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Geophys. J. Int. (2020) 223, 22–44 doi: 10.1093/gji/ggaa273
Advance Access publication 2020 June 03
GJI Heat Flow and Volcanology
The 2018–2019 seismo-volcanic crisis east of Mayotte, Comoros
islands: seismicity and ground deformation markers of an
exceptional submarine eruption
Anne Lemoine ,1Pierre Briole ,2Didier Bertil,1Agathe Roull´
e,1Michael Foumelis,1
Isabelle Thinon,3Daniel Raucoules,1Marcello de Michele,1Pierre Valty4and
Roser Hoste Colomer1
1BRGM- French Geological Survey, Direction Risques et Pr´
evention, 3, av. C. Guillemin, 45060 Orl´
eans, France. E-mail: a.lemoine@brgm.fr
2UMR 8538 CNRS - Ecole Normale Sup´
erieure - PSL Research University, Laboratoire de G´
eologie, 24, r. Lhomond, 75231 Paris, France
3BRGM-French Geological Survey, Direction des G´
eoressources, 3, av. C. Guillemin, 45060 Orl´
eans, France
4IGN-Institut National de l’Information G´
eographique et Foresti`
ere, Service de G´
eod´
esie et M´
etrologie, 73, av. de Paris, 94165 St Mand´
e, France
Accepted 2020 May 25. Received 2020 May 18; in original form 2019 October 22
SUMMARY
On 10 May 2018, an unprecedented long and intense seismic crisis started offshore, east
of Mayotte, the easternmost of the Comoros volcanic islands. The population felt hundreds
of events. Over the course of 1 yr, 32 earthquakes with magnitude greater than 5 occurred,
including the largest event ever recorded in the Comoros (Mw=5.9 on 15 May 2018).
Earthquakes are clustered in space and time. Unusual intense long lasting monochromatic
very long period events were also registered. From early July 2018, Global Navigation Satellite
System (GNSS) stations and Interferometric Synthetic Aperture Radar (InSAR) registered a
large drift, testimony of a large offshore deflation. We describe the onset and the evolution of
a large magmatic event thanks to the analysis of the seismicity from the initiation of the crisis
through its first year, compared to the ground deformation observation (GNSS and InSAR) and
modelling. We discriminate and characterize the initial fracturing phase, the phase of magma
intrusion and dyke propagation from depth to the subsurface, and the eruptive phase that starts
on 3 July 2018, around 50 d after the first seismic events. The eruption is not terminated
2 yr after its initiation, with the persistence of an unusual seismicity, whose pattern has
been similar since summer 2018, including episodic very low frequency events presenting a
harmonic oscillation with a period of ∼16 s. From July 2018, the whole Mayotte Island drifted
eastward and downward at a slightly increasing rate until reaching a peak in late 2018. At the
apex, the mean deformation rate was 224 mm yr−1eastward and 186 mm yr−1downward.
During 2019, the deformation smoothly decreased and in January 2020, it was less than
20 per cent of its peak value. A deflation model of a magma reservoir buried in a homogenous
half space fits well the data. The modelled reservoir is located 45 ±5 km east of Mayotte, at
a depth of 28 ±3 km and the inferred magma extraction at the apex was ∼94 m3s−1.The
introduction of a small secondary source located beneath Mayotte Island at the same depth
as the main one improves the fit by 20 per cent. While the rate of the main source drops by
a factor of 5 during 2019, the rate of the secondary source remains stable. This might be a
clue of the occurrence of relaxation at depth that may continue for some time after the end
of the eruption. According to our model, the total volume extracted from the deep reservoir
was ∼2.65 km3in January 2020. This is the largest offshore volcanic event ever quantitatively
documented. This seismo-volcanic crisis is consistent with the trans-tensional regime along
Comoros archipelago.
Key words: Indian Ocean; Seismicity and tectonics; Submarine tectonics and volcanism;
Effusive volcanism; Remote sensing of volcanoes; Volcano seismology.
22
C
The Author(s) 2020. Published by Oxford University Press on behalf of The Royal Astronomical Society. This is an Open Access
article distributed under the terms of the Creative Commons Attribution License (http://creativecommons.org/licenses/by/4.0/), which
permits unrestricted reuse, distribution, and reproduction in any medium, provided the original work is properly cited.
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The Mayotte 2018–2019 seismo-volcanic crisis 23
1 INTRODUCTION
The volcanic Comoros archipelago (Fig. 1) is composed of four
main volcanic islands, from west to east: Grande Comore, Moh´
eli,
Anjouan and Mayotte, and further east, the Geiser and Leven vol-
canic banks (Daniel et al. 1972). Volcanism spans between the
Miocene and Holocene. In Mayotte Island, the earliest phase of
magmatic activity is dated 11 Ma (Debeuf 2004; Pelleter et al.
2014), whereas the latest reported volcanic event occurred within a
few 1000 yr in Petite Terre, East of Mayotte (Zinke et al. 2003a,b).
At the present time, there is a diffuse and moderate seismicity in
the archipelago and few historically felt events.
The Mayotte volcano-seismic sequence began on 10 May 2018.
Over 1 yr, we recorded 32 earthquakes with magnitudes greater
than 5, the largest one reaching Mw=5.9 on 15 May 2018, the
largest event ever recorded in Comoros. Such intense and long seis-
mic activity had never been recorded before in the archipelago.
More than 1 yr after the onset of the crisis, and after ∼2 months of
intense seismic unrest at the beginning of the sequence, the seismic-
ity was persisting with a large number of small earthquakes (Bertil
et al. 2019) and still ongoing ground deformations are observed
by Global Navigation Satellite System (GNSS) and Interferomet-
ric Synthetic Aperture Radar (InSAR). There were also frequent
episodic monochromatic very long period (VLP) earthquakes (Poli
et al. 2019; Satriano et al. 2019; Cesca et al. 2020; Feuillet et al.
in review) typical of active magmatic or hydrothermal processes as
described in other cases (e.g. Chouet 2003;McNutt2005; Shapiro
et al. 2017).
In an active volcanic context, magma movement can cause seis-
mic crisis and volcano-tectonic earthquakes migrations can lead to
potential eruptions (e.g. Toda et al. 2002; Battaglia et al. 2005;
L´
opez et al. 2012;Mart
´
ıet al. 2013; Gudmundsson et al. 2014;
Sigmundsson et al. 2015;´
Ag´
ustsd´
ottir et al. 2016). There are sev-
eral past examples of intense seismic activity associated with de-
formation in volcanic contexts. Such observations can be markers
of magma intrusion especially through propagating and opening
dykes, as in the Gulf of Aden, between November 2010 and March
2011 (Ahmed et al. 2016), where the spatio-temporal evolution of
intense seismicity, with 29 M≥5 earthquakes, was interpreted as
the sign of magma ascent, followed by the propagation of dykes.
Regarding the intensity of the largest events, in Izu islands (Japan)
in 2000 (Toda et al. 2002), five M≥6 events occurred within a few
weeks associated with the eruptions of Miyakejima volcano and
dyke intrusion.
During seismo-volcanic crises and magma intrusion into dykes
processes, swarms of small to moderate earthquakes (M<6) are
commonly observed. As for example in Afar (Ethiopia) where a
phenomenon that began in 2005 involved seismicity organized in
clusters, including M>5 earthquakes: between 2005 and 2010, 14
dyke intrusions were identified along a 60-km-long rift segment and
earthquakes are organized in clusters, that is distributed at spatially
reduced areas close to the dykes (e.g. Ayele et al. 2007; Grandin et al.
2011). Another example illustrating both seismicity and ground
deformation patterns as markers of magmatic phenomenon is the
rifting event near the B´
ardarbunga volcanic system (Iceland) that
lasted from 16 August to 6 September 2014. As revealed by seismic-
ity and ground deformation, it involved a 45-km-long segmented
dyke intrusion delineated by earthquakes, magma source deflation
and slow collapse of a caldera, magnitude M>5 earthquakes and
effusive fissure eruption (Gudmundsson et al. 2014; Sigmundsson
et al. 2015;´
Ag´
ustsd´
ottir et al.2016,2019). Finally, the El Hierro
eruption (Canary Islands) initiated on 10 October 2011 and had
been preceded since July 2011 by intense and fluctuant seismicity
(including acceleration and migration, not exceeding ML=4.3) and
by the concomitant evolution of deformation patterns. Pre-eruptive
phases were distinguished, especially according to seismicity and
displacement monitoring (L´
opez et al. 2012). Analysis of earth-
quake migration and surface deformation have shown that, after the
initial fracturing and intrusion episode, a migration phase appeared
and extended for ∼20 km, preceding a submarine eruption (Mart´
ı
et al. 2013).
The 2018–2019 seismo-volcanic activity at Mayotte is unusual
due to its duration and the intensity of its seismic sequence and
deformations, marking the end of a volcanic quiescence period in
the area. The crisis has been monitored since the beginning of the
unrest with an evolving seismic network, originally not designed
for accurate monitoring volcano seismicity. From 10 May 2018 to
15 May 2019, the seismic monitoring network remained sparse, but
allowed for the analysis of the evolution of the seismicity pattern.
Permanent GNSS stations and InSAR data registered different dis-
placement patterns that we identified as different eruptive phases.
Here we present an analysis of the first year of the Mayotte
2018–2019 seismo-volcanic episode combining both seismological
analysis and deformation observations (GNSS and InSAR) before
the significant improvement of the initial seismic monitoring net-
work in May 2019 operated through the installation of ocean bottom
seismometers (OBS). We propose a model of scenario that explains,
with different phases, the spatio-temporal evolution of the seismic-
ity and the geodetic observations.
2 GEOLOGICAL SETTING AND
REGIONAL CONTEXT
The Comoros Islands form an overall east–west (E–W) trending
archipelago located north of the Mozambique Channel, between
the eastern coast of Mozambique and the northern tip of Mada-
gascar (Fig. 1). North of the Comoros Archipelago, most authors
agree that the crust of the Somali basin is oceanic dating to the
Mesozoic time (S´
egoufin & Patriat 1981; Rabinowitz et al. 1983;
Malod et al. 1991; Sauter et al. 2018) from 152 to 120 Ma (Davis
et al. 2016). South to the archipelago, the nature of the Comoros
basin is debated between a thinned continental crust (e.g. Flower
& Strong 1969; Bassias & Leclaire 1990;Roachet al. 2017)
and an oceanic crust setting during the Jurassic Magnetic Quiet
Zone (Talwani 1962; Recq 1982; Rabinowitz et al. 1983; Klimke
et al. 2016). The present-day morphology of this area arose from
the Permian-Triassic Karro northwest–southeast (NW–SE) rifting,
which resulted in the separation between Gondwana and Madagas-
car continental blocks (e.g. Malod et al. 1991;Daviset al. 2016).
The Mozambique and Somalia basins opened while Madagascar
Island drifted southward along the Davie ridge (Fig. 1). The tec-
tonic regime in the Mozambique Channel is dominated by overall
east–northeast by west–southwest (ENE–WSW) extension, which
is also identified with the East African Rift System and Madagascar
(e.g. Bertil & Regnoult 1998; Piqu´
e1999;Delvaux&Barth2010).
The Comoros archipelago began emplacement during the Ceno-
zoic in the northern Mozambique Channel. Michon (2016) proposed
that magmatic activity first appeared in Mayotte, ∼20 Ma and then
in Anjouan, Moh´
eli, and Grande Comore ∼10 Ma. Comoros vol-
canism seems synchronous with volcanism in surrounding areas,
that is in Madagascar and in some provinces linked to the East-
African Rift System (e.g. Nougier et al. 1986; Debeuf 2004; Michon
2016). According to rock geochronology (e.g. Nougier et al. 1986;
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24 A. Lemoine et al.
Figure 1. Velocity map for 36 GNSS stations around Mayotte (set to zero) from the time-series published by the Nevada geodetic laboratory (Blewitt et al.
2018). The values are found in Table S1. A clockwise correction of 2.1◦Myr−1was applied to cancel the overall rotation around Mayotte. Main features in
the area of the East African Rift are plotted (Davie Ridge, Eastern Rift, Main Ethiopian Rift, Western Rift, Malawi Rift and Assoua Fracture Zone). The plate
boundaries in white lines are modified from Saria et al. (2014) and Stamps et al. (2018). Elevation grid is from Gebco 2014 (Weatherall et al. 2015). Inset: The
Comoros islands surrounded by the Somali and Comoros basins. Bathymetry from Gebco 2014 completed around Mayotte by SHOM (2015) data. J. represents
Les Jumelles: bathymetric edifices lying NE from Mayotte (visible on maps from Fig. 4).
Debeuf 2004; Pelleter et al. 2014), the youngest volcanic activ-
ity is in Grande Comore (0.13 ±0.02 Myr to present) with the
active Karthala volcano, and the oldest was reported in Mayotte
(∼11 Myr), with a similar age reported in Anjouan. In Mayotte,
these authors identify different eruptive phases separated by quies-
cent periods. The northeast area of Mayotte is composed by a more
recent volcanic complex dated from 500–150 kyr in Grande-Terre
and up to 4 kyr in Petite-Terre (Nehlig et al. 2013; Pelleter et al.
2014). According to Zinke et al.2003a,b), the latest volcanic events
documented are some ash deposits in the barrier reef between 4 and
7 kyr. Offshore, on the insular slope of Mayotte, tens of small vol-
canic seamounts are observed in the north, northwest and northeast
parts (Audru et al. 2006) and is mainly distributed off Petite Terre
along a WNW–ESE ridge where a new active volcanic structure
has recently been documented (Paquet et al. 2019; Feuillet et al. in
rev.).
The origin of the Comoros volcanism is not well understood yet.
It could be due to the presence of a hot spot, based on the west-
ward migration of volcanism age (e.g. Emerick & Duncan 1982),
or to the influence of lithospheric fractures (Nougier et al. 1986).
Recent work (Debeuf 2004; Michon 2016) suggests a mixed so-
lution; regional extensive tectonics interacting with asthenospheric
processes.
A catalogue of instrumental seismicity of the Comoros was
extracted for 1964–2018 from the International Seismological
Centre catalogue (ISC 2016). For 1978–1995, it was completed
using the Malagasy catalogue (Bertil & Regnoult 1998), which
lowers the detection threshold, depending on the distance from the
Tananarive observatory (from mb≈4.0 eastward from Mayotte
to mb≈4.3 around Grande Comore). The location accuracy is
poor, up to 30 km for many events in this poorly instrumented
region. The Mozambique Channel seismicity is marked by a north–
south (N–S) band of activity (Fig. 2), associated with the Davie
ridge (Fig. 1) showing normal faulting mechanisms (Grimison &
Chen 1988). A zone of diffuse seismicity, ∼100 km wide, runs
E–W from 48◦E and 42.5◦E along the volcanic line of the Co-
moros islands (Fig. 2). From 1982 to 2016, eight moderate earth-
quakes occurred (Mw5.0–5.3), spread throughout this band pre-
senting normal faulting and strike-slip earthquakes, especially east
of Mayotte (Fig. 2). In Madagascar, a moderate and distributed
seismicity is associated with inherited structures and an overall
E–W extension (Bertil & Regnoult 1998; Rindraharisaona et al.
2013).
Several branches of the Cenozoic East African Rift System
(EARS) form a distributed plate boundary zone between the Nu-
bia and Somalia plates (e.g. Calais et al. 2006;Delvaux&Barth
2010;D
´
eprez et al. 2013; Stamps et al. 2018) where several authors
(Hartnady 2002; Horner-Johnson et al. 2007;Sariaet al. 2014;
Stamps et al. 2018) defined three plates: the Victoria, Rovuma and
Lwandle blocks (Fig. 1). The boundaries of the Lwandle block are
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The Mayotte 2018–2019 seismo-volcanic crisis 25
Figure 2. The 1901–2018 seismicity for M>4 is created by merging the International seismological centre (ISC 2016) catalogue and regional catalogue (Bertil
& Regnoult 1998; Bertil et al. 2019). The focal mechanisms for 1976–2019 are from the Global centroid moment tensor catalogue (G-CMT, Dziewonski et al.
1981;Ekstr
¨
om et al. 2012). The plate boundaries represented by white lines are modified from Saria et al. (2014) and Stamps et al. (2018). The bathymetry
and topography are from Gebco 2014 (Weatherall et al. 2015). Green triangles show the location of seismic stations forming our monitoring network (RESIF
1995;Edusismo; Scripps Institution of Oceanography 1986; GEOFON Data Centre 1993; the Karthala Observatory and the Bureau Central Sismologique
Franc¸ais, Table S2).
potentially diffuse and poorly constrained, especially with Soma-
lia (Horner-Johnson et al. 2007; DeMets et al. 2010;Sariaet al.
2014).
The overall actual trend of E–W extension over the East African
region is illustrated by seismicity pattern (Fig. 2), geodetic obser-
vations, and the expression of extensive tectonics from EARS to
Madagascar, including Mozambique Channel and Comoros vol-
canic axis (e.g. Bertil & Regnoult 1998; Delvaux & Barth 2010;
Rindraharisaona et al. 2013; Stamps et al. 2018). Deville et al.
(2018) reported offshore extensional and trans-tensional fault sys-
tems, respectively, around the Davie Ridge and South of it in the
Mozambique Channel. They linked these active fault zones as-
sociated with volcanic activity with the southernmost part of the
eastern branch of the EARS. Along the Comoros axis, the few
available focal mechanisms show normal faulting and strike slip,
compatible in terms of orientation to a NE–SW tensional axis. From
onshore field observations on Mayotte, Anjouan and Moh´
eli islands
(Fig. 1), Famin et al. (2020) proposed that the Comores archipelago
delineates the northern boundary between Somalia and Lwandle
plates, undergoing strike-slip stress field associated with a NW–SE
shortening (where reverse faulting, strike slip and en ´
echelon ex-
tensional fractures coexist). Whereas, from offshore data, Feuillet
et al. (in rev.) suggested that this diffuse E–W striking area accom-
modates the deformation between offshore East African rifts system
and Madagascar grabens, thanks to major extensional features in a
trans-tensional area.
The NE–SW extension across the line formed by the Comoros
islands is also revealed by GNSS data. Fig. 1shows the velocities
of the GNSS stations in the area (Table S1) with respect to the
station MAYG in Mayotte. North of 19◦S where the Davie Ridge
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26 A. Lemoine et al.
is bending, both GNSS (Fig. 1) and focal mechanisms (Fig. 2)
confirm the presence of a NE–SW tensional axis. The GNSS sta-
tion distribution is uneven and particularly sparse around Comoros.
The relative velocity of SEYG (Seychelles) with respect to the sta-
tions PMBA and NMPL located in Mozambique, on the Rovuma
Plate, is of 2.6 ±0.5 mm yr−1in the azimuth N50◦.AstheCo-
moros axis is associated with volcanism since 20 Ma and seis-
mically active, while there is no identified seismicity within the
Somalia basin and between Seychelles and the Comoros, a sig-
nificant part of this NE–SW extension between Seychelles and
Mozambique could well be accommodated across the volcanic
line of the Comoros, through oblique extension along the volcanic
axis.
This relative motion possibly accommodated in the Comoros
would fit with the focal mechanisms reported in Fig. 2along the
Comoros archipelago: normal faulting and strike-slip events, W
and SSE of the Comoros that could indicate a trans-tensional stress
regime.
Furthermore, VHMR and MTVE GNSS stations (Fig. 1), re-
spectively, at the northern tip of Madagascar and in the Tanza-
nia coastal area, point in opposite directions (relatively to Mayotte
station MAYG), which suggest an overall and significant counter-
clockwise rotation in the area. South of the Davie ridge is dom-
inated by left lateral shear at a rate of 1 mm yr−1, according to
the difference of velocities between PMBA and NMPL on the
Rovuma Plate and ABPO, VOIM and ZMBT on the Lwandle Plate
(Fig. 1).
Of note, between the Neogene and Quaternary, tectonic and vol-
canic activities occurred in several branches in EARS (e.g. D´
eprez
et al. 2013), but also in several part of Madagascar, especially in
its northern part (e.g. Rindraharisaona et al. 2013; Michon 2016)
and in Comoros (Nougier et al. 1986; Debeuf 2004; Michon 2016).
Michon (2016) proposed that the Comoros archipelago could poten-
tially coincide with the eastern border of the EARS. Debeuf (2004)
discussed the influence of extension linked to the East African
Rift on Comoros volcanism. She mentioned the presence of an an-
cient regional lithospheric structure: the Assoua fracture zone (east
Africa), a Precambrian lineament of lithospheric dimension that,
despite its age, could have localized the ascent of asthenospheric
plumes in the recent geological past.
3 SEISMICITY OF MAYOTTE AND THE
SEISMIC CRISIS OF 2018–2019
3.1 Historical seismicity
There are only a few testimonies of historical seismicity in the Co-
moros before the 20th century. Hachim (2004) identified damaging
earthquakes in Mayotte in 1606, 1679 and 1788 from oral transmis-
sion without being able to constrain testimonies from bibliographi-
cal data. Gevrey (1870) and Vienne (1900) quote felt earthquakes in
the Comoros Archipelago in 1808, 1829 and 1865, without any de-
tail. Moreover, Lambert (1997) associated the 1829 event not with
an earthquake but with a cyclone. From 1910 to 1960, the Mala-
gasy academy published a catalogue of locally felt earthquakes
made by the Tananarive observatory. The 16 January 1936 earth-
quake, west of Mayotte, was felt across the whole Comoros, causing
moderate damage in several municipalities (La D´
epˆ
eche de Mada-
gascar, 19/02/1936, BNF). More recently, regional felt earthquakes
have been reported in Sisfrance Oc´
ean Indien (2010). No historical
chronicle documents any seismic cluster in Comoros.
3.2 Instrumental seismicity
Regarding the instrumental period, the catalogue of seismicity we
built for the Comoros (see preceding chapter) contains some events
within 150 km near the Mayotte edifice. Some events occur red south
of Mayotte, including a mb=5.0 event, on 23 April 1993, located
40 km south of the island. In 2007, at ∼150 km east of Mayotte, a
patch of seismicity seems to be located around Z´
el´
ee bank, includ-
ing a magnitude 5.1 event on 23 June 2007 showing a strike-slip
fault mechanism processed by the Global Centroid Moment Tensor
(G-CMT, Dziewonski et al. 1981;Ekstr
¨
om et al. 2012). The most
significant event close to Mayotte before 2018 occurred on 1 De-
cember 1993, located 40 km west of the island with mb=5.2 and
associated with moderate damage (Lambert 1997). Close to this
event, a normal faulting earthquake occurred on 9 September 2011
(Mw=5.0). No significant event was reported near the 2018–2019
seismic sequence during the instrumental period, in a poorly instru-
mented region associated with high detection magnitude threshold
(∼4.5). Moreover, no long lasting seismic sequence is reported in
the instrumental catalogue of the Comoros.
3.3 The 2018–2019 seismic sequence
At the beginning of the sequence, the closest broad-band sta-
tions from international seismic networks operating in this re-
gion were more than 650 km away (Fig. 2). In Mayotte, the ac-
celerometric station, YTMZ from the French RESIF-RAP network
(R´
eseau Acc´
el´
erom´
etrique Permanent, RESIF 1995) was active
in real time since 2016. To monitor the crisis, a virtual network
was set up, composed of YTMZ and regional broad-band stations.
This network evolved in time with the addition of local and re-
gional stations until mid-July 2018 (Supporting Information and
Tab le S 2) .
Here we present an analysis of the swarm of volcano-tectonic
earthquakes from the beginning of the crisis until 15 May 2019,
when new onshore and offshore stations allow a good azimuthal
coverage of the seismic sequence and a better analysis of seismicity
distribution (Saurel et al. 2019; Feuillet et al. in rev.).
The YTMZ station allows finer monitoring of the crisis than
global networks. Thanks to it, from 10 May 2018 to 15 May 2019,
1872 events with a local magnitude ML(Supporting Information)
greater than or equal to 3.5 could be detected (161 and 32 of them
corresponding, respectively, to magnitude greater than or equal to
4.5 and 5, Fig. 3). Only 1163 of these 1872 events could be located
with crustal phase data for distance less than 1400 km (Figs 4
and 5).
The location configuration (described in Supporting Information)
have been built and tested on earthquakes that occurred after the
improvement of the seismic network, that is after the period analysed
in this paper. It makes it possible to limit epicentre shifts with
the seismicity pattern observed beyond May 2019 thanks to the
seismic network improvement (Lemoine et al. 2019;Saurelet al.
2019).
Accuracy of the location depends directly on the network config-
uration that evolved with time and on the magnitude (on azimuthal
coverage of the network and on number of manually picked seis-
mic phases available for location). Epicentral location errors range
from 10 km in the most unfavourable configuration to 1 km in the
most favourable. Depth could not be retrieved from hypo71 (Lee
&Valdes1985) for some events associated to a reduced number
of picked phases, especially until latest network improvement (i.e.
15 July 2018). In such cases, location is made fixing the depth at
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The Mayotte 2018–2019 seismo-volcanic crisis 27
(a)
(b)
(c) (d)
(e)
Figure 3. (a) Magnitude distribution for ML>3.5 events. Three clusters are identified from S–P time evolution (d): cluster 1 (blue) is mainly active from 10
May 2018 to early July 2018. Cluster 2 (yellow) starts on 26 June 2018 and remains active until the end of the observation period. Cluster 3 (red) begins inJuly
2018 and is still active on 15 May 2019. (b) Daily number of ML≥3.5 events for each cluster (blue for cluster 1, yellow for cluster 2, and red for cluster 3).
(c) S–Ptime distribution at YTMZ station by cluster. (d) Evolution of S–P time from YTMZ station with time. (e) Cumulative seismic moment for the three
clusters.
30 km (depth average in the best network configuration is around
32 km). 30 km is also in the range of depth values from G-CMT
locations (Fig. 6) and permit to retrieve the seismicity pattern during
the whole crisis (Saurel et al. 2019). In other cases, average vertical
location error is estimated to be equivalent to horizontal error (i.e.
few km). However, depth distribution depends on the velocity model
which is far from being constrained around Mayotte and surround-
ing areas (e.g. Jacques et al. 2019). Finally, we could state that the
locations including more than 10 manually picked seismic phases,
are stable, especially horizontally whereas depth distribution should
be considered with caution.
The crisis started on the morning of 10 May 2018 with small
events detected by the YTMZ accelerometer only. In the evening, a
ML=4.3 earthquake, not detected by the global networks, was the
first event felt in Mayotte. From May 13 to 15, the activity increased
gradually with four events with MLbetween 4.5 and 5 (Fig. 3).
The main event, and largest ever recorded in the Comoros, with
magnitude ML=5.8 and Mw=5.9 (G-CMT), occurred on 15 May
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28 A. Lemoine et al.
Figure 4. Top panel: localized earthquake map with ML≥3.5 from 10 May 2018 to 15 May 2019. The first cluster (light to dark blue) is active mainly from
10 May 2018 to early July 2018, the second (yellow to light orange) from 26 June 2018, and the third (red to pink) from mid July 2018. Within each cluster,
colour of epicentres evolves with time (days from the beginning of the crisis, that is 10 May 2018). Green points represent earthquakes registered during
the observation period but out of the clustering seismicity (one ML≥3.5 event is reported westward from Mayotte). The bathymetry is from the Homonim
project (SHOM 2015) and Gebco 2014 (Weatherall et al. 2015). Bottom panel: seismicity plotted on a vertical cross section between points A and B, only
resolved depths are plotted (otherwise, the depth is fixed at 30 km). Yellow triangles indicate the GNSS stations. Epicentres marked by a light gray circle and
transparency were localized with less than 10 phases and considered as less relevant (map and cross section).
2018. Then the seismicity was relatively steady for the following
∼15 d with 10 to 30 ML>3.5 events per day. Two strong events
occurred on May 20 and 21 (Mw=5.5). Between June 1 and 7, the
seismicity culminated with 52 to 87 ML≥3.5 events daily, with the
highest activity on 1 June 2018. After 7 June 2018, the seismicity
decreased. It increased again between 19 and 27, June 2018 with
several ML≥5.0 events with a new peak of activity on 23 June
2018 (50 ML≥3.5 events). The period between July 10 and August
26 was quiescent with a sharp drop in the number and magnitude of
events (Fig. 3). On August 26, eight events with magnitude between
ML=4 and 4.5 occurred (27 ML≥3.5 events), initiating a renewal
of activity with two ML=4.8 events occurring on October 16
and November 7, marking the last day associated with such a high
level of activity (afterward, the maximal daily number of ML≥3.5
events were 14). However, their contribution to the total moment
release remains low compared to the first weeks of activity (Fig. 3).
Since October 2018, a relatively steady level of active seismicity is
observed that began to lower at the end of April 2019 (Fig. 3).
The cumulated seismic moment for the whole period corresponds
to a Mw=6.5 event, with most of the seismic moment released
before early July 2018. The most active period, from 1 to 7 June
2018, corresponds to a Mw=6.2 event, more than the Mw=5.9
of the largest earthquake of May 15, whereas the largest magnitude
registered in this period is ML=5.3. In comparison, the cumulative
seismic moment from 1977 to 2017 for the entire Comoros (between
42.5◦E and 48◦E) represents less than Mw=6.0. Between 1 August
2018 and 15 May 2019, the cumulative seismic moment released
corresponds to a Mw=6.1 event. After a quiescent phase from 13
July to 26 August 2018, there was a resurgence of seismic activity
at the end of August with lower intensity and low contribution to
the moment release (Fig. 3).
In order to follow the temporal evolution of the seismicity, we
choose to keep S–Ptime at YTMZ as a reference, as it is the only pa-
rameter common to the entire seismic sequence that is independent
to location uncertainties or network evolution.
The seismicity is characterized by a long lasting sequence of
events localized in a reduced area. Even if the largest earthquake to
have been recorded in the Comoros belongs to this seismic sequence,
the seismicity cannot be described from a main shock followed by
aftershocks. S–Ptime and magnitude distributions (Fig. 3) highlight
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The Mayotte 2018–2019 seismo-volcanic crisis 29
Figure 5. From top to bottom panel: longitude, latitude and depth evolution with time within the first year of the Mayotte seismic sequence (ML≥3.5 from 10
May 2018 to 15 May 2019). Vertical lines mark principal steps in the seismic network improvements. Events marked by a light gray circle and transparency
were localized with less than 10 phases and considered as less relevant. Before mid-July 2018, depth is particularly badly constrained (consequentlyfixedat
30 km when not retrievable): hypocentral evolution must be considered accordingly.
the presence of different clusters. The first seismic cluster began on
10 May 2018 and is associated with an increase of S–Ptime until
1 July 2018. Two other clusters then appeared associated to distinct
values of S–Ptime. These three main clusters are coherent in time
(Fig. 3) and space, as they are restricted to limited area (Fig. 4).
In the first one, active mainly from May 10 to beginning of July
2018, the events (blue dots) occurred in a ∼20 km diameter area,
corresponding to S–Ptime at YTMZ from 5.2 to 8.1 seconds (i.e.
east of YTMZ, ∼30 to 50 km considering hypocentral distance).
This first cluster gathers all the largest events of the crisis. During
May 2018, most events were concentrated in an around 12 km
diameter area located 30 km eastward from the island after the
very first earthquakes occurred closest to the island (Figs 4and 5).
The whole period shows a gradual increase of S–Ptime at YTMZ,
especially from the beginning of June 2018 (Fig. 3). From 1 June
to 1 July 2018, S–Ptime at YTMZ increased from 5.8 to 8.1 s,
associated with a migration SE, with a total of 20 km shift for the
period (Fig. 5). Depths could be determined only for few events
during June 2018 (Fig. 5), but they are compatible with an upward
migration of the seismicity. SE migration is confirmed by G-CMT
locations for the largest events (Fig. 6) associated in parallel with
a depth change from 35 km to 12 km from early June 2018 (depth
from G-CMT catalogue is decreasing from June 4, Fig. 6), which
is beyond the error bars and confirms the hypothesis of upward
migration. Depth migration is also confirmed by Cesca et al. (2020)
from moment tensor inversions and analysis of teleseismic depth
phases. We distinguished two phases in this first cluster: cluster 1a
from May 10 to the beginning of June 2018, and cluster 1b covering
early June to beginning of July 2018, that started at the beginning
of the seismically very intense first week of June. Cluster 1b is
associated with the increase and spreading of S–Ptime reported at
YTMZ and preceded its lowering at initial values at the end of June
2018 (Fig. 3).
From the end of June 2018 (around June 26) to July 2018, S–
Ptime retrieves lower values equivalent to the beginning of the
first cluster (between 5.2 and 6.2 s); the second cluster (yellow dots)
starts westward from the latest events of the firstcluster (Fig. 4), that
is closer to YTMZ (hypocentral distance to YTMZ: ∼40–49 km).
This second cluster is close to the initial position of the first cluster,
before its migration (i.e. corresponding to cluster 1a). Clusters 1
and 2 overlap in time until the former almost vanishes around 8 July
2018. In July 2018, a third cluster starts (red dots), with S–Ptime
at YTMZ from 3.3 to 5.1 s, overlapping in time with cluster 2, but
more energetic (Fig. 3). Clusters 2 and 3 are spatially separated by an
aseismic gap around longitude 45.49◦(Figs 4and 5), and distinction
between these two families of events is sharper considering this
border than values of S–Ptime (Figs 3and 5). After mid July 2018,
there is no evidence of overall earthquake migration within clusters
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30 A. Lemoine et al.
Figure 6. Top panel: 27 G-CMT locations corresponding to the first cluster (4.7 ≥Mw≥5.9) added by one event belonging to cluster 3 (14 May 2019). As
colours depend on date, SE migration appears in parallel to the upward migration. The G-CMT focal mechanisms correspond to ML>4.7 events. (bottom)
Depth distribution of the seismicity as revealed by the depths determined by G-CMT for events listed in Table S3. Upward migration is observed from early
June 2018 (4 June 2018 is marked by the upper arrow).
2 and 3 (Figs 4and 5). However, 26 August 2018 marked the end
of the period of seismicity quiescence and Fig. 5illustrates the
associated aligned hypocentres distribution (likely along NW–SE
axis). This particular pattern marked the beginning of an intense
seismicity within the cluster 3. Inside clusters 2 and 3, depth ranged
mainly between 25 and 40 km with a sparse seismicity shallower,
except on 26 August 2018, when the aligned hypocentres seem to
be associated to depth ranging between 5 and 25 km.
The earthquakes of the first cluster seem to be shallower than
those of clusters 2 and 3 (Figs 4and 5); however, the depth distribu-
tion of earthquakes must be interpreted with caution, due to the lack
of an ad hoc velocity model and to the unfavourable geometry of the
network, especially between May and June 2018. As a whole, the
three clusters depict an activity migrating westward (Fig. 4) during
the onset of the crisis until July 2018 (whereas inside the first cluster,
a SE migration is observed, confirmed by G-CMT locations for the
largest events), after what a steady seismicity pattern is observed.
Westward from the Mayotte edifice, one ML>3.5 earthquakes
is reported in Fig. 4(green circle): on 6 April 2019 (ML=3.7,
with 12 phases). It occurred close to the epicentres of both the 1
December 1993 damaging earthquake (Lambert 1997,mb5.2) and
9 September 2011 event (Mw5, normal faulting, G-CMT).
The 32 largest events (ML≥5.0) are listed in Table S3 with their
G-CMT moment when it exists. All together they correspond to
45 per cent of the total moment release from 10 May 2018 to 15
May 2019. All occurred during the first part of the seismic sequence,
between 14 May and 27 June 2018, except three events (28 March
2019, ML=5.0; 5 April 2019, ML=5.2; and 14 May 2019, ML
=5.1). Except one reverse faulting mechanisms for 14 May 2019
event belonging to cluster 3, most of them have strike slip focal
mechanisms and belong to cluster 1. The G-CMT nodal planes are
very similar for these events with averages and mean scatters of
352 ±9◦,63±9◦and 5 ±13◦for plane 1 and 260 ±12◦,85±11◦
and 152 ±9◦for plane 2 (Fig. 6). The corresponding tension axis
is oriented N34 ±15◦, which corresponds to a most favoured fault-
ing azimuth N304◦. Both nodal planes indicate strike slip on a steep
fault, indicating, during May and June 2018, either right lateral shear
along ∼E–W structures or left lateral shear along ∼N–S structures
(Fig. 6). The former is more consistent with the relative movements
of the plates from either side of the archipelago, that is counter-
clockwise rotation of the Lwandle Plate with respect to the Somalia
Plate (Saria et al. 2014), but potential pre-existing active structures
that could be responsible for those mechanisms are still not iden-
tified (if we consider an influence of the regional tectonics on the
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The Mayotte 2018–2019 seismo-volcanic crisis 31
seismo-volcanic crisis). N–S structures would have the same orien-
tation as ancient transform faults linked to ancient spreading in the
Somali basin along the N–S direction (e.g. Rabinowitz et al. 1983).
In the surrounding areas, there are other strike-slip focal mecha-
nisms reported in the G-CMT catalogue since 1976 (Fig. 2), making
the whole consistent with a present-day trans-tensional regime in
the area and a ∼NE–SW regional extension. Moreover, strike-slip
focal mechanisms are commonly observed during dyke intrusions
(e.g. Toda et al. 2002). 14 May 2019 occurred within cluster 3,
at 34 km depth, its reverse focal mechanism is incoherent with
regional trans-tensional stress regime and cannot be explained by
dyke intrusion. As suggested by Cesca et al. (2020) who reported
more events of this type located in our Cluster 3, such thrust mech-
anisms should be linked to local processes, probably of magmatic
origin.
4 VERY LOW FREQUENCY
MONOCHROMATIC SEISMIC EVENTS
During the crisis, long lasting monochromatic very long period
(VLP) events were recorded by both local and international net-
works (Table S5 describes few of them). After the initial emer-
gent onset of the signal, they show decaying harmonic oscillations
comparable to Long-Period (LP) waveform patterns observed, for
example, at Kusatsu-Shirane or Galeras volcanoes (Kumagai &
Chouet 1999). However, they presented much larger period of os-
cillations (∼16 s) and long lasting signals. Systematic detection
and analysis of such events are reported by Poli et al. (2019), Sa-
triano et al. (2019)orCescaet al. (2020) who reported more than
400 of such VLP events and located them in the vicinity of the
cluster 3. The largest one, on 11 November 2018, well recorded
worldwide, awoke the interest of the seismological community and
the media (e.g. Wei-Haas 2018; Lacassin et al. 2020). This long
lasting monochromatic VLP event was not preceded by a strong
earthquake, but there are earthquakes embedded in the signal, dur-
ing the occurrence of the VLP event, especially at the beginning of
the signal (Figs 7and S2). As observed in Polynesia in 2011 and
2013 by Talandier et al. (2016), such VLP wave train are associated
with elliptical particle motion in the vertical plane, in the direction
of propagation, between the source and the receiver. It can be at-
tributed to long monochromatic very low frequency Rayleigh waves.
As shown in Fig. S3: elliptical particle motion is observed in the
vertical plane and can provide information in the back-azimuth, di-
rection of propagation of local station should converge in the source
area.
For the largest very low frequency event of 11 November 2018,
a first signal appeared at 9:27:27 UTC, followed by a moderate
earthquake (ML3.1 at 9:28:00), and then the VLP signal that lasts
more than twenty minutes emerged (Fig. 7), as shown by the E–W
raw record of the local MCHI seismometer from G30 (Fig. S2).
Others embedded earthquakes appeared in the VLP signal, two of
them could be localized within cluster 3, the one that marked the
beginning of the very low frequency signal (ML=3.1 at 9:28:00,
45.3990◦E, 12.7317◦S, depth 21 km, 12 phases) and the smaller
one that followed it (ML=3 at 9:29:33, 45.3683◦E, 12.9208◦S,
depth 12 km, 10 phases). It should be mentioned that depth for
such events should be considered with caution. A third, smaller
earthquake occurs at 9:30:50, followed by two very small ones that
we could not localize, but that are well visible in the raw trace shown
in Fig. S2(a). The spectrogram allows us to separate the long VLP
signal from standard volcano-tectonic earthquakes. The spectrum
of the signal (Fig. S2b) shows a peak at 0.062 Hz, thus a period of
∼16 s corresponds to the VLP event integrated in the signal (green
dot).
Fig. 7shows the three displacement components for the 11
November 2018 event at seven local and regional stations (Fig. 2),
including the YTMZ accelerometer and velocimeters corrected by
instrumental responses. After integration, the signal was filtered
with a band of 0.05 to 1 Hz. Three distinct phases can be distin-
guished in the records. At the event onset, the first low frequency
phase starts almost at the same time as the small high frequency sig-
nal that initiates the sequence (black dots, especially that of MCHI,
on Fig. 7). For the local stations, this initial phase lasts around
one minute. The second phase is the one with the oscillations and
highest amplitude monochromatic waves (green dots). In the decay
part of the harmonic oscillations, there are two embedded decay en-
velopes starting at the green and purple dots and lasting 20–30 min.
The duration of the monochromatic very low frequency wave train
seems not to depend on distance from the source: on the eastern
and vertical component of stations on Fig. 7very low frequency
oscillations from local stations (YTMZ and MCHI) has the same
duration than the ones from regional stations (CAB in Grande Co-
more, or SBV northern Madagascar). This observation suggests that
the long duration very low frequency signal is linked to a source
process rather than to a propagation effect. 11 November 2018 sig-
nal could be related to an oscillating source, in agreement with the
model of resonance of a deep magma reservoir proposed by Cesca
et al. (2020)
Long period seismicity is considered to be related to magmatic
and hydrothermal processes in a volcanic context. Their signal can
be linked to oscillations within an excited fluid filled oscillator (e.g.
Chouet 2003). Physical properties of the resonator can be deduced
from the signal characteristics. We modelled the decay envelopes of
the very low frequency event of 11 November 2018 (Fig. S4) with
a model of a damped oscillator following the equation exp (-t/τ)
with τ=360 s. From the physics perspective, this is compatible
with the excitation followed by the free damping of a viscous fluid
oscillating within a reservoir.
Figs 7and S3 show that, especially at local stations YTMZ and
MCHI, the VLP tremor affects mainly the vertical and E–W com-
ponents, and much less the N–S component, as indicated by particle
motion plotted from the decaying part of the signal in order to avoid
noise induced by embedded events. The particle motion calculated
for YTMZ and MCHI shows how the oscillation is polarized dur-
ing the monochromatic decay phase. Horizontal polarizations of
the closest stations are used in order to estimate the position of
the source: in Fig. S3, the green lines from YTMZ and MCHI on
the map are converging towards the aseismic gap between clusters
2 and 3, with the azimuth ∼N85◦from MCHI and ∼N100◦from
YTMZ. An origin of this long lasting very low frequency oscilla-
tion in the vicinity of cluster 3 would be consistent with the location
of the small located earthquakes present at the onset of the signal.
Those earthquakes could have excited a fluid filled resonator, yet
their locations need to be considered with caution.
5 GROUND DEFORMATION
5.1 GNSS data
Six GNSS stations (Table S6) were operating in Mayotte at the
beginning of the crisis: one (MAYG) installed in 2014, the five
others in early 2018. Unfortunately, there were no GNSS stations
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32 A. Lemoine et al.
Figure 7. Displacements at local and regional stations, in meters, for the 11 November 2018 very low frequency event. Black dots represent the initial
excitation, green and purple dots are end and the beginning of the two embedded monochromatic damping phases.
and data available from the surrounding islands of Anjouan, Moh´
eli,
Grande Comore, Glorieuse. Fig. 8shows the time-series at the
station MAYG. Only this station has long enough time-series to
estimate and remove a seasonal term (GNSS data processing is
described in the Supporting Information).
The time-series (Fig. S5) show a major component of common
mode. Four phases are visible, we call them: A, B, C and D. Phase
A spans from the origin of the GNSS observations to the main
earthquake of May 15 it shows the background deformation. No
precursory deformation is observed at any of the stations. Then,
from 15 May 2018, three distinct deformation anomaly phases are
observed link to the seismo-volcanic crisis. During phase B, from
15 May to 30 May 2018, there is a small yet well visible defor-
mation on the east component (Fig. S5). During phase C, from
May 30 to July 3 there is little horizontal deformation at the sta-
tions and a slight global subsidence. During phase D, starting on
July 3, all stations show a large and steady drift mostly east and
downwards.
Table S7 contains the velocity anomalies during phase D, ob-
tained by linear fit. The time-series of the N–S component of the
two southern stations, PORO and BDRL, show significant north-
wards velocity during phase D unlike the other stations. As the E–W
velocities at all stations are very similar, we averaged them and, at
the first order, we consider that the average vector represents the
motion of the entire island of Mayotte at the barycentre of the six
points located at 45.164◦E and 12.813◦S. Fig. 9shows how the pat-
tern of the cumulated seismic moment correlates with the average
eastern velocity of Mayotte.
Fig. S6 shows the evolution of the east versus north compo-
nent, and east versus vertical. This figure is dominated by the large
drift that occurs during phase D and shows clearly that this drift is
linear at the first order during this first investigated period (early
July–mid November 2018). Vertical versus east component plots
for phase B, C and D are aligned, showing that the subsidence
is not restricted to phase D, but occurs also, with lower intensity,
during phases B and C. For north versus east, the situation is dif-
ferent and phase B, C and D occupy spaces not aligned together
(Fig. S6).
Using the time-series of Fig. S5, we estimate the best-fitting
velocity anomalies, thus removing the long term International Ter-
restrial Reference Frame (ITRF14) velocity and the seasonal com-
ponent of MAYG—both assumed to be valid at all stations. The
velocities anomalies are summarized in Table S7.
We finally analysed the data until mid-January 2020, and we split
phase D in six periods of three months (D1 to D6) in order to assess
the spatio-temporal evolution of the deformation during the crisis.
Fig. S7 shows the time-series at four of the GNSS stations for the
periods D1–D4 (July 2018–July 2019) and Fig. 8MAYG for the six
periods. The deformation evolves with time and at a lowering rate
during phases D3, D4, D5 and D6 (January 2019–January 2020).
Table S8 reports this observation quantitatively.
5.2 Interferometric analysis
In order to determine a dense pattern of deformation for the entire
island of Mayotte we performed interferometric data analysis. This
analysis involved the processing of Copernicus Sentinel-1 mission
Synthetic Aperture Radar (SAR) data from both ascending and
descending acquisition geometries for the periods of December
2017 to April 2019 (29 scenes) and June 2017 to April 2019 (56
scenes), respectively (the processing is described in the Supporting
Information).
InSAR measurements well-outlined the ground displacement pat-
tern of Mayotte Island, showing a ramp subsiding east (Figs 10
and S8), as a part of the overall concentric ground deforma-
tion common to volcano-tectonic unrests. The observed maxi-
mum rate change from one side of the island to the other (ap-
proximately along an E–W direction) reached −115 and −63 mm
yr−1for the ascending and descending tracks, respectively, which
are perfectly consistent with the relative motion measured by
the GNSS network. Differences of line of sight (LoS) measure-
ments from opposite observation points indicate the presence of
significant horizontal motion, as depicted by GNSS vectors. In
that case, subsidence combined with motion away from the satel-
lite for the ascending orbit results in higher displacement rates
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The Mayotte 2018–2019 seismo-volcanic crisis 33
1 Jan. 2014 1 Jan. 2015 1 Jan. 2016 1 Jan. 2017 1 Jan. 2018 1 Jan. 2019 1 Jan. 2020
Date
-100
0
100
200
300
400
500
600
M illim etres
Ve-ITRF = 21.9 mm yr -1 Amp -yearly = 0.6 mm
Vn-ITRF = 14.8 mm yr -1 Amp -yearly = 0.9 mm
Vu-ITRF = 0.8 mm yr -1 Amp -yearly = 2.6 mm
D1 D2 D3 D4
GNSS station M AYG
East
North
Vertical
D5 D6
Phase B
Phase C Phase D
Phase A
Figure 8. Full time-series of station MAYG, including the evaluation of the seasonal signal that must be removed to derive the real ground deformation
produced by the eruption.
Figure 9. The phases of the crisis and the cor responding evolution of the cumulated seismic moment and average E–W displacement of Mayotte. The scale
on the left is for both the deformation (mm) and the seismic moment (×1016).
compared to the descending one. The compatibility between InSAR
and GNSS and agreement to the proposed phases of the crisis can
also be seen by examining the InSAR LoS displacement time-series
(Fig. 10).
It should be noted that as the entire island is deforming, In-
SAR can provide only information regarding the relative motions
within Mayotte, while the knowledge of the absolute motions is
coming from the GNSS. Using the LoS changes estimated from
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34 A. Lemoine et al.
Figure 10. Sentinel-1 line of sight (LoS) displacement rate maps from (a) ascending and (b) descending acquisition geometries over the common period of
observation. Selected local reference is marked as a black rectangle. Vertical blue bars delineate deformation phases A, B, C and D. Displacement history of
MAYG GNSS station projected into the InSAR descending LoS, after compensation for regional motion trends and adjustment to the InSAR local reference.
the GNSS time-series at nearby pixels allow us to overcome this
limitation and thus derive absolute LoS time-series at all InSAR
pixels.
Of high interest, and not achievable with GNSS, is the obser-
vation made by InSAR of the homogeneity of the displacement
field within the whole island. This uniformity is an expected con-
sequence of the distance of the source that we calculate in the next
section. There is no evidence of landslide or any other land motion
that could have been triggered by the repeated moderate seismic
events during the crisis and could constitute a potential threat for the
inhabitants.
6 GROUND DEFORMATION
MODELLING
6.1 Evidence of three phases of deformation
Based on the characteristics of the seismicity and deformation,
we split the crisis in three unrest phases (B, C, D) preceded by
the quiescent pre-crisis phase A. The three phases, issued from
deformation anomalous phases, are interpreted as follows: phase
B, from 15 May to 30 May 2018: seismic crisis and fracturing
between a deep reservoir and the surface; phase C, from 30 May
to 3 July 2018: magma ascent from the reservoir to the surface and
SE migration; phase D, starting on 3 July 2018 and still ongoing in
early 2020: eruption.
In the following, we first model phase D and then move backward
in time to C and B. Indeed we need first to decipher the dominant
phase D accurately as we will use its implications for the correct
interpretation of the initial phases C and B.
6.2 Model of the phase D (eruption)
For modelling the ground deformation observed during phase D,
we use the classical Mogi (1958) model that assumes a deflating
point source buried in an elastic half space. Our data well constrain
the source azimuth, pitch, and deflation, but not its distance to the
stations. This is because all GNSS stations are in the same azimuth.
Consequently the GNSS displacements are fit almost equally well
with reservoirs in the range of longitude 45.55◦E–45.75◦E. As the
ratio vertical/horizontal displacement indicates exactly the direction
of the source centre, the modelled depth varies from 25 to 40 km
when browsing the range of possible longitudes. On the other hand,
if we do not use the vertical information but only the horizontal
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The Mayotte 2018–2019 seismo-volcanic crisis 35
one, the preferred location of the source is at the average crossing
place of the horizontal vectors at a longitude at around 45.50◦E
thus slightly outside of the range permitted by the 3-D modelling.
If now we blend in our inversion 50 per cent of cost of fit of the
Mogi model with 50 per cent of cost based on the crossing place of
the horizontal vectors (Fig. 11), using as metrics the minimization
of the cross product of the displacement vectors and the vectors
source–station, we find a best-fitting source at 45.590◦E, 12.777◦S
and 28 km depth (depth of the deflation source is discussed in the
Supporting Information). The blending of the two criteria in the
inversion is interesting because the unconstrained Mogi modelling
tends to pull the model further east (to give a better fit at one single
observation, the E–W motion of MAYG, which means that these
data have to be highly trusted), while the crossing of the horizontal
vectors pulls the solution the other way. The combination of both,
with similar weight acts as a joint inversion and gives a stable
solution with mean residuals of 25, 5, and 20 mm yr−1in the east,
north, and vertical components, respectively (Table S7). With that
geometry we infer an almost steady emission rate of 82 m3s−1
during the first months of the phase D.
6.3 Model of the phase C (magma ascent)
During the phase C, we interpret the deformation (presented in Table
S9, left) as due to the ascent of magma in a conduit connecting a
deep reservoir to the subsurface, combined with the deflation of
this reservoir of a similar volume. This scenario is supported by
the geodetic observations as, during that phase, the E–W motion
is nearly zero while there is a slight vertical motion (Fig. S5) of
7±4 mm. Such ground motion can be explained with a magma
body migrating upward without emission at the surface, thus without
a global volume change. As the above-mentioned subsidence rate
is 55 per cent of that of phase D and mostly attributable to the
deflation of the reservoir (because a shallow source cannot produce
vertical changes at GNSS stations at that distance), we assume that
the horizontal rate associated with the deflation is also 55 per cent
of that of phase D. Consequently, this allows us to evaluate what we
call the C reservoir contribution in Table S9. The remaining part
is what we call C conduit. We interpret it as the elastic response
measured at the GNSS stations to the opening of the feeding magma
conduit (Fig. 11) from the top of the chamber to the subsurface and
the further opening of shallow dykes and sills. As the emission
rate is 82 m3s−1in phase D, rescaled by 55 per cent, it is 45 m3
s−1during phase C, which means that the volume injected in the
channel, shallow dykes, and sills is 0.16 109m3during phases B
and C.
We model the C conduit deformation by using the Okada (1992)
formalism and the inversion code developed by Briole et al. (1986)
and Briole (2017). As the GNSS stations are far from the source,
we cannot resolve the shape of the conduit. We assume that it is an
elongated dyke buried at a depth of 22 km, at the top of a deflating
chamber that we assume is 12 km in diameter (consistent with the
extend of cluster 1a and with dimension of numerous well-known
calderas, such as the ones reported by Geyer & Mart´
ı(2008), e.g.
Campi Flegrei in Italy) with its tip 2 km beneath the seafloor. As a
priori azimuth for the conduit, we use N304◦(Section 3.3), which
is perpendicular to the tension axis. The best-fitting azimuth found
by our inversion is N318◦, though loosely constrained because of
the distance of the GNSS stations. The best-fitting solution for the
width of the conduit is 4 ±2 km, and its opening is 1.62 ±0.7 m.
Thus, the best-fitting volume of the modelled conduit is
0.13 ×109m3, 20 per cent less than the apriorivalue. The root
mean square (rms) scatter of the residuals is 3 mm for the horizontal
components and 6 mm for the vertical. In the uppermost part of the
crust, the conduit could extend laterally in a shallow dyke ∼20 km
long, along the same axis ∼N318◦. This dyke could accommodate
the missing 0.03 ×109m3. The less active seismic cluster 2 might
correspond to earthquakes induced by redistribution of the stress in
the rock surrounding the magma reservoir deflation, as it is located
close to its centre (e.g. Gargani et al. 2006).
There is no means with the geodetic data to resolve precisely the
characteristics of the shallow channels (dyke and or sill), because
their volume is small for the distance and they do not affect the
vertical component. From the vertical GNSS data, we can only
assess that there cannot be a large horizontal injection at depth (e.g.
10 km long at 10 km depth) because this would give a geodetic
signature that is not observed.
6.4 Model of the phase B (fracturing of the crust)
For phase B, during which a small yet clear deformation is observed
from 15 to 30 May 2018, we used an Okada (1992) model with the
following parameters: same location and geometry as phase C, a
rectangular fault elongated vertically of 20 km length and 4 km
width (Fig. 11, Table S10). The best-fitting slip is 0.85 m of pure
strike-slip. Assuming a rigidity of 3.3 1010 Pa, the corresponding
moment is M =224 1016 Nm, a value twice as large as the total
seismic moment cumulated on May 30, but compatible with the
cumulated moment a few days later on June 2. When relaxing the
azimuth from its initial value of N318◦(inherited from phase C
modelling), the best fit is obtained with N341◦, which corresponds
to an azimuth with the tension axis oriented at N26◦, slightly smaller
than those found before, but consistent with the G-CMT focal mech-
anisms. With both geodesy and seismology, we have no means to
discriminate between the conjugate fault planes, and the upward
fracturing during phase B might well involve ruptures that blend the
two families of planes.
6.5 Temporal evolution of the deformation during the
eruption (phase D)
In the section 6.2, we model the ground deformation by assuming,
during phase D, a single source of deformation and assuming a
constant rate of deformation during period D from 3 July until
15 November 2018. Below, we use the available data until early
January 2020 (Fig. 8for MAYG), and we split in six periods of
three months (D1–D6), the evolution of the deformation during the
eighteen months of the deflation signal. The modelling, made by
assuming the same location (horizontal and depth) of the deflating
reservoir, shows a clear evolution of the effusive rate during the
eruption. The rate increases during the first phase of the eruption to
reach a maximum around the end of 2018. This is the period during
which the largest very low frequency tremor were observed on 11
November 2018, and during which clusters 2 and 3 reach steady
activity (Figs 3and 4). Then, the deformation gradually decreases
during the first semester of 2019 (blue line on Fig. 12). Table S11
gives the effusive rate values for the six periods. According to this
model, the cumulated volume extracted from the deep reservoir
has reached ∼2.65 km3in January 2020, 1 yr and a half after the
start of the eruption. This is less than the value reported from the
estimated volume of the new submarine volcanic cone discovered
at the sea floor (∼5km
3,REVOSIMA2019; Feuillet et al. in rev.)
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36 A. Lemoine et al.
Figure 11. Yellow and green vectors are proportional to the vertical (bottom panel) and horizontal (top panel) observed and modelled displacement anomalies
during phase B (fracturing from 15 to 30 May 2018); phase C (magma ascent from 30 May to 3 July 2018) and phase D (intense deflation mapped from 3
July 2018 to 14 November 2018). Only a part of the intense deflation phase D was represented in order to keep the same scale for displacement with phases
B and C. Clusters 1, 2 and 3 with ML>3.5 events for each period are plotted in blue, yellow, and red colour scales. Locations of events with a clear outline
and transparency should be less consistent (less than 10 manually picked phases). On the vertical cross section, only resolved depths are plotted (otherwise,
the depth is fixed at 30 km). Yellow triangles indicate the GNSS stations. Light orange circle corresponds to the location of the magmatic chamber considered
in this model. Yellow surfaces (bottom panel) represent the proposed geometry of the conduit connecting the deep reservoir to the subsurface and potential
laterally extended shallow dyke zone.
Single sour ce (rate in k m3 /y r)
r.m.s. f it with single sour ce (cm)
Dual source (main reservoir)
Dual source (seconda ry reser voir)
r.m.s. f it with dual sourc e (cm)
2 Jul. 1 Oct. 31 Dec. 1 Apr. 1 Jul.
Date (2018-2019)
1 Oct. 31 Dec.
0.50
1.00
1.50
2.00
2.50
3.00
3.50
Effusion rate (km/yr)
3
0.00
Figure 12. Temporal evolution of the effusion rate for single and dual source simple models.
but consistent if one considers a 40 per cent increase of volume
between magma at depth and the products at the surface.
7 DISCUSSION
7.1 Fracturing of the crust and magma injection
The intense seismic sequence begins offshore Mayotte with a first
cluster, active from 10 May 2018 to early July 2018. During May
2018 (cluster 1a), intense seismicity (11 ML≥5.0 events) is con-
centrated in a limited area with a diameter around 12 km. The first
week of June represents the most intense week for the seismic se-
quence (equivalent to Mw=6.2), and the beginning of seismicity
migration towards the surface and SE until the end of June 2018
(forming cluster 1b).
The GNSS data provide information on the timing of the creation
of the feeding channel that opened the way for the magma to migrate
from the reservoir to the surface.
This occurs during phases of the deformation corresponding to
fracturing linked to strike-slip faulting, and then to magma ascent,
between 15 May and 3 July 2018. Phase B, associated with mod-
erate deformation from 15 May to 30 May 2018, is interpreted as
the fracturing phase, and linked to the intense cluster 1a. Moreover,
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The Mayotte 2018–2019 seismo-volcanic crisis 37
the change in the deformation signature suggests that the filling
of the conduit started at the end of phase B, during the peak seis-
micity (and rock fracturing) that occurs between 30 May and 7
June, 2018. During the 2014 B´
arðarbunga dyke intrusion, most of
the released seismic moment occurred during the rapid propagation
phases ( ´
Ag´
ustsd´
ottir et al. 2016). At the same time, cluster 1b shows
intense seismicity migrating upward and SE (Figs 5and 6)thatat-
tests to upward fracturing, while the conduit started being filled
by magma until early July 2018. The fracturing of phase B can
be explained geodetically without the need of filling the ruptured
faults, that is it can be modelled thanks to coseismic displacements
associated with strike-slip faulting.
We have no strong constraint on the geometry of the feeding
conduit. The best-fitting model is an elongated vertical conduit
starting arbitrarily at the depth of 22 km on top of the reservoir and
ending at a depth of 2 km, yet the latter value is loosely estimated by
the inversion. What is well constrained by the inversion, and almost
independent on the shape of the source, is the volume of the conduit,
0.13 ±0.01 109m3, and therefore we can infer its section, 6500 m2.
Assuming a pipe shape, its diameter would be 90 m. It is difficult
to figure out how such a cylinder could appear in a few days within
10–20 km of oceanic crust not having been perforated by magma
in recent geological time. Therefore, it is much more likely that the
feeding conduit is a dyke. The best fit for this dyke is a horizontal
width of 4 km, an opening of 1.62 m, and an azimuth of N318◦.
The azimuth is loosely constrained because the conduit is narrow,
but consistent with the overall SW–NE extension inferred from the
focal mechanisms and the orientation of the extension from GNSS
regional observations (Fig. 1).
The steadiness of the deformation in June 2018 implies a steady
rate (within our observational uncertainties) of the amount of
magma released from the reservoir during that month. This sug-
gests that the cracks moved towards the surface and started being
filled according to a relatively smooth process.
The G-CMT focal mechanisms are associated with the largest
events of May and June 2018. Those events belonging to cluster 1,
during which there is an apparent ∼20 km SE epicentral migration
(Fig. 5). This migration was confirmed by location of the largest
events by G-CMT (Fig. 6). However, there is evidence of upward
migration: the event depths determined by G-CMT show a clear up-
ward migration from early to late June 2018 (Fig. 6) that is coherent
with depth evolution with time from our location, even if during
this period, seismic network sparsity do not permit accurate depth
determination for the whole earthquakes. During the 2011–2012 El
Hierro unrest, a migration extending over 20 km and ending in an
offshore eruption was precisely located (L´
opez et al. 2012;Mart
´
ı
et al. 2013). The B´
arðarbunga-Holuhraun dyke intrusion (Iceland)
in 2014, is associated to magma propagation along 48 km during
2weeks(
´
Ag´
ustsd´
ottir et al.2016,2019). White and McCausland
(2016) compiled data from seismic swarms linked to eruptions and
intrusions. They noted that, generally, seismicity began at a distance
that can reach tens of km from the eruptive location.
Focal mechanisms (from G-CMT) show the dominance of strike
slip events during the fracturing period. In Fig. 6, the strike-slip
focal mechanisms represent earthquakes belonging to the period
from 14 May to 27 June 2018, that is corresponding to deformation
phases B (fracturing) and C (magma ascent). Dyke related seis-
micity also sometimes produces the so called ‘dogbone seismicity
distribution’, where normal faulting events around the dyke (nor-
mal faults parallel to dyke) coexist with strike slip events around it
(e.g. Toda et al. 2002). In the case of the Mayotte seismo-volcanic
phenomenon, only the largest events are associated with a focal
mechanism, leading to a lack of accuracy in the description of the
seismicity pattern.
We interpret cluster 1 as corresponding to fracturing of the crust
followed by triggering by magma ascent. The seismicity for the
first period of cluster 1 (cluster 1a during fracturing phase B) is
roughly above the location of the reservoir (Fig. 11), the inten-
sity of the seismicity during this period should be linked to the
resistance of the crust probably not associated to a significant frac-
ture density (e.g. Rivalta & Dahm 2004). Then, the upward and
SE migration is observed from early June 2018 (cluster 1b, earth-
quakes triggered by magma ascent during phase C). Magma in-
trusions phenomena are associated with seismic swarms, as ob-
served for example in 2000 in Izu islands (Japan) were an energetic
seismic swarm was accompanied by dyke intrusion and eruptions
of Myakejima (Toda et al. 2002; Rivalta & Dahm 2004). As re-
ported during the B´
arðarbunga-Holuhraun (Iceland) dyke intru-
sion in 2014 by ´
Ag´
ustsd´
ottir et al. (2016), the seismicity located
close to the front of a dyke is mainly characterized by strike-slip
faulting, as it was observed in Mayotte in May and June 2018.
The submarine eruptive observations highlighted in May 2019 dur-
ing MAYOBS campaigns (doi:10.18142/291; 10.17600/18 001 222;
10.17600/18 001 230; 10.17600/18 001 238) support the proposed
interpretation regarding cluster 1: during June 2018, beginning with
a very intense seismic phase, upward and SE migration was ob-
served until the end of June before the eruption began in early July
2018 (Fig. 13).
7.2 Eruption onset and the intense ground deformation
The seismic and geodetic data suggest that the proper eruption starts
associated with a second seismic cluster, much less active than the
first one, which appears at the end of June 2018. This cluster is
centred NW in comparison with preceding events. It appears during
a short intense period of seismicity (end of June 2018), marking the
last stage of migration. Then, in July and August 2018, there is a
period of relative seismic quiescence. Some days after the begin-
ning of this period, a third seismic cluster appears, located to the
west of the second one. Simultaneously, early July 2018 is charac-
terized by the onset of an intense deformation recorded by GNSS,
with large eastward and downward drifts of the entire island of
Mayotte.
Those two clusters were still active in early 2020, with the seis-
micity of the third one remaining more intense. This third cluster is
closest to Mayotte. Clusters 2 and 3 present b-values significantly
higher than the one of cluster 1 (Fig. S1), which is consistent with
values observed at some volcanoes (e.g. Wyss et al. 1997).
Because of the poor azimuth coverage of the GNSS stations
(Fig. 11), the longitude of the deflation source is not easy to assess.
However, as our two criteria (best-fitting Mogi source and hori-
zontal crossing of the vectors) are leading in opposite directions,
we believe that our localization of the deflating source is resolved
within a few kilometres in a robust way. Subsequently, its depth
(28 km) is constrained accurately by the pitch of the deformation
and the trade-off between distance and depth in the model (Fig.
S6). Its azimuth is well constrained by the azimuths of the GNSS
vectors. Moreover, this reservoir is localized in an area were level
of seismicity is very low whereas it is surrounded by earthquakes
(Figs 11 and 13). A similarly deep magmatic reservoir can be found
in El Hierro (∼25 km, Mart´
ıet al. 2013;Kl
¨
ugel et al. 2015)orin
the Klyuchevskoy volcano group in Kamchatka (z∼30 km, Levin
et al. 2014; Shapiro et al. 2017), linked to shallower sources.
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38 A. Lemoine et al.
Figure 13. Schematic summary. Yellow full filled circle represents the main deep reservoir inferred by the GNSS data. Best located ML≥3.5 events of clusters
1, 2 and 3 are plotted in blue, yellow and red colour scales. The eruption takes place on the yellow triangle, as observed during marine surveys (Deplus et al.
2019; Feuillet et al. 2019, in review). It is connected to the vertical conduit by a ∼20-km-long dyke. The model includes a main deep reservoir and a secondary
source beneath the south end of Mayotte island. A moderate flux of material (∼3m
3s−1) is inferred from that source to the main reservoir. Large dashed line:
proposition for Moho position (e.g. Jacques et al. 2019 from Dofal et al. 2018 and Coffin et al. 1986).
Our modelling assumes that the source is a point, which is of
course not the case, but we cannot resolve the geometry of the
source with GNSS network as, at such distance, a source with a
radius of several kilometres is indistinguishable from a point source.
We assumed that the source diameter is 12 km, because this is a
typical value used in literature for deep reservoirs, based on the
size of many calderas that can be seen on Earth [e.g. the Campi
Flegrei caldera in Italy as described in the database from Geyer
&Mart
´
ı(2008)]. It also ranges with the dimension of the area
where earthquakes of cluster 1a were concentrated (interpreted as
the fracturing phase preceding any magma ascent, Fig. 11).
7.3 Speed of magma ascent during the eruption
With 6500 m2for the best-fitting section of the feeding channel and
82 m3s−1for the effusive rate, if we assume a constant emission rate
between 3 July and 15 November 2018, the speed of magma ascent is
0.013 m s−1. The speed is half this rate between 30 May and 30 June
2018, when migration of seismicity revealed the magma migration
(during June 2018). Such a low speed of magma ascent has been
observed elsewhere and is discussed by Gonnermann & Manga
(2013) through the analysis of the crystals in the erupted magmas.
This evaluation of the speed could be further compared and tested
with the one derived from the analysis of the crystal growth in
the fresh rocks sampled on the new volcano (e.g. Bach`
elery et al.
2019).
7.4 Rates and emitted volume
The rate of magma released is less in June, approximately half,
than the one estimated after 3 July 2018. In other words, the flux
in the feeding channel doubled in July once the path to the sur-
face opened, and the eruption started. The timing of the start of
the eruption, estimated on 3 July 2018, comes from the beginning
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The Mayotte 2018–2019 seismo-volcanic crisis 39
of the large ground deformations, thus, when there is no longer a
balance between the volumes collected at large depth in the reser-
voir and those stored at shallow depth. Early July corresponds also
to the beginning of a seismicity quiescence period, and the on-
set of a new pattern of seismicity: only few small shallow events
around the dyke and the onset of a new and steady seismicity pattern
within clusters 2 and 3. In the same way, during the B´
arðarbunga–
Holuhraun rifting event (Iceland) in 2014, a rapid decrease of the
seismicity followed the beginning of the main eruption and a lack
of shallow seismicity is observed ( ´
Ag´
ustsd´
ottir et al.2016,2019),
like in other contexts (e.g. during the 1983 Kilauea dyke intrusion,
Rubin et al. 1998). Persistent seismicity is not associated to the
propagation of a dyke (as during phase of magma ascent in May-
otte), but to processes of continual fracturing that maintains magma
flow (e.g. ´
Ag´
ustsd´
ottir et al. 2016). Moreover, Ruch et al. (2016)
highlighted importance of pre-existing fractures in rifting events as
they can control the magma propagation and new distribution of
stresses.
With an effusion rate of 45 m3s−1over 42 d (magma ascent
phase C) and then, after the estimated beginning of the eruption
early July 2018, six phases over 3 months with rates 72.9, 92.4,
77.2, 40.7, 27.0 and 12.2 m3s−1, the total volume extracted from
the chamber is 2.65 km3, according to our modelling. White &
McCausland (2016) proposed an empirical tool for estimating the
intruded magma volume from the analysis of a large range of seis-
micity patterns that preceded past eruptions or intrusions. Consid-
ering the cumulative seismic moment between 10 May 2018 and
15 May 2019, the estimated magma volume should be of the order
of 1.55 km3, lower than what we predict at this date from GNSS
data (∼2.21 km3). Moreover, recent marine surveys (R/V Marion
Dufresne) performed bathymetric and water column observations
that confirm recent submarine volcanic activity, 50 km eastward
from Mayotte and differential bathymetry allowed to estimate an
important volume emitted at the seafloor of ∼5km
3(REVOSIMA
2019; Feuillet et al. in rev.). The volume of released magma from
the reservoir that we estimated without aprioriassumption is lower
than the volume deduced from direct seafloor observations of the
eruption. This discrepancy can have different origins, including po-
tential volume change of material between its extraction at 28 km
depth and the seafloor (Bach`
elery et al. 2019).
Such volume ranks the Mayotte eruption among the major events
of the last 1000 yr. It can be compared only to other very large
eruptions such as the Lanzarote one between 1730 and 1736, that
emitted 3–5 km3, associated with an effusion rate five times lower
than the one estimated for Mayotte with comparable volumes (Car-
racedo et al. 1992). The Laki eruption spanned from June 1783 to
February 1784 and emitted 14.7 km3of magma, with an effusive
rate nearly four times higher than the one from Mayotte (Thordarson
&Self1993).
7.5 11 November 2018 very low frequency tremor
Only two events reported in 2011 and 2013, close to Rocard sub-
marine volcano, Society hot-spot, Polynesia (Talandier et al. 2016)
could be compared to Mayotte unusual 11 November 2018 signal:
both 2011 and 2013 intense Rayleigh wave trains were monochro-
matic (the period observed was 17.0 s) and long-lasting (50–60 min
in 2011 and 30–40 min in 2013). Talandier et al. (2016) interpreted
these events as due to the strong hydrostatic pressure undergone by
the volcanic edifice at 3200–4000 m depth below the sea surface.
Under high hydrostatic pressure, magma movement can generate
the resonance of a shallow opened conduit considered as a fluid
filled reservoir after punctual excitation.
If those unusual events belongs to VLP seismicity (regarding
frequency range), their waveforms look like of some LP events as-
sociated with the decaying oscillation in a resonator. Among them,
some earthquakes are described by decaying harmonic oscillations
after a short onset in the period range 0.2–2 s, characterized by
one or more frequency peaks related to a source effect (e.g. Chouet
1996; Chouet & Matoza 2013). This pattern is interpreted as a time-
localized pressure excitation mechanism followed by resonance of a
fluid filled cavity in response to the excitation. Damped oscillations
can be described by its dominant frequency and quality factor that
depend on the source properties: fluid types and resonator geome-
try (e.g. Kumagai & Chouet 2000). Characteristic frequencies seem
to be different from one volcano to another. Signatures of signals
recorded at Kusatsu-Shirane or Galeras volcanoes present a typical
signature of long harmonic codas following a brief onset (Kumagai
and Chouet 1999) that exhibit some similarities with the Mayotte 11
November 2018 event, but with a lower range for the characteristic
period and duration. Among the family of VLP events (presenting
longer periods), signals observed at Hachijo volcano in 2002 present
a decaying harmonic oscillation lasting 300 s with period ∼10 s, a
pattern that is similar to features of Long Period (LP) events with
shorter period ∼1 s. These VLP signals at Hachijo volcano were
interpreted as the resonance of a basaltic magma in a dyke (Ku-
magai et al. 2003). Kumagai (2006) proposed that the variation of
frequency and quality factor of the harmonic oscillation are due to
the gradual expansion of a crack containing basalt mixed with gas.
Kawakatsu et al. (2000) reported that Long Period Tremors (LPT)
with a dominant period around 15 s were observed at Aso volcano
(Japan), with a short duration (less than one minute), both during pe-
riods of volcanic quiescence and unrest. These LPT are interpreted
as a response of resonating fluid filled cracks to pressurization within
the hydrothermal system at Aso volcano (e.g. Hendriyana & Tsuji
2019).
The 11 November 2018 signal emitted offshore Mayotte (and
similar reported VLP events) is associated with a frequency peak
included in the range of VLP seismology, but its waveforms are
typical of some long period events attributed to a decaying oscilla-
tion in a resonator. This event can be considered as an end-member
of sources linked to resonator oscillation after excitation, similar to
the two events reported in 2011 and 2013 in Polynesia (Talandier
et al. 2016).
During the Mayotte seismo-volcanic crisis, the source process of
long-lasting VLP events is non-destructive as several hundreds of
them were reported since the beginning of the unrest and even few
months before (Poli et al. 2019; Satriano et al. 2019; Cesca et al.
2020). The characteristic period of such events increased until late
2018 before decreasing (e.g. Satriano et al. 2019; Cesca et al. 2020).
This period of highest period of resonance corresponds to the high-
est modelled effusion rate (Fig. 12). In parallel, volcano-tectonic
seismicity pattern follows a steady pattern for several months from
this period: activity is mainly concentrated in cluster 3 and cluster
2 is still active (Fig. 3). Geometry of the volume of the resonator
and characteristics of fluid filling should influence VLP signals.
The constraints on the geometry of the deep feeding channel, and a
possible shallow dyke, can be used as apriorigeometric model for
the interpretation and modelling of the reported very low frequency
tremors. Those could be generated by an oscillation, potentially
in a long horizontal shallow dyke or in the magma reservoir, but
not in the vertical conduit as the particle motion does not support
the hypothesis of a vertical vibrating source. Cesca et al. (2020)
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40 A. Lemoine et al.
linked the long-lasting VLP events to a 15 km long subhorizontal
crack forming a deep reservoir. However, constraining the source of
11 November 2018 event does not belong to the objectives of this
work, geodetic network do not allow us to constrain the geometry
and the complexity of the plumbing system: future analysis of such
unusual events and onshore and offshore acquisitions could help us
understand it.
Other tremors are recorded before and after this major tremor,
weaker but with very similar characteristics (Poli et al. 2019;Sa-
triano et al. 2019;Cescaet al. 2020), suggesting a common source
and a non-destructive process.
7.6 Spatial distribution of the source of deformation
The GNSS time-series (Fig. 8) show a change of deformation rate
and an increase of the pitch (ratio N–S/E–W velocities) of the
deformation with time. To take into account this observation, we
refine our model with a dual source. The small secondary source
can be regarded as a perturbation of the main one. The localization
and evolution of this secondary source is expected to provide clues
on the dynamic of the process at depth.
For each of the subphases D1–D6 we now invert for the main
source plus the small secondary source. Whatever are the initial
conditions set for the secondary source, the inversions bring this
source towards a position which is located beneath the south of
the Mayotte island around coordinates 45.18◦E and 12.94◦S(Fig.
13). Only the depth of the source cannot be constrained by our
model, so we decided to fix it at the same depth as the main
source.
Table S12 gives the values estimated for this dual source model.
With respect to the single source model, the overall volume of
magma released is unchanged with most of the magma (100–94
per cent) still extracted from the main reservoir and the secondary
reservoir providing a small amount of the total (0.1 per cent in phase
D1, 4.1 per cent in phase D2, 4.5 per cent in phase D3, 5.8 per cent
in phase D4, 6.6 per cent in phase D5 and 14.5 per cent in phase D6).
At the same time, this small perturbation of the model improves the
fit to the data by 23, 9, 23, 22 and 14 per cent, respectively, during
phases D2–D6, which is remarkable. In terms of emitted volumes,
the relative importance of the secondary source is increasing with
time.
The modelled secondary source might correspond, at the first
order, to a flux of material extracted from beneath Mayotte where
the main asthenospheric plume that has built the island should be
located. This flux might contribute dynamically to maintain the
isostatic equilibrium of the whole system and in that case its time
span will be controlled by the average viscosity of the deep rocks
involved in the process. This relaxation might continue well after
the end of the eruption that could occur by the mid of 2020 based
on the trend shown by Fig. 12.
Other authors have studied the connexion between reservoirs at
depth. For instance, Bato et al. (2018) proposed deep connections
between the two volcanic systems of Gr´
ımsv¨
otn and B´
arðarbunga
(Iceland), ∼25 km away explained by lateral flow hypothesis at
depth or shared magma reservoir at depth ∼30 km. From the 2011
to 2012 submarine El Hierro eruption analysis, Kl¨
ugel et al. (2015)
suggested that lateral magma movements at depth are typical of
mature oceanic intraplate volcanoes. They proposed that between
2011 and 2014, pre-, syn- and post-eruptive seismic swarms testify
to the presence of magma paths in the uppermost mantle and lower
crust.
7.7 Interpretation of the seismic clusters
Phase D of intense deflation, which has run for more than 1 yr at
the time of this paper, is the phase with the lowest seismic rate.
This indicates that the fracturing from the reservoir depth to the
subsurface is terminated. The second cluster, which spans from the
end of June 2018 until now, is located in the vicinity of the defla-
tion centre, likely bellow it, as far as depth distribution is reliable
(Fig. 11). During the El Hierro eruption (2011–2012), Mart´
ıet al.
(2013) interpreted the syn-eruptive seismicity as a response to read-
justments of the plumbing system following magma withdrawal. As
the plumbing system in the case of Mayotte is not described, it is
difficult to propose a reliable explanation for the sources of clus-
ters 2 and 3, both of which are syn-eruptive. A new swarm or
geometry in the seismicity pattern can mark changes in magma cir-
culation (e.g. segmented dyke growth in 2014 in the B´
arðarbunga
volcanic system; Sigmundsson et al. 2015). In the present case, it
seems that there is a seismicity gap between clusters 2 and 3, ex-
cept maybe at larger depths (Figs 4,5and 11), however, they are
both syn-eruptive and almost steady during months. Cluster 2, due
to its vicinity with the centre of deflation, should be linked to the
redistribution of stress around the reservoir due to magma with-
drawal. Post-eruptive clusters could also reflect regions of magma
circulation. Moreover, the third cluster could be linked to fluid
movement maybe between two reservoirs (Fig. 13). Cesca et al.
(2020) interpreted earthquakes observed offshore Mayotte the fol-
lowed the onset of the eruption as due to the progressive failure of
the roof of a magma reservoir (then inducing its resonance and VLP
events).
The three seismic clusters show an overall migration of activity
SE and then NW and W, indicating a coaxial complex and extended
activity ranging up to ∼40 km depth. All focal mechanisms pre-
ceding the eruption (May–June 2018) made by G-CMT indicate an
extension axis oriented ∼N50◦. The perpendicular azimuth N310◦
is therefore the most favourable axis for a feeding dyke. The cen-
tre of deflation, is ∼20 km from the new volcanic edifice (Deplus
et al. 2019; Feuillet et al. in rev.) in the azimuth N316◦(Fig. 13),
and thus, close to the abovementioned N310◦and to the azimuth
of the feeding conduit found by inversion of GNSS data which
is N318◦. This azimuth is also consistent with some topographic
features observed in the bathymetry east of Mayotte, for example
the Jumelles seamounts (NE from Mayotte, Fig. 1), and with that
of dykes exposed onshore north of Mayotte (Nehlig et al. 2013).
It is also consistent with the overall axis of the Comoros volcanic
Archipelago and is roughly perpendicular to the relative motion
measured by GNSS (Fig. 1) in azimuth N50◦across the Comoros.
So, the episode of magma ascent and dyke intrusion is consistent
with regional kinematics and tectonics. Feuillet et al. (in rev.) sug-
gested that the new volcanic edifice belongs to a volcanic ridge,
extending between Mayotte up to the new volcanic edifice, which
is part of a regional en ´
echelon system where dextral movement
between EARS and Madagascar is accommodated. Famin et al.
(2020) described Comores archipelago as the northern boundary of
Lwandle Plate were volcanic cones are associated to Riedel shear
and en ´
echelon extension fractures.
The 2018 Mayotte crisis shows some similarities with past telluric
crises involving dyke intrusions. For example, Ahmed et al. (2016)
suggested a dyking event related to intense seismicity involving
29 M>5.0 earthquakes in the Gulf of Aden over a few months,
where seismic migration is interpreted as magma ascent before
dyke propagation. As in the case of Mayotte, cluster 1b (June 2018)
may correspond to a phase of shallow injection (upward and lateral
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The Mayotte 2018–2019 seismo-volcanic crisis 41
migration). Regarding the reservoir depth (28 km) or distance of mi-
gration, the Mayotte seismo-volcanic crisis shares some similarities
with that of 2011–2012 at El Hierro, Canary Island (deep reservoir
depth: 20–25 km), where the activity was characterized by several
phases, including phases of migration and submarine eruption (e.g.
Mart´
ıet al. 2013).
8 CONCLUSION
We proposed a scenario for the onset and the first 18 months of
the volcanic eruption that started off-shore Mayotte in early July
2018 and was eventually confirmed by direct observation 10 months
later (Feuillet et al. in rev.). Our scenario is mostly based on the
available seismological, GNSS and InSAR data acquired at Mayotte
and surroundings during and before the crisis.
The first emergence of the seismic crisis was on 10 May 2018
and the main shock on May 15.
Two phases preceded the proper eruption that initiated around
July 3 according to the GNSS data.
First, from May 10 to May 31, the intense seismic activity, with
most of the largest events of the crisis (29 among the 32 ML≥5.0
events reported), is interpreted as a phase of fracturing of the crust
but without the intrusion of magma in the fractures. This is strongly
supported by the pattern of the deformations during those 3 weeks.
Then, during the entire month of June 2018, a second and distinct
phase occurred, with the injection of the magma in the fissures. This
injection phase started with a week of intense seismicity (equivalent
to Mw=6.2) and the migration of the seismicity towards southeast
and the surface. The GNSS data are also recording the magma
ascent during this period at an average speed that we can estimate
around 0.013 m s−1. This speed might be confirmed by the analysis
of the crystals present in the fresh basalts collected at the sea bottom
on the slopes of the new volcano, or could be used to constrain the
analysis of those crystals.
Based on the clear and abrupt change in the GNSS time-series,
we guess that the proper start of the eruption is on 3 July 2010 ±
2 d, with a flux of magma that doubles within thefeeding cracks with
respect to June, and the beginning of building of the new volcano
at the sea floor.
During the 2 months of eruption that follow, in July and August,
the seismicity is low, while the slope of the GNSS changes is totally
stable, thus the flux of magma perfectly steady ±5 per cent.
The modelling of the ground deformation is robust but with a
20 km range of possible models along the longitude axis, due to
the peculiar geometry of the GNSS stations. Our preferred model is
with a deflating reservoir located ∼30 km east of Mayotte at 28 km
depth.
At the end of June, approximately at the same time the proper
eruption starts, a new cluster of seismicity appears, mostly NW
from latest events of the first cluster and close to its initial events.
Its activity is moderate in comparison with the first cluster, but it
remains active during the entire crisis.
A third seismic cluster starts also around the beginning of the
deflation, west from the second one. This third cluster is more
active than the second one and was still active at the end of the
observation period. Those two clusters might correspond to the
propagation of part of the magma at depth in dykes or sill, but
it might also correspond to cracks triggered by the on-going crisis
without the need of invoking the injection of magma in those cracks.
There is no signature of those clusters in the GNSS data that are
totally dominated by the intense deflation.
An intense long-lasting monochromatic very long period
(∼16 sec.) seismic event occurred on 11 November 2018, last-
ing more than 20 min, corresponding presumably to the oscilla-
tion of fluid in the system, either magma in the chamber or in
a long dyke. This tremor was roughly located between clusters 2
and 3.
The peak of the effusion rate, 93.8 m3s−1is dated around the end
of 2018. Then during the entire year 2019 the effusion rate decreased
smoothly. Between October 2018 and April 2019, a sustainable
increased seismicity appears, essentially regarding cluster 3 but also
cluster 2. Then, at the end of April 2019, seismicity enters a new
slowdownphase. The total magma released from depth, according to
our model, and for the period July 2018–January 2020, is estimated
between 2.5 and 3 km3.
Whereas the plumbing system and local structure are poorly
known, adding a smaller secondary source in our model of deflation
improves very significantly the fit of the GNSS data. This second
source that we localize, quite robustly, beneath the south of the
Mayotte main island might represent the area where material is
predominantly collected to partly refill the major depleted area in a
process that might be a viscous relaxation flow.
The seismo-volcanic crisis of 2018–2019 offshore Mayotte has
shown activity of a WNW–ESE aligned system. It marks the end of
a period of quiescence several 1000 yr of the Mayotte island edi-
fice and occurred in a poorly known and poorly instrumented area.
However, it confirms the alignment of active seismicity and vol-
canism around Comoros archipelago, between the Somalian Plate
and Lwandle block in an overall extensive regional context, con-
sistent with trans-tensional regime in the area. The ongoing crisis
in Mayotte might be important for teaching us more about the dy-
namics between the Somali Plate and the Lwandle block and the
related volcanism. Improving this knowledge and improving the
knowledge of the distribution, alignment, and ages of the offshore
volcanic features, especially around the main islands, may lead to a
better understanding of the behaviour, evolution, and related hazards
of this peculiar area.
ACKNOWLEDGEMENTS
We are grateful to Jean Battaglia, Julien Bernard, Nicolas Chamot-
Rooke, Alison Colombain, Christine Deplus, Anne Deschamps,
Andrea Di Muro, C´
ecile Doubre, R´
emi Dretzen, Val´
erie Ferrazzini,
Nathalie Feuillet, Bruno Garayt, C´
ecile Gracianne, Marc Grunberg,
Eric Jacques, the Gempa team, the Guralp hotline, Philippe Jousset,
Philippe Kowalski, Fr´
ed´
eric Lacquement, J´
erˆ
ome Lambert, Yohann
Legendre, Arnaud Lemarchand, H´
el`
ene Lyon-Caen, Estelle Maison-
neuve, Kristel Meza Fajardo, Fabien Paquet, Aline Peltier, S ´
ebastien
Saur, Jean-Marie Saurel and J´
erˆ
ome van der Woerd for their help and
the fruitful discussions. IPGP and R´
eNaSS performed the observa-
tory work during August 2018. Didier Bertil reprocessed the initial
locations and magnitudes. IPGP shared the data from the Karthala
Observatory in Grande Comore. IGN/RGP (http://rgp.ign.fr)and
its partners TERIA, L´
el@sarl and CNES provided the GNSS data,
RESIF-RAP network hosted the YTMZ accelerometer (RESIF is a
national Research Infrastructure, recognized as such by the French
Ministry of Higher Education and Research. RESIF is managed by
the RESIF Consortium, composed of 18 Research Institutions and
Universities in France. RESIF is additionally supported by a public
grant overseen by the French National Research Agency (ANR) as
part of the ‘Investissements d’Avenir’ program (reference: ANR-
11-EQPX-0040) and the French Ministry of Ecology, Sustainable
Development and Energy), EDUSISMO (www.edusismo.org)was
Downloaded from https://academic.oup.com/gji/article-abstract/223/1/22/5850758 by guest on 10 July 2020
42 A. Lemoine et al.
responsible for the existence of the station MCHI, BCSF (Bu-
reau Central Sismologique Franc¸ ais) provided two RaspberryShake
stations. We are grateful to ISC (www.isc.ac.uk) for its cata-
logue of seismicity. This study was funded by the French Min-
istry for the Ecological and Inclusive Transition (MTES) and by
BRGM fundings from the French Ministry for Higher Educa-
tion, Research and Innovation (MESRI). We used the software
GMT to produce maps (Wessel & Smith 1995), SAC (Gold-
stein et al. 2003), Seiscomp3 (Hanka et al. 2010), GIPSY Oasis
6.4 (https://gipsy-oasis.jpl.nasa.gov/) and bathymetric data from
Homonim project (SHOM 2015), and Gebco 2014 (Weatherall
et al. 2015). Author contribution statement: DB built the monitoring
tools, DB, AR and AL analysed the seismological data assisted by
RHC, PV performed GNSS computations, PB analysed the GNSS
data and made the deformation model, MF performed InSAR data
analysis, assisted by MM and DR, IT brought her knowledge on
the local context, and PB and AL wrote the paper. Detailed com-
ments provided by two anonymous reviewers greatly improved the
manuscript.
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9 SUPPORTING INFORMATIONS
Supplementary data are available at GJI online.
Figure S1. b-Values for the three clusters.
Figure S2. The 11 November 2018 very low frequency tremor and
embedded events.
Figure S3. The 11 November 2018 very low frequency event: par-
ticle motions.
Figure S4. Modelling of the tremor decay.
Figure S5. Time-series for the coordinates of the six GNSS stations
during phases A, B, C and D.
Figure S6. Ratio between horizontal and vertical motion during
phase D.
Figure S7. Time-series since January 2018 at stations BDRL,
GAMO, KAWE and MAYG.
Figure S8. InSAR (Sentinel-1) vertical and E–W components rate
maps.
Figure S9. Parametric analysis of the depth and source type.
Tab le S1 . ITRF2014 horizontal velocity of GNSS stations around
Mayotte.
Tab le S2 . The stations of the Mayotte monitoring seismic network.
Tab le S3 . The 32 events with ML≥5.0.
Tab le S4 . Regional velocity model
Tab le S5 . Three intense Very Long Period tremors.
Tab le S6 . The GNSS stations in Mayotte.
Tab le S7 . Velocity anomalies at the Mayotte GNSS stations during
deformation phase D.
Tab le S8 . Temporal evolution of the velocity anomalies during
phase D.
Tab le S9 . Deformation during phase C.
Table S10. Deformation during phase B.
Table S11. Evolution of the effusion rate.
Table S12. Evolution of the effusion rate estimated over periods of
3 months each by assuming a dual source.
Table S13. Analysis of the fits and effusive volumes as a function
of source depth and model
Please note: Oxford University Press is not responsible for the con-
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authors. Any queries (other than missing material) should be di-
rected to the corresponding author for the paper.
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