# Resolving mid- to upper-crustal exhumation through apatite petrochronology and thermochronology

Gilby Jepson, Barbara Carrapa, Sarah W.M. George, Antoine Triantafyllou, Shana Egan, Kurt Constenius, George Gehrels, Mihai Ducea

► To cite this version:

Gilby Jepson, Barbara Carrapa, Sarah W.M. George, Antoine Triantafyllou, Shana Egan, et al. Resolving mid- to upper-crustal exhumation through apatite petrochronology and thermochronology. Chemical Geology, 2021, 565, pp.120071. 10.1016/j.chemgeo.2021.120071. hal-04880662

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# Resolving mid- to upper-crustal exhumation through apatite petrochronology and thermochronology

Gilby Jepson*<fnref n="1" />, Barbara Carrapa<fnref n="1" />, Sarah W. M. George<fnref n="1" />, Antoine Triantafyllou<fnref n="1" />, <fnref n="2" />, <fnref n="3" />, Shana M. Egan<fnref n="1" />, Kurt N. Constenius<fnref n="1" />, George E. Gehrels<fnref n="1" />, and Mihai N. Ducea<fnref n="1" />, <fnref n="4" />

November 25, 2020

## Abstract

Double-dating using the apatite U-Pb and fission-track systems is becoming an increasingly popular method for resolving mid- to upper-crustal cooling. However, these thermochronometers constrain dates that are often difficult to link through geological time due to the large difference in temperature window between the two systems (typically >250°C). In this study, we apply apatite U-Pb, fission-track, and apatite and whole rock geochemistry to fourteen samples from four tectonic domains common in Cordilleran orogenic systems: (1) basement cored-uplifts, (2) plutons intruded through a thick crustal column, (3) metamorphic core complexes and associated detachment faults, and (4) rapid, extrusive volcanic cooling, in order to provide a link between *in situ* geochemical signatures and cooling mechanisms. Comparisons of trace element partitioning between apatite and whole rock geochemistry provide insights into initial apatite-forming processes and/or subsequent modification. Apatite trace element geochemistry and the Th/U and La/Lu<sub>N</sub> ratios provide tools to determine if an apatite is primary and representative of its parent melt or if it has undergone geochemical perturbation(s) after crystallization. Further, we demonstrate that by using a combined apatite U-Pb, FT, trace element, and whole rock geochemistry approach it is possible to determine if a rock has undergone monotonic cooling since crystallization, protracted residence in the middle crust, and provide unique structural

<footnote n="1">¹Department of Geosciences, University of Arizona, Tucson, Arizona, USA</footnote>

<footnote n="2">²Department of Geology, University of Liège, 4000 Sart Tilman, Belgium</footnote>

<footnote n="4">⁴Faculty of Geology and Geophysics, University of Bucharest, Bucharest, Romania.</footnote>

<footnote n="3">³Laboratoire de Géologie de Lyon : Terre, Planètes, Environnement, Université de Lyon, France</footnote>

<footnote>*corresponding author (gjepson@arizona.edu)</footnote>
information such as the history of detachment faulting. Insights provided herein offer new applications for apatite thermochronology.

# 1 Introduction

Due to the low-solubility of phosphorous in silicate-rich melts and its incompatible behavior with many major rock-forming mineral phases, apatite is a common accessory mineral found in a wide range of crustal rocks (Chew & Spikings 2015). Apatite occurs as an accessory phase in many igneous and metamorphic rocks, and can be abundant in clastic sedimentary rocks (e.g., Henrichs et al. 2018, O'Sullivan et al. 2020, Chew et al. 2020). As well as being a readily available mineral phase, apatite incorporates trace amounts of radiogenic elements such as $^{238}$U, $^{235}$U, and $^{232}$Th, making it a useful geo- and thermo- chronometer (Carrapa et al. 2009, Cherniak et al. 1991). Apatite is used as a low-temperature thermochronometer through fission-track and (U-Th-Sm)/He) thermochronometric techniques, which constrains the thermal evolution of a sample through the upper crust thus providing insights into erosional and tectonic processes (e.g., Fitzgerald et al. 1993, Ehlers 2005, Reiners & Brandon 2006, Braun et al. 2006)

Improvement in Laser Ablation-Inductively Coupled Plasma-Mass Spectrometry (LA-ICP-MS) instrumentation and standards have facilitated progress in the development and application of apatite uranium-lead thermochronology (AUPb, e.g., Carrapa et al. 2009, Thomson et al. 2012, Chew et al. 2014). The apatite U-Pb system is controlled by the thermally activated volume diffusion of radiogenic lead at temperatures between (~375 and 570°C; e.g., Cochrane et al. 2014), thus, AUPb acts as a medium-temperature thermochronometer (e.g., Cherniak et al. 1991, Watson & Cherniak 2013). However, apatite is also susceptible to recrystallization, emphasizing the need for thorough interpretation of apatite U-Pb ages (e.g., Spear & Pyle 2002, Kusebauch et al. 2015, Odlum & Stockli 2020). The ~375-570°C temperature window is valuable as it provides information on the cooling history of a rock through the middle crust making the apatite U-Pb system a useful chronometer to a variety of geological questions such as tectonic exhumation (e.g., Odlum & Stockli 2019, 2020), ore formation (e.g., Belousova et al. 2002a, Glorie et al. 2019), and metamorphism (e.g., Schneider et al. 2015). Through the application of LA-ICP-MS, U-Pb and FT/(U-Th-Sm)/He (~60-120°C and ~80-45°C, respectively) apatite thermochronology can be integrated together in the form of double and triple-dating of a single mineral to constrain a rocks middle to upper crustal evolution (e.g., Jepson et al. 2018, Horne et al. 2019). Despite the use of the single-grain double and triple-dating method, it remains difficult to connect the ~500-60°C thermal and crustal history
of a rock in complex tectonic settings due to substantial ambiguities in thermal pathways (i.e. monotonic cooling *versus* more complex pathways such as reheating, Figure 1).

In addition to U-Pb dating, *in situ* trace and rare earth elements (REE) concentrations can also be analyzed simultaneously on high resolution single collector ICP-MS. While similar to procedures have been applied to zircons (Belousova et al. 2002b, Grimes et al. 2015, Chapman et al. 2016, Balica et al. 2020), trace elements in apatite are far more diagnostic of their host rock and a faithful recorder of the parent melt geochemistry (e.g. Prowatke & Klemme 2006, Jennings et al. 2011, Mark et al. 2016, O'Sullivan et al. 2018, Gillespie et al. 2019). In particular, trace elements can be used in conjunction with apatite U-Pb thermochronology to understand magma formation and storage (Nathwani et al. 2020) and have even been shown to preserve complex diffusion profiles, demonstrating the application of apatite as a novel middle-crust petrochronometer (Seymour et al. 2016, Smye et al. 2018). The variety in radiometric dating and geochemical applications suggests that apatite can not only provide information on timing of cooling, but shed light on the mechanisms driving cooling and on the original whole rock chemistry (e.g., Bruand et al. 2016, 2017, O'Sullivan et al. 2020).

Here, we apply U-Pb, FT, and trace element analyses to single apatite grains, along with whole rock geochemistry to samples taken from different tectonic domains within the North American Cordillera: (1) cratonic-basement from Laramide basement cored-uplifts (Stevens et al. 2016); (2) Laramide-aged plutons related to a thickened crust (DeCelles 2004); (3) metamorphic core complex (MCC) and associated detachment faults (Lister & Davis 1989), and (4) syn- to post- core-complex extrusive volcanism (Best et al. 2009); with the aim of exploring the relationships between apatite U-Pb, apatite fission-track thermochronology and their trace element signatures as related to different tectonic processes (Figure 1 and 2).
<figure><img src="image_2.png" /></figure>

Figure 1: Schematic cross-section illustrating typical tectonic processes in a Cordilleran type margin and commonly associated thermal histories associated with each tectonic process. A) prolonged emplacement at mid-to-low temperatures in the shallow crust along basement-cored uplifts, B) steady monotonic cooling from mid-to-low temperatures via erosion of an intrusive pluton that has formed in thickened crust, C) multiple phases of mid-temperature reheating and cooling associated with tectonic exhumation along detachment faults, and D) extrusive volcanic rocks preserving rapid monotonic cooling.

## 2 Tectonic setting of the North American Cordillera

The North American Cordillera extends from Alaska in the north down to Mexico in the south and is part of the Circum-Pacific orogenic belt (e.g., DeCelles 2004, DeCelles et al. 2009). The North American Cordillera formed largely in response to continuous Late Paleozoic to Cenozoic subduction deforming thick passive margin strata atop cratonic basement (e.g., Coney et al. 1980, Dickinson 2004, Lawton 2008, Yonkee & Weil 2012, Coney et al. 1987, Wernicke et al. 1987, Sonder & Jones 1991). The well-documented tectonic record of deformation, thickening, magmatism, and extensional collapse preserved along the North American Cordillera (e.g. Price 1986, Coney 1987, Wernicke et al. 1987, Sonder & Jones 1991) make it and ideal setting for fingerprinting cooling and geochemical changes characteristic of different tectonic settings.
## 2.1 Archean Craton

The Wyoming craton is a segment of Archean crust which preserves predominantly rocks of ca. 3.5-2.6 Ga and up to ca. 4.0 Ga (Hoffman 1988, Frost et al. 1998, Mueller & Frost 2006). The Wyoming craton collided with Laurentia in the Paleoproterozoic during the Trans-Hudson Orogeny to form the core of the North American continent (e.g., Mueller & Frost 2006, Whitmeyer & Karlstrom 2007). The western margin of Laurentia experienced Neoproterozoic and Paleozoic passive margin sedimentation prior to disruption by the Antler orogeny during the Late Devonian-Mississippian and the Sonoman orogeny in the Triassic (e.g., Burchfiel et al. 1992, Dickinson 2004, Gehrels & Pecha 2014). Following mid-Paleozoic orogenesis the western Laurentian margin underwent a mix of Triassic continental and marine sedimentation and Lower to Middle Jurassic terrane accretion consolidating the North American margin by the Late Jurassic (Dickinson & Lawton 2001, DeCelles 2004).

## 2.2 North American Cordilleran deformation and crustal thickening

The onset of Farallon plate subduction along the western margin of North America occurred during the Late Jurassic (e.g., Burchfiel et al. 1992, Dickinson 2004, DeCelles 2004). Oblique, east-dipping subduction increased from 5 cm/yr in the Jurassic to ~10 cm/yr in the Cretaceous (SW-NE, e.g., Engebretson et al. 1985, Weil & Yonkee 2012). The rapidly converging Farallon plate generated the Sevier fold-thrust belt, a region of thin-skinned deformation in the western North American Paleozoic-early Mesozoic sedimentary sequences (e.g., DeCelles 2004, Weil & Yonkee 2012, Yonkee & Weil 2015, Saylor et al. 2020, Figure 2). In the late Cretaceous, the Farallon slab shallowed resulting in flat slab subduction from California inboard (e.g., Coney & Reynolds 1977, Henderson et al. 1984, Bird 1998, Saleeby 2003, Humphreys 2009, Liu et al. 2010, Bé). Slab flattening has been variably attributed to interactions with the cratonic lithospheric keel or shallowing due to subduction of a Shatsky Rise conjugate (Liu et al. 2010, Jones et al. 2011, Carrapa et al. 2019). Regardless of the precise mechanism, flat-slab subduction during the late Cretaceous-early Paleogene caused deformation to propagate inboard, generating basement-involved uplifts known as the Laramide uplifts, which were exhumed and eroded during the late Cretaceous and early Cenozoic (e.g., Dickinson & Snyder 1978, Schmidt & Garihan 1978, Brown 1988, Erslev 1993, Peyton et al. 2012, Carrapa et al. 2019). In addition to the Laramide orogeny, flat slab subduction significantly thickened the North American crust leading to regionally widespread metamorphism and magmatism (e.g., Terrien 2012, Fornash et al. 2013, Whitney et al. 2013, Behr & Smith 2016). Thus, the Wyoming craton remained at shallow crustal depths prior to being exposed via reverse faults, whereas the precondition for tectonic exhumation
115 of metamorphic core complexes was thick-crust and extensive mid-crustal magmatism and associated
116 metamorphism.

## 2.3 Slab rollback and magmatism

118 Foundering and subsequent rollback of the Farallon slab initiated in the late Paleogene (Bird 1979, Consten-
nius 1996, Smith et al. 2014, Cassel et al. 2018). Removal of the flat slab and the change from subduction
119 to transform plate boundaries between the Pacific and North American plate caused crustal collapse of
120 the thickened North American crust and induced regional-scale extension across North America exhum-
ing Metamorphic Core Complexes (MCC) along large-wavelength detachment faults (e.g., Atwater 1970,
121 Davis 1980, Dewey 1988, Lister & Davis 1989, Spencer & Reynolds 1990, Spencer et al. 2019). The thick-
122 ened crust and elevated geothermal gradient in the lower to middle crust during the Laramide orogeny set
123 the precondition for “flareup” magmatism, characterized by extensive extrusive silicic magmatism that
124 is suggested to have swept westward from the Eocene to Miocene (ca. 49-18 Ma) tracking the rollback
125 of the Farallon slab and associated asthenosphere mantle upwelling (e.g., Coney 1978, Constenius et al.
126 2003, Best et al. 2016).
127
128

# 3 Materials and methods

## 3.1 Sample sites

131 Five samples (KC82317-1, KC82317-3, KC82816-1, KC82816-2 and KC82816-3) were collected from
132 Archean to Paleoproterozoic rocks exposed along basement-cored uplifts in southwestern Montana (Table
133 1 and Figure 2, Carrapa et al. 2019). Zircon U-Pb dating on exposed Archean and Paleoproterozoic
134 crystalline basement constrain ages ranging from 2.5-1.6 Ga with a prominent distribution at ca. 1.7-1.8
135 Ga (Foster et al. 2006).

136 Two samples (KJJ09-7 and KJJ09-8) were collected from Laramide-aged plutons exposed in the Santa
137 Catalina mountains, Arizona (Table 1 and Figure 2). Samples KJJ09-7 and KJJ09-8 were identified as
138 Laramide intrusive rocks with Cretaceous zircon U-Pb ages of 69.1 and 73.0 Ma, respectively (Fornash
139 et al. 2013).

140 Three samples (MG1655, MG2465 and MG3150) were collected along a transect in the Pinaleño
141 Mountain core complex, Arizona (Table 1 and Figure 2). Samples MG2465 and MG3150 were collected
142 from an orthogneiss and a granodiorite, respectively, comprising the Proterozoic metamorphic core (Long
et al. 1995, Johnson & Arca 2010). Sample MG1655 was collected from a Proterozoic-mapped granite which hosts a strong mylonitized fabric (Johnson & Arca 2010). In addition, sample KJJ09-3 is a sill of Laramide-age which hosts a mylonitic fabric related to the Catalina detachment fault and a Paleogene zircon U-Pb age of 56.6 Ma (Fornash et al. 2013).

Samples GM-03, GM-05 and GM-08 were collected from the Galiuro Mountain, Arizona (Table 1). The Galiuro Mountains contain a section of late Oligocene volcanic rocks typical of the large ignimbrite flare-up of the southwestern US (Best et al. 2009); with predominantly dacitic ignimbrites reaching up to 2 km thick in thickness and are dated by zircon U-Pb as 29-24 Ma (Creasey & Krieger 1978). These ignimbrites overlay a Precambrian basement section dominated by mid Proterozoic metasedimentary rocks intruded by 1.1 Ga mafic dikes of suspected hot spot origin (Dickinson 1991, Arca et al. 2010).

<table><caption>Table 1: Samples collected from the North American cratonic basement, Laramide-aged plutons, Metamorphic Core Complexes/Detachment faults, and Paleogene volcanic rocks: Age is the published age of the rock, Lat is the north latitude and Long is the west longitude using the WSM 84 coordination system, Elev is elevation in meters above sea level. Lithology notation; MA is Middle Archean, LA is Late Archean, PP is Paleoproterozoic, K is Cretaceous, and $P_e$ is Paleogene. AFT is apatite fission-track, AUPb is apatite uranium-lead and trace and rare earth elements, and WR is whole rock geochemistry.</caption><thead><tr><td>Sample</td><td>Formation</td><td>Region</td><td>Lithology</td><td>Age</td><td>Lat</td><td>Long</td><td>Elev</td><td>Method</td></tr></thead></table>

### Basement-Cored Uplifts

<table><tr><td>KC82317-1</td><td>Cratonic<br>basement</td><td>Little Belt,<br>Montana</td><td>Gneissic<br>granite</td><td>PP</td><td>46.942</td><td>-110.745</td><td>1733</td><td>AFT/AUPb</td></tr><tr><td>KC82317-3</td><td>Intrusion</td><td>Little Belt,<br>Montana</td><td>Diorite</td><td>PP</td><td>46.958</td><td>-110.754</td><td>1742</td><td>AFT/AUPb</td></tr><tr><td>KC82816-1</td><td>Cratonic<br>basement</td><td>Gravelly Range,<br>Montana</td><td>Schist</td><td>LA</td><td>44.885</td><td>-111.694</td><td>2350</td><td>AFT/AUPb</td></tr><tr><td>KC82816-2</td><td>Cratonic<br>basement</td><td>Gravelly Range,<br>Montana</td><td>Gneiss</td><td>MA</td><td>44.888</td><td>-111.643</td><td>2120</td><td>AFT/AUPb</td></tr><tr><td>KC82816-3</td><td>Cratonic<br>basement</td><td>Gravelly Range,<br>Montana</td><td>Granite/<br>Gneiss</td><td>PP</td><td>45.318</td><td>-111.846</td><td>1929</td><td>AFT/AUPb</td></tr></table>

### Thickened Crust/Laramide Plutons

<table><tr><td>KJJ09-7</td><td>Leatherwood<br>Intrusion</td><td>Santa Catalina,<br>Arizona</td><td>Granodiorite</td><td>K</td><td>32.452</td><td>-110.752</td><td>2337</td><td>AFT/AUPb/<br>WR</td></tr><tr><td>KJJ09-8</td><td>Rice Creek<br>Porphyry</td><td>Santa Catalina,<br>Arizona</td><td>Hypabyssal<br>granite</td><td>K</td><td>32.479</td><td>-110.697</td><td>1424</td><td>AFT/AUPb</td></tr></table>

### Metamorphic Core Complex/Detachment Fault
<table><tr><td>MG1655</td><td>Core Complex Granite</td><td>Pinaleño Mts, Arizona</td><td>Mylonitized Granite</td><td>PP</td><td>32.745</td><td>-109.840</td><td>1655</td><td>AFT/AUPb/ WR</td></tr><tr><td>MG2465</td><td>Core Complex Granite</td><td>Pinaleño Mts, Arizona</td><td>Foliated Granite</td><td>PP</td><td>32.717</td><td>-109.855</td><td>2465</td><td>AFT/AUPb/ WR</td></tr><tr><td>MG3150</td><td>Core Complex Gneisses</td><td>Pinaleño Mts, Arizona</td><td>Granodiorite</td><td>PP</td><td>32.692</td><td>-109.864</td><td>3150</td><td>AFT/AUPb/ WR</td></tr><tr><td>KJJ09-3</td><td>Catnip Sill</td><td>Santa Catalina, Arizona</td><td>Peraluminous granite</td><td>P<sub>ε</sub></td><td>32.354</td><td>-110.723</td><td>1608</td><td>AFT/AUPb/ WR</td></tr></table>

## Extrusive Volcanic Rocks

<table><tr><td>GM-03</td><td>Galiuro Volcanics</td><td>Galiuro Mts, Arizona</td><td>Welded Tuff</td><td>P<sub>ε</sub></td><td>32.657</td><td>-110.357</td><td>1534</td><td>AFT/AUPb/ WR</td></tr><tr><td>GM-05</td><td>Galiuro Volcanics</td><td>Galiuro Mts, Arizona</td><td>Welded Tuff</td><td>P<sub>ε</sub></td><td>32.658</td><td>-110.316</td><td>1885</td><td>AFT/AUPb/ WR</td></tr><tr><td>GM-05</td><td>Galiuro Volcanics</td><td>Galiuro Mts, Arizona</td><td>Welded Tuff</td><td>P<sub>ε</sub></td><td>32.688</td><td>-110.441</td><td>1620</td><td>AFT/AUPb/ WR</td></tr></table>
<figure><img src="image_3.png" /></figure>

Figure 2: A schematic map of the western US Cordillera illustrating the key tectonic and magmatic processes: 1) Basement cored uplifts, 2) emplacement of Laramide-related plutons, 3) formation of meta-morphic core complex and associated detachment faults, and 4) emplacement of Paleogene extrusive volcanic rocks. Map modified from Dilek & Moores (1999).

## 3.2 Apatite fission-track thermochronology

Apatite fission-track (AFT) analyses were performed using the external detector method (Tagami 1987).

Apatite grains were mounted in epoxy and polished, and spontaneous fission tracks were revealed by etching with 5.5-M nitric acid for 20 s at 21°C before irradiation (after Donelick et al. 2005). The
neutron fluence was monitored using CN5 U-doped glass (Bellemans et al., 1995). The irradiation was
performed at Oregon State University. After irradiation, mica external detectors were etched in 40%
hydrofluoric acid for 45 min at 21°C (after Donelick et al. 2005). Analyses were conducted for optical
identification of fission-tracks using an Olympus microscope at 1600X magnification with a drawing tube
located above a digitizing tablet and a Kinetek computer-controlled stage driven by the FT Stage program
provided by Trevor Dumitru of Stanford University. The fission-track analyses were performed at the
Arizona Fission Track Laboratory in the University of Arizona (Table 3 and 4, and Supplementary Table
2). Confined tracks were measured to provide information on the time spent in the 120-60°C apatite
partial annealing zone (APAZ), with longer mean confined track lengths (>13.5µm) defining rapid cooling
through the APAZ and shorter mean confined track lengths indicating prolonged residence in the APAZ
(e.g., Laslett et al. 1982, Gleadow et al. 1986, Donelick & Miller 1991, Tagami & O'Sullivan 2005).

Samples KC82317-1, KC82317-3, KC82816-1, KC82816-2 and KC82816-3 were selected for apatite
U-Pb and FT single-grain double dating. The concentration of uranium (²³⁸U) in the counted areas
of apatite was determined using LA-ICP-MS. Age calculation was carried out using in-house R script
following equations as described in Hasebe et al. (2004) and Vermeesch (2017), using the Durango apatite
(McDowell et al. 2005) to perform a session-zeta calibration (Vermeesch 2017). For details on trace
element acquisition, see section 3.4.

## 3.3 Apatite U-Pb thermochronology

The AUPb method relies on the thermally activated volume diffusion of Pb within the crystal lattice of an
apatite grain to provide information about its thermal history cochrane2014high, seymour2016tectonics, glorie2017the
Estimates of the closure temperature of the AUPb system are between *ca.* 550 and 370°C depending on
cooling rates and apatite crystal sizes (Chew & Donelick 2012, Thomson et al. 2012, Cochrane et al.
2014, Gillespie et al. 2018), although temperatures of *ca.* 450-550°C are more typically reported (e.g.,
Schoene & Bowring 2007, Blackburn et al. 2011). Apatite grains that were selected for AFT analysis were
also targeted for same-grain AUPb and *in situ* trace element analyses. The AUPb dates for this study
were acquired using a Photon Machine Analyte G2 Excimer 193 nm laser ablation system connected to
a Thermo Element2 single-collector High Resolution ICP-MS at the Arizona LaserChron Center. Spot
ablations were performed using a 30 µm spot size and a 5 Hz laser repetition rate. Each analysis com-
prised 20s background collection (laser off) and 40.2 s apatite ablation (laser on; see detailed operational
conditions in Table 2). During apatite U-Pb analysis, the Madagascar (MAD2) apatite primary reference
material (reference age 474.2±0.4 Ma, Thomson et al. 2012) was measured repeatedly throughout the
session to correct for instrumental drift and down hole fractionation, and the Mt. McClure (523.98 ±
0.12 Ma Schoene & Bowring 2006) reference material was analyzed as a secondary standard for accuracy
checks (Mt. McClure = 514.6 ± 2.0 Ma, Supplementary File 5). Data reduction was performed using
the "VizualAge_UcomPbine" Data Reduction Scheme (DRS) in Iolite (Paton et al. 2011, Chew et al.
2014). More details on this DRS can be found in Chew et al. (2014). Isotopic data (238/206 and 207/206
ratios) were plotted and apatite U-Pb ages were obtained using a Tera-Wasserburg lower concordia inter-
cept based on a unique upper (²⁰⁷Pb/²⁰⁶Pb) of a relevant non-radiogenic Pb regression line Figures 3-6.
Tera-Wasserburg plots were calculated using the IsoplotR package v.2.3 (Vermeesch 2018) in R v.3.5.2.

3.4 Apatite trace element analysis

A range of trace elements were measured simultaneously to the acquisition of U and Pb isotopes through
LA-ICPMS (Table 2 Supplementary File 4, Chapman et al. (2016)). These elements were selected based
on their potential to inform on source rock processes (e.g. Belousova et al. 2002b, Chapman et al. 2016,
Bruand et al. 2017, Henrichs et al. 2018, Gillespie et al. 2018). Trace element data reduction were per-
formed using the 'X_Trace_Elements_IS' DRS in Iolite (Paton et al. 2011) following Chew et al. (2016).
Instrumental drift was corrected using NIST610 as the primary standard and elemental concentrations
were calculated using ⁴³Ca for internal standardization using stoichiometric abundance of Ca at 39.74
wt% (Chew et al. 2014). Individual element concentrations for standards NIST610 glass, MAD2 apatite,
Durango apatite, and McClure Mountain apatite can be found in Supplementary File 3.
<table><caption>Table 2: Analytical parameters for apatite U-Pb and trace and rare earth element (REE) analyses using Laser Ablation Inductively Coupled Plasma Mass Spectrometry (LA-ICP-MS). MFC is mass flow controller for He gas. Acquisition time and settling time are listed in order of isotopes measured. For a detailed list of LA-ICP-MS parameters, see Supplementary File 4.</caption><thead><tr><th>Parameters</th><th>Apatite U-Pb and trace and REE</th></tr></thead><tbody><tr><td colspan="2"><em>Laser Parameters</em></td></tr><tr><td><strong>Instrument</strong></td><td>Photon Machines Analyte G2 excimer laser 193 nm</td></tr><tr><td><strong>Washout</strong></td><td>20s</td></tr><tr><td><strong>Background</strong></td><td>9s</td></tr><tr><td><strong>Analysis Duration</strong></td><td>40.2s</td></tr><tr><td><strong>Laser Repetition</strong></td><td>5 Hz</td></tr><tr><td><strong>Spot Size</strong></td><td>30μm</td></tr><tr><td><strong>Energy</strong></td><td>3.9 J/cm²</td></tr><tr><td><strong>MFC1, MFC2<br>(L/min)</strong></td><td>0.11, 0.29</td></tr><tr><td colspan="2"><em>Mass Spectrometer</em></td></tr><tr><td><strong>Instrument</strong></td><td>Thermo Element 2</td></tr><tr><td><strong>Number of Isotopes<br>measured</strong></td><td><sup>29</sup>Si, <sup>35</sup>Cl, <sup>43</sup>Ca, <sup>51</sup>V, <sup>55</sup>Mn, <sup>88</sup>Sr, <sup>89</sup>Y, <sup>91</sup>Zr, <sup>139</sup>La, <sup>140</sup>Ce, <sup>141</sup>Pr, <sup>146</sup>Nd, <sup>147</sup>Sm, <sup>153</sup>Eu, <sup>157</sup>Gd, <sup>159</sup>Tb, <sup>163</sup>Dy, <sup>165</sup>Ho, <sup>166</sup>Er, <sup>169</sup>Tm, <sup>172</sup>Yb, <sup>175</sup>Lu, <sup>202</sup>Hg, <sup>204</sup>Pb, <sup>206</sup>Pb, <sup>207</sup>Pb, <sup>208</sup>Pb, <sup>232</sup>Th, <sup>238</sup>U</td></tr></tbody></table>

## 3.5 Whole rock geochemistry

Samples were cut and finely molded in agate jars. Approximately 50 mg of each rock powder was dissolved by two successive acid attacks. Samples were firstly dissolved in a HF:HNO₃ concentrate mixture (2:1) in 6 mL Savillex® standard octagonal bodied PFA Teflon vials. These Teflon vials, generally referred as Teflon bombs, were sealed using a wrench top closure and Ultem® socket cap. Teflon vials were then set into an oven at 130°C for 5 days (see similar approach described in Inglis et al. 2018). Solutions were then evaporated on hot plates, and residues were re-dissolved in a HCl concentrate (16M) solution and put on hot plate (120°C) for 24h. The solution was then re-evaporated, and residue dissolved in 2% nitric acid for chemical analysis. Samples were centrifuged to remove undissolved material from the solution. Samples run for trace element analysis were diluted to 10 mL with 2% nitric acid. Trace element analyses were conducted in solution on a ThermoFisher X-Series II Quadrupole Inductively Coupled Plasma Mass Spectrometer (Q-ICP-MS) at the Geosciences Department at the University of Arizona, following the procedures outlined in Rossel et al. (2013). Trace element results were standardized to the Columbia River Basalt (BCR-2) USGS rock standard (Schudel et al. 2015). Standard error for samples run in solution is expected to be less than 5% for major elements and less than 3% for REEs. A blank composed of the same 2% nitric acid that was used to dilute the samples and standards was run before and after each
sample and standard to account for any material added by our matrix. Detailed whole rock geochemistry
results can be found in Supplementary File 6.

# 4 Results

## 4.1 Basement-cored uplifts

### 4.1.1 Apatite U-Pb dating

All five samples from basement-cored uplifts clearly define linear arrays in Tera-Wasserburg concordia space (Figure 3A, D, G, J, and K). Using the lower intercept discordia, samples KC82317-1, KC82317-2, KC82816-2, and KC82816-3 yielded Paleoproterozoic ages of 1840 ± 14 Ma (MSWD = 0.42), 1858 ± 13 Ma (MSWD = 0.42), 1811 ± 31 Ma (MSWD = 8.0), and 1675 ± 17 Ma (MSWD = 0.7), respectively. Sample KC82816-1 yielded an Archean age of 2417 ± 32 Ma (MSWD = 3.3, Figure 3 and Supplementary Table 1).

### 4.1.2 Apatite fission-track thermochronology

Two samples (KC82317-1 and KC82317-3) from the Paleoproterozoic core of the Little Belt Laramide uplift in central Montana yielded Cretaceous AFT ages of 85 ± 4 Ma and 64 ± 3 Ma, respectively (Table 1 and Table 3). Confined fission-tracks from sample KC82317-3 provided a mean track length (MTL) of 14.2 ± 1.0 µm.

Three samples (KC82816-1, KC82816-2 and KC82816-3) from Archean and Paleoproterozoic rocks exposed along the Gravelly Range Laramide uplift in southwest Montana yielded Cretaceous AFT ages of 71 ± 7 Ma, 95 ± 12 Ma, and 74 ± 4 Ma, respectively (Table 1 and Table 3). Confined fission-tracks from samples KC82816-2 and KC82816-3 yielded relatively long MTLs of 14.2 ± 1.2 µm and 14.0 ± 0.9 µm, respectively.

<table><caption>Table 3: Summary of laser ablation apatite-fission track data from this: n = number of grains dated, Ns = number of spontaneous tracks counted, ρs = average spontaneous track density (x10⁵cm⁻²), ²³⁸U = average uranium concentration of the analyzed grains, P(χ²) = χ² probability after (Galbraith 1981), t is the central age (Ma), Dpar = mean track etch pit diameter parallel to the crystallographic c-axis, P = pooled age, # is the number of confined tracks counted, MTL is the mean of measured confined tracks in µm and ±1σ is also in µm.</caption><thead><tr><th>Sample</th><th>n</th><th>Ns</th><th>ρs</th><th>²³⁸U<br>(ppm)</th><th>P(χ²)</th><th>t<br>(Ma)</th><th>±1σ</th><th>Dpar<br>(μm)</th><th>P<br>(Ma)</th><th>±1σ</th><th>#</th><th>MTL<br>(μm)</th><th>±1σ</th></tr></thead><tbody></tbody></table>
Basement-Cored Uplift

<table><tr><td>KC82317-1</td><td>23</td><td>1183</td><td>16.4</td><td>44.2</td><td>0.17</td><td>64</td><td>3</td><td>2.2</td><td>52</td><td>8</td><td>-</td><td>-</td><td>-</td></tr><tr><td>KC82317-3</td><td>26</td><td>1391</td><td>27.7</td><td>59.9</td><td>0.15</td><td>85</td><td>4</td><td>2.5</td><td>85</td><td>14</td><td>50</td><td>14.2</td><td>1.0</td></tr><tr><td>KC82816-1</td><td>10</td><td>239</td><td>12.4</td><td>33.5</td><td>0.12</td><td>71</td><td>7</td><td>2.0</td><td>92</td><td>16</td><td>-</td><td>-</td><td>-</td></tr><tr><td>KC82816-2</td><td>11</td><td>148</td><td>5.0</td><td>10.2</td><td>0.06</td><td>95</td><td>12</td><td>2.2</td><td>129</td><td>24</td><td>28</td><td>14.2</td><td>1.2</td></tr><tr><td>KC82816-3</td><td>25</td><td>839</td><td>10.7</td><td>26.7</td><td>0.76</td><td>75</td><td>4</td><td>2.3</td><td>73</td><td>12</td><td>50</td><td>14.0</td><td>0.9</td></tr></table>

244

### 4.1.3 Apatite trace elements

Samples KC82317-1 and KC82317-3 both display REE patterns enriched in light REE (LREE) and depleted in heavy REE (HREE), as well as high mean values of La/Lu$_N$ (N is for chondrite normalized) of 86.0 and 176.2 for KC82317-1 and KC82317-3, respectively (Figure 3. Sample KC82317-1 shows more variation in its LREE content (Figure 3 and 8); La/Sm$_N$ ranging from 0.4 to 2.4 (median 1.3); as well as a distinctive negative Eu anomaly ($Eu/Eu^* = Eu_N/[Sm_N * Nd_N]^{1/2}$) ranging from 0.13 to 0.80 (median 0.27, compared to 0.78 for sample KC82317-3). Samples KC82816-1, KC82816-2, and KC82816-3 display various REE patterns (Figure 3). KC82816-1 shows a relatively flat REE pattern (La/Lu$_N$ ranging from 0.9 to 18.2) with consistent concentrations of REE, 100 to more than 1000x chondritic values, along with a moderate negative Eu anomaly (Eu/Eu* from 0.49 to 0.63). Sample KC82816-2 shows convex-upward REE patterns (Figure 3), with depletion in L- and H-REE compared to Medium-REE (MREE; La/Sm$_N$: 0.05-0.35 and Dy/Yb$_N$: 2.21-3.94). Negative Eu anomaly ranging between 0.15 to 0.31 (median of 0.18). Sample KC82816-3 shows a similar pattern as KC82816-1 with a moderate Eu anomaly (Eu/Eu* median value: 0.48; Figure 3). However, it shows a pronounced depletion in HREE compared to MREE (Dy/Yb$_N$ ranging from 1.20 to 2.97, median 1.54), similar to KC82816-2.

246

247

248

249

250

251

252

253

254

255

256

257

258

259

## 4.2 Laramide-aged plutons

260

### 4.2.1 Apatite U-Pb dating

261

262

263

264

Both samples from Laramide plutons clearly show linear arrays in Tera-Wasserburg concordia space (Figure 4A and D). Samples KJJ09-07 and KJJ09-08 yield lower intercept discordia ages of 40 ± 2 Ma (MSWD = 1.3), and 62 ± 21 Ma (MSWD = 2.3), respectively (Figure 4 and Supplementary Table 1).
<figure><img src="image_4.png" /></figure>

Figure 3: Tera-Wasserburg (T-W) concordia diagrams (A, D, G, J, and M), chondrite normalized REE profile plots (B, E, H, K, and N), and apatite fission-track radial plots (C, F, I, L, and O) for samples KC82317-1, KC82317-2, KC82816-1, KC82816-2, and KC82816-3. Apatite U-Pb analyses are plotted using IsoplotR (Vermeesch 2018) and colored according to their La/Lu<sub>N</sub> ratio. On the T-W, Ellipses are plotted at 2σ, solid lines represent discordia projections used to obtain an estimate of the initial <sup>207</sup>Pb/<sup>206</sup>Pb ratio. Chondrite values from Pourmand et al. (2012) and analyses are colored according to their <sup>207</sup>Pb corrected apatite U-Pb age based on the sample's unique discordia. Circles in the radial plots (Vermeesch 2009) are colored according to their La/Lu<sub>N</sub> ratio, with blank circles representing grains that no geochemical data was obtained.
<figure><img src="image_5.png" /><figcaption>Figure 3: Continued.</figcaption></figure>

### 4.2.2 Apatite fission-track thermochronology

Two samples (KJJ09-7, and KJJ09-8) from Laramide-aged plutons within the Santa Catalina Meta-morphic Core Complex, Arizona yielded Paleogene-Neogene AFT ages of 24 ± 2 Ma and 41 ± 6 Ma, respectively (Table 1 and Table 4). Confined fission-tracks from sample KJJ09-07 yielded a long MTL of 13.7 ± 1.2 µm.

### 4.2.3 Apatite and whole rock trace elements

Samples KJJ09-07 and KJJ09-08 display similar REE patterns, with MREE content ~200x chondritic values. Apatites from both samples display positive LREE/HREE ratios (La/Lu<sub>N</sub> ranging between 18.8-280.4 and 160.3-307.6, respectively) and moderately negative Eu anomaly (Eu/Eu* median values of 0.54 and 0.67, respectively). Whole rock REE for sample KJJ09-07 displays a REE profile enriched in LREE and depleted in HREE (La/Lu<sub>N</sub> of 11.84) similar to *in situ* chemical analyses on constitutive apatites (Figure 3).
<figure><img src="image_6.png" /><figcaption>Figure 4: Tera-Wasserburg (T-W) concordia diagrams (A, D, and G), chondrite normalized REE spider-plots (B, E, and H), and apatite fission-track radial-plots (C, F, and I) for samples KJJ09-03, KJJ09-07, and KJJ09-08. Apatite U-Pb analyses are plotted using IsoplotR (Vermeesch 2018) and colored according to their chondrite normalized La/Lu ratio. On the T-W, Ellipses are plotted at $2\sigma$, solid lines represent discordia projections used to obtain an estimate of the initial $^{207}\text{Pb}/^{206}\text{Pb}$ ratio. Chondrite values from Pourmand et al. (2012) and analyses are colored according to their $^{207}\text{Pb}$ corrected apatite U-Pb age based on the sample's unique discordia. Solid black lines in the spider-plots are whole rock REE patterns. Circles in the radial plots (Vermeesch 2009) are colored according to their La/LuN ratio, with blank circles representing grains that no geochemical data was obtained.</figcaption></figure>
## 4.3 Metamorphic Core Complex and detachment fault

### 4.3.1 Apatite U-Pb dating

Of the three MCC and detachment fault samples, two samples, MG3150 and MG2465 displayed scattered ellipses through Tera-Wasserburg concordia space (Figure 5A, D, and G). Sample MG3150 displays oscillatory zoning in cathodoluminescence (CL) images, whereas sample MG2465 preserves recrystallization textures (Supplementary File 8). Thus, we define two and one clear discordia arrays, respectively. Sample MG3150 yields a prominent Mesoproterozoic lower intercept discordia age (n = 17) of 1087 ± 13 Ma (MSWD = 2.0), with a secondary Devonian discordia age (n = 9) of 373 ± 18 Ma (MSWD = 3.9, Figure 5). Additionally, sample MG3150 preserves two ellipses that yield a ca. 700 Ma age, however two ellipses were deemed too few for a robust discordia age calculation (Supplementary Table 1). Sample MG2465 yields a prominent Neoproterozoic lower intercept discordia (n = 14) of 734 ± 23 Ma (MSWD = 1.6) with younger ellipses indicative subsequent of recrystallization. The two detachment faults MG1655 and KJJ09-03 both preserve homogeneous CL textures, a clear, linear array in Tera-Wasserburg concordia space and yield a lower intercept discordia ages of 22 ± 5 (MSWD = 0.5) and 39 ± 7 Ma (MSWD = 2.9), respectively (Figure 5G and J).

### 4.3.2 Apatite fission-track

Three samples (MG1655, MG2465, and MG3150) from Paleoproterozoic rocks exposed within the Pinaleño Mountain core complex, Arizona yielded Paleogene-Neogene AFT ages of 20 ± 2 Ma, 25 ± 3 Ma, and 29 ± 3 Ma, respectively (Table 1 and Table 4). Confined fission-tracks from samples GM3150 and MG2465 yielded MTLs of 13.8 ± 1.6 µm and 13.0 ± 1.7 µm, respectively. Sample KJJ09-3 from the detachment mylonitized granite along the Santa Catalina Metamorphic Core Complex, Arizona yielded Paleogene-Neogene AFT age of 21 ± 3 Ma (Table 4).

### 4.3.3 Apatite and whole rock trace elements

Sample MG3150 displays a variably fractionated REE profile (Figure 5), marked by minor to moderate LREE/HREE ratios (La/LuN ranging from 5.88 to 341.84, median of 94.84); and a minor depletion in Eu (Eu/Eu* median at 0.80). Fractionation of HREE compared to MREE is distinct in this sample (Dy/YbN range: 0.65-2.37). Sample MG2465 displays relatively flat REE patterns (in particular for HREE), with minor depletion in LREE (La/SmN ranging from 0.15 to 1.90, median of 1.12) and a noticeable depletion
<figure><img src="image_7.png" /></figure>

Figure 5: Tera-Wasserburg (T-W) concordia diagrams (A, D, G, and J), chondrite normalized REE spiderplots (B, E, H, and K), and apatite fission-track radialplots (C, F, I, and L) for samples MG3150, MG2465, MG1655, and KJJ09-03. Apatite U-Pb analyses are plotted using IsoplotR (Vermeesch 2018) and colored according to their La/Lu ratio. On the T-W, Ellipses are plotted at 2σ, solid lines represent discordia projections used to obtain an estimate of the initial $^{207}$Pb/$^{206}$Pb ratio. Dashed lines represent a more uncertain discordia projection used to obtain an estimate of the initial $^{207}$Pb/$^{206}$Pb ratio, dashed ellipses represent grains that were corrected using the more uncertain discordia. Chondrite values from Pourmand et al. (2012) and analyses are colored according to their $^{207}$Pb corrected apatite U-Pb age based on the sample's unique discordia, mixed ages were calculated using an assumed initial $^{207}$Pb/$^{206}$Pb ratio based on (Stacey & Kramers 1975). Solid black lines in the spiderplots are whole rock REE patterns. Circles in the radial plots (Vermeesch 2009) are colored according to their La/Lu<sub>N</sub> ratio, with blank circles representing grains that no geochemical data was obtained.
<figure><img src="image_8.png" /><figcaption>Figure 5: Continued.</figcaption></figure>

in Eu (Eu/Eu* median value of 0.25). Whole rock REE composition for sample MG2465 mimics the REE pattern of constitutive apatites, with a flat trend and a distinctive Eu anomaly (Eu/Eu* of 0.43, Figure 5). In contrast, sample MG1655 shows variable fractionation trend of LREE in comparison to MREE (La/Sm<sub>N</sub> from 0.10 to 1.35, median of 0.25) and relatively flat HREE distribution (Dy/Yb<sub>N</sub> median value of 1.04; Figure 7). Eu anomaly is negative and moderate ranging from 0.10 to 0.69. In comparison, whole rock REE composition for sample MG1655 shows a similar trend (M- and HREE and Eu anomaly) except for LREE that appears enriched for the bulk rock composition (La/Sm<sub>N</sub> of 3.11). All three samples display a large degree of variance between individual grain REE concentrations (Figure 5 and Figure 7).

Sample KJJ09-03 displays a flat to slightly convex-upward REE pattern (Figure 5K), with a moderate depletion in LREE (La/Sm<sub>N</sub> range: 0.16-0.67) and in HREE (Dy/Yb<sub>N</sub> range: 0.85-2.07, median of 2.04). Patterns are marked by a minor negative Eu anomaly (Eu/Eu* median at 0.36). In contrast, the whole rock REE composition for sample KJJ09-03 shows a high LREE/HREE ratio (La/Lu<sub>N</sub> of 37.85) and no Eu anomaly (Eu/Eu* of 1.30).
# 4.4 Volcanic rocks

## 4.4.1 Apatite U-Pb dating

All three samples from the volcanic rocks clearly define linear arrays in Tera-Wasserburg concordia space (Figure 6A, D, and G). Samples GM-03, GM-05, and GM-08 yield Eo-Oligocene lower intercept discordia ages of 41 ± 12 Ma (MSWD = 1.6), 28 ± 31 Ma (MSWD = 1.2), and 29 ± 12 Ma (MSWD = 1.1), respectively (Figure 6 and Supplementary Table 1).

## 4.4.2 Apatite fission-track thermochronology

Three samples (GM-03, GM-05, and GM-08) from Paleogene volcanic rocks along the Galiuro Mountains in Arizona yielded AFT ages of 31 ± 4 Ma, 31 ± 6 Ma, and 29 ± 5 Ma, respectively (Table 1 and Table 4).

<table><caption>Table 4: Summary of external detector apatite fission track data from the North American Cordillera: n = number of grains dated, Ns = number of spontaneous tracks counted, ρs = average spontaneous track density (x10⁵cm⁻²), Ni = number of induced tracks counted, ρi = average induced track density (x10⁵cm⁻²), P(χ²) = χ² probability after (Galbraith 1981), Nd = number of dosimeter tracks counted, ρi = average dosimeter track density (x10⁵cm⁻²), t is the central age (Ma), Dpar = mean track etch pit diameter parallel to the crystallographic c-axis, # is the number of confined tracks counted, MTL is the mean of measured confined tracks in µm and ±1σ is also in µm. G. Jepson Zeta: 341.7 ± 8.</caption><thead><tr><th>Sample</th><th>n</th><th>Ns</th><th>ρs</th><th>Ni</th><th>ρi</th><th>P(χ²)</th><th>t (Ma)</th><th>±1σ</th><th>ρd</th><th>Nd</th><th>Dpar<br>(μm)</th><th>#</th><th>MTL<br>(μm)</th><th>±1σ</th></tr></thead></table>

### Laramide Plutons

<table><tr><td>KJJ 09-7</td><td>20</td><td>144</td><td>2.1</td><td>1249</td><td>18.1</td><td>0.50</td><td>24.1</td><td>2.24</td><td>12.3</td><td>5533</td><td>2.0</td><td>50</td><td>13.7</td><td>1.2</td></tr><tr><td>KJJ 09-8</td><td>20</td><td>79</td><td>1.1</td><td>364</td><td>5.2</td><td>0.66</td><td>40.6</td><td>5.66</td><td>12.1</td><td>5533</td><td>2.0</td><td>-</td><td>-</td><td>-</td></tr></table>

### Metamorphic Core Complex

<table><tr><td>MG1655</td><td>20</td><td>78</td><td>1.5</td><td>939</td><td>18.3</td><td>0.54</td><td>19.6</td><td>2.38</td><td>13.8</td><td>5533</td><td>2.1</td><td>-</td><td>-</td><td>-</td></tr><tr><td>MG2465</td><td>20</td><td>106</td><td>1.4</td><td>970</td><td>13.0</td><td>0.60</td><td>25.0</td><td>2.65</td><td>13.4</td><td>5533</td><td>2.4</td><td>14</td><td>13.0</td><td>1.7</td></tr><tr><td>MG3150</td><td>20</td><td>156</td><td>3.0</td><td>1193</td><td>23.2</td><td>0.25</td><td>29.2</td><td>2.97</td><td>13.0</td><td>5533</td><td>2.0</td><td>38</td><td>13.8</td><td>1.6</td></tr><tr><td>KJJ 09-3</td><td>20</td><td>80</td><td>1.0</td><td>770</td><td>10.0</td><td>0.47</td><td>21.4</td><td>2.58</td><td>12.1</td><td>5533</td><td>2.0</td><td>-</td><td>-</td><td>-</td></tr></table>

### Extrusive Volcanic Rocks

<table><tr><td>GM-03</td><td>20</td><td>77</td><td>1.2</td><td>697</td><td>11.2</td><td>0.84</td><td>30.8</td><td>3.8</td><td>16.4</td><td>5533</td><td>2.2</td><td>-</td><td>-</td><td>-</td></tr><tr><td>GM-05</td><td>20</td><td>32</td><td>1.0</td><td>287</td><td>8.9</td><td>0.96</td><td>30.5</td><td>5.7</td><td>16.0</td><td>5533</td><td>2.4</td><td>-</td><td>-</td><td>-</td></tr><tr><td>GM-08</td><td>20</td><td>38</td><td>1.0</td><td>345</td><td>9.4</td><td>0.92</td><td>28.6</td><td>5.0</td><td>15.2</td><td>5533</td><td>2.2</td><td>-</td><td>-</td><td>-</td></tr></table>
<figure><img src="image_9.png" /></figure>

Figure 6: Tera-Wasserburg (T-W) concordia diagrams (A, D, and G), chondrite normalized REE spider-plots (B, E, and H), and apatite fission-track radial-plots (C, F, and I) for samples GM-03, GM-05, and GM-08. Apatite U-Pb analyses are plotted using IsoplotR (Vermeesch 2018) and colored according to their La/Lu ratio. On the T-W, ellipses are plotted at $2\sigma$, solid lines represent discordia projections used to obtain an estimate of the initial $^{207}\text{Pb}/^{206}\text{Pb}$ ratio. Chondrite values from Pourmand et al. (2012) and analyses are colored according to their $^{207}\text{Pb}$ corrected apatite U-Pb age based on the sample's unique discordia. Solid black lines in the spider-plots are whole rock REE patterns. Circles in the radial plots (Vermeesch 2009) are colored according to their La/Lu<sub>N</sub> ratio, with blank circles representing grains that no geochemical data was obtained.
### 4.4.3 Apatite and whole rock trace element results

All three volcanic samples (GM-03, GM-05, and GM-08) display comparable REE patterns with MREE showing 200 to 2000 the chondritic values (Figure 7). All three samples show an enrichment in LREE compared to MREE (La/Sm<sub>N</sub> median values at 10.71, 2.78 and 2.38, respectively) and a moderate negative Eu anomaly (Eu/Eu* median values at 0.55, 0.40 and 0.48, respectively). Sample GM-03 shows a relatively flat HREE trend (Dy/Yb<sub>N</sub> ranging from 1.00 to 1.58) compared to samples GM-05 and GM-08 that are marked by more fractionated trends, with Dy/Yb<sub>N</sub> ranges of 2.21-3.11 and 1.71-3.23, respectively.

The whole rock REE profiles for each sample mirrors the apatite REE profile with an enrichment in LREE (GM-03, GM-05, and GM-08 La/Sm<sub>N</sub> values of 4.57, 4.11, and 3.58, respectively) and depletion in HREE (GM-03, GM-05, and GM-08 Dy/Yb<sub>N</sub> values of 1.23, 1.30, and 1.30, respectively), but with a more modest depletion in Eu (GM-03, GM-05, and GM-08 Eu/Eu* values of 0.77, 0.73, and 0.74, respectively, Figure 6-7). All three samples show very little variation between individual grain REE concentrations (Figure 6-7).

## 5 Discussion

The combination of apatite U-Pb dating, fission-track thermochronology, and apatite and whole rock trace element analysis, applied to well constrained tectonic environments across the North American Cordillera make it possible to distinguish thermal pathways and trace element patterns indicative of various crustal histories. Such discriminant chemical ratios for primary (geochemically representative of whole rock) *versus* secondary apatite (geochemically altered) are crucial for single-grain LA-ICP-MS U-Pb and FT analysis, as it is important to minimize the amount of material ablated as collecting the full range of REEs is material intensive. Finally, our results encourage the usage of apatite geochemistry to elucidate middle to upper crustal thermal histories across a range of cordilleran tectonic processes.

### 5.1 Primary *versus* secondary thermal history

By using apatite trace element chemistry it is possible to determine the nature of its host rock and to discuss their potential chemical reequilibration (e.g. Gillespie et al. 2018, O'Sullivan et al. 2020). Through trace element and REE composition; secondary (metamorphic or metasomatic) apatites can
be distinguished from their primary (magmatic) counterparts based on their depletion in LREE and/or HREE. These depleted trends are generally attributed to the growth of cogenetic mineral phases, with epidote and monazite sequestering LREE, and garnet sequestering HREE (Janots et al. 2008, El Korh et al. 2009, Henrichs et al. 2018, Glorie et al. 2019, Henrichs et al. 2019). Such metamorphic apatite composition is illustrated in samples KC82816-1, KC82816-2, and KC82816-3, marked by a larger spread of their ΣREE content but above all, showing concave downward REE pattern with unaffected MREE content, a large depletion of LREE and variable depletion of HREE (Figure 3). The apatite host-rock samples consist of an intermediate to granitic orthogneisses and a micaschist composition forming the cratonic basement which underwent medium- to high-grade metamorphic conditions (2-10 kbar and 400-800°C, Table 1 and Figure 3, e.g. Harms et al. 2004, Mueller et al. 2005). Additionally, apatites extracted from mylonitized granitic samples KJJ09-03 and MG1655 both display similar REE patterns, marked by moderate to major depletion in LREE. Depleting L- or H-REE alters the sum of an apatite's REE content, which is used as a sensitive proxy for distinction between igneous and metamorphic apatite; with the sum of REE in apatite decreasing with increasing metamorphic gradient (Figure 7A, e.g. Belousova et al. 2002a, El Korh et al. 2009, Henrichs et al. 2018).

Low LREE/MREE ratios (samples MG1655 and KJJ09-03) are interpreted as subsequent chemical re-equilibration during deformation and recrystallization at medium temperature conditions (~400-300°C); likely alongside the formation of LREE rich phases (Table 1 and Figure 5). More specifically, apatite recrystallization occurs simulatenously with monazite (LREE- and Th-rich) and/or epidote group minerals (LREE-rich and Sr-rich) in medium- to low-grade metamorphic felsic and mafic rocks (e.g. Grapes & Hoskin 2004, Glorie et al. 2019). These two reactions are a possible controls on the apatite geochemistry, with syn-crystallization of monazite influencing sample KC082816-2. Whereas epidote growth controls the apatite geochemistry observed in the mylonitic samples (MG1655 and KJJ09-03, Figure 7B). The Sr/Y ratio in apatite is variably used to distinguish igneous and metamorphic host rock or interpreted to indicate of increased fluid involvement (Nishizawa et al. 2005, Prowatke & Klemme 2006, Henrichs et al. 2018, Odlum & Stockham 2019). Igneous apatites show strong negative correlation between Sr/Y and ΣREE due to the similar compatibility of REE and Y in apatite (Figure 7A) as well as lower ΣREE and Sr/Y ratio values. These trends agree with the co-recrystallization of Ca-rich (and thus Sr-rich) phases like epidote during metamorphism.

The Th/U ratio in zircon is a reliable proxy to distinguish metamorphic and magmatic zircon (e.g. Rubatto 2002) and here we apply it to apatite. Cordilleran apatites from metamorphic samples displaying a decreasing trend in Th/U coeval with decreasing La/SmN (Figure 7B) attributed to the co-recrystallization of
Th-rich phases like monazite. Uranium depletion has been observed in many metamorphic apatites (O'Sullivan et al. 2018, Henrichs et al. 2018, 2019), however, apatites in these studies were sourced from coarse-grained rocks. Large diffusion domains in coarse grained rocks can obscure precise metamorphic processes. There was no noticeable difference between U concentrations of metamorphic (e.g. sample KC82816-1 yielded 3-66 ppm) and igneous apatite (e.g. sample GM-03 yielded 3-15 ppm, Supplementary Table 1) in this study. Rather, U abundances are thought to be controlled by the nature and composition of the host whole rock U concentration. As this study largely focuses on felsic crustal rocks, little variation between samples is expected (O'Sullivan et al. 2020).

The apatites extracted from mylonitized granitic samples (KJJ09-03 and MG1655) show a notable enrichment in HREE uncorrelated to the whole rock composition (Figure 5J). A possible source of these HREE could be the retrogression of garnet, destabilized during low-grade metamorphism (Rubatto 2002, Rubatto & Hermann 2007, Odlum & Stockli 2020). Here we suggest that the retrogression of garnet during low-grade metamorphism is losing HREE for which apatite is providing a sink whilst undergoing diffusion between ~450-375°C (Figure 7A and Supplementary File 8).

Several authors use the Eu anomaly as a proxy to track oxygen fugacity in igneous systems (e.g. Bau 1991, Trail et al. 2012). In our study, the variability in Eu anomaly is likely related to preexisting Eu anomaly in the whole rock composition (e.g. Figure 5E and H). This is supported by single sample inter-grain variability in the Eu anomaly, Sr/Y, and sumREE for either magmatic or metamorphic host rocks (Figure 7A and C). Thus, the Eu anomaly for a given apatite is likely inherited from the source composition of the parental magma of its respective rock and associated pre- or cogenetic crystallization of plagioclase during apatite crystallization (Figure 7D, e.g. Lu et al. 2016, O'Sullivan et al. 2020). This, further emphasizes the applicability of Eu anomaly in individual apatite grains to decipher its host rock composition, degree of crustal reworking, and depth of magmatic source (e.g. Henrichs et al. 2018, Nathwani et al. 2020).

Samples MG3150 and MG2465 both display low LREE/HREE ratios and flat REE patterns that are typical of metamorphic apatite, despite both rocks coming from weakly metamorphosed and non-metamorphosed rocks, an orthogneiss and granodiorite, respectively (Figure 5 and Table 1). This suggests that it is possible to have some degree of diffusion of REEs in apatite at temperatures sufficient for thermally-activated volume diffusion, but without causing the rock and apatites to undergo metamorphic recrystallisation, demonstrating the applicability of apatite as a useful tool for tracking petrochronological
<figure><img src="image_10.png" /><figcaption>Figure 7: Apatite geochemistry diagrams plotting A) Sr/Y against sum of rare earth elements (REE), B) Th/U against sum of light REEs (La-Nd), C) Eu/Eu* against sum of light REEs (La-Nd) and D) Eu/Eu* whole rock against Eu/Eu* apatite. Plotted for comparison are igneous and metamorphic apatite from O'Sullivan et al. (2020)'s apatite geochemistry database.</figcaption></figure>
processes (Smye et al. 2018).

The ratio of light to heavy REE can be used to distinguish between primary, igneous apatite which is reflective of its original melt conditions and metamorphic apatite or apatite which has experienced subsequent geochemical alteration. For detailed discrimination of apatite geochemistry, the reader is referred to O'Sullivan et al. (2020) and references therein. High La/Lu<sub>N</sub> ratios (>50) correlate with apatite of igneous origin (i.e. samples KC82317-1, KC82317-3, KJJ09-07, KJJ09-08, GM-03, GM-05, and GM-08) and low La/Lu<sub>N</sub> ratios (<10) define apatite which has undergone either metamorphic recrystallization or mid-crustal fluid alteration (i.e. samples KC82816-1, KC82816-2, KC82816-3, KJJ09-03, MG3150, MG2465, and MG1655, Table 1 and Figures 3-5). Among primary igneous apatite, the samples for which whole rock were obtained (KJJ09-08, GM-03, GM-05, and GM-08) show that primary igneous apatite REE patterns are in good agreement with associated whole rock REE patterns but with far greater ΣREE values (Figure 4 and 6). Thus, apatite with high La/Lu<sub>N</sub> ratio can be used to investigate the composition of the primary melt from which it equilibrates. Finally, the range of magmatic and metamorphic processes observable in apatite indicates that apatite geochemistry has the potential to provide complementary information on melt-forming conditions in conjunction with commonly used accessory minerals such as monazite, titanite and zircon (e.g., Kylander-Clark et al. 2013, Chapman et al. 2016, Kirkland et al. 2020).

## 5.2 Link between *in situ* geochemical signature and cooling mechanisms

Apatite trace and REE geochemistry reflects crystallization processes as well as medium-temperature crustal processes such as exhumation, hydrothermal alteration, and magma-storage (e.g., Glorie et al. 2019, Odlum & Stockli 2020, Nathwani et al. 2020). However, apatite trace and REE geochemistry can also provide insight into low-temperature exhumation mechanisms. Specifically, the intra-sample variation in single grain trace elements suggests that REEs are undergoing modification via hydrothermal activity at temperatures that are not resetting the apatite U-Pb system. Here we integrate AUPb, AFT, and trace and REE geochemistry from well constrained tectonic settings in order to define the thermo-tectonic pathway of apatite from the mid- to upper crust (Figure 8 and 9).

### 5.2.1 Volcanic cooling

The combined U-Pb, FT, and trace element approach was applied to three volcanic rock samples in the Galiuro Mountains ignimbrite complex (Table 1). The Galiuro Mountains ignimbrite complex is dated at
its base as Oligocene (27.3 Ma, Dickinson 1991), which within error of the Eocene-Oligocene apatite U-Pb
and FT dates obtained in this study (41 ± 12 Ma, 28 ± 31 Ma, and 29 ± 12 Ma, Table 4 and Figure 6).
All three volcanic samples display high uncertainty which is likely a product of their rapid crystallization,
cooling, and geochemical homogeneity making it difficult to fit a non-radiogenic Pb discordia, however,
as they are within error of published constraints we view their ages as representative. The REE patterns
obtained for the Galiuro Mountain samples (GM-03, GM-05, and GM-08) all display little intra-sample
single-grain variation and give profiles which mirror their respective whole rock geochemistry (Figure 7 and
Figure 8). In addition, elevated Th/U and La/Sm ratios with low intragrain variability are characteristic
of the volcanic rock samples (Figure 7C). The Galiuro Mountains ignimbrite complex is interpreted to
be rapidly cooled from the melt during several Oligocene eruptions and has remained undisturbed at
surface temperature until present day (e.g., Dickinson 1991, Arca et al. 2010). Thus, we suggest that the
consistency between apatite U-Pb and FT ages, combined with elevated Th/U and a homogeneous trace
and REEs across single-grain analyses is indicative of a thermal history characterized by rapid cooling
from ~575°C to upper crustal temperatures of ~60°C without any secondary thermal or fluid alteration
(Figure 9).

**5.2.2 Laramide intrusions**

A relationship between medium to low temperature cooling and trace elements is also observed in Laramide-aged intrusions from the Catalina Mountain. Sample KJJ09-08 was collected from granite which yields a zircon U-Pb age of 73 ± 2 Ma which is in close agreement with the apatite U-Pb age of 62 ± 21 Ma and is consistent with the apatite fission-track age of 41 ± 6 Ma, suggesting a single, monotonic cooling history (Table 4 and Figure 4, Terrien 2012, Fornash et al. 2013). The close agreement of these high to low temperature thermochronometric systems is complimented by a homogeneous trace and REE distribution and consistently high Th/U and La/Lu ratios Figure 7B-8. This further emphasizes that homogeneous, elevated Th/U and La/Lu ratios is indicative of a monotonic cooling pathway to the upper crust (Figure 9).

Sample KJJ09-07 was collected from a granodiorite which yielded a zircon U-Pb age of 69 ± 3 Ma, a relatively precise apatite U-Pb age of 40 ± 4 Ma, and younger 24 ± 2 Ma apatite-fission track cooling age (Table 4 and Figure 4, Terrien 2012, Fornash et al. 2013). These ages are in relatively close agreement and could be interpreted as a single cooling history with discrete cooling pulses, however, the single-grain spread in apatite trace and REE (Th/U and La/Lu, Figure 7-8) suggests that this sample underwent
<figure><img src="image_11.png" /><figcaption>Figure 8: Plots showing A) Log-log of apatite U-Pb age minus apatite fission-track age against La/Lu<sub>N</sub> ratio, and B) Log-log of apatite U-Pb age minus apatite fission-track age against the median La/Lu<sub>N</sub> ratio divided by the standard deviation. Samples which have experienced protracted residence in the middle-crust correlate with an increased spread of La/Lu<sub>N</sub> ratios and higher La/Lu<sub>N</sub> coefficient of variance.</figcaption></figure>
an additional period of mid-crust alteration. Expanding on this, previous studies in the Catalina Core Complex have found a distinctive phase of Eocene-aged dikes throughout the range and Eocene zircon rim growth in the Laramide-aged intrusions (Fornash et al. 2013, Ducea et al. 2020). These samples contain moderate to high Th/U and La/Sm<sub>N</sub> ratios (Figure 7C). Increased magmatic activity in the Eocene would have elevated the geothermal gradient and resulted in hydrothermal fluids which percolated and cooled through the middle-to-upper crust resulting in a varying degree of alteration across single-grain REE geochemistry, but insufficient to deviate the apatite chemistry from the whole rock REE pattern (Figures 4 and 8). Thus, the variance in single-grain REE geochemistry, and Th/U, La/Sm<sub>N</sub>, the La/Lu<sub>N</sub> ratios, can be indicative of fluid alteration, medium-temperature thermal perturbation or prolonged residence (Figure 9).

### 5.2.3 Metamorphic core complex and detachment faulting

Apatite U-Pb dating has been shown to constrain the timing of faulting along major detachment fault systems as the apatite U-Pb temperature window (350-550°C) is ideally suited to constraining crustal evolution through the brittle-ductile transition zone (~10 km depth, e.g., Odlum & Stockli 2020). Samples KJJ09-03 and MG1655 were both collected from mylonitized basement exposed along the base of the Catalina and Pinaleño Core Complexes, respectively (Table 1). Zircon U-Pb dates from sample KJJ09-03 constrain crystallization to 56.6 ± 3 Ma (Fornash et al. 2013) whereas sample MG1655 is mapped as Paleoproterozoic in age (Johnson & Arca 2010). Apatite U-Pb ages from those same samples define a discordia age of 21 ± 5 Ma for sample KJJ09-03 and 39 ± 7 Ma for sample MG1655. Albeit with relatively large uncertainty (Figure 5), the AUPb age for sample KJJ09-03 is within error of the age for the onset of detachment extension of *ca.* 26.4 Ma along the Catalina detachment and sample MG1655 is slightly older than the *ca.* 29 Ma constraint placed along the Pinaleño detachment (Long et al. 1995, Peters et al. 2003, Terrien 2012). The close match between AUPb age and the hypothesized timing of mylonitization emphasizes the applicability of the apatite U-Pb to constrain timing of detachment faulting (Odlum & Stockli 2020). The REE geochemistry for both mylonitized samples display significantly altered profiles (depleted in LREE and enriched in HREE) when compared to their respective whole rock REE patterns (Henrichs et al. 2018), suggesting that the AUPb ages are reflective of apatite recrystallization during mylonitization. Zircons from sample KJJ09-3 displayed disturbed Hf isotopic values, which were attributed to mobilization of REEs during mylonitization (Fornash et al. 2013) demonstrating the distinct AUPb age and associated geochemistry are indicative of detachment faulting. Further, the La/Lu<sub>N</sub> ratio
<figure><img src="image_12.png" /><figcaption>Figure 9: A summary diagram illustrating the relationship between apatite U-Pb (AUPb), apatite (ap) and whole rock (wr) trace and rare earth element (REE) geochemistry, and apatite fission-track (AFT) across four Cordilleran tectonic terranes.</figcaption></figure>
is far more consistent across single-grains for sample KJJ09-03 than MG1655 (Figure 8), suggesting a
prolonged residence time in the middle crust leads to increased variation in apatite geochemistry.

Research from a Metamorphic Core Complex in the Pyrenees found that AUPb ages and geochemistry
were able to constrain multiple phases of hydrothermal fluid involvement associated with fault movement
(Odlum & Stockli 2020). In their study, the upper and lower limits of the discordia spread dated the
crystallization and exhumation ages, respectively (Odlum & Stockli 2020). In contrast, the two samples
from beneath the detachment in the Pinaleño Mountains MCC (MG3150 and MG2465) preserved two
distinctive discordia ages of *ca.* 1100 Ma, and *ca.* 400 Ma (Figure 5). These ages considerably pre-
date the onset of the Oligocene Pinaleño detachment (29 Ma) and are temporally unrelated to Paleogene
phase of exhumation along a detachment fault (Long et al. 1995). However, the *ca.* 1.1 Ga apatite U-Pb
age is within error of Rb-Sr whole rock and baddeleyite U-Pb of dikes within the Pinaleño Mountains
MCC (Swan 1976, Bright et al. 2014). Thus, we suggest that the younger discordia age relate to a
subsequent period of magmatism or hydrothermal flux in the region (Odlum & Stockli 2020). Further,
the preservation of multiple apatite U-Pb discordia within a single sample expands the range of potential
petrochronology applications for apatite (Figure 9, Garber et al. 2017, Smye et al. 2018). Finally, both
MG3150 and MG2465 demonstrate a broad range of Th/U, La/Sm<sub>N</sub>, and La/Lu<sub>N</sub> ratios, which could be
indicative of either protracted residence in the middle crust or due to multiple emplacement stages during
associated magmatic activity (Figure 7 and 8).

**5.2.4 Basement cored uplifts**

Previous studies have attempted to elucidate long-term cooling and exhumation from cratonic environ-
ments, finding that cratonic settings preserve distinct, stable low-temperature thermal histories (e.g.,
Flowers et al. 2006, Hall et al. 2018, Kohn & Gleadow 2019, McDannell et al. 2019). Cratonic basement
exhumed along basement cored uplifts (KC82317-1, KC82317-3, KC82816-1, KC82816-2, and KC82816-3)
all record Cretaceous AFT ages, suggesting that unlike cratonic environments in the Canadian Shield,
South Australia and Baffin Island, the western Wyoming Craton experienced sufficient burial to reset the
AFT low-temperature thermochronometer prior to the onset of the Laramide orogeny (Orme et al. 2016).
Apatite U-Pb ages from the western Wyoming craton yield three discordant ages at *ca.* 2.4, *ca.* 1.8, and
*ca.* 1.6 Ga (Figure 3).

Samples KC82317-1 and KC82317-3 display REE patterns that are indicative of primary igneous
apatite (O'Sullivan et al. 2018, Gillespie et al. 2018) and their AUPb ages of *ca.* 1.8 Ga are within error of
zircon U-Pb and hornblende and biotite <sup>40</sup>Ar/<sup>39</sup>Ar ages obtained from the Little Belt Mountains (Mueller et al. 2002, Holm & Schneider 2002, Foster et al. 2006). Therefore, we interpret these rocks (KC82317-1 and KC82317-3) to have remained at middle-to-upper crustal temperature from the Paleoproterozoic. This protracted thermo-tectonic history is reflected in the apatite REE composition (Figure 7), with high single-grain REE variability, low Th/U concentrations and a high La/Lu coefficient of variance (CV = (σ/μ) * 100, Figures 3, 7C and 8).

Similar to the two Little Belt samples, sample KC82816-2 (Gravelly Range, Montana, Table 1) preserves an AUPb age of *ca.* 1.8 Ga (Figure 3). However, the concave REE pattern suggests that sample KC82816-2 underwent metamorphism during the Little Belt arc accretion at *ca.* 2.0-1.8 Ga (e.g., Foster et al. 2006, Whitmeyer & Karlstrom 2007). In contrast, the nearby sample KC82816-1 preserves an Archean age of *ca.* 2.4 Ga and an apatite REE pattern indicative of meta-pelitic rocks (Henrichs et al. 2018), suggesting that the apatite U-Pb age records evidence of cratonic assembly along the western boundary of the Wyoming Craton (Kellogg et al. 2003, Mueller & Frost 2006). Sample KC82816-3 records the youngest cratonic age of *ca.* 1.6 Ga, these form a single discordia and a low La/Lu ratio (Figure 8). The 1.6 Ga discordia age likely reflects a widespread tectono-thermal event along southwestern Montana between *ca.* 1.65-1.63 Ga (Mueller et al. 2005, Foster et al. 2006), which would have reset the apatite U-Pb chronometer. However, sample KC82816-3 preserves both an igneous and metamorphic REE pattern (Henrichs et al. 2018), suggesting that thermal conditions were not sufficient to completely re-equilibrate the apatite and the sample underwent selective alteration in response to hydrothermal fluid interaction (Figure 3 and Supplementary File 9).

Finally, samples KC82816-1, KC82816-2, and KC82816-3 (Gravelly Range, Montana, Table 1) all display low La/Lu and Th/U ratios, relatively high La/Lu coefficient of variance, and high variability along single-grain apatite REE patterns (Figure 3 and 7). Considering that this region of the Wyoming Craton is interpreted to have undergone very little exhumation since the Proterozoic (Foster et al. 2006, Whitmeyer & Karlstrom 2007), and the low-temperature thermochronometric systems are completely reset in the Cretaceous (Table 3), we suggest that the high La/Lu coefficient of variance and variable apatite REE patterns are evidence for protracted residence in the middle crust (Figure 9).
562 6 Conclusions

563 Through the combined application of apatite U-Pb and FT thermochronology with *in situ* and whole rock
564 trace elemental geochemistry this study shows the viability of this petrochronological and thermochrono-
565 logical approach on a range of high to low temperature crustal settings. We demonstrate that the REE
566 patterns and the La/Lu<sub>N</sub> ratio can be interpreted as indicators to whether an apatite grain preserves a
567 primary (whole rock representative; typically La/Lu<sub>N</sub> >50) or secondary (geochemically modified; typ-
568 ically La/Lu<sub>N</sub> <50)) thermal history. If the apatite geochemistry preserves a primary thermal history,
569 then apatite geochemistry acts as a representative record of whole rock geochemistry similar to other com-
570 monly used accessory minerals. Apatite geochemistry applied to extrusive volcanic rocks demonstrates
571 that consistent single-grain REE patterns reflect monotonic cooling. In contrast, samples which have
572 undergone medium temperature (>250°C) thermal perturbation or samples that experience protracted
573 residence in the middle crust host considerable single-grain REE variation. Thus, in absence of double
574 dating, homogeneous intergrain REEs and elevated Th/U and La/Lu<sub>N</sub> ratios are indicative of a monotonic
575 cooling pathway through the upper crust. We further show the utility of apatite U-Pb thermochronol-
576 ogy in constraining the timing of detachment faulting (sensu Odlum & Stockli 2020). Finally, apatite
577 geochemistry has the potential to be an important complementary indicator of melt-forming conditions.
578 Applied in conjunction with other commonly used accessory minerals (monazite, titanite, and zircon),
579 the complex relationship between apatite thermochronology and geochemistry expands the applicability
580 of apatite as a middle-crust thermochronometer.

581 7 Acknowledgements

582 Daniel Alberts is thanked for his work in establishing the apatite U-Pb and trace element method at
583 the University of Arizona. James Worthington is thanked for his assistance with thin-section mineral
584 identification. Reviewers Margret Odlum and David Chew, and editor Catherine Chauval are thanked
585 for their insightful comments which improved the manuscript. Barbara Carrapa acknowledges US Na-
586 tional Science Foundation grant EAR Tectonics 1919179. Mihai N. Ducea acknowledges US National
587 Science Foundation grant EAR 1725002 and the Romanian Executive Agency for Higher Education, Re-
588 search, Development and Innovation Funding project PN-III-P4-ID-PCCF-2016-0014. George E. Gehrels
589 acknowledges US National Science Foundation EAR 1649254 for support of the Arizona LaserChron Cen-
590 ter. Antoine Triantafyllou is a FRS-FNRS post-doctoral research fellow for the PROBARC project (Grant
CR n°1. B. 414.20F). Antoine Triantafyllou thanks the Rotary Club de Mons and the University of Mons for providing their financial support via the Pierre Jacobs post-doctoral grant (2018).

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