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Earth and Planetary Science Letters 408 (2014) 296–306
Contents lists available at ScienceDirect
Earth and Planetary Science Letters
www.elsevier.com/locate/epsl
Thermal imprint of rift-related processes in orogens as recorded
in the Pyrenees
A. Vacherat a,b,c,∗, F. Mouthereau a,b,1, R. Pik c, M. Bernet d, C. Gautheron e, E. Masini f,
L. Le Pourhiet a,b, B. Tibari c, A. Lahfid g
aSorbonne Universités, UPMC Univ Paris 06, UMR 7193, Institut des Sciences de la Terr e Paris (iSTeP), 4 Place Jussieu, F-75005 Paris, France
bCNRS, UMR 7193, Institut des Sciences de la Te rre Paris (iSTeP), 4 Place Jussieu, F-75005 Paris, France
cCRPG, UMR 7358, CNRS-Université de Lorraine, BP20, 15 rue Notre-Dame des Pauvres, 54500 Vandoeuvre-lès-Nancy, France
dInstitut des Sciences de la Te rre (ISTerre), Univ Joseph Fourier, CNRS, 1381 rue de la Piscine, Grenoble 38041, France
eUniv Paris Sud, UMR GEOPS-CNRS 8148, Bâtiment 504, Rue du Belvédère, 91405 Orsay, France
fTOTAL, CSTJF, Avenue Larribau, 64016 Pau, France
gBRGM/MMA/MIN, 3 avenue Claude Guillemin, 45060 Orléans Cedex 2, France
a r t i c l e i n f o a b s t r a c t
Article history:
Received 1 July 2014
Received in revised form 8 October 2014
Accepted 10 October 2014
Avail able online xxxx
Editor: T.M. Harrison
Keywords:
thermal inheritance
detrital thermochronology
margin inversion
inverse modeling
Pyrenees
The extent to which heat recorded in orogens reflects thermal conditions inherited from previous rift-
related processes is still debated and poorly documented. As a case study, we examine the Mauléon basin
in the north-western Pyrenees that experience d both extreme crustal thinning and tectonic inversion
within a period of ∼30 Myrs. To constrain the time–temperature history of the basin in such a scenario,
we provide new detrital zircon fission-track and (U–Th–Sm)/He thermochronology data. The role of rift-
related processes in subsequent collision is captured by inverse modeling of our thermochronological
data, using relationships between zircon (U–Th–Sm)/He ages and uranium content, combined with
thermo-kinematic models of a rift-orogen cycle. We show that the basin recorded significant heating at
about 100 Ma characterized by high geothermal gradients (∼80 ◦C/km). Our thermo-kinematic modeling
and geological constraints support the view that subcontinental lithospheric mantle was exhume d at that
time below the Mauléon basin. Such a high geothermal gradient lasted 30 Myr after onset of convergence
at ∼83 Ma and was relaxed during the collision phase from ∼50 Ma. This study suggests that heat
needed for ductile shortening during convergence, is primarily inherited from extension rather than
being only related to tectonic and/or sedimentary burial. This should have strong implications on tectonic
reconstructions in many collision belts that resulted from inversion of hyper-extended rift basins.
©2014 Elsevier B.V. All rights reserved.
1. Introduction
The steady-state thermal structure of collisional orogenic belts
is controlled by upward advection of heat through the coupling
between crustal shortening and erosion (Royden, 1993; Stüwe et
al., 1994; Willett and Brandon, 2002). However, considering typ-
ical thermal relaxation time of several 100 Myrs for thick litho-
spheres (Jaupart and Mareschal, 2007), transient effects might not
be negligible for continental margins that experienced both ther-
mal resetting during thinning and structural inversion over a rela-
tively short period of time (Mouthereau et al., 2013). This process
*Corresponding author.
E-mail address: arnaud.vacherat@upmc.fr (A. Vacherat).
1Now at: Université Toulouse III – Paul-Sabatier, Laboratoire Géosciences Envi-
ronnement Toulouse, UMR 5563, 14 av. Edouard Belin, F-31400 Toulouse, France.
might be even more significant for inverted distal margins that
have experienced extreme crustal thinning and mantle exhuma-
tion (Manatschal, 2004). The pre-orogenic temperature anomalies
caused by crust/subcontinental lithospheric mantle (SCLM) thin-
ning, may therefore significantly impact the thermal history and
thermal-dependent ductile mechanisms of deformation in oro-
gens, but their magnitude has yet to be constrained. For instance,
Mesalles et al. (2014) using low-temperature thermochronological
data in southern Taiwan demonstrated that an originally hot distal
margin may record cooling only ∼20 Myrs after the end of rifting
due to the onset of continental accretion.
Here, we focus on the Pyrenees, where geochronological and
petrographic constraints indicate that rifting exhumed the SCLM
in Albian times (ca. ∼110 Ma) (Vielzeuf and Kornprobst, 1984;
Lagabrielle and Bodinier, 2008; Jammes et al., 2009; Lagabrielle et
al., 2010; Clerc et al., 2012) while plate convergence initiated at
∼83 Ma (Rosenbaum et al., 2002). The Mesozoic Mauléon basin in
http://dx.doi.org/10.1016/j.epsl.2014.10.014
0012-821X/©2014 Elsevier B.V. All rights reserved.
A. Vacherat et al. / Earth and Planetary Science Letters 408 (2014) 296–306 297
Fig. 1. A) Geologic map of the study area. Red stars and circles depict the position of studied samples and samples from which RSCM temperatures were obtained (Clerc and
Lagabrielle, 2014), including one new estimate from this study (Lu-1), re spectively. The ex tent of cleavage domain is shown as re d dashed area. B) Synthetic lithostratigraphy
of meso-cenozoic successions of the Mauléon basin and layer thickness inferred from borehole data (Fig. 5B). C) Geological cross-section of the western part of the Mauléon
basin, including the location of samples and the extent of cleavage domain, same as shown in A. Note that the ductile deformation domain is observed at the base of the
basin. NPFT: North Pyrenean Frontal Thrust, SPT: Saint-Palais Thrust, GRH: Grand-Rieu High. (For interpretation of the references to color in this figure legend, the reader is
referred to the web version of this article.)
the north-western Pyrenees (Fig. 1A) is interpreted as a preserved
hyper-extended rift system formed during the Late Aptian–Early
Albian, above a low-angle detachment system (Johnson and Hall,
1989; Jammes et al., 2009; Masini et al., 2014). This is supported
by field evidences of breccias of mantle peridotites reworked in
syn-/post-rift sediments of Albo-Cenomanian age, or tectonically
overlying the granulitic complex of the Labourd–Ursuya Massif
(Jammes et al., 2009).
To establish the time–temperature history of the Mauléon basin,
we inverse modeled detrital zircon fission-track and (U–Th–Sm)/He
thermochronological data collected for this study. Model results
were then compared to thermal patterns predicted from a 1D
thermo-kinematic numerical model of the evolution of hyper-
extended rift basins that are inverted during collision. Our re-
sults reveal that high geothermal gradients, inherited from hyper-
extension, are maintained over 30 Myrs after convergence initiated.
2. Hyper-extension in the Pyrenees and thermal constraints
The Pyrenean belt resulted from the inversion of previously ex-
tended domains of the Iberian and European plates from the Late
Cretaceous to the Late Oligocene (Choukroune et al., 1989 and ref-
erences therein). The Mauléon basin, located in the western part of
the North Pyrenean Zone, consists of folded Mesozoic sedimentary
units, thrust northward during the Tert iar y along the Saint-Palais
Thrust and the North-Pyrenean Frontal Thrust (Fig. 1). The basin
is a Late Aptian to Albo-Cenomanian sag basin interpreted as a
hyper-extended rift basin formed above a low-angle extensional
detachment system, which is identified on the northern flank of
the Labourd–Ursuya massif (Jammes et al., 2009) and at the base
of the Igountze–Mendibelza massif (Johnson and Hall, 1989). In
such a hyper-extended system, middle–lower crustal rocks and the
SCLM were exhumed (Jammes et al., 2009; Masini et al., 2014),
but age constraints on the timing of exhumation are still lacking.
The only age associated to this extension phase is obtained in the
eastern part of the Mauléon basin, where a gabbroic dyke, intrud-
ing the exhumed mantle body of Urdach, is sealed by Cenomanian
sediments (Jammes et al., 2009; Debroas et al., 2010), and yields
a relative flat Ar–Ar on biotite spectrum in the 105–108 Ma range
(Masini et al., 2014).
Alkaline magmatism, high-temperature (up to 600 ◦C) low-
pressure (HT-LP) metamorphism (Montigny et al., 1986; Golberg
and Leyreloup, 1990), and hydrothermal fluid circulation associ-
ated with talc-chlorite mineralization (Boulvais et al., 2006)are
observed elsewhere in the North Pyrenean Zone. These constraints
indicate a heating episode from 110 Ma to 85 Ma (Montigny et
al., 1986). Raman spectroscopy of carbonaceous material (RSCM)
shows that the Albian to Cenomanian series of the Mauléon basin
experienced shallow to mid-crustal temperatures of 180 to 295 ◦C
(Clerc and Lagabrielle, 2014). Determining whether these tempera-
298 A. Vacherat et al. / Earth and Planetary Science Letters 408 (2014) 296–306
tures reflect thinning processes is not trivial and requires a thermal
modeling approach.
An important delay occurred between the onset of plate con-
vergence at ∼83 Ma, (chron A34) used in plate reconstructions
(Rosenbaum et al., 2002) and the exhumation in the belt recorded
from ∼50 Ma to ∼20 Ma, as constrained by low-temperature ther-
mochronological data mainly in the Axial Zone (Yelland, 1990,
1991; Morris et al., 1998; Fitzgerald et al., 1999; Sinclair et al.,
2005; Jolivet et al., 2007; Maurel et al., 2008; Gunnel et al., 2009;
Metcalf et al., 2009). Time–temperature history (burial, heating
and cooling) during the initial accretionary stage are therefore
largely unknown. It may involve underestimated competing cool-
ing processes such as syn-orogenic thermal relaxation or cooling
by underthrusting as suggested recently in Taiwan (Mesalles et al.,
2014).
3. Sampling and methods
3.1. Strategy
Determining thermal histories of crustal rocks is classically
done using multiple low-temperature thermochronometers on
bedrock samples, but, in the Pyrenees, published bedrock low-
temperature data are only consistent with episodes of collision-
related cooling. To gain resolution on syn- to post-rift time–
temperature history, a direct approach relies on examining syn-rift
basins that recorded both extension and compression in the North
Pyrenean Zone. In this aspect, the Mauléon basin is suitable as it
experienced temperatures in the 110–295 ◦C range (Fig. 1A). Its
time–temperature evolution may therefore be resolved by com-
bining zircon fission-track (ZFT) and (U–Th–Sm)/He (ZHe) ther-
mochronology, which have their Partial Annealing Zone and Par-
tial Retention Zone (PRZ) between 160–270 ◦C and 140–220 ◦C,
respectively (e.g. Brandon et al., 1998; Guenthner et al., 2013).
When zircons reside in these temperature intervals, the result-
ing ages are highly dependent on their time–temperature histories
and diffusion kinetics. Recently published helium diffusion mod-
els reveal that apparent (U–Th–Sm)/He ages are controlled by the
amount of α-recoil damage, which is proportional to the effective
U concentration [eU] (Flowers et al., 2009; Gautheron et al., 2009;
Guenthner et al., 2013). Following these models, the thermal his-
tory explaining these ZHe age and eU correlations can be deduced
from inverse modeling.
The dataset consists of five detrital sandstone samples from
deep-water turbidites of the western part of the Mauléon basin,
deposited at 101 ±4Ma in Albo-Cenomanian (Su-1, Ar-2, Lu-1,
Ch-1, and Mi-1, Fig. 1A, B). Two basement samples from a granitic
intrusion (Itx-1) and gneiss (Lag-1) were collected (Fig. 1) to com-
plete these data with apatite (AHe) and ZHe analyses, and to
add independent constraints on collision-related cooling. Lu-1 was
analyzed with Raman Spectroscopy following the protocol devel-
oped by Lahfid et al. (2010). Samples were prepared at CRPG
(Nancy, France). They were crushed and zircon and apatite grains
were separated for low-temperature thermochronological analy-
ses using standard heavy-liquid and magnetic separation from the
61–280 μm fraction.
3.2. Zircon Fission track analysis
Zircon grains were handpicked for fission track dating and an-
alyzed at the thermochronology laboratory of ISTerre (Université
Joseph Fourier, Grenoble). We used standard ZFT preparation pro-
cedures as described by Bernet and Garver (2005). Grains were
mounted in a teflon sheet, polished to expose internal surfaces and
etched with NaOH–KOH at 228 ◦C between 8 and 13 h. Irradiation
was performed in the FRM II Research Reactor at the Technische
Universität München (Germany) with CN1 dosimeter glasses and
Fish Canyon Tuff age standards. Mica detectors used for external
detector method (Gleadow et al., 1976) and standards were then
etched in 48% HF at 21 ◦C for 18 min. Counting was carried out on
an Olympus BH2 optical microscope using FTStage 4.04 system of
Dumitru (1993). Individual fission-track ages were obtained using
zeta factors following approach of Hurford and Green (1983).
3.3. Zircon (U–Th–Sm)/He analysis
For (U–Th–Sm)/He dating, we handpicked 5 to 10 zircon grains
per bedrock sample and between 60 and 110 zircon grains per
detrital sample. Detrital zircons were mounted in epoxy and pol-
ished for future U/Pb analyses. Among those zircons, we retrieved
from the epoxy mounts between 20 and 35 zircons per detri-
tal samples for (U–Th–Sm)/He analysis, chosen so as to represent
the main peaks individualized in the U/Pb age distributions. All
these detrital zircons were measured as single grains. Addition-
ally, 4 to 5 replicates of 5 to 10 zircon grains were analyzed
in bedrock samples (Itx-1 and Lag-1). Zircon grains (prismatic
to round-shaped, 0 to 2 pyramids, with an equivalent spheri-
cal radius ranging from 35 to 60 μm) were then measured, and
loaded in Pt capsules for He extraction at CRPG. They were out-
gassed at 1500 ◦C for 20 min, and analyzed for He concentra-
tions with a VG603 noble gas mass spectrometer (Pik et al., 2003;
Godard et al., 2009). After total helium extraction, Pt capsules
containing zircon grains were retrieved for U, Th, and Sm con-
tent measurements at SARM (Nancy, France). Pt capsules were
opened, and loaded in Pt crucibles along with ultra-pure LiBO2and
ultra-pure B(OH)3for 2hat 990 ◦C in an automatic tunnel oven.
Then, the Pt crucibles were digested 12 h into acid. The prepa-
ration was then analyzed using an inductively coupled plasma
mass spectrometer. The overall precision of He ages determined
with this procedure is within 5–6% (1σ). Zircon grains whose He
and/or U concentrations are too close from the blank (e.g. for He
content less than 1.10−13 moles and for U concentration in the
solution less than 100 ppb after blank correction) were not con-
sidered for this study. Zircon ages were corrected for α-ejection
(FT) following Ketcham et al. (2011) (Tables DR2, DR3). To account
for the abrasion of the detrital zircon single grains, we consider
that we deleted ∼20 μm (the mean stopping distance, Ketcham
et al., 2011) and the half of the mean width of our zircon grains
(∼45 μm). Following Reiners et al. (2007), we corrected FTconsid-
ering an abrasion of 45 μm for each detrital grain.
3.4. Apatite (U–Th–Sm)/He analysis
Apatites were prismatic, with 0 to 2 pyramids, and with an
equivalent spherical radius ranging from 60 to 160 μm. We per-
formed AHe analyses at Paris-Sud University (Orsay, France) on
bedrock samples (Lag-1 and Itx-1, Fig. 1) following the procedure
described by Fillon et al. (2013). Four single grain replicates were
analyzed for Itx-1 and Lag-1 with 8% precision (1σ). AHe ages
were corrected for α-ejection following Ketcham et al. (2011) and
apatites with outlier Th/U ratios were excluded (Table DR4).
4. Results
ZFT analyses performed on samples Su-1, Ar-2, Ch-1, and Mi-1
yielded 23 to 63 dated grains per sample (Table DR1). Each sample
with an identical depositional age shows a similar age distribu-
tion. We therefore only present age component distributions for
the combined samples (n =171) (Fig. 2A). Most of the grains (97%)
are older than the depositional age, indicating very minor resetting
after deposition.
A. Vacherat et al. / Earth and Planetary Science Letters 408 (2014) 296–306 299
Fig. 2. Resul ts of thermochronological analyses and decomposition of statistically
representative age component of Albo-Cenomanian detrital samples. A) ZFT results
from Su-1, Ar-2, Ch-1, and Mi-1 samples. B) ZHe age distribution from Su-1, Ar-2,
Ch-1, and Lu-1 samples.
We decomposed our age distribution into age components us-
ing DensityPlotter (Vermeesch, 2009, 2012). The software repre-
sents distribution of ages using KDE (Kernel Density Estimation),
which is determined by stacks of Gaussian curves on top of each
measurement, whose standard deviation is determined by the lo-
cal probability density. Deconvolution for combined data returned
three age components (errors are given as ±2σ): two majors at
134 ±46 (P1, 17%) and 236 ±40 (P2, 79%), considered as cool-
ing events, and a minor population at 1005 ±886 Ma (4%) char-
acterized by a too important error to be statistically meaningful
(Fig. 2A).
ZHe analyses were carried out on the same Su-1, Ar-2, Ch-1
samples and on Lu-1. Ten to 27 grains were dated per sample
(Table DR2) and show similar ages and eU distributions. The age
distribution from the combined data ranges from 36 Ma to 131 Ma
(n =75, Fig. 2B) and yields four age peaks at 39 ±4 (16%), 50 ±3
(38%), 68 ±4 (33%), and 116 ±7 Ma (13%). Most of these detri-
tal ZHe grain ages (87%) are younger than the depositional age,
suggesting that they have been, at least, partially reset by post-
deposition burial. We will test the timing and amount of burial
and exhumation through numerical inversion of the data in the
next section.
ZHe analyses on two bedrock samples from the Labourd–Ursuya
massif (Lag-1 and Itx-1) give ages ranging from 51 ±5to 74 ±
7Ma, and from 61 ±6to 86 ±9Ma, respectively (Table DR3). AHe
single grain analyses performed on the same samples yield ages
ranging from 42 ±3to 49 ±4 Ma and from 35 ±3to 43 ±4Ma
for Lag-1 and Itx-1, respectively (Table DR4).
5. Thermal modeling of partially reset ages
It has been demonstrated that α-recoil damages associated to U
and Th decay, and their respective concentration (eU) could affect
He diffusion in apatites (Shuster et al., 2006; Flowers et al., 2009;
Gautheron et al., 2009). For high eU concentrations, the amount of
α-recoil damages increases with He retentivity and closure tem-
perature. Guenthner et al. (2013) highlighted the same trend in
zircons but only for relatively low concentrations of eU. For very
high eU concentrations, the He retentivity rapidly decreases. These
authors hypothesized that for very high eU content, the amount
of α-recoil damage is high enough so that damaged areas in the
crystal are interconnected and form through-going fast diffusion
pathways for He. Guenthner et al. (2013) showed that the evolu-
tion of He retentivity in zircons, which depends on the eU content,
Fig. 3. The statistical distribution of our ZHe age-eU data (blue empty squares) is
resolved from a 2D Kernel probability density function using a Parzen window ap-
proach (Matlab code available on request). Low eU content from 0 to 1100 ppm
correspond to a large distribution of ZHe ages from 65 Ma to 131 Ma (group B
in blue). These oldest ages are associated to a closure temperature of ∼220 ◦C. By
contrast, high eU content (>1100 ppm) only show young ZHe ages from 36 Ma to
65 Ma (group A in red), corresponding to lower closure temperatures (<140◦C).
controls both variations of the closure temperature and individual
annealing behaviors. This effect can lead to large ZHe age distri-
butions, under a given time–temperature path. The non-random
distribution of our ZHe age-eU dataset is supported by their sta-
tistical distribution in Fig. 3 (see caption for details concerning
the density function used) and suggests such a control. Originating
from a dense zone of young ZHe ages and low-eU grains two oppo-
site trends can be identified as ZHe-eU groups. A first group A (red
area) consists in young ZHe ages (from 36 to 65 Ma) associated
with a large eU distribution (from 400 to 4000 ppm). The group
B (blue area) corresponds to older ZHe age (from 65 to 131 Ma)
associated with low eU values only (from 0 to 1100 ppm).
Following Guenthner et al. (2013), the oldest ages could corre-
spond to zircon grains that have been less resetted due to a higher
closure temperature (∼220 ◦C). Such zircons require a longer resi-
dence time in the PRZ to be reset. In contrast, the young ZHe grain
ages that display eU >1100 ppm would correspond to a lower
closure temperature of <140 ◦C.
To determine the time–temperature paths of these zircon
grains, we used the HeFTy soft (Ketcham, 2005) that includes
the kinetic model of Guenthner et al. (2013). The limited number
of grains (seven) that can be input in the HeFTy inverse model-
ing procedure do not allows direct inversion of the entire dataset
and requires to identify representative individual ZHe age-eU pairs
within the two groups observed in Fig. 3. These two trends which
originate in the red high density zone of Fig. 3 can be easily and
robustly described by a couple of representative samples. In order
to describe the entire range of age-eU distribution, seven repre-
sentative samples have been taken along the A and B groups and
used for distinct sets of inversion. Various tests demonstrated that
the use of representative samples is not an issue in this inversion
procedure.
Because the A and B groups have been potentially controlled
by distinct closure temperatures linked to the amount of radiation
damage accumulated in the zircon grains (Guenthner et al., 2013)
it is crucial to take into account the ZFT data obtained for these
zircons in the inversion modeling. The ZFT data also exhibit two
distinct populations characterized by peak ages at P1 (∼134 Ma)
and P2 (∼236 Ma) that represent two independent cooling histo-
ries prior to deposition (at ∼100 Ma). At that time, the amount of
300 A. Vacherat et al. / Earth and Planetary Science Letters 408 (2014) 296–306
Fig. 4. A) Time–temperature histories extracted from HeFTy inverse modeling constrained by ZFT data, ZHe age-eU pairs, and depositional ages, for each model (P1-A,
P1-B, P2-A and P2-B). B) ZHe age-eU statistical distribution of Albo-Cenomanian detrital zircon grains. We compare the ZHe age-eU correlation corresponding to the best
time–temperature path of each model with the data.
accumulated damages was therefore significantly higher for the P2
population and could have triggered differential He diffusion when
the sediments have been subsequently buried and re-heated. It is
however not possible to directly relate one ZFT population with
one ZHe ages group. Consequently both P1 and P2 ZFT populations
have been used alternatively as input parameters for the inversion
modeling.
Four sets of inversion models have therefore been tested
(Fig. 4), corresponding to the various combinations using ZFT pop-
ulations (P1, P2) and ZHe ages groups (A, B). Models are char-
acterized by: (i) different pre-deposition histories constrained by
the P1 and P2 ZFT characteristics (134 ±15 Ma/240 ±40
◦C and
236 ±20 Ma/240 ±40
◦C respectively), (ii) identical depositional
age at 110 Ma, (iii) a free post-depositional time–temperature his-
tory inverted for 7 representative ZHe ages of groups A and B
independently. To reproduce the partial reset signature of the
ZHe data, we constrained the software to search post-deposition
time–temperature paths from shallow to mid-crustal temperatures,
corresponding to a range from 20 ◦C and 200 ◦C. The inversion
consisted of randomly testing 300,000 time–temperature paths for
each model.
The P1-A model returned 439 acceptable and 180 good solu-
tions. The P1-B model returned only 70 acceptable and no good
solutions. The P2-A model only returned 35 acceptable and no
good solutions. The P2-B model returned 405 acceptable and 62
good solutions. The best time–temperature path of each model
A. Vacherat et al. / Earth and Planetary Science Letters 408 (2014) 296–306 301
Fig. 5. A) Total decompacted thickness of sediments in the Mauléon basin, as obtained by combining seven boreholes within or close to the Mauléon basin for maximum,
minimum and weighted averaged (red curve) estimations. Tem pora l influence intervals resolved from each borehole is shown in grey. Vertical dashed lines represent the
period of heating highlighted in model of Fig. 4 and the horizontal dashed lines correspond to the thickness of sediments deposited during this period. A mean valu e of
2km of sediments was deposited between 105 and 70 Ma. B) Map showing the location of the different boreholes. (For interpretation of the references to color in this figure
legend, the reader is referred to the web version of this article.)
corresponds to a ZHe age-eU correlation, which is compared to
the data in Fig. 4B. The P1-A model better fits the group A than
the P2-A model, which fails to reproduce the data with eU >
2000 ppm. The P2-B model is in better agreement with the group
B than the P1-B model. Models P1-A and P2-B return the best
correlations between eU and ZHe ages that reproduce the data
(Fig. 4B). We infer that all Albo-Cenomanian ZHe detrital data are
obtained by the combination of these two time–temperature mod-
els (Fig. 4).
These models show a consistent post-100 Ma thermal his-
tory. In particular, time–temperature paths of our zircon grains
are consistent with heating to temperatures of ∼180 ◦C soon after
∼100 Ma at an average heating rate of ∼5◦C/Myr. Several of these
pathways show a nearly isothermal stage established at ∼80 Ma
which lasted maximum ∼30 Myr. This heating/isothermal stage
was followed by a relatively rapid cooling stage from 50 Ma to
present (∼3◦C/Myr). This cooling path is not very precisely de-
scribed by the inverse modeling because of the lack of additional
lower temperature thermochronometers.
6. 1D thermo-kinematic modeling of rift-to-collision evolution
The results of modeling (Fig. 4) show that the Albo-Cenomanian
zircon grains were heated to a temperature of ∼180 ◦C during
the post-breakup evolution of the Mauléon basin. To constrain the
geothermal gradient associated with this heating episode, a few
Myrs only after deposition of the sampled Albo-Cenomanian rocks
(see Fig. 4A), we first need to determine the thickness of the en-
tire pile of sediments in the Mauléon basin. The complete burial
history shown in Fig. 5A was resolved by combining well data
from different boreholes, including boreholes in the Arzacq basin,
North of the Mauléon basin (for instance Lacq 301, Brunet, 1984),
and wells drilled in the Mauléon basin. They are from South to
North: Ainhice 1, Chéraute 1, Uhart-Mixe 1 and Saint-Palais 1 for
the Triassic to Late Cretaceous history, and Lahontan 1bis, Lacq 301,
and Nassiet 1 for the Late Cretaceous to the Late Eocene deposits
(Fig. 5B). Estimates of minimum and maximum thicknesses at time
of deposition of the studied samples from 105 to 70 Ma are indi-
cated in Fig. 1B. We estimate that a mean sediment thickness of
∼2km was deposited above the studied samples during this pe-
riod of time (Fig. 5A). A temperature of 180 ◦C at ∼2km depth, as
suggested from the ZHe data, allows to define a geothermal gra-
dient as high as ∼80 ◦C/km (assuming a surface temperature of
20 ◦C).
To examine the tectonic conditions that led to the observed
cooling history, we consider two different end-member thinning
processes (Figs. 6A, B, and 7) that are thought to embody most of
the fundamental characteristics of rifted margins as summarized
by Huismans and Beaumont (2011). A stepwise tectonic evolu-
tion from 130 Ma to 0 Ma of a lithosphere section below the
Mauléon basin involves thinning through a rifting phase from the
Early Barremian (130 Ma) until the Late Cenomanian (95 Ma)
with a breakup occurring at 110 Ma. This is followed by inver-
sion and underthrusting of the thinned lithosphere from 83 Ma
until the Early Eocene (50 Ma) that marks the onset of crustal
thickening and thrust-related exhumation. In order to test these
two hypotheses, we adopt a 1D forward thermo-kinematic model-
ing approach. The thermochronological modeling showed that the
Albo-Cenomanian series in the Mauléon basin were heated to a
temperature of ∼180 ◦C, which was maintained during ∼30 Myrs
302 A. Vacherat et al. / Earth and Planetary Science Letters 408 (2014) 296–306
Fig. 6. 1D thermal–kinematic models tested for the Mauléon basin. A) Model A:
crustal breakup at 110 Ma and SCLM is thinned and exhumed to the base of the
basin until 95 Ma. Dashed red lines on model A correspond to tests consider-
ing SCLM thickening from onset of convergence at 83 Ma to mature collision and
exhumati on after 50 Ma. B) Model B: SCLM breakup occurs at 110 Ma and the con-
tinental crust thins until 95 Ma, lying in contact with the asthenosphere. Tes ted
thinning factors βAand βBare 4, 10, and ∞(breakup) in both models. (For inter-
pretation of the references to color in this figure legend, the reader is referred to
the web version of this article.)
(Fig. 4). We therefore attempt to retrieve from the model the time-
depth evolution of the 180◦C isotherm by varying thinning factors
for crust and mantle.
The thickness of the Mauléon sedimentary layers at time of de-
position of Albo-Cenomanian is constrained by the subsidence his-
tory of the Mauléon basin defined in Fig. 5A. We consider an initial
thickness of 2km of sediments above the continental crust that
increases to a final thickness of 6km. The current Moho depth is
estimated to 32 km in the region of the Mauléon basin (Daignières
et al., 1982; Jammes et al., 2010; Chevrot et al., 2014) leading to
consider a final thickness of continental basement of 28 km. We
hypothesize that the thickness of the continental basement was
the same before the onset of rifting. At the initial and final stages
we consider that the lithosphere is in equilibrium stage and adopt
a typical thickness value for a Phanerozoic continental lithosphere
of 130 km (e.g., Poudjom Djomani et al., 2001) (Fig. 6).
In model A, the crust is thinned until breakup occurs (Fig. 6A).
After crustal breakup, the SCLM is thinned and exhumed at the
base of the Mauléon basin, according to variable amount of thin-
ning factors βAranging from 4, 10 and higher (SCLM breakup). In
model B, the SCLM is thinned until its breakup, leading to the rise
of hot asthenosphere below the continental crust (Fig. 6B). Sim-
ilarly to model A, model B is run for the same variable amount
of thinning factor βBfor the continental crust. We also test the
impact of the thickening of the SCLM (model A) or the crust
(model B), after convergence initiated at 83 Ma, on the thermal
evolution of Mauléon basin. This was performed by taking into
account the accumulation of syn-orogenic sediments under local
isostatic conditions.
For both scenarios, the role of transient diffusive heat relax-
ation and advection related to basin subsidence history, isostasy
and rock uplift is quantified. It accounts for realistic diffusivity and
heat production distribution. To simulate the effect of a high ther-
mal conductivity layer represented by the Triassic evaporites, the
basin rests above a 1 km-thick Triassic salt layer, which thickness
is kept constant during the simulation. Although fluid circulations
may play a key role during extension by maintaining high tem-
peratures below the basin and favoring heat transfers, we kept
the model as simple as possible so as to depend on a minimum
of unknown parameters, as the basin evolved from extensional to
compressional tectonic settings.
Model results show that the depth of the 180 ◦C isotherm is
controlled to first order by the amount of thinning of the SCLM
(Fig. 7). This effect is most significant for model A (SCLM exhuma-
tion) in which the depth of the 180 ◦C isotherms is seen to vary
between 1 and 5 km as a function of the amount of thinning. In
the model B, this is less apparent because the asthenosphere is
kept closer to the surface (from 7km depth to surface depending
on the βconsidered) during all the experiment.
Prior to crustal breakup at 110 Ma, model A and B show very
different thermal responses to rifting. An upward deflection of
isotherms is observed for the model B, while model A indicates
a cooling phase before heating. Thermal evolution in model B re-
flects the upward advection of the base of the lithosphere during
thinning, which is maximum when the SCLM breakup is achieved.
In model A, a delay is observed between the onset of crustal and
SCLM breakup. This reflects the loss of the radiogenic heat source
caused by crustal breakup, which is not instantaneously compen-
sated by advection of heat caused by SCLM thinning.
During the inversion phase, the 180 ◦C isotherms are main-
tained to the same depth from 95 to 50 Ma for both models (red
Fig. 7. Comparison of the Mauléon basin burial history with the depth of the 180◦C isotherm predicted from 1D rift-to-collision thermal models (A and B) shown in Figs. 6A
and 6B, respectively. Depth evolution of the Albo-Cenomanian deposits (grey) is distinguished from the Meso-Cenozoic successions (green) and water (blue). Depth of the
180 ◦C isotherms produced by different thinning factors (βAand βB) meets the position of the studied samples relatively soon after breakup of the crust (models A and A)
or mantle (model B) at 110 Ma. The isotherm is kept at a constant depth after onset of tectonic inversion. C: Crust, S: SCLM, A: Asthenosphere. (For interpretation of the
references to color in this figure legend, the reader is referred to the web version of this article.)
A. Vacherat et al. / Earth and Planetary Science Letters 408 (2014) 296–306 303
and black curves in Fig. 7) corresponding to a limited linear in-
crease of heat with respect to the base of the Mauléon basin. When
we account for the thickening of the SCLM or the crust below the
basin during the underthrusting phase, the 180 ◦C isotherms re-
main flat from 83 Ma to 50 Ma for both models A and B. The
progressive deepening of the 180 ◦C isotherms after 50 Ma reflects
the primary effect of thermal relaxation and the deepening of the
SCLM, relative to the heat advection due to erosion.
7. Discussion
7.1. Comparison between thermochronological data and
thermo-kinematic modeling
In this study, we focus on the thermal history of the Mauléon
basin from Albo-Cenomanian times until today. However, because
our ZFT ages are only slightly reset (10%) with no significant in-
fluence on age populations, we can assume that P1 (∼134 Ma)
and P2 (∼236 Ma) reflect two cooling events that occurred prior
to deposition. The P1 cooling event appears to be consistent with
an extension event recorded in the southern Pyrenees in the Early
Cretaceous at ca. 145–132 Ma (Vergés and García-Senz, 2001). Zir-
cons cooled between 150 Ma and 100 Ma (P1) from mid-crustal to
surface temperatures may reflect denudation in the footwall of a
rolling-hinge normal fault (Axen and Bartley, 1997). The P2 event is
also coherent with a magmatic (Rossi et al., 2003) and/or exhuma-
tional event during the Triassic, as recognized in Albian sediments
in the southern Pyrenees (Filleaudeau et al., 2011).
On the other hand, ZHe age data show a complex and large
distribution from 36 to 131 Ma with most of them younger than
depositional age. This is typical of partial resetting and, for a
given duration of thermal event, its amplitude could have been
controlled by various factors including: (i) the size of the grains,
(ii) the initial age distribution of grains, (iii) the position in the
PRZ during re-heating, (iv) the residence time above the PRZ and
the amount of α-recoil damages accumulated before re-heating
(Guenthner et al., 2013). Inversion of ZHe data with thermochrono-
logical models (Fig. 4) suggests that zircon grains have been heated
to temperatures up to ∼180 ◦C soon after deposition ∼100 Ma ago
(Fig. 4A, B). This is consistent with our thermo-kinematic mod-
els A and B showing that the basin was already hot at the end
of the rifting phase (95 Ma), due to upward deflection of the
180 ◦C isotherm reaching the depth of the Albo-Cenomanian series
at 2km for β=10 or higher (Fig. 7). After this heating phase, both
our thermochronological models P1-A and P2-B require that zircon
grains were maintained at this temperature of 180 ◦C through a
nearly isothermal stage until 50 Ma. This period corresponds to
the inversion phase of the thermo-kinematic models, where the
180 ◦C isotherm depth remains constant from 95 Ma to 50 Ma.
The youngest ZHe population from 60 to 40 Ma is associated to
the largest eU concentration distribution (from ∼0 to 4000 ppm,
Fig. 3) and corresponds to the lower limit of the He-PRZ (clo-
sure temperature lower or equal to 140 ◦C, Guenthner et al., 2013).
These youngest ages are directly related to the main episode of
cooling that affected the Mauléon basin since the Eocene. This
is consistent with our thermo-kinematic models that indicate a
progressive cooling driven by mantle subduction and thermal re-
laxation during the orogenic phase (Fig. 6). This directly led to the
compensation of the hot thermal anomaly previously emplaced, as
plate collision and crustal thickening initiated at 50 Ma. In the
absence of very-low-temperature thermochronological constraints,
the results of the inversion models (Fig. 4A) do not lead to precise
t–Tscenario concerning this late phase of cooling. Whether such
cooling through the He-PRZ of zircons was mostly achieved early
(50–40 Ma) and driven by thermal relaxation or whether part of
this cooling occurred later in the Pyrenean orogenesis (40–25 Ma)
and was driven by exhumation is not precisely expressed in
the models. However thermo-kinematic models conducted in this
study clearly demonstrate that thermal relaxation during exhuma-
tion, following transient upward deflection of isotherms, represent
a significant cooling process that must be taken into account in the
interpretation of thermochronological data in this range of temper-
ature. In the Pyrenean belt this is particularly true for the North
Pyrenean Zone which experienced large-scale hyper-extension re-
lated high geothermal gradients.
Our simple approach did not allow the evaluation of the role
of the fluids effect in the Mauléon basin, but the good agreement
between model and data suggests its role might be minor at least
from a thermal perspective. However, fluid flow and serpentiniza-
tion of the exhumed mantle in such settings may be prominent
processes allowing the localization of deformation during exten-
sion.
7.2. Implications for the evolution of the Pyrenees
After deposition, Albo-Cenomanian zircon grains were heated
to a temperature of ∼180 ◦C during the post-breakup evolution of
the Mauléon basin. At this time, the basin was presumably floored
by the exhumed mantle as shown by geological evidences sum-
marized in Jammes et al. (2009): reworked granulites and man-
tle peridotites in Albian sediments, and tectonic relationship with
SCLM exhumation. These geological data best support a model A
hyper-extended rift basin (Fig. 6A) even if both models A and B
rift basins reproduce the thermal history of the basin (Fig. 6A, B).
This heating phase was characterized by a geothermal gradient as
high as ∼80 ◦C/km consistent with RSCM temperatures (180 ◦C to
295 ◦C) and HT-LP metamorphism of pre-Cenomanian sedimentary
units (Fig. 1).
Heating in the basin ceased rapidly from ∼80 Ma on. This stage
was followed by a rather isothermal period that initiated coevally
with the onset of plate convergence at 83 Ma. Both tempera-
ture and geothermal gradient were then kept at a high level for
30 Myrs, until 50 Ma when cooling/exhumation started associated
with mountain building. The persistence of high surface thermal
flow and geothermal gradients 18 Myrs after sea-floor spreading
has been reported in present-day rifted margins of the Gulf of
Aden (Lucazeau et al., 2010; Rolandone et al., 2013). In the case
of the Mauléon basin, the temperature structure acquired during
the rift phase prevailed at the earliest stage of continental accre-
tion. This is in marked contrast with thermal evolution reported,
e.g., in Taiwan (Mesalles et al., 2014) where rapid underthrusting
of the lower plate (50–80 mm/yr) at onset of continental accretion
led to downward deflection of isotherms. This cooling phase is not
detected in the early accretionary prism stage of the Pyrenees. We
interpret this difference as a consequence of limited lateral heat
advection induced by a much slower plate convergence of only
3–4 mm/yr (Mouthereau et al., 2014).
Our result reveals that onset of shortening in the Mauléon basin
occurred in an abnormally hot basin. Due to the absence of signifi-
cant nappe stacking in the region, we argue that ductile shortening
documented in the inverted rifted basin results from high temper-
atures inherited from rifting rather than syn-convergence burial. It
is characterized by axial-planar and crenulation cleavages in folded
Albian to Cenomanian units of the Mauléon basin that reveal am-
bient temperatures of 100–200 ◦C (Choukroune, 1974), consistent
with our models.
Fission track analyses on the Labourd–Ursuya Massif yield two
ages at 42.2 ±2.4 Ma and 48.3 ±2.3Ma on apatites, and an
age at 81.8 ±3.1Ma on zircons (Yelland, 1991). Thus, ZFT and
ZHe ages from the Labourd–Ursuya Massif indicate initial cooling
from probably deeper crustal temperatures at 80–50 Ma, show-
ing a different thermal history from the Mauléon basin. AFT and
304 A. Vacherat et al. / Earth and Planetary Science Letters 408 (2014) 296–306
Tabl e A.1
Thermal and mechanical parameters considered for each type of rock in the model.
Sedimentary cover Basement
Deposits Triassic salt Continental crust SCLM Asthenosphere
Thermal conductivity k
(W/(m K))
2.25 6.5 2.25 3.3 3.3
Heat capacity Cp
(m2/(m s2))
900 840 900 750 750
Heat production H
(μW/m3)
0.9 0 0.6 0.009 0.009
Density ρ
(kg/m3)
2500 2170 2800 3300 3300
AHe ages (ranging from 49 ±4 Ma and 35 ±3Ma) in the Western
Pyrenees suggest that cooling/thermal relaxation of high tempera-
tures after 50 Ma occurred synchronously with the North Pyrenean
massifs in the Eastern and Central Pyrenees (Morris et al., 1998;
Fitzgerald et al., 1999; Yelland, 1991), as a result of crustal thick-
ening and erosion. The thermal relaxation observed after 50 Ma
in the Mauléon basin therefore appears related to a major and re-
gional exhumational phase in the Pyrenees. Because erosion is one
of the main agent in orogenic belts bringing heat closer to the
surface, it may seem counterintuitive that thermal relaxation oc-
curred during the main exhumational phase. Processes other than
erosion may therefore explain the thermal relaxation. Heat advec-
tion recorded in the Mauléon basin remained limited first because
only 2 km of basin sediments were eroded since 50 Ma. In ad-
dition, our thermo-kinematic experiments (Fig. 6) show that the
emplacement the Mauléon basin onto a thicker and colder fore-
land lithosphere compensates heating due to exhumation.
8. Conclusions
This study demonstrates that the analysis of low-temperature
thermochronological constraints performed on pre-/syn-rift sedi-
ments preserved in mountain belts is effective in resolving the
long-term post-rift and syn-convergence thermal evolution of
rifted margins and hyper-extended rift basins. When combined
with thermal–kinematic models of rift-to-collision evolution, our
data allowed to test hypotheses on the thinning processes between
crust and the lithospheric mantle that cause the reconstructed
time–temperature history.
Our low-temperature thermochronological data show that the
sediment succession of the Mauléon basin recorded a phase of
heating following breakup in the Albo-Cenomanian as a result
of extreme extension. The Albo-Cenomanian sandstones reached
temperatures of 180 ◦C at only ∼2km depth, corresponding to a
geothermal gradient of ∼80 ◦C/km.
Using this approach we demonstrate that the thermal structure
of the Mauléon basin is consistent with extreme thinning, although
the relative thermal effect of breakup of the SCLM and crustal
breakup can hardly be differentiated. The temperature anomaly in-
herited from extreme thinning lasted 30 Myrs, from ∼80 Ma to
∼50 Ma. This inherited thermal anomaly explains ductile shorten-
ing identified in the inverted basin. It provides a mechanism for
explaining the observations of abnormally high temperatures (rel-
ative to inferred burial), syn-convergence MT or HT metamorphism
and ductile deformation in post-rift sediments. On the other hand,
these tectono-metamorphic characteristics are diagnostic of highly
extended rift basin inverted relatively soon after its emplacement.
Thermal relaxation of the rift-related heat anomaly occurred dur-
ing the main stage of the orogenic development, when the hyper-
extended rift basin was thrusted onto a colder and thicker Euro-
pean plate ∼50 Myrs ago. The Pyrenees give us a vivid example of
how high temperatures inherited from the rifting can affect the
thermal structure of the early stages of the collision, and how
these temperatures are relaxed during the late stage of orogenic
processes.
Acknowledgements
This study was supported by French National Research Agency
(ANR Project PYRAMID, ANR-2011-BS56-0031). We thank G. Man-
atschal, Peter Reiners, and R.S. Huismans for constructive discus-
sions. We also thank the staff of the different laboratories (CRPG,
SARM, ISTerre, IDES) for welcoming us and for their precious help
during data acquisition. Comments by two anonymous reviewers
have been helpful in clarifying some points of the manuscript. This
is CRPG contribution number 2344.
Appendix A
The code of our 1D thermo-kinematic model solves the tran-
sient heat advection diffusion equation (A.1), including heat pro-
duction in one dimension:
−∂
∂zk∂T
∂z+ρCpvz
∂T
∂z=ρCp∂T
∂t+ρH(A.1)
Density, ρ, heat capacity, Cp, heat production Hand the heat con-
ductivity, k, are given constant values for each rock type and are
listed in Table A.1. The solution is obtained using a standard im-
plicit in time centered finite difference scheme at each time step.
However, in order to allow for advection of the 1330 ◦C isotherm,
or to allow for erosion and sedimentation, the model domain is
remeshed at every time step.
The material advection parameter is treated independently of
the mesh using pre-computed level-set functions that define the
limit between each material phase (sediment, basement crustal
rocks, mantle rocks), excluding artificial diffusion of material prop-
erties with time.
In the models, we assume that velocity, vz, in the rock column
can be interpolated linearly between each petrologic interface. Be-
neath the lowest interface, velocity is constant and equal to the
velocity of that interface. This ensures that the 1330 ◦C isotherm
imposed at the base of the model is not tight to rock uplift and al-
lows for thermal relaxation to occur. Similarly, to enable effects of
rock uplift or sedimentation, the 20 ◦C isotherm is imposed at the
surface of the Earth, but velocity at the surface is equal to that of
the shallowest rock interface.
Initial conditions are obtained by solving the heat diffusion
equation (A.2) at steady state using defined material properties:
−∂
∂zk∂T
∂z=ρH(A.2)
This avoids artificial thermal re-equilibration, which would relate
to ill-defined initial geothermal gradients that would not be consis-
tent with the material properties and particularly heat production
distribution.
A. Vacherat et al. / Earth and Planetary Science Letters 408 (2014) 296–306 305
Appendix B. Supplementary material
Supplementary material related to this article can be found on-
line at http://dx.doi.org/10.1016/j.epsl.2014.10.014.
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