# Revisiting the geodynamical history of the so-called uplifted atolls, north-west Tuamotu, French Polynesia, central South Pacific

Edwige Pons-Branchu, Lionel Marié, Lucien F. Montaggioni, Bertrand Martin-Garin, Bernard Salvat

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## To cite this version:

Edwige Pons-Branchu, Lionel Marié, Lucien F. Montaggioni, Bertrand Martin-Garin, Bernard Salvat. Revisiting the geodynamical history of the so-called uplifted atolls, north-west Tuamotu, French Polynesia, central South Pacific. Earth-Science Reviews, 2023, 244, pp.104532. 10.1016/j.earscirev.2023.104532. hal-04180415

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1 Revisiting the geodynamical history of the so-

2 called *uplifted* atolls, north-west Tuamotu,

3 French Polynesia, central South Pacific

4 Lucien F. Montaggioni a*, Edwige Pons-Branchu b, François

5 Fournier a, Bertrand Martin-Garin a, Arnaud Dapoigny b,

6 Lionel Marié c and Bernard Salvat d

7 a Aix Marseille Univ, CNRS, IRD, INRAE, CEREGE, 13331

8 Marseille, France

9 b LSCE-IPSL, CEA, CNRS, UVSQ, Université Paris-Saclay,

10 91198 Gif-sur-Yvette, France

11 c Aix Marseille Univ, CNRS, IRD, INRAE, CEREGE, 13545

12 Aix-en-Provence, France

13 d PSL-École pratique des hautes études, UAR 3278, EPHE,

14 CNRS, UPVD, CRIOBE, Labex Corail, université de

15 Perpignan, 66860 Perpignan, France

16 Highlights

17 – Uplift occurred sequentially from atoll to atoll, over the
18 last 2 Myr, along a north-west to south-east axis,
19 according with the drift direction of the Pacific plate.

20 – Hotspot thermal rejuvenation as resulting in an
21 asthenospheric bump is considered to be the controlling
22 factor of uplift.

23 – In the northwesternmost atolls, uplift was followed by
24 atoll stabilization, then subsidence resumption.
25 Abstract

26 In the Tuamotu Archipelago, central South Pacific, some north-
27 western atolls – e.g. Makatea, Mataiva, Tikehau, Rangiroa,
28 Kaukura, Niau, and Anaa – exhibit raised reef limestones,
29 locally at unusual elevations above present sea level. These
30 islands were formerly interpreted to have been primarily
31 uplifted by a lithospheric flexure generated by the load of
32 nearby Society high volcanic edifices (Tahiti, Moorea). A
33 reappraisal of previous field data and additional uranium-
34 series, ESR and strontium isotope dating of the exposed reef
35 limestones reveal the flexure-induced load model failed to
36 account for the tectonic evolutionary history of the 7
37 uplifted north-west Tuamotu atolls, mentioned above. A
38 new geodynamical scenario is proposed, implying the
39 episodic thermal rejuvenation of the Society hotspot swell,
40 thus generating an asthenospheric bump. At Makatea,
41 Mataiva, Tikehau, Rangiroa and Kaukura, the basements
42 and reef-rims were known to be Miocene in age. At
43 Makatea, the closest island to the bump centre, the upper
44 plateau peaks at present at 113 m above present mean sea
45 level (pmsl), while on the other atolls, in peripheral
46 positions, remnants of reef limestones are at 2.5 to 12 m
47 above pmsl. At Niau, the atoll reef-rim, made up of a
48 Pliocene core, locally overlapped by Late Pleistocene reef
49 units, peaks at +7.5 m. At Anaa, the reef-rim is strictly
50 composed of Late Pleistocene reef deposits at elevations of
51 about +4–5 m. Makatea, Mataiva, Tikehau, Rangiroa and
52 Kaukura, resting on the Pacific plate, have shifted
53 northwestwards and as such entered the influence zone of
54 the hotspot swell, approximately two million years (Ma)
55 ago. The hotspot reheating phase that generated Moorea and
56 Tahiti edifices, resulted in the uplift of Makatea – at
57 estimated long-term rates of about 0.16 to 0.20 mm/yr – and
of the nearer atolls. By contrast, the most south-eastern ones (Niau and Anaa), at that time, located further south-east from the hotspot swell, are assumed to have continued subsiding according to expected standard crust cooling conditions. From approximately 0.5 Ma to 0.3 Ma, due to their continuous motion northwestwards, the most north-western islands ended up escaping the hotspot influence and started to re-subside. At the same time, Niau and Anaa Atolls, as subjected increasingly to the hotspot reheating effects, presumably in relation to the activity of sub-aerial Tahiti-Iti and Mehetia volcanoes, stopped subsiding and presumably were slightly uplifted prior to being stabilized. This is supported by the occurrence of Last Interglacial high-stand reefs at present-day elevations close to their presumed eustatic position. The uplift-stabilization-subsidence resumption scenario is consistent with a northwestward motion of islands carried away on a treadmill, i.e. the drifting Pacific plate, running over a rejuvenated mantle plume.

## Keywords

Atolls

Uplifting

Subsidence

Thermal hotspot rejuvenation

Tuamotu

Central South Pacific

## 1. Introduction

The evolution of oceanic volcanic islands in midplate settings is known to generally involve long-lasting subsidence, in the absence of geodynamic constraints to ultimately form atolls (Darwin, 1842; Droxler and Jorry, 2021) and, if moving out
of reef-building latitudes, finally guyots (Grigg, 1982).
Although the actual drowning process of oceanic volcanoes
still remains not completely understood (Huppert et al.,
2020), this trend is classically regarded as driven by various
mechanisms: thermal cooling of the underlying lithospheric
plate (Parsons and Sclater, 1977; Marty and Cazenave,
1989; Stein and Stein, 1992); lithospheric flexure due to
volcanic loading (Walcott, 1970; Watts and Cochran, 1974;
Lambeck, 1981a,b; Menard, 1983; Zhong and Watts, 2013)
or hotspot swell decay (Morgan et al., 1995). The end-
member control would be the displacement of volcanoes
away from past mantle melting sources by rapid plate
motion (Huppert et al., 2020).

However, in the tropics, a long-standing question in midplate
volcano geodynamics is that of the origin of uplifted ocean
islands as evidenced by the presence of raised reefal
limestones at unusual elevations, and not interpreted to have
resulted from deposition during eustatic high sea levels.
These islands were used to test the response of the oceanic
crust and underlying mantle to thermal and surface load
constraints generated by relatively young (less than 2–4
Myr-old) volcanoes (McNutt and Menard, 1978; Lambeck,
1981 a, b; Huppert et al., 2015; Ramalho et al., 2017 and
references herein). Such islands are encountered in a number
of archipelagoes in the Pacific Ocean: Hawaii (Grigg and
Jones, 1997), Cook (Stoddart et al., 1985, 1990; Spencer et
al., 1987; Woodroffe et al., 1991; Hein et al., 1997; Gischler
et al., 2022), Austral (Pirazzoli and Veeh, 1987;
Montaggioni, 2015), Pitcairn (Spencer, 1989; Blake, 1995;
Blake and Pandolfi, 1997) and Tuamotu (Gabrié and Salvat,
1985). In addition, high carbonate islands in the same areas
are frequently composed of emerged coral-reef limestone
piles and show associated geomorphic features at various
elevations – few meters to up to +100 m. Such island types
include Bermuda (Hearty and Kindler, 1995; Muhs et al., 2002) in the Atlantic, Glorioso (Guillaume et al., 2013) and the Aldabra Group (Braithwaite, 2020) in the western Indian Ocean, Nauru (Jacobson et al., 1997), Banaba (Taylor, 2005 in Teaiwa, 2015), and Makatea (Montaggioni, 1985, 1989; Montaggioni and Camoin, 1997) in the south-central Pacific. It is noteworthy that there are other uplifted, reef-capped islands related to fore-arc bulges (e.g. Loyalty Islands, Tonga and Niue) in the Pacific.

In the Tuamotu Archipelago (French Polynesia), it has been long recognized that some islands of the north-west group have been affected by significant uplift (Fig. 1). This is supported by the presence of raised reef tracts, marine terraces and/or tidal notches at unusual elevations (Ranson, 1953, 1957; Obellianne, 1963; Montaggioni, 1985). The so-called *uplifted atolls* form a continuous belt along the north-western side of the Tuamotu Archipelago. Oriented north-west to south-east, between latitudes 14°55' and 17°35' south and longitudes 148°46' and 145°20' west, this belt is currently just facing the Society Archipelago, at a distance of approximately 250 to up to 300 km from the eastern coasts of Mehetia, Tahiti, and Moorea islands. This belt is composed, from north-west to south-east, of the islands of Mataiva, Tikehau, Rangiroa, Makatea, Kaukura, Niau, and Anaa. All these islands exhibit remnants of raised reefal limestones – *feo* in Polynesian language –, locally associated with other marine geomorphic features – tidal notches, abraded terraces, beaches –, at elevations of about 1 m to up to 110 m above present mean sea level (pmsl). Different geophysical scenarios, expected to be able to reproduce most of field observations, have been proposed to explain the present-day elevation of the emerged reef limestones at the regional scale. Emergence would result from sea-floor
uplift in relation to lithospheric flexuring generated by the
volcanic Tahiti-Moorea load (McNutt and Menard, 1978;
Lambeck, 1981 a; Spencer et al., 1987) or to a combination
between load flexuring and asthenospheric bump due to
reheating of the Society swell (Montaggioni, 1989;
Montaggioni and Camoin, 1997) according to the models
proposed by Crough (1978, 1983, 1984) and Menard (1983).

The recent revival of interest in vertical motion mechanisms of
intraplate oceanic volcanic islands (Ramalho et al., 2010
a,b,c, 2015, 2017; Klein, 2016; Marques et al., 2020)
prompted us to review the issues of the reference uplift
models applied in the north-west Tuamotu islands. Their
geodynamical evolutionary history thus is revisited from a
fresh perspective with respect to the timing and magnitude
of successive vertical motions, by presenting arguments
against certain previous reference models and by placing
constraints on other plausible ones. Reassessment of local
field observations and data and inclusion of additional
chronological results bring new insights on the timing of
uplift and subsidence events according to island location.

## 2. Geological and geophysical background

French Polynesia lies in the central South Pacific between
latitudes 5°S and 30°S and longitudes 130°W and 160°W,
extending over about 5,500,000 km² (Fig. 1). It has the
highest concentration of volcanic features, including a wide
range of isolated seamounts, midplate islands, aseismic
plateaus and en-echelon ridges. The whole region belongs to
the so-called *South Pacific Superswell* zone (McNutt and
Fischer, 1987; Goodwillie and Watts, 1993; McNutt, 1998;
McNutt et al., 1996; Sichoix et al., 1998; Jordahl et al.,
2004; Suetsugu et al., 2007; Adam et al., 2014), i.e. an
anomalous seafloor, 250 to 1,000 m shallower than the
depth predicted for its age by the classical seafloor
subsidence models that assume a decaying effect increasing
with age (Parsons and Sclater, 1977). It is associated with a
low shear-wave velocity suggesting the presence of a
shallow mantle source (Suetsugu and Hanyu, 2013), a geoid
low (McNutt and Judge, 1990) and a low elastic plate
thickness (Calmant and Cazenave, 1986) with respect to the
empirical elastic thickness versus age relationship (Calmant
and Cazenave, 1987). These attributes were regarded as
reflecting a combination of lithospheric reheating, thinning
and dynamic uplift due to ascending mantle flow, i.e. a
superplume (McNutt and Judge, 1990; Koppers et al., 2003;
Adam et al., 2010). The age of its sea floor ranges from 58
to 70 million years (Ma) (Talandier and Okal, 1987;
Munschy et al., 1996). The depth anomalies resulting in the
shallow bathymetry are thought to arise from physically or
chemically induced decrease in the density and thickness of
the upper mantle beneath the regional hotspot swells
(McNutt and Judge, 1990; Sichoix et al., 1998), such that the
lithosphere is easily penetrated by hotspot volcanism
(Monnereau and Cazenave, 1988; Bonneville et al., 1997).
The region contains a number of volcanic island chains,
namely, the Cook-Austral, the Marquesas, the Pitcairn and
the Society, regarded to be related to local, stationary
hotspot swells and aligned along approximately NW-SE
linear trends, clearly falling into the pattern of absolute
motion direction (N115 ± 15) of the Pacific plate (Duncan
and Clague, 1985; Clouard et al., 2000). By contrast, the
underlying oceanic crust exhibits structural features,
including south-west to north-east-oriented, transform faults
and magnetic lineations (Patriat et al., 2002) and volcanic
ridges (Jordahl et al, 2004), oblique to the hot-spot-
generated volcano trails (Ito et al., 1995). According to
Clouard and Gerbault (2008), some of these features could
have been generated by restricting kinetics of the south Pacific plate at its boundaries. Due to a slowing down of the westward drift of the south Pacific plate triggered by the Kermadec-Tonga Trench retreat westwards, the plate in its central part would be deforming internally, with localized deviatoric and rotating shear stresses. This view is supported by the occurrence of differentially oriented (N110° to N145°) ridges, hotspot tracks and island chains, from Line Islands to the north-west, to Pukapuka and Crough Ridge to south-east (Fig. 1).

This region subsides less rapidly than predicted by thermal subsidence models of oceanic lithosphere (Parsons and Sclater, 1977; McNutt and Fischer, 1987). The evolution of single hotspots would be characterized by intermittent, short-lived volcanic activity, their longevities being shorter than 20 Ma. This idea led Koppers et al. (2003) to argue that such hotspots were related to multiple closely-spaced, thinner plumes (plumelets) originated at very shallow mantle depths from the top of oscillating superplumes. By contrast, Yoshida (2014) claimed that some hotspot plumes in the South Pacific do not originate from specific, large low shear-wave velocity areas, but are isolated features originated from the base of the mantle transition zone (about 600 km deep).

Situated between latitudes 16–19° South and longitudes 147–153° West, the Society Archipelago is administratively composed of two island groups, namely, the Leeward islands – e.g. Bellingshausen and Scilly Atolls; Tupai, Maupiti, Bora Bora, Raiatea, Tahaa, Huanine as volcanic high islands – and the Windward Islands – Moorea, Tetiaroa, Tahiti, Mehetia high Islands. The island chain spreads across a 700 km-long and 200-km wide band nearly
oriented in the spreading direction of the present Pacific plate. This is supported by the observed 11.1 cm/yr rate along-chain migration of volcanic activity, in good agreement with the absolute motion of the plate relative to a fixed hotspot (Morgan, 1971; Duncan and McDougall, 1976). It is considered to be a typical hotspot-generated volcanic alignment which is believed to have originated from a still seismically active intraplate hotspot zone located close to Mehetia island and nearby seamounts, including Teahitia, Rocard, Cyana, Moua Pihaa (Talandier and Okal, 1987; Cheminée et al., 1989; Devey et al., 1990, Hekinian et al., 1991; Nolasco et al., 1998), south-east of the main island of Tahiti through an already cool and strong oceanic crust, 90 Ma-old in the north to 65 Ma-old in the south (White and Duncan, 1996; Grevemeyer et al., 2001). Consistent with the hotspot theory (Morgan, 1971), the Society Islands display a westward increase in age from Maupiti to Mehetia and probably to Scilly and Bellingshausen Atolls, at the northwestern end of the Society chain trail. The radiometrically dated rocks from sub-aerially exposed edifices have provided ages that range from 5 to about 0 Ma: Maupiti (4.79 to 3.9 Ma), Bora Bora (4.01 to 3.1 Ma), Tahaa (3.29 to 2.57 Ma), Raiatea (2.77 to 2.29 Ma), Huahine (3.19 to 2.02 Ma), Moorea (2.15 to 1.36 Ma), Tahiti-Nui (1.12 to 0.51 Ma), Tahiti-Iti (0.51 to 0.25 Ma), Mehetia (0.38 to 0 Ma) (Duncan and McDougall, 1976; Brousse and Léotot, 1988; Guillou et al., 1998; Blais et al., 2000, 2002; Yamamoto et al., 2002; Uto et al., 2007 and references herein), consistent with the plate motion at a constant rate of 11.1 cm/yr. Volcanism is thought to be related to deep-rooted upwelling pulses separated by pauses of less than 1 Myr duration (Uto et al., 2007). The duration of subaerial volcanic activity in a given island ranges from about 0.30 to 1.17 Myr. From Maupiti to Huanine, the activity spanned
288 2.78 Myr and from Moorea to Mehetia, only 1.90 Myr. The
289 subaerial eruptions in Moorea and Tahiti occurred within an
290 area of about 30 km wide and 110 km long, whose
291 orientation is broadly consistent with the drift of the Pacific
292 plate (Uto et al., 2007).

293 The NW-SE extent of the active Society hotspot zone is
294 thought to be about 300 km (Nolasco et al., 1998), or closer
295 to 500 km (French and Romanovicz, 2015; Isse et al., 2016).
296 The eastern boundary of the Society hotspot area would be
297 delineated by a bulge east of Mehetia, which is also the
298 frontier of the recent volcanic activity in the area (Hekinian
299 et al., 1991). Further east, the seafloor deepens to about
300 4,000 m and shows a typical abyssal hill topography
301 (Binand et al., 1991), pratically devoid of sediment cover.
302 The age of the seafloor surrounding island basements at the
303 time of loading is defined as its *effective thermal age*
304 (McNutt, 1984). At Tahiti, located at about 150 km north-
305 west of the active hotspot zone, the effective thermal age of
306 the underlying oceanic crust is 5.4 Ma (Calmant and
307 Cazenave, 1986). At Maupiti and Bora Bora, situated at
308 about 480 km NW of the hotspot, the effective thermal crust
309 ages are 13.5 and 9 Ma respectively. Beneath Mehetia and
310 the nearby seamounts, the age is 2.75 Ma (Calmant and
311 Cazenave, 1986). Assuming that the thermal crust ages
312 reflect the actual times of edifice emplacement throughout
313 the Society chain, the mean eruptive rates of the edifices
314 during their submarine growth phases can be estimated,
315 knowing the total edifice volumes (*Supplementary material*,
316 Table S1 and Figure S1; Table S2). These rates tend to
317 increase from the north-west to the south-east: 0.12-0.32
318 km³/kyr from Maupiti to Raiatea, 0.47 km³/kyr at Moorea,
319 1.95 km³/kyr at Tahiti-Nui. This suggests that the activity of
320 the regional hotspot may have varied through time, as
321 sustained by the similar decreasing trend observed in the
subaerial eruptive productivity. The values range from 0.0042 to 0.22 km³/kyr in the older Leeward Islands over a duration of 2.78 Myr, while Moorea and Tahiti volcanoes were built at rates of 0.070, 0.86, and 0.26 km³/kyr respectively over a duration of 1.9 Myr. Over the last hundred thousand years, while the basement of Mehetia was produced at a rate of 0.43 km³/kyr prior to 0.38 Ma, the hotspot has become less active, building the subaerial portion of Mehetia at a mean rate lower than 0.0010 km³/kyr. Relatively lower between 5 to about 3 Ma, the volcanic intensity in the Society increased significantly after 3 Ma and peaked between 2 and 0.50 Ma.

The active volcanic zone at the surface extends downwards into an anormally conductive lithosphere whose bottom, between 70 and 130 km deep, is partly melted (1-2%). Beneath this zone, from 130 to 300 km deep, the mantle is highly resistive, probably due to a melted and volatile depleted fraction of the mantle (Nolasco et al., 1998). At about 400-450 km deep, there would be a significant drop in mantle resistivity, interpreted as an increase of temperature of about 350 °C. Despite the fact there is no clear evidence for intense reheating and thinning of the lithosphere in this area, temperature increase may be indicative of upwelled hot magmatic material from the depth (Nolasco et al., 1998). Measurements of buoyancy fluxes (i.e. flux of hot material from the mantle) under the Society (1.53-1.58 Mgs⁻¹) indicate that there is a source of a negative density anomaly between 0 and 240 km deep, assumed to reflect a vertical mantle upwelling directly reaching the lithosphere beneath the active volcanic zone (Nolasco et al., 1998).

The Tuamotu islands are separated from the Society Archipelago by basins exceeding 5,000 m in depth (Fig. 1). They rise from a vast submarine plateau, namely, the
Tuamotu Plateau, part of the Pacific Superswell. This plateau is approximately 1,500 km long and 800 km wide and is less than 2,000 m in average depth. The plateau is made up of basaltic flows, thought to have been produced most probably by near-axis volcanism associated with propagating rifts developed near the crest of the ancestral East Pacific Rise (Talandier and Okal, 1987; Ito et al., 1995). Gravity field analysis revealed a strong negative anomaly beneath the plateau interpreted as evidence of crustal thickening beneath (Panet et al., 2006). Except for Moruroa and Fangataufa Atolls (Bablet et al., 1995), located at the south-east extremity of the Tuamotu Plateau, none of the Tuamotu atolls have been depth-drilled to penetrate the basaltic basements. The scarcity of radiometric dates from volcanic basements therefore did not allow a linear volcanic migration pattern to be validated in the Tuamotu. However, volcanic rocks dredged along the south-western flank of the northern Tuamotu plateau at depths of 1,000–1,500 m yielded Ar/Ar ages of 47.40 ± 0.90 to 41.80 ± 0.90 Ma (Schlanger et al., 1984). Ages of seamounts between 14°5' south and 17°0' south fall into 53.5 to 37.5 Ma (Jarrard and Clague, 1977). The carbonates recovered by depth-drilling at about 85 km east of Rangiroa seamount flanks provided maximum ages of 55 Ma (Burckle and Saito, 1966; Martini, 1976). All these ages agree with that of volcanic loading, ranging from 55 to 50 Ma (Ito et al., 1995). According to Okal and Cazenave (1985), Talandier and Okal (1987), the formation of the plateau may have started as soon as 70 million years ago, in agreement with the oldest age of the underlying lithosphere (Munschy et al., 1996). The north-western parcel is underlain by an oceanic crust with ages of 64 to 67.5 Ma (Ito et al., 1995; Patriat et al., 2002). The Tuamotu is supplied by the weakest flux (0.53–0.72 Mgs⁻¹) in the Polynesian region (Adam et al., 2010), as evidence of
a significant decrease in volcanism. The emplacement of the Tuamotu volcanoes is thought to result from the passage of two or three hotspots 50 to 30 Ma ago as they migrated beneath the Pacific plate (Ito et al., 1995). However, the Tuamotu island group, like the Line islands, appear to result from a polyphase formation, as composed of multiple, short, sub-parallel, irregularly progressive subchains and ridges, oriented from N100° (Pukapuka Ridge), N115–120° – Pitcairn-Gambier, northern and central atoll chains –, N135° (Anaa Ridge) to N140° – Reao-Tepoto alignment –, closely-spaced, about 50–100 km apart from one another with a great number of isolated seamounts, and not a single system showing the uniform progression required by the hotspot model (Jackson and Schlanger, 1976). Owing to the spacing between the chains and the absence of age-progression pattern, their origin is still questionable. Pre-existing crustal discontinuities may have influenced the loci of volcanism (Binard et al., 1992). Alternative models can be put forward to account for the formation of linear volcanic chains, with no clear evidence of linear age progression (Ballmer et al., 2007), particularly, to be in agreement with the presence of linear, prominent, discontinuous, sub-parallel to radial aseismic ridges clearly built by the superposition of small, discrete, flat-topped volcanic cones (Jordahl et al., 2004). Note that, in response to shearing stresses at its western boundary, the Pacific plate has experienced a rotating deviation to the east (Clouard and Gerbault, 2008). This has resulted in a marked reorientation of the Line Island chain and the Boudeuse Ridge further east (N150°), relative to the northern Tuamotu segment (N120°–125°).

Seismic investigations across the Society and northern Tuamotu areas allowed to best image the crustal structure
regionally (Fig. 2). The structural model reveals a three-layer crust including a volcano-sedimentary, a basaltic and a gabbroic layer (Talandier and Okal, 1987; Patriat et al., 2002). Beneath Tahiti and surrounding crustal areas, sediments are about 0.4 km thick, basaltic flows 4 to 1.8 km thick and gabbros 6.2 km thick. The total thickness of the oceanic crust is 12–13 km. Tahiti is associated with a flexural deepening of the crust/mantle boundary (Moho). This situation is that of a regional isostatic compensation, with a long flexural wavelength, apparently not compatible with flexure-generating uplift models. The Society basements would be supported by a strong lithosphere (Watts and Ten Brink, 1989). Beneath the Tuamotu Plateau, at Rangiroa Atoll, the total crust thickness is markedly thicker (21 km) than under Tahiti, with a significant thickening of the lower, gabbroic layer and an about 2 km-thick carbonate cap. The Tuamotu Plateau could be close to isostatic equilibrium since its formation (Sandwell and McKenzie, 1989; Mutter and Mutter, 1993). Differences between the Society and Tuamotu structures could have been caused by differences in the thermal state of the oceanic crust at the time of volcanic activity (Patriat et al., 2002). By contrast, beneath the seafloor in-between the Society and the Tuamotu Plateau, on which Makatea Island is sitting, the lithosphere does not exceed 10 km thick.

In the Tuamotu, reef-related carbonate deposition atop of the oldest volcanoes is likely to have begun as early as 51 to 49 Ma (Jackson and Schlanger, 1976). The volcanic basement of the atolls is not accurately dated, but are at least 50 Ma old, based on a variety of paleontological evidence (summarized by Jarrard and Clague, 1977; Ito et al., 1995) and at most 63 Ma based on the age of the underlying seafloor. According to the standard curve, the sea floor in
the Tuamotu region would subside approximately at rates of 0.025 mm/yr (Crough, 1984). Estimates of long-term subsidence rates derived from the expected thickness of carbonates accumulated at Rangiroa (2,000 m) over the last 55 Ma (Talandier and Okal, 1987) provide higher values of 0.040 mm/yr. Similarly, long-term mean subsidence rates of 0.035-0.047 mm/yr were obtained from Takapoto Atoll, north-west Tuamotu, based on the expected present-day position of the Last Interglacial reef terrace (Montaggioni et al., 2019). These values are in good agreement with those established from regularly subsiding atolls in the central Pacific (0.020-0.050 mm/yr; Detrick and Crough, 1978; Menard and McNutt, 1982; Schlanger et al., 1987; Paulay and McEdward, 1990; Toomey et al., 2016) especially, on Cook Islands (0.045 mm/yr; Gray et al., 1992) and on Eniwetok, Marshall Islands (0.020-0.040 mm/yr; Ladd and Schlanger, 1960; Quinn and Saller, 1997) based on the present stratigraphic positions of Miocene to Pleistocene sea-level unconformities. A low amount of subsidence would result from the additional flexure of the underlying oceanic crust due to the increasing growth of the reef limestone cap over time (Detrick and Crough, 1978).

### 3. Current knowledge of the geological history of the so-called *uplifted atolls*

Reconstructing the geodynamical evolution of the north-west Tuamotu Islands requires determination of the spatial distribution, age, stratigraphy and original elevations reached by the raised reef limestones (Fig. 3) and associated sea-level indicators. Makatea Island was visited in April 1982 and October 2021, Mataiva in May 1981 and October 2021, Tikehau and Rangiroa in April-May 2021, Kaukura and Anaa in May 2022 and Niau in October 2022. The
487 geographic coordinates of collected samples for radiometric
488 dating (See Supplementary material for descriptions of
489 dating methods) and thin-section analysis were determined
490 using a handheld GPS Garmin Etrex 30X. Elevations of sea-
491 level indicators relative to pmsl were estimated using a
492 measuring stick and level, resulting in an error range
493 estimated of about ± 0.50 m.

494 **3.1. Makatea**

495 Located at 15°50' south and 148°15'west, 245 km north-east of
496 Tahiti, Makatea is a high carbonate island, 7 km long from
497 north to south and 4.5 km wide from west to east (Fig. 3).
498 The age of the underlying oceanic crust ranges from 65 to
499 66 Ma. The island has a typical atoll, i.e. ring-shaped,
500 morphology. The outer margin displays prominent vertical
501 cliffs over 40-45 to about 75 m high (Fig. 4A). The cliff-top
502 has an abrasion plateau-like topography, with pitted
503 pinnacles (feo) locally culminating at 113 m high. The
504 plateau is deeply dissected by a dense karstic network, in the
505 form of two central basins with floors ranging from 20 to
506 35 m elevation above pmsl, a great number of close-set pot-
507 holes, about 75 m in maximum depth and residual high
508 reliefs. The coastline is surrounded almost continuously by
509 necrotic fringing reefs, about 30 to 120 m wide.

510 Detailed lithological and stratigraphical examination of the pile
511 deposits revealed a clear concentric zonation of
512 subhorizontal, flat-topped, shallow-water carbonate deposits
513 with a peripheral coralgal rim enclosing a series of internal
514 sandy to muddy zones rich in macrofauna (corals, mollusks,
515 echinoids) and microfauna, including Miogypsina,
516 Miogypsinoides, Austrotrillina, Borelis (Montaggioni,
517 1985). Based on the composition of foraminiferal
518 assemblages, the entire carbonate pile that forms the bulk of
519 the island was found to be of Early Miocene age
520 (Aquitanian, about 23 to 20 Ma) (Montaggioni, 1985;
521 Montaggioni et al., 1985; Bourrouilh-Le Jan and Hottinger,
522 1988). The Miocene carbonates are calcitized (Fig. 12A) and
523 locally deeply dolomitized (Fig. 12B).

524 The presence of geomorphic features, such as several metres
525 high, jagged pinnacles and deep (tens of metres) pot-holes
526 indicates unambiguously that the dissected central upland
527 has resulted from sub-aerial erosion and the modern
528 topography with central depressed areas is clearly karst-
529 inherited. Stratigraphic and lithologic features, such as
530 subhorizontal depositional beds with small-scale, low-angle
531 cross-bedding, suggest that originally sedimentary beds do
532 not slope away from the platform margin inwards
533 (Montaggioni, 1985). The shape of the platform at the time
534 of growth has been broadly flat-topped, resembling to most,
535 flat-topped Miocene carbonate platforms found in the
536 tropical Indo-Pacific areas (Wilson, 2002; Teillet et al.,
537 2020). The depositional system was regarded as having
538 functioned mainly as a rimmed platform (Montaggioni,
539 1985, 1989; Montaggioni and Camoin, 1997), probably
540 since its emplacement by the lower Eocene as suggested by
541 the age of carbonates deposited at the base of the nearby
542 Rangiroa Atoll (Schlanger and Premoli-Silva, 1981; Ito et
543 al., 1995).

544 Karstic cavities within the central island depression are infilled
545 with heterogeneous phosphate deposits. These occur as
546 consolidated and unconsolidated material, composed of
547 oolitic to pisolitic sands and gravels (Obelliane, 1953.
548 Montaggioni, 1985; Bourrouilh-Le Jan, 1990). Whatever the
549 source of phosphorus (guano, marine organisms or waters),
550 phosphatogenesis was assumed to have occurred through
551 diagenetic recycling of phosphorus by microbial activity
552 resulting in the formation of cyanobacterial mats – *kopara* in
553 the Polynesian language – within shallow-water, lagoonal
554 pools (Jehl and Rougerie, 1995; Rougerie et al., 1997;
555 Trichet and Fikri, 1997). This means that the main phase of
556 phosphatogenesis took place prior to island uplift. U–Th
557 dating of apatite fraction in phosphates at Makatea provided
558 age estimates of > 200 ka (years x 1000) (Aharon and Veeh,
559 1984) and > 300 ka (Roe and Burnett, 1985) poorly
560 consistent with the depositional history of the island. Veeh
561 (1985), on the basis of <sup>230</sup>Th/<sup>231</sup>Pa data, pointed out that the
562 obtained ages referred to those of the apatite precipitation,
563 but not necessarily to the time of phosphate deposition on
564 the island, because insular phosphorites frequently exhibit
565 evidence of reworking or redistribution from meteoric
566 waters. These phosphorites are strictly restricted to the upper
567 plateau interior and clearly predate the Pleistocene reefal
568 deposits. Accordingly, uranium-series ages of insular apatite
569 should be regarded as minimum age limits. The actual ages
570 are likely to be at least 800 ka or older (Veeh, 1985),
571 probably referring to the Pliocene (Montaggioni, 1989).

572 The carbonate pile has experienced at least five distinct phases
573 of dolomitization, thought to have been related to periods of
574 sea-level stabilization over the Pleistocene. Dolomite occurs
575 as isolated, broadly lenticular masses embedded within low-
576 magnesian calcite deposits (Montaggioni and Camoin,
577 1997).

578 The coastal cliffs display multiple levels of former sea-level
579 notches and cave lines respectively at elevations of 1.0–1.5
580 m, 6–8 m, 25–30 m, and 43–47 m above pmsl (Figs. 4A &
581 B). In places where the notches are missing, there are locally
582 indications of isolated caves. At their southwestern end, the
583 cliffs change into two stepped abrasional terraces at levels of
584 +10–12 m and +27–29 m respectively.
The lowest notch line is in planimetric continuity with an irregularly pitted, necrotic coralgal reef platform, 0.3 to 1 m in maximum elevation. This is primarily constructional in origin and appears to be the remnants of fringing reef flats developed during the Mid-Holocene higher sea level (Fig. 4C) identified at similar elevations in most islands throughout the French Polynesia (Pirazzoli and Montaggioni, 1988; Hallmann et al., 2018, 2020) and Cook Islands (Gischler et al., 2022) interpreted as resulting from the Global Isotatic Adjustment -GIA- (Hallmann et al., 2020). Calibrated radiocarbon age of the exposed reef flats range between 3,966 and 5,995 yr BP (Supplementary material, Table S6). This suggests that Makatea Island has experienced no thermal uplifting over the last 6,000 years. Visual inspection of the seaward edge of the platform indicates that the Mid-Holocene reef is a veneer over a karstified limestone landscape.

The two notch lines at +6–8 m and +25/+30 m are in close connection with two distinct generations of apron reefs, adjoined to the coastal cliffs, ranging respectively from 0–6 m and from 8–10 to about 25 m above pmsl (Fig. 4D). These reef units have previously provided U/Th ages of 100 ± 20 ka to 140 ± 30 ka and of ≥ 200 ka (Veeh, 1966) respectively. New uranium-series dating confirm the fact that the lower apron reefs were built during the last interglacial high sea stand (MIS 5e) (Supplementary material, Table S3: Samples MK 32, MK 46, MK 48). These results are consistent with ESR dating indicating that the lower reef aprons have developed between 91 +7/-4 and 139 +12/-5 ka (Supplementary material, Table S4: Samples MK 7, MK 9, MK 15, MK 18, MK 38). The sea-level indicators for the MIS 5e highstand at Makatea range from about 5 m (reef crests) to a maximum of 6–8 m (notches). There is a
broad consensus that the Last interglacial highstand was higher than today, between 6 m and 9 m higher than present (Kopp et al., 2009; Dutton and Lambeck, 2012; Medina-Elizalde, 2013; Polyak et al., 2018) although recent estimates from Bahamian islands indicate that the global mean sea level peaked at least 1.2 m (95 % probability) higher than present, but probably not exceeding +5.3 m (5 % probability) (Dyer et al., 2021). However, as emphasized by Creveling et al. (2015), local sea-level position within a given time frame is typified by marked departure from eustasy due to GIA. In the northern Tuamotu, during MIS 5e, isostasy effect was predicted to reach up to about ± 2 m. Accordingly, the current position of apron-reef tops and associated notches at about +6 m relative to pmsl strongly supports the idea that Makatea Island has undergone no significant vertical motion since the Last Interglaciation.

Furthermore, ESR dating revealed that the upper apron-reef generation was emplaced between 260 +22/-11 ka and 293 +25/-13 ka (Supplementary material, Table S4: Samples MK 13, MK 20). On a number of other uplifted islands in the central South Pacific, reef units apparently belonging to the similar generation have been found. On Atiu and Mauke, southern Cook Island Group, Woodroffe et al. (1991) described well-developed reefs, ranging in ages between 261 ka and 205 ka, at heights of 1 to 15 m above pmsl. Similarly, in the far-east Polynesian region, on Henderson Island (Pitcairn Group), there are prolific reef units flanking the coastal cliff at maximum elevations of about 26 m that provided U/Th ages of 230-215 to 285-275 ka (Blake, 1995). Accordingly, at Makatea, the upper reef aprons could be tentatively interpreted as deposited during the penultimate (MIS 7e) interglacial substage (245 to 230 ka). Findings by Toscano and Lundberg (1999), Camoin et al.
651 (2001), Bard et al. (2002), Schellmann and Radtke (2004),
652 Schellmann et al. (2004), Dutton et al. (2009 a,b) and
653 Colleoni et al. (2014) based on coral-reef and speleothem
654 records suggested that during the whole MIS 7 duration (245
655 to 190 ka), eustatic sea level was expected to have
656 maintained continuously several metres below present level,
657 at about -6/-9 to -18/-20 m. Although some estimates
658 placed MIS 7 sea-levels close to or above present (Chappell,
659 1974; Murray-Wallace, 2002; Thompson and Goldstein,
660 2005; Rowe et al., 2014), other studies based on planktonic
661 and benthic foraminiferal oxygen isotopes (Shackleton et al.,
662 1993; Rohling et al., 2009; Spratt and Lisiecki, 2016 and
663 references herein) and on coastal sites with low and high
664 uplift rates (Muhs et al., 2011) seem to confirm that sea
665 level during the MIS 7 episode was significantly lower than
666 present. Assuming that the eustatic, ice-equivalent, paleosea
667 level during MIS 7e was positioned several metres below
668 pmsl, mean rates of island uplifting between MIS 7e and
669 MIS 5e (over a time interval of about 120 kyr) should have
670 been ranged between 0.25 and 0.30 mm/yr to bring any MIS
671 7e-related apron reef and notch line at the present-day
672 maximum elevation of +25 m.

673 The uppermost, locally, well preserved notch line, at about +43
674 m in height (Fig. 4B), could be attributed tentatively to the
675 MIS 9 (335.6 to 316.8 ka) high sea stand as known to have
676 been close to the present (Bowen, 2010), +1 m or higher
677 than present (Murray-Wallace and Woodroffe, 2014). If
678 correct, this suggests that, since the MIS 5e substage, island
679 uplifting took place at mean rates of 0.17-0.19 mm/yr.

680 As for the flat-topped, outer rim surfaces of the upper island
681 plateau at about +70 m high, interpreted to marine
682 abrasional features by Bourrouilh-Le Jan (1985), they could
683 have been formed during an older Pleistocene high sea
684 stand. It should be noted that, due to their specific location,
685 the apron-reef crests in planimetric continuity with the base
686 surfaces of the two lowest notches seem relatively well
687 preserved, partly covered by flowstone floors and
688 concretions and, thus, not directly affected by significant
689 surface loss.

690 Scuba diving investigations seaward of the north-north-eastern
691 coast (Teapoto coast) revealed that there is a gentle-dipping
692 (5°) terrace between 9 and 11 m deep, about 30 m wide
693 (Montaggioni et al., 1987) (Fig. 5). This terrace might be
694 interpreted as resulting from constructional or erosional
695 processes. In the first case, the terrace would be the top
696 surface of a former reef tract built right up against the
697 Miocene pile. In the second case, it would be an abrasional
698 platform dug through the pile. In both cases, the terrace is
699 indicative of a former, Pleistocene, lower sea level.

700 3.2. *Mataiva*

701 Mataiva Atoll is located by 14°55' latitude south and 148° 36'
702 latitude west, 300 km N-NE of Tahiti. It is 10 km long and 5
703 km wide. The atoll basement sits on an underlying plate,
704 aged 66.4 to 64.3 Ma. Remains of the carbonate pile (feo)
705 are cropping out locally through the modern inner reef flat,
706 on the northern coast of the atoll, at maximum elevations of
707 about +2.50 m (Figs. 3 and 6A). As on Makatea, the rock is
708 deeply dolomitized (Fig. 12C). Seawards, the gentle-sloping,
709 upper fore-reef terrace is thought to be the top surface of the
710 carbonate atoll basement, presumably covered by a thin
711 coralgal veneer. This is supported by subsurface drilling
712 surveys revealing that, beneath the modern rim-islet and
713 lagoon topography, the basement paleosurface occurs in the
714 form of a basin-shaped structure and pot-holes distributed
715 between depths of 5 and 23-30 m (Rossfelder, 1990; this
study), interpreted as residual karstic features (Fig. 7). These are at present buried beneath 4 to 15 m-thick, Holocene reefal material. Calibrated radiocarbon ages of coral material (Supplementary material, Table S6) ranges from 305 to > 35,250 yr BP (Pirazzoli and Montaggioni, 1986). Due to its very poor state of preservation and by analogy with the nearby island of Makatea, the basement was believed to be of Miocene age (Delesalle, 1985; Pirazzoli and Montaggioni, 1986). As observed on Makatea, the basins and pot-holes within the lagoon are infilled with phosphate deposits. Age estimates of phosphate deposition, based on $^{230}$Th/$^{238}$U and $^{226}$Ra/$^{238}$U ratio, range from 32 ± 6, 53 ± 2, 80 ± 10, 90 ± 6, 127 ± 6, 137 ± 23, 210 ± 15, and > 300 ka, suggesting that several periods of phosphatogenesis would have taken place over the Late to Mid Pleistocene (Roe and Burnett, 1985). It is unlikely that phosphatogenesis has mainly proceeded relatively recently, according to the historical context in which the island has evolved. Phosphates appear in the form of hard, light-brown to friable, chalky, white phosphatized limestones, indicating contamination with younger coral material. An Early Pleistocene to Pliocene age appears to be more consistent with the island geological history. Indeed, the issue concerning the values of isotope activity ratios as probably artifacts due to rock weathering, leaching and nuclide migration has been raised by the authors themselves. The closed system is regarded as having been violated and any ages derived from Th/U and Ra/U are suspected to be refreshed (Roe and Burnett, 1985).

There is no evidence of elevated reef remnants of Pleistocene age on the atoll rim. Interestingly, some mega-blocks derived from the adjacent foreereef zone are currently trapped into the modern sandy beach and conglomerates along the northern coast (Fig. 6B). U/Th dating of one of them yielded
an age of 119,143 ± 1,498 and 116,864 ± 1,730 calendar yr
BP (Supplementary material, Table S3: Sample MTV 11).
This indicates that remnants of the reef tract belonging to the
MIS 5e stage occur presently along the outer reef slope, thus
implying island subsidence at least since the Last
Interglacial.

### 3.3. Tikehau

Tikehau Atoll is situated at latitude 18° south and longitude
148° west. Subcircular in shape, the atoll is nearly 28 km in
widest diameter. The atoll apron rises from a 65.1 to 63.5
Ma-old lithosphere.

Surveys by scuba diving along the west-north-west upper
foreereef zone (Pahitomo area), showed that, forward of the
spur-and-groove system, from 6 m deep, there is a
subhorizontal to gentle-dipping (< 10°), upper terrace less
than 10-12 wide, and interrupted at around 11 m deep by a
+5 to +7 m high cliff. The latter ends on a lower, sub-
horizontal, flat-topped terrace partly covered by fallen
blocks. The terrace is connected at around 18 m deep to a
poorly developed system of buttresses and valleys ending
into a drop-off (Fig. 8). Similar features were encountered in
other places along the atoll foreslopes at similar depths
(Harmelin-Vivien, 1985). Such features may be regarded as
indicating the position of a former sea stand, believed to be
of Late Pleistocene age.

There are a number of exposed old reefal remnants (feo)
cropping out through the southern, eastern and north-
western parts of the atoll rim, culminating at elevations of
10-12 m above pmsl (Fig. 3). These are typical sharp-
pointed, jagged, karstic morphology (Fig. 6C), with
calcitized to deeply dolomitized facies (Figs. 12D and 13A).
By analogy with the nearby Makatea Island and due to the
occurrence of *Austrotrillina foraminifera* (Fig. 12D), these reliefs are thought to be of Cenozoic age, probably Miocene.

Reef drilling was carried out in the south-south-west areas of the atoll (Rougerie and Wauthy, 1993; Rougerie et al., 1997). Four drillholes (P1, P2, P4 and P5), 30 to 33 m long, penetrated the rim reef-flat zone. An additional hole (P3), 17 m long, was drilled through an emergent, reef patch within the lagoon, at a distance of about 5 km inwards from the reef flat (Fig. 9). Examination of the cores extracted from the reef-flat zone revealed a continuous sequence of karstified, calcitized to dolomitized reef carbonates, with an apparent porosity of 30%, overlain by a thin (a few cm-thick) living coralgal veneer. The lagoonal reef-patch sequence was composed of soft, reef-derived, skeletal sediments (Rougerie and Wauthy, 1993; Rougerie et al., 1997). By analogy with lithology of old limestones from Makatea Island, the *feo* from Mataiva Atoll and from the adjacent beaches, the outcropping limestones and subsurface ones beneath the modern reef flat are considered to be of Miocene age. By contrast, the lagoonal patch reefs appear to be made up of post-glacial to Holocene deposits, at least at their upper parts. Indeed, coral samples extracted from the core P3 provided U/Th ages ranging from 16,697 ± 54 calendar yr BP at base to 4,520 ± 12 calendar yr BP at top (Supplementary material, Table S3: Samples T5, T8, T15, T28, T39, T44).

Fossil microatolls of massive *Porites* colonies, U/Th-dated between 3.213 and 5.747 calendar yr BP, are found on the western reef flat zone, exposed at elevation of +0.04 m to up to +0.80 m with reference to the elevation of living microatolls at the same study site. This has contributed to document a Mid to Late Holocene sea-level highstand less than 1 m high, preceding a sea-level fall (Hallmann et al., 2020).
816 As on Mataiva, on Tikehau, there is no evidence of exposed
817 Pleistocene reef limestones.

818 **3.4. Rangiroa**

819 Located at 15°10' south and 147°40' west, Rangiroa, the
820 largest Polynesian atoll, is 78 km long and 33 km wide
821 (Ricard, 1985). The atoll basement sits on an ocenic crust,
822 64.5 to 61.4 Ma old. Along the outermost parts of the
823 southern atoll rim (Île aux Récifs), there are deeply pitted,
824 jagged, pinnacle-shaped limestones, up to +3 m in elevation
825 (Ricard, 1985; Stoddart 1989) (Figs. 3 and 6D). Analysis of
826 their mineralogical composition and microfacies indicated
827 that they are composed of low-magnesian calcite, coralgal
828 boundstones to grainstones with local enrichment (55-70 %)
829 in dolomite (Figs. 13B and C). As for Mataiva and Tikehau,
830 such rock remnants were interpreted as pieces of ancient
831 reef-related deposits of Late Neogene age (Vahrenkamp and
832 Ginsburg in Ricard, 1985).

833 Investigations using electrical resistivity tomography conducted
834 through the north-eastern atoll rim have identified the top
835 surface of the carbonate pile, at depths ranging from 10 to
836 about 50 m beneath the modern reef surface and interpreted
837 as Miocene in age (Montaggioni et al., 2022). Indeed, the
838 pile surface appeared to be very similar to that of the upper
839 plateau from the nearby Makatea Island, with irregular-
840 shaped and large, karstic cavities and to pinnacle-shaped
841 dolomitized remnants found in the southern atoll rim. The
842 sub-surface identified cavities appear to be infilled with
843 loose to poorly consolidated material, believed to be
844 phosphates. The uppermost, subhorizontal layers, less than
845 10 m in total thickness, were considered to be composed of
846 Late Pleistocene to Holocene reefal sediments (Montaggioni
847 et al., 2022). However, as observed on Mataiva and Tikehau,
848 exposed remains of Pleistocene reef limestones are missing.

849 There are emergent coral micro-atolls at many reef-flat sites
850 around the atoll rim, at the elevation of about +0.70 m
851 relative to the elevation of modern, nearby analogues. U/Th
852 ages of fossil corals range from 2,966 to 5,487 calendar yr
853 BP, confirming a sea-level highstand over the Mid-Late
854 Holocene in the north-west Tuamotu (Hallmann et al.,
855 2020).

856 3.5. Kaukura

857 Situated at latitude 15°45' south and longitude 146°41' west,
858 Kaukura rises from a plate aged 62.3 to 60.1 Ma. It is
859 elliptical and, as Rangiroa, elongated along a WNW-ESE
860 axis. Its size is 48 km long and 15 km wide. From the north-
861 western to the south-eastern atoll-rim end, there are almost
862 continuous lines of karstified, calcitized to dolomitized
863 reliefs (Figs. 3 and 10A) culminating at about +3.5 m high,
864 rich in dolomite (Figs. 13D and 14A). These are considered
865 to be the analogs of those found on atolls further
866 northwestwards (Rangiroa, Tikehau, Mataiva). As in these
867 atolls, at Kaukura, there is no evidence of emerged reef
868 deposits of Pleistocene age.

869 3.6. Niau

870 Niau is an enclosed atoll-like island, located at 16°10' south
871 and 146°20' west. It rises from a 59.3 to 56.8 Ma-old
872 oceanic crust. Subcircular in shape, it is 8 km in mean
873 diameter with a total surface area of 60 km² (Salvat et al.,
874 2018).

875 The reef-rim core consists of massive coral-reef beds. Two reef
876 generations were identified inland (Fig. 3). The oldest reefal
877 remains reach a maximum elevation of +7.5 m high on the
878 northern cliffed coast. Although apparently well-preserved
879 locally (Fig. 10B), these remains appear usually in the form
880 of deeply karstified, pinnacle-shaped rocky structures (Fig.
881 10C), calcitized and/or dolomitized (Figs. 14B & C). Isotope
882 Sr dating reveal that these are Early Pleistocene to Pliocene
883 in age, averaging 1.200 to 3.814 Ma (Supplementary
884 material, Table S5), at times when sea level was episodically
885 at elevations of less than +20 m relative to pmsl during
886 interglaciations and not deeper than -26 m during glacial
887 episodes (Raymo et al., 2011; Rovere et al., 2014; Berends
888 et al., 2021). The youngest reef generation was identified as
889 forming fringing tracts deposited alongside the Pliocene to
890 Early Pleistocene pile, in its western and eastern sides (Fig.
891 10D). Former U/Th dating of coral specimens from the
892 youngest reefs collected at 2 and 4 m above pmsl provided
893 ages of 120 ± 20 and 160 ± 40 ka (Veeh, 1966), indicating
894 that the relevant reef tract has developed during the last
895 interglacial (MIS 5e) highstand. This age is confirmed by
896 new U/Th dating: 135,077 ± 2,289 to 110,866 ± 1,256
897 calendar yr BP (Supplementary material, Table S3:
898 Samples N27, N56, N72). Locally, as observed in some
899 other northwest Tuamotu atolls, there are exposures of mid-
900 Holocene reef features (algal crests) at elevations of 1 m
901 above pmsl (Fig. 11A), dated at 4,623 ± 0.052 calendar yr
902 BP (Supplementary material, Table S3: Sample N18).

903 3.7. *Anaa*

904 Situated at 17°23′ south and 145°30′ west, Anaa is a closed
905 atoll, 28 km long and 7 km wide. It rises from a NW-SE
906 elongated aseismic ridge as a part of the Tuamotu seamount
907 trail. This ridge is apparently formed by a number of small,
908 discrete volcanoes in close proximity.
909 The age of the underlying oceanic crust ranges from 59.6 to 52.5 Ma. The main atoll axis is itself oriented along the NW-SE direction as are most of the north-western Tuamotu atolls. In the north-eastern rim tip, there are well-preserved, emergent reef tracts (Figs. 11B and 14D), about +4 to +5 m high. Based on U/Th dating, their ages range between 150 ± 40 and 110 ± 20 ka (Veeh, 1966). In addition, an ESR age of 109 +8/-4 ka has been obtained from a coral sample containing less than 3 % of calcite (Pirazzoli et al., 1988). These indicate clearly that the fossil reefs that form the modern atoll reef-rim have grown during the MIS 5e stage. There is no evidence of older reef exposures.

921 Elevated, *Porites* microatolls have been found in growth position at different sites on the atoll rim. Samples collected in the northern area, at elevations of 0.75 to 0.85 m above pmsl provided calibrated radiocarbon ages of 1,890, 2,226 and 3,235 yr BP (Pirazzoli et al., 1988).

## 4. The previous regional lithospheric flexure models: uncertainties and inconsistencies

928 The flexural response of the lithospheric plate to volcanic island loading was believed to be the primary mechanism of vertical motion for neighbouring islands in midplate oceanic areas. Volcanic loading is known to produce downward deformation of the underlying elastic plate, as well as uplift of adjacent lithosphere due to the resulting flexure. Deformation results in the formation of a proximal moat and a distal lithospheric arch at the periphery of the loading centre. Although Zhong and Watts (2002) claimed that these models cannot produce the observed uplift and mantle plume-plate interactions and thus cannot be the primary cause of the uplift in the Hawaiian islands, several rheological models related to the lithospheric flexure
generated by the weight of midplate volcanic seamounts were released over the last four decades, especially from observations in the Hawaiian Archipelago (McNutt and Menard, 1978; Watts and ten Brink, 1989; Grigg and Jones, 1997; Huppert et al., 2015), regarded as the archetype for flexure-induced volcanic loading patterns. The picture is that of a deflection under the load of the island chain and a resulting arch of the lithospheric plate peaking at a radius of about 100-140 km from the loading swell locus (Walcott, 1970, 1976; Watts and Cochran, 1974; Watts and ten Brink, 1989; Watts and Zhong, 2000; Watts, 2001; Klein, 2016).

Based on the observations that some north-west Tuamotu atolls with raised limestone deposits are found in close proximity to the younger volcanic islands of Moorea, Tahiti, and Mehetia, McNutt and Menard (1978) and Lambeck (1981a) attributed the observed emergence of the limestones to the position of the atolls on an expected arch generated by a lithospheric deflection in compensation of the mass of one or more of the recently active volcanoes. Bathymetric surveys across Tahiti effectively showed a near-island moat and a peripheral bulge in the surrounding island topography (Lambeck, 1981a). Makatea, the highest island, was believed to be located atop of the arch axis, while the nearby, significantly lower-lying atolls, including Mataiva, Tikehau, Rangiroa, Kaukura, Niau, and Anaa, were thought to rise along the eastern, gentle-dipping flanks of the arch.

The regional flexure models were constrained on the basis of present-day key geometric and geophysical parameters, including gravity or geoid height anomalies, the distribution density of the load, the distance of the peripheral arch and uplifted atolls to the loading axial centre, the degree of infilling of the moat by detrital sediments, the amount of the
so-called *observed uplift* derived from the current elevations
of limestones as a measure of uplift, the mechanical
properties of the underlying lithosphere. Predictions of uplift
amount by McNutt and Menard (1978) were based using a
large arch at a radius of 245–430 km centered on Tahiti and
Mehetia as the loading loci, a load density of volcanic rocks
of 2,800 kg/m³, a lithospheric flexural rigidity of
2.1 × 10²² Nm, a density of the sedimentary material in the
moat of 2,300 kg/m³, and a degree of moat infilling of 1,
present elevations of raised reef remnants – +70 m for
Makatea; 0 to +6 m for other uplifted atolls. As a result, they
assumed that about two third of the 70 m uplift at Makatea
resulted from the load of Tahiti and the remaining one third
from Mehetia alone, an island ten times smaller than Tahiti.
Using the same geophysical and geometric variables as
McNutt and Menard (1978), but with different values –
geoid height anomaly = 10 m; flexural rigidity = 3 × 10²⁹
Nm; load density = 2,500 kg/m³; degree of moat infilling =
0.85 –, Lambeck (1981a) considered that the uplift at
Makatea was truly expressed by an average elevation of
45 m, assuming that sea level at the time of the building of
the Tahiti-Moorea complex was about 10 m higher than
present. As showed in the Hawaiian Islands, the predicted
flexure effects on a multi-layered viscoelastic plate depends
upon the degree of viscoelasticity of both layers. In the case
of Tahiti Island, formed on a 70 Ma-old, two-layer low-
viscosity oceanic plate, with a mean viscosity of 5 to
10 × 10²⁵ Pa (Cazenave et al., 1988), the model indicated
that the load-induced flexure operating over a period of up
to one million years should generate an arch at a distance of
less than 150 km from the loading centre, as observed in the
Hawaiian Islands (Watts and Zhong, 2000). This is in
contrast with the distance of up to 250 km, at present
1006 separating Tahiti to Makatea and Kaukura.

1007 The flexural uplift model defended by McNutt and Menard
1008 (1978), Lambeck (1981a) and later adopted by Spencer et al.

1009 (1987) to explain the presence of uplifted islands in the
1010 Cook and the Tuamotu archipelagoes was challenged by
1011 Crough (1983) and Montaggioni (1989). As emphasized by

1012 Carlson and Johnson (1994), any viable geophysical model
1013 must fit the field data with an acceptable degree of
1014 systematic misfit. Large misfits indicate that the model
1015 cannot explain the observations. Indeed, whatever the age of
1016 the limestones, it appears difficult to follow the views of
1017 McNutt and Menard (1978) and Lambeck (1981a). Spencer
1018 et al. (1987) already have made several points of
1019 controversy regarding the pertinence and the goodness-of-fit
1020 of these models applied to the north-west Tuamotu atolls.

1021 Estimates of the amount of uplift were significantly biased
1022 because based on apparent elevations of limestone reliefs,
1023 not considering the long-term, possibly substantial sub-
1024 aerial erosion and expected reversals in the vertical motion
1025 trends since the end of the uplift events. As emphasized by
1026 Jarrard and Turner (1979), the maximum elevations of the
1027 limestone surfaces can only provide a minimum estimate for
1028 the amount of uplift due to surface lowering since the time
1029 of island emergence. In addition, in the north-west uplifted
1030 Tuamotu islands, there are found distinct generations of
1031 raised reef limestones, differing in age, state of preservation
1032 and elevation significance. All these observations violate all
1033 the assumptions of the flexural island load hypothesis
1034 proposed previously.

1035 The regional flexural models considered that all of the uplifted
1036 atolls in the north-west Tuamotu have been elevated through
1037 one single continuous event, believed to have been
1038 synchronous at the regional scale, but having started at
1039 different epochs according to different authors: from
1040 Miocene times (Chevalier, 1973), from about 0.30 Ma
1041 (McNutt and Menard, 1978) or from an indeterminate time
1042 coeval with notches at present at expected elevations of +35
1043 or +55 m at Makatea (Lambeck, 1981a). Actually, if uplift
1044 would have resulted from the load-induced flexure by Tahiti
1045 island, the required volcanic critical mass required to
1046 generate the flexure could not be reached not prior to 1 to
1047 about 0.7 Ma, when both shield volcanoes (Tahiti-Nui,
1048 Tahiti-Iti) were about to be built. Moreover, according to
1049 McNutt and Menard (1978), the uplift would be still active
1050 as controlled by the recent building of the nearby Mehetia
1051 Island.

1052 Furthermore, these models were static, based on the currently
1053 assessed rheologic properties of the lithosphere, while the
1054 latter have probably evolved through time due to changing
1055 geodynamic adjustments, in response of changes in a
1056 number of environmental parameters through time: island
1057 configuration due to denudation, geographic location in
1058 relation to plate motion and changes in volcanic activity.
1059 This has probably led to an incorrect and rigid perception of
1060 the tectonic evolution of the north-west Tuamotu atolls.

1061 Similarly, in the southern Cook Islands, the emergence of
1062 Mangaia, Atiu, Mitiaro, and Mauke, all volcanic islands
1063 bearing raised reef limestones has been ascribed to the
1064 loading by Rarotonga alone (Lambeck, 1981b) or by other
1065 surrounding volcanic islands, including Aitutaki and
1066 Manuae (McNutt and Menard, 1978). As in the case of the
1067 north-west Tuamotu, troubling inconsistencies emerge when
1068 comparing field data and geophysical assumptions. The
1069 calculated amount of island uplift was unsatisfactory
1070 because of the choice of the supposed maximum elevations
1071 reached by many of the geomorphic features regarded as
1072 significant in terms of paleo-sea levels. A more accurate
1073 determination of the elevations, extent and significance of
1074 these features would have been required for validating the
1075 reference models. Discussing the value and limits of these
1076 models applied to the north-west Tuamotu and southern
1077 Cook Islands, Spencer et al. (1987) and Stoddart et al.
1078 (1990) concluded that the interpretations were not sustained
1079 by the poorly constrained field data and geophysical
1080 specifications. The newly revisited chronological frame of
1081 raised reef limestones in some southern Cook Islands by
1082 Gischler et al. (2022) may help locally resolving this
1083 dilemma. Although the flexural loading model by Lambeck
1084 (1981b) yielded a more consistent fit to known island
1085 heights, large uncertainties remain due mainly to the lack of
1086 information about the amount of solutional surface
1087 downwering (Jarrard and Turner, 1979).

1088 Thus, the regional flexuring models fall short of explaining
1089 uplift in the considered region. These run into difficulties in
1090 the face of the validity of the physical parameters used, of
1091 the distribution along the entire sequence of stratigraphic
1092 data relevant to the timing of volcanic edifice formation and
1093 of periods of limestone deposition. Based on studies from
1094 Hawaiian Islands, the flexure loading model (Zhong and
1095 Watts, 2002) hypothesizes the existence of an axi-
1096 symmetrical disc load, able to generate significant uplift,
1097 particularly, of the islands located *downstream* of the load,
1098 i.e. in the direction of the plate motion. Within the
1099 considered French Polynesian area, the test is to view if the
1100 regional flexure is accompanied by upwarping of the islands
1101 located further north-west of the loading locus (Tahiti) at
1102 similar distances than the uplifted, north-west Tuamotu
1103 islands. According to the presumed flexural wavelength of
1104 the local lithosphere from the locus to the arch (245 km
1105 from the centre of Tahiti to Makatea), most of the north-
1106 western Society Islands surrounding Tahiti would lie along
1107 the flexural bulge supporting Makatea and the other uplifted
1108 north-west Tuamotu islands. In other words, some of these
1109 islands, including Bora-Bora, Maupiti and Tupai, located at
1110 the respective distances of 275 to 300 km downstream of the
1111 Tahiti-Moorea complex should have been also uplifted. In
1112 addition, the other nearby islands (Moorea, Maiao, Tetiaroa)
1113 residing within the flexural half-wavelength of the Tahiti
1114 loading centre should be also uplifted. By contrast, it is
1115 observed that, in all of the north-west Society Islands near
1116 Tahiti, the subsidence caused by island loading clearly
1117 dominates over uplift. The general trend of the
1118 northwesternmost Society Islands appears to correspond to
1119 long-term subsidence. The rates were estimated at Bora-
1120 Bora from the present depth of MIS 5e reef tracts, between
1121 about 4 m and 30 m below pmsl. These ranges from 0.030 to
1122 0.14 mm/yr (Gischler et al., 2019), consistent with the mean
1123 values (0.09 mm/yr) proposed by Menard (1983). Similar
1124 subsidence rates were calculated at Raiatea (0.12-
1125 0.14 mm/yr). The subsidence of the islands is quite rapid
1126 and, as that, is believed to have been accelerated, due to a
1127 local load-induced stress flexure (Menard, 1983).

1128 This independent observation disproves the idea that the
1129 loading-induced flexure may have caused a differential
1130 uplift of several tens of meters on its north-eastern flank,
1131 while its northwestern side extending in the direction of the
1132 plate motion, *downstream* of Moorea and Tahiti, and would
1133 have escaped large-scale uplift. This is not consistent with
1134 the loading flexure scheme as established from the Hawaiian
1135 islands (Walcott, 1970; Watts and Cochran, 1974; Watts and
1136 ten Brink, 1989; Grigg and Jones, 1997; Zhong and Watts,
1137 2013; Huppert et al., 2015; Klein, 2016) in which Lanai,
1138 Molokai and Oahu, the closest islands to the flexure-
1139 centered Hawaii Island have been significantly uplifted in
1140 the direction of the plate motion.

1141 At Makatea, the amount of uplift inferred from present-day
1142 elevations of raised limestones, net of ground-lowering, is
1143 an order of magnitude significantly larger than the presumed
1144 flexural effect which predicts elevations of only a few tens
1145 of metres (McNutt and Menard, 1978; Watts and ten Brink,
1146 1989; Grigg and Jones, 1997; Huppert et al., 2015). Even if
1147 the lithospheric flexure induced by the load of the volcanic
1148 Tahiti-Moorea-Mehetia complex had contributed later to
1149 the uplift of the northernmost Tuamotu islands, its
1150 contribution would have been at the least modest. The
1151 resulting loading flexure seems to be unable to produce the
1152 observed uplift at Makatea and other surrounding atolls. As
1153 claimed by Minshull and Charvis (2001), load densities of
1154 the oceanic islands have been systematically overestimated
1155 in studies of lithospheric flexure.

1156 **5. Regional uplift mechanism: the hotspot**

1157 rejuvenation model

1158 Aside from the lithospheric load-induced flexure by nearby
1159 younger islands, several competing mechanisms have been
1160 put forward to explain midplate ocean-island uplift in a
1161 context of fast-moving plate. These mechanisms may
1162 contribute in greater or lesser degree to the observed uplift
1163 trends. For uplift acting at broad regional scales, hotspot-
1164 swell growths by either spreading of melt residue or
1165 dynamic topography were regarded as the most plausible
1166 mechanisms (McNutt and Judge, 1990; Ribe and
1167 Christensen, 1994; Morgan et al., 1995; Zhong and Watts,
1168 2002; Ramalho et al., 2010b; Wilson et al., 2010) or
1169 incremental dynamic pressure of buoyant plume material
1170 diapirically impinging at the base of the underlying
1171 lithosphere as well (Davies, 1988; Sleep, 1990; Hoggard et
1172 al., 2020).

1173 At smaller regional scales, uplift may be generated by flexural
1174 uplift induced by subsurface loads – crustal *underplating*
1175 loads; Watts and ten Brink, 1989; Ali et al., 2003. However,
1176 these uplift mechanisms still act upon a wide area which
1177 largely depends on large-scale plate rheology and thus
1178 cannot be used to explain contrasting uplift histories for
1179 edifices spatially close together (Ramalho et al., 2010a,
1180 2010b, 2010c) as in the north-west Tuamotu. Flexural
1181 rebound driven by mass wasting or erosion was also
1182 regarded as a potential uplift process (Smith and Wessel,
1183 2000; Menéndez et al., 2008). Tahiti-Nui has experienced
1184 giant flank landslides and mass wasting (about 1,000 km³)
1185 along its northern area (Hidebrand et al., 2006; Gargani,
1186 2020) spreading out over a distance of 40 km in a north-
1187 north-west direction. These events have not been effective
1188 enough to account for atoll uplift in the north-west Tuamotu.
1189 Significant uplift by erosive flexural rebound is problematic
1190 because it requires large volumes of eroded/mass wasted
1191 material transported over large distances, or because the
1192 effects of redistribution of wasted materials over wider areas
1193 need to be accounted for (Smith and Wessel, 2000).

1194 Another alternative mechanism has been proposed by Menard
1195 (1973), Detrick and Crough (1978) and Crough (1979,
1196 1984), and, recently by Huppert et al. (2020) that involves a
1197 thermally driven asthenospheric bump as plate motion
1198 carries volcanoes over mantle hotspots. At the scale of the
1199 considered north-west Tuamotu parcel, bathymetric swells
1200 present over spatially well bounded hotspot plumes can be a
1201 source of uplift and subsidence for intraplate, ocean volcanic
islands and seamounts. Thus, thermal uplift due to the upwelling of the Society hotspot plume may offer a viable alternative. This mechanism explains more of the available field data from the uplifted islands at the local scale than the load-induced flexure model. Island emergence is assumed to prominently result from upper-mantle plume-lithosphere interactions, from an ascending mantle plume and associated thermal rejuvenation (reheating) of the underlying oceanic crust beneath the hotspot zone sheared by the high-velocity moving Pacific plate as it produces an asthenospheric bump. The swell support is continuously produced at the hotspot location and penetrative mantle diapirism occurring at or close to the hotspot centre can transfer heat at a rate regarded as almost instantaneously at the geological scale (Crough, 1983). The depth of the 450-500 °C isotherms, at the base of the elastic lithospheric layer under non-reheating conditions, is moved upward, likely to be achieved on short-time scales (Calmant et al., 1986). Associated with hotspot reheating, isotherms are pushed upwards, the oceanic crust is thinned and mantle material replaces denser lithospheric material at some depth causing uplift of the surface. This is supported by large positive gravity and geoid anomalies over the Society hotspot swell zone (Cazenave et al., 1988). These may move vertically with the underlying seafloor as islands override existing swells during plate motion-swell migration (Zhong and Watts, 2002). According to the hotspot rejuvenation model, the thermal state of the oceanic upper mantle depends mainly upon whether, when and how close to the hotspot any area of the sea floor passes over (Carlson and Johnson, 1994).

The hotspot plume strength, i.e. the flow of heat material from the mantle, can be estimated from the measure of buoyancy fluxes, computed from swell morphology. The Society
1235 hotspot plume is the strongest in French Polynesia with
1236 buoyancy higher than 1 Mg·s⁻¹ (Adam et al., 2010: 1.53–
1237 1.58Mg·s⁻¹; King and Adam, 2014: 1.79–2.72 Mg s⁻¹) to
1238 be compared with the lowest buoyancy values of the
1239 Tuamotu (0.53-0.72 Mg·s⁻¹) (Adam et al., 2010). In
1240 addition, lithospheric low elastic thickness was related to
1241 patterns of mantle convection (Cazenave et al., 1987),
1242 interpreted as indicative of thermal rejuvenation (McNutt
1243 and Menard, 1982; Calmant and Cazenave, 1986).

1244 There is no independent evidence to either support or refute
1245 thermal rejuvenation. However, some authors have favoured
1246 a dynamic origin (Courtney and White, 1986; Liv and
1247 Chase, 1989) rather than a thermal one (Dietrick and
1248 Crough, 1978; von Herzen et al., 1982) in which upwelling
1249 mantle causes uplift effects (Stein and Stein, 1992; Adam et
1250 al., 2010) in agreement with the low effective elastic
1251 thickness (15–25 km) of the lithosphere near the Society
1252 hotspot (Patriat et al., 2002).

1253 The magnitude of uplift is likely to depend upon the intensity
1254 and duration of eruptive pulses, itself controlled by
1255 variations in mantle melting. Only a large and persistent
1256 mantle plume activity is thought to release larger volumes of
1257 lava flows (Delinger, 1997; Foulger, 2007). Over the last
1258 5 Myr, since the building of Maupiti, the highest intensive
1259 period of volcanism coincides with the formation of Moorea
1260 and Tahiti edifices. The main phase of uplift in the north-
1261 west Tuamotu therefore is presumed to have operated during
1262 the time of maximum melt and lava production from about 2
1263 to 1 Ma, in relation to the subaerial edification of Moorea
1264 and Tahiti volcanoes.

1265 A number of southern Cook Islands, including Mangaia, Atiu,
1266 Mitiaro, and Mauke, previously thought to have been
1267 uplifted by load-induced lithospheric flexuring, in conflict
1268 with most field data (Spencer et al., 1987) were later
1269 considered to have been uplifted by renewed thermal effects
1270 associated with anomalous volcanism out of the trail
1271 sequence (Dickinson, 1998). Similarly, the uplift of the high
1272 carbonate Henderson Island in the Pitcairn group (south-
1273 eastern end of the Tuamotu) was previously believed to be
1274 the product of a lithospheric flexure under the load of
1275 Pitcairn Island, located some 190 km southwestwards
1276 (Spencer, 1989). Due to the respective volumes of Pitcairn
1277 (about 2,300 km³) and Henderson (about 2,500 km³), it
1278 seems difficult to ascribe Henderson uplift to the mass effect
1279 of Pitcairn alone.

1280 Additional evidence supports the idea that islands and atolls are
1281 elevated as they passed over active hotspot swells. There are
1282 a number of other volcanic islands with raised limestones
1283 and high carbonate islands throughout the Pacific, including
1284 Rurutu and Rimatara in the Austral Island group, Nauru and
1285 Banaba in the central Pacific and Christmas, part of the
1286 northern Line Island group. These are not located in close
1287 proximity to large and massive volcanic insular complexes
1288 and, consequently, could not have been elevated by
1289 lithospheric flexuring caused by volcanic loading. Only an
1290 asthenospheric bump generated by local hotspot swell and
1291 thermal rejuvenation (Calmant and Cazenave, 1986) can
1292 explain the presence of raised Miocene to Pleistocene reef
1293 limestones on Rurutu and Rimatara Islands (Pirazzoli and
1294 Veeh, 1987; Stoddart and Spencer, 1987; Dickinson, 1998,
1295 2000). Rurutu has passed some 50 km north of the active
1296 Arago hotspot about 1.2 Ma ago, while Rimatara has passed
1297 pratically over this hotspot, about 2 Ma ago, far from any
1298 large shield volcano.
1299 **6. Reconstruction of the regional uplift events**

1300 The progressive northwesterly migration of the Pacific plate at
1301 the rate of 11.1 cm/yr has resulted in the passage of the
1302 north-west Tuamotu islands near the south-eastern end of
1303 the Society hotspot swell and thus has subjected them to its
1304 influence. Such islands are believed to have been elevated
1305 within an area whose diameter is equivalent to the width of
1306 the heat source (Carlson and Johnson, 1994), that is about
1307 300 km or larger in the considered region (Nolasco et al.,
1308 1998; Isse et al., 2016). Uplift caused by interactions
1309 between hotspot activity and plate motion is a process that
1310 could produce sequenced uplift phases. In the north-west
1311 Tuamotu, there are two distinct island groups, the
1312 northernmost group (Mataiva, Makatea, Tikehau, Rangiroa,
1313 Kaukura) and the southernmost one (Niau, Anaa) that have
1314 experienced different vertical motion trends, at different
1315 times and at different rates.

1316 Prior to 4 Ma ago, the Society hotspot activity has resulted in
1317 the building of the submarine parts of the Leeward Society
1318 Islands – Maupiti, Bora Bora, Raiatea, Tahaa, and Huahine.
1319 At that time, the northwesternmost atolls were located at
1320 about 450 km further south-east, too far away to be
1321 subjected to the influence of the hotspot swell.

1322 From about 3 to 2.0 Ma, Makatea, and Mataiva were positioned
1323 about 380 to 330 km further south-east than today, close to
1324 the eastern boundary of the Society hotspot swell, that is
1325 approximately close to the present position of Niau. At that
1326 time, the regional hotspot plume was typified by a relatively
1327 low eruptive activity, giving rise to the subaerial cones of
1328 Maupiti, Bora Bora, Tahaa, Raiatea and Huahine. A total of
1329 about 203 km³ of lava flows were produced during the time
1330 interval, at a mean rate of 0.073 km³/kyr. As a response, the
1331 long-term subsidence of Makatea and Mataiva may have
1332 slowed down. A few hundreds of thousands of years later,
1333 Tikehau, Rangiroa and Kaukura have followed the same
1334 movement. These three islands were forced to be
1335 tectonically stabilized. By contrast, Niau and Anaa Atolls,
1336 located further south-east, continued to sink at rates
1337 consistent with the standard plate cooling model.

1338 From about 2 Ma, after some few thousand years of
1339 quiescence, the Society hotspot was reactivated, erupting the
1340 largest lava flow volumes that has never released before in
1341 the region, thus achieving the building of Moorea and Tahiti
1342 volcanoes. At that time, these islands occupied the same
1343 geographic location as that at present occupied by Mehetia
1344 and nearby seamounts (Figs. 15A & C). However, following
1345 Clouard and Gerbault (2008), the Pacific would have
1346 experienced a north-eastward rotating deviation during its
1347 displacement. In addition to the reorientation of the Line
1348 Islands and the Boudeuse Ridge (Schlanger et al., 1984),
1349 some Tuamotu atoll and seamount alignments would have
1350 experienced changes in orientation: from Reao to Napuka
1351 Atolls, from Tureia to Marokau Atolls (Fig. 1). Similarly,
1352 about 2 Ma ago, the fact that the northwesternmost atolls
1353 should have been closer to the eastern boundaries of the
1354 hotspot swell than currently, prior to deviate further east,
1355 cannot be ruled out.

1356 The regional mantle plume maintained high melt production for
1357 about 1.9 million years, from 2.15 to about 0.50 Ma. A total
1358 of about 600 km³ subaerial lava volumes were erupted.
1359 Assuming that the erupted lava flow volumes reflect the
1360 degree and rate of mantle melting underneath the hotspot
1361 site and that maximum lava production due to enhanced
1362 melting dynamics can promote uplift of surrounding islands
1363 (Foulger, 2007), the renewed intensive volcanic activity
1364 should have been responsible for initiating uplift of the
1365 surrounding islands easterly of the Society hotspot swell.
1366 According to the trajectory of the north-west Tuamotu islands
1367 relative to the hotspot, uplift of Makatea, Mataiva, Tikehau,
1368 and Rangiroa islands would have started as soon as these
1369 penetrated into the the hotspot influence zone. Kaukura
1370 stopped subsiding. Makatea and the nearest atolls were some
1371 250 to 450 km respectively east-north-east of Moorea and
1372 Tahiti edifices. Makatea was running over the north-eastern
1373 hotspot zone, just above the axis of the mantle bump, while
1374 the other most northern atolls were situated around 100 km
1375 further east, within the hotspot boundary. As a consequence,
1376 Makatea has experienced a higher-amplitude upwarping,
1377 while Mataiva, Tikehau, Rangiroa, and later, Kaukura,
1378 located more peripherally, have been less raised (Figs. 15A
1379 & C and 16). By approximately 2 Ma ago, Niau and Anaa
1380 were still eccentrically positioned relative to the hotspot
1381 swell and continued subsiding. The Tahiti-Nui volcano
1382 started erupting subaerially from about 1.37 Ma and
1383 continued to build up uninterruptedly until 0.87 Ma, thus
1384 indicating that the hotspot was active permanently over 0.50
1385 Myr.

1386 By 1 Ma ago, the Pacific plate had shifted 110 km farthest
1387 north-west. Moorea was no longer active and most parts of
1388 Tahiti volcanoes were built. Makatea and the closest atolls,
1389 from Mataiva to Rangiroa, were respectively at a distance of
1390 some 300 to 450 km west-north-west of the Moorea-Tahiti
1391 complex, at a distance apparently incompatible with any
1392 flexure-induced uplift driven by the load of both volcanoes.
1393 Kaukura has penetrated the active hotspot swell zone and
1394 commenced to be uplifted. As moving along the eastern
1395 border of the active hotspot zone, all these islands have been
1396 uplifted over the mantle bump generated by the hotspot.
1397 Niau was positioned at about a few tens of km only away
1398 from the external hotspot boundary. As the north-
1399 westernmost atolls have started uplifting, Niau has
1400 continued subsiding. Anaa was still at a distance of about
1401 150-200 km far from the eastern, hotspot boundary (Figs.
1402 15A & C).

1403 After a short quiescence and a partial collapse of the Tahiti-Nui
1404 volcano, between 0.87 and 0.85 Ma, there was a sudden
1405 acceleration of the magmatic production during 0.10 Myr
1406 (Hildebrand et al., 2004). From 0.79 to 0.62 Ma, the
1407 subaerial part of Tahiti-Iti was built up. A second phase of
1408 activity occurred between 0.47 and 0.45 Ma. Has this event
1409 promoted the uplift of Niau? Whatever the answer, the Niau
1410 reef-rim of Pliocene to Early Pleistocene age is significantly
1411 younger than the Makatea pile and the outcropping reef-rim
1412 limestones from Mataiva, Tikehau, Rangiroa, and Kaukura.
1413 This implies that Niau has been uplifted later and thus has
1414 suffered karstic dissolution later, during a shorter period. As
1415 the north-westernmost atolls have started uplifting, Niau is
1416 thought to have continued subsiding (Fig. 16), thus possibly
1417 allowing successive Early-Mid Pleistocene reef generations
1418 to be superimposed. After uplift, subaerial erosion may have
1419 operated over a time span of 0.60 to 0.45 Myr, likely
1420 resulting in the removing of Pleistocene deposits older than
1421 the Last Interglacial generation and, finally, in the
1422 denudation of the Pliocene bodies. At present, Niau seems to
1423 have reached equilibrium, at least since the Last Interglacial.
1424 This is suggested by the current position of the MIS 5e reefs
1425 at 3-4 m above pmsl, an elevation close to the level
1426 regarded as eustatic in origin for the period (Kopp et al.,
1427 2009; Deyer et al., 2021), assuming the original reef
1428 surfaces having been lowered by a few meters due to karstic
1429 solution.

1430 From 0.3 Ma, the rejuvenation of volcanic activity at the
1431 hotspot site resulted in the building of Mehetia Island
1432 (Cheminée et al., 1989). Anaa Atoll occupied a position
1433 substantially similar to present, very close to the axis of the
asthenospheric bump (Figs. 15B & D). The occurrence of a
reef-rim, strictly made up of a Last Interglacial reef unit at a
height similar to that of Niau, gives evidence of a late
entering of the atoll into the influence zone of the hotspot
swell. This event may have produced coevally a cessation of
sinking of the atoll, if not a slight upraising. At the same
time, in response to deflation of the mantle bump, the other
north-western Tuamotu islands are likely to have
commenced re-subsiding (Fig. 16).

It is noteworthy that the asthenospheric bump is likely
asymmetrical in form and effect, only efficient in terms of
uplift on its eastern and north-eastern flank. Indeed, Tahiti,
the closest western Society island, is clearly not affected by
this bumping, not exhibiting any evidence of uplift, rather
the opposite. Since at least the Last Interglacial times, Tahiti
was shown to be subsiding at a rate ranging from 0.25
mm/yr (Thomas et al., 2012) to 0.50 mm/yr (Fadil et al.,
2011). Similarly, Moorea appears to undergoing active
subsidence, as suggested by the presence of several
submarine terraces at depths of 13 to 40 m, interpreted to be
Pleistocene reef top surfaces (Montaggioni, pers. observ.).
The bump was apparently not efficient along its western
boundary. This may be explained by the fact that, in addition
to partial deflation, at its western periphery, the related
plume would be stretched and thinned as a result of the
migration of the Pacific plate north-westwards. The
remaining dense and active part of the plume would
maintain further east.

7. Island uplift versus ground lowering

Geomorphic re-examination was undertaken to redefine the
patterns of vertical motion and sub-aerial erosion as a proxy
to geodynamical island evolution. Attention was focused on
1466 age and altimetric distribution of geomorphic features,
1467 including Tertiary reef limestone remnants (*feo*) and diverse
1468 paleo-markers (tidal notches, reef-tracts) of Pleistocene and
1469 Holocene emergence. The present heights of geomorphic
1470 features and surfaces atop of the islands are only minimum
1471 measures of post-depositional uplift. The north-western
1472 uplifted islands have undergone sub-aerial erosion as they
1473 were carried northwestwards by the motion of the Pacific
1474 plate. Estimated maximum elevations of *feo*, reef-tracts and
1475 notches, with added ground-lowering amplitudes, can
1476 therefore be regarded to be valid approximations of net
1477 island uplift.

1478 In the previous models (McNutt and Menard, 1978; Lambeck,
1479 1981a), karstic solution would have been thought to have
1480 just converted uplifted surfaces initially regular and smooth
1481 to jagged surfaces without significant ground lowering.
1482 Whatever the processes that have caused island uplift, how
1483 does one explain the absence of Mid Miocene to Pliocene
1484 marine deposits on Makatea and nearby uplifted atolls
1485 without invoking intense sub-aerial erosion? The only way
1486 to explain such an absence is to ensure that sub-aerial
1487 erosion has the ability of severely lowered emerged
1488 limestone surfaces from originally much higher levels
1489 (Jarrard and Turner, 1979; Paulay and Edward, 1990;
1490 Winterer and Purdy, 2001). At Makatea, this is supported by
1491 both the elevation difference between the deepest internal
1492 basin (20 m) and the highest island point (113 m) and the
1493 current location of phosphates. The elevation difference is
1494 the result of an active meteoric solution at the expense of an
1495 originally flat-topped reef platform. Similarly, the present-
1496 day location of phosphate material implies a redistribution
1497 of the original deposits from bottoms of shallow-water
1498 ponds to deep erosional pot-holes at slope angles greater
1499 than 40°. Such a depositional mode argues for a large-scale
1500 collapse of the wall-rock limestones due to severe karstic
1501 solution.

1502 Deep drilling through subsiding Pacific atolls developed from
1503 pre-Neogene volcanoes revealed that the overlying
1504 carbonate deposits can reach several hundreds of metres. At
1505 Eniwetok, Marshall Islands, the carbonate pile approximates
1506 1,400 m in total thickness (Lincoln and Schlanger, 1991;
1507 Quinn and Saller, 1997). At Rangiroa, based on seismic
1508 records, it was thought to be as thick as 2,000 m thick
1509 (Talandier and Okal, 1987). The sequences from the top of
1510 the Early Miocene to the top of the latest Pliocene are
1511 around 200 m thick at Bikini (Ladd and Schlanger, 1960;
1512 Ladd et al., 1967) and Eniwetok (Lincoln and Schlanger,
1513 1991; Quinn and Saller, 1997) and probably at Rangiroa as
1514 well. At Makatea, the coeval carbonate sequences,
1515 postulated to be of a similar order of thickness, have been
1516 removed within the last 2 Myr. This suggests that karstic
1517 solution should have operated at a mean rate of about 0.10
1518 mm/yr.

1519 According to the methods used, estimates of carbonate surface
1520 loss over time provided conflicting findings, varying by
1521 several orders of magnitude. Values of sub-aerial erosion
1522 rates ranging from 0.16 to 0.10 mm/yr were proposed by
1523 Paulay and McEdward (1990), Winterer and Purdy (2001),
1524 Kennedy et al. (2012), Mylroie and Mylroie (2017) and
1525 Montaggioni et al. (2019) from various exposed tropical reef
1526 sites worldwide. Based on reef drilling, Marshall and Davies
1527 (1984) estimated that the Last Interglacial reef tracts at
1528 present overlain by Holocene reefs in the Australian Great
1529 Barrier Reef have suffered a surface loss at rates of 0.07 to
1530 0.14 mm/yr over the Late Pleistocene to post-glacial low
1531 sea-stand interval. This appears consistent with values
1532 derived from the modeled atoll-rim dissolution factor
1533 (Toomey et al., 2016), assuming a 27.5 m/Myr of vertical
1534 dissolution for each mm/day of rainfall. Accordingly, at
1535 Makatea where mean rainfall is 4.90 mm per day, the rate of
dissolution can be estimated at 135 m/Myr (0.135 mm/yr).
1536
1537 When compared to radionucleide-based erosion rates, all
1538 these values seem to overestimate erosion rates by an order
1539 of magnitude of 2 or 3. Indeed, measurements of
1540 cosmogenic radionucleide ($^{10}$Be, $^{26}$Al, $^{36}$Cl) concentrations
1541 in exposed Pleistocene coral reefs in the Caribbean yielded
1542 significantly lower lowering rates, varying from 0.026 to
1543 0.045 mm/yr. Finally, selecting mean values of ground
1544 lowering of 0.10–0.12 mm/yr for the north-western Tuamotu
1545 carbonates appears to be an appropriate compromise. It is
1546 noteworthy that, at Makatea, lowering is likely to have
1547 operated at varying rates, higher in the deepest inland areas,
1548 where at present elevations do not exceed 20 to 35 m above
1549 pmsl, and lower along the plateau margins at present
1550 peaking between +75 to +113 m.

1551 Applying ground-lowering rates of 0.10–0.12 mm/yr over the
1552 past 2 Myr results in a surface loss of 200–240 m at
1553 Makatea. If correct, given its present maximum elevation
1554 (+113 m), the island would have been uplifted theorically by
1555 about +310 to +350 m relative to present sea level.
1556 Accordingly, over the last 2 Myr, island uplift should have
1557 occurred at mean rates ranging between 0.16 and 0.18
1558 mm/yr. Assuming that the 6-m and 25-m high notches refers
1559 respectively to MIS 5e (about 0.12 Ma-old) and to MIS 7e
1560 (about 0.23 Ma-old), the uplift event between the two high-
1561 stand events lasted approximately 110,000 years, at a mean
1562 rate of 0.17 mm/yr. Moreover, if the 43-m high notch line is
1563 postulated to have formed during MIS 9 (about 320 ka old),
1564 the mean rate of uplift between these two interglaciations
1565 would be 0.20 mm/yr, maybe in relation with the activity of
1566 the emerged Mehetia volcano. Accordingly, it seems
1567 reasonable to consider that uplift has taken place at rates
1568 varying between about 0.18 and 0.20 mm/yr, at times prior to 0.32 Ma, which is during the last 1.68 Myr, in relation with the activity of the emerged Tahiti volcanoes. By contrast, over the last 0.23 Myr, the rate of island uplift appears to have decreased progressively, to be close to zero from 0.12 Ma to present (Fig. 17), presumably with the cessation of activity of Mehetia. Thus, through time, the uplift rates were likely to have varied in response to variations in the activity of the hotspot swell. Since at least the Last Interglacial, Makatea appears to be tectonically stabilized, as evidenced by the current peak elevations of both the MIS 5e and mid-Holocene reefs, respectively close to their expected eustatic, GIA-corrected positions.

1581 As soon as the island has started to emerge, sub-aerial erosion has acted concomitantly with uplift. Accordingly, Makatea has never culminated at about 310-350 m high at any time, due to the interplay between uplift and lowering. Based on rates of the two respective processes, almost 70% of the exhumed reliefs at Makatea appear to have been destroyed incrementally as the island has been rising up. The difference in present altitudes of the raised Miocene limestone reliefs between Makatea and the closest Tuamotu atolls (Mataiva, Tikehau, Rangiroa, Kaukura) is around 100 m. Assuming a constant lowering rate from atoll to atoll in the region, this indicates that, from the central axis of the bump, the adjacent flanks are very gentle dipping eastwards (less than 0.1°).

1595 Niau and Anaa exhibit Last Interglacial reef flats that peak at about +3 to +5 m high, at elevations regarded as close to their initial eustatic level. Assuming a lowering at a mean rate of 0.10-0.12 mm/yr and a resulting surface loss of about 10-12 m since emergence suggests that both atolls have experienced a very slight uplift, if any, before reaching equilibrium. Their current steady state is evidenced by the
occurrence of exposed mid-Holocene fringing reef flats and associated features at heights of around 1 m, as close to those observed on Makatea and other islands throughout the region.

Although uplift was demonstrated to result regionally from a different process, a similar decreasing trend in uplift rates were observed in the Hawaiian Archipelago: 0.06 to 0.02 mm/yr at Oahu over the last 500 ka; 0.19 to 0.04 mm/yr at Molokai over the last 300 ka (Grigg and Jones, 1997; McMurtry et al., 2010), probably resulting from recent lithospheric relaxation (Klein et al., 2016).

It is noteworthy that the uplift of Makatea seafloor at mean rates of about 0.20 mm/yr poses important constraints on the dynamics of the Society mantle plume, especially on the effects of plume-plate interactions. A long-term uplift as rapid as 0.20 mm/yr is possible only if the hotspot heat flux is ten times normal (Zhong and Watts, 2002), strongly supporting the thermal rejuvenation hypothesis.

Finally, it should be noted that karstic erosion has acted differentially across the carbonate pile, at various rates, from margins to interiors. If originally, the considered islands have been flat-topped or very shallow-water carbonate platforms, the end-products are today reef-rimmed structures with central basins. The margins composed mainly of coralgal boundstones have been less sensitive to solution than the internal parts dominated by skeletal detritus, from grainstones to mudstones, as observed at Makatea (Montaggioni, 1985).

## 8. Post-uplift stabilization and subsidence

The expected scenario of plate motion over hotspot swells presented herein involves subsidence resumption following emergence (Fig. 16). After rapid uplift along the steeper
1634 flanks of the bump, upstream of the hotspot zone, once each
1635 increment of the chain has passed beyond the influence of
1636 mantle reheating, the reheated crust cooled and thickened,
1637 accompanied by thermal contraction and gradual subsidence
1638 along the gentle dipping, downstream side. Depending on its
1639 new thickness related to its new thermal age, any reheated
1640 oceanic crust was thought to re-subside at unusually rapid
1641 rates, compared to those typifying a standard cooling plate
1642 (Crough, 1978, 1984; Detrick and Crough, 1978; Calmant
1643 and Cazenave, 1986; Huppert et al., 2015). This scenario is
1644 that observed in the most northwestern Tuamotu area.

1645 Vivid evidence for subsidence resumption is the lack of
1646 exposed to Late Pleistocene reefs on the northwesternmost
1647 atolls, indicating that the uplift episode has been interrupted
1648 and these atolls have started re-subsiding rapidly (Fig. 16).
1649 The MIS 5e and older highstand reefs are expected to be at
1650 present submerged, in the form of reef aprons lying against
1651 the submarine atoll flanks, as shown by the deposition of
1652 Last Interglacial reef remains as mega-blocks along the
1653 northern Mataiva coast (Fig. 6B). The 18 m-depth terrace
1654 found at base of the 7m-high cliff on the northwestern
1655 foreslope at Tikehau may be indicative of a former (Late
1656 Pleistocene?) higher sea stand. Assuming that MIS 5e
1657 features in the most north-western Tuamotu atolls are
1658 currently submerged at depths greater than 10 m below
1659 pmsl, the inferred subsidence rate approximates 0.15 mm/yr.
1660 This value is significantly greater (by a factor of 3 to 10)
1661 than that predicted from the standard cooling curve of about
1662 65 to 60 Ma-old oceanic crusts (Parsons and Sclater, 1977:
1663 0.045-0.035 mm/yr; Dickinson, 1998: 0.014-0.020 mm/yr).
1664 Such a current rapid drowning could be ascribed to
1665 relaxation of the ocean crust as the plate and
1666 northwesternmost atolls have drifted away from the hotspot
1667 zone. However, it is difficult to define the time at which the
1668 most north-western atolls, Mataiva to Rangiroa, have started
1669 to subside again, presumably not before 0.30 Ma, at the time
1670 when the rate of uplift at Makatea has dropped off rapidly
1671 (Fig. 16). On this island, the closest to the active hotspot
1672 locus, the uplift effects have been recently minimized if not
1673 aborted. As pointed out above, the island is at present
1674 tectonically stabilized at least since about the last 0.12 Myr.
1675 .

1676 Once the two southernmost atolls (Niau and Anaa) have
1677 entered the bump zone, subsidence has stopped at a time
1678 presumably just prior to the settlement of MIS 5e reefs. This
1679 is supported by the presence of exposed Last Interglacial
1680 reef flats and the absence of any older Pleistocene reef
1681 outcrop (Fig. 16)

1682 9. Conclusions

1683 The so-called *uplifted*, northwest Tuamotu atolls feature clear
1684 evidence of uplift and subsidence based on paleosea level
1685 markers, providing a record through the Cenozoic of atoll
1686 evolution. The overall consistency of the field data with the
1687 newly proposed hotspot rejuvenation (*asthenospheric bump*)
1688 model encourages confidence in it. The interpretations based
1689 on available field data depart markedly from the previous
1690 modelled expectations about the role of lithospheric load-
1691 induced flexure by the Tahiti–Moorea–Mehetia complex in
1692 island emergence. The flexure model does not account for
1693 the observed field patterns and magnitudes of vertical island
1694 motions in both the north-western Tuamotu and the north-
1695 western Society islands close to Moorea and Tahiti. In
1696 agreement with the flexure model, the islands of Huahine,
1697 Raiatea, Tahaa, and Bora Bora, at distances from Tahiti
1698 similar to those of the north-west Tuamotu islands would
1699 have been uplifted, while these are rather rapidly subsiding
1700 since their formation.

1701 A key issue is that of discrepancies in *feo* (raised limestones)
1702 distribution, nature and age along the northern *versus*
1703 southern sections of the uplifted Tuamotu atoll belt. Except
1704 for Makatea which displays both emergent, Early Miocene
1705 to Late Pleistocene reefal remnants, the most northern atolls
1706 (Mataiva, Tikehau, Rangiroa, and Kaukura) exhibit
1707 limestone exposures of Miocene age only, apparently
1708 covered directly by Holocene coralgal veneers. Conversely,
1709 the reef-rims of the most southern Niau and Anaa Atolls are
1710 made up of Plio-Pleistocene and/or Late Pleistocene reef
1711 limestones and lack any Miocene reef exposure. This clearly
1712 reveals that the two islands groups have experienced
1713 different vertical motion trends at different times. The most
1714 northwestern group has experienced probably a nearly
1715 continuous linear uplift of several tens of metres at a mean
1716 rate of about 0.20 mm/yr prior to slow down from the last
1717 0.30 Myr. The main uplift events are presumed to have
1718 occurred at the time when the hotspot swell has been fully
1719 active through strong eruptive pulses that have given rise to
1720 the building of Moorea and Tahiti volcanoes from 2 to about
1721 0.50 Ma, when Makatea and the nearby atolls passed over
1722 the stationary heat source of the Society hotspot. Consistent
1723 with an initial uplift rate of 0.20 mm/yr, Makatea is
1724 presumed to have been theoretically elevated to a maximum
1725 of about 310-350 m above pmsl, but incrementally
1726 compensated for by a denudation rate of about 0.10-0.12
1727 mm/yr.

1728 Moreover, the Makatea record suggests that, later, over at least
1729 the past 0.30 Myr, the uplift rate has progressively decreased
1730 from 0.20 mm/yr to zero. Probably at the same time, after
1731 cessation of uplift, the atolls closer to Makatea have
1732 commenced re-subsiding at rates of around 0.15 mm/yr,
1733 significantly higher than those predicted by the standard
1734 crust cooling curve, so that Pleistocene highstand reefs and
1735 associated geomorphic features have been submerged at
1736 least since the Mid Pleistocene. By contrast, the most
1737 southeastern group, including Niau and Anaa, has
1738 undergone a reversal in the subsidence trend, resulting in
1739 maintenance of the two islands under relatively stable
1740 conditions. This could have been caused by the renewal of
1741 volcanic activity at the hotspot site over the last 0.30 Myr
1742 that gives rise to sub-aerial Mehetia volcano and nearby
1743 submarine volcanoes.

1744 The asthenospheric bump model explains satisfactorily the
1745 successive evolutionary events that have driven the
1746 geodynamical history of some northwestern Tuamotu atolls.
1747 The picture of a treadmill, i.e. the drifting Pacific plate,
1748 running over a rejuvenated mantle plume is consistent with
1749 graduated uplift, followed by stability and, finally, by a
1750 comeback to subsidence.

1751 **Declaration of competing interest**

1752 The authors declare that there is no competing financial interest
1753 or personal relationships liable to influence the work
1754 published herein.

1755 **Acknowledgements**

1756 The field work was partly financially supported by Fondation
1757 de France (2019-08602), ANR-19-CE34-0006-Manini and
1758 ANR-19-CE14-0010-SENSO. It has benefited from
1759 logistical support from a number of local institutions and
1760 people. Special thanks go from the leading author (LFM) to
1761 the following ladies and gentlemen: to Prof. Motoharu
1762 Koba, Kansai University, Japan, for ESR dating of Makatea
1763 samples; at Makatea, in October 2021, to Philippe Bacchet,
1764 Norbert Apicilla, Pascal Lecointre, Jacky Ioane, Sylvanna
1765 Nordman Haoa; at Mataiva in October 2021, the owners of
1766 Mataiva Village Hotel; at Niau in April 2022, to Philippe
1767 Bacchet for photo acquisition; at Anaa in May 2022, to
1768 Patricia Frébault, Joël Dexter, Michel Muraton and Rustik
1769 and at Kaukura in May 2022, to Philippe Bacchet, Ludia and
1770 Robert Bellais for rock collection and photo acquisition. BS
1771 thanks Annie Aubanel and Paul Niva who helped collecting
1772 samples on Rangiroa and Tikehau in April-May 2021, and
1773 Guillaume Molle and Vahine Rurua who organized
1774 archeological research on Niau in October 2022 under the
1775 auspices of CIRAP (International Center for Archaeological
1776 Research in Polynesia). The authors are indebted to Prof.
1777 Colin Woodroffe and an anonymous reviewer for revising
1778 an early version of the manuscript.

1779 Author contribution

1780 LFM designed research, conducted field work and sampling
1781 collection at Makatea and Mataiva. BS collected rock
1782 samples at Tikehau, Rangiroa and Niau. EPB performed U-
1783 Th and isotope Sr dating analyses and interpretations in
1784 association with AD. FF participated in description and
1785 interpretation of rock facies and provided data on island
1786 volumes and lava production rates. BMG participated in the
1787 data interpretations, designed and provided illustrations. LM
1788 provided rock thin-sections. LFM wrote the manuscript with
1789 input from EPB, FF and BMG.

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2544 **Figure captions**

2545 **Fig. 1.** Structural and location map of the Society and Tuamotu
2546 Islands, French Polynesia, central South Pacific (from
2547 Google Earth map), showing the volcanic island-seamount
2548 chains mainly oriented N115°–120° according to the motion
2549 direction of the Pacific lithospheric plate. The Line Island
2550 Group and the Boudeuse Ridge are regarded as belonging to
2551 the same structural system as the Tuamotu, but having
2552 experienced a rotating motion north-eastward during north-
2553 westward plate displacement in response to shearing stresses
2554 suffered by the plate at its western boundary. Note that some
2555 Tuamotu chains have also undergone a rotating motion
2556 north-eastward, including the Reao–Pukapuka, Tureia–
2557 Marokau, Pinaki–Nukutavake bearing chains (yellow lines).
2558 The islands with raised reef limestones (*feo*) are green-
2559 coloured.
2560 Symbol letter codes of islands and seamounts: An: Anaa; Be:
2561 Bellinghausen; Bo: Bora Bora; Cr: Crough; Du: Ducie; He:
2562 Hereheretue; Hu: Huahine; Ka: Kaukura; Ga: Gambier; Hn:
2563 Henderson; La: Larson; Ma: Maupiti; Me: Mehetia; Mk:
2564 Makatea; Mo: Moorea; Mr: Marokau; Mu: Moruroa; Mv:
2565 Mataiva; Ni: Niau; Nu: Nukutavake; Oe: Oeno; Pi: Pinaki;
2566 Pt: Pitcairn; Pu: Puka Puka; Ra: Rangiroa; Re: Reao; Ri:
2567 Raiatea; Sc: Scilly; Ta: Tahaa; Te: Tetiaroa; Th: Tahiti; Ti:
2568 Tikehau; Tp: Tupai; Tu: Tureia
2569 **Fig. 2.** Deep crustal structure beneath Tahiti and Rangiroa Atoll
2570 (western part of the Tuamotu Plateau) showing the three
2571 layers making up the oceanic crust underlain by the upper
2572 asthenospheric mantle. The numbers refer to mean layer
2573 densities. Modified from Patriat et al., 2002.

2574 **Fig. 3.** Location map and age of the raised reef limestones from
2575 the north-western Tuamotu Islands, with location of cross-
2576 sections in Figures 5 and 8.

2577 **Fig. 4.** Views from Makatea Island, Tuamotu. A/ Aerial view
2578 of the cliff, showing three notch lines, respectively at about
2579 1–2 m, 6–8 m, and 25–30 m above pmsl, east-south-east of
2580 Moumou Bay. Locally, there are late and mid Pleistocene
2581 apron-reef remnants deposited against the lower parts of the
2582 cliff – photography courtesy of P. Bacchet. B/ Cliff,
2583 showing the uppermost notch lines, at 25–30 m and 43–
2584 45 m above pmsl, north-north-east of Moumou Bay –
2585 photography courtesy of P. Bacchet. C/ Exposed mid-
2586 Holocene reef flat, Temao area, north-north-west coast –
2587 photography by L. Montaggioni. D/ View of the 6–8 m high
2588 notch and the late and mid-Pleistocene reef aprons, Moumou
2589 Bay, north-north-east area – photography by L.
2590 Montaggioni.

2591 **Fig. 5.** Cross-section of the coastal cliff and adjacent forereef
2592 slope, Teapoto area, north-north-east of Makatea Island,
2593 showing the successive geomorphic features. Note the
2594 occurrence of a submarine platform at about 9–11 m deep,
2595 interpreted as erosional in origin and formed during a
2596 Pleistocene sea stand lower than today. See Figure 3 for
2597 location.

2598 **Fig. 6.** Views from Mataiva, Tikehau and Rangiroa Atolls,
2599 Tuamotu. A/ Raised remnant of the expected Miocene
2600 carbonate pile – namely, *Rocher de la Tortue*, Turtle Rock –
2601 , northern reef-rim area – photography courtesy of P.
2602 Bacchet. B/ Mega-block dated from the last interglacial
2603 stage, trapped into the modern beach, northern reef-rim area
2604 – photography by L. Montaggioni. C/ Raised remnants of
2605 the expected Miocene carbonate pile, south-south-west reef-
2606 rim area, Tuherahera area, Tikehau Atoll – photography by
2607 B. Salvat. D/ Raised remnants of the expected Miocene
2608 carbonate pile, southern reef-rim area – *Îles aux Récifs*
2609 area –, Rangiroa Atoll – photography courtesy of P.
2610 Bacchet.

2611 **Fig. 7.** Reconstructed subsurface structure of Mataiva Atoll,
2612 Tuamotu, based on lithology and radiometric dating of
2613 drillholes – see Figure 3 for location of drillholes. Modified
2614 from Pirazzoli and Montaggioni, 1986.

2615 **Fig. 8.** Morphology of the upper forereef slope, west-north-
2616 west coast (Pahitomo area), Tikehau Atoll, showing the
2617 occurrence of a submarine cliff ending into a sub-horizontal
2618 terrace, between 11 m and 18 m deep. This feature is
2619 interpreted as a former sea stand, presumably of Late
2620 Pleistocene age. See figure 3 for location. Modified from
2621 Harmelin-Vivien 1985, Fig. 55.

2622 **Fig. 9.** Reconstructed subsurface structure of Tikehau Atoll,
2623 Tuamotu, based of lithology (Rougerie and Wauthy, 1993;
2624 Rougerie et al., 1997) and radiometric dating of the drillhole
2625 extracted from a lagoonal reef-patch. See Figure 3 for
2626 location.
2627 **Fig. 10.** Views of raised reef limestones from Kaukura and
2628 Niau Atolls, Tuamotu. A/ Raised remnants of the expected
2629 Miocene carbonate pile, south-west reef-rim, area Kaukura
2630 Atoll - photography courtesy of P. Bacchet. B/ Northern
2631 cliffed reef-rim of Early Pleistocene age, Niau Atoll -
2632 photography by B. Salvat. C/ Pinnacle-shaped remnant of
2633 Pliocene age, central part of the reef-rim, Niau Atoll -
2634 photography courtesy of Philippe Bacchet. D/ Exposed last
2635 interglacial reef tract, eastern outer reef-rim, area, Niau
2636 Atoll - photography courtesy of P. Bacchet.

2637 **Fig. 11.** Views of exposed reef limestones from Niau and Anaa
2638 Atolls. A/ Exposed Mid-Holocene algal crest, southern outer
2639 reef-rim area, Niau Atoll - photography by B. Salvat. B/
2640 Exposed remnants of the Last Interglacial reef tracts, south-
2641 south-west reef-rim area, Anaa Atoll - photography courtesy
2642 of P. Bacchet.

2643 **Fig. 12.** Thin-sections photomicrographs from raised
2644 limestones on north-west Tuamotu Islands. A/ Calcified,
2645 foraminiferal wackestone, dominated by miliolids, with
2646 *Borelis*, centre of the upper plateau, Early Miocene,
2647 Makatea Island. Early Miocene. Scale = 500 µm B/
2648 Dolomitized, coral boundstone, north-east, outer margin of
2649 the upper plateau, Makatea Island. Early Miocene. Red
2650 alizarine coloured: white surfaces refer to dolomicrite. Scale
2651 = 200 µm. C/ dolomitized, coral boudstone, *Rocher de la
2652 Tortue*, north reef-rim, Mataiva Atoll. Miocene. Note the
2653 occurrence of scattered dolomite rhombs. The original
2654 texture is partly obliterated by dolomite growth. Scale = 500
2655 µm. D/ Dolomitized, coralline-foraminiferal wackestone,
2656 showing an altered *Austrotrillina* test, south-south-west reef-
2657 rim area, Tikehau Atoll. Miocene. Scale = 200 µm.

2658 **Fig. 13.** Thin-section photomicrographs from raised limestones
2659 on north-west Tuamotu Islands. A/ Dolomitized coral
2660 boundstone, south-south-west reef-rim area, Tikehau Atoll.
2661 Miocene. Scale = 1 mm. B/ Dolomitized coral boundstone,
2662 south reef-rim area, Rangiroa Atoll. Miocene. Scale = 1 mm.
2663 C/ Weakly dolomitized coralline-foraminiferal grainstone,
2664 south reef-rim area, Rangiroa Atoll. Miocene. Scale = 500
2665 µm. D/ Dolomitized coral boundstone, north-north-west
2666 reef-rim, central area, Kaukura Atoll. Miocene. Scale = 1
2667 mm.

2668 **Fig. 14.** Thin-section photomicrographs from raised limestones
2669 on north-west Tuamotu Islands. A/ Weakly dolomitized
2670 coralline-foraminiferal grainstone, south-east reef-rim area,
2671 Kaukura Atoll. Miocene. Scale = 500 µm. B/ Weakly
dolomitized coralgal boundstone, southern reef-rim end,
2672 Niau Atoll. Pliocene. Scale = 1 mm. C/ Weakly dolomitized
2673 coral boudstone, south-south-east reef-rim, central area,
2674 Niau Atoll. Pliocene. Scale = 1 mm. D/ Coralgal-
2675 foraminiferal boudstone, north-north-east outer reef-rim,
2676 Anaa Atoll. Last Interglacial. Scale = 1 mm.
2677

2678 **Fig. 15.** Interpretative reconstruction of the extended Society
2679 hotspot swell and of the surrounding deep crustal features
2680 through time, with the location of the related asthenospheric
2681 bump and of the so-called uplifted north-western Tuamotu
2682 Islands - Mk: Makatea; Mv: Mataiva; Ti: Tikehau; Ra:
2683 Rangiroa; Ka: Kaukura; Ni: Niau; Aa: Anaa. A/ B/ Expected
2684 areal extent of the influence zone of the hotspot swell, and
2685 of the bumping axis from 2 to 1.50 Ma and from 0.30 Ma to
2686 present respectively. The x-y and x'-y' lines refer to the
2687 location of the cross-sections in Figs. 15 C&D. The light
2688 pink subcircular zones refer to the hotspot influence zone.
2689 The dark pink elliptical areas refer to the axial zone of the
2690 asthenospheric bump. C/ D/ Cross-sections beneath the
2691 north-western Tuamotu area, showing the deep crustal
2692 structure from 2 to 1.50 Ma and from 0.30 Ma to present
2693 respectively. Both cross-sections step through Makatea to
2694 Anaa. These two islands are in the forefront of the drawing,
2695 while the others are located in the background, setting back
2696 from the x-y and x'-y' lines. Note the reduction in size of the
2697 plume and in areal extent of the hotspot zone and in vertical
2698 amplitude of the related bump through time, in relation with
2699 the regional decreasing volcanic activity.

2700 **Fig. 16.** Geodynamical evolution of the so-called *uplifted* atolls, north-west Tuamotu, showing the successive tectonic events – uplift, stabilization, subsidence – and their relative timing over the last 2.5 Myr. The sea-level curve over the period is given.

2705 **Fig. 17.** Geodynamical behavior of Makatea Island over the last 0.40 Myr. Note the regularly decreasing trend in uplift rates from about 0.30 Ma to present and the resulting tectonic island stabilization.
<figure><img src="image_3.png" /></figure>

250 km

140°W

PITCAIRN ISLANDS

Crough Ridge

Oe

Hn

Du

Cr

Pt

GAMBIER ISLANDS

Austral Fracture Zone

He°

Tu

Mu

Ga

Nu, Pi

Re

Mr

An

Ridge
TUAMOTU ISLANDS

Me

Th

Hu

Ri Mo

Bo
Tp

Mw. Ti Ra

Mk

Ka

Ni

SOCIETY ISLANDS

AUSTRAL ISLANDS

20°S

La

11.1 cm yr⁻¹

Line Islands

MARQUESAS ISLANDS

Marquesas Fracture Zone

Pukapuka Ridge

Pu

Ra

Ka

Ni

Te

Hu

Ma

Be

Sc

<figure><img src="image_4.png" /></figure>

250 km
<figure><img src="image_5.png" /></figure>

TAHITI

of MAKATEA

RANGIROA

NNE

VOLCANOES + LIMESTONES

OCEANIC CRUST

UPPER MANTLE

THREE LAYERS CRUST

MANTLE

<figure><img src="image_6.png" /></figure>

Volcano-
sediments

<figure><img src="image_7.png" /></figure>

Basalts

<figure><img src="image_8.png" /></figure>

Gabbros
Puutiare Peak Teapoto

Temao

MAKATEA

2 km

Figure 5

Moumou Bay

Vaiau

2 km

Rocher de la Tortue

<figure><img src="image_9.png" /></figure>

Drillholes

Figure 8

Pahitomo

TIKEHAU

<figure><img src="image_10.png" /></figure>

A

10 km

Figure 9

Tuherahera

<figure><img src="image_11.png" /></figure>

10 km

KAUKURA

<figure><img src="image_12.png" /></figure>

2 km

Avatoru

RANGIROA

<figure><img src="image_13.png" /></figure>

île aux

Récifs

10 km

<figure><img src="image_14.png" /></figure>

2 km

ANAA

<figure><img src="image_15.png" /></figure>

Isolated feo

<figure><img src="image_16.png" /></figure>

Late Pleistocene

<figure><img src="image_17.png" /></figure>

Mid-Pleistocene

<figure><img src="image_18.png" /></figure>

Early-Pleistocene

<figure><img src="image_19.png" /></figure>

Pliocene

<figure><img src="image_20.png" /></figure>

Early Miocene
<figure><img src="image_21.png" /></figure>

<figure><img src="image_22.png" /></figure>

<figure><img src="image_23.png" /></figure>

<figure><img src="image_24.png" /></figure>
ENE

MAKATEA

WSW

<figure><img src="image_25.png" /></figure>
<figure><img src="image_26.png" /></figure>

<figure><img src="image_27.png" /></figure>

<figure><img src="image_28.png" /></figure>

<figure><img src="image_29.png" /></figure>
MATAIVA

<figure><img src="image_30.png" /></figure>

SSW

Modern reef

<figure><img src="image_31.png" /></figure>

BE

Y1

X2

AW

E1

CNE

NNE

WNW

ESE

U7

27C

Motu

Modern reef

<figure><img src="image_32.png" /></figure>

-6,013 ?

-4,234

-8,608

-6,676

-7,361

-12,181

-6,598

305

-6,375 + 5,956

?

+6,188

[33,500

35,250

?

?

1 km

<figure><img src="image_33.png" /></figure>
<figure><img src="image_34.png" /></figure>

Spur-and-groove system
Algal ridge
TIKEHAU

WSW

ENE

<figure><img src="image_35.png" /></figure>
<figure><img src="image_36.png" /></figure>

<figure><img src="image_37.png" /></figure>

<figure><img src="image_38.png" /></figure>

<figure><img src="image_39.png" /></figure>
<figure><img src="image_40.png" /><figcaption>B</figcaption></figure>

<figure><img src="image_41.png" /></figure>
<figure><img src="image_42.png" /></figure>

<figure><img src="image_43.png" /></figure>

<figure><img src="image_44.png" /></figure>

<figure><img src="image_45.png" /></figure>
<figure><img src="image_46.png" /></figure>

<figure><img src="image_47.png" /></figure>

<figure><img src="image_48.png" /></figure>

<figure><img src="image_49.png" /></figure>
<figure><img src="image_50.png" /></figure>

<figure><img src="image_51.png" /></figure>

<figure><img src="image_52.png" /></figure>

<figure><img src="image_53.png" /></figure>
<caption>A</caption>

-2.0 to -1.5 Ma

<figure><img src="image_54.png" /><figcaption>B</figcaption></figure>

-0.3 to 0 Ma

<figure><img src="image_55.png" /><figcaption>C</figcaption></figure>

-2.0 to -1.5 Ma

Motion of the Pacific Plate

Asthenospheric bumping

Uplift

Subsidence

Subsidence

Lithosphere
Upper mantle

Mk

Mv

Ti

Ra

Ka

Ni

Aa

Plume

100 km

Uplift

Subsidence

Not scaled
vertically

<caption>D</caption>

-0.3 to 0 Ma

Motion of the Pacific Plate

Asthenospheric bumping

Uplift-stabilization

Subsidence

Lithosphere
Upper mantle

Mk

Mv

Ti

Ra

Ka

Ni

Aa

Plume

100 km

Uplift-stabilization

Subsidence

Not scaled
vertically
<figure><img src="image_56.png" /></figure>
<figure><img src="image_57.png" /></figure>

Margin of the
upper plateau

Stability

0.17 mm/yr

Decreasing uplifting rate

MIS 11

MIS 13