ANDEAN FOREARC DYNAMICS, AS RECORDED BY DETRITAL...

Journal of Sedimentary Research, 2015, v. 85, 646–659

Research Article

DOI: http://dx.doi.org/10.2110/jsr.2015.45

ANDEAN FOREARC DYNAMICS, AS RECORDED BY DETRITAL ZIRCON FROM THE EOCENE TALARABASIN, NORTHWEST PERU

ANGELA M. HESSLER1AND ANDREA FILDANI2

1The Deep Time Institute, 8711 Bluegrass Drive, Austin, Texas 78759, U.S.A.2Statoil Research, Development and Innovation Center, 6300 Bridge Point Parkway, Suite 2–500 Austin, Texas 78730, U.S.A.

e-mail: [email protected]

ABSTRACT: Eocene strata of the Talara forearc basin of northwest Peru contain detrital-zircon (DZ) populations that, whenintegrated with other provenance indicators, record paleogeographic shifts in the Andean forearc that can be linked to larger-scale tectonic drivers such as subduction erosion, plate convergence rate, and regional accretion events. Because forearc basinsalong the western margin of South America are particularly difficult to preserve, the new DZ results provide rare insight intothe trench–arc dynamics associated with the Andean orogenic cycle. The Talara basin tapped into both the Andean volcanic arcand older basement, and the DZ populations subsequently reflect long-lived plate interactions and progressive crustal growthalong the northwestern edge of Amazonia, Gondwana, and South America. Dominant U–Pb zircon age populations include:36–110 Ma (Andes convergence); 215–285 Ma (western Gondwana extension); 460–625 Ma (Rodinia extension and westernGondwana convergence); 950–1250 Ma (Grenville orogeny); and . 1250 Ma (Amazonia assembly). The data corroborateprevious geochronology on pre-Andean basement and support the existence of a Famatinian arc (, 480 Ma) andCarboniferous–Triassic magmatism along northwestern Gondwana. Pre-Andean and Amazonian DZ likely was recycled fromPaleozoic metasedimentary rocks of the Andean Cordillera and Amotape terrane, although the data do not preclude delivery ofAmazonian detritus directly into the forearc via a large-scale, westward drainage. The Andean DZ population constrains thetiming of arc activity and quiescence, with two peaks at , 75 Ma and , 50 Ma, and a gap at , 70 Ma, linked to accretion ofthe Caribbean Oceanic Plateau against northwest South America. Arc volcanism was concurrent with Eocene forearcdeposition, and syndepositional DZ is used to revise maximum depositional age for the upper Eocene Helico Member andVerdun Formation. Progressive subsidence and arc shut-down in the late Eocene coincide with arc collision to the north, upliftof the Amotape block, subduction erosion, and slab shallowing.

INTRODUCTION

Forearc basins preserve the stratigraphic record of first-order interac-tions among magmatic arcs, forearcs, and accretionary prisms abovesubduction zones (see review in Dickinson 1995). Subduction parametersare controlled by changes in (1) convergence rate and orientation, (2) age,thickness, and density of the down-going slab, and (3) subduction-channel shear, all of which impact processes on overriding plates that canbe recorded as major stratigraphic events and provenance patterns inforearc basins. In particular, detrital-zircon (DZ) patterns in forearcstrata have allowed the correlation between well constrained paleogeo-graphic changes and major subduction adjustments. For instance, theCretaceous–Paleogene forearc of California, one of the best preservedand most studied basins in the world (Ingersoll 1982; Dickinson 1995),preserves provenance patterns related to low-angle subduction thatbroadly affected western North America during the Laramide orogeny(Sharman et al. 2014). Provenance analysis of the Paleocene–Eoceneforearc of southern Alaska reveals a unique magmatic history associatedwith subduction of an oceanic ridge and migration of the slab window(Kortyna et al. 2013). In the Cenozoic Moquegua forearc in southernPeru, a major late Eocene DZ shift marks onset of uplift of the CentralAndes associated with flat-slab subduction and strong intraplate

coupling, rather than tectonic compression or magmatic underplating(Decou et al. 2013).

Most forearc strata are ultimately not preserved (Ingersoll 2012). Up to70% of active convergent margins are non-accreting, and the process ofupper-plate tectonic erosion (or subduction erosion) transfers crustalmaterial from all convergent margins (both accreting and non-accreting)at a total average of 1.35–1.4 km3/yr (Von Huene and Scholl 1991; Clift etal. 2009). This is particularly true along western South America, wherePeru’s western coastline has been narrowed by up to 150 kilometers dueto subduction erosion (Clift et al. 2003).

At the western tip of South America sits the Talara basin, a remnant ofthe Eocene Andean forearc now exposed along Peru’s northern coast(Fig. 1A). Its forearc position gives a proximal, though difficult topreserve, perspective into subduction-zone tectonics along the ever-changing convergent margin that is the Peru–Chile arc–trench system(Pardo-Casas and Molnar 1987; Jaillard et al. 1995; Noblet et al. 1996;Silver et al. 1998). To this end, Talara’s stratigraphy has helped testconcepts such as subduction erosion and wrench tectonism for the Eocenemargin (Fildani et al. 2008) and holds provenance information regardingAndean and pre-Andean magmatic events.

Perhaps not serendipitously, the Talara basin is preserved at animportant juncture along the Andean Cordillera: the boundary between

Published Online: June 2015

Copyright E 2015, SEPM (Society for Sedimentary Geology) 1527-1404/15/085-646/$03.00

the Central and Northern Andes, across which occur major shifts inbasement composition, structural grain, and volcanic activity (Shepherdand Moberly 1981). Pre-Cordilleran basement shows evidence of oldertectonic regimes reaching back one billion years, as this region has been atthe locus of several cycles of supercontinent formation and disaggregation(Chew et al. 2008; Miskovic et al. 2009). Structural overprint and recentvolcanic cover has mostly obscured this older history.

Detrital zircon from the Talara basin retains evidence for the timing ofmagmatic and/or metamorphic events that have acted upon theCordillera, its predecessor basement, and (depending on drainagepatterns) the craton interior. We present the first SHRIMP-RG DZU-Pb ages from the northern Andean forearc; this dataset captures, inaddition to pre-Cordillera events, the Mesozoic and early Cenozoicmagmatism and paleogeographic changes related to Farallon subduction

FIG. 1.—A) Talara basin within the context of South America’s geochronological provinces. B) Sampling location and regional tectonic features, including theCarnegie Ridge (CR), Guayaquil Fracture Zone (GFZ), Dolores–Guayaquil fault system (DGFS), the Ocala Ridge (OR), and the Sarmiento Ridge (SR). Samplelocations are marked in black: V, Verdun Fm.; CB, Cabo Blanco Mbr.; H, Helico Mbr.; PG, Pale Greda Fm.; M, Mogollon Fm. White circles mark locations of nearbyDZ or igneous-zircon (IZ) studies (Fig. 2; CA, Cardona et al. 2010; C1, Chew et al. 2008; C2, C3, Chew et al. 2007). C) Cross section through Eocene trench–forearc–arccomplex in northern Peru (modified from Fildani et al. 2008). ‘‘C–T’’ indicates Carboniferous–Triassic plutonic rock. Arrows mark potential sediment routes into Talarabasin. D) Simplified stratigraphic column for the Talara basin, with intervals sampled for DZ analysis. Chronostratigraphy is based on biostratigraphy (Fildani et al.2008) using ICS absolute ages (Cohen et al. 2013).

ANDEAN FOREARC DYNAMICS RECORDED BY DETRITAL ZIRCON IN THE EOCENE TALARA BASIN 647J S R

and the shaping of modern South America. We address the followingquestions: (1) Does the Precambrian DZ population in the forearc derivefrom Cordilleran basement and/or directly from the shield? (2) Do the DZdata provide new evidence for Paleozoic to Triassic magmatism alongwestern Gondwana? (3) Can we better constrain the timing ofCretaceous–Eocene Andean arc activity, quiescence, and major accretionevents (i.e., the Caribbean Oceanic Plateau)? (4) Do we need to revisemaximum depositional ages in the Talara basin based on the new DZdata? (5) How do DZ patterns refine our understanding of forearcdynamics and subduction mechanisms?

THE TALARA BASIN

Regional Geology

The 5000-km-long Peru–Chile arc–trench system is the result ofsubduction of the oceanic Farallon–Nazca plate beneath western SouthAmerica. Late Cretaceous calc-alkaline arc activity in Peru marks theonset of Farallon subduction (Benavides-Caceres 1999), following theAlbian opening of the South Atlantic and westward drift of SouthAmerica (Nurnberg and Muller 1991). Along the Peruvian andEcuadorian margin, a combination of flexure, wrench tectonics, andsubduction erosion created several Cretaceous–Tertiary forearc basins,which have since been uplifted due to partial subduction of oceanic ridges(Fig. 1B; Gutscher et al. 1999). Inland lies the Andean Cordillera,a Neogene uplift composed of pre-Andean basement, as well asMesozoic–Cenozoic arc-related intrusions and volcanic cover. Today,flat-slab subduction and a volcanic gap characterize the Cordillera at thelatitude of Talara (Bernal et al. 2002).

The Talara basin sits oceanward of the transition between the Centraland Northern Andes, where the Cordillera makes a pronounced bendtoward the northeast, across what is known as the Huancabambadeflection (Fig. 1B; Mourier et al. 1988). The change in orientation, theresult of differential post-Oligocene rotation on either side of thedeflection (Mitouard et al. 1990), is accompanied by a marked changein pre-Andean basement composition, namely the occurrence ofCarboniferous–Triassic intrusions into Paleozoic metasedimentary rocksnorth of the deflection, versus Carboniferous and older Paleozoicintrusions to the south.

In the associated forearc, the Dolores–Guayaquil fault system (Fig. 1B)defines another break in basement composition. To the south of thislineament, the Talara and other basins rest on autochthonous,continental-type basement (Shepherd and Moberly 1981). To the north,the forearc is underlain by oceanic plateau material (Reynaud et al. 1999),presumed to be part of the now-Caribbean plate that was accreted to theEcuadorian margin in the Late Cretaceous or Paleogene (Hughes andPilatasig 2002; Spikings et al. 2005; Toro Alava and Jaillard 2005).

Outlying Cordilleran basement punctuates the forearc of northern Peru(Fig. 1B). The Amotape Mountains are the topographic expression of thesouthern Amotape–Tahuin block, a para-autochthonous sliver ofCordilleran basement sheared and rotated to its present position startingin the Late Cretaceous (Mourier et al. 1988; Aspden and Litherland1992). Today, basement highs associated with the Amotape–Tahuin blockisolate the Talara basin from the Lancones basin to the east (Fig. 1C), theSechura basin to the south, and the Tumbes basin to the north. However,these basement structures were not always depositional barriers (Fildaniet al. 2008); the Talara basin shares some Eocene stratigraphy with boththe Sechura and Lancones basins (Caldas et al. 1980; Fildani 2004).

Forearc Positioning

The Talara basin sits in a precarious position in the Andean forearc,along which up to 150 km of Peru’s continental margin has receded sincethe Eocene due to subduction erosion (Clift et al. 2003), thinning of the

forearc wedge just above the subducted plate. This is not uncommonalong convergent margins; forearc retreat occurs along up to 80% of theworld’s convergent ocean margins at a rate between 30–90 km3/km/My(Von Huene and Scholl 1991; Clift and Vannucchi 2004; Scholl and VonHuene 2007). The Talara basin owes its present exposure to uplift relatedto post-Eocene slab shallowing (Pardo-Casas and Molnar 1987) andultimately Pliocene collision and partial subduction of oceanic ridges(Gutscher et al. 1999; Pedoja et al. 2006).

Today, flat-slab subduction and a volcanic gap characterize the Andesin the region of Talara (Bernal et al. 2002). Since convergence began inthe Cretaceous (Nurnberg and Muller 1991), arc-related volcanic activityhas been intermittent, depending on the nature (rate and direction) ofconvergence between the Farallon–Nazca and South American plates(Pardo-Casas and Molnar 1987; Jaillard et al. 1995; Noblet et al. 1996;Silver et al. 1998). Specific timing of volcanic activity versus quiescence isnot well documented for this segment of the Andes. However, LateCretaceous through Miocene calc-alkaline, arc-related plutonic andvolcanic rocks occur along the western flank of the Andean Cordillera(Fig. 1B), with the few reported radiometric ages clustering at 23–28 Ma(Bristow and Hoffstetter 1977), 35–39 Ma (Noble et al. 1990; Steinmann1997), 43–55 Ma (Bristow and Hoffstetter 1977; Noble et al. 1990;Steinmann 1997), 61–68 Ma, as well as 73, 83, 93, and 110 Ma (Bristowand Hoffstetter 1977).

Stratigraphy

The Talara basin contains 9000 meters of fluvial to deep-marine stratathat accumulated while the basin deepened by 1000 meters during themiddle Eocene (Fig. 1D; Fildani et al. 2008). The oldest alluvial–fluvialMogollon Formation rests unconformably on Mesozoic and oldermetasedimentary and granitoid rocks of the Amotape block. Upsection,the marginal-marine Pale Greda and Echinocyamus formations includeminor unconformities and fluctuating shallow-water depths (, 200 m),including the distinctly deltaic Cabo Blanco Member.

These strata are capped by a major unconformity, above which weredeposited the middle Eocene deepwater (. 1000 m) Talara and Verdunformations. The Verdun Formation caps the Eocene section. Thecomplex, tectonostratigraphic history of the Talara basin is described inFildani et al. (2008).

Candidate Source Terranes

The source areas for Talara detritus, while spatially constrained bypaleocurrent analysis, are not well studied due to difficult field access.One of the goals of this work is to use DZ to better understand thegeochronology of Andean and pre-Andean terranes surrounding theTalara basin.

Sediment dispersal into the Talara basin was to the northwest duringdeposition of the Mogollon Formation through the Helico Member,whereas the overlying Verdun Formation shows consistent southwest-ward paleocurrents (Fig. 1C; Fildani et al. 2008). Therefore, the earlyTalara basin tapped into the Amotape block and, depending on localrelief, the Lancones basin and Cordillera farther to the east-southeast(Fig. 1C). The later Talara basin received sediment from the Amotape–Raspas block and Cordillera to the northeast, as well as intra-formationaluplifts of the Verdun Formation itself (Fildani et al. 2008).

Conglomerate clast counts and sandstone petrography (Fildani et al.2008) suggest the following zircon-bearing source rocks: quartzite, schist,metasandstone–metasiltstone, intermediate to felsic volcanic rock, tuff,granite, and recycled sandstone. Sandstone source rocks could alsoinclude uplifted and eroded Talara or Lancones sedimentary beds.Volcanic source rocks would have occurred as arc-related flows and tuffsin the Cordillera and the Lancones basin, and to a lesser degree asvolcanic or metavolcanic components of the pre-Andean basement in the

648 A.M. HESSLER AND A. FILDANI J S R

Cordillera and Amotape block. Granite and metamorphic source rocksoccur as part of pre-Andean basement in the Amotape block andCordillera.

Direct radiometric constraints for pre-Andean basement are lacking innorthern Peru and southern Ecuador. However, there are DZ ages fromPaleozoic metasedimentary units in the Cordillera east and southeast ofthe Talara basin (Fig 1B; Chew et al. 2007, 2008; Cardona et al. 2010) andzircon ages from Permo–Triassic granitoids just south of the Huanca-bamba Deflection (Fig. 1B; compiled from Miskovic et al. (2009) inCardona et al. (2010)). These zircon populations suggest that thefollowing remnants of pre-Andean magmatic and metamorphic basementoccur in the Cordillera: (1) Mesoproterozoic (0.9–1.2 Ga); (2) Ordovicianto Neoproterozoic (450–700 Ma); (3) Carboniferous to Permian (280–350Ma); and (4) Permian to Triassic (190–260 Ma). Geochemical andradiometric constraints on pre-Andean granitoids and metamorphicrocks from the Amotape block point to its affinity to the main Cordillera(Noble et al. 1997); therefore, we do not attempt to distinguish betweenthe Amotape block and the Cordillera as contributors of pre-Andean DZin the Talara basin.

DETRITAL-ZIRCON METHODS

Detrital zircon (n 5 267) was extracted from five sandstone samples,each from a unique formation and covering diverse Eocene depositionalenvironments preserved in the Talara basin (Fig. 1D). The sampledformations are, from oldest to youngest: (M) alluvial to fluvial MogollonFormation; (PG) marginal marine Pale Greda Formation; (CB) deltaicCabo Blanco Formation; (H) mid-bathyal Helico Formation; and (V)bathyal Verdun Formation. Sample preparation procedures are outlinedin DeGraaff-Surpless et al. (2002).

Uranium–lead ages were determined using the Sensitive HighResolution Ion Microprobe–Reverse Geometry (SHRIMP-RG) at theStanford–U.S. Geological Survey Microscopic Analytical Center (SU-MAC). We analyzed 50–60 grains per sample, to achieve . 90%probability of identifying one representative grain for each subpopulationmaking up more than 5% of the total population (per Dodson et al. 1988;Andersen 2005).

U-PB AGE DISTRIBUTIONS

207Pb-corrected 206Pb/238U populations for each sample are presentedin Figure 2. All samples are dominated by a peak between 30 and 110 Ma(Andean source), and include lesser peaks at , 250 Ma, , 550 Ma,, 1100 Ma (pre-Andean source), and a scattering of Amazonian(. 1250 Ma) ages. The abundance of Andean grains is highest for theMogollon Formation and Helico Member. Pre-Andean and Amazoniagrains are most abundant in the Pale Greda Formation and Cabo BlancoMember. The Verdun Formation contains moderate populations of bothAndean and pre-Andean grains.

An expanded view of the Andean (, 110 Ma) population shows thefollowing pattern common to all samples: (1) abundant between 110 and80–75 Ma; (2) a clear gap between 70 and 65 Ma; (3) an increase between60 and 50 Ma; and (4) a decline from 45 to the youngest grain at 35 Ma.

The youngest measured DZ age for each sample is consistent withpublished depositional age (Fildani et al. 2008), with the exception of theHelico Member and the Verdun Formation, where the youngest DZ age isyounger than published depositional age. This result encourages a re-interpretation of depositional age for these units, as presented in theDiscussion.

CALIBRATION TO REGIONAL U-PB STUDIES

The DZ populations here, despite the proximity of the Talara basin toan active volcanic arc, include solid representation from pre-Andean

(. 110 Ma) sources, including prominent peaks at , 250, , 550, and, 1100 Ma, as well as a cluster of Amazonian (. 1250 Ma) ages(Fig. 3A). Ours are the first Andean (, 110 Ma) DZ ages in the region,but we can calibrate the older, pre-Andean and Amazonian populationswith zircon datasets based on Paleozoic to early Mesozoic metasedimen-tary and igneous rocks of the Cordillera that occur to the east andsoutheast of the Talara basin (Fig. 3B; Chew et al. 2007, 2008; Miskovicet al. 2009; Cardona et al. 2010).

Together, these data help us understand persistent age populations inbasement rocks of the Andean Cordillera and the Amotape terrane. Wedescribe our data as they correspond (or do not) with major tectonicphases observed in regional zircon populations and known from previouswork, in the context of the following supercontinent cycles: the assemblyof Amazonia (3000–1250 Ma), the Rodinia cycle (1250–625 Ma), and theGondwana–Pangea cycle (625–110 Ma).

Assembly of Amazonia (3000–1250 Ma)

This oldest population (Fig. 3A) originally derived from periods ofsignificant growth of the continent Amazonia (Fig. 1A; see review inCordani and Sato 2000), whose westernmost extension is thought toreside beneath the Andes, based on inherited ages from Cordilleranbatholiths and metasedimentary rocks (Chew et al. 2008; Miskovic et al.2009; Cardona et al. 2010). By , 1250 Ma, western Amazonia (whichwould have included basement of the Andean Cordillera in northernPeru) was nearing collision with eastern Laurentia to form thesupercontinent Rodinia (Tohver et al. 2005; Chew et al. 2008).

Rodinia Cycle (1250–625 Ma)

This population records the period when western Amazonia was thesite of continental collision during the assembly of Rodinia, and was alsothe site of rifting as Rodinia broke apart in the late Precambrian. The ca.1250–950 Ma cluster (Fig. 3A) derives from the Grenville–Sunsasorogeny, when western (then northern) Amazonia collided with eastern(then southern) Laurentia (Dalziel 1991; Hoffman 1991) to form Rodinia.The peak is prominent in DZ populations from Paleozoic metasedimen-tary rocks in the region (Fig. 3) and also corresponds with episodesduring the Sunsas orogeny constrained by basement exposure insouthwest Amazonia (Tohver et al. 2005) and the North AmericanGrenville Province (Rivers 2008). Relative abundance of Grenvillian agesin Cordilleran basement likely relates more to the high zircon content ofGrenville rocks, rather than to the spatial abundance of Grenvillebasement in northwest Peru (Moecher and Samson 2006).

After a 150-million-year gap (995–843 Ma), the 850–725 Ma peak maymark initial separation of Laurentia and Amazonia during the breakup ofRodinia, which has been constrained by exposures in Argentina (Omariniet al. 1999). This population is not observed in Permo-Triassic granitoidsof Peru’s Eastern Cordillera (Miskovic et al. 2009), but does appear in DZsignatures of Paleozoic metasedimentary rocks in the region (Fig. 3B).

The ca. 685–625 Ma cluster is characterized by increasingly abundantzircon (Fig. 3A). This age population is neither common nor wellunderstood in the northern Andes (Chew et al. 2007) but may correspondwith felsic extensional magmatism known to have occurred around thesame time in the Arequipa–Antofalla terrane of southern Peru (Loewyet al. 2004).

Gondwana–Pangea Cycle (625–110 Ma)

This population records the long-lived active western margin ofGondwana through its assembly and breakup, as well as the interveningassembly and breakup of Pangea. There is no known arc-intrusive sourcefor the 600–550 Ma peak in the north-central Andes, but the population isstriking in Paleozoic metasedimentary rocks of the Eastern Cordillera and


Olmos terrane and may have been sourced by the buried northernextension of the well documented Sierra Pampeanas arc intrusives innorthern Argentina (Chew et al. 2008).

A peak between 488 and 459 Ma (Fig. 3A) supports existence ofa Famatinian arc in the north-central Andes (Chew et al. 2008). However,unlike other studies in the region (e.g., Chew et al. 2008; Miskovic et al.2009), we do not report a long-lived (, 100 My) magmatic gap after theFamatinian, but note a steady supply of zircon between ca. 450 Ma andca. 350 Ma. DZ profiles from Permo-Triassic and older rocks in theOlmos terrane (Cardona et al. 2010) include grains from this period,suggesting that western Gondwana was not a passive margin, as proposedby Bahlburg and Herve (1997), but may have shifted from orthogonalconvergence to an oblique or even transform plate margin.

A 348–322 Ma peak (Fig. 3A) reflects the onset of the pan-PacificGondwanide orogeny (Chew et al. 2008; Miskovic et al. 2009), followedby late Pennsylvanian magmatic quiescence (ca. 310 Ma) that could havebeen caused by an increase in plate convergence and slab shallowingrelated to the Pangea-forming collision between northern Gondwana,southeastern North America, and western Africa (see Miskovic et al. 2009and references herein).

Magmatism flared up again during the Permo-Triassic transition (285to 220 Ma) in the northern proto-Andean region (Fig. 3A), this timerelated to strike-slip extension (Miskovic et al. 2009). The whole westernmargin of Gondwana appears to have been active through this period,as suggested by Early Permian to Late Triassic batholiths along theChilean Frontal Cordillera (Herve et al. 2014), the reported Permianvolcanism of the Choiyoi province of Argentina (Rocha-Campos et al.2011), and Permian to Triassic ashfalls straddling South America intosouthern Africa (Lopez-Gamundı 2006; Lopez-Gamundı et al. 2013;McKay et al. 2015).

Following peak magmatism at , 240 Ma, felsic to intermediatemagmatism gave way to more mafic melts as rifting evolved (Miskovic etal. 2009), thus giving rise to the Late Triassic–Early Jurassic gap betweenca. 215 to ca. 195 Ma (Fig. 3A).

By the Early Jurassic, ca. 190 Ma, a narrow ocean separated North andSouth America, and the latter began to drift away from southern SouthAfrica, setting up oblique ocean–continent convergence along westernGondwana (Nurnberg and Muller 1991). In the Talara basin, a few grainsof Jurassic zircon herald renewed arc-related magmatic activity (Fig. 3A).These are coeval with Middle Jurassic intrusive and extrusive rocksrelated to backarc extension in the Eastern Cordillera (184 Ma, 172 Ma(U–Pb); Miskovic et al. 2009).

A paucity of ages between 170 and 110 Ma is noted in our data. Wemeasured one Late Jurassic age (156 Ma), which roughly coincides witha single Late Jurassic to Early Cretaceous magmatic age reported byNoble et al. (1997) in the Eastern Cordillera of southernmost Ecuador(143 Ma (Pb–Pb)). The 156–110 Ma gap is consistent with the generallack of evidence for magmatism during this time in northwest Peru.

A PRE-ANDEAN SOURCE FOR ZIRCON

Basement Rocks of the Andean Cordillera and Amotape

The similarity between DZ profiles from the Talara basin and nearbyAndean basement points to the Andean Cordillera as a likely source forpre-Andean DZ in the Eocene forearc (Fig. 3A). Because the Amotapeterrane is considered a para-autochthonous sliver of Cordilleran

basement (Mourier et al. 1988; Aspden and Litherland 1992), we assumeits pre-Andean zircon population to be comparable to that of Cordilleranbasement, and that the Amotape is equally likely to have contributedolder zircons to the Talara basin, depending on its exposure duringdeposition.

Whether the pre-Andean zircon was derived directly from magmaticsource rocks or was recycled from older sedimentary or metasedimentaryrocks relates in part to known exposures of age-specific rock types in theCordillera and Amotape. In particular, Triassic igneous rocks are welldocumented in the Amotape and Olmos terranes as well as the Cordillera.Noble et al. (1997) measured Triassic Pb-Pb zircon–monzonite ages(227.3 6 2.2 Ma; 227.5 6 0.8 Ma) for granitoid rocks in the Amotape andCordillera, and Chew et al. (2008) determined a Triassic U-Pb age (230.46 3.2 Ma) for migmatite in the Olmos terrane. Carboniferous igneousrocks are not known in the Cordillera near Talara but are commonlyexposed in the northern Maranon Complex just south of the Huanca-bamba Deflection (U–Pb zircon ages: 309–320 Ma; Miskovic et al. 2009).Abundant Triassic DZ (and likely the less common Carboniferous DZ) inthe Talara basin could have been derived directly from granitoidexposures in the Amotape and Cordillera.

As for the prominent , 550 and , 1100 Ma DZ population in theTalara basin, there are no known Neoproterozoic exposures in thevicinity, although the abundance of both populations in nearby Paleozoicmetasedimentary rocks suggests that Neoproterozoic rocks may be buriedbeneath the Andes, as possible extensions of the Sierra Pampeanas (, 550Ma) and Sunsas (, 0.9–1.3 Ma) orogenies exposed in northern Argentinaand southwestern Brazil, respectively (Chew et al. 2007). While it ispossible that the Talara basin tapped directly into Neoproterozoic rocksexposed during the Eocene, it seems more likely that these DZpopulations were recycled from Paleozoic metasedimentary rockscommon to this part of the Andean Cordillera (Fig. 1B).

Likewise, there are no known older Proterozoic or Archean rocks in theCordillera or Amotape. The presence of inherited Amazonian (. 1250Ma) zircon in Cordilleran batholiths and metasedimentary rocks impliesthat the westernmost boundary of the Amazon craton exists beneath thispart of the Andes (Chew et al. 2008; Miskovic et al. 2009; Cardona et al.2010). Perhaps the original craton was exposed in the Cordillera duringthe Eocene, to provide Amazonian zircon to the Talara basin. Morelikely, however, these grains were recycled from the Paleozoic meta-morphic rocks of the Cordillera and Amotape, and also from inheritedzircon known in Carboniferous–Triassic granitoids (Noble et al. 1997;Miskovic et al. 2009).

A Reversed Proto-Amazon Drainage?

The presence of Amazonian DZ in the Eocene forearc of Peru can beexplained in part by the erosion and recycling of zircon from Paleozoicmetasedimentary rocks (Chew et al. 2007, 2008; Cardona et al. 2010) andCarboniferous–Triassic granitoids (Noble et al. 1997; Miskovic et al.2009) that occur in the Andean Cordillera and Amotape, depending ontheir exposure during the time of deposition. However, we raise a secondpossibility: that a westward, continental-scale drainage may havedelivered Amazonian zircon directly from the shield to the forearc regionthrough breaks in the low-lying Cordillera. We consider our data in lightof regional observations that suggest (1) a large-scale westward drainage(Costa et al. 2001; Ruiz et al. 2004; Mapes 2009; Horton et al. 2010) and

rFIG. 2.—Kernel density, frequency and probability curves (Ludwig 2003; Vermeesch 2012) for 207Pb-corrected 206Pb/238U ages of detrital zircon from each sampled

interval. Inset plots show the Andean (, 110 Ma) population and include youngest DZ single-grain (YSG) age and published maximum depositional age (solid grayvertical bar; from Fildani et al. 2008). Concordia diagrams are included for each sample.


FIG. 3.—A) Compiled 207Pb-corrected 206Pb/238U DZ ages for this study. Major tectonic cycles affecting northwest Peru are indicated by dotted bars. Inset highlightsthe Andean (, 110 Ma) population; horizontal gray bar marks period of sedimentary filling of the Talara basin. B) Previously published 206Pb/238U zircon agepopulations from metasedimentary and igneous pre-Andean basement rocks of the Cordillera in the vicinity of the Talara basin (see Fig. 1 for locations). CA, Cardonaet al. (2010), where Permo-Triassic igneous rocks are compiled from Miskovic et al. (2009); C1, Chew et al. (2008); C2, C3, Chew et al. (2007).


(2) a lower-relief Cordillera through the early Cenozoic (Gregory-Wodzicki 2000; Horton et al. 2010).

From the Cretaceous into the Miocene, a reverse ‘‘proto-Amazon’’system may have developed in response to opening of the equatorialAtlantic and elevation of rift shoulders along eastern South America,which gave a westward tilt to the new continent (Costa et al. 2001; Potter1997). Fold–thrust–belt stratigraphy in Colombia (e.g., Horton et al.2010) demonstrates a Late Cretaceous shift in large-scale drainagepatterns, from predominately east-to-west (tapping the Guyana shield) towest-to-east (tapping the Cordillera). Large-scale, eastward drainages offthe eastern Cordillera (i.e., the modern Amazon) may not have beenestablished until the mid-Miocene (McDaniel et al. 1997), and these wereconfined to an interior wetland by the Purus arch (Fig. 1A) until itsbreach in the late Miocene, based on DZ trends along the Amazon trough(Mapes 2009) and in the foreland of Ecuador and Colombia (Ruiz et al.2004; Horton et al. 2010), as well as stratigraphic comparisons betweenwestern and eastern Amazonia (Cunha et al. 1994; Eiras et al. 1994). Thefirst appearance of Andean-derived grains in the Atlantic Amazon fan isdated to , 11 Ma (McDaniel et al. 1997; Figueiredo et al. 2009) to mimicthe modern Amazon drainage.

For large-scale rivers to have delivered sediment directly from theshield to the Cretaceous–Paleogene forearc, these drainages would havehad to traverse the Andean Cordillera, an active volcanic arc at the time.However, during the Cretaceous and early Cenozoic, the centralCordillera may have been low-lying, even at sea level in some regions(Gregory-Wodzicki 2000); accelerated uplift of the northern Andes didnot occur before the Oligocene (Horton et al. 2010). Therefore, there mayhave been opportunity for topographic breaks, like the modernHuancabamba deflection, through which fluvial systems could transportcratonic sediment into the forearc region (i.e., the Cretaceous Lanconesand/or the Eocene Talara basins). In this case, we would expectAmazonian DZ in the Talara basin to be proportionately high comparedto the Amazonia DZ in potential source rocks like the Paleozoicmetasedimentary rocks studied by Chew et al. (2007, 2008). For thePaleozoic metasedimentary rocks (Fig. 3B), the Amazonian (n) zirconpopulation (n 5 34) is 15% of the total (N) zircon population (N 5 223);the proportions for individual samples range from 27% (C1; n 5 26;N 5 96) to 3% (C2; n 5 2; N 5 66) to 10% (C3; n 5 6; N 5 61).

For comparison, we calculate the proportion of Amazonian DZ to thetotal Talara DZ population that is older than 330 Ma (i.e., the full agespectrum available in the Paleozoic metasedimentary rocks). Takentogether, the Amazonian population for all Talara samples is 24%(n 5 29) of the total (N 5 122; Fig. 3A). If we consider each Talara sampleseparately (Fig. 2), the Amazonian DZ populations are: 14% Mogollon(n 5 3; N 5 22); 16% Pale Greda (n 5 5; N 5 31); 33% Cabo Blanco (n 5 10;N 5 30); 40% Helico (n 5 8; N 5 20); and 16% Verdun (n 5 3; N 5 19). Theproportion of Amazonian DZ in the Mogollon, Pale Greda, and Verdunformations is in line with that calculated for the Paleozoic metasedimentaryrocks (, 15%). On the other hand, the Cabo Blanco and Helico membersinclude a higher proportion of Amazonian DZ (33–40%) compared witheven the most ‘‘Amazonian-rich’’ Paleozoic rock (C1; 27%). If theenrichment of Amazonian DZ is real, then perhaps the Cabo Blanco andHelico members include a pulse of shield-derived rather than recycledAmazonian DZ, delivered by westward drainages through gaps in thevolcanic arc. However, the current dataset does not allow us tounequivocally promote the idea of a westward, trans-Cordillera drainageover simple recycling of old zircon from Paleozoic metasedimentary rocks(and Carboniferous–Triassic granitoids) of the Cordillera and Amotape.

THE ANDEAN (, 110 MA) DZ POPULATION: REGIONAL TECTONICS

Between 110 and 104 Ma, the ‘‘Equatorial Gateway’’ opened as theSouth Atlantic became a full-fledged ocean basin (Nurnberg and Muller

1991). The final separation of South America from Africa occurred by 100Ma (Granot and Dyment 2015) and set up South America’s westward driftand accelerating convergence with the Farallon plate (Pardo-Casas andMolnar 1987). Early Cretaceous (110–100 Ma) volcanic rocks in theLancones basin (Fig. 1B) include back-arc or wrench extensional felsic tomafic rocks, shifting to continental-arc intermediate igneous rocks at theEarly–Late Cretaceous boundary (, 100 Ma) (Winter et al. 2010) as theLancones and areas west became part of the Andean forearc proper.Cretaceous DZ in the Talara basin includes two presumably ‘‘back-arc’’ages (100–110 Ma) followed by steadily increasing numbers of ‘‘continen-tal-arc’’ ages (, 100 Ma) (Fig. 3A). Continental-arc magmatism innorthern Peru may have preceded arc magmatism to the north;compressional tectonics and subduction-related magmatism in theColombian Andes appear to have started in the Campanian (Hortonet al. 2010).

The Caribbean Oceanic Plateau (COP) collided with and partlyaccreted onto Ecuador and Colombia sometime in the Late Cretaceous,although the timing for this event (or series of events) is cloudy. It hasbeen associated with a period of rapid cooling in the Amotape Complexand Andean Cordillera of Ecuador between 85 and 80 Ma (Hughes andPilatasig 2002; Spikings et al. 2005). However, our data show increasingDZ from 90 to 70 Ma, and no important shift at , 85 Ma. The notablelull between 70 and 66 Ma could reflect an accretion event and aligns withrecent stratigraphic and COP geochemical work (Toro Alava and Jaillard2005), which suggests latest Campanian and/or Maastrichtian (75–65 Ma)accretion. Our data might constrain this event closer to 70 Ma,concurrent with initiation of the Dolores–Guayaquil fault system(Fig. 1A; Shepherd and Moberly 1981).

By the time the Talara basin began subsiding in the late Paleocene(, 58 Ma), the Farallon plate began to converge in a less obliquedirection with respect to South America (Pardo-Casas and Molnar 1987),at a slightly faster rate (Somoza and Ghidella 2012). Subductionmagmatism revived in northern Peru, as evidenced by coeval intermediateigneous rocks identified along the Cretaceous–Tertiary volcanic arc(Fig. 1B; Bristow and Hoffstetter 1977; Noble et al. 1990; Steinmann1997) and the notable DZ populations identified in the Talara forearcbasin (Fig. 3A).

DYNAMICS OF THE EOCENE ANDEAN FOREARC

Up-section variation of DZ spectra in a forearc basin is a reflection ofdynamic paleogeographic adjustment (i.e., subsidence, uplift) andvolcanic activity, themselves the response to changes in plate convergenceand slab orientation. Here, we interpret the up-section variation of DZ inthe Eocene Talara basin with respect to: (1) depositional environment andsource area; (2) syndepositional magmatism and chronostratigraphy; and(3) regional tectonic drivers.

Depositional Environment and Source Area

In the Talara basin, DZ populations clearly vary upsection (Fig. 2).Like any single provenance indicator, DZ offers a filtered view of thesource area. Therefore, we consider upsection DZ variation not inisolation, but together with prior provenance work using paleocurrents,conglomerate clasts, sandstone petrography, and shale geochemistry(Fildani et al. 2008) to understand where DZ variation indicates a changein source area, or could be dictated in part by depositional environmentand grain-size fractionation.

Mogollon Formation.—In the case of the alluvial–fluvial MogollonFormation, DZ ages are likely self-limited by depositional environmentgrain-size fractionation. The Mogollon is dominated by Andean(, 110 Ma) zircon with few older grains. However, quartzite is the mostcommon clast type in Mogollon conglomerate (, 80%; Fildani et al. 2008),


and metamorphic lithic grains are less common in Mogollon sandstone (5–14%; Fildani et al. 2008), suggesting that the Mogollon Formation waspartially derived from pre-Andean source rock but that those moreresistant rocks did not readily yield sand grains, including DZ. Because theMogollon Formation is a short-transport, alluvial–fluvial sequence, therarity of pre-Andean DZ could be more a function of inhibitedsedimentary fractionation of resistant pre-Andean rock types (i.e.,quartzite), not necessarily a lack of pre-Andean rocks in its source area.

Prior provenance work has suggested that during deposition of theMogollon Formation, detritus was sourced by the Lancones basin andAndean Cordillera to the east (Fig. 1C; Fildani et al. 2008). TheMogollon Formation unconformably overlies Cretaceous and older unitsof the central and lowermost Amotape Mountains, implying that theAmotape were low-lying and not a barrier to sediment coming from theLancones or farther east. Conglomerate clasts, sandstone petrography,and shale geochemistry all point to mixed sources for the MogollonFormation, including felsic to intermediate arc (i.e., Andean) andrecycled-orogen (i.e., pre-Andean, Amazonian) rocks. The prominentAndean DZ population in the Mogollon Formation could have derivedfrom: (1) intermediate Cretaceous lava flows in the Lancones basin(Winter et al. 2010); (2) syndepositional (Eocene) volcanic ashfalls intothe Talara or Lancones basins; and (3) the Cretaceous–Tertiary Andeanarc itself (Fig. 4).

Pale Greda Formation.—The DZ spectrum for the overlying PaleGreda Formation is more balanced between Andean and pre-Andeanages, compared with the Mogollon Formation (Fig. 2). The Pale Gredarests conformably on the Mogollon and is marked by the oldestoccurrence of marine fossils; it represents a period of relative sea-levelrise in this part of the basin and deposition of shallow-marine shale,siltstone, and intercalated sandstone and coquina beds. Sandstonepetrography suggests more influence from ‘‘recycled-orogen’’ sourcesthan for the Mogollon Formation, but paleocurrent direction wasunchanged between the two formations (both west- to northwest-directed) and the Amotape terrane was not yet a topographic high. Toaccount for the increase in recycled-orogen sandstone and correspondingpre-Andean DZ in the Pale Greda, compared with the underlyingMogollon, two processes may have been at work. First, the shallow-marine environment may have allowed advanced mechanical breakdownof resistant, recycled-orogen rocks (i.e., quartzite) into sand or smallergrains, thereby increasing the proportion of pre-Andean DZ in PaleGreda sandstone. Second, rising sea level could have pushed headwatererosion eastward beyond the Lancones basin and into the AndeanCordillera, to tap directly into pre-Andean basement (Fig. 4).

Cabo Blanco Member.—The Cabo Blanco Member is a single fluvio-deltaic unit within the sequence of stacked, prograding fluvial to shallow-marine deposits that unconformably overly the Pale Greda Formationand make up the Echinocyamus Formation (Fig. 1D). The Cabo BlancoMember sandstone is quartz-rich with feldspars, metamorphic lithicgrains, and few lithic volcanic grains. Fildani et al. (2008) concluded thatafter deposition of the Pale Greda Formation, the Amotape Mountainswere uplifted to become a proximal source for the fluvio-deltaic depositsof the Cabo Blanco Member (Fig. 4). This is a significant source-areashift that is not clearly reflected in the DZ populations of the CaboBlanco Member versus the underlying Pale Greda Formation (Fig. 2).There are two likely explanations for the similarity. First, deltas (and theirfeeder rivers) of the Cabo Blanco Member could have reworked sedimentof the underlying Pale Greda Formation, which may have been partly

uplifted along with the Amotape terrane. Second, the Amotape terranewas the proximal source for the Cabo Blanco Member; as noted above,the Amotape terrane is considered compositionally and geochronologi-cally equivalent to the pre-Andean basement of the Cordillera, which waslikely a significant source for Pale Greda sediment. Interestingly, theAndean DZ population in the Cabo Blanco Member is strongly weightedtoward the older (, 75 Ma) peak (Fig. 2), because the uplifted Amotapeserved as a barrier to the younger, syndepositional volcanic rocks of theLancones basin and Cretaceous–Tertiary arc (Fig. 4).

Helico Member.—The spectrum of DZ for the Helico Member isunique, with a strong Andean population (including a prominentsyndepositional peak at , 45 Ma) and a minor pre-Andean component(Fig. 2). The Helico Member is a mid-bathyal turbiditic sandstone andconglomerate sequence deposited soon after the middle Eocene basin-wide deepening event. Like the underlying formations, paleocurrents aretoward the northwest (Fig. 4; Fildani et al. 2008).

Helico Member conglomerate is rich in quartzite and intermediatevolcanic clasts, the sandstone has strong volcanic-arc affinity withabundant intermediate lithic volcanic grains, and shale geochemistryshows moderate LREE/HREE fractionation typical of intermediatevolcanic source rocks (Fildani et al. 2008). All provenance data pointto a source terrane dominated by Andean arc volcanics with only minorinput of pre-Andean detritus. In contrast to the paleogeography of theCabo Blanco Member, the Amotape terrane does not appear to have beena topographic barrier to sediment transport from the Lancones basin andCretaceous–Tertiary volcanic arc during deposition of the HelicoMember (Fig. 4). Deepening of the basin would have dissected theAmotape and forced headwaters into the Lancones basin and perhapsback into the Cordillera to tap arc-related volcanic rocks and to alesser extent, pre-Andean basement. The oldest Andean DZ population(, 90–110 Ma) is more prominent in the Helico Member than in othersamples (Fig. 2) and may represent evolved erosion into older volcanicflows of the Lancones basin and arc itself.

Verdun Formation.—The DZ spectrum for the Verdun Formation issimilar to that of the underlying Helico Member, aside from a slightlydiminished Andean population, particularly in terms of the older(, 75 Ma) population (Fig. 2). The similarity is notable because, whilethe Verdun Formation is also a mid-bathyal turbiditic sandstone andconglomerate sequence, it was deposited after a significant (, 1 My)depositional hiatus, signifies a major change in paleo-dispersal to thesouthwest, and lacks the volcanic detritus so prevalent in the underlyingHelico Member (Fildani et al. 2008).

Conglomerate clasts, sandstone petrography, shale geochemistry, andpaleodispersal patterns all point to a northern source terrane rich inquartzite and granite, most likely the Amotape–Raspas terrane (Fig. 4).The pre-Andean DZ population of the Verdun Formation, therefore,derived from the Permo-Triassic granitoids and older metasedimentaryrocks in that part of the Amotape block (Noble et al. 1997). Before andduring deposition of the Verdun Formation, the northern Amotape blockappears to have been uplifted along with intrabasinal units like theVerdun itself, based on the occurrence of intraformational clasts andgrains in the Verdun Formation (Fildani et al. 2008). The Andean DZpopulation may have been: (1) recycled from older formations like theHelico Member exhumed during uplift of the northern Amotape block;(2) derived from ashfall deposits related to coeval volcanic activity in thearc; and/or (3) eroded and transported from the volcanic arc into the‘‘Verdun Deep’’ via transverse tributary drainages (Fig. 4).

RFIG. 4.—Paleogeographic model and provenance highlights for each sampled unit. Black arrows mark proposed sediment routes, and brown arrows indicate regions of

uplift and/or subsidence.


Syndepositional Magmatism and Chronostratigraphy

Abundant Andean DZ in the Eocene Talara basin implies that theCretaceous–Tertiary volcanic arc was active in this region, beginning inthe Late Cretaceous and continuing through the end of Talara basindeposition (Fig. 3). That the Andean DZ derived largely from a volcanicarc, rather than from volcanism related to crustal extension related towrench tectonics, is supported by (1) the occurrence of coeval calk-alkaline volcanic rocks in the Cordillera (Bristow and Hoffstetter 1977;Noble et al. 1990; Steinmann 1997) and (2) the prevalence of intermediatevolcanic and plutonic clasts and grains in Eocene conglomerate andsandstone as well as moderately fractionated REE in Eocene shales of thestudy area (Fildani et al. 2008). Overall, the Andean DZ populationsuggests that zircon-yielding magmatism was not constant but fluctuated,peaking twice at , 75 Ma and , 45 Ma, with a , 10 My period ofquiescence between.

The younger population (, 35–56 Ma) corresponds with Eocenestratigraphy of the Talara basin and demonstrates a progressive upsectionshift toward younger DZ ages (Fig. 2). We consider the younger DZpopulation, therefore, to be ‘‘syndepositional’’ and to derive frommagmatic rocks that crystallized contemporaneously with Eocene forearcdeposition. Only the Cabo Blanco Member lacks syndepositional DZ; thepaucity does not appear to represent a magmatic gap, as the overlyingHelico Member and Verdun Formation contain DZ that crystallizedduring deposition of the Cabo Blanco Member (, 46–47 Ma). Instead,young volcanic detritus from the arc was sequestered behind the AmotapeMountains, which became a topographic barrier during Cabo Blancodeposition (Fig. 4). For the remaining sampled units, however, thepresence of syndepositional DZ prompts a review and refinement of priorchronostratigraphy for the Eocene section and suggests a significantrevision for maximum depositional age for the youngest units (the HelicoMember and Verdun Formation).

Revision of Maximum Depositional Age.—Eocene chronostratigraphyfor the Talara basin (Fig. 1D) is based on biostratigraphy (i.e.,palynology, calcareous nannoplankton, benthonic forams) with varyingdegrees of control (Fildani 2004; Fildani et al. 2008). Using availablebiostratigraphic data, a best estimate for the depositional age range ofeach unit is marked on Figure 2, as well as the ‘‘youngest single grain’’(YSG) U–Pb age. We also calculate the ‘‘youngest 1s grain cluster (n $

2)’’ (YC1s2+) for each formation as an additional proxy for maximumdepositional age, using the approach of Dickinson and Gehrels (2009).

The age of the Mogollon Formation does not require revision, as theYSG (51.6 6 1.0 Ma) and YC1s2+ cluster (52.6 6 2.2 Ma; n 5 5) agreewith previously defined maximum depositional age of , 50–52 Ma. Theoverlying Pale Greda Formation shows close agreement between YSG(48.3 6 0.9 Ma), YC1s2+ cluster (48.7 6 1.0 Ma, n 5 2), andbiostratigraphy (, 50 Ma). The YSG DZ for the Cabo Blanco Member isconsiderably older than its depositional age, and therefore is not used torevise maximum depositional age.

The Helico Member has a biostrat-constrained depositional age of , 43Ma (6 , 0.5 Ma; Fig. 2). Notably, this unit, as well as the immediatelyoverlying and underlying units, are constrained only to the middle Eocene(37.8–47.8 Ma; ICS chronostratigraphy of Cohen et al. 2013) based oncalcareous nannoplankton (Fildani et al. 2008), and so the , 43 Maestimate in Figure 2 is based on the stratigraphic placement of the HelicoMember within the middle Eocene section. The YSG U–Pb age is39.7 6 1.0 Ma, while the YC1s2+ cluster includes two older grains andgives another measure for maximum depositional age of 43.6 6 1.9 Ma.Use of the YSG age for maximum depositional age requires an , 3 Mychronostratigraphic shift for the Helico Member (and immediatelyoverlying units) from the middle middle Eocene to the late middle Eocene,while the YC1s2+ cluster agrees with the current biostrat-constrainedestimate of maximum depositional age for the Helico Member.

Finally, prior interpretation placed the Verdun Formation in the latemiddle Eocene Bartonian Age (37.8–41.2 Ma from ICS 2013; Fildaniet al. 2008). The YSG U–Pb age provides a measure of maximumdepositional age at 35.7 6 0.7 Ma (Fig. 2), while the next youngestYC1s2+ cluster that includes two grains with overlapping ages (based on1s) gives a much older age of 43.9 6 1.5 Ma, older even than theconservative estimate provided by the YC1s2+ cluster in the underlyingHelico Member. Use of the YSG pushes maximum depositional age ofthe Verdun Formation , 2 My from the middle to upper Eocene, roughlyequivalent to the shift required by using YSG in the underlying HelicoMember. On the other hand, using the YC1s2+ cluster requires a shiftback , 3 My to the middle middle Eocene, where the Helico Member iscurrently assigned. Based on the YSG constraint provided by the Helicosample (, 40 Ma), and because the Verdun Formation includes a40.0 6 0.6 Ma DZ that is not included in the YC1s2+ (lack of 1s-overlapwith next-oldest ages), we favor a maximum depositional age range forthe Verdun Formation of 37–40 Ma, approximately 1–2 million yearsyounger than ages based on biostratigraphy.

Whether the YSG U–Pb ages for the Helico Member (39.7 6 1.0 Ma)and the Verdun Formation (35.7 6 0.7 Ma) are viable measures ofmaximum depositional age could be clarified by a larger dataset,assuming that more DZ data will offer more young (, 40 Ma) ages.On the other hand, there may be a real scarcity of ages , 40 Ma thatreflects a magmatic slow-down caused by tectonic reconfiguration duringthe latter stages of Talara deposition.

Tectonic Drivers

The Eocene Talara basin experienced tectonic events specific to itsforearc position, each of which influenced paleogeography and theavailability of DZ: (1) pre–Upper Mogollon emergence due to platereconfiguration; (2) Pale Greda subsidence due to volcano-sedimentaryloading; (3) pre–Cabo Blanco SE Amotape uplift due to platereconfiguration; (4) pre-Helico subsidence due to terrane approach andwrench tectonism; and (5) Verdun subsidence due to terrane accretion tothe north and subduction erosion. We consider these tectonic events inlight of DZ patterns and forearc response to changing plate convergence.

Pre–Upper Mogollon (, 52 Ma) Emergence Due to Plate Reconfigura-tion.—The Upper Mogollon Formation (‘‘Mogollon Superior’’) com-prises arc- and basement-sourced alluvial to braided-fluvial deposits thatprograded over marginal-marine deposits. Volcanic activity in the arc wasat a peak after the lull between 70 and 60 Ma (Fig. 2) and, together withpresumed emergence of Cordilleran basement, set up the relief necessaryfor a period of uplift and erosion (unconformity at top of underlyingmarginal-marine units) and eventual progradation of Mogollon Superiorfluvial systems (Fig. 4). Both the increased volcanic activity and basementuplift were likely triggered by a change at , 58 Ma from N to NNEconvergence to more direct, NE convergence of the Farallon–Nazca platetoward South America (Pardo-Casas and Molnar 1987) at a slightly fasterrate (Somoza and Ghidella 2012). By this time, the Amotape terrane wasin place (though of low relief) between the Lancones and Talara basins(Fig. 4), having been sheared off and rotated away from the mainCordillera in the Late Cretaceous (75–65 Ma; Mourier et al. 1988; Aspdenand Litherland 1992).

Pale Greda (, 50 Ma) Subsidence Due to Volcano-SedimentaryLoading.—Marginal-marine to outer-shelf sediments of the Pale GredaFormation were deposited into a continually subsiding basin. During thisperiod, forearc subsidence was likely a response to volcanic and sedimentaryloading in the arc and forearc in addition to thermal flexure (Fig. 4), as thereare no indications for faulting, uplift, or other paleogeographic adjustmentssuggestive of compressional or wrench tectonism.

Pre–Cabo Blanco Uplift Due to Plate Reconfiguration (, 47 Ma).—Fluvio-deltaic progradation of the Cabo Blanco Member followed


a period of erosion (top of the Pale Greda Formation) associated withlocal uplift of the Amotape terrane to the southeast (Fig. 4). This eventcorresponds to a major reconfiguration of Farallon–South Americaconvergence at , 47 Ma, characterized by a notable rate increase anda nearly head-on approach in northern Peru (Somoza and Ghidella 2012).Greater trench-perpendicular stress in the forearc could have reactivatedwrench faults to uplift the Amotape terrane and segregated the Talarabasin from the Lancones basin and volcanic arc to the east. Notably, theDZ population in the Talara basin peaks around this time (45–50 Ma;Fig. 3). The subsequent decrease in DZ may result from changes in thevolcanic arc related to the increased rate of convergence, such as (1)migration of the arc to the east and/or (2) magmatic slow-down related toslab shallowing.

Pre-Helico Subsidence Due to Terrane Approach and Wrench Tectonism(, 43–40 Ma).—Deposition of the deepwater Helico Member followed. 1000 m of subsidence in the forearc (Fig. 1D; Fildani et al. 2008). Thisdeepening forced drainages eastward to tap the volcanic arc rocks of theLancones basin and Cordillera (Fig. 4) and perhaps even farther east tocapture shield-draining fluvial systems. Pre-Helico subsidence was likelythe result of wrench tectonism associated with the oblique approach ofthe Macuchi Arc in the north (Hughes and Pilatasig 2002), recorded bythe onset of cooling (i.e., exhumation) of the Amotape block at , 43 Ma(Spikings et al. 2005).

Verdun Subsidence Due to Terrane Accretion and Subduction Erosion(, 40–37 Ma).—Deposition of the Verdun Formation followed a , 1 Mydepositional hiatus, after which paleo-dispersal was predominantly to thesouth to fill a north–south–elongated, fault-bounded basin (Duerichen2005). To the north, the Macuchi Arc had collided with the Ecuadorianmargin (Hughes and Pilatasig 2002), and final uplift and cooling of thenorthern Amotape were occurring (, 39 Ma; Spikings et al. 2005),establishing this region as the source terrane for Verdun detritus.

Subsidence in the forearc after Amotape uplift, marked by the thick,basin-wide deepwater deposits of the Verdun Formation, is thought to belargely related to subduction erosion and resultant thinning of theoverriding plate (Fildani et al. 2008), as has been observed for the centralEcuador forearc (Sage et al. 2006), the Lima basin to the south (Clift et al.2003), and the modern Peruvian trench offshore Talara (Bourgois et al.2007). In the late Eocene, plate convergence remained head-on, but anincreased rate of convergence (Pardo-Casas and Molnar 1987) andshallowing of the Farallon slab could have exacerbated the processes ofsubduction erosion and forearc subsidence as the ‘‘Verdun deep’’continued to fill. With continued shallow-slab subduction, we wouldexpect decreased and/or eastward migration of arc magmatism, perhapsmanifest in the diminishing DZ after , 40 Ma (Fig. 2), and the eventualemergence and shoaling of the forearc that is documented for the upperVerdun Formation (Fildani et al. 2008).

CONCLUSIONS

Detrital zircon in Eocene strata of the Talara basin, combined withother provenance indicators, records paleogeographic shifts in theAndean forearc that are linked to larger-scale tectonic drivers such assubduction erosion, plate convergence rate, and regional accretion events.Pre-Andean DZ ages (. 110 Ma) corroborate recent geochronologicalcharacterization of Triassic and older basement in the Cordillera ofnorthern Peru but permit speculation regarding delivery of Amazonian(. 1250 Ma) zircon directly from the shield to the forearc through breaksin the Cordillera. The Andean DZ population (, 110 Ma) suggests peakarc activity at , 75 Ma and , 50 Ma, and a period of quiescence at , 70Ma linked to accretion of the Caribbean Oceanic Plateau. Syndeposi-tional volcanism throughout the Eocene allows revision of maximumdepositional age for upper Eocene units of the Talara basin. Forearcsubsidence and magmatic decrease in the late Eocene coincided with arc

collision to the north, uplift of the Amotape block, subduction erosion,and slab shallowing.

SUPPLEMENTAL MATERIAL

A dataset is available from the JSR Data Archive: http://sepm.org/pages.aspx?pageid5229.

ACKNOWLEDGMENTS

We are grateful to Grand Valley State University and to the StanfordProject on Deepwater Depositional Systems (SPODDS) for funding. Manythanks go to J. Wooden and M. Coble at the SUMAC Stanford-USGSSHRIMP-RG laboratory, and to R. Ingersoll, J. Trop, and an anonymousreviewer for their thorough and constructive reviews. A.F. and A.M.Hdedicate this manuscript to their son, Adriano.

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Received 7 October 2014; accepted 24 March 2015.


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