same tal lithospheric mantle, trending toward a depleted mantle at mafi c com- subcontinental lithospheric mantle magma source region. This implies positions. We also plotted isotopic ratios versus longitude for mafi c Puna that at least a thin portion of the subcontinental lithospheric mantle has Plateau volcanics and mafi c coastal Jurassic samples to illustrate the clear remained intact beneath NW Argentina during the past ~485 Ma. Radio- separation in Sr, Nd, and Pb isotopic values (Fig. 10). Thus, Figures 8, 9, genic growth of 87Sr and 143Nd over the ~500 Ma time period elapsed since and 10 show that the lithospheric magma source region for the Puna Pla- the formation of the Famatinian arc is minor, if concentrations typical of teau is chemically decoupled from the depleted mantle directly to the west the subcontinental lithospheric mantle are used for Rb, Sr, Sm, and Nd and east. This has been the case for hundreds of millions of years despite (Carlson et al., 2005). We conservatively estimate an ~0.0001 increase in long-term tectono-magmatic activity along the proto-Andean and Andean 87 86 ε the Sr/ Sr and ~0.5 Nd of the mantle over 500 Ma. margin and within the continental interior. Our silicic andesite and dacite samples, P07-ANT02 and P07-J201, respectively, and the more evolved samples from the literature have Sr and IMPLICATIONS FOR RECENT TECTONISM, MAGMATISM, AND Nd isotopic ratios that trend toward and overlap granodiorite samples in LITHOSPHERIC FOUNDERING Figure 7. The evolved Puna lavas have a strong elemental crustal signature. Because Puna lavas and Famatinian gabbro and intermediate rocks Elemental and isotopic data from Puna Plateau lavas indicate that produce an obvious mixing trend in Sr-Nd isotopic space (Stair, 2008), we parts of the subcontinental lithospheric mantle and presumably the lower suggest that the relationship between Famatinian mafi c and intermediate crust are in place beneath the plateau. Comparison to isotopic data from magmas was magma mixing during the Early Ordovician and the rela- Early Ordovician Famatinian plutonic rock suggests that the preserved tionship between evolved Puna lavas and Famatinian arc rocks is recent lithospheric mantle domain has been inherited from the Famatinian crustal assimilation involving contamination of a basaltic magma derived arc. However, the lithosphere is anomalously thin, and the lithospheric from the subcontinental lithospheric mantle. mantle has been reduced to <10 km, as evidenced by geophysical data Binary mixing between our most primitive basalt P07-SUR101 and (e.g., Tassara et al., 2006; McGlashan et al., 2008). Recent foundering of a high-87Sr/86Sr, low-143Nd/144Nd Famatinian granodiorite sample (CC1, the mantle lithosphere and potentially lowermost crust, whether through Fig. 7) produces a crustal assimilation line that explains the evolution of gravitationally driven convective removal or delamination, has been many Puna lavas. Puna basaltic andesite magmas have mantle:crust mix- proposed to explain the thinned lithosphere beneath the Puna Plateau, tures of ~80:20. More evolved andesite and dacite magmas show mixtures N-S extension within the plateau, and the late Miocene to present mafi c with greater amounts of granodiorite crust. This contamination may result volcanism (Kay and Kay, 1993; Kay et al., 1994; Whitman et al., 1996). from interaction of ascending basaltic magma with lower crust or a mid- This is supported by tomographic data, which suggest the presence of crustal melt layer (e.g., Yuan et al., 2000) that is at least partly composed a detached lithospheric block beneath the Puna Plateau (Schurr et al., of Famatinian granodiorite. Extremely evolved magmas are likely crustal 2006). This apparent confl ict between geochemical data and geophysical melts of Famatinian granodiorite, granite, or metamorphosed rock. and structural data requires an explanation. A Famatinian granodiorite crustal end member is not appropriate for One possible explanation is that the radiogenic Sr and Pb ratios and explaining the evolution of all Puna magmas. Therefore, we created a low Nd ratios in the Puna Plateau basalts could refl ect the signature of hybrid crustal end member that contains the lowest Sr (0.713478) and Nd crustal material emplaced along the base of the remaining lithosphere as (0.512038) isotopic ratios among Famatinian samples (CC2, Fig. 7). The the result of Neogene subduction erosion (Kay et al., 2008). The occur- hybrid end member better explains the evolution of some Puna magmas rence of this tectonic process has been discussed for the Chilean margin with relatively low 87Sr/86Sr. This isotopic composition could be part of (e.g., von Huene and Scholl, 1991; Kay et al., 2005). Subduction erosion the lowermost crust, possibly a more mafi c layer (thus the less radiogenic involves the removal of trench sediment and forearc crust from the over- 87Sr/86Sr) that was contaminated prior to Puna magmatism by ascending riding plate and introduction of this material into the subduction zone. melts with low 143Nd/144Nd compositions. For the Puna Plateau, the eroded component should be Jurassic igneous Primitive basalts have not erupted in the active Andean arc at latitudes material from the continental margin immediately to the west. Figures 8,

similar to the Puna Plateau. Only one volcanic sample with <57% SiO2 9, and 10 show isotopic data from Jurassic volcanic and plutonic rocks and >5% MgO has been analyzed for both Sr and Nd isotopes. This sam- located along the central South American margin at latitudes equivalent 87 86 ′ ′ ple has 56.2 wt% SiO2, 5.81 wt% MgO, and isotopic ratios of Sr/ Sr to the Puna Plateau (21º40 S–27º0 S) (Lucassen et al., 2006). Clearly, 0.70624 and 143Nd/144Nd 0.51247 (Matthews et al., 1994). Volcanic sam- mixing between these isotopic values and depleted mantle cannot pro- ples from the Puna Plateau with similar major-oxide chemistry tend to duce the isotopic ratios present in primitive lavas from the Puna area. have more radiogenic Sr isotope values and less radiogenic Nd isotope Therefore, although material from subduction erosion could be altering values. Assuming crustal assimilation or AFC magma evolution, this sug- the composition of the mantle wedge beneath the active Andean arc, it is gests that the magma source region composition beneath the active arc not entering the mantle beneath the Puna Plateau. is depleted compared to that for Puna magmas. However, arriving at a This brings us back to an explanation involving removal of the man- sound conclusion based on this inference is impossible due to lack of arc tle lithosphere. Geophysical and mass balance arguments (e.g., Beck and basalts in the region. We have therefore compiled published data from Zandt, 2002) make lithospheric foundering inescapable, and delamina-

Jurassic volcanic (47–70 wt% SiO2) and plutonic (43%–78% SiO2) rocks tion/convective removal of some form is favored in the area due to the from the continental margin to the west (Lucassen et al., 2006) and Cre- occurrence of maﬁ c volcanism with approximately OIB trace-element taceous mantle xenolith data from the Salta rift system to the east and characteristics and a rapid change in tectonic regime (Kay and Kay, southeast (Lucassen et al., 2005) for comparison to our Puna volcanic 1993; Kay et al., 1994). As discussed in the following paragraphs, how- samples (Figs. 8, 9, and 10). Figure 9 shows that the Mesozoic volcanic ever, this loss may have occurred partially or in a piecemeal fashion and plutonic rocks used here as a proxy for the active arc have lower rather than as rapid removal of the entire mantle lithosphere. 87Sr/86Sr and higher 143Nd/144Nd at a given wt% MgO. Coastal Jurassic Evidence against wholesale removal of the subcontinental lithospheric samples and mantle xenoliths from the Salta rift are more similar in mantle includes the lack of maﬁ c lavas erupted with asthenospheric isotopic

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compositions postdating the hypothesized major delamination event. If magmatic events cannot be ruled out. What can be ruled out, however, is either convective dripping or delamination had occurred, then the isotopic that no mantle-derived melt products with asthenospheric compositions composition of primitive basaltic volcanism should have a depleted mantle are found on the Puna Plateau. The striking similarity of isotopic ratios chemical signature, perhaps similar to that of mantle xenoliths from the between widespread Ordovician surface igneous rocks and all late Ceno- Salta rift system (e.g., Lucassen et al., 2005). Sudden removal of large zoic mantle-derived mafi c magmas further suggests that the mantle source amounts of lithospheric material via convective dripping should also be was removed from convection for at least ~500 Ma; the subplateau mantle accompanied by upwelling of asthenospheric mantle and relatively large that supplied Cenozoic Puna region volcanics must therefore be enriched, amounts of decompression-related melting. For example, if the thickness old, subcontinental lithospheric mantle. of the lithosphere was 120 km and sudden foundering of a spherical body Further, this domain has remained chemically distinct from the depleted of the lower lithosphere (60 km in diameter) were to occur, one would mantle. Recently proposed lithospheric foundering for the central Andean expect about a 15-km-thick melt column from decompression melting of Puna region (Kay et al., 1994) seems inescapable based on thermal, mass the asthenosphere replacing the space alone (using the parameterization of balance, and other constraints. However, we show here that only a portion McKenzie and Bickle, 1988). There is no indication of such large volumes of the subcontinental lithospheric mantle has been removed, and we infer of basalt either at the surface or at depth. Interestingly, the phenomenon of that lithospheric loss took place as a piecemeal or partial process rather only minor amounts of surface basalts derived from the lithosphere also than wholesale removal of a large, coherent lithospheric body. Finally, occurs in other places in which delamination has been proposed (Farmer et although subduction erosion may be contaminating the mantle wedge to al., 2002—Sierra Nevada, California; Manthei et al., 2009—Coast Moun- the west, it is clearly not affecting the chemical composition of the Puna tains Batholith, British Columbia and southeast Alaska). Plateau subcontinental lithospheric mantle domain. The issue of only minor volumes of surface basalt is diffi cult to address, as it is possible that a large volume of more mafi c magmas may ACKNOWLEDGMENTS be present at depth as intrusive bodies. The isotopic data are more helpful in providing an alternative explanation to large-scale lithospheric founder- This project was funded by National Science Foundation grant EAR- ing. The exclusively enriched isotopic composition of the existing mafi c 0635584 to L. Schoenbohm. M. Ducea acknowledges generous support product has two explanations: (1) the melting portion is the downgoing for this project from ExxonMobil’s COSA fund to the University of Ari- mantle lithosphere because it is hydrated (Elkins-Tanton, 2005), and/or zona. S. Drew thanks J.A. Palavecino for assistance in the fi eld during (2) the size of delaminating domains is small (e.g., 1 km or less), which sample collection. We wish to acknowledge K. Putirka and B. Cousens for in turn requires dripping to be slow, thus not allowing major upwelling helpful reviews that improved the manuscript considerably. and partial melting of asthenosphere. 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