Xenolith constraints on Altiplano architecture and assembly
Constraints on plateau architecture and assembly from
deep crustal xenoliths, northern Altiplano (SE Peru)
Alan D. Chapman1,2,†, Mihai N. Ducea1,3, Nadine McQuarrie4, Matthew Coble5, Lucian Petrescu3, and
Derek Hoffman1
Department of Geosciences, University of Arizona, Tucson, Arizona 85721, USA
Department of Geological Sciences and Engineering, Missouri University of Science and Technology, Rolla, Missouri 65409, USA
3
Facultatea de Geologie Geofizica, Universitatea Bucuresti, Strada N. Balcescu Nr 1, Bucuresti 010041, Romania
4
Department of Geology and Planetary Science, University of Pittsburgh, Pittsburgh, Pennsylvania 15260, USA
5
Department of Geological and Environmental Sciences, Stanford University, Stanford, California 94305, USA
1
2
ABSTRACT
Newly discovered xenoliths within Pliocene and Quaternary intermediate volcanic
rocks from southern Peru permit examination of lithospheric processes by which thick
crust (60–70 km) and high average elevations (3–4 km) resulted within the Altiplano,
the second most extensive orogenic plateau
on Earth. The most common petrographic
groups of xenoliths studied here are igneous or meta-igneous rocks with radiogenic
isotopic ratios consistent with recent derivation from asthenospheric mantle (87Sr/ 86Sr =
0.704–0.709, 143Nd/144Nd = 0.5126–0.5129). A
second group, consisting of felsic granulite
xenoliths exhibiting more radiogenic compositions (87Sr/ 86Sr = 0.711–0.782, 143Nd/144Nd =
0.5121–0.5126), is interpreted as supracrustal rocks that underwent metamorphism
at ~9 kbar (~30–35 km paleodepth, assuming a mean crustal density of 2.8 g/cm3) and
~750 °C. These rocks are correlated with
nonmetamorphosed rocks of the Mitu Group
and assigned a Mesozoic (Upper Triassic or
younger) age based on detrital zircon U-Pb
ages. A felsic granulite Sm-Nd garnet wholerock isochron of 42 ± 2 Ma demonstrates that
garnet growth took place in Eocene time.
Monazite grains associated with quenched
anatectic melt networks in the same rocks
yield ion microprobe U-Pb ages ranging from
3.2 ± 0.2 to 4.4 ± 0.3 Ma (2σ). These disparate
geochronologic data sets are reconciled by a
model wherein Mesozoic cover rocks were
transferred to >30 km depth beneath the plateau in the Eocene and progressively heated
until at least Pliocene time. Isothermal de†
Current address: Macalester College; chapman@
macalester.edu.
compression and partial melting ensued as
these rocks were entrained as xenoliths in volcanic host magmas and transported toward
the surface. Mafic granulites and peridotites
from the same xenolith suite comprise the
basement of the metasedimentary sequence,
exhibiting isotopic characteristics of Central
Andean crust. Calculated equilibrium pressures for these basement rocks are >11 kbar,
suggesting that the basement-cover interface
lies beneath the northernmost Altiplano at
~30–40 km below the surface. Together, these
results indicate that crustal thickening under
the northernmost Altiplano started earlier
than major latest Oligocene and Miocene
uplift episodes affecting the region and was
coeval with a flat slab–related regional episode of deformation. Total shortening must
have been at least 20% more than previous
estimates in order to satisfy the basement
to cover depth constraints provided by the
xenolith data. Sedimentary rocks at >30 km
paleodepth require that Andean basement
thrusts decapitated earlier Triassic normal
faults, trapping Paleozoic and Mesozoic
rocks below the main décollement. Magma
loading from intense Cenozoic plutonism
within the plateau probably played an addi­
tional role in transporting Mesozoic cover
rocks to >30 km and thickening the crust beneath the northern Altiplano.
INTRODUCTION
The formation of large orogenic plateaus
remains a central topic of research in continental tectonics. High-temperature (igneous and
metamorphic) petrology plays a large role in
answering critical questions such as: (1) When
and how did plateaus achieve their elevation
(Saylor­and Horton, 2014)? (2) How much
of the crustal thickening is due to shortening
(Gotberg et al., 2010) versus magmatic additions from the mantle­(Haschke and Günther,
2003)? (3) Where in the lithosphere and sublithospheric mantle did melting occur (e.g.,
Schilling et al., 2006)? (4) What role does the
presence of partial melt play in weakening the
midcrust and forming subdued plateau surface
relief (e.g., Clark and Royden, 2000; Gotberg
et al., 2010; Jamieson et al., 2011)? Much of
the tectonic research on the deep crustal processes under high plateaus is based on geophysical observations (e.g., Wigger­et al., 1994;
Myers et al., 1998; Graeber and Asch, 1999;
Beck and Zandt, 2002; Yuan et al., 2002; Zandt
et al., 2003). Indirect geologic evidence for the
high-temperature processes occurring at depth
is derived from the volcanic record sampled at
the surface (Kay et al., 1994; Davidson and de
Silva, 1995; Mamani et al., 2010). On rare occasions, lithospheric xenoliths are entrained in
volcanic rocks and transported to the surface
from beneath orogenic plateaus (e.g., Esperanca
et al., 1988; Smith et al., 1994; Davidson and de
Silva, 1995; Hacker et al., 2000; McLeod et al.,
2012, 2013). Such nodules provide direct constraints on lithospheric compositions beneath
the surface, the thermal state of the lithosphere,
the timing of crustal thickening and partial melting in the crust, as well as the origin of the material involved in the thickening of crust.
The Altiplano-Puna Plateau in the Central
Andes, one of the world’s largest modern plateaus, has been magmatically active since its
birth to present (Isacks, 1988; Allmendinger
et al., 1997). Magmatic rocks associated with
the Altiplano-Puna resemble compositions
found in subduction-related arcs, but they are
not confined to linear “arcs”; instead, they
are distributed across the width of the plateau.
Magma compositions range from basaltic to
GSA Bulletin; November/December 2015; v. 127; no. 11/12; p. 1777–1797; doi: 10.1130/B31206.1; 14 figures; 2 tables; Data Repository item 2015205;
published online 10 June 2015.
For permission to copy, contact editing@geosociety.org
Geological
Society of America Bulletin, v. 127, no. 11/121777
© 2015 Geological Society of America
Chapman et al.
rhyolitic (e.g., de Silva, 1989; Kay and Kay,
1993; Kay et al., 1994; Francis and Hawkesworth, 1994; Carlier et al., 2005; Mamani
et al., 2008, 2010; Drew et al., 2009), and the
volume of magmatic material ranges from
<0.01 km3, for example, in the volcanic centers
investigated here, to ~300 km3 at Cerro Galan,
one of the largest felsic calderas on the planet
(Francis­et al., 1983). The plateau also hosts the
Altiplano-Puna magma body, one of the largest
(>10,000 km3 in volume) zones of active partial
melting in a continental setting (Chmielowski
et al., 1999; Babeyko et al., 2002; Zandt et al.,
2003; Schilling et al., 2006).
Lower-crustal and deeper xenoliths are
extraordinarily rare in volcanic rocks from the
Altiplano-Puna, perhaps due to the thick crust
beneath the region (locally >70 km; e.g., Beck
and Zandt, 2002). To date, mantle peridotites
have not been described from the plateau region,
despite the presence of numerous mafic rocks
within various plateau volcanic fields (e.g., Kay
et al., 1994; Ducea et al., 2013). In this study, we
provide new major- and trace-element chemistry, Sr, Nd, and Pb whole-rock isotopic ratios,
U-Pb zircon and monazite geochronology, and
Sm-Nd garnet geochronology from six newly
discovered xenolith localities from the northern
part of the Altiplano in southern Peru. These
localities feature crustal and mantle xenoliths
hosted in small-volume Pliocene and Quaternary back-arc volcanic centers (Carlier and
Lorand, 1997; Carlier et al., 2005). The purpose of this investigation is: (1) to characterize
the rock types and estimate their distribution
within the sub-Altiplano lithospheric column; (2) to compare Peruvian xenoliths with
recently described crustal xenoliths from the
Bolivian Altiplano (Davidson and de Silva,
1995; McLeod et al., 2012, 2013); (3) to place
geochemical and geochronologic constraints on
how the crustal structure beneath the northern
Altiplano has evolved since the onset of thickening; and (4) to evaluate the relative importance
of magmatic additions versus tectonic shortening in thickening the crust beneath the northern
Altiplano.
BACKGROUND
Regional Geology
Southern Peru (13°S–16°S latitude) is located
above a subduction zone in the Central Andes
and is composed of, from west to east, a deep
trench, a forearc, a frontal magmatic arc (the
Western Cordillera), a high, internally drained
plateau (the Altiplano), and a foreland foldand-thrust belt (the Eastern Cordillera; Isacks,
1988; Allmendinger et al., 1997). Crustal thickness increases from 30 km in the forearc to over
1778
70 km under the Cordilleras and the Altiplano,
and decreases eastward beneath the sub-Andean
foreland and the craton to ~35–40 km (James
and Sacks, 1999; Beck and Zandt, 2002; Phillips and Clayton, 2014). Average elevations are
above 4 km in the Western Cordillera and the
Altiplano.
The Central Andes have resided in the upper
plate of an active convergent margin for much
of the time following the breakup of Rodinia
(Ramos, 2008, and references therein). Over
the course of this protracted convergent history, the locus of calc-alkaline arc magmatism
has shifted several times due to either changes
in slab dip and/or migrations of the subduction
trench relative to the upper plate (e.g., Mamani
et al., 2010). Arc magmas were emplaced into
a framework comprising: (1) ca. 1.0–1.5 Ga
crust intruded by younger Paleozoic (0.4–
0.55 Ga) arcs along the southwestern edge
of the Amazon craton (Tassinari et al., 2000;
Loewy et al., 2004), (2) the Paleo­protero­zoic
to early Paleozoic Arequipa terrane within
the modern forearc (Wasteneys et al., 1995;
Loewy et al., 2004; Ramos, 2008; Casquet
et al., 2010), and (3) ~5–7 km of Paleozoic
low-grade metamorphic rocks and another
~5 km of Mesozoic and Tertiary cover rocks
(Gotberg et al., 2010; Reimann et al., 2010;
Bahlburg et al., 2011; Reitsma, 2012). The
boundary between the Amazonia and Arequipa
domains is obscured by Paleozoic and younger
cover rocks and various magmatic arcs. These
rocks include Carboniferous to Early Jurassic
“Gondwanide” granitoids and related Triassic rift basin deposits­of the Mitu Group, both
produced during the assembly and breakup
of western Gondwana, (Sempere et al., 2002;
Ramos, 2008; Mišković et al., 2009; Reitsma,
2012; Perez and H
­ orton, 2014).
Subduction-related magmatism started
regionally during the Carboniferous and has
been more or less continuous until today
(Sande­man et al., 1995; Mamani et al., 2010).
Several arc domains are recognized, based on
composition, age, and location, in southern
Peru (Fig. 1): the Chocolate arc (301–90 Ma),
the Toquepala arc (90–45 Ma), the Anta arc
(45–30 Ma), the Tacaza arc (30–24 Ma), the
Huaylillas arc (24–10 Ma), the lower Barroso
arc (10–3 Ma), the upper Barroso arc (3–1 Ma),
and the active volcanic arc (<1 Ma). Of these,
the Chocolate arc consists of submarine volcanics and volcaniclastic deposits, whereas all
subsequent arcs starting with the Toquepala arc
at ca. 90 Ma were formed in a subaerial environment (Mamani et al., 2008, 2010). The Chocolate arc consists primarily of basaltic rocks, with
a thickness of 1 to >3 km, that overlie Paleozoic
passive-margin (meta)sedimentary rocks (Pino
et al., 2004). Calc-alkaline, dominantly mafic
magmatism continued well into the Cretaceous
over a >100 km width of the margin throughout
the Central Andes, suggesting that these areas
represent frontal arc, in contrast to back-arc,
magmatic products within an attenuated upper
plate (Oliveros et al., 2006; Charrier et al., 2007;
Rossel et al., 2013). The Late Cretaceous marks
a significant change in the character of magmatism, which switched from extension-related
and mafic-dominated volcanic and intrusive
rocks to a more typical Cordilleran intermediate
and silicic arc (Mukasa, 1986). This also corresponds to a well-known shift in the sedimentary
record from marine to continental deposition
(Sempere et al., 2002). Together, these observations suggest that the upper plate (South
America) changed from an extensional to compressional stress regime, marking the onset of
crustal thickening at ca. 90 Ma.
Volcanic rocks provide additional support
for significant Cenozoic crustal thickening in
the Central Andes. Geochemical proxies for
crustal thickness (e.g., Sr/Y, La/Yb, La/Sm,
Sm/Yb, and Dy/Yb) increase with time in these
rocks, particularly between 30 Ma and today
(Mamani et al., 2010). The pattern of crustal
thickening is complicated, however, as these
arcs migrated during this time period. The
Toquepala arc, the first continental arc sensu
stricto, is manifested in the geologic record
by the emplacement of the Coastal Peruvian
batholith and related ignimbrite activity (Pitcher
et al., 1985; Mukasa, 1986; Clark et al., 1990;
Martínez and Cervantes, 2003). The Toquepala arc activity waned at ca. 50 Ma (Mamani
et al., 2010), and the arc migrated inboard by
>150 km. The 45–30 Ma Anta arc consists of
the Andahuaylas-Yauri batholith and the Anta
volcanic rocks. These rocks are interpreted as
products of an inboard magmatic sweep, probably related to shallowing of the subducting
slab (James and Sacks, 1999). Initiation of the
Anta arc also corresponds to the Incaic shortening event that began at 45 Ma (Noble et al.,
1979; Sandeman et al., 1997; Roperch et al.,
2006) and is thought to represent an episode of
flat slab–related crustal thickening. However,
substantial crustal thickening did not begin until
the mid-Oligocene (ca. 30 Ma) and was accompanied by formation of the Tacaza, Huaylillas,
and lower Barroso and upper Barroso arcs and
strato­vol­canoes of the modern arc front. The
volcanic products of these arcs record a progressive thickening of the crust regionally (Mamani
et al., 2010) and are overall more silicic than
older arc rocks in the area.
While most of the Quaternary arc products
are located in the frontal arc and are defined by
a linear array of stratovolcanoes (de Silva, 1989),
Geological Society of America Bulletin, v. 127, no. 11/12
Xenolith constraints on Altiplano architecture and assembly
Figure 1. Geologic map modified after Mapa
Geológico del Perú (1975) showing locations
of Rumicolca (black stars) and Puno suite
(white stars) exposures, and sampled xenolith localities (H—Huarocondo; N—Tipon;
O—Oropesa; P—Puno; R—Raqchi; Y—
Taray). Cuzco-Vilcanota fault system runs
from Cuzco to Lake Titicaca. Triassic Mitu
Group clastic sediments are shown in light
gray with gravel pattern. Other abbreviations: Cz—Cenozoic; Mz—Mesozoic; Pz—
Paleozoic. Color version of figure available
in the GSA Data Repository (see text footnote 1).
small-volume potassic volcanic rocks crop out
in the back arc within the northernmost Altiplano and appear on regional geologic maps as
the Quaternary Rumicolca formation (Carlier
et al., 1997, 2005; Carlotto et al., 2011). They
are high-K calc-alkaline lava flows and domes,
ranging in composition from basalt to rhyolite,
which erupted along the NW-striking CuzcoVilcanota fault system. The slip history of the
Cuzco-Vilcanota fault system is poorly understood, most likely originating as a pre-Andean
suture zone between two distinct lithospheric
blocks (Carlier et al., 2005); it appears to
have been most recently active as a left-lateral
transtensional fault, based on surface morphology along the fault (field observations; Fig. 2;
e.g., Sebrier et al., 1985). The volcanic rocks
(described in more detail below) that crop out
along the Cuzco-Vilcanota fault system host
crustal and rare mantle xenoliths.
Xenolith Localities
This study focuses on six xenolith localities discovered during the 2010 and 2011 field
seasons (Figs. 1 and 3). Xenolith-bearing vol­
canic rocks are concentrated in two areas: (1) in
flows of Quaternary age found mostly along the
Cuzco-Vilcanota fault system, near Cuzco, and
(2) in Pliocene lava flows near Puno, along the
western shore of Lake Titicaca (Fig. 1). Both
the Cuzco and Puno suites are high-K calcalkaline lavas that range from basalt to dacite.
On average, the Puno flows are less silicic than
the Cuzco field, although both suites have typical calc-alkaline major-element trends (Carlier
et al., 2005; Chapman and Ducea, 2013). Below
we provide a general summary of the petrog­
raphy and geochemistry of xenolith-bearing
volcanic rocks in southern Peru.
The Cuzco volcanic suite appears on geologic
maps of Peru as the Quaternary Rumi­colca formation (e.g., Carlotto et al., 2011), which is
composed of high-K intermediate material,
referred to as shoshonitic rocks (Carlier et al.,
2005; Mamani et al., 2010). The Rumicolca
formation crops out as lava flows and domal
structures less than 1.5 km in diameter scattered along the Cuzco-Vilcanota fault system
(Figs. 1 and 2; Carlier et al., 2005). In detail,
exposures of the Rumicolca formation contain
a variety of compositions from basaltic ande­
site to dacite and are best classified as lampro­
phyres with K2O/Na2O = 1–2. These flows
contain a ubiquitous phenocryst assemblage of
biotite, amphibole, clinopyroxene, Fe-Ti oxides,
and moderately to strongly aligned plagioclase
laths, with subordinate sanidine, orthopyroxene,
and apatite phenocrysts (Fig. 4A). Nepheline
phenocrysts are present in one locality (Raqchi­,
described later herein) in samples lacking
orthopyroxene. Cuzco suite lavas are extremely
enriched in incompatible trace elements, lack
Geological Society of America Bulletin, v. 127, no. 11/121779
Chapman et al.
72˚ W
Y
H
Cuzco
O
FS
CV
14˚ S
N
50 km
R
Cuzco basin
fau
achay
m
o
b
m
Ta
lt
Figure 2. Field view of the Tambomachay fault, a member of the Cuzco-Vilcanota fault system, seen toward the
west from near point O on Figure 1. Field of view along range front is ~5 km wide. Inset shows location of CuzcoVilcanota fault system (CVFS) relative to exposures of Rumicolca back-arc volcanic rocks (shown in black). Other
inset abbreviations are as in Figure 1.
Eu anomalies among rare earth elements
(REEs), and have high Sr/Y (Chapman and
Ducea, 2013). Carlier et al. (2005) interpreted
these rocks as partial melts of clinopyroxenerich veins in the uppermost mantle. Alternatively, they could be small-degree partial melts
of the lowermost crust beneath the Altiplano,
which is dominated by clinopyroxene-rich arc
cumulates with or without amphibole and garnet. These rocks yield negative eNd values and Sri
> 0.705, indicating a lithospheric origin (Carlier
et al., 2005). Cuzco suite xenoliths, consisting
of (in order of decreasing abundance) gabbros
and diorites, granulites, clinopyroxenites, and
peridotites, were recovered from Huarocondo,
Oropesa, Raqchi, Taray, and Tipon (Fig. 1;
Table 1). Of these localities, the flow east of
Oropesa contains the highest xenolith abundance and the most variety in the composition
and size of xenoliths (up to 10 cm). Most other
localities contain smaller (<3 cm) nodules.
Granulite xenoliths reported from a Pliocene–
Quaternary lamproite dike ~15 km northeast
of Cuzco (Carlier and Lorand, 1997) were not
investigated here.
A second suite is represented by more mafic
trachyandesites and basaltic trachyandesites
1780
in Pliocene lava flows (Mamani et al., 2010)
exposed ~15 km southeast of Puno (Fig. 1).
These rocks are part of the western Altiplano
volcanic suite of Carlier et al. (2005) and were
classified as shoshonites by those authors. Puno
lavas are best classified as high-K calc-alkaline basanites with K2O/Na2O ratios between
1 and 2. Phenocryst assemblages found in the
Puno suite are distinct from those found in
the Cuzco suite, consisting of plagioclase laths,
clino­pyroxene, olivine, nepheline, and opaques,
and lacking biotite and amphibole. The Puno
volcanic suite yields lower eNd (<–3) and higher
Sri (>0.706) baseline values than the Cuzco
suite (Carlier et al., 2005), also suggesting an
origin by partial melting of the South American lithosphere (see following). The Puno suite
is also characterized by elevated Sr/Y (>40)
and La/Yb (>20) values with no Eu anomalies
(Carlier et al., 2005), suggesting a source rich
in clinopyroxene and garnet, but not in feldspar. Small (<5 cm) fresh xenoliths of clinopyroxenite and subordinate weathered granitic
assemblages were recovered from a single Pliocene–Pleistocene flow in the Umayo volcanic
complex (Kaneoka and Guevara, 1984) of the
Puno suite (Table 1).
RESULTS
Details of sample preparation and the analytical techniques used for whole-rock (major
and trace elements, and Sr-Nd-Pb isotopes) and
mineral geochronology (garnet Sm-Nd, zircon
U-Pb, and monazite U-Th-Pb) are provided in
the GSA Data Repository.1
Xenolith Petrography
Xenoliths recovered from the Cuzco suite
are divided into five categories: spinel-bearing
perido­tites (Fig. 4B), clinopyroxenites (Fig. 4C),
gabbros and diorites (Fig. 4D), mafic granulites
(Fig. 4E), and felsic peraluminous granulites
(Figs. 4F, 4G, and 4H). The main conclusions of
this study are based on data from the lattermost
group, but petrographic descriptions of all rock
types are provided here, as it is important to note
the relative abundance of igneous (peridotites
1
GSA Data Repository item 2015205, analytical
techniques, images of garnet mineral separates, and
data tables, is available at http://​www​.geosociety​
.org​/pubs​/ft2015​.htm or by request to editing@​
geosociety​.org.
Geological Society of America Bulletin, v. 127, no. 11/12
Xenolith constraints on Altiplano architecture and assembly
A
B
Figure 3. Field photographs showing xenoliths from the lamprophyric Rumicolca formation. Coin is 22 mm across: (A) diorite xenolith exhibiting primary foliation; (B) weathered
surface of peridotite xenolith. Color version of figure available in the GSA Data Repository
(see text footnote 1).
and gabbros) versus metasedimentary (clino­
pyroxenite skarns and both mafic and felsic granulites) xenolith types. For each lithology, average
modal mineralogical abundances are given in the
GSA Data Repository (see footnote 1), and key
petrographic features are shown in Figure 4.
Two fresh peridotite inclusions, both ~1 cm
in diameter, were collected from a single location in the Huarocondo flow. These nodules are
mineralogically and texturally homo­
geneous,
exhibiting a medium- to coarse-grained xeno-
morphic granular texture with common 120°
grain boundary triple junctions (Fig. 4B)
and containing mainly olivine (Fo89-Fo92),
orthopyroxene, minor (<3% modal abundance)
primary clinopyroxene and chromite-spinel, and
secondary (presumably metasomatic in origin)
phlogopite and vein calcite and quartz.
Gabbro and diorite xenoliths constitute a
second, more abundant and widespread, igneous xenolith group recovered from the Huarocondo, Oropesa, Taray, and Tipon flows of the
Cuzco suite. Plagioclase, and hornblende and/or
clinopyroxene are the predominant minerals
in these rocks, which contain varying amounts
of orthopyroxene, olivine, biotite, K-feldspar,
titanite, Fe-Ti oxides, apatite, quartz, rutile, and
magmatic epidote. Gabbro and diorite xenoliths
lacking biotite are medium to coarse grained
with allotriomorphic granular textures, whereas
biotite-rich varieties are medium grained with
well-developed igneous foliation. Biotiteand K-feldspar–rich diorites probably represent fragments of a gradational contact zone
between gabbro/diorite and metasedimentary
xenolith types, described later. Quenched glass,
ubiquitous in metasedimentary xenoliths, is not
observed in gabbro and diorite nodules.
Clinopyroxene-rich assemblages were studied from the Cuzco suite (Oropesa and Raqchi
flows) and in the Puno suite. Clinopyroxenites
from the Oropesa flow contain 40%–90% diopsidic clinopyroxene, with plagioclase and quartz
constituting the remainder of the major phases
present, and with minor amounts (in order of
decreasing abundance) of Fe-Ti oxide, kelyphitized grossular garnet, biotite, K-feldspar, calcite, titanite, apatite, and zircon, likely reflecting
a metasedimentary (calc-silicate skarn) origin.
Puno suite clinopyroxenites generally have
a higher proportion of clinopyroxene (84%–
92%), also diopsidic in composition, with
minor amounts of Fe-Ti oxides, plagioclase,
and quartz. All clinopyroxenite xenoliths are
coarse grained with weakly to moderately foliated fabrics.
Mafic granulite xenoliths were recovered
from the Oropesa and Raqchi flows in the Cuzco
suite and are distinguished from the clinopyrox­
enite suite based on less abundant clinopy­
roxene, and a higher proportion of ortho­pyrox­
ene, plagioclase, biotite, and K-feldspar. Minor
phases of the granulite suite include olivine­
(< 1%), hornblende, garnet, titanite, Fe-Ti
oxides, apatite, and up to 1% quenched felsic
melt phase along grain boundaries. No quartz
was detected in these nodules. Mafic granulite
xenoliths are coarse grained and show welldeveloped gneissic foliation with respect to
the clinopyroxenite suite, defined by aligned
biotite-rich layers.
A more felsic suite of granulite xenoliths
occurs as inclusions in the Oropesa flow, dominated by plagioclase, K-feldspar, and quartz
(generally 80%–90% of studied assemblages),
with lesser amounts of biotite, garnet, kyanite,
clinopyroxene, titanite, Fe-Ti oxides, apatite,
and zircon. Garnet-bearing granulites reported
from ~15 km northeast of Cuzco (Carlier and
Lorand, 1997) probably represent correlative
assemblages. This suite of xenoliths exhibits anastomosing networks of rhyolitic glass
Geological Society of America Bulletin, v. 127, no. 11/121781
Chapman et al.
Figure 4. Photographs of
structural and petrologic features in xenoliths and volcanic
host rocks. (A) Lamprophyre
showing biotite, clinopyroxene, opaques, and plagioclase
(~20-µm-long laths, flow aligned
parallel to long dimension of
photograph) phenocrysts set in
a glassy matrix, Oropesa flow;
crossed-polarized­light (xpl).
(B) Spinel harzburgite (~10–
80 µm spinel grains not visi­ble at
magnification of image­), Huaro­
condo flow; xpl. (C) Foli­
ated
clinopyroxenite, Puno flow; xpl.
(D) Foliated hornblende-clinopyroxene diorite, Oropesa flow;
xpl. (E) Mafic granulite, Raqchi
flow; xpl. (F) Garnet-bearing
felsic granulite, Oropesa flow;
plane-polarized light (ppl).
(G) Felsic granulite showing
evidence for biotite breakdown
by dehydration melting. Note
melt pockets in textural equilibrium with monazite along
biotite cleavage, Oropesa flow;
backscattered electron image.
(H) Felsic granulite showing
neocrystallization of monazite,
spinel, Fe-Ti oxides, ortho­
pyroxene, orthoamphibole, and
z i rc o n f ro m a p o c k e t o f
quenched melt. Note radiating needles of orthoamphibole,
Oropesa flow; backscattered
electron image. Mineral abbreviations: Bt—biotite; Cpx—
clinopyroxene; Ged—gedrite;
Grt—garnet; Hbl—hornblende;
Hc—hercynite; k—kelyphite;
Kfs—K-feldspar; Mag—mag­
netite; Mnz—monazite; Plag—
plagioclase; Ol—olivine; Opx—
orthopyroxene; Qtz—quartz;
Zrn—zircon.
1782
A
B
C
D
E
F
G
H
Geological Society of America Bulletin, v. 127, no. 11/12
Xenolith constraints on Altiplano architecture and assembly
TABLE 1. XENOLITH LOCALITIES AND HOST ROCK PETROGRAPHY
Locality
Location
Age of flow*
Lithology†
Phenocrysts§
Xenoliths present#
Huarocondo
13.42871°S, 72.20319°W
Q
d; lamp
pl, san, cpx, hb, bt, opq
hz, gbd
Oropesa
13.59760°S, 71.74499°W
<0.7 Ma1
b-a; lamp
bt, hb, cpx, pl, opq, cal, ap
gbd, fgr, cpx, mgr, qz
1
b-a; bsnt
pl, cpx, ol, nph, opq
cpx
Puno
15.88321°S, 69.90642°W
5.6 Ma
b-a; teph
cpx, opx, pl, nph, opq
mgr, cpx
Raqchi
14.15833°S, 71.38109°W
<0.03 Ma1
Taray
13.42040°S, 71.88109°W
Q
b-a; lamp
bt, hb, cpx, pl, opq, ap
gbd
Tipon
13.56919°S, 71.78333°W
Q
b-a; lamp
bt, pl, cpx, hb, ap, opq
gbd
*Quaternary (Q; age not available from studied exposure) adjacent flows younger than 2 Ma (Carlier et al., 2005); 1—Kaneoka and Guevara (1984).
†
Lithologic abbreviations: b-a—basaltic andesite; bsnt—basanite; d—dacite; lamp—lamprophyre; teph—tephrite.
§
Mineral abbreviations: ap—apatite; bt—biotite; cal—calcite; cpx—clinopyroxene; hb—hornblende; nph—nepheline; ol—olivine; opx—orthopyroxene; pl—plagioclase;
opq—Fe-Ti oxide; san—sanidine.
#
Xenolith abbreviations: cpx—clinopyroxenite; fgr—felsic granulite; hz—harzburgite; gbd—gabbro/diorite; mgr—mafic granulite; qz—quartzite.
(average wt% Na2O = 1.9, K2O = 6.16, SiO2 =
72.6, and Al2O3 = 13.6) along grain boundaries.
These quenched melts are most voluminous in
areas within and surrounding biotite grains. A
secondary mineral assemblage of orthopyroxene, orthoamphibole (gedrite), hercynitic spinel,
corundum, and monazite (Figs. 4F, 4G, and 4H)
is associated with quenched melts. Felsic granulite xenoliths are also referred to here as para­
gneisses, given the ubiquitous gneissose textures
and the likely pelitic protolith for these rocks.
The ~1–2 cm average diameter of these xenoliths is comparable to the spacing of alternating
quartzofeldspathic leucosomes and biotite ±
garnet-rich melanosomes; as a result, xenoliths
of leucocratic and melanocratic material may
be misinterpreted as having been derived from
different rocks. Garnet grains are ~0.5–2 mm
diameter, generally inclusion-poor, subidioblastic porphyroblasts enclosing mainly apatite and
quartz grains <100 µm in diameter. Kelyphite
coatings up to 50 µm thick containing mainly
glass, ilmenite, and spinel are present on ~10%
of garnet porphyroblasts studied here and
appear to be restricted to contacts with biotite
and/or melt (Figs. 4 and 5).
Thermobarometry and
Petrogenetic Modeling
Representative mineral compositions used in
thermobarometric calculations are reported in
the GSA Data Repository (see footnote 1). Equilibrium pressure-temperature (P-T ) conditions
(Table 2) were calculated for garnet-bearing felsic granulite xenoliths using the internally consistent average P-T mode in THERMOCALC,
version 3.26, assuming a water activity of unity,
due to the abundance of hydrous phases (Fig. 6;
Powell and Holland, 1994; Holland and Powell,
1998). The P-T conditions for felsic granulite
xenoliths were calculated from garnet + ­kyanite +
plagioclase + quartz assemblages interpreted,
based on chemical and textural evidence, to have
equilibrated during peak metamorphism. Garnet
grains from these rocks are rich in almandine,
with lesser amounts of pyrope, spessartine, and
grossular components (Xalm ~ 0.69, Xprp ~ 0.17,
Xsps ~ 0.09, and Xgrs ~ 0.05; Fig. 5; see also GSA
Data Repository material [footnote 1]), and are
compositionally homogeneous with flat interior
profiles (Fig. 5). Calculated temperatures and
pressures and 1s uncertainties for felsic para­
gneiss xenoliths are ~760 ± 45 °C at 9.1 ± 1.4–
9.5 ± 1.3 kbar (Table 2).
The Gibbs free energy minimization software package THERIAK-DOMINO (de Capitani and Brown, 1987) and the thermodynamic
end-member and solution models of the accompanying “tcdb55c2d” database (THERMOCALC database as distributed in version 3.30)
were used to construct a P-T pseudosection
for the bulk composition of sample 11MA20A
(Fig. 7). The peak mineral assemblage of garnet + plagio­clase + K-feldspar + biotite + quartz
present in this sample is quite common in other
samples of the felsic granulite suite and is consistent with those predicted from the pseudosection. The P-T results from THERMOCALC plot
within the field of the peak assemblage and are
in close proximity to the biotite-consuming and
orthopyroxene-producing dehydration melting
reaction (Fig. 7).
Diorite xenoliths containing the equilibrium assemblage hornblende + plagioclase +
K-feldspar + quartz + titanite + Fe-Ti oxide
were analyzed for aluminum-in-hornblende
(Al-in-hbl) igneous barometry (Hammarstrom
and Zen, 1986; Hollister et al., 1987; Johnson
and Rutherford, 1989; Schmidt, 1992; Anderson and Smith, 1995; Ague, 1997). We report
hornblende-plagioclase temperatures as well
as Al-in-hbl pressures for calibrations based
on Hammarstrom and Zen (1986), Hollister
et al. (1987), Johnson and Rutherford (1989),
Schmidt (1992), and Anderson and Smith
(1995) in the GSA Data Repository (see footnote 1; Table 2). Diorite xenoliths yield Anderson and Smith (1995) temperatures and pressures ranging from 710 °C to 839 °C and from
6.4 to 9.2 kbar, respectively.
Two samples of mafic granulite yielded significantly higher P-T estimates than either felsic
paragneiss or diorite xenolith suites, ranging
from 978 °C to 1377 °C and from 10.9 to 11.7
kbar. These estimates are based on two-pyrox-
ene thermometry (Wood and Banno, 1973) and
clinopyroxene-plagioclase- quartz barometry
(Ellis, 1980).
Whole-Rock Major- and TraceElement Geochemistry
Felsic granulite xenoliths are the most silicic
assemblages studied here, with SiO2 ranging
from 59.51 to 66.51 wt% (Fig. 8A [see footnote 1 for data]). One sample containing abundant garnet (16.4 modal%) and biotite (28.5
modal%), most likely representing a nodule
of melanosome material, yielded lower silica
values of 49.59 wt% (see footnote 1). Relative to the other xenoliths studied here, this
suite of xenoliths shows: (1) the highest weight
percent values of Al2O3 (16.66–20.60), Na2O
(2.13 – 5.19), and K2O (3.94–7.09) and the
lowest MgO (0.44–4.01), FeO* (1.20–13.65),
and CaO (1.58–2.75); (2) the most pronounced
enrichments in fluid-mobile large ion lithophile
elements (LILEs) relative to high field strength
elements (HFSEs), with notable trace-element
peaks at K and Pb and depletion in Nb, Zr,
and Ti, although the Zr deficit may be due to
incomplete dissolution of zircon during sample
preparation; and (3) negative Eu anomalies and
heavy (H) REE depletion in samples with relatively low modal plagioclase (<30%) and garnet
(<5%), respectively (Fig. 8B [see footnote 1]).
Compositionally, leucogneiss xenoliths are
similar to the average composition of the North
American shale composite (Gromet et al., 1984);
they are modally similar to a granodiorite.
Diorite and gabbro xenoliths are considerably more mafic than felsic granulite nodules, with lower weight percent values of SiO2
(42.24–52.49) and K2O (0.48–3.08), higher
values of MgO (2.43–9.92), FeO* (4.48–15.51),
and CaO (8.04–14.28), and comparable ranges
of Al2O3 (10.88–20.29) and Na2O (1.68–4.43).
With respect to felsic granulite xenoliths, diorite and gabbro xenoliths are depleted in LILEs
and show similarly variable HFSE concentrations (Fig. 8B). Hornblende-rich xenoliths show
concave-downward middle rare earth element
(MREE) trends with maxima between Gd and
Geological Society of America Bulletin, v. 127, no. 11/121783
Chapman et al.
Figure 5. (A) X-ray maps and (B) a quantitative zonation profile in garnet from felsic granulite sample 11MA20A. Location
of detailed garnet X-ray maps is shown in
thin section Mg map. Location of monazite
in Figure 12 is shown as black star in thin
section Mg map. Location of quantitative
profile is shown in Ca map. Bright areas in
X‑ray images represent relatively high concentration. Mineral abbreviations are as in
Figure 4. Ap—apatite; Ky—kyanite.
Ho, in contrast to clinopyroxene-rich gabbroic
xenoliths, which show monotonically decreasing light (L) REE to HREE patterns. Samples
lacking significant clinopyroxene and containing relatively low modal plagioclase (<30%)
and relatively high modal hornblende (>60%)
show positive Eu anomalies; samples containing
abundant clinopyroxene (>20%) show negative
Eu anomalies; and remaining diorite and gabbro
xenoliths lack Eu anomalies.
Clinopyroxenite xenoliths show a wide range
of silica values, with two distinct clusters from
41.57 to 55.66 and from 69.33 to 76.13 wt%
SiO2. The large variation in silica values derives
from the former of the two clusters being clinopyroxene-dominated assemblages, whereas the
latter of the two clusters contain >40% modal
quartz. These xenoliths have relatively low
Al2O3 (1.53–7.40), K2O (0.03–0.51), and Na2O
(0.38–1.97), high values of MgO (6.94–16.68)
and CaO (10.39–21.46), and a wide range of
FeO* (2.19–18.68). This suite is depleted in
LILEs, HFSEs, and REEs relative to diorite/
gabbro and felsic granulite xenoliths. Two distinct REE patterns are apparent in this xenolith
suite: (1) Nodules containing quartz and the
lowest clinopyroxene abundances exhibit negative Eu anomalies and La/Yb values of 4–15,
and (2) clinopyroxene-rich (>80%) assemblages
yield La/Yb of <5 and show concave-downward
patterns with maxima between Nd and Sm and
lack significant Eu troughs.
Mafic granulite xenoliths generally overlap the major-element compositions of diorite/
gabbro­xenoliths but give lower CaO (4.18–
8.13) and Na2O (1.56–2.57) and higher K2O
(2.92–7.28) values. These assemblages display
similarly “spiky” trace-element patterns to those
of the felsic granulite suite and are depleted in
HREEs relative to LREEs (La/Yb = 7–62).
Peridotite xenoliths are the most mafic of
the assemblages studied here, with wt% SiO2
of 40.90–45.52. These xenoliths show the
lowest weight percent oxide values of Al2O3
(0.59–2.14) and CaO (0.03–1.24), with negligible Na2O and K2O, the highest values of
1784
A
B
Geological Society of America Bulletin, v. 127, no. 11/12
Xenolith constraints on Altiplano architecture and assembly
TABLE 2. SUMMARY OF THERMOBAROMETRIC CALCULATIONS FROM PERUVIAN XENOLITHS
Sample
Lithology*
Mineralogy*
T (°C)†
P (kbar)†
Corr§
11MA20A
grt-lg
grt ky pl qtz
760 ± 45
9.5 ± 1.3
0.491
0.461
11MA20H
grt-spl-lg
grt ky pl qtz
758 ± 45
9.1 ± 1.4
11MA2A
pxite/chkt
pl opx cpx qtz
1118 ± 35
11.7 ± 0.6
NA
11MA6C
pxite/chkt
pl opx cpx qtz
1081 ± 30
10.9 ± 0.6
NA
10MA7
bt di/gb
hb pl bt kfs spn opq qtz
760 ± 41
7.8 ± 0.6
NA
11MA19G3
hb-cpx gb cum
hb pl kfs opq qtz
787 ± 41
7.8 ± 0.6
NA
11MA18B
hb-cpx gb cum
hb pl kfs opq qtz
719 ± 42
7.8 ± 0.6
NA
11MA19F
hb-cpx gb cum
hb pl kfs opq qtz
781 ± 40
6.5 ± 0.6
NA
11MA21
px-hb gbn
hb pl bt kfs opq qtz
817 ± 41
9.2 ± 0.6
NA
11MA20D
di/gb
hb pl kfs opq qtz
772 ± 40
6.9 ± 0.6
NA
11MA20Em
hb-di
hb pl bt kfs opq qtz
836 ± 40
6.4 ± 0.6
NA
*Mineral and rock type abbreviations: bt—biotite; chkt—charnockite; cpx—clinopyroxene; cum—cumulate; di—diorite; gb—gabbro; gbn—gabbronorite; grt—garnet;
hb—hornblende; kfs—K-feldspar; ky—kyanite; lg—leucogneiss; opq—Fe-Ti oxides; opx—orthopyroxene; pl—plagioclase; px—two pyroxenes; pxite—pyroxenite; qtz—
quartz; spl—spinel; spn—titanite.
†
1σ uncertainties for samples 11MA20A and 11MA20H based on propagation of uncertainties on thermodynamic data and activity-composition relationships through
thermobarometric calculations. Uncertainties reported for remaining samples are standard deviations in calculated temperatures and pressures for multiple mineral pairs.
§
Correlation coefficient between pressure and temperature from THERMOCALC.
MgO (42.99–46.23), and intermediate FeO*
(7.92–10.34). These xenoliths show prominent
Nb troughs and Pb peaks and LILE to HFSE
depletion in trace-element variation diagrams,
and they show slight enrichment in LREEs with
respect to MREEs and HREEs (La/Yb ~ 5) and
small negative Eu anomalies.
Whole-Rock Isotope Data
Sr and Nd isotopic data from 27 xenoliths and
Pb isotopic data from 29 xenoliths are provided
in the GSA Data Repository (see footnote 1).
Felsic granulite xenoliths define distinct and
enriched arrays in isotope plots (Fig. 9), showing elevated 87Sr/86Sr spanning a wide range
from 0.7109 to 0.7820, 143Nd/144Nd values from
0.51258 to 0.51219, 207Pb/204Pb from 15.618 to
15.679, 208Pb/204Pb from 38.577 to 38.921, and
206
Pb/204Pb values from 18.731 to 19.679. Isotopic compositions of the felsic granulite suite
overlap values reported from “garnet granulite”
xenoliths from Bolivia (McLeod et al., 2013)
and are dissimilar to gneisses of the Arequipa
terrane, which exhibit higher 207Pb/204Pb and
208
Pb/204Pb values (Fig. 9; Loewy et al., 2004).
Cuzco suite metasedimentary clinopyroxenites
lie along the Sr-Nd and Pb isotope array defined
by felsic granulite xenoliths at lower 87Sr/86Sr
(0.7086–0.7094) and Pb (i.e., 206Pb/ 204Pb,
207
Pb/204Pb, and 208Pb/204Pb) values, and higher
143
Nd/144Nd (0.51261–0.51277). Puno suite
clinopyroxenites plot at lower 87Sr/86Sr and Pb
values and higher 143Nd/144Nd than Cuzco suite
clinopyroxenite xenoliths. Diorite and gabbro
xenoliths show the most radiogenic Nd isotope
ratios (0.51258–0.51292) and the least radiogenic Sr values (0.7041–0.7089) over a wide
range of Pb values. This suite of xenoliths shows
a high degree of overlap in isotopic composition
with the high-K intermediate volcanic rocks that
entrain them (Carlier et al., 2005; Mamani et al.,
2010). Peridotite and mafic granulite xenoliths
exhibit intermediate Sr-Nd-Pb isotopic values
(Fig. 9), approximating the average xenolith
value for the entire suite.
Geochronology
Zircon U-Pb
Zircon grains from a single diorite xenolith
from the Oropesa flow (11MA20E) are subhedral, 100–150 µm in length, and inclusion
poor, and they exhibit simple oscillatory zoning
patterns in cathodoluminescence (CL) images
(Fig. 10A). In situ laser-ablation–inductively
coupled plasma–mass spectrometry (LA-ICPMS) analysis of nine zircon grains from this
sample yielded a weighted mean age of 0.9 ±
0.4 Ma (2s; Fig. 10A [see footnote 1 for data]),
which is identical within uncertainty to published ages from the volcanic host rocks (Carlier
et al., 2005).
Figure 6. Calculated pressures and temperatures for felsic granulite,
diorite/gabbro, and mafic granulite samples, showing 30 °C/km gradient and aluminosilicate stability fields for reference. 1σ uncertainty
ellipses (felsic granulite samples) are based on propagation of uncertainties on thermodynamic data and activity-composition relationships through thermobarometric calculations. 1σ error bars (­diorite/
gabbro and mafic granulite) are from Table 2. Al—aluminum;
Hbl—hornblende; Plag—plagioclase; PX—pyroxene; CPX—clinopyroxene; Qtz—quartz; GASP—garnet-aluminosilicate-plagioclasequartz; Ky—kyanite; And—andalusite; Sil—sillimanite.
Geological Society of America Bulletin, v. 127, no. 11/121785
Chapman et al.
11
4
2
2
Bt
8
11
Grt Plag Kfs
Bt Qtz melt
10
4
Grt Plag Kfs
Opx Qtz melt
7
3
8
melt
7
8
8
Ms
P (kbar)
12
6
2
9
Opx
t
10
15
Gr
5
1
14
12
10
14
Grt Plag Kfs
Cpx Ms Bt Qtz
12
8
6
4
10
13
14
9
6
4
Plag Kfs
Ms Bt Qtz H2O
2
6
Plag Kfs
Bt Qtz melt
5
Grt Plag Kfs
Opx melt
16
6
17
4
4
8
12
3
500
600
700
13
20
18
800
2
Gr
t
19
900
T (˚ C)
Figure 7. Calculated pressure-​temperature pseudosection for the bulk composition of
sample 11MA20A with contours of garnet volume percent and THERMOCALC-​based
thermobarometric results overlain. Evidence for biotite breakdown as a result of dehydration partial melting suggests that felsic granulite samples probably traversed the
Grt-​Plag-​Kfs-​Opx-​Qtz-​melt field during their evolution. Thick univariant lines show appearance or disappearance of important phases. Shaded field shows stability of free H2O.
Numbered fields: 1—Grt-​Plag-​Kfs-​Cpx-​Ms-​Bt-​Zo-​Qtz; 2—Grt-​Plag-​Kfs-​Ms-​Bt-​Zo-​Qtz;
3—Plag-​Kfs-​Ms-​Bt-​Zo-​Qtz; 4—Plag-​Kfs-​Ms-​Bt-​Zo-​Qtz-​H2O; 5—Plag-​Kfs-​Cpx-​Ms-​Bt-​
Zo-​Qtz; 6—Plag-​Kfs-​Ms-​Bt-​Qtz-​Zo-​melt; 7—Plag-​Kfs-​Ms-​Bt-​Qtz-​melt; 8—Plag-​Kfs-​Bt-​
Qtz-​H2O; 9—Grt-​Plag-​Kfs-​Cpx-​Ms-​Bt-​Qtz-​melt; 10—Grt-​Plag-​Kfs-​Ms-​Bt-​Qtz-​Zo-​melt;
11—Grt-​Plag-​Kfs-​Ms-​Bt-​Qtz-​melt; 12—Plag-​Kfs-​Bt-​Crd-​Qtz-​melt; 13—Grt-​Plag-​Kfs-​Bt-​
Crd-​Qtz-​melt; 14—Grt-​Plag-​Kfs-​Cpx-​Qtz-​melt; 15—Grt-​Plag-​Kfs-​Qtz-​melt; 16—Grt-​Plag-​
Kfs-​Opx-​Bt-​Qtz-​melt; 17—Grt-​Plag-​Kfs-​Opx-​Bt-​melt; 18—Grt-​Plag-​Kfs-​Opx-​Crd-​melt;
19—Plag-​Kfs-​Opx-​Crd-​Spl-​melt; 20—Grt-​Plag-​Kfs-​Opx-​Bt-​Crd-​melt. Mineral abbreviations not listed in Figure 4: Crd—cordierite; Ms—muscovite; Zo—zoisite.
Forty-six concordant zircon grains were
analyzed from a felsic granulite (11MA20A)
within the same flow (Fig. 10B [see footnote 1]). Zircon grains from this and other felsic
granulite samples are subrounded, ~50–200 µm
in length, and inclusion poor, and they preserve
a faint and broad zoning texture in approximately half of studied grains. The remainder
show two distinct grain domains: (1) oscillatory
zoned CL-dark cores overgrown by (2) oscilla-
1786
tory zoned to patchy and convolute zoned CLbright rims up to ~50 µm wide. Targeted analyses of core and rim domains yielded exclusively
Mesoproterozoic (1017–1529 Ma) and Neoproterozoic to Paleozoic (593–472 Ma) ages,
respectively. Of these calculated rim ages, only
Ordovician ages are concordant, suggesting
that calculated Neoproterozoic ages represent
mixed ages where the beam overlapped rim and
core domains.
A normalized probability plot comparing detrital zircon age spectra from sample
11MA20A with Bolivian granulite xenoliths
(McLeod et al., 2013) and with Paleozoic and
Mesozoic cover assemblages of the Coastal
Cordillera, Altiplano, and the Eastern Cor­
di­
llera of Peru (Reimann et al., 2010; Bahlburg
et al., 2011; Reitsma, 2012) is shown in Figure
10B. The normalized probability plot was constructed using 206Pb/238U ages for grains younger
than 800 Ma and 207Pb/206Pb ages for grains
older than 800 Ma. Analyses with greater than
10% uncertainty, 30% discordance, and/or 5%
reverse discordance (17 grains) were excluded.
Considering all analyzed grain domains,
sample 11MA20A has major age peaks at ca.
415 Ma and 470 Ma, scattered ages between 215
and 400 Ma and 1000 and 1550 Ma, and one
Cretaceous grain. The spectrum of ages from
this sample suggests a detrital origin for zircon
grains within it. A maximum protolith depositional age is difficult to assign to this sample,
given the presence of a single Cretaceous grain.
Based on a cluster of three grains that overlap in
age within analytical error, a conservative maximum depositional age of 229.2 ± 7.3 (2s) Ma
(Late Triassic) was calculated for this sample.
Garnet Sm-Nd
Two-point Sm-Nd garnet and whole-rock
isochrons were calculated from felsic granulite
samples 13 and 11MA20A from the Oropesa
flow, yielding ages and 2s uncertainties of
42 ± 2 and 17 ± 39 Ma, respectively (Fig. 11).
Taken together, garnet and whole-rock Sm-Nd
data from both analyzed samples yield a fourpoint isochron age of 41 ± 7 Ma (2s). Mineral
data and whole-rock Sm and Nd isotope data
are provided in the GSA Data Repository (see
footnote 1). Given an average garnet grain size
of ~1 mm and assuming a slow cooling rate
of 3 °C/m.y. (typical of orogenic plateaus), a
Sm-Nd closure temperature of ~760 °C was calculated, using available diffusion parameters of
Ganguly and Tirone (1999). This calculated closure temperature is similar to peak metamorphic
temperature estimates from these rocks (Fig. 6;
Table 2), which may provide an explanation for
the wide range in precision in our two Sm-Nd
ages. In other words, elevated metamorphic
temperatures may have disturbed the Sm-Nd
systematics in sample 11MA20A relative to
sample 13.
Monazite U-Th-Pb
Monazites in samples 11MA20A and
11MA20C show two distinct compositional
domains, based on electron microprobe X-ray
mapping and spot analyses (Fig. 12 [see footnote 1]): grain interiors with average wt%
Geological Society of America Bulletin, v. 127, no. 11/12
Xenolith constraints on Altiplano architecture and assembly
A
Figure 8 (on this and following
page). (A) Major-element variation diagrams for Peruvian
xenoliths compared to back-arc
lavas (Chapman and Ducea,
2013) exposed along the CuzcoVilcanota fault system (CVFS).
FeO* denotes total Fe as FeO.
Geological Society of America Bulletin, v. 127, no. 11/121787
Chapman et al.
Sample/N-MORB
1000
B
ion beam. Uncertainties were multiplied by
the square root of the MSWD to reflect the
higher degree of scatter for individual analyses.
Monazite U-Pb ages from the two compositional
domains recognized in monazite from each
sample overlap within analytical error. Monazite U-Th-Pb ages may reflect either the timing of growth or cooling through ~800–900 °C
(Cherniak et al., 2004; Gardes et al., 2006).
100
10
DISCUSSION
Xenolith Classification Based
on Radiogenic Isotopes
1
.1
.01
.001
Cs Rb Ba Th U Nb K La Ce Pb Pr Sr P Nd Zr Sm Eu Ti Dy Y Yb Lu
Sample/C1 Chondrite
1000
100
10
1
.1
La
Ce
Pr
Nd Pm Sm Eu
Gd Tb Dy Ho Er
Tm Yb
Lu
Figure 8 (continued). (B) Normal mid-ocean-ridge basalt (N-MORB)–normalized traceelement and chondrite-normalized rare earth element (REE) diagrams comparing xeno­lith
sample groups. N-MORB and chondrite values are from Sun and McDonough­(1989).
ThO2 = 5.43%, UO2 = 1.08%, and Y2O3 = 3.46%
and rims with ThO2 = 5.08%, UO2 = 1.04%, and
Y2O3 = 2.93%. Monazites show sensitive highresolution ion microprobe (SHRIMP)–determined REE patterns with strong enrichment
of the LREEs relative to the HREEs (average
La/Yb = 605) and show strong negative Eu
anomalies (Eu/Eu* = 0.06; Fig. 13).
1788
Weighted mean monazite ages of 4.4 ± 0.3 Ma
(mean square of weighted deviates [MSWD] =
6.3, 2s) and 3.2 ± 0.2 Ma (MSWD = 4.6, 2s)
were calculated from samples 11MA20A and
11MA20C, respectively (Fig. 13 [see footnote 1). Calculated weighted mean ages do not
include three analyses in which both monazite
and adjacent minerals were overlapped by the
There are two groups of crustal xenolith
samples based on their measured Nd and Sr
isotopes: Group 1 rocks show 143Nd/144Nd >
0.5125 and 87Sr/86Sr < 0.710, and group 2 rocks
yield 143Nd/144Nd < 0.5125 and 87Sr/86Sr > 0.710
(Fig. 9). Group 1 rocks generally have igneous
textures, with mostly mafic to intermediate compositions (SiO2 < 55 wt%), and are interpreted to
represent Cenozoic additions to the crust. Group
2 assemblages show metamorphic textures and
are interpreted here to represent basement and
metamorphosed cover samples of the Altiplano,
e.g., Paleozoic and Mesozoic rocks from the
Central Andes, including the Famatinian arc,
the Ollantaytambo formation, the Mitu Group,
and other regionally extensive units. Similar distributions of isotopes exist among igneous and
metamorphic basement rocks from the Central
Andes (Wörner et al., 2000; Lucassen et al.,
2001; Loewy et al., 2003, 2004; Fig. 9).
The majority of group 1 xenoliths are ultramafic, mafic, or intermediate in composition,
and they were most likely derived from relatively primitive mantle magmas. The extent to
which some of them represent asthenospherederived additions to the crust or mixtures of
asthenospheric and lithospheric materials is
debatable, but it is clear that they are more narrowly grouped in terms of radiogenic isotopes
and have younger model ages that the second
group of crustal xenoliths (generally <1.0 Ga vs.
>1.0 Ga [see footnote 1]). Next, we will focus on
the significance of group 2 xenoliths, which we
interpret to represent basement and metamorphosed cover rocks of central South America.
Origin of Crustal Xenoliths
Metasediments
Garnet leucogneisses (i.e., felsic granulites)
are interpreted as the metasedimentary cover to
the South American basement, similar to highgrade intermediate to felsic granulite facies
xenoliths recovered from the Bolivian Altiplano
(Davidson and de Silva, 1995; McLeod et al.,
2012, 2013). These deposits must be younger
than the 550 Ma Rb-Sr whole-rock error
Geological Society of America Bulletin, v. 127, no. 11/12
Xenolith constraints on Altiplano architecture and assembly
A
B
Figure 9. (A) Sr and Nd isotopic compositions for group 1 and group
2 Peruvian xenoliths compared to Bolivian xenoliths (McLeod et al.,
2013), back-arc lavas (Chapman and Ducea, 2013), and Paleozoic
basement rocks (Lucassen et al., 2001) exposed along the CuzcoVilcanota­fault system (CVFS). Nd isotopic values from the northern Arequipa terrane (average ~0.5117) lie outside the plotted range.
(B) Pb isotopic compositions for Peruvian xenoliths compared to
Bolivian xenoliths (McLeod et al., 2013), back-arc lavas (Chapman
and Ducea, 2013) exposed along the Cuzco-Vilcanota fault system
(CVFS), the northern Arequipa terrane (Wörner et al., 2000; Loewy
et al., 2004), and the Amazon craton (Loewy et al., 2003).
chron defined by these rocks. Furthermore, the
­detrital distributions of age peaks corroborated
by the maximum depositional age of 229.2 ±
7.3 Ma from zircon U-Pb constraints on sample
11MA20A (Fig. 10B) suggest that the age is
Late Triassic or younger, and possibly as young
as Cretaceous, based on the presence of a single
grain of this age.
Mesozoic volcanic and sedimentary rift and
postrift cover rocks of the mid- to Late Triassic Mitu Group, the Early Cretaceous Huancane
Formation, and the mid- to Late Cretaceous
Yuncay­
pata Group commonly comprise the
country rocks into which Rumicolca back-arc
volcanic rocks intruded (e.g., Kontak et al.,
1990a; Sempere et al., 2002; Mišković et al.,
2009; Reitsma, 2012). Two key similarities
between Mesozoic strata of SE Peru, in particular the Mitu Group and Huancané Formation,
and leucogneiss xenoliths lead us to suggest that
the latter are the highly metamorphosed equivalents of the former. First, detrital zircon age
spectra from Mesozoic strata (Reitsma, 2012;
Perez and Horton, 2014; Fig. 10B) overlap the
spectrum of leucogneiss sample 11MA20A,
each containing distinct late Paleozoic to Late
Triassic, early Paleozoic, and Mesoproterozoic
age peaks. These age peaks are associated with
0.16–0.35 Ga Gondwanide (e.g., Kontak et al.,
1990a; Mišković et al., 2009), 0.44–0.53 Ga
Famatinian-Pampean, and 0.9–1.3 Ga Sunsas/
Grenville magmatic-orogenic events. Mesozoic
strata, as well as leucogneiss xenoliths studied
here, lack significant proportions of the Neo­
protero­
zoic and Paleoproterozoic and older
zircon grains that are abundant in South American cratonic crust. Grains of these ages are
present in Paleozoic strata overlying the Altiplano, Eastern Cordillera, and Arequipa terrane
(Reimann et al., 2010; Bahlburg et al., 2011).
Second, volcano-sedimentary assemblages of
the Mitu Group (interbedded sandstones, mudstones, conglomerates, and intermediate lavas)
represent suitable protoliths for leucogneiss
xenoliths.
The inference that felsic xenoliths in the
Rumicolca Formation represent the metamorphic equivalents of the Mitu Group implies
that portions of the Mitu Group must somehow
have been transferred to ~30–35 km depth. As
mentioned previously, the volcano-sedimentary
strata of the Mitu Group are thought to represent
rift deposits produced during the initial stages of
Gondwana disassembly (Kontak et al., 1990a;
Sempere et al., 2002; Perez and Horton, 2014).
Inversion of basins containing Mitu Group
strata is thought to have commenced at a slow
pace in Late Jurassic–Early Cretaceous time and
culminated during the Andean orogeny (Sempere et al., 2002), although the precise timing of
Geological Society of America Bulletin, v. 127, no. 11/121789
Chapman et al.
uplift is not well understood. Our garnet wholerock Sm-Nd age of 42 ± 2 Ma from leucogneiss
sample 13 indicates that garnet growth in this
sample, and hence a major pulse of Mitu basin
inversion and accompanying crustal thickening
beneath the plateau, took place in the Eocene.
Burial of the Mitu basin through thrust inversion requires that the vertical throw on Triassic
rift structures was in places significant (greater
than the ~7-km-thick Paleozoic section) and
that the normal fault basins were not completely
inverted. Complete reactivation of some rift
structures, with limited reactivation of others,
provides a mechanism for burial of Mitu rocks
preserved in the hanging walls of normal faults
(Fig. 14). Burial by basement thrust structures
can produce Mitu paleodepths no greater than
25–26 km, which is below the lower limit of
1s uncertainty in calculated peak metamorphic
pressures of 9.1 ± 1.4–9.5 ± 1.3 kbar (Table 2),
assuming a mean crustal density of 2.8 g/cm3.
Therefore, it is considered highly likely that
additional processes, discussed in the following
sections, were responsible for conveying Mesozoic sediments to ~9 kbar conditions.
Following burial of Mesozoic sediments to
midcrustal levels, peak temperatures of 750 °C
and higher (and a thermal gradient of 30 °C/km;
Fig. 6) were likely achieved through some combination of thermal relaxation in the crust and
Cenozoic magma input. The time lag between
Eocene garnet Sm-Nd ages and Pliocene monazite U-Th-Pb ages provides insight into the
relative importance of thermal relaxation versus
magmatism in the metamorphism of these rocks.
Pliocene monazite U-Th-Pb ages probably
represent the timing of mineral growth rather
than cooling ages, considering textural evidence
for monazite growth during partial melting (Figs.
4G and 4H), calculated peak metamorphic temperatures below ~800–900 °C U-Th-Pb closure,
and the fact that monazite must spend considerable periods of time at or above 800 °C to affect
U, Th, and Pb systematics (e.g., Cherniak et al.,
2004; Gardes et al., 2006; Williams et al., 2007).
Monazite growth during Pliocene time indicates that ~40 m.y. had elapsed since garnet
growth before partial melting of felsic leuco­
gneisses took place. Recent thermochronologic
results from plutons in the Eastern Cordillera of
SE Peru suggest similarly protracted shortening
beginning in the middle Eocene and continuing until at least middle Miocene time (Perez
et al., 2013).
Monazite chemistry, specifically the observed
decline in Y concentration from monazite core
to rim, reveals additional information regarding
garnet stability and equilibration with monazite. Recall that garnet grains in felsic leuco­
gneiss xenoliths exhibit thin, melt-containing,
1790
A
Figure 10 (on this and following page). (A) U-Pb zircon age spectrum (after Ludwig, 2003) for laser ablation–inductively coupled
plasma–mass spectrometry data on one diorite xenolith (sample
11MA20E). Data point errors are 1σ. Cathodoluminescence image
(inset) shows two of the analyzed grains. MSWD—mean square of
weighted devi­ates.
kelyphitic rims in the vicinity of biotite grains
that are breaking down to glass, orthopyroxene,
spinel­, monazite, Fe-Ti oxides, orthoamphibole,
and zircon. Similar textural and phase relationships are commonly observed in granulite-grade
metapelitic rocks in which garnet-involved
incongruent melting of biotite has taken place
(e.g., Cesare, 2000; McLeod et al., 2012). The
biotite-out reaction has a positive slope in the
vicinity of the calculated peak conditions in P-T
space (Fig. 7), indicating that dehydration partial melting is expected under decreasing pressure and/or increasing temperature conditions.
We suggest that biotite dehydration melting
with garnet as a reactant and monazite, spinel,
and orthopyroxene as products occurred during
decompression, perhaps accompanied by a thermal pulse, due to entrainment of these rocks in
back-arc magmas as xenoliths. We further speculate that garnet decomposition contributed Y to
the reaction volume, which was sequestered into
monazite at the onset of the reaction, with diminishing Y as the reaction progressed. Given that
the volcanic field containing these xenoliths is
~3 m.y. younger than monazite U-Pb ages from
xenoliths entrained within it (Table 1; Kaneoka
and Guevara, 1984), these xenoliths were likely
plucked from ~30–35 km depth and resided in
the mid- to upper crust prior to ultimate eruption.
Meta-Igneous Assemblages
Mafic granulite samples (e.g., 11MA2A) have
measured Sr isotopic ratios of 0.707–0.708, eNd
of ~–5, and equilibration temperatures exceeding
1000 °C and pressures of over 11 kbar (Table 2;
Fig. 6). These rocks have mafic chemistry, have
metamorphic textures, and are characterized
by isotopic ratios similar to the Famatinian arc
(e.g., Otamendi et al., 2009, 2012). Equivalent
mafic granulites were not described from the
Bolivian xenolith suites (McLeod et al., 2013).
We interpret these rocks to represent fragments
of the South American lithosphere upon which
Mesozoic strata were deposited. Peridotite nodules from the Huarocondo flow also probably
represent xenoliths of the sub-Altiplano mantle
lithosphere or pieces of peridotite lenses that
crop out within the Mitu rift zone (Jacay et al.,
1999; Sempere et al., 2002). It could be argued,
based on the small size of peridotite fragments
studied here, that these nodules represent
glomerocrysts. However, foam textures and the
lack of glomeroporphyritic textures­preclude the
possibility that these nodules formed by syn­
neusis (Hogan, 1993).
Nd ratios and depleted mantle model ages
(TDM) values of 0.46–1.4 Ga in mafic granulite
and peridotite xenoliths (see footnote 1) suggest that these rocks represent mafic fragments
of the Famatinian arc. An Amazonian origin is
ruled out because Amazonian basement rocks
are expected to have lower eNd ratios (<–15) and
older TDM values (>1.5 Ga). Another possibility is that mafic granulites represent relatively
young (Permian to Cenozoic) mafic magmatic
additions to the lower crust that were subsequently metamorphosed. However, the majority
of mafic magmatic products with late Paleozoic
Geological Society of America Bulletin, v. 127, no. 11/12
Xenolith constraints on Altiplano architecture and assembly
B
Figure 10 (continued). (B) Normalized probability plot comparing the spectrum of concordant zircon ages from
felsic granulite sample 11MA20A, plotted alongside spectra from Bolivian granulite xenoliths (McLeod et al., 2013)
and composite curves from Paleozoic and Mesozoic cover assemblages of the Coastal Cordillera, Altiplano, and the
Eastern Cordillera of Peru (Reimann et al., 2010; Bahlburg et al., 2011; Reitsma, 2012). Cathodoluminescence image
(inset) shows ana­lyzed core and rim age domains (1σ ages of each spot given in Ma).
Geological Society of America Bulletin, v. 127, no. 11/121791
Chapman et al.
Figure 11. Garnet (Grt) whole-rock (wr) Sm-Nd isochron for samples 11MA20A (labeled wr-11) and 13 (wr-13). Error bars and calculated two-point isochron ages (Ludwig, 2003) are 2σ.
Figure 12. Monazite X-ray maps.
Location of two analyzed spots is
shown in backscattered electron
image (BSE); monazite core and
rim regions overlap in age (1σ uncertainties shown).
1792
Geological Society of America Bulletin, v. 127, no. 11/12
Xenolith constraints on Altiplano architecture and assembly
A
B
Figure 13. Sensitive high-resolution ion microprobe (SHRIMP) data:
(A) U-Pb geochronologic and (B) chondrite-normalized (e.g., Sun
and McDonough, 1989) rare earth element (REE) data for selected
monazite grains from samples 11MA20A and 11MA20C. Error
ellipses­are 2σ. MSWD—mean square of weighted deviates.
to pre-Eocene ages that are found on the plateau
within 250 km of the xenolith localities have
more depleted isotopic signatures (Mamani
et al., 2010). Specifically, their eNd values range
from –2 to +4, which are distinctive from the
western South American orogenic margin range
of eNd (~–10 in the Central Andes). In conclusion, we interpret mafic granulites as basement
rocks to the Mesozoic and older cover sequence,
represented by garnet leucogneisses.
A corollary to this study is that all igneoustextured basement rocks (hornblende diorites,
gabbros, etc.), which are the predominant xenolith populations among the localities from southern Peru and Bolivia (McLeod et al., 2013), are
magmatic additions newer than the mid-Ceno-
zoic metamorphic event described here. In fact,
there is U-Pb zircon evidence that at least one of
the diorites from Oropesa is Quaternary in age
(Fig. 10A).
Shortening Constraints
What drove crustal thickening in the Central Andes and uplift of the northern Altiplano?
Shortening was clearly underway in southern
Peru by the Eocene (Lamb and Hoke, 1997),
during what is regionally referred to as the Incaic
orogeny (e.g., Steinmann, 1929; Noble et al.,
1979; Sandeman et al., 1997), and possibly even
before (e.g., the Campanian Peruvian orogenic
episode; Steinmann, 1929). The Incaic orogeny
was originally defined by well-dated unconformities and pluton crosscutting relationships
that mark the onset of deformation at 38–41 Ma
(Sandeman et al., 1995). Biotite Ar/Ar cooling
ages of 37 Ma, combined with 18–30 Ma apatite fission-track ages (Kontak et al., 1990b),
suggest that shortening and crustal thickening
continued in the Eastern Cordillera until ca.
20 Ma. Magmatic arc geochemistry (Mamani
et al., 2010) indicates a gradual thickening of
the crust, which may have started as early as
100 Ma, although REE patterns become steep,
requiring garnet in the lower-crustal residue,
only after 30 Ma.
Gotberg et al. (2010) estimated 123 km of
shortening (40%) across the eastern Altiplano,
Eastern Cordillera, and Subandean zone at the
latitude of southern Peru, only 14% of which
was estimated to be younger than 15 Ma.
Although most of the shortening is thought to
have occurred early (40–20 Ma), the total shortening only accounts for 40% of what is needed
to thicken the plateau to 70 km. The updated
shortening estimates presented here provide
the geometry necessary to bury Mitu rocks to
depths no greater than 25 km (Fig. 14), falling
short of calculated peak equilibration depths of
33–35 km (Table 2), assuming a mean crustal
density of 2.8 g/cm3. Furthermore, the shortening estimate of 185 km presented here can
account for only 62% of the current crustal thickness. These shortening and burial deficits leave
room for speculation as to whether magmatic
underplating is more involved in thickening the
crust than previously thought, or whether there is
significant shortening buried under the Western
Cordillera magmatic arc.
Implications for Plateau Uplift
and Crustal Flow
The history of surface uplift of the Central
Andean orogen is debated. Some workers have
suggested that major (>2 km over 2–3 m.y.)
surface uplift of the Altiplano occurred during
the mid- to late Miocene based on temperaturesensitive paleoaltimeters (Ghosh et al., 2006;
Garzione et al., 2008; Saylor and Horton, 2014).
These postulated episodes of rapid uplift are
explained by wholesale removal of dense lower
crust and/or mantle lithosphere (e.g., Gar­
zione et al., 2006; Hoke and Garzione, 2008).
In contrast to the hypotheses of rapid pulses of
uplift, others have favored long-lived and steady
isostatic rise in topography due to protracted
crustal shortening (e.g., Barnes and Ehlers,
2009; Isacks, 1988).
Our results demonstrate that low-density
sedimentary cover rocks and their metamorphosed equivalents have occupied the upper
Geological Society of America Bulletin, v. 127, no. 11/121793
Chapman et al.
Figure 14. Balanced cross-section A-A′ from the eastern edge of the Western Cordillera, across the Altiplano, and into the Eastern Cor­
dillera. Cross-section trace is shown in Figure 1. Deformed length is 185 km, and restored undeformed length is 370 km (185 km, or 50%,
shortening). Additional burial of Mitu Group sediments from ~25 km depth, perhaps by magma loading, to >30 km depth is necessary to
explain thermobarometric data. Approximate location of xenoliths recovered in this study is shown on deformed section.
35 km of the northern Altiplano crust since the
Eocene. This is consistent with models for early
Cenozoic shortening and crustal thickening and
accompanying rock uplift from beneath the
northern Altiplano. However, our results do not
preclude the possibility that regional uplifts were
driven by short-lived lithospheric delamination
events, as Eocene–Oligocene crustal thickening
and magmatism could have preconditioned the
lower crust and/or mantle beneath the northern Altiplano to foundering and renewed uplift
during and following Miocene time. A major
influx of primitive magmas represented by diorites, gabbros, and cumulate xenoliths in these
suites could be related to a younger convective
1794
removal event that took place in the Miocene or
more recently.
Monazite U-Pb data and textural observations
from metasedimentary assemblages, as well
as zircon U-Pb data from dioritic rocks, indicate that a significant proportion of the middle
crust (~20–25 km depth) beneath the northern
Altiplano has either undergone some degree of
partial melting or has been added to the crust
as a magma since ca. 5 Ma. This observation is
pertinent to discussions of the Bolivian orocline,
which possibly resulted from along-strike variations in shortening due to variations in thickness of the colliding Nazca plate (Isacks, 1988;
Sempere et al., 2002; Capitanio et al., 2011).
The orocline is also characterized by extremely
thick (locally ~75 km; Beck et al., 1996; Beck
and Zandt, 2002), mechanically decoupled
crust. While the crustal thickness is fairly uniform along the orocline, shortening estimates
are not. The total amount of crustal shortening
accumulated since Eocene time is estimated to
have been 115 km (this study) to 177 km (Gotberg et al., 2010) greater in central Bolivia compared to southern Peru. The Bolivian shortening
estimates can account for crustal thickness in
that portion of the orocline; however, shortening
estimates from southern Peru cannot account
for the extreme crustal thickness beneath the
northern Altiplano. A possible explanation for
Geological Society of America Bulletin, v. 127, no. 11/12
Xenolith constraints on Altiplano architecture and assembly
the shortening to thickness disparity is lateral
flow of ductile mid- to lower-crustal material­
from Bolivian to Peruvian sections of the Altiplano (Yang and Liu, 2003; Ouimet and Cook,
2010; Gotberg et al., 2010). Peruvian xenoliths provide evidence for a zone of partial
melt situated between 35 km (based on leuco­
gneiss pressure estimates) and 20 km (based on
diorite estimates) depth beneath the northern
Altiplano. However, the degree to which this
relatively weak crustal zone has been mobilized
parallel to the strike of the orogen is uncertain.
Similarities between garnet leucogneisses and
Mesozoic sediments in the vicinity of Cuzco
suggest that, aside from being brought from the
surface to ~30–35 km depth, the protoliths for
garnet leucogneisses were deposited proximal
to their current position (i.e., it is unlikely that
these rocks flowed northward in the mid- to
deep crust). Hence, partially molten material
beneath the northern Altiplano was more likely
generated during a thermal pulse associated
with the Pliocene and later influx of primitive
magmatic additions. Such an influx may have
played a significant role in thickening the crust
under the northern Altiplano through magmatic
additions and possibly convective overturning
of the crust, thereby transporting supracrustal
rocks downward as has been suggested beneath
the Altiplano-Puna volcanic complex (Babeyko
et al., 2002).
SUMMARY AND CONCLUSIONS
Crustal and probable mantle xenoliths discovered in Pliocene and Quaternary high-K intermediate volcanic rocks from the southern Peruvian
Altiplano show that the Mesozoic regional basement-cover contact is located ~30–40 km below
the modern surface. Equilibrium P-T estimates
of >1000 °C and >11 kbar from mafic granulite
xenoliths indicate that basement assemblages
below this interface are predominantly mafic
and are isotopically similar to Central Andean
crust. Granulite grade metasedimentary rocks
with radiogenic compositions (87Sr/86Sr >0.711
and 143Nd/144Nd <0.5126) comprise a cover
sequence—correlative to Triassic Mitu Group
rift basin deposits—that has been transferred
to >30 km depth beneath the plateau at temperatures >750 °C. Based on our new shortening
estimates, a maximum of ~25 km of Mitu basin
tectonic burial could have been accomplished
through basement-involved thrusting, falling
short of thermobarometric estimates requiring that these rocks equilibrated at ~30–35 km
paleodepth. Therefore, additional processes
such as magma loading must have been in part
responsible for conveying Mesozoic sediments
to ~9 kbar conditions.
A garnet Sm-Nd age of 42 ± 2 Ma, interpreted
to reflect the timing of garnet growth in Mitu
Group protoliths, suggests that metamorphism
associated with burial of these rocks and accompanying major crustal thickening and uplift of
the northern Altiplano began in the Eocene,
probably during the Incaic regional shortening
episode. Following burial, a protracted phase
(~40 m.y.) of thermal relaxation–related heating
and lower-crustal residence with limited unroofing occurred. Monazite neoblasts, yielding
U-Pb ages of 3.2 ± 0.2–4.4 ± 0.3 Ma, associated
with incongruent melting of biotite and involving garnet as a reactant, probably resulted from
near-isothermal decompression resulting from
xenolith entrainment and rapid transport toward
the surface. Significant primitive magmatism
with 87Sr/86Sr <0.709 and 143Nd/144Nd >0.5126,
as evidenced by Pliocene–Quaternary dioritic
to gabbroic xenoliths emplaced at ~7 kbar and
their extrusive equivalent volcanic host rocks,
is inferred to be responsible for further regional
crustal thickening by magma loading. The presence of these primitive (asthenospheric) mafic
melts in the crustal section may reflect later
removal of parts of the lithosphere.
ACKNOWLEDGMENTS
This work was supported by U.S. National Science Foundation grants EAR0907880 (to Ducea), and
EAR0908972 (to McQuarrie), as well as Romanian
National Sciences Consiliul National al Cercetarii
Stiintifice din Invatamantul Superior grant PN-II-IDPCE-2011-3-0217. Victor Carlotto is acknowledged
for providing information on volcanic rocks from
southern Peru, and Ally Thibodeau and Robinson
Cecil­are thanked for help with sample collection in
the field. The manuscript benefited from discussions
with S. Beck, J. Hogan, G. Mitchelfelder, N. Perez,
and G. Zandt and thoughtful reviews by Fernando
Ortega Gutiérrez and an anonymous reviewer.
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Science Editor: David Ian Schofield
Associate Editor: Luca Ferarri
Manuscript Received 9 September 2014
Revised Manuscript Received 13 March 2015
Manuscript Accepted 5 May 2015
Printed in the USA
Geological Society of America Bulletin, v. 127, no. 11/121797