# Geodynamic and Climatic Forcing on Late-Cenozoic Exhumation of the Southern Patagonian Andes (Fitz Roy and Torres del Paine massifs)

Veleda a P Muller, Christian Sue, Pierre G Valla, Pietro Sternai, Thibaud Simon-Labric, Cécile Gautheron, Kurt M Cuffey, Djordje Grujic, Matthias Bernet, Joseph Martinod, et al.

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## ► To cite this version:

Veleda a P Muller, Christian Sue, Pierre G Valla, Pietro Sternai, Thibaud Simon-Labric, et al. Geodynamic and Climatic Forcing on Late-Cenozoic Exhumation of the Southern Patagonian Andes (Fitz Roy and Torres del Paine massifs). Tectonics, 2024, 43 (7), pp.e2023TC007914. 10.1029/2023tc007914. hal-04752526

HAL Id: hal-04752526

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Submitted on 24 Oct 2024

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Tectonics®

RESEARCH ARTICLE

10.1029/2023TC007914

**Key Points:**

* Mantle upwelling in Southern Patagonia is the most likely mechanism forcing rock exhumation between 10 and 8 Ma in the Fitz Roy massif

* Apatite (U-Th)/He data reveal glacial erosion as the main driver to the exhumation of the Fitz Roy and Torres del Paine between 7 and 4 Ma

* Apatite ⁴He/³He data reveal intensified fluvio-glacial erosion in Torres del Paine as a result of the Plio-Pleistocene climate transition

**Supporting Information:**

Supporting Information may be found in the online version of this article.

**Correspondence to:**

V. A. P. Muller,
veledamuller@arizona.edu

**Citation:**

Muller, V. A. P., Sue, C., Valla, P. G., Sternai, P., Simon-Labric, T., Gautheron, C., et al. (2024). Geodynamic and climatic forcing on late-Cenozoic exhumation of the Southern Patagonian Andes (Fitz Roy and Torres del Paine massifs). *Tectonics*, 43, e2023TC007914. https://doi.org/10.1029/2023TC007914

Received 3 MAY 2023

Accepted 3 JUL 2024

**Author Contributions:**

**Conceptualization:** Veleda A. P. Muller, Christian Sue, Pierre G. Valla, Pietro Sternai, Thibaud Simon-Labric, Cécile Gautheron, Kurt M. Cuffey, Djordje Grujic, Matthias Bernet, Joseph Martinod, Matias C. Ghiglione, Jean Braun

**Data curation:** Veleda A. P. Muller, Pierre G. Valla, Thibaud Simon-Labric, Cécile Gautheron, Frédéric Herman

**Formal analysis:** Veleda A. P. Muller, Pierre G. Valla, Cécile Gautheron, Kurt M. Cuffey

**Funding acquisition:** Christian Sue, Pierre G. Valla, Thibaud Simon-Labric,

© Wiley Periodicals LLC. The Author(s). This is an open access article under the terms of the Creative Commons Attribution License, which permits use, distribution and reproduction in any medium, provided the original work is properly cited.

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# Geodynamic and Climatic Forcing on Late-Cenozoic Exhumation of the Southern Patagonian Andes (Fitz Roy and Torres del Paine massifs)

Veleda A. P. Muller¹,²,³ ID, Christian Sue²,⁴ ID, Pierre G. Valla², Pietro Sternai¹ ID, Thibaud Simon-Labric²,⁵, Cécile Gautheron²,⁶, Kurt M. Cuffey⁷ ID, Djordje Grujic⁸ ID, Matthias Bernet², Joseph Martinod², Matias C. Ghiglione⁹, Peter Reiners³ ID, Chelsea Willett⁷ ID, David Shuster⁷, Frédéric Herman¹⁰ ID, Lukas Baumgartner¹¹ ID, and Jean Braun²,¹²,¹³ ID

¹Dipartimento di Scienze dell'Ambiente e della Terra (DISAT), Università degli Studi di Milano-Bicocca, Milan, Italy,

²Institute des Sciences de la Terre (ISTerre), Université Grenoble Alpes, Université Savoie Mont Blanc, CNRS, IRD,

IFSTTAR, Université Gustave Eiffel, Grenoble-Chambéry, France, ³Department of Geosciences, University of Arizona,

Tucson, AZ, USA, ⁴Université de Franche-Comté, Besançon, France, ⁵Centre de Géologie Oisans Alpes, Musée des

Minéraux, Bourg-d'Oisans, France, ⁶Université Paris Saclay, CNRS, GEOPS, Orsay, France, ⁷Department of Geography,

Department of Earth and Planetary Science, University of California - Berkeley, Berkeley, CA, USA, ⁸Department of Earth

and Environmental Sciences, Dalhousie University, Halifax, NS, Canada, ⁹Instituto de Estudios Andinos "Don Pablo

Groeber", Universidad de Buenos Aires, CONICET, Buenos Aires, Argentina. ¹⁰Institute of Earth Surface Dynamics

(IDYST), Université de Lausanne, Lausanne, Switzerland, ¹¹Institut des Sciences de la Terre (ISTE), Université de

Lausanne, Lausanne, Switzerland, ¹²Helmholtz Centre Potsdam, GFZ German Research Centre for Geosciences, Potsdam,

Germany, ¹³Institute of Earth and Environmental Sciences, University of Potsdam, Potsdam, Germany

**Abstract** High-relief glacial valleys shape the modern topography of the Southern Patagonian Andes, but their formation remains poorly understood. Two Miocene plutonic complexes in the Andean retroarc, the Fitz Roy (49°S) and Torres del Paine (51°S) massifs, were emplaced between 16.9–16.4 Ma and 12.6–12.4 Ma, respectively. Subduction of oceanic ridge segments initiated ca. 16 Ma at 54°S, leading to northward opening of a slab window with associated mantle upwelling. The onset of major glaciations caused drastic topographic changes since ca. 7 Ma. To constrain the respective contributions of tectonic-mantle dynamics and fluvio-glacial erosion to rock exhumation and landscape evolution, we perform inverse thermal modeling of a new data set of zircon and apatite (U-Th)/He from the two massifs, complemented by apatite ⁴He/³He data for Torres del Paine. Our results show rapid rock exhumation recorded only in the Fitz Roy massif between 10 and 8 Ma, which we ascribe to local mantle upwelling forcing surface uplift and intensified erosion around 49°S. Both massifs record a pulse of rock exhumation between 7 and 4 Ma, which we interpret as enhanced erosion during the beginning of Patagonian glaciations. After a period of erosional and tectonic quiescence in the Pliocene, increased rock exhumation since 3–2 Ma is interpreted as the result of alpine glacial valley carving promoted by reinforced glacial-interglacial cycles. This study highlights that glacial erosion was the main driver to rock exhumation in the Patagonian retroarc since 7 Ma, but that mantle upwelling might be a driving force to rock exhumation as well.

**Plain Language Summary** The isotopic system (U-Th)/He in apatite and zircon record the ages in which a rock experiences relatively low temperatures (200–60°C) at shallow crustal depths (6–1 km). We present a new data set of low-temperature thermochronometers for rocks of the Fitz Roy and Torres del Paine mountains in the Southern Patagonian Andes. Fast rock cooling can be forced by intensified surface erosion, and/or tectonic and mantle activity. An episode of fast cooling between 10 and 8 Ma was identified in the Fitz Roy mountains, and mantle upwelling forcing surface uplift, combined with high fluvial erosion may have caused fast rock exhumation. A regional episode of fast rock cooling between 7 and 4 Ma causing 1–3 km of exhumation in the Fitz Roy and Torres del Paine is coincident with the onset of Patagonian glaciations, which would have enhanced erosion and, thus, rock exhumation. An episode of fast rock cooling in the Quaternary is recorded in Torres del Paine rocks, interpreted as enhanced fluvio-glacial erosion during the Plio-Pleistocene climate transition toward faster glacial/interglacial cycles. Therefore, we were able to quantify separately the effects of tectonics and climate changes on rock exhumation, what is usually difficult due to simultaneously occurring processes.
Cécile Gautheron, Kurt M. Cuffey, Matias
C. Ghiglione, Lukas Baumgartner

**Investigation:** Veleda A. P. Muller,
Christian Sue, Pierre G. Valla,
Pietro Sternai, Cécile Gautheron, Kurt
M. Cuffey, Joseph Martinod, Jean Braun

**Methodology:** Veleda A. P. Muller, Pierre
G. Valla, Kurt M. Cuffey, Djordje Grujic,
Matthias Bernet, Peter Reiners,
Chelsea Willett, David Shuster,
Frédéric Herman, Lukas Baumgartner

**Project administration:** Christian Sue,
Pierre G. Valla, Pietro Sternai

**Resources:** Christian Sue,
Thibaud Simon-Labric, Cécile Gautheron,
Kurt M. Cuffey, Djordje Grujic,
Matthias Bernet, Matias C. Ghiglione,
Peter Reiners, Chelsea Willett,
David Shuster, Frédéric Herman,
Lukas Baumgartner

**Supervision:** Christian Sue, Pierre
G. Valla, Pietro Sternai

**Visualization:** Veleda A. P. Muller

**Writing – original draft:** Veleda
A. P. Muller, Pierre G. Valla,
Pietro Sternai, Cécile Gautheron, Kurt
M. Cuffey, Djordje Grujic,
Matthias Bernet, Matias C. Ghiglione

**Writing – review & editing:** Veleda
A. P. Muller

# 1. Introduction

Orogens that form along subduction and continental collision zones grow and evolve according to a long-term balance between incoming and outgoing material through tectonic accretion and climate-driven erosion (Beaumont et al., 2001; Dahlen, 1990; Egholm et al., 2009; Ruddiman et al., 1997; Whipple, 2009; Willett, 1999; Willett et al., 2001). Compression is primarily accommodated by folds, nappes stacking, thrusts, and transpressive faults, which lead to lithospheric shortening and thickening as well as surface uplift (Dahlen, 1990; Willett, 1999). At depths, lithospheric slab subduction and upper mantle dynamics modulate the stress and thermal states of the crust, thereby affecting rock and surface uplift (e.g., Conrad & Husson, 2009; Faccenna et al., 2013; Guillaume et al., 2010; Heuret & Lallemand, 2005; Molnar et al., 1993; Sternai et al., 2019). At the surface, erosion shapes the relief of orogens, exhuming rocks and generating sediments, which subsequently fill topographic depressions, that is, basins (Brocklehurst, 2010; Champagnac et al., 2014; England & Molnar, 1990; Whipple, 2009; Whipple & Tucker, 1999; Willett, 1999). Climate-controlled erosion and lithospheric, deep-seated processes are intrinsically linked and operate at various spatial and temporal scales. Quantifying their relative contributions to rock exhumation and landscape evolution in orogenic settings is of prime importance in current geoscience research (e.g., Cloetingh et al., 2023; DeCelles & Carrapa, 2023; Sternai, 2023).

The erosional mass outflux from partly-glaciated orogens is influenced by the partitioning between glacial, fluvial, and hillslope processes that widen and deepen valleys and shape mountain peaks (e.g., Braun & Willett, 2013; Sternai et al., 2011; Wang et al., 2024). Hillslope processes, which include soil creep and landslides, are controlled by the hillslope gradient and hillslope transport flux (Roering et al., 2001; Wang et al., 2024). The main erosional fluvial process is river incision, forming “V” shaped valleys and rounded mountain summits, primarily controlled by the local slope, drainage area, water discharge, and rock erodibility (Braun & Willett, 2013; Whipple, 2004; Whipple & Tucker, 1999; Willett, 1999). Glacial erosion tends to form “U” shaped valleys and sharp rock summits and is primary controlled by the ice sliding velocity, which in turn depends on the local slope and subglacial hydrology (Brocklehurst & Whipple, 2006; Champagnac et al., 2014; Flowers & Clarke, 2002; Fox et al., 2014; Herman et al., 2011, 2018; Shuster et al., 2005, 2011; Sternai et al., 2011, 2013, 2016). Late Cenozoic global cooling and associated onset of glaciations generated cyclic shifts in fluvial/glacial erosional processes with associated transience in mountain landscape and surface uplift (Egholm et al., 2009; Herman et al., 2013, 2018; Koppes & Montgomery, 2009; Molnar, 2004; Peizhen et al., 2001; Valla et al., 2011). At orogenic scale, the glacier's Equilibrium Line Altitude (ELA) may limit the elevation of glaciated mountain ranges, independent of the tectonic uplift rate (Egholm et al., 2009; Montgomery et al., 2001; Sternai et al., 2011). Other studies, however, suggested that tectonic uplift modulates both fluvial and glacial erosion rates in orogens (Koppes & Montgomery, 2009). The impact of glacial erosion on the elevation and topographic relief of mountain ranges is thus complex and depends on factors such as bedrock physical properties, pre-glacial mountain topography, basal ice thermal regime, etc. (Flowers & Clarke, 2002; Shuster et al., 2011; Sternai et al., 2011, 2013, 2016; Pedersen & Egholm, 2013, and references therein).

The typical steep and high topographic relief of glaciated mountain belts exhumes and exposes crystalline rocks formed at high crustal depths (e.g., Willett, 1999). Low-temperature thermochronological bedrock ages often vary systematically with elevation, with low elevation rocks being younger, that is, more recently cooled. Age-elevation relationships, thus, facilitate estimations of regional exhumation rates (Braun, 2002; Fitzgerald & Gleadow, 1988; Fitzgerald et al., 1995; Reiners & Brandon, 2006; Wagner, 1979; Wagner & Reimer, 1972) and provide useful information regarding the exhumation history of orogenic regions, such as the balance between regional erosion and tectonic rock uplift (Berger et al., 2008; Spotila, 2005; Tomkin & Roe, 2007; Willett et al., 2001).

In the Southern Patagonian Andes, glacial landscapes and related sedimentary deposits have been proposed to reflect the onset of major glaciations during late Miocene at ca. 7 Ma (Christeleit et al., 2017; Georgieva et al., 2016, 2019; Lagabrielle et al., 2010; Mercer & Sutter, 1982; Rabassa, 2008; Ronda et al., 2022; Thomson et al., 2001, 2010; Willett et al., 2020). Currently, the Northern Patagonian (NPI, 46–47°S), the Southern Patagonian (SPI, 48–52°S) and the Cordillera Darwin (54°S) icefields are located along the strike of the internal domain of the orogen (Figure 1). Our study focuses on the eastern proximity of the SPI (Figure 1), where mountain peaks rise above the current glaciers' ELA, which has been oscillating between 0.5 and 2 km above sea level since the Last Glacial Maximum (LGM, ~21,000 years before present) (Broecker & Denton, 1990; Davies et al., 2020). Miocene plutonic complexes that intrude deformed Mesozoic sedimentary rocks constitute some of
<figure><img src="image_6.png" /><figcaption>**Figure 1.** Geodynamic context of the Southern Patagonian Andes. The orange region is the surface projection of the slab window extent, and the Chile Triple Junction (CTJ, black circle) is located where the oceanic Chile Ridge is currently subducting beneath the South American Plate. The black and gray lines between the Nazca and the Antarctic plates are the present day, and the ancient positions (at ca. 16 and 12 Ma), respectively, of the oceanic ridges and transform faults separating the oceanic plates (Breitsprecher & Thorkelson, 2009). The white arrows show the velocity and relative direction of convergence between the Nazca or Antarctic plates and the South American Plate (DeMets et al., 2010). Red triangles show Quaternary volcanoes (Global Volcanism Program, 2023). Low-temperature thermochronometric data presented in this study are from the Torres del Paine (TdP) and Fitz Roy (FzR) massifs (yellow circles), located on the eastern border of the Southern Patagonian Icefield (SPI, white line). The blue line delimits the ice-covered region during the Last Glacial Maximum (LGM) at ~21 ka (adapted from Thorndycraft et al., 2019). Other abbreviations: NPI: Northern Patagonian Icefield, CDI: Cordillera Darwin Icefield, MFFZ: Magallanes-Fagnano Fault Zone, LOFZ: Liquiñe-Ofqui Fault Zone.</figcaption></figure>

the highest peaks, which are surrounded by steep valleys close to the ice fields, and by gentle valleys toward the eastern continental foreland (Fosdick et al., 2013; Ramírez de Arellano et al., 2012).

The entire region of southern Patagonia is lying above a slab window (Figure 1) characterized by a spatial gap between the Nazca and the Antarctic slabs where warm mantle upwells beneath the South America continent. The slab window is detected by Rayleigh and P-wave tomography (Mark et al., 2022; Russo et al., 2010), heat flow data (Ávila & Dávila, 2018), azimuthal anisotropy (Ben-Mansour et al., 2022), and GPS-derived rock uplift rates (Ávila & Dávila, 2020; Lange et al., 2014; Russo et al., 2022). The northern extent of the slab window is located at the latitude Chile Triple Junction (CTJ, at 46°S) where the Nazca, Antarctic and South America plates meet (Breitsprecher & Thorkelson, 2009; Cande & Leslie, 1986). The slab window opened through subduction of oceanic spreading ridge segments approximately parallel to the trench, migrating northward from an initial junction around 54°S at 16 Ma to about 49°S at 12 Ma, and 47°S at 3 Ma (Figure 1) (Breitsprecher & Thorkelson, 2009; Cande & Leslie, 1986; Fosdick et al., 2013; Georgieva et al., 2016, 2019; Guillaume et al., 2009, 2013; Haschke et al., 2006; Lagabrielle et al., 2010; Ramos, 2005; Scalabrino et al., 2010; Stevens Goddard & Fosdick, 2019; Thomson et al., 2001). Among the effects of oceanic ridge subduction are the inception of retroarc volcanism and cessation of arc volcanism, reduction in shortening, and regional surface and rock uplift (Ávila & Dávila, 2020; Ávila et al., 2023; Ding et al., 2023; Guillaume et al., 2009, 2013; Lagabrielle et al., 2004, 2010; Ramos, 2005; Scalabrino et al., 2010; Stevens Goddard & Fosdick, 2019). Mantle upwelling during oceanic ridge subduction is expected to have generated long-wavelength dynamic and isostatic surface uplift and eastward tilting of the continent, following the northward motion of the CTJ (Ávila & Dávila, 2018, 2020; Ávila et al., 2023; Ding et al., 2023; Guillaume et al., 2009, 2013).

Using low-temperature thermochronology, we aim to constrain the geo-dynamic and climate-driven forcing on rock exhumation in the Southern Patagonian Andes. We present new apatite and zircon (U-Th)/He (AHe and ZHe) thermochronological data sets from two of the most emblematic massifs of the Southern Patagonia Andes: the Fitz Roy (FzR, 49°S) and Torres del Paine (TdP, 51°S) (Figures 1-4). These massifs were emplaced around 16.7 ± 0.3 Ma (Ramírez de Arellano et al., 2012) and 12.5 ± 0.1 Ma (Leuthold et al., 2012), respectively, in the regions swept by the CTJ passage during the

Late Miocene, and thus may have recorded the earlier effects of ridge collision and subduction, besides those of late Cenozoic glaciation (Andrić-Tomašević et al., 2021; Ávila et al., 2023; Christeleit et al., 2017; Fosdick et al., 2013; Georgieva et al., 2016, 2019; Guillaume et al., 2009, 2013; Ramos, 2005; Ronda et al., 2022; Stevens Goddard & Fosdick, 2019; Thomson et al., 2001; Willett et al., 2020). Additionally, we present new apatite ⁴He/³He thermal modeling for the TdP massif, which resolves the latest stage of rock cooling and provides additional insights into the Pliocene-Pleistocene rock exhumation history of the area. Because these two massifs are about 200 km apart along the strike of the orogen, comparing our new thermochronological data with previously published data sets in the Southern Patagonian Andes (e.g., Stevens Goddard & Fosdick, 2019) enables us to interpret their local and/or regional exhumation histories including transient processes. Interpretations of our thermochronological records are combined with stratigraphical, structural and geophysical information from the literature.

## 2. Geological Context

### 2.1. Geodynamic Setting and Tectonic Evolution

The FzR and the TdP massifs are satellite plutons of the Southern Patagonian Batholith (SPB) located in the retroarc of the Southern Patagonian Andes. The SPB concentrated the locus of Late Jurassic to Miocene
<figure><img src="image_7.png" /><figcaption>**Figure 2.** Geological map of the Fitz Roy massif (FzR), Argentina (modified from Giacosa et al., 2013, background satellite image is from ©Google Earth), with sample location. Major topographic peaks are indicated by red triangles, samples for (U-Th)/He thermochronometry are located by white (apatite) and black (zircon) circles. See Figure 1 for location within the Southern Patagonian Andes.</figcaption></figure>

subduction-related magmatism intruding Paleozoic metamorphic complexes and the Late Jurassic–Early Cretaceous Rocas Verdes Basin in the hinterland (Calderón et al., 2012; Hervé et al., 2007; Muller et al., 2021), and the Late Cretaceous–Cenozoic meta volcano-sedimentary units of the Magallanes-Austral Basin in the foreland (Fosdick et al., 2011, 2013; Ghiglione et al., 2009; Giacosa et al., 2013; Leuthold et al., 2012; Ramírez de Arellano et al., 2012). The closure of the ocean-floored backarc Rocas Verdes Basin in the Late Cretaceous ensued crustal shortening and thickening in the retroarc, and foreland subsidence with deposition of marine siliciclastic basal deposits of the Magallanes-Austral Basin (Calderón et al., 2012; Fildani et al., 2003; Fosdick et al., 2011; Ghiglione et al., 2009; Malkowski et al., 2017; Maloney et al., 2013; Muller et al., 2021). After the break-up of the Farallon Plate into the Nazca and the Cocos plates at ~25 Ma, the increase in convergence velocity generated eastward thrust propagation into the foreland basin, contributing to the formation of the N-S oriented Patagonian fold-and-thrust belt (Figures 2 and 3; Betka et al., 2015; Fosdick et al., 2011, 2013; Ghiglione et al., 2009; Giacosa et al., 2013; Kraemer, 2003; Suárez et al., 2000). Small Miocene plutonic complexes intrude the Patagonian fold-and-thrust belt, distributed for 800 km along the strike of the orogen (Ramírez de Arellano et al., 2012). In the FzR massif, the Chaltén Plutonic Complex (49°S) is composed of granitic to gabbroic rocks crystallized between 16.9 ± 0.05 and 16.4 ± 0.02 Ma (Ramírez de Arellano et al., 2012). Hornblende-thermobarometry from this plutonic complex indicates magmatic emplacement at 8–10 km depth and
<figure><img src="image_8.png" /><figcaption>**Figure 3.** Geological map of the Torres del Paine massif (TdP), Chile (modified from Fosdick et al., 2013, background satellite image from ©Google Earth), with sample location. Major topographic peaks are indicated by red triangles, samples for apatite (U-Th)/He thermochronometry are located by colored circles (see legend for details) and samples with <sup>4</sup>He/<sup>3</sup>He data are marked by a red asterisk. Plutonic samples in the metasedimentary units are from small-scale dikes, which are not distinguishable in this map scale. See Figure 1 for the location within the Southern Patagonian Andes.</figcaption></figure>

exhumation to 6–4 km depth during the syn-magmatic phase (Ramírez de Arellano, 2011). Currently, this plutonic complex is exposed, for example, in the Mount Fitz Roy (3,405 m a.s.l.) and Cerro Torre (3,128 m a.s.l.) (Figure 2). The Torres del Paine Plutonic Complex, located at 51°S, is a laccolith with feeder dikes of granitic to gabbroic composition, emplaced at 2–4 km depth, as constrained from contact-metamorphic assemblages such as prehnite-anorthite (Putlitz et al., 2001), and crystallized between 12.4 ± 0.006 Ma and 12.6 ± 0.009 Ma (Leuthold et al., 2012). The culminant peaks of the TdP massif are the Cerro Paine Grande (2,884 m a.s.l.), composed of Late Cretaceous metasedimentary siliciclastic rocks encompassing the pluton and the granitic Torre Central (2,460 m a.s.l., Figure 3).

The older and faster Nazca Plate subduction with respect to the Antarctic Plate generated a much wider Nazca slab than the Antarctic slab (DeMets et al., 2010; Hayes et al., 2018), with space for mantle upwelling in a slab window beneath Southern Patagonia (Figure 1; Ávila & Dávila, 2018; Ben-Mansour et al., 2022; Breitsprecher & Thorkelson, 2009; Cande & Leslie, 1986; Lagabrielle et al., 2004, 2007; Mark et al., 2022; Ramos, 2005; Ramos & Kay, 1992; Russo et al., 2010, 2020). Between 12 and 6 Ma, the slab window broadened between latitudes 51
<figure><img src="image_9.png" /><figcaption>**Figure 4.** Sample spatial distributions within the FzR and TdP study regions. (a) Fitz Roy massif and samples with AHe data (white circles) and ZHe data (black circles). (b) Torres del Paine massif and samples with AHe data from the Central sector (yellow circles), samples with additional ⁴He/³He data are marked by a red asterisk, North sector (white circles), and the West sector (green circles). Sample CH15-TP14 is a low-elevation sample from the West sector out of the figure area, see Figure 3 for location.</figcaption></figure>

and 48°S, triggering a dynamic surface uplift related to mantle flow, estimated at up to ~800 m (Guillaume et al., 2009). More recent studies argue that isostatic uplift, due to lithospheric thinning in response to the hot mantle upwelling, predominates with respect to dynamic uplift, and can have generated up to 0.3 mm/yr of regional uplift since the Late Miocene (V. A. Muller et al., 2024; Ávila & Dávila, 2020; Ávila et al., 2023). This thermal effect contributes to the 10–40 mm/yr rate of rock uplift currently measured in the SPI, which occurs as an isostatic response to glacier retreat over the last hundreds to thousands of years (V. A. Muller et al., 2024). Around the current latitude of the CTJ (46–47°S), the maximum elevations of the Southern Patagonian Andes drastically increase (from 2 to 4 km a.s.l.) with respect to the orogen just north of the CTJ (Georgieva et al., 2016; Ramos, 2005; Stevens Goddard & Fosdick, 2019; Thomson et al., 2010). South of 47°S, retroarc tectonic deformation associated with episodes of oceanic ridge collision with the subduction trench over the last 6 Myr generated local faulting and associated topographic variations (Andrić-Tomašević et al., 2021; Georgieva et al., 2016, 2019; Lagabrielle et al., 2004). North of the CTJ, the transpressive Liquiñe-Ofqui Fault Zone (LOFZ) accommodates a great part of the lithospheric deformation and is associated with fast rock exhumation in the last 10 Myr (Thomson, 2002). Immediately south of the CTJ, transpressive and extensional faults active in the last 3–2 Myr were assigned as the main causes of rock uplift and exhumation around the NPI (Georgieva et al., 2016). The FzR and TdP, instead, are distant from the damage areas of the active LOFZ and Magallanes-Fagnano Fault Zone (Figure 1), enabling us to dismiss any potential thermal/deformational effect from these major strike-slip faults on the cooling and exhumation of the study region (Georgieva et al., 2016, 2019; Guillaume et al., 2013; Lagabrielle et al., 2010; Thomson et al., 2001).
Magmatic effects of oceanic ridge subduction include the cessation of arc volcanism (Ramos, 2005), and extensive plateau basaltic volcanism in the retroarc region, recording the episodes of slab window opening since ~14 Ma (Breitsprecher & Thorkelson, 2009; Gorring et al., 1997; Guivel et al., 2006; Ramos & Kay, 1992). Amongst the six Quaternary volcanoes of the Austral Andes Volcanic Zone (Global Volcanism Program, 2023; Stern et al., 1984), the closest Lautaro (49°S) and Reclus (51°S) volcanoes are located more than 20 km away from the study regions (Figure 1), which leads us to assume negligible thermal influence from recent volcanism on the investigated massifs.

## 2.2. Paleoclimatic Setting and Geomorphologic Evolution

The Southern Patagonian Andes are approximately perpendicular to the main wind trend dominated by the Westerlies, acting as an orographic barrier since at least the late Oligocene (Blisniuk et al., 2006; Fosdick et al., 2013; Thomson et al., 2001). This implies enhanced precipitation and erosion at the windward (western) side of the orogen, and little erosion on the leeward (eastern) side due to the so-called "rain shadow" (Blisniuk et al., 2006; Fosdick et al., 2013; Herman & Brandon, 2015; Thomson et al., 2001). South of 42°S, the present-day mean annual precipitation rates are ~1-4 m/yr on the windward side, and less than 0.5 m/yr on the leeward side of the orogen (Garreaud et al., 2013). As a result of sustained orography and tectonic convergence, an eastward migration of the orographic barrier between 30 and 8 Ma is proposed to be related with an eastward migration of the locus of retroarc deformation (Fosdick et al., 2013; Thomson et al., 2001). Between 42°S and 46°S, orography and glacial-periglacial erosion are proposed to enhance denudation, tectonic deformation, and volcanism around the LOFZ (Muller et al., 2022; Thomson, 2002).

The onset of Patagonian glaciations at around 7 Ma is supported by sedimentary and geomorphologic evidence, including glacial troughs, striations, and till and moraine deposits around the paleo-glacial lakes of the retroarc, fed by ice sheets from the orogen interior (Mercer and Sutter, 1982; Singer et al., 2004; Rabassa et al., 2005, 2011; Lagabrielle et al., 2010; Zachos et al., 2001). Extensive ice cover formed the "Patagonian Ice Sheet" between 7 and 4 Ma, terminating in glacial lobes reaching approximately the LGM limit (Figure 1) (Lagabrielle et al., 2010; Mercer & Sutter, 1982; Rabassa, 2008). At 3 Ma, a drastic change in the retroarc landscape near the glacier terminations turned into a series of W-E/SW-NE deep incised glacial valleys, interpreted as drainage reorganization in response to tectonic inversion and uplift (Andrić-Tomašević et al., 2021; Georgieva et al., 2019; Lagabrielle et al., 2010; Willett et al., 2020). During the Quaternary, several advancements and retreats of the Patagonian glaciers are recorded in the sedimentary record, likely linked with a Plio-Pleistocene increase in the frequency of glacial-interglacial cycles (A. M. De Boer et al., 2010; Lisiecki, 2010; Lisiecki & Raymo, 2007). Periods of maximum extent of the ice sheets are known as the "Great Patagonian Glaciation," at ~1 Ma, and the LGM, at ~21 ka (Davies & Glasser, 2012; Davies et al., 2020; Glasser & Jansson, 2008; Kaplan et al., 2004; Singer et al., 2004; Thorndycraft et al., 2019). Negative feedback in response to an increase in fluvial-glacial erosion rates is the reduction in accumulation area and consequent shrinking of the ice sheets to the present-day icefields (Kaplan et al., 2009; Singer et al., 2004). The FzR and TdP massifs are located in the rain shadow at the eastern side of the SPI. These regions host some of the highest peaks of Southern Patagonia (~2,500-3,500 m a.s.l.) and important glaciers that have shaped the landscape and feed the major lakes and fluvial drainage systems running through the eastern semi-arid region up to the Atlantic Ocean (Figures 1-4).

## 2.3. Available Low-Temperature Thermochronology and Regional Exhumation Record

Low-temperature thermochronological ages (apatite and zircon (U-Th)/He and fission track) from the windward side of the SPB range from 60 to 10 Ma (Thomson et al., 2001, 2010). Regional eastward younging ages were interpreted as an across-strike eastward migration of the topographic divide, a possible effect of subduction erosion enhancing compressive deformation in the retroarc (Thomson et al., 2001). In the TdP massif, recorded cooling ages ranging between 22 and 18 Ma were ascribed to erosional exhumation during thrusting in the Patagonian fold-and-thrust belt evolution (Fosdick et al., 2011, 2013). Younger ages between 7 and 3 Ma were ascribed mainly to enhanced glacio-fluvial erosion after the onset of Patagonian glaciations (Fosdick et al., 2011, 2013). The geographic position of the TdP at the narrow termination of South America, close to the Atlantic and Pacific oceans, might be the cause of intensified erosion and younger cooling ages even inside the rain shadow (Fosdick et al., 2013).
Around the CTJ region (46–47°S), low-T thermochronological ages of 7–3 Ma were associated with both changing glacial/interglacial cycles, and transform and thrust faulting due to oceanic ridge collision (Andrić-Tomašević et al., 2021; Christeleit et al., 2017; Georgieva et al., 2016, 2019; Lagabrielle et al., 2010; Ronda et al., 2022; Scalabrino et al., 2010; Thomson et al., 2001; Willett et al., 2020). Exhumation rates can be averaged as 0.5 km/Myr since ca. 12–10 Ma at the eastern boundary of the NPI, but can reach up to 2.2 km/Myr between ca. 5 and ca. 3 Ma, assigned to a combination of intensified glacial erosion and tectonic activity (Andrić-Tomašević et al., 2021; Georgieva et al., 2016, 2019), and potentially with dynamic uplift due to mantle upwelling (Guil-laume et al., 2013).

Regional southward older low-temperature thermochronological ages are associated with a decrease in the long-term erosional efficiency linked with bedrock shielding by cold-based glaciers at high latitudes (Thomson et al., 2010). There is no low-temperature thermochronological data younger than 3 Ma in the area, thus limiting our knowledge of the most recent exhumation history in the region.

# 3. Materials and Methods

## 3.1. Sample Locations and Processing

Sampled bedrock outcrops are distributed within the FzR and TdP massifs, with a sampling strategy along elevation profiles when possible (Figures 2–4). The FzR profile covers an elevation segment of 660 m (from 1,410 to 2,070 m) over 3–4 km of horizontal distance of magmatic rocks of the Chaltén Plutonic Complex (Figures 2, 4a, and 5a). We collected seven samples for AHe dating, and four samples for zircon ZHe dating (Tables 1–2). In the TdP massif, AHe samples were collected and grouped into three sectors (Central, North and West, Figures 3 and 4b), according to their location and age-elevation relationship (Figure 5b). Two samples from the Central sector also have apatite ⁴He/³He ages (Tables 3–5 and Figure S1 in Supporting Information S1). The Central sector covers an elevation segment of 1,600 m (from 206 to 1,802 m) over 20 km of horizontal distance, while the West and North sectors extend over 630 m of elevation (from 112 to 746 m) and 30 km of horizontal distance, and 550 m elevation (from 369 to 916 m) and 5 km of horizontal distance, respectively (Figures 3 and 4b). The Central sector is comprised of 15 plutonic samples from the Torres del Paine Plutonic Complex and three metasedi-mentary samples from the Patagonian fold-and-thrust belt located near the Nordenskjöld Lake (Figures 3, 4b, and 5b). The West sector is comprised of three samples of metasedimentary rocks near the Gray and Tyndall lakes (although samples CH15-TP15 and CH15-TP22 are located close to the Central sector, their respective AHe ages are different from the age-elevation relationship of the Central sector, leading us to consider them separately for thermal modeling). The North sector has two samples of metasedimentary rocks near the Dickson Lake (Figures 3 and 4b).

Apatite and zircon crystals were extracted from bedrock samples using crushing followed by standard magnetic and heavy-liquid separation techniques (Kohn et al., 2019). For apatite and zircon, preparation included selection of crystals with euhedral shape, preferably two terminations, and absence of inclusion. Based on length and width measurements, equivalent spherical radius ($R_s$, between 40 and 100 µm for all analyzed crystals) and alpha ejection correction factor ($F_T$) were calculated (Tables 2 and 3).

## 3.2. AHe, ZHe, and Apatite ⁴He/³He Thermochronology Data

ZHe and AHe dating methods used in this study are based on ⁴He production and accumulation within a crystal coming from alpha radioactive decay of the parent nuclides ²³⁸U, ²³⁵U, ²³²Th, and ¹⁴⁷Sm, and their daughter products in the respective decay chains (Farley, 2002; Zeitler et al., 1987). Helium accumulates in a crystal depending on the diffusion coefficient, crystal size, and possibly the presence of crystal broken pyramids for apatite (e.g., Brown et al., 2013; Reiners & Farley, 2001). He diffusion in apatite and zircon is strongly controlled by radiation damage accumulating in crystals with time (Reiners, 2005; Shuster et al., 2006). For apatite, it has been shown in addition that radiation damage anneals as a function of temperature (Shuster & Farley, 2009) and the chemical composition (Gautheron et al., 2013). In consequence, He retention is complex and the associated effective closure temperature ($T_c$) varies with alpha damage dose production and annealing through time (Dodson, 1973, 1979). The $T_c$ represents the temperature where 50% of the produced He atoms are retained in a crystal structure for a monotonic cooling. This notion is a purely mathematic formulation but illustrates the temperature sensitivity of a thermochronological system. As a result, (U-Th)/He age varies depending on the crystal size (diffusion domain), and diffusion coefficient that change as a function of the alpha damage dose
<figure><img src="image_10.png" /><figcaption>**Figure 5.** Age-elevation profiles for the FzR and TdP massifs. (a) ZHe and AHe data for the FzR massif: single-crystal apatite and zircon (U-Th)/He ages are indicated by circles, and mean ages with $1\sigma$ errors (standard deviation of the single-crystal ages) by diamonds. (b) AHe data for the TdP massif: single-crystal apatite (U-Th)/He ages are indicated by circles, and mean ages with $1\sigma$ errors (standard deviation of the single-crystal ages) by diamonds. Two samples with apatite $^{4}$He/$^{3}$He data are marked by a red asterisk. Colors correspond to different sectors of the TdP massif (Central, West and North sectors), gray samples are sedimentary (sed.) and yellow are plutonic (plut.) samples from the Central sector, according to Figure 3 and explained in the main text. Full analytical details for AHe and ZHe data are given in Tables 1–3 and in Supporting Information S1.</figcaption></figure>

accumulation through time. For apatite, the $T_c$ increases with alpha damage dose from 50 to 120°C (Djimbi et al., 2015; Flowers et al., 2009; Gautheron et al., 2009; Shuster et al., 2006). For zircon, the $T_c$ at low alpha damage dose (<10$^{18}$ $\alpha$/g) is around 100°C (Gautheron et al., 2020; Gérard et al., 2022) and increases with the alpha damage dose up to 200°C (Guenthner et al., 2013). For alpha damage doses above 10$^{18}$ $\alpha$/g, the damage percolation threshold value, the $T_c$ decreases up to temperatures <100°C (Guenthner et al., 2013; Ketcham et al., 2013).

For apatite, He diffusion models exist that consider damage production and annealing, and here we used the model of Flowers et al. (2009) with predicted $T_c$ for this study ranging between around 50 to 70°C. This model was chosen because the radiation damage present in the apatite crystals of our study is relatively low, as the plutonic rocks are young (FzR: 16.7 ± 0.3 Ma, Ramírez de Arellano, 2012; TdP: 12.5 ± 0.1 Ma; Leuthold et al., 2012). For ZHe from the FzR plutonic complex, we used an adapted He diffusion coefficient, with He diffusion parameters calculated based on Gautheron et al. (2020), taking into account the plutonic rock age and U-Th content (Table 1 and Table S2 in Supporting Information S1). The calculated closure temperatures $T_c$ of the ZHe system range between ~87 and ~108°C for the FzR samples (See Supporting Information S1 and Table S2 for details). In $^{4}$He/$^{3}$He thermochronometry (Shuster & Farley, 2004) the natural spatial distribution of radiogenic $^{4}$He is constrained by stepwise degassing and $^{4}$He/$^{3}$He analysis of a sample containing synthetic, homogeneously distributed, proton-induced $^{3}$He.

AHe thermochronometry was performed following standard procedures (House et al., 2000) at the ARHDL of the University of Arizona (USA) for plutonic samples of TdP, and at the Berkeley Geochronology Center (USA) for metasedimentary samples of TdP and $^{4}$He/$^{3}$He data. AHe thermochronometry of FzR plutonic samples was performed in the GEOPS laboratory at the Paris-Saclay University (Orsay, France), following the methods of Gautheron et al. (2021) and ZHe thermochronometry was performed in the UTHHE Laboratory of Dalhousie University (Halifax, Canada) following the methods of Reiners et al. (2004) and Reiners (2005). Full analytical details for ZHe, AHe and $^{4}$He/$^{3}$He data production are given in Supporting Information S1. The detailed reduced database for AHe and ZHe is available at V. Muller et al. (2024).

### 3.3. Inverse Thermal Modeling

We used inverse thermal modeling to interpret our new AHe and ZHe data in terms of bedrock cooling histories, and eventually to estimate the timing and spatial differences in exhumation histories between the TdP and FzR massifs. To this aim, we used the QTQt model (Gallagher, 2012), which is based on a Bayesian Markov-Chain Monte-Carlo approach to statistically explore different temperature-time (T-t) paths for multiple samples distributed along an elevation profile. For predicting He diffusion in an apatite or zircon crystal, the model uses the raw contents of U, Th and Sm, and the equivalent spherical radius ($R_s$) calculated from crystal measurements in the laboratory (Tables 1–3). For AHe data, we used the He diffusion kinetic parameters from the radiation damage and annealing model of Flowers et al. (2009). For ZHe data, we estimated and input the He diffusion parameters (activation energy, $E_a$, and frequency factor, $D_0$) for each individual zircon crystal to further investigate He diffusion in zircon (see Supporting Information S1 and Table S2 for details; Gérard et al., 2022). Then, we conducted several thermal QTQt inversions for FzR massif and TdP sectors (Central, West and North sectors, Figure 5) with shared input modeling parameters in QTQt. First, we prescribed a geothermal gradient of 35 ± 10°C/km, according to the 70–90 mW/m$^2$ regional thermal
<table><caption>**Table 1**
Zircon (U-Th)/He (ZHe) Data for Fitz Roy (FzR) Samples</caption><thead><tr><th>Sample N°</th><th>Latitude/Longitude<br>(°S/°W)</th><th>Elevation (m)</th><th>U (ppm)</th><th>Th (ppm)</th><th>Sm (ppm)</th><th><sup>4</sup>He (nmol/g)</th><th>R<sub>s</sub> (μm)</th><th>F<sub>t</sub></th><th>Raw age<br>(Ma)</th><th>Raw error<br>(Ma)</th><th>Corrected age<br>(Ma)</th><th>1σ error<br>(Ma)</th></tr></thead><tbody><tr><td>FZR3-1</td><td></td><td></td><td>205.8</td><td>91.5</td><td>6.0</td><td>7.914</td><td>54.35</td><td>0.78</td><td>6.4</td><td>0.04</td><td>8.2</td><td>0.1</td></tr><tr><td>FZR3-2</td><td></td><td></td><td>284.3</td><td>100.8</td><td>2.5</td><td>11.135</td><td>56.52</td><td>0.79</td><td>6.7</td><td>0.04</td><td>8.5</td><td>0.1</td></tr><tr><td>FZR3-3</td><td></td><td></td><td>506.7</td><td>156.9</td><td>3.4</td><td>18.834</td><td>53.10</td><td>0.78</td><td>6.4</td><td>0.04</td><td>8.2</td><td>0.1</td></tr><tr><td>FZR3-4</td><td></td><td></td><td>516.9</td><td>138.7</td><td>2.5</td><td>19.664</td><td>53.24</td><td>0.78</td><td>6.6</td><td>0.04</td><td>8.5</td><td>0.1</td></tr><tr><td>FZR3-5</td><td></td><td></td><td>513.1</td><td>198.4</td><td>2.6</td><td>21.002</td><td>51.87</td><td>0.77</td><td>6.7</td><td>0.04</td><td>8.7</td><td>0.1</td></tr><tr><td>FZR3</td><td>49.2566/73.0265</td><td>2,070</td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td><strong>8.4 ± 0.2</strong></td><td></td></tr><tr><td>FZR4-1</td><td></td><td></td><td>371.3</td><td>140.9</td><td>3.5</td><td>13.674</td><td>47.51</td><td>0.76</td><td>6.3</td><td>0.04</td><td>8.3</td><td>0.1</td></tr><tr><td>FZR4-3</td><td></td><td></td><td>309.6</td><td>109.3</td><td>2.5</td><td>11.613</td><td>52.78</td><td>0.78</td><td>6.4</td><td>0.04</td><td>8.3</td><td>0.1</td></tr><tr><td>FZR4-4</td><td></td><td></td><td>1,128.8</td><td>326.6</td><td>3.9</td><td>34.597</td><td>45.21</td><td>0.75</td><td>5.3</td><td>0.03</td><td>7.1</td><td>0.4</td></tr><tr><td>FZR4-5</td><td></td><td></td><td>715.6</td><td>234.1</td><td>2.9</td><td>23.664</td><td>52.90</td><td>0.78</td><td>5.7</td><td>0.03</td><td>7.3</td><td>0.4</td></tr><tr><td>FZR4</td><td>49.2550/73.0281</td><td>1,955</td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td><strong>7.7 ± 0.6</strong></td><td></td></tr><tr><td>FZR5-1</td><td></td><td></td><td>904.2</td><td>321.4</td><td>7.4</td><td>35.659</td><td>52.99</td><td>0.78</td><td>6.7</td><td>0.04</td><td>8.7</td><td>0.5</td></tr><tr><td>FZR5-2</td><td></td><td></td><td>271.4</td><td>96.8</td><td>3.1</td><td>11.343</td><td>46.23</td><td>0.75</td><td>7.1</td><td>0.04</td><td>9.6</td><td>0.6</td></tr><tr><td>FZR5-3</td><td></td><td></td><td>412.4</td><td>310.9</td><td>5.7</td><td>12.060</td><td>45.33</td><td>0.74</td><td>4.6</td><td>0.03</td><td>6.2</td><td>0.4</td></tr><tr><td>FZR5-4</td><td></td><td></td><td>527.9</td><td>246.5</td><td>6.5</td><td>22.495</td><td>43.24</td><td>0.73</td><td>7.1</td><td>0.04</td><td>9.7</td><td>0.6</td></tr><tr><td>FZR5-5</td><td></td><td></td><td>529.4</td><td>268.6</td><td>5.2</td><td>14.285</td><td>43.28</td><td>0.73</td><td>4.5</td><td>0.03</td><td>6.1</td><td>0.4</td></tr><tr><td>FZR5</td><td>49.2540/73.0311</td><td>1,758</td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td><strong>8.0 ± 1.8</strong></td><td></td></tr><tr><td>FZR6-1</td><td></td><td></td><td>1,157.2</td><td>380.0</td><td>8.4</td><td>66.359</td><td>49.97</td><td>0.77</td><td>9.9</td><td>0.06</td><td>12.9</td><td>0.8</td></tr><tr><td>FZR6-2</td><td></td><td></td><td>1,221.7</td><td>476.3</td><td>9.3</td><td>62.142</td><td>45.48</td><td>0.75</td><td>8.6</td><td>0.05</td><td>11.6</td><td>0.7</td></tr><tr><td>FZR6-3</td><td></td><td></td><td>1,021.0</td><td>290.1</td><td>4.2</td><td>41.435</td><td>43.47</td><td>0.74</td><td>7.1</td><td>0.04</td><td>9.6</td><td>0.6</td></tr><tr><td>FZR6-4</td><td></td><td></td><td>403.0</td><td>141.7</td><td>2.8</td><td>11.628</td><td>44.43</td><td>0.74</td><td>4.9</td><td>0.03</td><td>6.7</td><td>0.4</td></tr><tr><td>FZR6</td><td>49.2487/73.0327</td><td>1,569</td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td><strong>10.2 ± 2.7</strong></td><td></td></tr></tbody></table>

Note. F<sub>t</sub> = alpha-ejection correction factor (Farley et al., 1996). Bold numbers are mean ages calculated from the single-crystal replicates; 1σ error for mean ages is standard deviation of replicate ages. Analytical uncertainties are around 6% of the corrected age.

flow predicted by Ávila and Dávila (2018). The geothermal gradient is allowed to vary with time within the 35 ± 10°C/km range, and no reheating was allowed due to the lack of evidence for a reheating event in the study FzR and TdP regions during the late Miocene to Plio-Quaternary (Ramírez de Arellano et al., 2012). We considered an atmospheric lapse rate of 6 ± 2°C/km, and a present-day surface temperature of 1 ± 1°C to ensure model simulations reaching surface temperature for modern conditions (all sample locations have a mean annual surface temperature below or around 0°C; Garreaud et al., 2013). We also constrained the initial thermal constraints based on pluton crystallization ages, using temperature-time constraints of 275 ± 25°C (i.e., hotter than the respective T<sub>c</sub> of the ZHe and AHe systems), and 16 ± 1 Ma and 12 ± 1 Ma for the FzR and TdP (Central sector) massifs, respectively. For the meta-sedimentary samples (West and North sectors of TdP massif), we did not impose any initial thermal constraint, and re-heating was not allowed given the lack of independent evidence for re-heating events after early Miocene basin thrusting (Fosdick et al., 2011; Klepeis et al., 2010). For thermal inversion, QTQt's modeling is based on a linear interpolation between the highest and the lowest elevation samples to predict thermal paths randomly and shared geothermal gradients for all samples in an elevation profile (Figures 6–8, Figures S1–S3 in Supporting Information S1) that best predicts observed thermochronological data in a consistent manner (Gallagher, 2012, 2021). QTQt inverse simulations were done for 10,000 iterations to ensure the model convergence and robustness of the inversion results (see Supporting Information S1).

Finally, we used thermal inversion modeling (Schildgen et al., 2010) to interpret ⁴He/³He thermochronological data in two TdP samples (Figure 3, Table S1 in Supporting Information S1). To explore possible changes in ⁴He diffusivity through time, all cooling paths (20–30 × 10,000 iterations for each sample) began at 150°C, well above the accumulation of radiation damage effects (Flowers et al., 2009) and ended after 10 Myr at the modern surface
<table><caption>**Table 2**
Apatite (U-Th)/He (AHe) Data for Fitz Roy (FzR) Samples</caption><thead><tr><th>Sample<br>N°</th><th>Latitude/Longitude<br>(°S/°W)</th><th>Elevation<br>(m)</th><th>U<br>(ppm)</th><th>Th<br>(ppm)</th><th>Sm<br>(ppm)</th><th><sup>4</sup>He<br>(nmol/g)</th><th>R<sub>s</sub><br>(μm)</th><th>F<sub>t</sub></th><th>Raw<br>age (Ma)</th><th>Raw<br>error (Ma)</th><th>Corrected<br>age (Ma)</th><th>1σ<br>error (Ma)</th></tr></thead><tbody><tr><td>FZR10-M</td><td></td><td></td><td>13.3</td><td>38.2</td><td>277.3</td><td>0.716</td><td>43.5</td><td>0.73</td><td>5.5</td><td>0.9</td><td>7.5</td><td>0.6</td></tr><tr><td>FZR10-N</td><td></td><td></td><td>53.3</td><td>168.0</td><td>436.1</td><td>3.646</td><td>42.2</td><td>0.73</td><td>7.1</td><td>1.2</td><td>9.7</td><td>0.8</td></tr><tr><td><strong>FZR10</strong></td><td><strong>49.2543/73.0319</strong></td><td><strong>1,780</strong></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td><strong>8.6 ± 1.6</strong></td><td></td></tr><tr><td>FZR11-A</td><td></td><td></td><td>3.5</td><td>12.0</td><td>103.0</td><td>0.164</td><td>55.4</td><td>0.80</td><td>4.3</td><td>1.0</td><td>5.4</td><td>0.4</td></tr><tr><td>FZR11-D</td><td></td><td></td><td>8.0</td><td>19.9</td><td>170.5</td><td>0.521</td><td>51.1</td><td>0.77</td><td>6.9</td><td>0.7</td><td>8.9</td><td>0.7</td></tr><tr><td>FZR11-E</td><td></td><td></td><td>23.5</td><td>38.6</td><td>122.4</td><td>0.812</td><td>60.7</td><td>0.80</td><td>4.5</td><td>0.7</td><td>5.7</td><td>0.5</td></tr><tr><td>FZR11-L</td><td></td><td></td><td>14.4</td><td>46.4</td><td>344.1</td><td>0.568</td><td>40.0</td><td>0.73</td><td>3.8</td><td>0.0</td><td>5.1</td><td>0.4</td></tr><tr><td><strong>FZR11</strong></td><td><strong>49.2538/73.0346</strong></td><td><strong>1,840</strong></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td><strong>6.3 ± 1.8</strong></td><td></td></tr><tr><td>FZR12-E</td><td></td><td></td><td>7.3</td><td>20.6</td><td>136.3</td><td>0.369</td><td>57.1</td><td>0.78</td><td>5.2</td><td>0.8</td><td>6.7</td><td>0.5</td></tr><tr><td>FZR12-M</td><td></td><td></td><td>23.6</td><td>68.0</td><td>381.9</td><td>1.332</td><td>41.8</td><td>0.71</td><td>5.9</td><td>1.0</td><td>8.3</td><td>0.7</td></tr><tr><td>FZR12-N</td><td></td><td></td><td>15.3</td><td>45.8</td><td>250.0</td><td>0.930</td><td>47.0</td><td>0.76</td><td>6.2</td><td>1.1</td><td>8.1</td><td>0.7</td></tr><tr><td><strong>FZR12</strong></td><td><strong>49.2493/73.0364</strong></td><td><strong>1,765</strong></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td><strong>7.7 ± 0.9</strong></td><td></td></tr><tr><td>FZR13-B</td><td></td><td></td><td>26.9</td><td>46.2</td><td>140.1</td><td>0.968</td><td>56.3</td><td>0.77</td><td>4.7</td><td>0.0</td><td>6.0</td><td>0.1</td></tr><tr><td>FZR13-F</td><td></td><td></td><td>7.9</td><td>22.8</td><td>83.5</td><td>0.472</td><td>77.0</td><td>0.84</td><td>5.0</td><td>0.0</td><td>6.0</td><td>0.1</td></tr><tr><td>FZR13-L</td><td></td><td></td><td>71.3</td><td>135.8</td><td>461.4</td><td>2.465</td><td>41.3</td><td>0.69</td><td>4.3</td><td>0.7</td><td>6.3</td><td>0.3</td></tr><tr><td><strong>FZR13</strong></td><td><strong>49.2553/73.0473</strong></td><td><strong>1,410</strong></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td><strong>6.1 ± 0.2</strong></td><td></td></tr><tr><td>FZR14-A</td><td></td><td></td><td>10.0</td><td>17.4</td><td>105.9</td><td>0.443</td><td>66.3</td><td>0.81</td><td>5.6</td><td>0.2</td><td>6.9</td><td>0.1</td></tr><tr><td>FZR14-D</td><td></td><td></td><td>11.1</td><td>30.5</td><td>186.5</td><td>0.631</td><td>54.2</td><td>0.74</td><td>6.0</td><td>0.2</td><td>8.1</td><td>0.1</td></tr><tr><td>FZR14-E</td><td></td><td></td><td>10.5</td><td>29.4</td><td>183.3</td><td>0.568</td><td>58.0</td><td>0.75</td><td>5.6</td><td>0.2</td><td>7.5</td><td>0.1</td></tr><tr><td>FZR14-F</td><td></td><td></td><td>6.7</td><td>19.5</td><td>76.1</td><td>0.397</td><td>79.4</td><td>0.85</td><td>6.3</td><td>0.2</td><td>7.4</td><td>0.1</td></tr><tr><td>FZR14-G</td><td></td><td></td><td>9.8</td><td>24.9</td><td>100.2</td><td>0.553</td><td>74.9</td><td>0.84</td><td>6.3</td><td>0.2</td><td>7.5</td><td>0.1</td></tr><tr><td><strong>FZR14</strong></td><td><strong>49.2586/73.0493</strong></td><td><strong>1,530</strong></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td><strong>7.5 ± 0.4</strong></td><td></td></tr><tr><td>FZR15-A</td><td></td><td></td><td>28.3</td><td>72.1</td><td>150.8</td><td>1.128</td><td>64.5</td><td>0.78</td><td>4.6</td><td>0.0</td><td>5.8</td><td>0.1</td></tr><tr><td>FZR15-D</td><td></td><td></td><td>27.3</td><td>78.1</td><td>130.8</td><td>1.180</td><td>70.5</td><td>0.80</td><td>4.7</td><td>0.0</td><td>5.9</td><td>0.1</td></tr><tr><td>FZR15-F</td><td></td><td></td><td>36.3</td><td>106.9</td><td>214.3</td><td>1.603</td><td>55.7</td><td>0.75</td><td>4.8</td><td>0.0</td><td>6.4</td><td>0.1</td></tr><tr><td><strong>FZR15</strong></td><td><strong>49.2715/73.0568</strong></td><td><strong>1,910</strong></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td><strong>6.0 ± 0.3</strong></td><td></td></tr><tr><td>FZR16-A</td><td></td><td></td><td>13.7</td><td>40.9</td><td>131.7</td><td>0.392</td><td>57.3</td><td>0.78</td><td>3.0</td><td>0.0</td><td>3.9</td><td>0.1</td></tr></tbody></table>
<table><caption>**Table 2**
*Continued*</caption><thead><tr><th>Sample N°</th><th>Latitude/Longitude (°S/°W)</th><th>Elevation (m)</th><th>U (ppm)</th><th>Th (ppm)</th><th>Sm (ppm)</th><th><sup>4</sup>He (nmol/g)</th><th>R<sub>s</sub> (μm)</th><th>F<sub>t</sub></th><th>Raw age (Ma)</th><th>Raw error (Ma)</th><th>Corrected age (Ma)</th><th>1σ error (Ma)</th></tr></thead><tbody><tr><td>FZR16-C</td><td></td><td></td><td>14.2</td><td>40.2</td><td>109.1</td><td>0.334</td><td>63.4</td><td>0.82</td><td>2.6</td><td>0.0</td><td>3.1</td><td>0.1</td></tr><tr><td>FZR16-E</td><td></td><td></td><td>12.2</td><td>38.5</td><td>68.9</td><td>0.443</td><td>78.0</td><td>0.83</td><td>3.8</td><td>0.0</td><td>4.6</td><td>0.1</td></tr><tr><td>FZR16</td><td>49.2654/73.0586</td><td>1,710</td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td>3.9 ± 0.7</td><td></td></tr></tbody></table>

*Note.* F_t = alpha-ejection correction factor (Farley et al., 1996). Bold numbers are mean ages calculated from the single-crystal replicates; 1σ error for mean ages is standard deviation of replicate ages. Analytical uncertainties are <1, ~3, and ~2% for respectively U, Th, and Sm measurements; and <1% for <sup>4</sup>He measurements.

temperature. Following each specified cooling path, the model first calculated an AHe age that was compared to the measured age (using averaged U-Th and R_s values for the effective model, Table S1 in Supporting Information S1). If the predicted age was within 1 standard deviation (SD) of the mean measured age (Table 3), a model <sup>4</sup>He/<sup>3</sup>He ratio evolution was calculated using the same analytical heating schedule as the sample and compared to observed ratios. This approach enables a random-search scheme to identify cooling histories that are compatible with the observations based on the computation of misfit statistics (M; mean of squared residuals weighted by the individual uncertainties in the ratio measurements; Schildgen et al., 2010); we set a misfit limit M ~ 2, which corresponds to the 99% confidence level. Thermal histories yielding M > 4 are excluded by the data, 2 < M < 4 are marginally acceptable (yellow lines in Figures 9b and 9d), and M < 2 are good fits to the data (green lines in Figures 9b and 9d).

# 4. Results

## 4.1. AHe and ZHe Thermochronological Data

We present all analytical details, as well as full AHe and ZHe data in Tables 1–3. For illustration, we report in Figure 5 both single-crystal and average AHe and ZHe corrected ages with 1σ standard deviation in age-elevation diagrams.

The FzR data set (Figure 5a) reveals single-crystal ZHe ages ranging from 6.1 ± 0.4 to 12.9 ± 0.8 Ma, and single-crystal AHe ages ranging between 3.1 ± 0.1 and 9.7 ± 0.6 Ma. ZHe ages are ~1–2 Myr older than AHe ages for samples at similar elevations, but there is no clear age-elevation relationship for each thermochronological system, which is likely linked with the small range of elevations in the elevation profile.

For the TdP massif, single-crystal AHe ages from the Central sector (Figure 5b) range from 3.1 ± 0.1 to 13.1 ± 0.3 Ma, and the metasedimentary samples reveal similar ages to the plutonic samples for this sector. There is an apparent break-in slope in the age-elevation relationship between high- (above 1,400 m a.s.l.) and mid- to low-elevation (below 1,400 m a.s.l.) samples of the Central sector around 7 Ma, potentially indicating an increase in rock cooling after the TdP pluton emplacement age (ca. 12.5 Ma). Samples from the North and West sectors have been collected at lower elevations (i.e., below 1,000 m a.s.l., Figure 5b), with AHe single-crystal ages ranging from 5.9 ± 0.1 to 11 ± 0.2 Ma in the West sector, and 6.3 ± 0.2 to 10.4 ± 0.1 Ma in the North sector. Interestingly, these two sectors show relatively older AHe ages compared to similar elevations for the Central sector (Figure 5b).

## 4.2. QTQt Thermal Inversion Results

The results of QTQt thermal inverse modeling and best fitting predicted versus observed ages are reported in Figures 6–8 with the relative probabilities of the expected (weighted mean) temperature-time (T-t) paths, as well as the maximum likelihood, maximum mode, and maximum posterior output thermal models for the lowest elevation sample of each profile presented in Figure 5. Thermal models of the lowest elevation sample, which we expect to have experienced the hottest thermal history amongst the multiple profile samples, are the most informative because they depend on the thermal history parameters for the inverse modeling (as well as for the highest elevation sample), and show the transition from the geothermal gradient to the atmospheric lapse rate (Gallagher, 2021), potentially bringing information about the late-stage relief evolution. The expected thermal
<table><caption>**Table 3**
Apatite (U-Th)/He (AHe) Data for Torres del Paine (TdP) Samples, Central Sector</caption><thead><tr><th>Sample N°</th><th>Latitude/ Longitude (°S/°W)</th><th>Elevation (m)</th><th>U (ppm)</th><th>Th (ppm)</th><th>Sm (ppm)</th><th><sup>4</sup>He (nmol/g)</th><th>R<sub>s</sub> (μ m)</th><th>F<sub>t</sub></th><th>Raw age (Ma)</th><th>Raw error (Ma)</th><th>Corrected age (Ma)</th><th>1σ error (Ma)</th></tr></thead><tbody><tr><td>04-JM-66ap1</td><td rowspan="4">50.97541/73.02521</td><td rowspan="4">1,802</td><td>8.2</td><td>10.1</td><td>442.4</td><td>0.506</td><td>53.6</td><td>0.73</td><td>8.6</td><td>0.3</td><td>11.7</td><td>0.4</td></tr><tr><td>04-JM-66ap2</td><td>5.1</td><td>11.2</td><td>291.3</td><td>0.288</td><td>54.8</td><td>0.73</td><td>6.7</td><td>0.2</td><td>9.2</td><td>0.2</td></tr><tr><td>04-JM-66ap3</td><td>8.8</td><td>13.3</td><td>397.6</td><td>0.504</td><td>39.8</td><td>0.64</td><td>7.6</td><td>0.3</td><td>11.9</td><td>0.4</td></tr><tr><td>04-JM-66ap4</td><td>32.4</td><td>40.9</td><td>575.7</td><td>2.183</td><td>55.5</td><td>0.74</td><td>9.6</td><td>0.2</td><td>13.1</td><td>0.3</td></tr><tr><td><strong>04-JM-66</strong></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td><strong>11.5 ± 1.6</strong></td><td></td></tr><tr><td>04-JM-67ap1</td><td rowspan="4">50.97492/73.02651</td><td rowspan="4">1,731</td><td>7.5</td><td>16.9</td><td>290.2</td><td>0.389</td><td>61.4</td><td>0.76</td><td>6.1</td><td>0.1</td><td>8.1</td><td>0.1</td></tr><tr><td>04-JM-67ap2</td><td>54.2</td><td>75.2</td><td>1,128.1</td><td>3.364</td><td>59.0</td><td>0.75</td><td>8.2</td><td>0.1</td><td>10.9</td><td>0.2</td></tr><tr><td>04-JM-67ap3</td><td>8.9</td><td>21.1</td><td>442.6</td><td>0.586</td><td>77.4</td><td>0.81</td><td>7.5</td><td>0.1</td><td>9.3</td><td>0.1</td></tr><tr><td>04-JM-67ap4</td><td>53.9</td><td>67.3</td><td>996.1</td><td>2.002</td><td>56.3</td><td>074</td><td>5.2</td><td>0.1</td><td>7.1</td><td>0.1</td></tr><tr><td>04-JM-67ap5</td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td>10.0</td><td>0.4</td></tr><tr><td><strong>04-JM-67</strong></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td><strong>9.1 ± 1.5</strong></td><td></td></tr><tr><td>04-JM-68ap1</td><td rowspan="2">50.97495/73.02651</td><td rowspan="2">1,731</td><td>4.2</td><td>11.6</td><td>162.0</td><td>0.278</td><td>76.7</td><td>0.80</td><td>7.3</td><td>0.2</td><td>9.1</td><td>0.3</td></tr><tr><td>04-JM-68ap2</td><td>8.4</td><td>15.4</td><td>309.2</td><td>0.327</td><td>54.3</td><td>0.73</td><td>5.0</td><td>0.2</td><td>6.8</td><td>0.3</td></tr><tr><td><strong>04_JM-68</strong></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td><strong>8.0 ± 1.6</strong></td><td></td></tr><tr><td>04-JM-71ap1</td><td rowspan="3">50.97405/73.02843</td><td rowspan="3">1,619</td><td>36.1</td><td>89.2</td><td>701.2</td><td>1.427</td><td>72.0</td><td>0.81</td><td>4.6</td><td>0.1</td><td>5.7</td><td>0.1</td></tr><tr><td>04-JM-71ap2</td><td>15.1</td><td>25.1</td><td>508.5</td><td>0.779</td><td>70.6</td><td>0.79</td><td>6.7</td><td>0.1</td><td>8.5</td><td>0.1</td></tr><tr><td>04-JM-71ap4</td><td>19.8</td><td>26.2</td><td>579.9</td><td>0.819</td><td>59.0</td><td>0.75</td><td>5.8</td><td>0.2</td><td>7.7</td><td>0.2</td></tr><tr><td><strong>04-JM-71</strong></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td><strong>7.3 ± 1.5</strong></td><td></td></tr><tr><td>04-JM-76ap1</td><td rowspan="3">50.97342/73.03557</td><td rowspan="3">1,310</td><td>15.2</td><td>21.6</td><td>466.1</td><td>0.567</td><td>46.2</td><td>0.69</td><td>5.1</td><td>0.3</td><td>7.3</td><td>0.4</td></tr><tr><td>04-JM-76ap2</td><td>21.1</td><td>41.2</td><td>503.4</td><td>0.678</td><td>41.1</td><td>0.65</td><td>4.0</td><td>0.2</td><td>6.1</td><td>0.4</td></tr><tr><td>04-JM-76ap3</td><td>41.2</td><td>42.2</td><td>539.9</td><td>1.238</td><td>46.8</td><td>0.69</td><td>4.4</td><td>0.1</td><td>6.4</td><td>0.2</td></tr><tr><td><strong>04-JM-76</strong></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td><strong>6.6 ± 0.6</strong></td><td></td></tr><tr><td>04-JM-87ap1</td><td rowspan="3">50.97449/73.03936</td><td rowspan="3">1,189</td><td>11.1</td><td>31.8</td><td>448.1</td><td>0.424</td><td>42.5</td><td>0.66</td><td>4.1</td><td>0.3</td><td>6.2</td><td>0.4</td></tr><tr><td>04-JM-87ap2</td><td>21.9</td><td>50.8</td><td>291.5</td><td>1.106</td><td>98.3</td><td>0.84</td><td>6.0</td><td>0.1</td><td>7.1</td><td>0.1</td></tr><tr><td>04-JM-87ap3</td><td>17.9</td><td>25.5</td><td>588.6</td><td>0.584</td><td>58.1</td><td>0.75</td><td>4.4</td><td>0.1</td><td>5.9</td><td>0.1</td></tr><tr><td><strong>04-JM-87</strong></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td><strong>6.4 ± 0.6</strong></td><td></td></tr><tr><td>04-JM-90ap1</td><td rowspan="3">50.97638/73.04372</td><td rowspan="3">1,042</td><td>28.0</td><td>22.8</td><td>254.3</td><td>0.675</td><td>50.5</td><td>0.71</td><td>3.8</td><td>0.1</td><td>5.3</td><td>0.2</td></tr><tr><td>04-JM-90ap2</td><td>12.7</td><td>28.5</td><td>352.8</td><td>0.631</td><td>66.7</td><td>0.77</td><td>5.9</td><td>0.2</td><td>7.7</td><td>0.2</td></tr><tr><td>04-JM-90ap3</td><td>22.3</td><td>29.9</td><td>312.2</td><td>0.823</td><td>63.7</td><td>0.77</td><td>5.2</td><td>0.2</td><td>6.8</td><td>0.2</td></tr><tr><td><strong>04-JM-90</strong></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td><strong>6.6 ± 1.2</strong></td><td></td></tr><tr><td>08-JL-385ap1</td><td rowspan="5">50.97669/73.10013</td><td rowspan="5">1,175</td><td>18.6</td><td>64.7</td><td>185.9</td><td>0.707</td><td>46.4</td><td>0.68</td><td>3.8</td><td>0.1</td><td>5.7</td><td>0.1</td></tr><tr><td>08-JL-385ap2</td><td>2.2</td><td>0.9</td><td>0.2</td><td>0.052</td><td>47.1</td><td>0.70</td><td>3.8</td><td>0.3</td><td>5.5</td><td>0.5</td></tr><tr><td>08-JL-385ap3</td><td>11.2</td><td>33.6</td><td>270.3</td><td>0.542</td><td>44.3</td><td>0.67</td><td>5.2</td><td>0.2</td><td>7.7</td><td>0.2</td></tr><tr><td>08-JL-385ap4</td><td>26.7</td><td>60.4</td><td>397.6</td><td>0.746</td><td>38.4</td><td>0.62</td><td>3.3</td><td>0.1</td><td>5.3</td><td>0.2</td></tr><tr><td>08-JL-385ap5</td><td>44.8</td><td>128.8</td><td>299.4</td><td>2.148</td><td>48.6</td><td>0.69</td><td>5.3</td><td>0.1</td><td>7.6</td><td>0.1</td></tr><tr><td><strong>08-JL-385</strong></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td><strong>6.4 ± 1.2</strong></td><td></td></tr><tr><td>07-JL-160ap2</td><td rowspan="3">50.97669/73.10013</td><td rowspan="3">1,175</td><td>5.1</td><td>23.4</td><td>208.1</td><td>0.269</td><td>57.1</td><td>0.73</td><td>4.6</td><td>0.1</td><td>6.2</td><td>0.1</td></tr><tr><td>07-JL-160ap3</td><td>5.3</td><td>21.8</td><td>145.7</td><td>0.231</td><td>57.7</td><td>0.74</td><td>4.0</td><td>0.1</td><td>5.4</td><td>0.1</td></tr><tr><td>07-JL-160ap4</td><td>9.1</td><td>30.8</td><td>181.9</td><td>0.392</td><td>51.1</td><td>0.71</td><td>4.4</td><td>0.1</td><td>6.2</td><td>0.1</td></tr><tr><td><strong>07-JL-160</strong></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td><strong>6.0 ± 0.4</strong></td><td></td></tr><tr><td>07-JL-165ap4</td><td rowspan="2">50.97677/73.10036</td><td rowspan="2">1,200</td><td>20.8</td><td>40.3</td><td>222.7</td><td>0.743</td><td>49.2</td><td>0.70</td><td>4.5</td><td>0.1</td><td>6.4</td><td>0.1</td></tr><tr><td>07-JL-165ap4</td><td>16.7</td><td>54.4</td><td>250.3</td><td>0.493</td><td>47.9</td><td>0.69</td><td>3.1</td><td>0.1</td><td>4.4</td><td>0.1</td></tr></tbody></table>
<table><caption>**Table 3**
Continued</caption><thead><tr><th>Sample N°</th><th>Latitude/Longitude (°S/°W)</th><th>Elevation (m)</th><th>U (ppm)</th><th>Th (ppm)</th><th>Sm (ppm)</th><th><sup>4</sup>He (nmol/g)</th><th>R<sub>s</sub> (μm)</th><th>F<sub>t</sub></th><th>Raw age (Ma)</th><th>Raw error (Ma)</th><th>Corrected age (Ma)</th><th>1σ error (Ma)</th></tr></thead><tbody><tr><td>07-JL-165ap4</td><td rowspan="2">50.97718/73.10087</td><td rowspan="2">1,275</td><td>42.7</td><td>93.2</td><td>317.6</td><td>1.749</td><td>50.6</td><td>0.71</td><td>5.0</td><td>0.1</td><td>7.1</td><td>0.1</td></tr><tr><td>07-JL-165ap4</td><td>21.3</td><td>43.4</td><td>285.7</td><td>0.606</td><td>42.0</td><td>0.65</td><td>3.5</td><td>0.1</td><td>5.4</td><td>0.2</td></tr><tr><td><b>07-JL-165</b></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td><b>5.8 ± 1.2</b></td><td></td></tr><tr><td>04-JM-49ap1</td><td rowspan="4">51.04051/73.08473</td><td rowspan="4">785</td><td>10.5</td><td>42.3</td><td>416.6</td><td>0.531</td><td>57.2</td><td>0.74</td><td>4.7</td><td>0.1</td><td>6.4</td><td>0.1</td></tr><tr><td>04-JM-49ap2</td><td>9.5</td><td>26.6</td><td>348.4</td><td>0.367</td><td>60.0</td><td>0.75</td><td>4.2</td><td>0.1</td><td>5.6</td><td>0.1</td></tr><tr><td>04-JM-49ap3</td><td>8.2</td><td>24.7</td><td>330.6</td><td>0.261</td><td>46.1</td><td>0.68</td><td>3.3</td><td>0.1</td><td>4.9</td><td>0.1</td></tr><tr><td>04-JM-49ap4</td><td>8.4</td><td>24.2</td><td>285.5</td><td>0.457</td><td>64.0</td><td>0.77</td><td>5.9</td><td>0.1</td><td>7.6</td><td>0.1</td></tr><tr><td><b>04-JM-49</b></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td><b>6.1 ± 1.2</b></td><td></td></tr><tr><td>04-JM-30ap1</td><td rowspan="2">50.94709/72.99015</td><td rowspan="2">1,830</td><td>9.1</td><td>12.8</td><td>501.3</td><td>0.488</td><td>68.8</td><td>0.79</td><td>7.2</td><td>0.0</td><td>9.18</td><td>0.4</td></tr><tr><td>04-JM-30ap2</td><td>15.8</td><td>38.0</td><td>671.1</td><td>0.821</td><td>51.2</td><td>0.71</td><td>6.0</td><td>0.2</td><td>8.46</td><td>0.2</td></tr><tr><td><b>04-JM-30</b></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td><b>8.8 ± 0.5</b></td><td></td></tr><tr><td>04-JM-23ap1</td><td rowspan="2">50.94713/72.99902</td><td rowspan="2">1,461</td><td>6.2</td><td>14.9</td><td>317.0</td><td>0.279</td><td>63.9</td><td>0.77</td><td>5.2</td><td>0.2</td><td>6.77</td><td>0.2</td></tr><tr><td>04-JM-23ap2</td><td>6.7</td><td>12.9</td><td>352.7</td><td>0.229</td><td>40.8</td><td>0.65</td><td>4.2</td><td>0.5</td><td>6.53</td><td>0.8</td></tr><tr><td><b>04-JM-23</b></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td><b>6.7 ± 0.2</b></td><td></td></tr><tr><td>13-TP-26-a</td><td rowspan="4">50.9728/72.8839</td><td rowspan="4">206</td><td>16.8</td><td>90.4</td><td>60.5</td><td>0.476</td><td>56.4</td><td>0.74</td><td>2.3</td><td>0.1</td><td>3.08</td><td>0.1</td></tr><tr><td>13-TP-26-z-BR</td><td>10.6</td><td>40.0</td><td>53.3</td><td>0.456</td><td>63.3</td><td>0.77</td><td>4.1</td><td>0.1</td><td>5.37</td><td>0.1</td></tr><tr><td>13-TP-26-year-BR</td><td>9.8</td><td>34.3</td><td>46.4</td><td>0.313</td><td>63.6</td><td>0.77</td><td>3.1</td><td>0.1</td><td>4.11</td><td>0.1</td></tr><tr><td>13-TP-26-x</td><td>11.3</td><td>45.6</td><td>51.8</td><td>0.386</td><td>59.2</td><td>0.75</td><td>3.2</td><td>0.1</td><td>4.22</td><td>0.1</td></tr><tr><td><b>13-TP-26</b></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td><b>4.2 ± 0.9</b></td><td></td></tr><tr><td><b>CH15-19</b></td><td>51.01057/72.94342</td><td>372</td><td>0.2</td><td>3.3</td><td>23.6</td><td>0.024</td><td>84.8</td><td>0.82</td><td>3.7</td><td>0.4</td><td><b>4.4 ± 0.5</b></td><td></td></tr><tr><td>CH15-TP21_x</td><td rowspan="3">50.98973/72.79723</td><td rowspan="3">268</td><td>4.9</td><td>30.1</td><td>51.4</td><td>0.179</td><td>69.6</td><td>0.78</td><td>2.7</td><td>0.1</td><td>3.42</td><td>0.1</td></tr><tr><td>CH15-TP21_y</td><td>11.3</td><td>13.1</td><td>34.1</td><td>0.240</td><td>55.4</td><td>0.73</td><td>3.0</td><td>0.1</td><td>4.14</td><td>0.1</td></tr><tr><td>CH15-TP21_z</td><td>3.7</td><td>16.7</td><td>13.1</td><td>0.126</td><td>67.1</td><td>0.78</td><td>3.0</td><td>0.1</td><td>3.91</td><td>0.1</td></tr><tr><td><b>CH15-TP21</b></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td><b>3.8 ± 0.4</b></td><td></td></tr><tr><td>CH15_TP30_x</td><td rowspan="3">50.94726/72.90560</td><td rowspan="3">1,053</td><td>1.7</td><td>11.3</td><td>16.1</td><td>0.142</td><td>78.9</td><td>0.81</td><td>5.7</td><td>0.1</td><td>7.19</td><td>0.1</td></tr><tr><td>CH15_TP30_y</td><td>1.5</td><td>11.3</td><td>19.5</td><td>0.125</td><td>66.2</td><td>0.77</td><td>3.4</td><td>0.1</td><td>7.01</td><td>0.1</td></tr><tr><td>CH15_TP30_z</td><td>1.1</td><td>11.826</td><td>12.423</td><td>0.097</td><td>60.4</td><td>0.75</td><td>4.5</td><td>0.3</td><td>6.1</td><td>0.4</td></tr><tr><td><b>CH15_TP30</b></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td><b>6.8 ± 0.6</b></td><td></td></tr></tbody></table>

Note. F_t = alpha-ejection correction factor (Farley et al., 1996). Bold numbers are mean ages calculated from the single-crystal replicates; 1σ error for mean ages is standard deviation of replicate ages. Analytical uncertainties are <1, ~3, and ~2% for respectively U, Th, and Sm measurements; and <1% for ⁴He measurements.

history for samples between the highest and the lowest elevation samples are obtained by linear interpolation based on the difference in elevation (Gallagher, 2021). These are shown in Figures S1–S3 in Supporting Information S1, together with the predicted output thermal gradients (within the imposed range of 35 ± 10°C/km, see Section 3.3). Using the output T-t paths and geothermal gradients, we can estimate exhumation rates for the different periods of time, that we have graphically determined based on major observed changes in output T-t paths (Figures 6–8 and Figures S1–S3 in Supporting Information S1).

QTQt inversion results for the FzR massif suggest a multi-stage cooling history (Figure 6 and Figure S1a in Supporting Information S1). The predicted versus observed ages shows a good fit of the AHe and ZHe data (Figure 6b), with some single-grain AHe ages not being well reproduced at low and high elevations (FZR-6, 12, -14, -15, and -16). Sample FZR-6, which presents the two most scattered single-crystals (FZR6-1 and FZR6-2) with respect to the best-fit modeled ages (Figure 6b), has suggested “parentless” He in excess (Vermeesch, 2008).
<table><caption>**Table 4**
Apatite (U-Th)/He (AHe) Data for Torres del Paine (TdP) Samples, West Sector</caption><thead><tr><th>Sample N°</th><th>Latitude/Longitude<br>(°S/°W)</th><th>Elevation<br>(m)</th><th>U<br>(ppm)</th><th>Th<br>(ppm)</th><th>Sm<br>(ppm)</th><th><sup>4</sup>He<br>(nmol/g)</th><th>R<sub>s</sub><br>(μm)</th><th>F<sub>t</sub></th><th>Raw<br>age (Ma)</th><th>Raw<br>error (Ma)</th><th>Corrected<br>age (Ma)</th><th>1σ<br>error (Ma)</th></tr></thead><tbody><tr><td>CH15-TP14_x</td><td></td><td></td><td>11.8</td><td>86.6</td><td>46.4</td><td>0.877</td><td>77.7</td><td>0.81</td><td>4.9</td><td>0.1</td><td>6.2</td><td>0.1</td></tr><tr><td>CH15-TP14_z</td><td></td><td></td><td>1.6</td><td>15.6</td><td>23.1</td><td>0.130</td><td>58.7</td><td>0.75</td><td>4.3</td><td>0.1</td><td>5.9</td><td>0.1</td></tr><tr><td><strong>CH15-TP14</strong></td><td><strong>51.25035/73.24379</strong></td><td><strong>112</strong></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td><strong>6.1 ± 0.2</strong></td><td></td></tr><tr><td>CH15-TP15_x</td><td></td><td></td><td>5.0</td><td>6.4</td><td>21.9</td><td>0.219</td><td>61.2</td><td>0.76</td><td>6.0</td><td>0.1</td><td>7.9</td><td>0.1</td></tr><tr><td>CH15-TP15_y</td><td></td><td></td><td>3.0</td><td>20.9</td><td>20.1</td><td>0.235</td><td>61.4</td><td>0.76</td><td>5.3</td><td>0.1</td><td>7.2</td><td>0.2</td></tr><tr><td>CH15-TP15_z</td><td></td><td></td><td>2.7</td><td>18.6</td><td>34.1</td><td>0.299</td><td>48.6</td><td>0.70</td><td>7.5</td><td>0.1</td><td>11.0</td><td>0.2</td></tr><tr><td><strong>CH15-TP15</strong></td><td><strong>50.98633/73.21632</strong></td><td><strong>235</strong></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td><strong>8.7 ± 2.0</strong></td><td></td></tr><tr><td>CH15_TP22_x</td><td></td><td></td><td>9.9</td><td>7.0</td><td>18.1</td><td>0.493</td><td>60.4</td><td>0.75</td><td>7.7</td><td>0.1</td><td>10.3</td><td>0.1</td></tr><tr><td>CH15_TP22_y</td><td></td><td></td><td>2.9</td><td>16.8</td><td>14.9</td><td>0.230</td><td>47.5</td><td>0.69</td><td>6.1</td><td>0.2</td><td>9.0</td><td>0.4</td></tr><tr><td>CH15_TP22_z</td><td></td><td></td><td>0.8</td><td>23.3</td><td>12.3</td><td>0.253</td><td>55.8</td><td>0.73</td><td>7.2</td><td>0.6</td><td>10.2</td><td>0.8</td></tr><tr><td><strong>CH15_TP22</strong></td><td><strong>51.00625/73.14149</strong></td><td><strong>746</strong></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td><strong>9.8 ± 0.7</strong></td><td></td></tr></tbody></table>

Note. Fᵢ = alpha-ejection correction factor (Farley et al., 1996). Bold numbers are mean ages calculated from the single-crystal replicates; 1σ error for mean ages is standard deviation of replicate ages. Analytical uncertainties are <1, ~3, and ~2% for respectively U, Th, and Sm measurements; and <1% for <sup>4</sup>He measurements.

Samples with excess He, however, can still hold geological significance, because radiogenic He can stay trapped in inclusions that are not susceptible to He diffusion (He et al., 2021). In this case, the excess He is probably prevenient from the magmatic cooling phase. Using both ZHe and AHe data, the expected, maximum mode, maximum posterior, and maximum likelihood thermal histories show an unconstrained cooling from the imposed initial magmatic temperature (275 ± 25°C) and age (16 ± 1 Ma) constraints, to a temperature range between 210 and 110°C at ca. 10 Ma (Figure 6a). Between ca. 10 and 8 Ma, the thermal models become well constrained, showing cooling from 150 to 100°C until 8.5–8 Ma by the expected models, probably indicating a cooling rate of 35°C/Myr, which varies between 30°C/Myr (maximum mode and posterior models), and 70°C/Myr (maximum likelihood model), depending on the thermal model (Figure 6a). Between 8 and 6 Ma, output thermal models predict a phase of slow cooling at 4–5°C/Myr. At ca. 6 Ma, thermal models predict an acceleration of cooling to ~70°C/Myr during a short period of ~1 Myr. The highest sample (FZR3, Figure S1 in Supporting Information S1) attained surface temperatures at Plio-Pleistocene times, while the lowest elevation sample (FZR13, Figure 6a) still experienced apparent monotonic cooling at ~5°C/Myr until the present day.

For the TdP massif, inverse thermal modeling using QTQt provided variable information on the regional exhumation history from the different sectors. For the Central sector, predicted AHe ages reproduce well the observed data (except for a few single-crystal ages of samples 04-JM-66, 04-JM-67, 04-JM-90, 04-JM-71, and 08-JL-385, Figure 7b and Figure S2b in Supporting Information S1). The two most scattered single-crystals (04-JM-

<table><caption>**Table 5**
Apatite (U-Th)/He (AHe) Data for Torres del Paine (TdP) Samples, North Sector</caption><thead><tr><th>Sample N°</th><th>Latitude/Longitude<br>(°S/°W)</th><th>Elevation<br>(m)</th><th>U<br>(ppm)</th><th>Th<br>(ppm)</th><th>Sm<br>(ppm)</th><th><sup>4</sup>He<br>(nmol/g)</th><th>R<sub>s</sub><br>(μm)</th><th>F<sub>t</sub></th><th>Raw<br>age (Ma)</th><th>Raw<br>error (Ma)</th><th>Corrected<br>age (Ma)</th><th>1σ<br>error (Ma)</th></tr></thead><tbody><tr><td>CH15-TP17_y</td><td></td><td></td><td>2.3</td><td>15.1</td><td>25.2</td><td>0.164</td><td>72.4</td><td>0.79</td><td>5.3</td><td>0.1</td><td>6.3</td><td>0.2</td></tr><tr><td>CH15_TP17_z</td><td></td><td></td><td>1.8</td><td>15.6</td><td>36.7</td><td>0.151</td><td>54.9</td><td>0.73</td><td>7.5</td><td>0.1</td><td>6.6</td><td>0.2</td></tr><tr><td><strong>CH15_TP17</strong></td><td><strong>50.81598/73.12218</strong></td><td><strong>369</strong></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td><strong>6.46 ± 0.2</strong></td><td></td></tr><tr><td>CH15_TP26_x</td><td></td><td></td><td>31.3</td><td>49.5</td><td>53.9</td><td>1.849</td><td>83.9</td><td>0.82</td><td>7.8</td><td>0.1</td><td>9.6</td><td>0.1</td></tr><tr><td>CH15_TP26_y</td><td></td><td></td><td>34.9</td><td>41.4</td><td>41.4</td><td>1.131</td><td>62.0</td><td>0.76</td><td>4.6</td><td>0.1</td><td>6.1</td><td>0.2</td></tr><tr><td>CH15_TP26_z</td><td></td><td></td><td>13.3</td><td>23.8</td><td>38.6</td><td>0.848</td><td>69.3</td><td>0.78</td><td>8.1</td><td>0.1</td><td>10.4</td><td>0.1</td></tr><tr><td><strong>CH15_TP26</strong></td><td><strong>50.85859/73.12883</strong></td><td><strong>916</strong></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td><strong>8.71 ± 2.27</strong></td><td></td></tr></tbody></table>

Note. Fᵢ = alpha-ejection correction factor (Farley et al., 1996). Bold numbers are mean ages calculated from the single-crystal replicates; 1σ error for mean ages is standard deviation of replicate ages. Analytical uncertainties are <1, ~3, and ~2% for respectively U, Th, and Sm measurements; and <1% for <sup>4</sup>He measurements.
<figure><img src="image_11.png" /><figcaption>**Figure 6.** QTQt thermal modeling outputs for the FzR massif (lowest elevation sample FZR-13, 1,410 m), derived from AHe and ZHe data (Tables 1 and 2). (a) Expected (weighted mean) model with relative probability of sample FZR-13 and its 95% confidence intervals (black solid lines), maximum-likelihood model (red line), maximum posterior model (brown line), and maximum mode model (white line). Gray dashed lines highlight key time periods with interpreted rapid cooling. The black box indicates the initial thermal constraint, and the red box represents general T-t priors. Note that output thermal histories for other FzR samples are linearly interpolated between the highest and the lowest elevation samples, shown in Figures S1a in Supporting Information S1. (b) Best-fitting predicted versus observed age diagram for the expected model with single-crystal AHe (dark green triangles) and ZHe (downward green triangle) ages of all samples of the FzR massif.</figcaption></figure>

67 ap2 and 04-JM-66 ap4), with respect to the best-fit, modeled ages (Figure 7b), have suggested excess “parentless” He, and their apparent older age is probably inherited from the magmatic cooling phase. Expected T-t paths from the dense AHe data set (Figure 5b) first show apparent rapid and unconstrained cooling from the imposed magmatic temperature/age constraints to the temperature range of 130–70°C (closure temperatures of the AHe system) at around 11.5 Ma, which is the oldest observed thermochronological AHe age in the profile and the time from which the output thermal models start to be constrained (Figure 7). We propose that this cooling signal reflects thermal relaxation after the shallow intrusion of the TdP Plutonic Complex, but no cooling rate can be measured since the initial temperature is imposed in the model (note that the Central sector includes two metasedimentary sample that we assume were re-heated by the same intrusion event, in accordance with Fosdick et al., 2013). From ca. 10.5 until ca. 6.5 Ma thermal models show slow cooling <1°C/Myr of TdP samples of the Central sector. From ca. 6.5 to 6 Ma thermal models predict an increase in cooling rate to 120°C/Myr (Figure 7a and Figure S2a in Supporting Information S1). This rapid cooling phase lasted for a relatively short time and at 6 Ma the highest elevation sample (04-JM-66, Figure S2a in Supporting Information S1) reached almost surface temperatures. The sample from the lowest elevation, 13-TP-26 (Figure 7a), remained in a steady thermal state until ca. 2 Ma, exhibiting a slow cooling rate of less than 1°C/Myr. However, it has experienced a sudden increase in cooling rate of up to 30°C/Myr from then to present-day. Due to the dense AHe data set and well-defined AHe age-elevation relationship (Figure 5b), the output thermal history of the late Miocene to Plio-Quaternary period is relatively well constrained.

For the North and West sectors of the TdP Massif (metasedimentary samples, Figure 8), the AHe data set is less dense and thus output model predictions are less constrained. For the North sector, the predicted AHe ages show an overall good fit to the observed data (Figure 8b and Figure S3b in Supporting Information S1), and there is no constraint on the thermal histories before 15 Ma (Figure 8a and Figure S3a in Supporting Information S1). The output thermal histories predict slow cooling between ca. 15 and 5 Ma, with an estimated cooling rate <2°C/Myr, followed by a faster cooling episode between ca. 5 and 4 Ma at 40–60°C/Myr. After 4 Ma the cooling rate decreases to 10°C/Myr for the low elevation sample (Figure 8a and Figure S3a in Supporting Information S1). For the West sector, the predicted AHe ages show an overall good fit to the observed data (Figure 8d and Figure S3d in Supporting Information S1). The expected T-t paths suggest a fast cooling at 24°C/Myr until ca. 11.5 Ma, which is not well constrained (Figures 8a and 8c). After ca. 11.5 Ma the output thermal models become well constrained predicting a slow

cooling at ca. 4°C/Myr. At around 7 Ma, cooling is predicted to accelerate to 25–30°C/Myr for a period of 2–2.5 Myr, bringing the high-elevation sample (Figure S3d in Supporting Information S1) near surface temperatures. After this time-interval, slow cooling of ~2°C/Myr is predicted until the present-day for the low-elevation sample (Figure 8c) with a potential faster late-stage cooling since ~0.5 Ma (with a cooling rate up to 80°C/Myr).

Potential cooling events before 15 Ma would not have been recorded by the AHe data set in the TdP massif, especially for the Central sector, given that the intrusion emplacement age is ~12.5 Ma. All TdP sectors show a short episode of fast cooling between 7 and 4 Ma, which is common for both the plutonic and the metasedimentary samples, and a potential delay (or lower precision in timing) for this event north of the TdP massif (North sector). A final stage of cooling is predicted for low-elevation samples of the Central and West sectors (Figures 7a and 8c), but the timing of onset for this episode is spatially variable, being earlier (at ~2 Ma) for the Central sector compared to the West sector (at 0.5 Ma). However, cooling histories below the closure temperature (T_c) of the
<figure><img src="image_12.png" /></figure>

**Figure 7.** QTQt thermal modeling outputs for the Central sector of the TdP massif (lowest elevation sample 13-TP-26, 206 m), derived from AHe data (Table 3). (a) Expected (weighted mean) model with relative probability of sample 13-TP-26 and its 95% confidence intervals (black solid lines), maximum-likelihood model (yellow line), maximum posterior model (brown line), and maximum mode model (green line). Gray dashed lines highlight key time periods with interpreted rapid cooling. The black box indicates the initial thermal constraint and the red box represents general T-t priors. Note that output thermal histories for other TdP samples are linearly interpolated between the highest and the lowest elevation samples, shown in Figure S2 in Supporting Information S1). (b) Best-fitting predicted versus observed age diagram for the expected model with single-crystal AHe ages of all samples of the Central sector of TdP massif (green triangles).

AHe system cannot be fully constrained (Shuster & Farley, 2004), and thus this late-stage cooling episode remains only a first-order estimate.

### 4.3. $^4$He/$^3$He Thermal Histories

Inverse thermal modeling of $^4$He/$^3$He data shows variable resolution for the two TdP samples (Central sector, Figures 5b and 9). $^4$He/$^3$He data resolution is relatively low for sample 13-TP-26 (due to large analytical uncertainties, Figure 9c), resulting in unconstrained output cooling histories over the last 10 Ma. However, significant cooling (40–50°C) had still occurred for this sample since 6-5 Ma (Figure 9d). $^4$He/$^3$He data resolution is much higher for sample 04-JM-90 (Figure 9a) and associated output cooling histories (Figure 9b) suggest fast cooling until any time between 10 and 7 Ma (poor information of the $^4$He/$^3$He data for temperatures higher than the AHe $T_c$; Shuster & Farley, 2004), before a quiescent period of low cooling rates. A late-stage cooling episode is recorded since 3-2 Ma, coherent with QTQt thermal predictions for the Central sector of TdP (Figure 7b). Observed differences between the AHe-constrained QTQt and $^4$He/$^3$He thermal model outputs are due to the sensitivity differences of the systems, with limited information from AHe data on constraining late-stage cooling episodes at temperatures lower than the AHe $T_c$ (Shuster & Farley, 2004).

## 5. Discussion

### 5.1. The Role of Ridge Subduction and Mantle Dynamics

In the FzR thermal models (Figure 6 and Figure S1 in Supporting Information S1), initial cooling between ca. 16 and 10 Ma potentially reflects post-magmatic thermal relaxation. However, our thermochronological AHe and ZHe ages are younger than 14 Ma (Figure 5a) and, thus, any earlier cooling/exhumation episodes cannot be quantified by our thermal modeling. An episode of rapid cooling between ca. 10 and 8 Ma is suggested by the thermal models (Figure 6a), which can be converted into a rock exhumation rate of ~1 km/Myr by dividing the cooling rate by the predicted geothermal gradient for that time interval (Figure S1 in Supporting Information S1). Additionally, the overlap between AHe and ZHe ages at 1,700–1,800 m elevation may also suggest fast cooling between ca. 10 and 8 Ma (Figure 5a). In this context, spreading ridge subduction beneath the FzR massif occurred sometime between 14 and 8 Ma (Breitsprecher & Thorkelson, 2009; Cande & Leslie, 1986; Guillaume et al., 2009, 2010, 2013; Guivel et al., 2006; Lagabrielle et al., 2010; Ramos, 2005; Thomson et al., 2001). During ridge subduction, dynamic and isostatic surface uplift may have occurred due to mantle upwelling and lithospheric thinning associated with increased heat flow (Ávila & Dávila, 2018, 2020; Ávila et al., 2023; Flament et al., 2015; Guillaume et al., 2009, 2010, 2013; V. A. Muller et al., 2024). With time, dynamic and isostatic uplift would have amplified the effects of surface erosion on rock exhumation over the incipient slab window, thus enhancing exhumation rate of the FzR massif during the Late Miocene (Figures 10c and 11a).

Interpreting rock exhumation as a signature of the opening of the slab window, however, has been debated in the literature (e.g., Georgieva et al., 2016, 2019; Guillaume et al., 2009, 2013; Stevens Goddard & Fosdick, 2019; Thomson et al., 2001). In the region near the CTJ (~46–47°S), for example, where faults were episodically active in the last 4–3 Myr, fast rock exhumation has been associated with compressive tectonics as a product of the collision of the oceanic ridge with the subduction trench (Andrić-Tomašević et al., 2021; Georgieva et al., 2016, 2019; Salze et al., 2018; Thomson et al., 2001). Recent studies of large scale low-temperature thermochronology modeling, however, showed that compressive events in the Patagonian Andes predate oceanic ridge collision with the subduction trench by 2–5 Myr (Stevens Goddard & Fosdick, 2019). In the FzR massif, the emplacement of the Chaltén Plutonic Complex around 16.5 Ma was synchronous with a phase of regional compression, which
<figure><img src="image_13.png" /><figcaption>**Figure 8.** QTQt thermal modeling outputs for the lowest elevation samples of the West and North sectors of the TdP massif, derived from AHe data (Tables 4 and 5). (a, c) Expected (weighted mean) model with relative probability of sample CH15-TP17 (369 m, a) and CH14-TP14 (112 m, c) and their 95% confidence intervals (black solid lines), maximum-likelihood model (red line), maximum posterior model (brown line), and maximum mode model (white line). Gray dashed lines highlight key time periods with interpreted rapid cooling. The red box represents general T-t priors. Note that output thermal histories for other TdP samples are linearly interpolated between the highest and the lowest elevation samples, shown in Figure S3a in Supporting Information S1. (b, d) Best-fitting predicted versus observed single-crystal AHe ages (green triangles) diagram for the expected models for all samples of North and West sectors of the TdP elevation profiles, respectively.</figcaption></figure>

occurred before the opening of the slab window at that latitude and is rather associated with a fast convergence rate (>13 cm/yr) between the Nazca Plate and the South American Plate (Breitsprecher & Thorkelson, 2009; Eagles & Scott, 2014; Pardo-Casas & Molnar, 1987; Somoza & Ghidella, 2005). The Chaltén Plutonic Complex, though, is not deformed, suggesting minor shortening after pluton emplacement, and implying that the opening of the slab window between 14 and 8 Ma did not generate important regional compression (Ramírez de Arellano et al., 2012). This is supported by a lack of evidence of compressive deformation in the Cenozoic foreland volcanic and sedimentary units (Espinoza et al., 2005; Ghiglione et al., 2009; Gorring et al., 2003). Instead, these units present discordant contacts and are tilted, characteristic of long-wavelength deformation from dynamic and isostatic surface uplift during mantle upwelling (Ávila & Dávila, 2020; Dávila et al., 2018; Flament et al., 2015; Guillaume et al., 2009). More recent numerical modeling using low-temperature thermochronology data predicts that dynamic and isostatic surface uplift forced by mantle upwelling are necessary to generate the present day topography, and calculate a rate of ~0.3 km/Myr of rock uplift for the last 12 Myr in the southern Patagonian Andes (Ávila & Dávila, 2020; Ávila et al., 2023). Agreeing with that, our results suggest that the calculated rock exhumation of ~1 km/Myr between 10 and 8 Ma could also be a product of mantle upwelling, which forces rock uplift and enhances surface erosion if the topography is already at high elevations (Figures 10c and 11a) (e.g., Ávila & Dávila, 2020; Ávila et al., 2023; Guillaume et al., 2009, 2010, 2013; Scalabrino et al., 2010). This is in accordance with a high topographic relief (>1 km) in the Patagonian Andes since at least the Late Miocene (Figure 10a; Blisniuk et al., 2006; Fosdick et al., 2013; Lagabrielle et al., 2007; Thomson et al., 2001). The delay between the age of the collision between the oceanic ridge and the subduction trench (~14–12 Ma), the increased surface and rock uplift, and consequent fast rock exhumation (~10–8 Ma) demonstrated by our study, is supported by field evidence of terraces tilting and numerical models (Guillaume et al., 2009), but could also be
<figure><img src="image_14.png" /><figcaption>**Figure 9.** $^{4}$He/$^{3}$He thermochronometry of the TdP massif (see Figures 3–5 for locations and details). $^{4}$He/$^{3}$He ratio evolution diagrams and model cooling paths are shown for 04-JM-90a (a, b) and 13-TP-26a (c, d). The measured $^{4}$He/$^{3}$He ratios of each degassing step (Rstep) are normalized to the bulk ratio (Rbulk) and plotted versus the cumulative $^{3}$He release fraction ($\Sigma\text{F}^{3}\text{He}$). Boxes indicate $\pm 1\sigma$ (vertical) and integration steps (horizontal). Colored lines show the predicted $^{4}$He/$^{3}$He ratio evolution diagrams (a, c) for arbitrary cooling paths between 150 and 10°C (b, d). Each colored path predicts the observed AHe age of the sample to within $\pm 1\sigma$ (cooling paths failing to predict the AHe age are not shown); red and yellow cooling paths are excluded by the $^{4}$He/$^{3}$He data, whereas green cooling paths are permitted (see Section 3.3 and Supporting Information S1, and Table S1 in Supporting Information S1 for analytical details).</figcaption></figure>

attributed to the uncertainty of dating the events of spreading ridge collision and subduction underneath the continent (Breitsprecher & Thorkelson, 2009; Guenthner et al., 2010; Ramos, 2005; Scalabrino et al., 2010).

A major change toward very slow cooling rates dominated the period between 8 and 6 Ma in the FzR massif, according to well constrained output thermal models, representing a converted exhumation rate of ~0.2 km/Myr (Figure 6a and Figure S1 in Supporting Information S1). Following the phase of oceanic ridge subduction and slab window opening sometime between 14 and 8 Ma at 49°S (Breitsprecher & Thorkelson, 2009), the slower subduction of the Antarctic Plate underneath the South American Plate (Eagles & Scott, 2014; Somoza & Ghidella, 2005) may have caused erosional quiescence between 8 and 6 Ma in both FzR and TdP regions (Figures 6–8).

In the TdP massif, tectonic compression with fold-and-thrust belt propagation and topographic growth occurred between 22 and 18 Ma (Fosdick et al., 2011, 2013; Lagabrielle et al., 2004), prior to the oceanic ridge collision with the trench at 51°S between 16 and 12 Ma (Breitsprecher & Thorkelson, 2009; Guillaume et al., 2009), a period that is not resolved by our thermochronological data sets. In the Central sector, fast cooling rate after pluton emplacement until ca. 11 Ma might be due to thermal relaxation of the pluton, although our AHe data have no resolution to constrain the duration and magnitude of this episode (thermal conditions higher than the AHe closure
<figure><img src="image_15.png" /><figcaption>**Figure 10.** Schematic cross-sections of the spatio-temporal evolution of pluton exhumation for the Fitz Roy and Torres del Paine massifs. (a) In the late Miocene around 12 Ma, the Chaltén Plutonic Complex was already emplaced and partly exhumed from its intrusion depth estimated at ~6 km (Ramírez de Arellano, 2011), this episode of exhumation is proposed to be forced by a combination of dynamic and isostatic uplift from the mantle upwelling underneath, and fluvial erosion. (b) The Torres del Paine laccolith was being emplaced around 12.5 Ma (Leuthold et al., 2012) at ~4 km depth (Putlitz et al., 2001) and the mantle upwelling was already attenuated due to the northward migration of the opening slab window. Bedrock deformation has occurred before the time constrained by the results of this study, as shown by the inactive faults. (c, d) Patagonian glaciation at the Miocene-Pliocene transition at ~6 Ma, which generated ~1.8 km of pluton exhumation in the Fitz Roy massif between 6 and 5 Ma (c), and 1.2–1.5 km of pluton exhumation between 7 and 4 Ma in the Torres del Paine massif (d), with a higher exhumation rate (~3 km/Myr) in the Central sector, compared to the 0.6–1.4 km/Myr exhumation rate in the West and North sectors of TdP. (e, f) Present day topography and geological stratigraphy and structures with the approximate location of AHe and ZHe samples from this study (white circles), and the Southern Patagonian Icefield outcropping in both regions. Inactive fault planes were interpreted from Giacosa et al. (2013) for the Fitz Roy massif (e), and Fosdick et al. (2013) for the Torres del Paine massif (f).</figcaption></figure>

temperature). The unconstrained initial cooling in the West and North sectors is due to the lack of AHe data older than 12 Ma and no arbitrarily-prescribed initial conditions for the thermal models. AHe ages between ca. 12 and 10 Ma (Figure 5b) suggest that high-elevation samples for TdP were in (or near) the AHe partial retention zone between ca. 12 and 9 Ma (Figures 5b, 7, and 10f), in agreement with similar AHe ages obtained in the TdP metasedimentary units (Fosdick et al., 2013) that were associated with thermal resetting during pluton emplacement. Additionally, the slow cooling interval predicted by our thermal models between ca. 12 and 7 Ma, is coincident with the slow convergence rate between the South American and the Antarctic plates (Eagles & Scott, 2014), and this period is interpreted as a tectonic quiescence. Any forcing from the oceanic ridge collision with the subduction trench and/or mantle upwelling might have been already attenuated when the Torres del Paine laccolith was being emplaced (Figure 10f).

## 5.2. The Role of Late Miocene to Plio-Quaternary Glacio-Fluvial Erosion

Inverted AHe ages from the FzR and the TdP massifs constrain an episode of abrupt acceleration in rock exhumation between 7 and 4 Ma, when the convergence rate of the Antarctic Plate with respect to the South America Plate was constant, and the tectonic regime was regionally stable (Eagles & Scott, 2014; Somoza & Ghidella, 2005). This is in agreement with the observed break in slope in the age-elevation profile of the Central sector of TdP at ~7 Ma and 1,400-m elevation (Figure 5b), potentially indicating an increase in rock exhumation rate at this time. The predicted onset of an increased exhumation rate is synchronous with the onset of major glaciations at 7-6 Ma in the Southern Patagonian Andes (Figures 10b, 10e, and 11b), supported by stratigraphic,
<figure><img src="image_16.png" /></figure>

**Figure 11.** Block diagram with the interpretation of the geodynamic and topographic evolution of the Southern Patagonian Andes from the late Miocene to the Quaternary. (a) During the late Miocene (~12 Ma), the spreading ridge between the Nazca Plate (NZ) and the Antarctic Plate (AT) was subducting beneath the South America Plate (SAM) at ~49°S, mantle upwelling caused dynamic and isostatic surface uplift, forcing part of the rock exhumation in the FzR massif. The greater part of the Southern Patagonian Batholith (SPB), and the Chaltén Plutonic Complex in the Fitz Roy region were already emplaced, whereas the Torres del Paine Plutonic Complex was being emplaced. (b) At the Miocene-Pliocene transition (7–6 Ma), the spreading ridge was subducting at ~48°S, and associated dynamic and isostatic uplift must be occurring north of the studied regions. The Antarctic Plate was subducting at the latitudes of the studied regions, and the fold-and-thrust belt and the intrusions were being exhumed mainly due to glacial erosion after the onset of Patagonian glaciations. Regional fast exhumation due to intensified glacial erosion encompasses also the regions around the CTJ (Andrić-Tomašević et al., 2021; Christeleit et al., 2017; Georgieva et al., 2016, 2019; Guillaume et al., 2009, 2013; Thomson et al., 2001; Willett et al., 2020). (c) During the Quaternary, the mountain belt was being exhumed mainly due to glacio-fluvial erosion, carving deep valleys and leaving the mountain peaks far above the bottom of the valleys. The studied plutonic complexes must be at or near the surface. The subducting oceanic ridge was around the CTJ latitudes (47–46°S), generating also important rock exhumation associated with tectonic activity and mantle upwelling at those latitudes (Andrić-Tomašević et al., 2021; Georgieva et al., 2016, 2019; Guillaume et al., 2009, 2013; Thomson et al., 2001). Black arrows highlight the regions exhuming rapidly. Orange arrows highlight the region of main mantle upwelling.
geomorphologic (Mercer and Sutter, 1982; Lagabrielle et al., 2004, 2007, 2009, 2010; Rabassa, 2008; Rabassa et al., 2005, 2011), and thermochronological evidence (Andrić-Tomašević et al., 2021; Christeleit et al., 2017; Fosdick et al., 2013; Georgieva et al., 2016, 2019; Glodny et al., 2008; Ronda et al., 2022; Thomson et al., 2001, 2010, 2010, 2010; Willett et al., 2020). The Andean topography quickly responded to the transition from fluvial-dominated to glacial-dominated erosional processes (Figures 10b–10f), as proposed for other alpine environments (Champagnac et al., 2014; Egholm et al., 2009; Herman et al., 2013; Shuster et al., 2005, 2011; Valla et al., 2011).

The magnitude and duration of this event depend on the analyzed sector of the TdP and FzR massifs, with estimated exhumation rates varying between 0.6 km/Myr (West sector of TdP), 1.4 km/Myr (North sector of TdP), 1.8 km/Myr (FzR), and 3 km/Myr (Central sector of TdP). In the Fitz Roy, the exhumation rate of 1.8 km/Myr lasts from 6 to 5 Ma (Figure 6a and Figure S1 in Supporting Information S1). In the North sector of TdP, the exhumation of 1.4 km/Myr lasts from 5 to 4 Ma (Figure 8a and Figure S3 in Supporting Information S1). Despite the limited exhumation (0.6 km/Myr) in the West sector of TdP, it lasted 2.5 Myr according to thermal modeling (Figure 8c and Figure S3 in Supporting Information S1), resulting in ~1.2 km of total exhumation between 7 and 4.5 Ma. In the Central sector of TdP, instead, the exhumation rate is higher (3 km/Ma) but for a shorter interval of time between 6.5 and 6 Ma (Figure 7a and Figure S2 in Supporting Information S1), resulting in 1.5 km of total exhumation. The slower Miocene-Pleistocene exhumation rate in the North and West sectors of the TdP massif compared to the Central sector (Figures 7 and 8) could reflect topography unfavorable to glacial erosion in the North and West sectors, such as broad valleys with longitudinal slopes, limiting ice sliding (Pedersen & Egholm, 2013; Shuster et al., 2011). This is the case for the present-day spatial topographic organization in the TdP massif and, thus, implies that topographic spatial distribution has not changed significantly over the last 7-4 Ma. In addition, a plausible explanation for a persistent differential topography is the higher mechanical strength of the granitic rocks of the Central sector compared with the metapelitic rocks of the North and West sectors (Bernard et al., 2019; Palumbo et al., 2010). The Central sector granitic rocks are likely more erosion-resistant and thus have not been eroded drastically since 6 Ma. The West and North sector metapelitic rocks, instead, tend to suffer more spread and continuous erosion, forming smooth and broad valleys. Slow ice sliding in broad valleys would be due partly to the lateral spreading of ice in wide channels, and partly to thick ice that favors flow by internal deformation (Figure 10e) (Pedersen & Egholm, 2013; Shuster et al., 2011). Headward propagation of glacial erosion can be more effective where the ice sliding is confined in narrow and deeply-incised valleys such as for the Central sector (Figure 10e) (Shuster et al., 2011; Sternai et al., 2013). Efficient erosion of high-elevation topography by glacial and periglacial processes during the late Miocene would have resulted in a net decrease of ice accumulation area and hence in ice extent, ice flux, and consequently in glacial erosion (Pedersen & Egholm, 2013; Shuster et al., 2011; Sternai et al., 2013). Such negative feedback has also been proposed to explain the late Quaternary gradual shrinking of the Southern Patagonian Icefield (Kaplan et al., 2009), and likely explains the short-lived erosion pulse that our results and previous studies identified in Southern Patagonia at the late Miocene/Pliocene transition (Christeleit et al., 2017; Willett et al., 2020).

Our estimated exhumation rates of 1–2 km/Myr for the 7–4 Ma interval in both TdP and FzR massifs are of similar magnitude compared with exhumation rates recorded in northern regions near the current position of the CTJ (46–47°S) (Andrić-Tomašević et al., 2021; Georgieva et al., 2016, 2019; Guillaume et al., 2013; Ronda et al., 2022; Thomson et al., 2001). At those latitudes, the onset of Patagonian glaciations was concurrent with tectonic activity in the Liquiñe-Ofqui Fault Zone (Georgieva et al., 2016, 2019), out-of-sequence thrusting (Andrić-Tomašević et al., 2021; Ghiglione et al., 2019; Ronda et al., 2022; Thomson et al., 2001), and/or mantle upwelling since 3 Ma (Guillaume et al., 2009, 2013) as previously discussed (Section 5.1). In our study regions (49°S and 51°S), the tectonic context has been relatively stable since at least 8 Ma (Breitsprecher & Thorkelson, 2009; Eagles & Scott, 2014; Somoza & Ghidella, 2005), with lack of evidence for tectonic shortening (Fosdick et al., 2011, 2013), and with attenuated mantle upwelling (Guillaume et al., 2009). Given that our constrained exhumation rates in the FzR and TdP regions are of similar magnitude as those reported for latitudes between 46 and 47°S, where the tectonic and climate forcing to rock exhumation are supposed to interact, we propose that glacial erosion is the main control on rock exhumation in the retroarc of the Patagonian Andes, at least for the period between 7 and 4 Ma. An alternative hypothesis would be that glacial erosion and imprint on the topography has increased southward since 7 Ma, likely related to cooler temperatures toward the South Pole, thus surpassing the combined effects of tectonic (i.e., mantle upwelling and/or deformation) and erosional (glaciations) processes in the northern sector (46–47°S) of the Southern Patagonian Andes.
Pliocene low exhumation rates (<0.1 km/Myr) recorded in the FzR and TdP massifs indicate erosional quiescence following transient high glacial erosion rates in the Southern Patagonian Andes (Figures 10a and 10d) (Christeleit et al., 2017; Willett et al., 2020). Quaternary fast exhumation rates of ~1 km/Myr is recorded by apatite ⁴He/³He and AHe data in the Central and West sectors of the TdP massif (especially for low-elevation samples), with a predicted onset varying between 3 and 0.5 Ma depending on the sample location (Figures 7–9, and Figures S2–S4 in Supporting Information S1). An increase in exhumation rates is recorded in several orogens worldwide since around 2 Ma (Herman et al., 2013), and has been interpreted as the onset of enhanced glaciations at mid latitudes, including Fjordland in New Zealand (Shuster et al., 2011), Alaska (Berger et al., 2008), British Columbia (Shuster et al., 2005), and the European Alps (Fox et al., 2015, 2016; Glotzbach et al., 2011; Haeuselmann et al., 2007; Valla et al., 2011). A possible climatic trigger of enhanced glaciations might be the observed global cooling at ca. 2 Ma (Lisiecki & Raymo, 2007) and/or increase in the duration and asymmetry of glacial-interglacial cycles at ~1.2 Ma (B. De Boer et al., 2010; Lisiecki, 2010; Lisiecki & Raymo, 2007). By periodical switching between glacial and fluvial conditions, and by changing associated vegetation and soil cover, geomorphic processes would be transient (Herman & Champagnac, 2016; Molnar, 2004), maintaining landscape disequilibrium and, in turn, enhancing erosion rates (Champagnac et al., 2014; Egholm et al., 2009; Pedersen & Egholm, 2013). North of the studied region (46–47°S), stratigraphic records from moraine deposits indicate a shift in the drainage network after 3 Ma (Georgieva et al., 2019; Lagabrielle et al., 2004, 2010; Willett et al., 2020), resulting in a major landscape change from a smooth piedmont surface with extensive icefields in the foreland to long west-east oriented and channelized glacial lobes (Lagabrielle et al., 2010). Tectonic uplift of the eastern Patagonian foreland might have influenced or facilitated the shift in the landscape features leading to the deepening of glacial valleys from west to east. However, this process was mainly associated with regions located north of 47°S, where compression, transpression, and mantle upwelling due to interactions between the oceanic ridge and the trench may have played a recent (<3 Ma) role (Andrić-Tomašević et al., 2021; Georgieva et al., 2019; Lagabrielle et al., 2010). In the Torres del Paine massif (51°S), instead, the observed late-stage acceleration in rock exhumation was most likely associated with Plio-Quaternary global cooling and a shift in glacial-interglacial cyclicity (A. M. De Boer et al., 2010; Lisiecki, 2010; Lisiecki & Raymo, 2007), enhanced glacial erosional processes, and icefield drainage reorganization in Southern Patagonia (Figures 10d and 11c). In the Fitz Roy massif, however, our results do not show any clear evidence for acceleration of rock cooling after 5.5 Ma. This might be associated with a lack of low-temperature thermochronological data from low elevations (in glacier ablation areas), and also with the lack of apatite ⁴He/³He data, which are more sensitive to late-stage cooling events (Shuster et al., 2011).

**6. Conclusions**

The Fitz Roy (49°S) and Torres del Paine (51°S) Miocene plutonic complexes in the retroarc of the Southern Patagonian Andes were exhumed in transient episodes with rates as high as 2–3 km/Myr, forced by large-scale tectonics and climate-driven erosion processes. The 200 km that separates these complexes along the North-South trend of the Southern Patagonian Andes allowed us to identify spatial and temporal variations of these processes. We found thermochronological evidence for the effects of oceanic ridge subduction and mantle upwelling on dynamic and isostatic surface uplift causing ~2 km of rock exhumation between ca. 10 and 8 Ma under pre-glacial fluvial erosion conditions in the Fitz Roy massif. Low-temperature thermochronological data from the Torres del Paine massif did not record this geodynamic event, possibly due to an already attenuated surface response to ridge subduction and mantle upwelling at 51°S.

The onset of major glaciations in Southern Patagonia generated a regional but short-lived signal of rapid rock exhumation in the Fitz Roy and Torres del Paine massifs (at 49–51°S) between 7 and 4 Ma. This was followed by a period of slow rock exhumation from the early Pliocene until the Quaternary, highlighting erosional quiescence and possibly reflecting bedrock shielding by extensive icefields covering the Southern Patagonian Andes. A Quaternary episode of accelerated rock exhumation is partly recorded in our thermochronological data sets, and coincides with worldwide increased mountain erosion ascribed to enhanced cooling and/or intense glacial-interglacial cycles. This climatic transition generated a geomorphological shift from smooth landforms to deeply incised glacial valleys, leaving the high elevations of the Torres del Paine and the Fitz Roy plutonic complexes standing far above nearby valley bottoms.
## Acknowledgments

This work also been supported by the Italian Ministry of Education, Universities and Research, MUR, (Project Dipartimenti di Eccellenza, 2023-2027, TECLA, Department of Earth and Environmental Sciences, University of Milano-Bicocca), Fondazione Cariplo and Fondazione CDP (Grant 2022-1546_001). The Université Grenoble Alpes, the French CNRS, the INSU SYSTER project, and the ECOS-SUD project A15U02 also supported this work. V.A.P.M. acknowledges the ERASMUS+ program for the mobility grant to visit the Université Grenoble Alpes, the TRB team in ISTerre, and the TANGO Project NSF Grant 2020935 at the University of Arizona. P.G.V. acknowledges funding support from the Swiss National Science Foundation SNSF (Grant PP00P2_170559) and the French ANR-PIA program (ANR-18-MPGA-0006). For permission to work and sample in P.N. Torres del Paine, K.C. is grateful to two Chilean agencies: Corporación Nacional Forestal (Resolución 15/2015), and Dirección Nacional de Fronteras y Límites del Estado. In the P.N. Los Glaciares (Fitz Roy massif), C.S., and M. C.G. are grateful to the Administración de Parques Nacionales Argentina (permit 31-DRPA to M.C.G.), and all the people who helped us on the field, specially the mountain guide Santiago Arias in El Calafate, Argentina. We acknowledge the support of The Martin Family Foundation to K.M.C., The Ann and Gordon Getty Foundation to D.S., and the Chilean Comisión Nacional de Investigación Científica y Tecnológica award to U.C. Berkeley. M. Salze and R. Pinna-Jamme are warmly thanked for the help during AHe analysis at GEOPS. We acknowledge M. Fox by reading and adding valuable comments to the first version of the manuscript, and S. Thomson by helping with the final revision process. We thank the editor L. Jolivet, associate editor M. Jolivet, and the two anonymous reviewers, for their constructive reviews that helped to improve the quality and clarity of our manuscript.

## Data Availability Statement

The reduced apatite and zircon (U-Th)/He data used in this paper are available in the Bicocca Open Archive Research Data repository: https://doi.org/10.17632/w92bbxbcgd.2 (V. A. Muller et al., 2024). Please access version 2 of this data set, because version 1 is outdated.

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