# Coral reef boulders as recorders of cyclone activity over the last 3500 years (Fakarava Atoll, Tuamotu, French Polynesia)

Lucien Montaggioni, Bernard Salvat, Éric Brunaud, Edwige Pons-Branchu, Gilbert Poli, Bertrand Martin-Garin

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## To cite this version:

Lucien Montaggioni, Bernard Salvat, Éric Brunaud, Edwige Pons-Branchu, Gilbert Poli, et al., Coral reef boulders as recorders of cyclone activity over the last 3500 years (Fakarava Atoll, Tuamotu, French Polynesia). Marine Geology, 2025, 488, pp.107605. 10.1016/j.margeo.2025.107605. hal-05198864

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1 Coral reef boulders as recorders of cyclone activity over the last 3,500 years

2 (Fakarava Atoll, Tuamotu, French Polynesia)

3 Lucien F. Montaggioni a*, Bernard Salvat b, Éric Brunaud c, Edwige Pons-

4 Branchu d, Gilbert Poli e, Bertrand Martin-Garin a

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6 a Aix Marseille Univ, CNRS, IRD, INRAE, Coll France, CEREGE, 13331 Marseille, France

7 b PSL-École Pratique des Hautes Etudes, UMR3278, EPHE, CNRS, UPVD, CRIOBE,

8 Université de Perpignan, 66860 Perpignan, France

9 c Environmental Consultant 44470 Carquefou, France

10 d LSCE/IPSL, CEA-CNRS-UVSQ, Université Paris-Saclay, 91198 Gif-sur-Yvette, France

11 e Environmental Consultant 75016 Paris, France

12 Highlights

13 – Cyclonic storms were the main drivers of mega-boulder detachment and transport, not
tsunamis.

15 – Reef mega-boulders record intensified cyclone activity between approximately 3,400 and 2,600 and 2,000 and 1,700 cal. yr BP, in relation with increasing ENSO activity.

17 Abstract

18 Extreme marine hazard events are expected to increase in the next future as a response to
19 global warming. Estimating their frequency and magnitude in the recent past is of prime
20 concern to better project their future impacts. In tropical reef provinces, coral boulders
21 deposited on reef top surfaces can be used as key proxies for defining the occurrence –
22 recurrence time and effects of paleo-storms. The present study defines origin, morphometry,
23 location, orientation, timing and mode of deposition of a number of coral reef boulders
24 present on reef-flats surfaces, along the south-eastern rim areas of Fakarava Atoll, north-west
Tuamotu, central South Pacific. These boulders, extracted from the upper forereef spurs, were projected onto the reef-flat zone, at distances of 51 to 235 m from the outer reef edge. Among the 44 described boulders > 20 m³, only 19 were revealed to be covered by well-preserved coral colonies in growth position and as such, were submitted to physical analysis and U/Th dating. Upper reef spurs have been built from about 5,000 to 3,000 cal. yr BP. Dominantly platy in shape, with sharp broken, lateral faces, the 19 investigated boulders range between 3 and 8.5 m long, 4 and 82 m³ in volume, 6.8 to 139.4 tons in weight. The minimum flow velocities to initially move and uplifted these blocks were estimated between about 4 to > 13 m/s. Morphometric comparison of the boulder set investigated at Fakarava with those from nearby atolls recognized as cyclone-moved strongly suggests that boulder detachment and deposition at Fakarava have resulted from cyclogenesis than tsunamigenic impacts. Contrary to cyclone-related boulder deposition that appears to have occurred mainly during the last centuries in the nearby atolls, U/Th dating of boulders at Fakarava reveals that cyclogenesis has increased between 3,400 and 2,300 cal. yr BP (15th–4th centuries BC) and between 2,000 and 1,700 cal. yr BP (1st–4th centuries AD) in the considered region, with an expected mean reccurrence time of around 30 to 50 years. This intensification of cyclone activity is interpreted as related to amplified ENSO phases in the south Pacific.

## Keywords

Coral reefs

Mega-boulders

Cyclones

ENSO

Fakarava Atoll

South Pacific.
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# 1. Introduction

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Defining the timing and frequency of marine high-energy events over the late Holocene is an important issue to better evaluate the risks to which low-lying islands, coasts and inhabitants will be exposed in the near future. These areas are known to be particularly prone to inundation triggered by storm- or tsunami-generated wave surges. Particularly, super-storms, including hurricanes, cyclones and typhoons are predicted to usually increase in magnitude and frequency due to future climate changes (IPCC, 2023) thus resulting in increasing reshaping of coastal environments. Erratic rocky coarse debris, also referred to as mega-clasts, mega-blocks, boulders, mega-boulders (Terry and Le Goff, 2014; Ruban et al., 2019) emplaced along coasts and further inland, were used as reliable indicators of past catastrophic flooding events. There are numerous evidences of mega-clasts displacement based on geological records and/or historical archives (See Scoffin, 1993; Onda, 1999; Nott, 1997, 2003a, b; Lau et al., 2016; Canavesio et al., 2023; and references herein). However, the use of boulder deposits for identification of such extreme events has been discussed at some length, since their significance remained controversial (Etienne et al., 2011 and references herein).

Identifying storms or tsunamis as actual causes of boulder emplacement could remain challenging in many cases (Nott, 2003a; Spiske et al., 2008; Barbano et al., 2010; Switzer and Burston, 2010; Frohlich et al., 2011; Weiss, 2012; Terry and Etienne, 2014; Roig-Munar et al., 2017; Cox et al., 2019; Scheffers and Kelletat, 2020; Kennedy et al., 2017, 2021; Weiss et al., 2022; Lario et al., 2023; Roig-Munar et al., 2023; Lloyd et al., 2024).

In the tropics, in response to regional geodynamic and climate constraints, coral reef systems are frequently impacted by tsunami- and/or cyclone-generated high waves. The presence of coral-reef mega-clasts deposited along reef-rim zones is relatively well documented in a number of reef sites throughout the Pacific Ocean. In areas close to lithospheric convergent plate boundaries, i.e. subduction zones, deposition of erratic, huge reef blocks was usually
postulated to have been detached by tsunami waves as observed in Taiwan (Ota et al., 2015), in the Ryukyus, Japan (Noji et al., 1993; Imamura et al., 2008; Araoka et al., 2013; Goto et al., 2010 a, b, 2013; Hisamatsu et al., 2014; Sato et al., 2014, 2019, 2022), in Honshu Island, Japan (Yamada et al., 2014; Iwai and Goto, 2021), in Thailand (Goto et al., 2007), in Tonga (Frohlich et al., 2009; Lavigne et al., 2021), in Fiji (Lau et al., 2018), in Samoa (Etienne et al., 2011), in Sumatra (Paris et al., 2010) although that typhoons could also be involved locally in reef boulder displacement (Taiwan: Nakamura et al., 2014, Terry et al., 2021; Ryukyus: Goto et al., 2009, Minamidate et al., 2020; Fiji: Etienne and Terry, 2012; Terry and Etienne, 2014; Thailand: Terry et al., 2016, 2018; Palau: Hongo et al., 2018; Philippines: Soria et al. (2018). These controversial interpretations highlighted the difficulty of easily recognizing the dynamic drivers of boulder genesis and transport. In areas relatively far from subduction zones, along passive margins such as eastern Australia, most detached coral reef blocks were usually interpreted to have been transported by cyclone-generated wave surges (Nott, 2003 a, b; Nott et al., 2014). However, the role of tsunami-generated waves was not ruled out (Bryant and Nott, 2001). In mid-plate island sittings, accumulations of coral megablocks onto reef-rim surfaces are common features. For instance, large boulders derived from coastal cliffed outcrops of fossil reefs were found along the southern coast of Mangaia Island, South Cook Island Group, projected onto the island platform at an elevation of up to 10 m. A tsunami generated by a submarine slope failure on 13 April 2010 was regarded as the most likely cause of this depositional event (Goff, 2011).

In French Polynesia, there is an extensive literature on the occurrence of coral reef boulders deposited on rim surfaces, especially from the Tuamotu atolls: Rangiroa (Bourrouilh-Le Jan and Talandier, 1985; Talandier and Bourrouilh-Le Jan, 1988; Etienne et al., 2014; Canavesio et al., 2023), Tikehau (Canavesio et al., 2023), Makemo (Etienne et al., 2014; Lau et al., 2014, 2016; Canavesio et al., 2018), Anaa (Etienne et al., 2014; Canavesio et al., 2023), Nukutipipi
99 (Salvat and Salvat, 1992), Hikueru, Marokau and Kauehi (Canavesio et al., 2018), Makatea and Mataiva (Montaggioni, persobs., 2021), Fakahina, Nihiru, and Fangatau (Aubanel, persobs), Tepoto, Napuka and Nukutavake (Salvat, persobs.). Volume and weight estimates of the largest boulders were about 100 to 500 m³ and 200 to > 1000 tons respectively.
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103 Available radiometric dating of boulders usually indicated that most of them have emplaced between the first and twentieth centuries AD. Additional information about clusters of non-dated coral boulder comes from the Society Island Group: Tetiaroa (Collin et al., 2013), Tahiti, Huahine (Etienne, 2012). In most cases, boulder deposition was interpreted as resulting from cyclogenic rather than tsunamigenic sources.
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108 The present paper aims at describing the morphometric attributes (size, volume, weight) and at defining the timing and mode of deposition of coral reef boulders that occur in the form of two distinct depocentres along the south-eastern reef-rim of Fakarava Atoll, northwest Tuamotu. Interestingly, while previous dating of reef boulders in French Polynesia illustrated depositional episodes of the Common Era only, age clustering of the studied boulders reveals that most boulders have emplaced before the beginning of the Common Era, thus giving evidence of local episodes of relatively frequent extreme marine hazard events during the late Holocene. In addition, following works by Goto et al. (2009), Lau et al. (2016), Hongo et al. (2018), Nansadena et al. (2011, 2022), Weiss et al. (2022), analyzing the displacement of coral reef boulders in the study area is a necessary approach to assess the wave flow pattern of past cyclone events.
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119 **2. Environmental site study**
120 **2.1. Atoll location and atoll reef-rim morphology**
121 Fakarava Island is located in northwest Tuamotu, one of the five French Polynesian archipelagoes (central South Pacific Ocean), at 16°18'south and 145°36'west (Fig. 1A).
Extending over about 1,500 km, the Tuamotu Archipelago is strictly composed of low-lying carbonate, coral reef islands (e.g. atolls) in the form of northwest to southeast-oriented linear volcanic chains (Gabrié and Salvat, 1985). Fakarava, as the second largest atoll in French Polynesia after Rangiroa, is about 55 km long and 25 km wide and covers a total marine area of 1,246 km². The lagoon is one of the deepest in the Tuamotu, with a maximum depth of approximately 60 m, and one of the second greatest with a surface area of 1,121 km².

The emerged land zones consist of low-lying, densely vegetated, islets totalizing about 27 km² (Duvat et al., 2020). Rectangular in shape, elongated along a northwest to southeast axis, Fakarava is typified by an asymmetrical reef-rim. This rim is widest on its western side (1,350 m) and narrowest on its northern (750 m), southern (about 450 m) and eastern (250 m) sides.

The islets mainly occur along the north to eastern reef-rim zones in the form of continuous, elongate deposits, about 2.50 to 4–5 m in maximum elevation. The islets are locally cut across by ocean-lagoon, exchange channels (*hoa*), as observed at the south-south-eastern reef rim end. These islets started to develop about 2,000 years ago (Montaggioni et al., 2023b). The reef rim is interrupted locally by two passes, respectively located at its northern (Garuae Pass) and southern (Temakohua Pass) ends. At ocean sides, the northern to south-eastern reef rim areas exhibit reef flats, about 60–80 m wide, while the western to south-western reef rim ones refer to open, tidally exposed reef flats, about 1,350 m in width. Inwards, their surfaces are occupied by dense networks of rubble sheets (Montaggioni et al., 2023b). In the studied areas, along the south-south-eastern reef rim, coral coverage ranges from 80% of the total substrate at the outer reef-flat settings to about 30% at the mid parts. This rate does not exceed 5 to 15% in the innermost reef-flat surfaces where *Acropora*, *Porites* and *Montipora* are present in small ponds.

As observed in other northern Tuamotu atolls (Kuhlmann and Chevalier, 1986; Montaggioni et al., 2019b), at Fakarava, the forereef zone is subdivided, from the uppermost part downwards,
148 by a spur-and-groove system, extending to about 6–10 m deep, and colonized by dense
149 communities of robust branching corals – *Pocillopora damicornis* and *P. verrucosa* – mainly,
150 with subordinate acroporids – *Acropora humilis*, *A. robusta*, encrusting *Montipora* – and
151 merulinids. The spurs from the southeast atoll rim sides range between 7 and 10 m in width
152 and 12 and 20 m in length. This zone is followed by gently sloping terraces to depths of 10–
153 15 m, covered by mixed coral associations dominated by robust branching (*Pocillopora*,
154 mainly), together with some massive (*Porites*) and platy (*Montipora*) forms. From about 20 m
155 deep, there are sub-vertical slopes to depths greater than 60 m, dominated by platy (*Porites*)
156 and foliaceous (*Pachyseris*) coral forms. The emergent, innermost parts of the spurs are
157 capped by an algal crest, about 20 m wide, reaching 0.20 m to 0.50m above low tide level
158 (Kühlman and Chevalier, 1986; Montaggioni et al., 2019b).

159 There are sets of coral reef boulders deposited in several zones of the emerged atoll rim. The
160 largest, composed of boulders of several tens of cubic metres, is found at the south-eastern
161 rim side (Fig. 1B). The others, mainly composed of boulders of only a few cubic metres,
162 occur along the northern rim side, respectively between Rotoava village and the northern
163 (Garuae) Pass and at the south of the village.

164 2.2. *Climate attributes and extreme marine hazards*

165 In the northwest Tuamotu, the climate is influenced by both the trade winds and El Niño-
166 Southern Oscillation (ENSO). The latter controls the activity of tropical to extra-tropical storms
167 (Andréfoüet et al., 2012). The eastward migration of warmer sea surface temperatures during
168 ENSO events results in the displacement of more storms further east in the central South
169 Pacific toward the Tuamotu. Particularly, in the region, the reinforcement of ENSO activity
170 was known to cause increasing cyclogenesis (Chand et al., 2017).

171 Fakarava Atoll is subjected to trade winds blowing from the eastern sectors for about 70% of
172 the year and from the south-east for about 20%. The hot, humid, austral summer occurs from
173 November to April and the relatively cooler and dryer, austral winter from May to October.

174 Air temperature varies between 22-23 °C and 31 °C. Mean annual rainfall ranges around 175 1,500 mm with extreme values between 900 and 2,200 mm. The tide regime, typically 176 microtidal and semi-diurnal, is 0.2-0.5 m in mean amplitude, 0.65-0.70 m at spring tide 177 (Laurent and Maamaatuaiahutapu, 2019; Montaggioni et al., 2023b).

178 During the summer, the trade winds are weak and usually generate low wave heights. Swell 179 direction, period and height are the same as those generated by the winter trade winds.

180 However, the wind regime can be disrupted by either tropical cyclones or depressions, or by 181 distant-source storms originated in the northern and southern latitudes (Canavesio, 2019).

182 During the summer, Fakarava is relatively sheltered from northern distant-source wave surges 183 by northern atolls due to the atoll *shadow effect*, but not from tropical depressions and 184 cyclones passing close to the atoll. Annual average wave height is 1.6 m (Andrefouët et al., 185 2012). Cyclogenesis is active usually during the hot season, and varies significantly in 186 intensity and frequency according to annual ENSO variability. The highest activity time occurs 187 in February with 40% of events (Laurent and Varney, 2014). The cyclone track pattern in 188 French Polynesia is typified by a tapering channel between the Society and the Austral 189 Archipelagoes through which 70% of the cyclone tracks pass (Larrue and Chiron, 2010) west 190 of Fakarava.

191 During the winter, the trade winds generate swells as high as 3-5 m. The region is hit by 192 strong distant-source swells coming from high southern and southwestern latitudes (> 40°).

193 Swell periods are greater than 12 s and heights of up to 4-5 m (Canavesio, 2019; Andrefouët 194 et al., 2023). Episodes of higher amplitude waves (up to 3.75 m) can be promoted by high 195 easterlies winds, western transient storms or distant cyclones (Andrefouët et al., 2012).

196 In the northwest Tuamotu, the pattern of extreme wave hazard events is relatively well 197 constrained during the last 200 years (Laurent and Varney, 2014; Canavesio, 2014; Canavesio
et al., 2023). By contrast, the events from the pre-European times are poorly documented, mainly based on analysis of reef sediment archives. The role of storms on atoll islet morphodynamics over the last 3 to 5 millennia was partly clarified previously in some Tuamotu atolls: Takapoto (Montaggioni et al., 2019a), South-Marutea (Montaggioni et al., 2023a) and Fakarava (Montaggioni et al., 2023b). Similarly, using coral boulders projected onto modern atoll reef tops as records of extreme marine hazard events, Etienne et al. (2014), Lau et al. (2016), Canavesio et al. (2018, 2023) have identified a number of cyclonic events which have hit Rangiroa, Tikehau, Makemo, Kauehi, Marokau and Anaa in the Tuamotu throughout the Common Era (last 2,000 years).

Cyclone activity may have been of greater magnitude during the past than the last century (Lau et al., 2016), especially during the Little Ice Age in the tropical Pacific (Bramante et al., 2020). Over intervals of several centuries, cyclonic events in the northwest Tuamotu have presumably not exceeded 2–3 per century (Canavesio, pers.com.). Non-cyclonic periods longer than 100 years were not uncommon. A total of 24 cyclones only would have hit the Tuamotu atolls over the past 192 years, from 1822 to 2014. Over the 20th century, about a dozen of cyclones and strong tropical depressions have passed across the northwestern Tuamotu region (see Figure 4 in Montaggioni et al., 2023 b: trajectories of tropical cyclones and depressions within the central French Polynesian region since the beginning of the twentieth century). At Fakarava, the January 1903 cyclone, the deadliest in French Polynesia (515 victims), completely destroyed the village and the plant cover in Rotoava area, close to the northern Garuae pass (Tessier, 1977). Storm coral block ridges, 4 m high and 50 m wide, found in south of the atoll rim, between the ancient Tetamanu village and the Tamakohua pass, are expected to have been emplaced during the 1903 cyclone. Although this cyclone was passed within 300 km east of the atoll, it could be able to generate a huge east-southeast swell over Fakarava. The 1905 cyclone has run north of the atoll while that of 1906 has flooded and
devastated once more the village with 5 m-high waves coming from the north (Tessier, 1977).

In 1906, on Anaa Atoll, cyclone-generated waves that have removed the western side of the island over a distance of more than 300 m, were estimated to have reached heights of about 15.5–18.5 m (Canavesio, 2014). The most active period in terms of cyclogenesis relates to the 1982-1983 austral summer, with 5 cyclones (Laurent and Varney, 2014). The last tropical cyclone that affected Fakarava was Orama (category 3: wind speed from 178 to 210 km/h; wave heights from 2.7 to 3.7 m) in February 1983. Swells as high as 10–15 m were reported during this season from the nearby atolls of Takapoto and Rangiroa, while the normal swell amplitudes during the cyclonic seasons range between 4 m and 12 m. The resulting disturbances of this cyclonic season were detailed by Harmelin-Vivien and Stoddart (1985), Laboute (1985), Dupon (1987) and partly reported by Lau et al. (2016). Large reef blocks extracted from outer reef slopes were projected onto adjacent reef flats and lagoons across *hoa*. In addition, coral rubble and skeletal sands were removed, either accumulated in islets or partly eroded from.

However, as most of the other Tuamotu islands, Fakarava has rarely suffered from tropical cyclones and depressions during the historical times (Laurent and Varney, 2014). Cyclone hazard records seem to reveal a 50-year return period for wave exceeding 12 m high (Canavesio, 2014). Fakarava is also known to experience the effects of low to moderate climate events, such as tropical lows, able to generate strong swells (Duvat et al., 2020).

Similarly, the history of tsunamis generated by earthquakes in the Tuamotu and their effects are poorly documented (Talandier, 1980; Etienne, 2012; Lau et al., 2016). According to Lau et al., (2016) and Canavesio et al. (2023), since the 16th century, only a few tsunamis with 0.3 m to 1.9 m-high waves have hit the region. This is in agreement with the findings by Goff et al. (2022) who confirmed, using numerical modelling, that earthquakes from Chile, Peru and Tonga plate boundaries generate tsunami-related waves < 1 m to about 2 m in
248 maximum height, propagating across French Polynesia. For instance, at Tahiti, Schindelé et al.
249 (2006) revealed that regional earthquakes of 7.3 magnitude generate waves of a few centimetres
250 amplitude only and those of 9.5 magnitude generate waves of 1.3 m amplitude, irrespective of the
251 areas of origin of earthquake-related tsunamis. Tsunamis in the Tuamotu would have been
252 significantly less numerous than in the other French Polynesian archipelagoes and suffered
253 less damage (Sladen et al., 2007). Tsunamis related to Aleutian Island 1946, Chile 1960,
254 South Pacific 2009, Chile 2010 and Tohoku-Oki 2011 earthquakes have generated waves not
255 exceeding about 1 m in amplitude. However, it is noteworthy that local to regional tsunamis
256 generated by landslides and collapses of volcano or atoll flanks may have occasionally
257 occurred in French Polynesia (Goff and Terry, 2016) and, more specifically, in the north
258 Tuamotu. This is supported by several features: huge landslides along the submarine flanks of
259 Tahiti shield volcanoes (Clouard et al., 2001; Hildenbrand et al., 2006; Gargani, 2022) and a
260 number of other French Polynesian high volcanic islands (Clouard and Bonnevelle, 2004);
261 scars along the seaward margins of some Tuamotu atolls (Chevalier et al., 1968; Faure and
262 Laboute, 1984; Montaggioni, 1985; Montaggioni et al., 2019b) interpreted as resulting from
263 large-scale slope collapses generated by flank failures; deposition of early Holocene coral
264 clasts onto reef top surfaces, extracted from adjacent reef slopes at a time when sea level was
265 several tens of metres below its present position (Montaggioni et al., 2022). Compared to the
266 Tuamotu, the Marquesas Islands would be hit more frequently and hardest by tsunamigenic
267 waves. However, strongest wave run-ups appear to rarely exceed 4 m high (Paris et al., 2020).
268 Following the post-glacial sea level rise, sea level outpaced its present position by around
269 6,000 yr BP, reached a highstand of about 0.80 m at around 4,000 yr BP prior to drop down to
270 its present position from about 3,000 yr BP (Pirazzoli and Montaggioni, 1988; Halmann et al.,
271 2020).
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**3. Material and methods.**

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Field investigations were carried out during two field trips: March 2019 and August–

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September 2022 along the southeast reef-rim side of the atoll on the so-called Pohatuteroki

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territory (Figs. 1B, 2). Up to 150 boulders of various sizes occurred along a 550 m-long reef-

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flat section. Oriented N96° in its western part, the section rotates to a N125°–N135° direction

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eastwards. Here, the reef rim line is convex-shaped, forming an exposed tip on which swell

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surges will be concentrated and thus increase in energy and height significantly as they

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converge to the reef front (Fig. 3).

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Of the boulders deposited along the south-east reef-flat area, around 60 are clustered into two

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specific zones, namely Zone A at the west side (Figs. 2A, 3) and Zone B at the east side (Figs.

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2B, 3). These boulders exhibit top surfaces no less than 2 m². The largest boulders up to 20

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m³, 44 in number, were subject to measurement of azimuts, directly in the field using a

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magnetic compass or from drone-acquired images. Close examination of their upper surfaces

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revealed that 19 boulders only exhibited well-preserved, coral colonies in upward growth

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position (Fig. 4), thus indicating only these were not deposited in an overturned position after

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detachment and during displacement. They were selected to detailed morphometric analysis

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and U/Th dating.

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**3.1. Measurement of boulder morphometric parameters and location**

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The selected 19 megaclasts were subjects of analysis of morphometric parameters, including

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measurements of longest axis – a-axis in the longest horizontal plane –, shortest axis – b-axis

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in the shorter horizontal plane, orthogonal to the a-axis –, height or thickness – c-axis in the

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vertical plane –, largest surface area in a horizontal plane and apparent volume.

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Measurements were conducted using both 10 to 20-m long tape measures and drone-based

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photogrammetry (Fig. 5) with photogrammetric processing software, Agisoft Metashape® –
version 1.8 on the field, versions 2.0 and 2.1 for subsequent data processing. The drone was a DJI mini 2 type, equipped with a satellite navigation system – Global Navigation Satellite System (GNSS), combining GPS, GLONASS and Galileo –, with 1/2.3" CMOS sensor of 12 M effective pixels. Throughout the field works, the drone was flown manually from 1 to 120 m above the ground level and the investigated boulders. Aerial pictures were taken about 3 m above the ground, using a 1 m-long reference scale, red-marked at both ends and a 2 m-long scale delineated by two tennis balls.

Size and surface areal measurements were refined using place-geometry software. For each investigated megablocks, a high-resolution digital surface model (DSM) and a 3D meshing were generated using Metashape®. The values of the upper surfaces are rounded up to the nearest square-metre.

Owing to its surface flatteness, the reef flat zone and adjacent conglomerate pavements were used as a baseline for the height calculations. The boulder volumes were calculated from the 3D meshing model, with an accuracy of about ± 10%. Finally, the orientation of longest axis was defined using boulder orthomosaic imagery. This data can be refined by measuring the azimuth along the longest boulder axis, from the orthomosaic images using Geogebra.

The distance between the boulder emplacement sites and the reef front line was determined from Google Earth satellite imagery. Due to a slight distortion of images, the error range was estimated at ± 3–5 m.

## 3.2. *Origins of the boulder material*

Identification of the source from which have derived coral reef boulders is an important issue in the study of boulder deposition processes (Etienne et al., 2011). At Fakarava, the rock boulder composition appeared to be quite similar to that of spurs from the uppermost forereef zone, previously described by Bouchon (1983, 1985), Külhmann and Chevalier (1986),
Montaggioni et al. (2019b) and Duce (2016) from various high-energy, reef locations in French Polynesia. Especially, the coral communities present at the upper boulder surfaces consist mainly of robust branching pocillloporids (Fig. 4A), encrusting merulinids (Fig. 4B), associated to massive poritids and branching acroporids. The source zones of boulders resting atop of reef-flat and adjacent zones were recognized previously by Onda (1999), Goto et al. (2009), Etienne et al. (2011, 2014), Lau et al. (2016), Canavesio et al. (2018, 2023). During the field investigations, the location of the upper forereef zones which boulders have been detached from can be clearly identified. Direct observations at low tide by calm sea and drone-acquired imagery allowed recording detailed views of the spur-and-groove system in the considered zones. Fault lines are showed to run parallel to the reef front line and to cut spurs (Fig. 6A). Indivual boulders appear as detached and isolated from adjacent spurs, submerged at depths of 5 to 10 m (Fig. 6B) or exposed on top of growing spurs (Fig. 6C), in a position to start moving and be projected onto the atoll reef-rim during a next high-energy hydrodynamic event. Truncated spur areas from which spur sections have been formerly removed remain visible locally (Fig. 6D).

## 3.2. *Estimates of boulder weight*

Establishing the biological composition, mean porosity and bulk density of boulders is a prerequisite for valubly approximate their weight. At seaward reef margins, reef rocks are coralgal frameworks composed of repeated alternations of encrusting organisms, with secondary pores resulting from the activity of boring organisms, deposition of sediment within primary skeletal cavities and bored pores, and precipitation of intergranular and intra-skeletal cements (Marshall, 1983; Perry and Hepburn, 2008; Macintyre, 1977, 2011). At Fakarava, the studied boulders are pieces made up of robust branching coral colonies (*Pocillopora*, *Acropora*), together with massive (*Porites*) and some minor encrusting (*Montipora*) forms, regarded as primary builders. Minor, secondary reef-building contributors
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are coralline algae, bivalves, gastropod vermetids, foraminifera and bryozoans. Secondary

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builders are trapped into or encrust detrital matrices entirely composed of skeletal rubble and

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rudstone. The resulting framework is firmly cemented.

348

Directly measured skeletal porosities and bulk densities of the above-mentioned coral forms

349

were provided by Lough and Barnes (1997), Bucher et al. (1998), Mwachireya et al. (2016),

350

Cuttler et al. (2017), Prada et al. (2021). Porosity and density of branching *Acropora* colonies

351

average 30–65% and 1.13–1.90 g/cm³. Colonies of branching *Pocillopora verrucosa* and *P.*

352

*damicornis* have porosity of 18–35% and density of 1.85–2.20 g/cm³. Skeletons of merulinid

353

corals (*Favia* and *Favites*) exhibit porosity of 40–69% and bulk densities ranging between

354

0.80 and 1.60 g/cm³ on average. *Porites* colonies are typified by porosity of 40–54% and

355

density of 1.1–2.0 g/cm³. Coralline algae are 50 to 60% porous. However, coralgal buildups

356

no longer preserve their original internal porosity and density. Intense taphonomic processes,

357

i.e. internal bioerosion (Perry and Hepburn, 2008; Hutchings, 2011) and early marine

358

cementation (Macintyre, 2011; Perrin, 2011) may alter biogenic structures. Organic

359

encrustation, sediment infilling and early cementation of primary cavities and borings result

360

in massive, rigid, close-work frames (see Martin-Garin and Montaggioni, 2023 for review),

361

the residual porosity of which was estimated to average 50% (Smith, 1983). In high-energy

362

reef sittings, early marine cements can precipitate rapidly within coral skeletons post-mortem

363

(Rashid et al., 2020). In addition, primary coral skeletal aragonite fibers continue developing

364

syntaxially through secondary inorganic toward the centre of intraskeletal voids, resulting in

365

their partial infilling (Perrin, 2011). Marine cements were shown to occupy about 3-5% of the

366

total rock volume (Macintyre, 1977; Montaggioni, 1978) and 20 to up to 50% of cavity and

367

pore volumes (Coudray and Montaggioni, 1986).

368

Values of reef boulder bulk densities used as references for calculating boulder weight vary

369

significantly from author to author while there was frequently no discussion and clarity as to
how these values were obtained. Values of 1.1–1.4 g/cm³ were provided by Paris et al. (2009), 1.3 g/cm³ by Lau et al. (2018), 1.44 g/cm³ by Morgan and Kench (2012), 1.85 g/cm³ by Morgan and Kench (2014), 1.99 to 2.4 g/cm³ (Soria et al., 2018), 2.0–2.01 g/cm³ by Goto et al. (2009), Spiske et al., (2008), Frohlich et al. (2009) and Hongo et al. (2018). Higher values range from 2.3–2.5 g/cm³ (Etienne, 2012; Etienne et al., 2014) to 2.70 g/cm³ (Bourrouilh and Talandier, 1985). Most of the previously applied densities for modern reef limestones appear to have been overestimated when compared to the average bulk densities and porosities of the representative boulder components, thus resulting in overestimates of actual boulder weights.
Some calculations of boulder masses were performed considering boulders as perfectly shaped polygons as cubes or parallelepiped rectangulars.
From the nearby Makemo Atoll, Lau et al. (2016), using measurements of both mass/volume ratios and Archimedean buoyancy from individual boulder samples, considered that a bulk density of 1.7 g/cm³ was a reasonable value for boulder mass determination. This value refers to an average boulder porosity of about 40% assuming a bulk density of 2.93 g/cm³ for aragonite and of 2.70 g/cm³ for calcite components. Such an amount of porosity is in agreement with estimates for recent biogenic limestones by Smith (1983). Accordingly, in the present study, a density value of 1.7 g/cm³ for coral boulders was used for the calculation of boulder weight.

### 3.3. Hydrodynamic model of boulder displacement

The boulders deposited along the south-eastern reef flat zone at Fakarava have proven to come from the adjacent uppermost forereef slope. These clasts are assumed to be detached pieces of upper forereef spurs, prior to be emplaced onto the outermost reef flat zone and displaced landwards. This is supported by the presence, at some parts of the spur-and-groove system, at depths of less than 10 m, of submerged, already detached blocks of same shapes
and of sockets of same or greater sizes as blocks present on the adjacent reef flat zone (Fig. 6).

Over the last two decades, the conditions of coral reef mega-clast displacement by hydrodynamic run-up flows were actively discussed in different ways – from sliding to rolling or lifting-saltation – strictly connected to pre-setting situations (Nott, 1997; Goto et al., 2009; Nandasena et al., 2011; Etienne, 2012; Etienne et al., 2014; Lau et al., 2016; Watanabe et al., 2016; Hongo et al., 2018; Iwai and Goto, 2021; Nandasena, 2011, 2022; Weiss et al., 2022 and references herein). The boulder size and hydrodynamic parameters of high-energy marine hazard events, including the wave flow velocity, were taken into account. In 2022, Weiss and associates presented a new model in which were underlined the importance of mass, slope and roughness in front of coastal boulders. Also considered were the shape of boulders – rectangular prismatic rather than spheric –, their possible submersion and vertical sliding by the flow. Similarly, Nandasena et al. (2022) provided a more realistic, new deterministic and theoretical approach. This led to revisit motion equations, expressing minimum flow velocities values for coastal boulder motion.

The formulas developed by Nandasena et al. (2022) were used herein for establishing a representative hydrodynamic model of coral reef boulder transportation. Non-locked boulders originally at submerged or emerged settings can be displaced through 3 basic transport modes (Fig. 7): lifting-saltation, sliding and/or rolling (overturning). The modes originate from drag, inertia, restraining and uplift forces, in response of application of wave flow to boulder faces. In the specific case of sliding, friction forces must be added. In the case of non-isolated boulders, trapped into joint-bounded sockets, especially, with their top surface only exposed, lateral motion by sliding and rolling remains restricted; lifting is the only movement to be achievable. The water flow and the upper boulder face only interplay, because the latter is the only to be exposed to the flow of the carrying fluid. Detached pieces of spurs initially
419 submerged at shallow depths can be projected onto the reef flat zone by lifting. After lifting
420 and deposition on the reef flat surfaces, boulders could be moved inwards by lower flow
421 velocities through sliding or rolling according to their form and geometry.

422 In the model used, a given boulder is regarded as a polygone defined by three main axes and
423 three projected surfaces. The three main axes are defined as follows: a-axis, the longest in the
424 horizontal plane and originally perpendicular to the displacement axis of the boulder and to
425 the flow direction when the boulder was not yet moved; b-axis, the shorter in the horizontal
426 plane and in the same direction as the displacement axis of the water flow and boulder; c-axis
427 defined in the vertical plane. The projected surfaces are delineated by *ab* and *ac* axes
428 respectively. The areas on which forces are acting on a given boulder are defined as follows:
429 area of drag force (*A*<sub>d</sub> = *ac*), area of lift force (*A*<sub>l</sub> = *ab*), and volume of the boulder (*V* =
430 *abc*).

431 The new parameters defined by Nandasena et al. (2022) are:

432 Virtual drag coefficient: $C_{dv} = \frac{A_c}{ac} C_d$ where *a* and *c* is the product of the values of *a*- and *c*-axes;

434 Virtual lift coefficient: $C_{lv} = \frac{A_l}{ab} C_l$ where *a* and *b* are the product of the values of the values
435 of *a*- and *b*-axes;

436 Virtual volume coefficient: $C_{vv} = \frac{v}{abc}$ where *a*, *b* and *c* are the product of the values of *a*-, *b*- and *c*-axes.

438 Based on the flow transport equations by Nandasena et al. (2022), for non-locked blocks in
439 exposed or submerged settings, flow velocities *u* can be calculated according to the
440 considered three transport modes:

441 – sliding:
$$u^2 \geq \frac{2 C_v v (\rho_s / \rho_w - 1) g c (\mu_s \cos \theta + \sin \theta)}{C_d v (c/b) + \mu_s C_l v}$$

– rolling (overturning):

$$u^2 \geq \frac{2 C_v v (\rho_s / \rho_w - 1) g c (\mu_s \cos \theta + (c/b) \sin \theta)}{C_d v (c^2/b^2) + C_l v}$$

– lifting-saltation:

$$u^2 \geq \frac{2 C_v v (\rho_s / \rho_w - 1) g c (\cos \theta)}{C_l v}$$

A coral reef boulder in a joint-bounded setting will be moved by lifting only if stating:

$$u^2 \geq \frac{2 C_v v (\rho_s / \rho_w - 1) g c (\mu_s \cos \theta + \sin \theta)}{C_l v}$$

In the equations, u is flow velocity in m/s; b and c are the b-and c-axes of boulders in metres;

$\rho_s$ is boulder density as defined above (1700 kg/m³); $\rho_w$ is seawater density (1020 kg/m³);

$C_v v$, virtual volume coefficient; $C_d v$, virtual drag coefficient (1.95); $C_l v$, virtual lift

coefficient (0.178); $\theta$, angle of the reef flat zone and conglomerate platforms on which

boulders were deposited originally (this angle in the studied areas is assumed to be equal to

zero; note that this value was also attributed to emerged spur top surfaces, as an

approximation based on drone-acquired imagery); $\mu_s$, coefficient of static friction (0.5); $g$,

acceleration of gravity (9.81 m/s²).

### 3.4. Dating procedures

Canavesio et al. (2018) claimed that radiometric dating of coral colonies collected in growth

position from the top side of reef boulders has limited application for defining the precise

time of boulder detachment, because some can be dead since an indefinite time. In order to

limit dating biases from sampling, two in-situ coral colonies per selected boulder, regarded as

well preserved – with visually identifiable corallites; see Fig. 13, photograph at top, were
usually sampled. The collected coral specimens were taken from fast-growing, branching –
*Pocillopora, Acropora* and massive forms (merulinids). A total of 33 coral specimens,
labelled FAK and F and respectively collected in 2019 and 2022, were analyzed using the
uranium-series method. From each selected boulder, the younger coral age was retained since
considered to be the closest to the time of boulder detachment.

Prior to dating, preparation required cutting of millimetric subsamples from original coral
material. The most clean and pristine subsamples were selected. Then, these pieces were
rinsed with mQ water and ultra-sonicated several times. After adding a triple ²²⁹Th ²³³U ²³⁶U
spike in a Teflon beaker, samples (from 150 to 350 mg) were dissolved with diluted HCl. The
U–Th separation and purification were performed after coprecipitation with Fe(OH)₃, on 0.6
ml columns filled with U-TEVA and prefilter resins, in nitric media – see Pons-Branchu et al.,
2022 for details. The U and Th isotopic compositions were analyzed at the Laboratoire des
sciences du climat et de l'environnement (LSCE, France), on a Multi-Collector inductively
coupled plasma source mass spectrometer (MC-ICPMS) Thermo Scientific™ Neptune Plus
fitted with a desolvating introduction system (aridus II), and a jet interface. For mass
fractionation correction, we used an exponential mass fractionation law – normalized to
natural ²³⁸U/²³⁵U isotopic ratio – and standard – sample bracketing. More details on the
analytical procedure – chemistry and MCEICPMS analysis – can be found in Pons-Branchu et
al. (2022). After corrections for peak tailing, hydrate interference and chemical blanks,
²³⁰Th/²³⁴U ages were calculated from measured atomic ratios through iterative age estimation
using the ²³⁰Th, ²³⁴U, and ²³⁸U decay constants of Cheng et al. (2013) and Jaffey et al. (1971).
In the text, ages are expressed in calendar years (cal. yr) before present (BP) corrected for
detrital thorium, all by reference to 2024 or in calendar years BC (before Jesus Christ) or AD
*(Anno Domini)* (Table 1). The error ranges associated with age measurements vary between ±
6 and ± 45 years. In order to make age results more easily comparable to previous research,
particularly that based on radiocarbon measurements or expressed relative to the Common Era, in Table S1 (Supplementary material), ages are also given in conventional years BP by reference to 1950.

Dating Boulder D from Zone A (west) provided the opportunity to define the age of building of forereef coral spurs from which boulders were detached. Seven coral samples were collected from three faces of the boulder at three different levels.

## 4. Results

### 4.1. Morphometric attributes and location of boulders

Azimuts of the longer (major) axis from 44 boulders were measured (Fig. 8). About 75% are oriented sub-parallel to the reef front line. Of the remaining 25%, half is aligned parallel, half is in a sub-orthogonal direction.

The subset of 19 U/Th dated boulders was investigated in terms of physical attributes (Table 2). According to the classification of Zingg (1935) and Blott and Pye (2008) in Dewey et al, (2021), most are rather platy, with sub-angular faces and sharp broken edges. Some are ellipsoidal and rounded with relatively smoothed faces. The investigated ones are 3.1 to 8.5 m in length, 1 to 6.1 m in width and 1.2 to 3.9 m in height. Their volumes range from 4 to 82 m³.

There is a weak linear correlation between the long (a) and short (b) axes, and between the long (a) and the height (c) axes (Fig. 9). Assuming a bulk carbonate rock density of 1.7 g/cm3, the corresponding masses were estimated to vary from 7 to 139.4 tons.

After detachment from the adjacent forereef spurs, the boulders were projected onto different reef-rim zones (Table 2). The majority of the boulders, including Boulders B, C, D, E in Zone A and Boulders Q, R, S, T, U, V, X in Zone B, are closer to the reef front line, within the 60 metres, setting on the inner reef-flat zone (Fig. 2A, B). These belong to the reef-top boulders category, according to the classification by Lau et al. (2016). More than 30 boulders,
including Boulders L, 14 to 18 in Zone A and 9 to 12, P1 to P7 in Zone B are located farther inwards, lying on conglomerate pavements in front of the vegetalized islets. These can be defined as *conglomerate-top* features. Others, including Boulders 1, 2 and 3, were encountered within the *hoa* zone. Most of these boulders have deposited as isolated elements, except for Blocks B, C and D, which crashed into one another as shown by the sharp break notching Block D (Fig. 10).

Among the investigated boulders, the farthest boulder (Boulder 1, Zone west, *hoa*) at a distance of 235 m is one of the smallest (9 m³) and of the lightest (15.3 tons). Surprisingly, the second farthest landward one (Boulder 2, Zone A), located within the south-south-west *hoa* zone at a distance of 215 m from the reef front, is also one of the heaviest (40 m³, 68 tons). Its present-day location may result from later removing further inland. The lighter ones – 14, 16, Zone A; 4 m³; 7 tons – are 120 and 110 m behind the reef edge. The heaviest clast – X, Zone B; 82 m³; 139.4 tons (Fig. 11A, B) – is at 75 m from the algal crest line. Boulder K (Zone A) does not exceed 28 m³ and 47.6 tons (Fig. 11C, D) is located at a distance close to that of Boulder X. One of the heaviest boulders – A, Zone A; 69 m³, 117 tons; (Fig. 11E, F) – is one of the closest to the reef edge (82 m). Such apparent inconsistencies may result from deposition controlled by marine hazard events of varying intensity or occasional removing during subsequent events. As shown in figure 12 A, there is a weak relationship between boulder weight and distance to the reef front line. This could be due to the small dataset investigated (n = 19) and local topographic conditions, such as boulder clustering near the edge of the conglomerate platform (as seen in Fig. 2) and/or by the occurrence of the heaviest blocks (A, X) located quite close to the reef edge. However, the distribution of the other lower heavy boulders seems to be controlled by their weight, the distance between boulder locations and the reef front line tends to increase with decreasing boulder weight (Fig. 12). The landward-fining boulder size distributional pattern is apparently controlled by a linear
decreasing trend, but mathematically barely significant, as expressed by a very low value of
the coefficient of determination (R² = 0,02487). Similar linear trends were observed from
boulder distributions in various reef sites (Goto et al., 2009; Hongo et al., 2018).

## 4.2. Estimation of flow velocities

At Fakarava, four boulders of varying volumes and weights (A, D, K and X) were selected as representatives of the studied boulder set for estimation of flow velocities (Table 3). The minimum flow velocity required to trigger the motion of the relevant boulders ranges between 3.05 and 13.66 m/s. Velocities of 12.11 to 13.66 m/s may have allowed detached and isolated blocks to be uplifted and deposited onto the outer reef flat zone. Their landward transport by sliding may have operated at velocities not exceeding 5.10 m/s, while transport by rolling required velocities ranging between 5.05 and 8.12 m/s. In the case of initially joint-bounded blocks, the velocity threshold values of 10.74 to 13.66 m/s were sufficient to initiate boulder movement. Calculating velocities related to sliding and rolling motions indicates that the c/b ratios play a significant role because contributing to the denominator of the Nandasena's equations. While consistent for boulder A, D and K, the respective velocities (3.35 and 5.05 m/s) experienced by the largest boulder X appear to be underestimated when compared with those applied to boulder A. Indeed, given sizes and masses of boulders A and X are very close, sliding and rolling velocities should be also very similar. The marked discrepancy in velocities between these blocks is probably due to the irregular, non-geometric shape of boulder X. Considering the estimated volumes (69 to 82 m³) and weights (117 to 139.4 tons) of the boulders, the minimum velocities obtained for uplifting are consistent with each other. The validity of the calculation method used to estimate flow velocities is altered by the reduction of blocks to elementary, virtual geometric forms. Although the boulder shape factor was not taken into account, analysis of the photograms from boulders A, K and X suggested a possible motion by rolling, contrary to boulder D, probably moved by sliding.
562 4.3. *Estimated emplacement ages of boulders*

563 The time at which a given segment of foreereef spurs was extracted and transported across
564 reef-rims can be estimated based on the U/Th ages of coral specimens found at boulder top
565 surfaces and regarded as ones among the lastest colonizers. Most of the investigated boulders
566 were subjected to dating of two coral samples (Tables 1, S1, Supplementary material). From
567 each investigated block, differences in age of the dated coral samples range between 4 and
568 751 years. It is noteworthy that there is a higher consistency – 4 to 82 years in age
569 differences – for the boulders detached most recently – less than 2,000 years old. By contrast,
570 age inconsistency tends to increase with the age of coral samples. From coral samples older
571 than about 2,000 cal. yr, age differences range between 226 and 765 years. This agrees with
572 the results by Canavesio et al. (2018) from dating of corals collected at the surface of coral
573 reef boulders in some Tuamotu atolls: ages of superficial coral colonies have to be regarded
574 as providing approximate, not absolute dates for boulder detachment, due to possible coral
575 death at an unknown date prior to detachment.

576 The returned coral ages range from 3,684 cal. yr BP – Boulder 18, Coral FAK 25 – to 282 cal.
577 yr BP – Boulder C, FAK 7. 42% of the dated corals, noted BC, have grown before the
578 beginning of the Common Era. The remaining 58% of the samples, noted AD, were living
579 over the last two millennia. The time interval of reef block emplacement is considered to
580 range from the year 1406 BC – Boulder G, FAK 10 – to the year 1742 AD – Boulder C, FAK
581 7 –, i.e. a period of about 3,200 years.

582 In Zone A (west), detailed examination of the error ranges for the U/Th coral ages indicated
583 that these vary from ± 8 yr (Boulder B), ± 20 yr (Boulder D) and ± 32 (Boulder C) (Table S1,
584 Supplementary material). The differences in age are 33 yr between Boulders B and D, 52 yr
585 between Boulders B and C, 79 yr between Boulders C and D. Boulders B, C and D would
586 have been transported at the end of the 17th century, if considering the dates obtained from
these blocks, respectively 1690 AD (FAK 9), 1742 AD (FAK 6) and 1663 AD (FAK 4).

While their error ranges of age measurements vary from ± 9 to ± 26 yr, the 8 blocks dated
between 21 and 350 AD in Zones A and B have differences in age which vary significantly:
for instance, 32 yr (between Boulders 5 and T), 34 yr (between Boulders A and K), 43 yr
(between Boulders 14 and X), 50 yr (between Boulders 2 and A). The minimum age
difference is 30 yr (between Boulders 16 and T).

The record of extreme marine hydrodynamic events during the period preceding the Common Era is more difficult to be established, due to larger age biases in coral dated samples from some investigated blocks (Fig. 13, Table 1). The depositional history of reef erratic blocks in the considered area would have started from the 15th BC with deposition of some clasts in Zone A: Boulder G deposited by 1406 BC (FAK 10), Boulder J by 1311 BC (FAK 15), Boulder E by 1269 BC (FAK 11). Boulder L (F 14), Boulder 18 (FAK 24) and Boulder 3 (FAK 29) have been deposited respectively at 1099 BC, 909 BC and 851 BC. The last blocks deposited prior to the beginning of the Common Era were Boulder 1 (FAK 62) at 586 BC and Boulder 14 (F 18) at 307 BC.

## 5. Discussion

### 5.1 Age of forereef spur building

From U/Th dating of the wall faces of Boulder D, growth initiation of forereef spurs in the investigated sector appears to have started approximately 5,000 years ago (Fig. 14). This is in agreement with the fact that, in the Pacific reef province, upper forereef spur accretion is known to have been active during the Holocene rise in sea level (Duce, 2016; Montaggioni et al., 2019a; Duce, 2016). The ages range from 4,603 to 2,616 cal. yr BP (Fig. 14; Table S1, Supplementary material). This reveals that spur growth was mainly active during the mid
Holocene, at the time when sea level was about 0.80 m above present position (Hallmann et al., 2020).

## 5.2. Boulder displacement modes

The distributional pattern of the investigated mega-clasts appears to be partly controlled by size. Their size distribution shows a linear landward decreasing trend. According to Goto et al. (2009) who observed a similar trend at Kudaka Island, Okinawa, Japan, such a trend could be explained by the fact that, after breaking on the reef edge, wave transport capacity of boulders across the reef flat zone tends to decrease inland in response to the decreasing wave height. However, the boulders may have been displaced not by a single wave, but by a series of successive waves during a given storm event or during several storm events. The wave force repeatedly applied to boulders can result in their occasional motion and repositioning inwards over greater distances. This is exemplified by the boulders, that are lying in the *hoa* at Zone A. The *hoa* area has very flat topography, is arased and pratically devoid of sediments in its inner parts. This strongly suggests that the channel is functioning only during extreme marine hazard events, allowing strong wave surges and transported sediments to extend towards the lagoon.

The best case scenario to date for estimating the cyclone-generated minimum flow velocities required for carbonate boulder dislocation and transport was established in eastern Samar, Philippines, during super-typhoon Haiyan in November 2023 (May et al., 2015; Kennedy et al., 2017; Soria et al., 2018). Wave runup exceeded 7 to 15 m high above mean sea level and moved large, emerged reef-top carbonate blocks of 83 m³ and 180 tons, at distances longer than 280 m inland. Boulder dislocation was interpreted as having operated with minimum flow velocities greater than 5 m/s and reaching values of 8.9–9.6 m/s at Fakarava, estimates of transport flow forces indicate that extracting and projecting isolated, submerged pieces of forereef spurs onto the reef flat surfaces has required minimum run-up flow velocities of
about 11 to higher than 13 m/s. Movement of boulders further landwards by sliding and rolling could have occurred at minimum velocities of about 3–5 m/s and 5–8 m/s respectively.

All these values are close, not only to velocities generated by supertyphoon Haiyan, but also to minimum flow velocities (Etienne et al., 2014) generated by the 1982–1983 ENSO-related cyclonic events at Anaa (4.2–14.1 m/s), Rangiroa (2–11.8 m/s) and Makemo (2.8–10.7 m/s) to move boulders similar in sizes to those described at Fakarava. During these hazard events, maximum swell height was 10–15 m (Laboute, 1985), as reported from eastern Samar, thus giving evidence of the run-up flow amplitude required to remove forereef spur fragments.

Boulders have been separated from spurs and projected onto reef rim surfaces throughout the last 3,500 years, from a time when sea level started to drop down to its present position.

Although the limited number of investigated boulders may bias the results and interpretation, boulder displacement is strongly suggested to have preferentially operated between 300 years before and 300 years after the beginning of the Common Era, at a time when sea level was close to its present position. This means that a seastand higher than present could not be a promoting factor of boulder motion.

## 5.3. Agents of block displacement: cyclones versus tsunamis

Nott (1997, 2003a), Spiske et al. (2008), May et al. (2015), Weiss (2012, 2022), Kennedy et al. (2016) and Wei et al., (2024) discussed the criteria to be used for distinguishing between tsunami- and cyclone-generated wave impacts in deposition of reef-derived boulders. One of the common useful criterium refers to the distance of boulder location to the expected source with respect to boulder weight. Boulders moved by tsunamis could be discriminated from those dislodged by storm waves, because the latter have shorter periods than tsunami-generated waves. Consequently, storm waves break further away from the shoreline (Goto et al., 2009, 2010a) and their transport capacity is significantly lower than that of tsunamis. In
the Ryukyu Islands, tsunami boulders of > 120 tons were displaced at distances of 1400–1500 m from their sources (Goto et al., 2010a), while storm boulders of similar weights were located at 40–50 m only inward of the reef edge. Similarly, at Fakarava, boulders of about 90 to 140 tons are located within a distance of 40 to 75 m away from the reef edge. However, the pattern imposed by the boulder depositional process may be locally obliterated since deposited clasts can be removed repeatedly by successive storms and reworked seawards by return flows. The available data about tsunamigenic wave events in the Tuamotu lead to think that the emplacement of reef mega-clasts along the south-eastern rim margins of Fakarava resulted from super-cyclones, not from earthquake-related tsunamis, for the following reasons. First, all the published data about the passage of tsunamigenic wave surges in the Tuamotu indicate that these waves do not exceed 2 m high (Shindelé et al., 2006). Even in the Marquesas Islands which are regarded as being hit hardest by earthquake-related tsunamis, wave runups rarely exceed 4 m high (Paris et al., 2020). The described dominant effects along the Marquesan coasts are deposition of sand washover fans and erosional features in coastal sand ridges, displacement of small sized boulders, ranging between 2 and 4 m3 in maximum volume, desintegration and reworking of conglomeratic beachrock flagstones (Paris et al., 2020). Second, contrary to the Marquesas Archipelago, the north-west Tuamotu atolls and, particularly the eastern reef-rim at Fakarava, are protected from tsunamigenic waves generated by motions of the plate boundaries and coming from the western american coasts, the north, north-west and west Pacific areas. This protection results from the *island shadow effect* (Andréfouët et al., 2012), meaning that most of the atolls located around Fakarava can serve as protective barriers against the impact of possible tsunamigenic waves. Third, the origin of mega-boulders deposited on nearby atolls was previously debated and attributed to extreme cyclonic events. At Rangiroa Atoll, projection of the largest mega-clasts onto the north-north-west reef-flat surface was previously postulated to have resulted from a
tsunami, although a cyclonic origin was not precluded (Bourrouilh-Le Jan and Talandier, 1985). Since in French Polynesia, waves from earthquake-related, tsunami events, at least from those historically recorded, can reach 2 m in maximum height (Schindelé et al., 2006; Sladen et al., 2007) and distant-source storm waves can occasionally exceed 4–5 m (Canavesio, 2019), it has been concluded that the investigated mega-blocks have been unlikely displaced by such low waves. Indeed, Hyvernaud (2009) claimed that these *cyclopean* boulders were emplaced by a cyclonic event dated at about 1715 ± 15 AD. Similarly, Etienne et al. (2014), Lau et al. (2016), Canavesio et al. (2019, 2023) did not retain the role of tsunamis in reef boulder extraction and displacement in French Polynesia. Using U/Th dating, the largest boulder (369 m³) found on the north-north-western rim zone at Rangiroa was demonstrated to have been moved during the AD 1906 cyclonic event (Canavesio et al., 2023). Similarly, the emplacement of the boulder field in north-north-western Anaa Atoll resulted from the same event. Wave heights of up to 15 m were expected to have been generated at that time (Canavesio, 2014). Although it is not possible to totally preclude the action of remote or local tsunamis in boulder transport in French Polynesia, the fact that largest known mega-clasts in the Tuamotu were displaced over distances of several tens of metres by cyclone storms, advocates for the dominating role of cyclogenesis in boulder detachment and motion. Finally, under the assumption that the detachment and transport of mega-boulders along the south-eastern coast of Fakarava may have been promoted by tsunamigenic waves, only failures of nearby atoll flanks (Goff and Perry, 2016; Montaggioni et al., 2019b) could have been responsible for the mega-boulder emplacement, not earthquakes generated from the Pacific plate boundaries.

## 5.4. Expected number of recorded extreme marine hazard events

Taking into account possible sampling biases (Canavesio et al., 2018), it was likely that all the blocks dated from the Common Era (AD) may have been emplaced during three distinct
extreme marine hazard events (Fig. 13). The early event could have occurred at the middle of the first century AD, resulting in the dislocation of Boulders 2 (FAK 63: 21 AD), Boulder A (FAK1: 71 AD) and Boulder K (FAK 16: 105 AD). The second event could have occurred at the middle of the third century AD; this is supported by the dislodgement of Boulder P3 (FAK 21: 177 AD), Boulder P5 (FAK 22: 212 AD) Boulder T (FAK 19: 244 AD), Boulder 16 (F 20: 274 AD) and Boulder X (F 30: 350 AD). As to the late event, it was regarded as responsible for the detachment of Boulder D (FAK 4: 1663 AD), Boulder B (FAK 8: 1690 AD) and Boulder C (FAK 6:1742 AD) at the limit between the 17 and 18th centuries AD. Boulder L (F 14; 1099 BC), Boulder 18 (FAK 24: 909 BC) and Boulder 3 (FAK 29: 851 BC) could have been dislocated during a same single event or several distinct events, adjusted around the 10th century BC. Deposition of Boulders G (FAK 10), J (FAK 15) and E (FAK 11), between 1406 BC and 1269 BC is expected to reflect one single or different events, that occurred around the 15th–13th centuries BC. Later, one or several additional extreme marine events are postulated to have occurred during the 6th–4th centuries BC as recorded by Boulders 1 (FAK 62: 586 BC) and 14 (F 18: 307 BC). As a summary, a minimum of six extreme marine hazard events (super-cyclones) would have hit the south-eastern coast of Fakarava over the last 3,500 years.

## 5.5. Comparisons with nearby atolls

On nearby atolls (e.g. Tikehau, Rangiroa and Anaa), the majority of the 28 U/Th dated boulders appear to have deposited between the 18th and 19th centuries, some during few centuries before the Common Era (Canavesio et al., 2023). On Makemo Atoll, U/Th dating of 20 reef boulders revealed that most have deposited during the 14th, 18th, and 19th centuries AD (Lau et al., 2016). By contrast, in the investigated areas at Fakarava, the youngest mega-block age is $1742 \pm 35$ AD (FAK 7, Boulder C). In French Polynesia, as the earliest records of cyclone date back only to 1822 and 1831 (Laurent and Varney, 2014), deposition of none
735 of the dated boulders at Fakarava in the investigated areas, can be linked to a recorded

736 historical cyclone event. However, some of them could have emplaced at the same time as

737 some blocks dated from Rangiroa and Anaa by Canavesio et al. (2023), from Marokau and

738 Kauehi by Canavesio et al. (2018) and from Makemo by Lau et al. (2016) respectively.

739 In Zone A, Boulder 14 (Sample F18: 307 ± 26 BC) would have been extracted during the

740 same hazard event as that has detached block n° 26 (281 ± 11 BC) studied at Rangiroa by

741 Canavesio et al. (2023). Boulder 2 (FAK 64: 21 ± 10 AD), Boulder A (FAK 2: 71 ± 6 AD)

742 and Boulder K (FAK 17: 105 ± 10 AD AD) may have emplaced coevally with several

743 boulders investigated from Kauehi Atoll (RCPB 160: 75 ± 9 BC, 7 ± 8 AD) and interpreted as

744 resulting from a marine event at 54 AD by Canavesio et al. (2018).

745 Similarly, Boulder C (FAK 6: 1742 ± 35 AD), Boulder D (FAK 4: 1663 ± 20 AD) and

746 Boulder B (FAK 8: 1690 ± 9) may have been deposited during the same storm events as block

747 E130 (1739 ± 6 AD), block C051 (1717 ± 10 AD) or block E129 (1722 ± 6 AD) at Makemo

748 and as block n°14 (1742 ± 9AD) at Anaa (Canavesio et al., 2023) as well. In both Zones,

749 Boulder P3 (FAK 21: 177 ± 14 AD), Boulder P5 (FAK 23: 212 ± 13 AD), Boulder T (FAK

750 19: 244 ± 9 AD), Boulder 16 (FAK 20: 274 ± 11 AD) and Boulder X (F 30: 350 ± 11 AD),

751 would have been detached coevally with block n° 24 (210 ± 12 AD) from Rangiroa

752 (Canavesio et al., 2023).

753 As a summary, our findings indicate that most of the studied megaclasts from south-east

754 Fakarava have been detached between about 3,500 and about 1,500 cal. yr BP. By contrast, in

755 some nearby atolls (Tikehau, Rangiroa, Makemo, Marokau, Kauehi, Anaa), the available data

756 revealed that deposition of boulders would have operated more recently, during the last

757 centuries. These conflicting results could be explained by differing trajectories of cyclones

758 over time with the occurrence of a single cyclonic event affecting a limited number of atolls,
759

possibly sparing the closest islands and by the limited number of investigated atolls and atolls

760

areas in the Tuamotu.

761

## 5.6. *Paleo-records of cyclone events connected with ENSO activity*

762

Age analysis of the megaclasts deposited along the south-eastern reef-rim at Fakarava offers

763

the opportunity to try extending the record of past cyclone events over the last 3,500 years in

764

the Tuamotu (Figs. 13, 15). At Fakarava, between the beginning of the 15th (1406 ± 12 BC)

765

and the 4th century BC (307 ± 45 BC) i.e. between 3,430 ± 12 (FAK 10) and 2,331 ± 26

766

cal. yr BP (F 18), over an interval of about 1,100 years, an earlier series of three huge cyclonic

767

events resulted in deposition of Boulders 1, 3, 14, 18, E, G, J and L. This may suggest that an

768

average of about 1 extreme marine hazard event (*supercyclone*) about every 250-yr long

769

interval in the considered area. Similarly, Boulders 2, 16, A, K, T, P3, P5 and X may have

770

been detached during the period of enhanced ENSO amplitude, during an about 350 yr-long

771

interval, between the beginning of the Common Era (21 ± 9 AD) and the 4th century AD

772

(350 ± 11 AD), i.e. between 2,003 ± 9 (FAK 64) and 1,674 ± 11 cal. yr BP (F 30). The ages of

773

boulder emplacement within this time span may suggest an average of about 1 extreme

774

marine hazard per every 150 years.

775

As a summary, in the considered area, cyclogenesis appeared to have markedly increased

776

between about 3,400 and 2,300 cal. yr BP and between about 2,000 and 1,700 cal. yr BP (Fig. 15A), interpreted as related to increasing ENSO activity. This scenario was confirmed by a

777

number of previous studies in the tropical south Pacific. During these spans, according to

778

Thompson et al. (2017), the frequency of ENSO events was amplified dramatically. In the

779

region, Chand et al. (2017) pointed out that cyclone activity increases by 20–40 % through

780

enhanced ENSO amplitude, expressing, according to Sadekov et al. (2012), a strong

781

mutidecadal to centennial scales variability. Similarly, in the Society Island group, grain size

782
analysis of sediment fractions > 250 mm in a vibrocore extracted from the lagoon of Taha'a Island suggested that cyclone activity has significantly increased at around 2,600–2,500 and 2,000–1,800 yr BP (Toomey et al., 2013; Bramante et al., 2020; Fig. 15B). These may provide an insight into short-term, individual ENSO-like episodes throughout longer-term, highly variable ENSO-like climate states and average ENSO behaviour. Accordingly, the boulder distributional pattern reported from the investigated sites is qualitatively consistent with several key ENSO-sensitive proxy records. The times of block detachment events could be linked to the identified active ENSO phases. In the literature, reconstructions of paleo-ENSO over the late Holocene are based on different proxy sources – corals, bivalves, planktonic foraminifera, lacustrine varved deposits, pollens, leaves, organic biomarkers – and global circulation models. From measurements of sediment laminae light-intensity in a core from Laguna Pallcacocha in Ecuador, Moy et al. (2002) claimed that several periods of intensification of ENSO activity occurred during the late Holocene, especially, peaking at around 2,500 cal. yr BP, between 2,000 and 1,500 cal. yr BP and at around 900–800 cal. yr BP (Fig. 15C). Using textural analysis of a core extracted from Junco Crater Lake in the Galapagos, Conroy et al. (2008) confirmed that there was an increased ENSO frequency and stronger El Niño events from about 3,200 cal. yr BP, with 2 two-step increases in intensity frequency between 2,000 and 1,500 cal. yr BP and at around 800–700 cal. yr BP (Fig. 15C). Similar findings were derived from coral-based reconstructions of ENSO variance in the central Pacific (Northern Line islands) over the Holocene by Cobb et al. (2013), underscoring an intensification of ENSO activity between around 3,000-2,500 cal. yr BP (Fig. 15D).

A close examination of additional ENSO paleo-records underpins this scenario or, at least, some of the identified events. In the western equatorial Pacific (Eastern Australia), analysing variations in carbone isotope composition of leaves deposited in lacustrine sediments (Swallow Lagoon), Barr et al. (2019) showed that, from about 3,200 cal. yr BP, there was an
amplification of the frequency and strenght of El-Niño phases, interpreted as triggered by
progressive orbitally-driven weakening of the Pacific Walker Circulation. In the central
tropical Pacific, coral-based reconstructions and climate simulations revealed that a
significant increase in ENSO amplitude and frequency from the late Holocene (Grothe et al.,
2019) with occasional extreme El-Niño events throughout a wide-ranging natural variance
(Cobb et al., 2013; Karamperidou and DiNezio, 2022; Lawman et al., 2022).

The extreme strom events resulting in deposition of the three boulders (B, C, D) during the
1663 ± 20 AD – 1742 ± 35 AD interval i.e. between 361 ± 20 (FAK 4) and 282 ± 35 cal. yr
BP (FAK 7), may be connected to highly active El Niño events reported to the second half of
the Little Ice Age (Henke et al., 2017; Prohaska et al., 2023).

Along the south-eastern reef-rim at Fakarava, several episodes of boulder non-deposition
appear to have occurred during the mid Holocene, before about 3,500 cal. yr BP – before the
15th century BC –, between about 2,500 and 2,000 cal. yr BP – between about the 6th and the
3rd centuries BC. These events could coincide with intervals of a relatively lower-baseline
cyclogenesis activity, in relation with an attenuation of ENSO amplitude as documented by
Moy et al. (2002), Koutavas et al. (2006), Conroy et al. (2008) and Grothe et al., (2019); (Fig.
15). Similarly, Emile-Geay et al. (2016) identified several intervals of reduced ENSO variance
from about 7, 000 yr BP, related to millennial-scale quiescent times in the tropical Pacific
climate. This idea was confirmed by Riedinger et al. (2002), Thompson et al. (2017) and
Carré et al. (2021) who stated a mid-Holocene ENSO minimum (HEM) from 7,000–5,000 to
about 4,600–3,000 cal. yr BP, with several intervals of reduced ENSO variabilty of about 80%
in the eastern to central Pacific. The non-deposition interval from about 1,500 yr BP (from the
4th–5th centuries AD) could result from a decrease in ENSO activity in the South Pacific, as
revealed by a marked weakening of regional cyclogenesis activity during the 1,500–1,000 yr-
832 BP interval (Moy et al., 2002; Conroy et al., 2008; Cobb et al., 2013; Toomey et al., 2013; Bramante et al., 2020).

834 However, it is obvious that a database including a relatively limited number of dated boulder specimens cannot provide a comprehensive record of past hazard events. Defining the number of catastrophic cyclonic events based on a limited number of collected boulders may be biased by sampling procedures including incomplete age determination, but expresses significant trends in ENSO variability.

839 **5.7. Frequency of cyclone impacts in the north Tuamotu over the late Holocene**

840 At Fakarava, from the investigated areas, the boulder distributional pattern reveals that there was a minimum of one catastrophic storm events per every three centuries on average 841 between the 15th and 4th centuries BC, and three events between the 1st and 18th centuries 842 AD, able to detach and uplift mega-boulders onto reef top surfaces (Fig. 13). These findings 843 do not seem to be in agreement with previous works in the same region. Based on dating of 844 coral reef mega-boulders, Lau et al. (2016) revealed that five to nine cyclones have impacted 845 Makemo Atoll between 1200 and 1900 AD, with three to five events between the 16th and the 846 19th centuries, thus suggesting that an average of one to five cyclonic events per century 847 occurred with a mean recurrence time of 50 years. The discrepancies between Lau's work and 848 the present study may be due to unidendified overlapping between some coral ages from 849 Makemo. The studies by Dupon (1987), Laurent and Varney (2014) and those based on 850 chronostratigraphy of coral clasts deposited within some atoll reef-rim islet sequences in the 851 Tuamotu (Montaggioni et al. 2018, 2019, 2023a, b) have also identified the occurrence of at 852 least three storms per century. For instance, at Fakarava, from U/Th dating of 33 within-islet 853 coral clasts spanning the Common Era, a maximum number of three to five marine hazard 854 events per century was recorded (Montaggioni et al., 2023b), mainly tooking place in the last 855 millennium, maybe promoted by intensification of cyclone activity during the Little Ice Age 856
857 (Bramante et al., 2020). Similarly, at Takapoto Atoll, during the first millenium BC, between
858 the 10th and 4th centuries, nine distinct storm events were registered, with a maximum
859 frequency of three events per century (Montaggioni et al., 2019a). Apparent discrepancies
860 between the present findings and previous historical records could be due to the fact that huge
861 cyclone hazards – so-called *supercyclones*, categories 4–5 in the Saffir-Simpson scale – able
862 to detach reef mega-clasts are likely to have been unfrequent when compared to most of less
863 powerful storms that have hit the Tuamotu.

864 6. Conclusions

865 Our findings strongly suggest that tsunamis have to be probably dismissed as the major
866 causative mega-boulder displacement mechanism in the Tuamotu and that boulder
867 detachment and transport are linked dominantly to activity of super-cyclones rather than to
868 tsunamigenic waves, even the occasional role of regional tsunamis generated by atoll flank
869 failures cannot be left out. At Fakarava, along the south-west reef-rim side, extraction and
870 deposition of coral reef boulders could be ascribed primarily if not exclusively to cyclone
871 impacts because (1) heights of historically recorded tsunami waves were less than 2 m, thus
872 precluding any detachment and transport of larger mega-clasts; (2) cyclones, especially those
873 of the last 150 years, are known to have generated waves up to 15 m high, minimun run-up
874 flow velocities of about 3–14 m/s and to have transported boulders of the sizes similar to or
875 larger than those in the studied zones. The convex shape of the south-east reef-rim front line
876 by driving swell energy to concentrate during extreme marine events is likely to have
877 promoted deposition of boulders along this specific rim band. By contrast, the fact that sea
878 level was about 1 m higher than today about 3,500 years ago appears clearly to have not been
879 a determining factor of boulder dislogdement and transport.

880 Ages of some investigated boulders at Fakarava revealed to have been emplaced before the
881 beginning of the Common Era. The regional marine hazard pattern seems successively to
have been typified by an higher activity from about the 14th to the 4th centuries BC (3,400 to 2,300 cal. yr BP), followed by a decrease between the 3th and the 1st century BC (2,300 to 2,000 cal. yr BP), by a recovery of activity over a period of 4 centuries from the 1st to 4th century AD (2,000 to 1,700 cal. yr BP). Accordingly, the present study gives evidence of times of increasing cyclogenesis in the late Holocene in the northern Tuamotu, typified by markedly infrequent extreme high-energy marine events (*super-cyclones*), apparently concentrated over a minimum of six periods. This activity is regarded as consistent with reconstructed variability of past ENSO dynamics across the tropical Pacific, probably in relation with a large shift of the Pacific Walker circulation (Bramante et al., 2020) and/or a northward migration of the Intertropical Convergence Zone – ITCZ (Koutavas et al., 2006). Similarly, weakening of storminess over several intervals within the Common Era could be caused by repeated migrations of the Walker circulation and/or the southward displacement of the ITCZ.

There was apparently any boulder deposition record of the last two centuries and, especially, of the major 1903 and 1906 cyclone events in the studied areas at Fakarava, although the Rotoava village in the northern rim zone was devastated and flooded at that time. This non-deposition was probably driven by the trajectories of both cyclones that have preferentially impacted the northern atoll rim sides. Reversely, the cyclone events that were causing deposition of boulders older than 1,500 and 2,500 cal. yr BP at Fakarava were apparently not recorded in the closest islands as revealed by the absence of contemporaneous mega-blocks. This is expected to be due to the preferential direction of cyclones paths passing over Fakarava or lack of field investigations in the nearby atolls.

Finally, this study confirms the value of coral reef boulders as one of the more efficient proxies to reconstruct paleo-cyclone activity in the central South Pacific.
906 Authorship contribution

907 Lucien Montaggioni: conceptualization, writing and reviewing of the original draft. Bernard

908 Salvat: conceptualization, field methodology, resources, contributing to write the original

909 draft. Eric Bruneaud: field methodology and mathematical approach. Edwige Pons-Branchu:

910 U/Th dating. Bertrand Martin-Garin: figure conceptualization and drafting, draft reviewing.

911 Gilbert Poli: field assistance.

912 Declaration of competing interest

913 The authors declare that there are no competing interests which may have influenced the work

914 presented in the present contribution.

915 Acknowledgements

916 This work was partly supported by the research project n°6538/MED/REC of 09/16/2019 –

917 Délégation à la recherche, Tahiti, leaders: Jean Yves Meyer and Serge Planes – and

918 n°9553/MCE du 08.12.2021 – Direction de l'Environnement, Tahiti leaders: Bernard Salvat

919 and Miri Tatarata. Warm thanks are given to Jean Yves Meyer and Miri Tatarata for financial

920 support. Major field research (BS) occurred in March 2019 – with help of Annie Aubanel and

921 Hugo Dayan – and August-September 2022 (BS, EB, GP). U/Th dating was performed by

922 EPB. LFM, the corresponding author, analyzed and interpreted all the collected data and

923 wrote the manuscript in collaboration with the co-authors. BS, EB and GP provided field

924 pictures. Figures were conceived by LFM and BMG and drawn by BMG. Special thanks are

925 given to Arnaud Davoigny, LSCE, Gif-sur-Yvette, France, for providing technical support in

926 U/Th dating procedures. Many thanks are expressed to the three anonymous reviewers for

927 allowing enriching debates and thus for helping to improve the original draft.
928 **Data availability**

929 Datasets related to this article are accessible through the dataDOI repository at:

930 https://doi.org/10.17632/cjfn8fg8yt.1

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1284 **Figure captions**

1285 **Figure 1.** A: Google Earth location map of Fakarava Atoll in the Tuamotu Archipelago, French Polynesia, central south Pacific. B: Google Earth map of Fakarava Atoll, with location of the studied areas (Fig. 1C) along the south-eastern reef rim margin. C: aerial view of the south-eastern rim part showing the studied boulder-bearing areas: Zone A (west) and Zone B (east).

1290 **Figure 2.** Google Earth detailed view of the spatial distribution of reef boulders in Zones A (west) and B (east), south-eastern reef rim margin, Fakarava Atoll. The boulders are identified by a specific letter or number. Twenty-two boulders are counted in each zone and their azimuth measured.

1294 **Figure 3.** Drone- acquired aerial oblique view of the both reef boulder fields, south-eastern reef rim margin, Fakarava Atoll: at the forefront, Zone A (east); in the background, Zone B (east). Note the convex-shaped morphology of the reef rim line on which swell surges have concentrated during extreme marine events, thus promoting dislocation of the upper forereef zone and detachment and transport of pieces of spurs. Courtesy of Hugo Dayan.
1299 Figure 4. Close-up views of in situ-coral colonies on top of upper boulder surfaces, south-eastern reef rim margin, Fakarava Atoll. A: Boulder D, Zone A (west). B: Boulder T, Zone B (east).
1300
1301

1302 Figure 5. Example of a drone-acquired photogrammetry 3D image of a coral reef boulder
1303 (Boulder X, Zone B, east).

1304 Figure 6. Drone-acquired aerial views of sections of upper fore-reef spurs, south-eastern reef
1305 rim margin, Fakarava Atoll. A: fault line cutting spurs and detached spur-derived boulders. B:
1306 dislocated section of the spur-and-groove system, with detached spur-derived boulders. C:
1307 partially preserved forereef spurs; note the occurrence of an isolated, spur-derived boulder
1308 resting on top of a spur. D: section of the spur-and-groove system, showing truncation scars
1309 affecting spurs.

1310 Figure 7. Morphology, zonation and interpretative sub-surface stratigraphy of the south-eastern, outer reef rim margin, Fakarava Atoll. The boulder pre-transport settings (submerged, emerged) and expected transport modes are indicated, with the forces respectively applied to the two location types of coral reef boulders – Fl: lift; Ff: friction; Fd: drag; Fr: restraining. The letters a, b and c plotted on the emerged boulder lying on the reef flalt zone refers respectively to its length (a-axis), width (b-axis) and height (c-axis). Pmlt: present mean low tide level. No scale is given.

1317 Figure 8. Orientation of reef boulders measured from magnetic north, south-eastern reef rim
1318 margin, Fakarava Atoll. A: Zone A (west), 22 measured boulders. B: Zone B (east), 22
1319 measured boulders. See figure 2 for location of boulders.

1320 Figure 9. Plotting of a-long axis versus b-short axis and c- height axis of the 19 U/Th dated
1321 coral reef boulders, south-eastern reef rim margin, Fakarava Atoll. Blue symbols: b-short
1322 axis; red symbols: c-height axis.

1323 Figure 10. Drone-acquired aerial view showing the fracturing of Boulder D, south-eastern reef
1324 rim margin, Fakarava Atoll. Postulated age of deposition: 361 ± 20 cal. yr BP. The fracturing
1325 is postulated to be due to the mechanic impact by Boulder C – postulated age of deposition:
1326 282 ± 35 cal. yr BP –) deposited about one century later or during the same event.
Figure 11. Selected side and top views of some coral reef boulders, south-eastern reef rim margin, Fakarava Atoll. A, B: Boulder X. C, D: Boulder K. E, F: Boulder A.

Figure 12. Weight (tons) versus distance (m) from the reef front line for the 19 investigated reef boulders, south-eastern reef rim margin, Fakarava Atoll. Note the apparently linear weight decreasing trend of most of the boulders with the increasing distance from the reef edge. Despite their location at distances less than 100 m from the reef front, the heaviest boulders (A and X) do not follow the same distributional trend.

Figure 13. Occurrences of expected extreme marine hazard events (super-cyclones) during the last 3,500 years as recorded along the south-eastern reef-rim side. Vertical red lines refer to the dated coral reef boulders and each line could represent one single hazard event. Assuming a minimum number of six hazard events, every shaded (grey) band refers to boulder clusters expected to have been deposited during a same extreme marine hazard event.

Figure 14. Stratigraphical location of U/Th dated coral samples from Boulder D, Zone A (west), south-eastern reef rim margin, Fakarava Atoll. The age of each sample, given in brackets, is expressed in cal. yr BP (before 2024).

Figure 15. Relationships between the ages of deposition for the 19 U/Th dated coral reef boulders (south-eastern reef rim margin, Fakarava Atoll) and ENSO-controlled cyclogenesis in the south tropical Pacific over the last 3,500 years. A. time distribution of the dated boulders. B: cyclone activity over the past 3000 years at Tahaa Island, Society archipelago, French Polynesia. Cyclogenesis is based on grain-size analysis, i.e. percentage of coarse fraction > 250 µm, in a core from lagoonal sediments (Toomey et al., 2013; Bramante et al., 2020). C: red curve: record of ENSO activity based on measurement of sand percentage in a core extracted from Juna Crater Lake, Galapagos, modified from Conroy et al. (2008); blue curve: record of ENSO activity based on sediment laminae light-intensity in a core extracted from Laguna Pallcacocha, Ecuador, modified from Moy et al. (2002). D: coral-based reconstructions of ENSO variance in the central Pacific (Northern Line islands) over the late Holocene, modified from Cobb et al. (2013).

Table 1. Uranium-thorium dating of selected coral samples from top surface reef boulders, south-eastern reef rim, Fakarava Atoll. Ages are given in calendar years before present (cal. yr BP), relative to 2024 and in years BC (Before Jesus Christ) or AD (Anno Domini). Age
discrepancies between the dated coral samples are given. See Table S1 (supplementary material) for error ranges associated with ages.

Table 2. Morphometric attributes (length, width, height, volume, weight and distance of location relative to the reef front line) of the 19 U/Th dated reef boulders, south-eastern reef rim, Fakarava Atoll, partly based on photogrammetric imagery analysis.

Table 3. Pre-displacement setting and flow velocities (m/s) of four selected boulders, from the south-eastern reef-rim side, Fakarava Atoll (north-west Tuamotu). As the bed angle of the boulders at the pre-displacement location was assumed to be $\theta = 0$, there is no difference between the values of lifting motion for both joint-bounded and non-joint bounded boulders.

Table S1. Supplementary material: uranium-thorium dating of coral samples extracted from reef boulders, south-eastern reef rim, Fakarava Atoll. Are given successively laboratory sample codes, field sample (F, FAK) numbers, uranium and thorium contents, isotopic composition with statistical errors – two standard deviations of the mean – and ages. Uranium and thorium contents and isotopic ratios $\delta^{234}$UM = {[(($^{234}$U/$^{238}$U)$_{sample}$ / ($^{234}$U/$^{238}$U)$_{eq}$]–1} $\times$ 1000, where ($^{234}$U/$^{238}$U) sample is the measured atomic ratio and ($^{234}$U/$^{238}$U)$_{eq}$ is the atomic ratio at secular equilibrium. $\delta^{234}$U$_{(T)}$ is the initial value and is calculated by the equation: $\delta^{234}$U(0) = $\delta^{234}$U$_{means}$ exp($\lambda_{234t}$), where t is the age in years and $\lambda_{234}$ is the decay constant for $^{234}$U. Ages were corrected for authigenic $^{230}$Th using $^{230}$Th/$^{232}$Th) of the detrital fraction = 10 $\pm$ 5. Ages are expressed in both calendar years before present (BP), relative to 2024, corrected for detrital content using a $^{230}$Th/$^{232}$Th isotopic ratio = 7 $\pm$ 50% and in conventional ages BP (Before Present, relative to 1950).
## A FRENCH POLYNESIA

### South Pacific Ocean

<figure><img src="image_3.png" /></figure>

B

Rotoava

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

Pass Garuae

Studied area
(Fig. 1C)

Tetamanu

Pass
Tumakohua

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

N
6 km

N
6 km

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

### Zone A
Fig. 2a

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

16°22'35.60"S
145°25'58.48"W

16°22'36.45"S
145°25'47.02"W

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

### Zone B
Fig. 2b

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

N
60 m
A

Reef flat zone

A

Conglomerate
platform

3

2
1

Ho'a

N

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

30 m

Lagoon

C

B

I

D

E

J

H

F

α

G

14

15

16

17

K

L

13

18

10 m

B

R

W

Q

S

U

T

8

6

7

12

9

10

11

P3

P1

P2

P4

P5

P7

P6

Motu

N

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

3 m

Reef flat zone

V

Conglomerate
platform

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

Zone B

Zone A

N

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

<figure><img src="image_14.png" /></figure>
<figure><img src="image_15.png" /></figure>

N

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

1 m

<figure><img src="image_17.png" /></figure>
<figure><img src="image_18.png" /><figcaption>Reef crest</figcaption></figure>

<figure><img src="image_19.png" /><figcaption>Reef crest</figcaption></figure>

<figure><img src="image_20.png" /><figcaption>Reef crest</figcaption></figure>

<figure><img src="image_21.png" /><figcaption>Reef crest</figcaption></figure>
<figure><img src="image_22.png" /></figure>

Spur-and-groove system

Emerged boulder
– sliding, rolling
or lifting

REEF FLAT ZONE

MOTU

Conglomerate
platform

HOLOCENE REEF UNIT

PRE-HOLOCENE
REEF UNIT
A

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

B

<figure><img src="image_24.png" /></figure>
<figure><img src="image_25.png" /></figure>

$$y = 0.5632x + 0.6169$$

$$R^2 = 0.5463$$

$$y = 0.3306x + 0.4679$$

$$R^2 = 0.4482$$
<figure><img src="image_26.png" /><figcaption>B</figcaption></figure>

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

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

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

N

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

1 m
<figure><img src="image_31.png" /><figcaption>B</figcaption></figure>

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

<figure><img src="image_33.png" /><figcaption>D</figcaption></figure>

<figure><img src="image_34.png" /><figcaption>E</figcaption></figure>

<figure><img src="image_35.png" /><figcaption>F</figcaption></figure>

<figure><img src="image_36.png" /></figure>
<figure><img src="image_37.png" /></figure>

$$y = 139,17 x^{-0,291}$$
<figure><img src="image_38.png" /></figure>

Number of
expected
supercyclone
events

Boulder
number

G J E

L 183

1

14

2 KP₅ 16

DBC
Section North

<figure><img src="image_39.png" /></figure>

Boulder D

<figure><img src="image_40.png" /></figure>

Section West

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

Section East

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

Section South

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

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

B

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

C

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

D

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