On the Early Permian shape of Pangea from paleomagnetism at its core

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On the Early Permian shape of Pangea from paleomagnetism at its core

Mathew Domeier

^a,

⁎ , Eric Font

^b,c

, Nasrrddine Youbi

^d,e,f

, Joshua Davies

^g

, Samantha Nemkin

^h

, Rob Van der Voo

^h

, Morgann Perrot

^g

, Mohamed Benabbou

ⁱ

, Moulay Ahmed Boumehdi

^c,d

, Trond H. Torsvik

^a,j

aCentre for Earth Evolution and Dynamics, University of Oslo, Norway

bDepartment of Earth Sciences, University of Coimbra, Portugal

cInstituto Dom Luiz, University of Lisbon, Portugal

dDepartment of Geology, Cadi Ayyad University, Morocco

eGeology and Sustainable Mining Department, Mohammed VI Polytechnic University, Morocco

fFaculty of Geology and Geography, Tomsk State University, Russia

gDepartment of Earth and Atmospheric Sciences, University of Quebec, Canada

hDepartment of Earth and Environmental Sciences, University of Michigan, United States

iDepartment of Geology, Sidi Mohamed Ben Abdellah University, Morocco

jSchool of Geosciences, University of Witswatersrand, South Africa

a b s t r a c t a r t i c l e i n f o

Article history:

Received 14 July 2020

Received in revised form 7 October 2020 Accepted 12 November 2020 Available online 21 November 2020 Keywords:

Pangea Paleomagnetism Paleogeography Permian Tectonics

Although central to an understanding of Earth's paleogeography and the myriad processes that it affects, the geometry of Pangea during the early phase of its lifetime has remained a topic of contention since the plate tectonic revolution. Despite decades of analysis and discussion, the crux of this debate still largely hinges on sparse, legacy paleomagnetic data derived from early Permian rocks of the Moroccan Meseta. In this work, we present the results of a study designed to revisit, update and expand on those key data, with the provision of new geochronologic and paleomagnetic results from six Permo-Carboniferous basins in central Morocco. New U-Pb zircon ages from volcanic rocks substantiate and refine existing geochronological data and reveal that volcanism among the studied basins spanned approximately 30 Ma, from at least 305 to 277 Ma, but was possibly punctuated by a more intense pulse in the mid-early Permian (~285 Ma). These new U-Pb data furthermore suggest that the age estimates previously assigned to the key poles from the Moroccan Meseta are probably too young, by some ~10 Ma. New paleomagnetic results from six basins yielded a common, well-defined remanent magnetization from 20 sites that is demonstrated to be pre-middle Permian in age. However, it remains unclear whether this magnetization (which is directionally similar to the previous paleomagnetic data from the same basins) is a primary magnetization acquired at ~285 Ma, or a syn-folding remagnetization that was acquired shortly there- after, at ~275 Ma. In adopting both interpretations as alternative hypotheses, we examine the corresponding pa- leogeographic implications of both. Through exhaustive comparisons with reference apparent polar wander paths and the direct reconstruction of Gondwana using these new paleomagnetic results, we demonstrate that in either case (i.e. whether the remanent magnetization is primary or was acquired soon after during folding) the data is compatible with a Pangea A geometry during the early Permian. We close with a topical discussion of some persisting arguments used to defend Pangea B and why we consider them to beflawed.

© 2020 The Authors. Published by Elsevier B.V. on behalf of International Association for Gondwana Research. This is an open access article under the CC BY license (http://creativecommons.org/licenses/by/4.0/).

1. Introduction

Despite its central importance to an understanding of the late Paleo- zoic and early Mesozoic, the paleogeography of Pangea remains a con- troversial topic. Although early debates included discussions of the paleogeography of Pangea throughout its lifetime (De Boer, 1965;

Irving, 1977), it is now widely agreed that Pangea, at least from late Permian until its Early Jurassic breakup, was organized in a form much like the original reconstruction ofWegener (1915), also called‘Pangea

A' (e.g.Domeier et al., 2011b;Kent et al., In Press). This reconstruction results from simple closure of the Atlantic, juxtaposing the eastern margin of North America against northwest Africa (Fig. 1). Vigorous debate continues, however, about whether Pangea had been directly assem- bled into this conﬁguration during its initial amalgamation at

~320 Ma, or if it rather started in a different form and was later trans- formed into a Pangea A geometry during the early to mid-Permian (Domeier et al., 2012;Gallo et al., 2017;Muttoni et al., 2009). Several alternative conﬁgurations of Pangea have been proposed for this late Car- boniferous to early Permian interval, but the only persistent alternative is‘Pangea B', which places Gondwana 3500+ km to the northeast of its position (relative to Laurussia) in Pangea A (Fig. 1).

⁎ Corresponding author.

E-mail address:[email protected](M. Domeier).

https://doi.org/10.1016/j.gr.2020.11.005

creativecommons.org/licenses/by/4.0/).

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The hypothetical transformation from Pangea B to Pangea A in the early to mid-Permian requires a megashear that accommodated 3500 + km of displacement along a ~6000 km long intracontinental boundary which bisected the supercontinent (Muttoni et al., 2003). Although there are numerous examples of Permo-Carboniferous transtensional and transpressional structures in the core of Pangea that have been attributed to relative motions between Laurussia and Gondwana (Arthaud and Matte, 1977), the summation of estimated displacements along those structures is signiﬁcantly less than that implied by the pur- ported megashear. In fact, the hypothesis of the Pangean megashear was not originally formulated from geological observations, but rather from paleomagnetic data—thereby raising critical questions about the reliability of the latter.

Several studies have demonstrated that the paleomagnetic necessity for Pangea B can be dismissed through paleomagnetic data qualityﬁl- tering exercises, as well as through improvements to the internal reconstruction parameters (Euler poles) used to re-assemble Pangea A (Domeier et al., 2012;Rochette and Vandamme, 2001;Van Der Voo and Torsvik, 2004). Nevertheless, the debate has endured, largely hing- ing on sparse late Carboniferous and early Permian data from Gond- wana, and in particular from parts of northern Italy that have controversially been interpreted to comprise part of a rigid promontory of Africa since at least late Paleozoic time (Muttoni et al., 1996;Muttoni et al., 2003).

Here we present the results of a new paleomagnetic and geochronologic study of latest Carboniferous and early Permian volcanics from

central Morocco. Paleomagnetic results from some of these units have already been reported, and comprise key existing constraints from the early Permian of Gondwana (Daly and Pozzi, 1976;Martin et al., 1978;

Westphal et al., 1979). But given the small sample sizes of the original work, as well as the improvements in laboratory techniques and equip- ment afforded in the decades since those results were published, we considered it important to revisit and re-study some of these localities, in addition to studying others that have not yet been reported on. The acquisition of new geochronological data from these rocks was also an important objective of this work. It has been shown that precise age data are crucial to the question of determining Pangea's paleogeography from paleomagnetism, in part because the entire supercontinent was slowly drifting northward through Permian time (Van Der Voo and Torsvik, 2004). We therefore collected samples for U-Pb geochronology both to accompany our paleomagnetic results, and to re-consider the existing paleomagnetic data in the context of improved age dates.

After presenting our new results, we consider their implications for Pangea's early Permian paleogeography and brieﬂy discuss other observations recently brought to bear on this debate.

2. Geological background and sampling

Structurally, the northern part of Morocco can be divided into three principal domains according to variations in the intensity and style of the late Paleozoic (Variscan) and Cenozoic (Alpine) orogenies (Fig. 2).

In the north, the Rif domain forms the westernmost extension of the Fig. 1.Alternative positions of Gondwana, relative to aﬁxed Laurussia (grey blocks) in a Pangea A (red Gondwana) vs. Pangea B (blue Gondwana) conﬁguration at ~300 Ma. European Variscan terranes have a hatched pattern. The dashed yellow line schematically traces the zone of the megashear that would be required to transform Pangea B to Pangea A.

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Maghrebide belt of northern Africa and comprises a SW-vergent nappe stack constructed during the late Eocene-Neogene. Variscan relics from the Rif domain are known only from allochthonous Paleozoic massifs in its northeastern Internal Zones, which were displaced hundreds of kilo- meters during the Alpine cycle (Chalouan et al., 2008;Leprêtre et al., 2018). In the south, the Anti-Atlas domain is recognized as the deformed, northernmost extent of the West African Craton. It comprises Precambrian basement and Ediacaran-Paleozoic cover rocks that were only weakly deformed during the Variscan, and only locally (to the north of the South Atlas Fault) deformed by the Alpine cycle (Michard et al., 2010b). In between the Rif and the Anti-Atlas, the Meseta domain is characterized by Paleozoic massifs that were strongly shortened during the Variscan, and also subjected then to syn-orogenic magmatism and metamorphism up to amphibolite facies (Chopin et al., 2014;

Lahﬁd et al., 2019;Michard et al., 2010a;Wernert et al., 2016). Those Pa- leozoic massifs host Carboniferous-early Permian continental basins (CPBs) that were developed and deformed in the waning phase of the Variscan cycle, and which are the focus of this study. Signiﬁcant transcurrent motion also occurred between the Meseta and Anti-Atlas domains during the Variscan, along their shared boundary (Houari and Hoepffner, 2003;Mattauer et al., 1972) that has been variously termed the‘South Meseta Fault Zone’or‘Atlas Paleozoic Transform Fault’, among other designations (Michard et al., 2010a). In the Triassic and Jurassic, the Meseta was affected by extension and its boundary zone with the Anti-Atlas evolved into an intracontinental rift (Frizon de Lamotte et al., 2000). Subsequently, during the Cenozoic Alpine orog- eny, the Meseta was again subjected to shortening and dextral wrenching, inverting the Triassic-Jurassic rift and creating the High and Middle Atlas (Frizon de Lamotte et al., 2008;Ellero et al., 2020;

Lanari et al., 2020).

Structurally and stratigraphically, the Meseta domain can be further sub-divided into three main‘blocks’: the Sehoul Block, the East- ern Meseta, and the Western Meseta, as well as a fourth domain, the‘Coastal Block’, which is usually considered part of the Western Meseta (Fig. 2). Variscan deformation appears to have occurred earliest in the Sehoul Block, followed by the Eastern Meseta, and then the Western Meseta and Coastal Block (Hoepffner et al., 2005; Lahﬁd et al., 2019;Michard et al., 2010b). The latter two are distinguished from one another according to the intensity of Variscan deformation

—the Coastal Block being weakly deformed and the rest of the West- ern Meseta being more strongly folded—and they are juxtaposed along the West Meseta Shear Zone (Chopin et al., 2014;Hoepffner et al., 2005;Michard et al., 2010b;Wernert et al., 2016). The CPBs were formed late in the Variscan orogenic cycle, when the regional orogenic edifice was collapsing. Orogenic collapse was largely accommodated by extensional processes, giving rise to a Basin-and-Range- type province that was marked by major low-angle detachment faults, the unroofing of large metamorphic core complexes, and the syn- extensional emplacement of plutonic bodies, dike and sill swarms, and related volcanic expressions. However, under that dominantly extensional regime, Europe and northwest Africa were affected by a complex system of conjugate strike-slip faults that further disrupted the Variscan edifice (Arthaud and Matte, 1975;Arthaud and Matte, 1977). It was under this complex regime that the CPBs formed as transtensional and pull-apart basins (Cailleux et al., 1983; Doblas et al., 1998;Youbi et al., 1995). As an aside, we would like to under- score that the presence of these CPBs and the abundant Permo- Carboniferous strike-slip structures of the region indisputably reflect the operation of important transcurrent tectonism at this time, which has been linked to dextral motion along an intracontinental Fig. 2.Principal tectonic domains of central Morocco (Hoepffner et al., 2005;Michard et al., 2010a;Michard et al., 2010b). Numbers 1–6 denote late Carboniferous-early Permian basins from which samples were collected (1: Tiddas; 2: Bou Achouch, 3: Khenifra, 4: Chougrane, 5: Mechra ben Abbou, 6: Nzalet el Hararcha). Abbreviations: MF, Maghrebide Front; RTFZ, Rabat- Tiflet Fault Zone; SAF, South Atlas Fault; SMF, South Meseta Fault; TTF, Tizi n’Tretten Fault Zone; WMSZ, Western Meseta Shear Zone.

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zone located between Gondwana and Laurussia (Arthaud and Matte, 1975;Arthaud and Matte, 1977); the outstanding question that remains debated is the magnitude of displacement that was accommodated along this shear system.

Although a few CPBs occur in the Eastern Meseta, they are mostly restricted to the Western Meseta, including the Coastal Block. The CPBs mostly occur as NE-SW elongate, rhomb-shaped structures and half- grabens developed along major NE-SW oriented lineaments cutting Pa- leozoic massifs (Cailleux et al., 1983;El Wartiti et al., 1990;Gonord et al., 1980;Saber et al., 2007;Youbi et al., 1995). These transtensive, pull- apart basins developed under a regime of N-S to NE-SW compression that was prevailing in the latest Carboniferous to earliest Permian, but by the late Cisuralian (Artinskian-Kungurian) the stress regime changed to one of E-W to ESE-WNW compression, and later NNW-SSE compression, which inverted the basins, moderately tilting and folding the CPB successions (Chopin et al., 2014;Hoepffner et al., 2005;Saidi et al., 2002;Wernert et al., 2016).

A largely common tectonostratigraphic package among the CPBs begins with latest Carboniferous (late Kasimovian-Gzhelian) terrestrial deposits (although the oldest deposits are generally thought to become progressively younger to the northeast), creating an angular unconformity above older Carboniferous units deformed during the main Variscan phase (Michard et al., 2008). The CPB successions are capped, with angular unconformity, by Late Triassic or younger units, following basin inversion during thefinal phase of Variscan deformation in the late Cisuralian (El Wartiti et al., 1990). Sedimentologically, the CPB successions are characterized by coarse, polymict conglomerates representing debris-flow deposits, and coarse sandstones and red mudstones, locally associated with lacustrine carbonates, or with marsh deposits and paleosol horizons that reflect fluvial and lacustrine environments (El Wartiti et al., 1990). The coarse conglomerates are sometimes restricted to the basin margins, but in other basins may occur intermittently within the sedimentary package, usually forming the base of upward-fining sequences. Thefluvial and lacustrine deposits are generally found in the central parts of the basins, where they reflect the dynamics of wet and dry cycles amidst the background of tectonically-driven basin subsidence (El Wartiti et al., 1990). Many of thefiner-grained units of these terrestrial basins have yielded faunas andfloras that have been used to constrain the age of their deposition, as will be further discussed below.

The CPBs also host volcanic units whose ascent to the surface is thought to have been facilitated by the major transcurrent faults associated with basin formation (Cailleux et al., 1983;El Wartiti et al., 1990).

Contemporaneous latest Carboniferous to early Permian magmatism is also known from the Meseta outside of the CPBs, and is expressed as granitoid plutons emplaced at shallow levels in shear zones (El Hadi et al., 2006;Gasquet et al., 1996;Lagarde et al., 1990;Mrini et al., 1992). Geochemically, both the volcanics within the CPBs and the granitoid plutons outside them generally exhibit calc-alkaline afﬁnities, although they possibly evolve with time toward a more alkaline composition (El Hadi et al., 2006;El Hadi et al., 2014;El Wartiti et al., 1990;Youbi, 1990;Youbi, 1998;Youbi et al., 1995). Compositionally, the volcanics are mostly acidic (rhyolites and rhyolitic breccias), but basalts, andesites and dacites are also observed, as well as volcanic tuffs.

In this study, we collected new paleomagnetic and geochronologic samples from six CPBs in the Western Meseta, namely from the basins of: 1) Bou Achouch, 2) Tiddas-Souk-es-Sebt des Ait Ikkou (hereafter

‘Tiddas’), 3) Khenifra, 4) Chougrane and Souk el Had Bouhsoussène (hereafter ‘Chougrane’), 5) Mechra ben Abbou, and 6) Nzalet el Hararcha. Theﬁrst four basins lie in the Central Massif of the Western Meseta, whereas the latter two lie in the eastern part of the Rehamna Massif (Fig. 2). In the following, we brieﬂy review the geological setting and existing age constraints from each of these basins, and describe our sampling of them.

2.1. Bou Achouch

Bou Achouch is a small (~1 km²) basin that occurs on the northeastern edge of the Central Massif (Figs. 2, S1). The latest Carboniferous- early Permian (C-P) sedimentary package is relatively thin (less than

~60 m thick) and lies unconformably above deformed, Visean (meta-) sedimentary rocks and unconformably below Pliocene deposits. The C- P stratigraphic succession is broken into numerous small fault blocks, and is locally disrupted by generally west-vergent thrusts. According to composite stratigraphic successions constructed byCailleux et al.

(1983),Youbi (1998)andBroutin et al. (1998), the C-P succession begins with conglomerates that include clasts of the underlying Visean rocks, followed by stratified conglomerates andfluvial sandstones alter- nating with horizons of fossil-rich, bluish-grey siltstones and mudstones of lacustrine origin. This initial sequence of conglomerates and siltstones is capped by a layer of felsic tuffs (‘cinerites’) that are overlain and locally incised by coarse conglomerates that include some minor clasts of rhyolite. Up section, the abundance of rhyolitic clasts increases in the conglomerates, andfiner-grained, fossil-rich siltstones and mudstones re-appear. This upper succession is blanketed by a second horizon of felsic tuffs, which are in turn overlain conformably by andesitic lavas that are thought to comprise the top of the C-P succession. On the basis of macroflora assemblages recovered from thefiner grained clastic units and reworked ash deposits,Broutin et al. (1998)estimated the age of the C-P sedimentary package of Bou Achouch to be Kungurian (late Cisuralian). On the other hand,Kerp et al. (2001)noted similarities between the Bou Achouchfloras and those of the Lower Rotliegend of the Saar-Nahe Basin (Germany), the age of which extends back into the latest Carboniferous.

Additional volcanic units of possible C-P age are recognized just beyond the basin margins. ~3 km to the northeast, andesitic lavas are found lying unconformably on the Visean basement, and a series of subvolcanic rhyolitic pipes are observed; the latter have been suggested to represent the vents from which the cinerites in Bou Achouch were erupted (Le Guern et al., 1983;Remmal et al., 1999). ~14 km to the east, andesitic dikes, possibly the source of the andesitic lavas of Bou Achouch, form an‘en-echelon’series intruding the Visean basement, and to the southwest of the basin, dolerite dikes and a small gabbro massif are inferred to be of Permian age (Borredon et al., 1986).

From the Bou Achouch basin, we collected paleomagnetic samples from an andesiteﬂow (BA1) and from a felsic tuff (cinerite; BA2), both of which are (according to the aforementioned composite stratigraphic model) considered to lie in the upper part of the sequence (Fig. S1). Although the sites are adjacent, they reside in different fault- bound blocks and have a slightly different structure. A sample for U- Pb geochronology was also collected from the andesite of BA1 (sample BA1G).

2.2. Tiddas

The Tiddas basin lies along the northern margin of the Central Massif and extends ~17 km along its elongate NE-SW axis, but is only ~2–3 km wide at its widest (Figs. 2, S2). Permian rocks of the basin lie unconformably above deformed, deep-marine siliciclastic rocks of Visean age, and are unconformably capped by Late Triassic and younger clastic sediments (Aassoumi, 1994;Cailleux et al., 1983;Gonord et al., 1980;

Hmich et al., 2006;Larhrib, 1996;Voigt et al., 2011a;Zouine, 1986b).

The Permian succession is characterized by a basal sequence of polymict conglomerates, containing clasts of the underlying (Visean) basement, and sandstones, and is overlain by repeating sequences of conglomerates, ﬂuvial sandstones and lacustrine mudstones in generally upward-ﬁning packages (Broutin et al., 1998). The mudstones include minor intercalations of coal, carbonate and felsic tuff, and bear abundant fossil remains and trace fossils (Voigt et al., 2011a). Collectively, this Permian sedimentary sequence has been variously estimated to

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comprise less than 200 m or up to 500 m (Broutin et al., 1998;Cailleux et al., 1983;Gonord et al., 1980;Voigt et al., 2011a). On the basis ofﬂoras and macrofaunas, and in particular tetrapod ichnofauna, the clastic succession is estimated to be of Artinskian-Kungurian age (Broutin et al., 1998;Hmich et al., 2006;Voigt et al., 2011a).

In the southwest of the basin, near the town of Tiddas, andesiteﬂows are exposed, and in the northeast of the basin (in the Souk-es-Sebt des Ait Ikkou sub-basin), rhyolitic domes and dacitic lavas are observed (Youbi, 1998). The relative stratigraphic positions of these volcanics are not directly observed, but on the basis of structural and stratigraphic studies,Gonord et al. (1980)andCailleux et al. (1983)distinguished three principal volcanic pulses: Pulse I was associated with the eruption of andesites around Tiddas, and is thought to pre-date the Permian sedimentary package. This is supported by the observation that clasts of the andesites are found among the conglomeratic units of the basin, imply- ing that volcanism either pre-dated or was at least contemporaneous with sedimentation (Broutin et al., 1998;Cailleux et al., 1983;Gonord et al., 1980). Volcanic pulse II is associated with the eruption of the pet- rographically distinct dacitic lavas, which were clearly emplaced at the same time as the sediments, as they are interbedded with them. Pulse III is associated with the intrusion of rhyolitic dikes and the‘Ari El Mahsar’rhyolitic dome (in the northeast of the basin), which likely post-date the emplacement of the daciticﬂows. Recently,Ait Lahna et al. (2018) reported a U-Pb SHRIMP age (on zircon) of 286.4 ± 4.7 Ma from the rhyolite dome of Ari El Mahsar. According to the stratigraphic scheme ofGonord et al. (1980)andCailleux et al. (1983), the andesites and dacites should be older than that date.

From the Tiddas basin we collected paleomagnetic samples from three sites (Fig. S2): from an andesiteﬂow (TI1) in the southwest of the basin, near the town of Tiddas, from the rhyolite dome of Ari El Mahsar (SA1), and from a dacite (SA2) of the‘Bled Bou Haouza’section in the northeast of the basin, near the town of Souk Ait Ikko.

2.3. Khenifra

The Khenifra basin lies along the eastern margin of the Central Mas- sif and is one of the largest of the CPBs of central Morocco, with an approximate area of 100 km² (Figs. 2, S3). The Permian succession unconformably overlies deformed Visean and older Paleozoic rocks, and in turn, is unconformably overlain by Late Triassic rocks (Broutin et al., 1998;Cailleux et al., 1983;Youbi, 1990;Youbi, 1998); the entire succession is estimated to be ~1.8 km thick (Cailleux et al., 1983;

Hmich et al., 2006;Voigt et al., 2011b). The clastic succession is characterized by coarse basal conglomerates and sandstones overlain by repeating sequences of generally upward-ﬁning successions comprised dominantly of sandstones and mudstones but also occasional conglomerates, followed at the top by another coarse succession of conglomerates and sandstones. On the basis of fossil assemblages, including tetrapod ichnofauna, the succession has been placed in the Cisuralian, but is thought to be older than the sedimentary successions of the Bou Achouch and Tiddas basins (Broutin et al., 1998;Hmich et al., 2006;

Voigt et al., 2011b).

Permian volcanic rocks are known both to the northwest of the main, clastic-dominated basin, as well as locally in the basin interior.

The volcanic series in the northwest, which has no direct contact with the basinal sediments, comprises rhyolites and rhyodacites reaching up to 150 m thickness, andesites and dacites that overlie and locally in- trude the earlier acidic series, ash-fall tuffs and pyroclastic rocks, and volcanic breccias (Youbi et al., 1995). The minor volcanic series that appears locally within the basin comprises daciteﬂows overlying the sedimentary succession at Jbel Taghat.

Prior age determinations on the volcanic rocks from Khenifra include a K-Ar whole rock age of 264 ± 10 Ma reported from a rhyolitic breccia at Jbel Bou Hayati (Jebrak, 1985). At the same locality, fossil wood was discovered in the volcanic pile that has been assigned to the early Perm- ian (Aassoumi et al., 1995;Youbi, 1990). More recently,Youbi et al.

(2018)reported LA-ICPMS U-Pb (on zircon) ages from several volcanic localities in Khenifra: a dacite from the lower lavaﬂows at Jbel Taghat was dated to 295.1 ± 2.9 Ma and was interpreted to comprise the oldest volcanic pulse of Khenifra. The composite, volcanic dome of Sidi Tir provided ages ranging from 290.3 ± 2.1 Ma (from rhyolite) to 287.9 ± 3.8 Ma (from‘hybrid lavas’, i.e. mixed and mingled andesite and rhyolite lavas), and so the emplacement of the dome was estimated to span

~2 Ma. Rhyolitic pyroclastic fall deposits from Sidi Tir yielded two pre- dominant age clusters at 307.3 ± 2.6 Ma and 290.6 ± 2.6 Ma, but the latter age was reported to be more concordant and taken as the age of eruption. From the upper dacitic lavas at Talat Mechtal,Youbi et al.

(2018)reported an age of 280.3 ± 2.1 Ma, which they interpreted to mark the last volcanic pulse in Khenifra, in agreement with the lithostratigraphic sequence ofYoubi et al. (1995)that inferred Talat Mechtal to post-date Cisuralian sedimentation in Khenifra.

From the Khenifra basin we collected paleomagnetic samples from both volcanic rocks in the basin interior and from the volcanic massif to the northwest (Fig. S3). In the basin interior, we collected two sites from dacites capping Jbel Taghat (TG1–2); the sites are approximately stratigraphically equivalent, but were collected from different sides of the ridge.Martin et al. (1978)reported paleomagnetic results from one site of‘Djebel Tarhat’(i.e. Jbel Taghat) redbeds which came from the sedimentary succession immediately below these dacites.

From the volcanic massif to the northwest, we collected sites from three areas. The northernmost site, TM1, was collected from andesites near the base of the volcanic pile of the Talat Mechtal massif. Site TA1 was collected from an intrusive andesitic neck that is thought to represent one of the earliest volcanic pulses of the Talat Mechtal massif. In the south, at the southeast extremity of Jbel Bou Hayati, four sites were collected from a vertical proﬁle through the volcanic pile, including two from andesiteﬂows at the base (GB3–4), a tuffaceous horizon above them (GB2), and a pyroclastic deposit at the top (GB1). Samples for U- Pb geochronology were taken from both site TA1 and TM1 (samples TA1G and TM1G, respectively).

2.4. Chougrane and Souk el Had Bouhsoussène

The Chougrane basin has thus far been comparatively under- studied, but represents one of the largest CPBs in central Morocco, with an approximate area of 110 km² (Figs. 2, S4). The basin is wedge-shaped and elongated in a NE-SW direction; along its NE-SW axis it spans ~20 km, whereas its NW-SE axis is only ~8 km at its widest (at the southwest end of the basin). To the northeast the C-P deposits pinch out, but the structural axis of the basin continues into a narrow

‘horst’of Ordovician sediments. Theﬂanks of the C-P basin are supported by deformed, Paleozoic basement, composed mostly by Carbon- iferous (meta-) sediments, but locally also Ordovician (meta-) sediments. According toCailleux et al. (1983)andYoubi (1998), the C-P basinal succession begins with basal conglomerates which are covered by andesite lavas and thenﬁner-grained, terrestrial redbeds, followed by a second pulse of andesite lavas that again are covered by redbeds which cap the succession. The age of the sedimentary succession is not well-established, but is broadly estimated to be early Permian on the basis of its strong resemblance to the facies of the Khenifra and Tiddas basins. The andesites of Bir el Gassaa, which overlie the sedimentary succession, have been dated to 270 ± 17 Ma by whole rock K-Ar (Van Houten, 1976).

At Souk el Had Bouhsoussène, ~15 km to the northeast of the Chougrane basin, an isolated succession of sediments and volcanics, similar to that observed in Chougrane and other CPBs of central Morocco, is inferred to be of Permo-Carboniferous age (Fig. S4). The sedimentary and volcanic succession is relatively thin (less than ~80 m thick), and lies unconformably above deformed, Visean (meta-) sedimentary rocks. It begins with polymict conglomerates that include clasts of quartzite and sandstone, followed by a succession of andesite lava ﬂows intercalated with autoclastic volcanic breccias (ﬂow

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breccias). Although the basinal successions of Chougrane and Souk el Had Bouhsoussène are isolated from one another, Cailleux et al.

(1983)andYoubi (1998)proposed, on the basis of lithostratigraphy, a correlative stratigraphic scheme wherein the volcanic series of Souk el Had Bouhsoussène was contemporaneous with the earliest volcanic pulse observed in Chougrane.

From Chougrane, we collected paleomagnetic samples from two andesites in the southwest part of the basin (TC1–2), closely associated with basinal redbeds (Fig. S4). Site TC1 was collected from a succession of lavaflows capping the sedimentary sequence (Bir el Gassaa), whereas the nature of TC2 is less clear; it has been interpreted both as a sill and as a lavaflow.Westphal et al. (1979)published paleomagnetic results from two andesites in the Chougrane basin that may correspond to TC1 and TC2, but a more detailed location for their sites was unfortunately not specified.Daly and Pozzi (1976)also published a paleomagnetic direction from one site sampled from C-P redbeds of this basin. A sample for U-Pb geochronology was also taken from site TC1 (sample TC1G).

From Souk el Had Bouhsoussène we collected a paleomagnetic site and a geochronology sample from an andesiteﬂow (site TZ1 and sample TZ1G, respectively) overlying conglomerates. Paleomagnetic results from andesites at this location (‘Taztot’) were previously published byDaly and Pozzi (1976). Again, according to the proposed stratigraphic scheme of Cailleux et al. (1983)andYoubi (1998), TZ1 should lie near the base of the C-P sequence, beneath TC1 and TC2. As will be discussed later, our new U-Pb geochronological results challenge this stratigraphic scheme.

2.5. Mechra Ben Abbou

Mechra Ben Abbou is located in the north of the Rehamna Massif of central Morocco, ~200 km southwest of Rabat (Figs. 2, S5). It comprises a clastic sequence including alluvial fans,fluvial deposits, andflood- plain and lacustrine deposits that include paleosols and minor carbonates. Locally, the clastic succession is estimated to reach a thickness of up to 3 km. A Cisuralian age for the sedimentary sequence has been estimated on the basis of ostracods found together with vegetal remains in the carbonates (Damotte et al., 1993;Freytet et al., 1999). The stratigraphic architecture of the basin has been divided into two‘mega-sequences’, both of which include volcanic expressions (El Kamel and Muller, 1987;Khounch, 1988;Muller et al., 1991). The volcanic occur- rences correspond to abundantflows of intermediate volcanic rocks that are locally intruded by subvolcanic, felsic dikes and sills and rhyodacites and rhyolite domes (Charif, 2001;El Kamel and Muller, 1987;Khounch, 1988;Muller et al., 1991;Youbi, 1998).

Hadimi et al. (2018)reported U-Pb SHRIMP ages on zircon from a rhyolite dome in the east of Mechra Ben Abbou (‘Bled Mekrach’), which yielded two age clusters of 611 ± 20 Ma and 285.3 ± 4.9 Ma.

The younger age was reported to be more concordant and interpreted as the age of crystallization of the rhyolites, whereas the older age of 611 Ma was attributed to inherited zircons sourced from the Pan- African basement underlying the Rehamna massif (Hadimi et al., 2018). A similar Permian age of 285.4 ± 6.1 Ma (U-Pb on zircon) was reported from a microgranitic/microdioritic dike swarm that intrudes the pre-Permian basement ~25 km south of the basin (Baudin et al., 2003;

Bensalah, 2012;Bensalah et al., 2018;Hadimi et al., 2018).

From this basin we collected paleomagnetic samples from multiple andesiteﬂows (MB1, MB2, MB3) assigned to the second‘mega-sequence’, and the rhyolite dome of Bled Mekrach (MB5) that is ascribed to theﬁrst‘mega-sequence’(Fig. S5). Samples for U-Pb geochronology were taken from site MB2 (sample MB2G) and from an andesite plug closely associated with sites MB1 and MB3 (sample MB4G).

2.6. Nzalet el Hararcha

The Nzalet el Hararcha basin lies on the southeast side of the Rehamna Massif and comprises a fault-bounded area of more than

150 km²(Figs. 2, S6). The basinal succession unconformably overlies older Paleozoic rocks, deformed and metamorphosed during the main phase of the Variscan. The sedimentary sequence begins with conglomerates that include clasts of the underlying basement, and continues upward intoﬁner-grained clastic rocks; together the sedimentary package presents a thickness in excess of 500 m (Saidi et al., 2002). The sediments are sealed by Permian lavas dominated by andesites and rhyolitic domes of calc-alkaline nature (Charif, 2001; Hadimi et al., 2018;

Hoepffner, 1982;Razin et al., 2003;Youbi, 1998). The entire basinal sequence is unconformably covered by Cretaceous and Cenozoic rocks (Razin et al., 2003).

From this basin we collected paleomagnetic samples from multiple andesiteﬂows (NE1, NE2, NE3; Fig. S6). A U-Pb geochronology sample was taken from a nearby rhyolitic dome that we observed the andesites to intermingle with (sample NE0G).

3. U-Pb geochronology methods and results

3.1. Methods

Zircon grains were separated using standard mineral separation techniques at the University of Lisbon, which involved crushing with a jaw crusher followed by sieving to <250μm. Further separation was carried out with use of a Wilfley table and sodium polytungstate heavy liquid, and magnetic grains were removedfirst with a hand magnet, and then using a Frantz magnetic separator. From the non- magnetic and dense mineral separates, zircon was hand-picked under a binocular microscope at the University of Geneva. The zircon minerals were annealed at 900 °C in a muffle furnace for 48 h to repair some of the radiation damage and also create a more homogeneous cathodoluminescence emission, which improves imaging. After anneal- ing, the grains were mounted in epoxy pucks either at the University of Geneva or the Université du Québec à Montréal (UQAM) and imaged on a JEOL JSM7001F Thermal Field Emission SEM with a Schottky electron gun using an accelerating voltage of 15 kV at the University of Geneva, or using a Hitachi S-3400 Variable Pressure SEM at an accelerating voltage of 20 kV at UQAM.

Laser ablation U-Pb dating was performed either at the University of Lausanne, using an element XR single collector mass spectrometer attached to a UP-193FX excimer laser ablation system, or at UQAM, using a Nu Attom single collector mass spectrometer attached to a pho- ton machines G2 193nm excimer laser. The operating parameters in Lausanne were similar to those described inUlianov et al. (2012), and for the Montreal lab, they were similar to those outlined inPerrot et al. (2017). In Lausanne, the data were processed using GJ1 as the primary reference material, and 91500 as the secondary reference material. The age for 91500 was 1064.32 ± 1.9 (2σ) Ma, which is in good agreement with the known age (Wiedenbeck et al., 2004). For the Mon- treal lab, the 91500 zircon was used as the primary reference material and GJ1 was used as the secondary reference giving an age of 595.79 ± 0.68 Ma, which is somewhat young, but in agreement with the TIMS data from this zircon (Jackson et al., 2004). All data were processed using iolite, and the ages were calculated and plotted using IsoplotR (Vermeesch, 2018).

3.2. Results by basin

In the following sections, the results of the geochronological analyses are presented by basin. The principal results are presented in abbre- viated form inTable 1(together with other relevant geochronological data from the literature), and displayed inFigs. 3 and 4; the complete data is available in Table S1. All ages are quoted with 2σuncertainties.

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3.2.1. Bou Achouch

Sample BA1G (andesite) contains zircon that have ages which span from >2Ga to ~300 Ma, indicating a distinct inherited component to the zircon population in this sample (Table S1). Most of the zircon (~60%) have ages which cluster around ~300 Ma, and the youngest part of this cluster gives a weighted mean²⁰⁶Pb/²³⁸U age of 301.50 ± 1.24 (n= 7, MSWD = 0.87;Fig. 3a). Five younger grains are considered to have been affected by Pb loss, since none of the ages overlap with each other. We therefore take the weighted mean age as the best estimate of the eruption age of the andesite. If this andesite indeed represents the top of the stratigraphic sequence at Bou Achouch (as the stratigraphic scheme ofCailleux et al. (1983)suggests), this latest Car- boniferous age would indicate that theﬂoras studied byBroutin et al.

(1998)are older than Kungurian (~283–273 Ma), like similarﬂoras found in the Lower Rotliegened (latest Carboniferous to early Permian) sediments of the Saar-Nahe Basin in Germany (Kerp et al., 2001). On the other hand, it is important to note that the stratigraphic scheme of Cailleux et al. (1983)is composite, as the basin stratigraphy is disrupted by thrusts. The possibility that the dated andesite (BA1G) actually lies stratigraphically below the fossil-bearing units therefore cannot be excluded.

3.2.2. Khenifra

Sample TA1G (andesite neck from Talat Mechtal) only yielded four zircon grains after mineral separation, of which, two were concordant and two were discordant with a suspected common Pb contribution (Fig. 3b). The two concordant grains yielded overlapping ages, giving a weighted mean²⁰⁶Pb/²³⁸U age of 282.76 ± 2.19 Ma (MSWD = 0.29).

The discordant grains yielded very similar²⁰⁶Pb/²³⁸U ages, but were excluded from the weighted mean calculation. The very sparse number of

analyses indicates that this weighted mean age should be treated with caution. With that in mind, we note that although this weighted mean age is indistinguishable from that of the youngest volcanic pulse in Talat Mechtal (280.3 ± 2.1 Ma), the lithostratigraphic scheme of Youbi et al. (1995)rather ascribes the TA1G andesite to one of the ﬁrst magmatic pulses of the Talat Mechtal massif (possibly correlative with TM1G; see below).

Sample TM1G (andesite lava from Talat Mechtal) only yielded three concordant ages around the expected age of eruption (out of 26 grains analyzed; Table S1). Two of these analyses have²⁰⁶Pb/²³⁸U ages that overlap within uncertainty, and these analyses also overlap with another slightly discordant analysis producing a²⁰⁶Pb/²³⁸U weighted mean age of 305.59 ± 2.68 Ma (MSWD = 0.89;Fig. 3c). The third concordant grain produced an age of 276.3 ± 8.1 Ma which we interpret as affected by Pb loss. Again, the paucity of zircon grains around the expected age means that the age interpretations should be treated with caution. The TM1G andesite lies at the base of the Talat Mechtal volcanic succession, and is overlain by pyroclastic fall deposits that have been dated to 290.6 ± 2.6 Ma (Youbi et al., 2018), so the weighted-mean age is not in violation of the existing stratigraphic-geochronologic constraints. Nevertheless, this age is distinctly older than anticipated, and would imply that volcanism at Talat Mechtal began ~15 Ma earlier than proposed byYoubi et al. (2018). It is worth noting that sample TM1G contains multiple zircon age populations that span back beyond 2 Ga (Table S1), and so inherited zircon is evidently abundant in these andesites. In the overlying ~290 Ma pyroclastic deposits,Youbi et al.

(2018)also reported a cluster of latest Carboniferous ages (307.3 ± 2.6 Ma), attributed to inherited zircon grains, which could have origi- nated from the TM1G andesite, if both populations were not otherwise inherited from some other common source.

Table 1

Overview of new and published geochronological data from Carboniferous-early Permian basins of central Morocco.

Basin/label Site/unit name Lat

(°N) Lon (°W)

Lithology System Best-est. age (Ma) Estimate type Reference

Bou Achouch

BA1G Bou Achouch lavas 33.673 5.743 andesite U-Pb on zircon 301.50 ± 1.24 weighted mean This work

Tiddas

G1 Ari el Mahsar dome 33.664 6.067 rhyolite U-Pb on zircon 286.4 ± 4.7 concordia Ait Lahna et al.

(2018) Khenifra

G2 Jbel Bou Hayati 32.981 5.701 rhyolitic

breccia

K-Ar whole rock

264 ± 10 whole rock Jebrak (1985)

G3 upper Talat Mechtal lavas 33.029 5.705 dacite U-Pb on zircon 280.3 ± 2.1 concordia Youbi et al. (2018)

TA1G Takarout intrusive neck 32.999 5.685 andesite U-Pb on zircon 282.76 ± 2.19 weighted mean This work

G4 Sidi Tiri lavas 33.035 5.712 hybrid lavas U-Pb on zircon 287.9 ± 3.8 concordia Youbi et al. (2018)

G5 Sidi Tiri lavas 33.036 5.713 rhyolite U-Pb on zircon 290.3 ± 2.1 concordia Youbi et al. (2018)

G6 Sidi Tiri pyroclastic fall 33.031 5.708 pyroclastic fall U-Pb on zircon 290.6 ± 2.6 concordia Youbi et al. (2018)

G7 lower Taghat lavas 32.995 5.643 dacite U-Pb on zircon 295.1 ± 2.9 concordia Youbi et al. (2018)

TM1G Talat Mechtal lavas 33.025 5.664 andesite U-Pb on zircon 305.59 ± 2.68 weighted mean This work

Chougrane

G8 Bir el Gassaa 32.926 6.315 andesite K-Ar whole

rock

270 ± 17 whole rock Van Houten (1976)

TZ1G Taztot lavas 33.148 6.092 andesite U-Pb on zircon 295.6 ± 2.9 to 267.9

± 3.9

range of single Zr ages

This work TC1G Bir el Gassaa 32.926 6.315 andesite U-Pb on zircon 307.8 ± 4 to 295.5 ± 3.4 range of single Zr

ages

This work

Mechra Ben Abbou

MB2G Douar Ouled Said Ben Ali lavas 32.660 7.812 andesite U-Pb on zircon 284.2 ± 4.6 concordia This work

G9 Bled Mekrach dome 32.639 7.751 rhyolite U-Pb on zircon 285.3 ± 4.9 concordia Hadimi et al. (2018)

G10 Dikes intruding Paleoz.

basement

32.362 7.924 microgranite U-Pb on zircon 285.4 ± 6.1 concordia Razin et al. (2003) MB4G Douar Ouled Said Ben Ali lavas 32.657 7.811 andesite U-Pb on zircon 294.63 ± 0.67 weighted mean This work Nzalet el Hararcha

NE0G Sidi Bou Yahia dome 32.266 7.651 rhyolite U-Pb on zircon 277.07 ± 0.61 weighted mean This work

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3.2.3. Chougrane

Samples TC1G (andesite from Bir el Gassaa) and TZ1G (andesite from Souk el Had Bouhsoussène) yielded zircon with potential inheritance and Pb loss effects, which make determining the eruption age

for these samples difﬁcult. Sample TC1G only yielded three concordant zircon grains with ages around the expected eruption age. The three ages span a range from 307.8 ± 4 to 295.5 ± 3.4 Ma, rather than forming a cluster (Fig. 3d). Although this range of ages is reasonable,

Fig. 3.Concordia diagrams and²⁰⁶Pb/²³⁸U weighted mean ages for samples dated by this work. Weighted mean ages were determined when multiple concordant zircon analyses produced the same age. Yellow data points highlight the analyses used to calculate the reported age. Dotted arrows indicate analyses affected by small amounts of common Pb. See the main text for interpretations of the reported ages.

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as TC1G lies at the top of the C-P stratigraphic sequence of Chougrane, it would imply that the C-P sediments below are of earliest Permian or late Carboniferous age.

According to existing stratigraphic models of the Chougrane CPB, the TC1G andesite is considered to lie stratigraphically above the TZ1G andesite of the Souk el Had Bouhsoussène sub-basin, and is thus expected to be younger. However, most of the TZ1G zircon grains (80%) have

206Pb/²³⁸U ages which span the early Permian from 295.6 ± 2.9 to 267.9 ± 3.9 Ma (Fig. 3e). The only U-Pb age interpretation that would be in agreement with the existing stratigraphic model would require interpreting the youngest age from TC1G (295.5 ± 3.4 Ma) and the oldest age (of the Permian grains) from TZ1G (295.6 ± 2.9 Ma) as the eruption ages. However, this interpretation is not parsimonious in that it treats the data from the two samples in different ways (i.e. selecting the oldest grain age from one sample and the youngest from the other). Thus, while these geochronological results are clearly inade- quate to draw strong conclusions from, they provisionally suggest that the Souk el Had Bouhsoussène sub-basin is younger than Chougrane, and therefore evolved independently of it, in contrast to the aforementioned stratigraphic models. Because Souk el Had Bouhsoussène was also tilted during thefinal phase of the Variscan, the youngest age of TZ1G (if geologically meaningful) could also place a constraint on this final phase of Variscan tectonism. We therefore emphasize that further geochronological work is needed to affirm both the provisional geochronological data presented here and the stratigraphic revisions they imply.

3.2.4. Mechra Ben Abbou

Samples MB2G (andesiteﬂow) and MB4G (andesite plug) provide

206Pb/²³⁸U ages of 284.2 ± 4.6 and 294.63 ± 0.67 Ma (n= 15, MSWD = 2.24), respectively (all based on concordant analyses;Fig. 3f,g). Notably, the result from MB2G is based on a single analysis, but it is in good agreement with the existing U-Pb geochronological constraints recently published from the volcanics of this basin (~285 Ma), as described above (Baudin et al., 2003;Hadimi et al., 2018). However, the more robust result from MB4G would suggest that volcanism could have started in this basin at least 10 Ma earlier, although this sample also bears evidence of inheritance (Fig. 3g).

3.2.5. Nzalet el Hararcha

Sample NE0G (rhyolite) provided a²⁰⁶Pb/²³⁸U weighted mean age of 277.07 ± 0.61 Ma (n= 9, MSWD = 2.33;Fig. 3h). This robust result represents theﬁrst absolute age constraint from this basin.

4. Rock magnetic methods and results

4.1. Methods

Rock magnetic experiments included the measurement of hysteresis, isothermal remnant magnetization (IRM) acquisition, and magnetic susceptibility versus temperature (χvs. T). Hysteresis loops and IRM acquisition curves were generated with use of a Lake Shore PMC MicroMag 3900 Vibrating Sample Magnetometer (VSM), andχvs. T curves were measured in air with an AGICO MFK1-FA Kappabridge with CS-4 Furnace, all housed at the Ivar Giæver Geomagnetic Labora- tory at the University of Oslo. Additional IRM acquisition curves were measured with use of an ASC Scientiﬁc pulse magnetizer and Molspin magnetometer at the University of Coimbra. Decomposition of the IRM curves to identify distinct magnetic components was performed with MAX UnMix (Maxbauer et al., 2016).

4.2. Results

Because the rock magnetic characteristics of the studied volcanics were very similar across the basins sampled, we describe their magnetic mineralogy collectively, but make references to speciﬁc behaviors and magnetic properties of individual sites when presenting the paleomagnetic results in the following section (Section 5).

Hysteresis measurements of the volcanic rocks revealed a spectrum of behaviors ranging from narrow, relatively sharp-shouldered loops dominated mostly by the ferromagnetism of a single, low-coercivity component (Fig. 5a) to more open loops with a relatively stronger para- magnetic signal and a ferromagnetic component with a higher coercivity (Fig. 5c). However, most samples exhibited an intermediate type of behavior associated with‘wasp-waisted’loops (Fig. 5b) that result from a combination of the low- and high-coercivity components in var- ious proportions. Such‘wasp-waisted’loops can be generated by combinations of single domain (SD) and superparamagnetic (SP) populations of magnetite (or low-titanium titanomagnetite), but owing to the coercivities observed in these volcanic samples, we ascribe the behavior to variable combinations of magnetite and hematite.

IRM acquisition curves (Fig. 5d–f) and the decomposition of their coercivity spectra (Fig. 5g–i) further demonstrate the occurrence of two remanence-bearing components with distinct coercivities in the volcanic rocks. Samples associated with narrow, sharp-shouldered hysteresis loops exhibit a likewise steep and sharp IRM acquisition curve (Fig. 5d) and exhibit a coercivity spectra dominated by a component with an average coercivity of ~30 mT (Fig. 5g), consistent with magnetite or maghemite. Samples associated with open hysteresis loops correspond- ingly yield more shallowly sloped IRM acquisition curves (Fig. 5f) and reveal high-coercivity components with average coercivities of

~0.5–2 T (Fig. 5i). Most of these high coercivity components are likely associated with hematite, although the very high coervicities could also indicate the presence of some other phase. The majority of samples exhibit some intermediate behavior that includes both low- and high- coercivity components (Fig. 5e,h).

Theχvs. T experiments further corroborate the presence of magnetite and hematite in variable proportions in the volcanic samples. In samples associated with a dominant low-coercivity component ascribed to magnetite,χvs. T heating curves exhibit a prominent drop in susceptibility between ~500 and 580 °C (Fig. 5j), consistent with the Curie temperature of magnetite. Those bearing evidence of a high- coercivity component show a drop in susceptibility above 600 °C and below 680 °C (Fig. 5l), consistent with hematite. Most samples exhibit indications of both minerals, and some χvs. T curves also show Fig. 4.Graphical summary of new and existing radiometric ages from volcanic rocks of the

studied basins. The vertical bars show absolute age estimates and their uncertainties from Table 1(blue = U-Pb weighted-mean or concordia ages; purple = range of single-grain U- Pb zircon ages; grey = K-Ar whole rock ages). Redﬁlled (open) circles denote directly- dated (estimated) ages of corresponding paleomagnetic sites. The pink band shows the age range of the early Permian, and the red-dashed line shows the approximate mean age of the early Permian paleomagnetic sites (~285 Ma).

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inﬂections in susceptibility during heating at temperatures between 300 and 400 °C, which could reﬂect the occurrence of maghemite (Fig. 5k). The disappearance of this phase in the cooling curves is consistent with the breakdown of maghemite with heating. In these samples, the low-coercivity components observed in the hysteresis and IRM experiments could thus be due to maghemite or a combination of maghemite and magnetite.

5. Paleomagnetic and magnetic fabric methods and results

5.1. Methods

Paleomagnetic samples were collected during twofield seasons and measured independently at the University of Michigan (‘UM’) and at the Ivar Giæver Geomagnetic Laboratory at the University of Oslo (‘UO’). At Fig. 5.Rock magnetic experiments: a–c) hysteresis loops before (blue) and after (red) correction for paramagnetism, d–f) isothermal remanent acquisition (IRM) curves, g–i) coercivity spectra (first-derivative of IRM curves above) withfitted components, and j–l) magnetic susceptibility (χ) vs. temperature curves (red = heating, blue = cooling, grey =first derivative of heating curve).

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the UM lab, specimens were stepwise demagnetized with either alter- natingfields (AF), using a Sapphire Instruments SI4 AF demagnetizer, or thermally, with an ASC model TD48 thermal demagnetizer. Between each demagnetization step the magnetization of the specimens was measured with a 2G three-axis superconducting magnetometer. For the UO collection, the anisotropy of magnetic susceptibility (AMS) was measured on a select set of specimens before they were demagnetized, in an effort to supplementfield measurements of the structural orientation of the volcanics (paleo-horizontal), which were sometimes difficult to ascertain on the basis of poor outcrop conditions, massive units or complexflow structures. AMS measurements were performed with an AGICO MFK1-FA Kappabridge. Specimens were then stepwise demagnetized with either AF, using an AGICO LDA-5 AF demagnetizer, or thermally, using a Magnetic Measurements TD80A thermal demagnetizer. After each demagnetization step the magnetization of the specimens was measured with either an AGICO JR-6 spinner magnetometer or a WSGI (2G) model 755 superconducting magnetometer.

Processing of the AMS data was performed with Anisoft (AGICO), and analysis of the paleomagnetic data was performed with PaleoMac (Cogné, 2003) and PmagPy (Tauxe et al., 2016), and included both the deﬁnition of linear components by principal-component analysis (Kirschvink, 1980) and theﬁtting of great circles by the approach of McFadden and McElhinny (1988).

The principal results of the AMS measurements are listed in Table S2 and displayed inFig. 6; summary results of the paleomagnetic measurements are provided inTable 2 and Table S3, and are displayed in Figs. 7–9. For reference, site-level magnetization directions reported by earlier studies are also plotted inFig. 8. In the following, we describe our new results by basin.

5.2. Results by basin 5.2.1. Bou Achouch

AMS results from BA1 (andesite) appear scattered, but the mean Kmaxdirection is close to the pole to the paleo-horizontal inferred fromfield measurements (Fig. 6a). This could be an inverse magnetic fabric produced by SD magnetite grains with some preferred alignment due toflow during emplacement, which is consistent with the hysteresis behavior observed in samples from this site. In BA2 (felsic tuff), the magnetic fabric is better defined, and exhibits aKmindirection that is sub-parallel to the measured (and well-determined) pole to the paleo-horizontal (Fig. 6b). Thus, the AMS data from BA2 (and possibly BA1) appear to corroborate our structuralfield measurements.

AF demagnetization of pilot samples from BA1 and BA2 was ineffective. Thermal demagnetization of BA1 samples was successful, and yielded entirely univectorial behavior deﬁning a single characteristic remnant component (ChRM) that we label component‘A’(Fig. 7a).

Most remanence loss occurred between temperatures of 300 to 590 °C, although in some samples a signiﬁcant proportion (>20%) of the natural remanent magnetization (NRM) survived beyond 600 °C, without an observed change in direction. This suggests that the magnetic remanence is held by both magnetite and hematite, both of which are present according to the rock magnetic experiments. Be- tween samples, component A is well-clustered and directed shallowly downward toward the south; because this direction is sub-parallel to the strike of the unit, structural correction does not signiﬁcantly modify the mean direction (Table 2;Fig. 8a).

Thermal demagnetization of site BA2 was unsuccessful. Upon heating to 100 °C, >90% of the NRM is lost and the samples become directionally unstable (erratic); no samples yielded interpretable results. Results from this site were therefore discarded and will not be discussed further.

5.2.2. Tiddas

From Tiddas, AMS measurements were collected from site SA1 (rhyolite dome), where the structure in thefield was ambiguous. The fabric appears‘normal’in the sense thatKminis clustered andKmaxandKintare distributed in a girdle, but the orientation ofKminis subhorizontal and approximately parallel to the axis of large prismatic columns measured in thefield (Fig. 6c). Given the large size of the dome from which SA1 was collected, and the fact that Permian rocks elsewhere among the CPBs are generally only mildly to moderately deformed and tilted, it seems unlikely that the orientation of the cooling columns and corresponding magnetic fabric reflects strong tilting of the dome, but rather sub-vertical intrusive emplacement.

Pilot samples from SA1 showed AF demagnetization to be relatively ineffective (30–70% of NRM remaining with 200 mT treatment), and thermal demagnetization was more successful. In most samples, a low temperature component (LTC) that we label component‘B' was removed up to 300 °C, followed by a univectorial decay up to 670 °C (ChRM), but often with signiﬁcant remanence loss below 575 °C (Fig. 7b). The mean direction of the LTC is oriented north and moderately down (Fig. 9), sub-parallel to the present-dayﬁeld (PDF), whereas the ChRM is oriented shallowly down to the SE (Figs. 8b). Because the structural orientation of SA1 is not known, a structural restoration of these directions is not possible.

Samples from TI1 (andesite) exhibit univectorial behavior with generally progressive unblocking between 200 and 680 °C, but sometimes with a sharper unblocking above 600 °C. In a few samples, a distinct LTC can be identiﬁed below ~400 °C, but it is generally poorly deﬁned and not directionally consistent, and so no site mean direction for it could be determined. The site-mean of the ChRM is directed shallowly upward to the SE in both geographic and tilt-corrected coordinates (Fig. 8b).

Samples from SA2 (dacite) behaved similarly to those of TI1, with remanence unblocking occurring progressively between ~200 and 680 °C. Some samples behaved univectorially, with no identiﬁable directional changes during treatment (Fig. 7c), whereas others exhibited a distinct LTC that was usually removed below ~400 °C. In geographic coordinates, the mean LTC is directed moderately down and to the north (Fig. 9), whereas the ChRM is oriented horizontally toward the SSE in both geographic and tilt-corrected coordinates (Fig. 8b).

5.2.3. Khenifra

AMS data from site TM1 (andesiteflow) are well-defined by a‘normal’oblate fabric withKminparallel to the pole to the paleo-horizontal (Fig. 6e). Results from site TA1 (andesitic intrusive neck) are also characterized by an oblate fabric, but with a highly inclinedKmin axis (Fig. 6f). The pole to the paleo-horizontal for this site, determined from the host of the andesitic neck, is close to the plane defined by KintandKmax, and the orientation of this plane could reflect sub- verticalflow associated with emplacement of the neck. Results from site TG1 (daciteflow) also exhibit a normal fabric, with a sub-vertical Kmin, sub-parallel to the pole to paleo-horizontal inferred fromfield measurements (Fig. 6g). In contrast, site TG2 (daciteflow), which has approximately the same structural attitude as TG1, yields a fabric with a sub-horizontalKmin(Fig. 6h). However,Kintis sub-parallel to the pole to bedding, and so it is possible that this fabric results from a mixed (intermediate) fabric formed by the superposition of a normal fabric and an inverse fabric imposed by a population of SD grains. Al- though we have not attempted to verify this through isolation of the postulated mixed fabrics, the structural attitude of TG2 is sufficiently clear in thefield that the possibility of it having a near-vertical attitude can be rejected.

With demagnetization, samples from both TM1 and TA1 exhibit two-component behavior. In TM1, a LTC was removed by ~250 °C, followed by univectorial decay to the origin by 670 °C, although 80%

of the NRM was usually removed by 590 °C (Fig. 7d). In TA1, the low

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stability component was removed by ~15–20 mT or ~300–400 °C, followed by univectorial decay to the origin by 200 mT or

~570–590 °C (Fig. 7e). The LTC from TA1 is badly scattered at the sample level, whereas in TM1 it is oriented northwards and moderately down in geographic coordinates, sub-parallel to the PDF (Fig. 9). The site- level ChRMs from both sites are oriented shallowly upwards toward the southeast (Fig. 8c).

In TG1 and TG2, both AF and thermal demagnetization reveal char- acteristically univectorial behavior, with ~95% of the NRM usually being removed by ~70 mT or ~550 °C (Fig. 7f). In a few samples a weak LTC is removed in theﬁrst demagnetization steps (~200 °C), but

is randomly oriented at the sample level. The mean ChRM direction of both sites is oriented horizontally to the southeast in geographic coordinates (Fig. 8c).

Thermal demagnetization of GB1 (rhyolitic ignimbrite) and GB4 (andesiteflow) exhibited similar behavior, characterized by linear decay up to 600 °C, which we define as the ChRM. In some samples, decay above 600 °C continues parallel to the direction of the ChRM defined below 600 °C, but in other cases the remnant direction changes above 600 °C. In the latter examples, the trajectory of this high temperature directional change is not consistent between samples, and great- circles fitted to the high-temperature segments of the individual Fig. 6.Anisotropy of magnetic susceptibility (AMS) results by site. All symbols are projections on the lower hemisphere. Orange arcs show the‘bedding’plane (inferred paleo-horizontal) as measured in thefield; stars mark the pole to the‘bedding’plane with the measurement uncertainty. For site SA1, the diamond (orange arc) marks the axis of (plane normal to) columnar jointing. The light (dark) blue arcs show the plane normal to theKminaxis of each measured specimen (the site-mean).

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Fig. 7.Examples of demagnetization behavior from samples of each basin. Filled (open) circles depict projections onto the horizontal (vertical) plane. Red (blue) arrows show the interpreted high (low) stability components. All results are shown in geographic coordinates.