# Reassessment of the 1907 Sumatra “Tsunami Earthquake” Based on Macroseismic, Seismological, and Tsunami Observations, and Modeling

Stacey Servito Martin, Linlin A Li, Emile Okal, Julie Morin, Alexander

Tetteroo, Adam Switzer, Kerry Sieh

► To cite this version:

Stacey Servito Martin, Linlin A Li, Emile Okal, Julie Morin, Alexander Tetteroo, et al. Re-assessment of the 1907 Sumatra “Tsunami Earthquake” Based on Macroseismic, Seismological, and Tsunami Observations, and Modeling. Pure and Applied Geophysics, 2019, 176 (7), pp.2831-2868. 10.1007/s00024-019-02134-2. hal-02529394

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¹ Earth Observatory of Singapore, Nanyang Technological University, Singapore

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² Department of Earth & Planetary Sciences, Northwestern University, Evanston, U.S.A.

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³ Laboratoire Magmas et Volcans, Université Clermont Auvergne, Clermont-Ferrand, France

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⁴ Institute for History, Leiden University, Netherlands

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⁵ Asian School of the Environment, Nanyang Technological University, Singapore

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† Corresponding author: [7point1@gmail.com](mailto:7point1@gmail.com)

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A reassessment of the 1907 Sumatra “tsunami earthquake” based on
macroseismic, seismological, and tsunami observations and modelling
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**Abstract**

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On 4 January 1907 an earthquake occurred off the west coast of Sumatra, Indonesia, with an instrumental surface-wave magnitude (Ms) in the range of 7.5 to 8.0 at periods of ~40s. The tsunami it generated was destructive on the islands of Nias and Simeulue where it killed hundreds and gave rise to the legend of the S'mong. This tsunami also appears in records in other parts of the Indian Ocean basin, and as far as La Réunion island. Relative to its instrumented magnitude, the size of the tsunami was anomalous qualifying it as a "tsunami earthquake." However, unusually for a tsunami earthquake the shaking on Nias was severe (6-7 EMS). We revisited the 1907 earthquake with a multidisciplinary approach by extracting evidence describing shaking effects or the tsunami from written documents and by acquiring new seismograms. Combining these, we discriminated two large earthquakes within an hour of each other with clear differences in seismological character. The first we interpret to be a tsunami earthquake with characteristic low levels of shaking and for which we estimate an average seismic moment (M₀) 2.5 x10²⁸ dynes*cm (Mw ≈ 8.2) in the frequency band 6–8 mHz. These records document a regular growth of moment with period: at the longest resolvable period (~170 s) the moment magnitude approaches Mw ≈ 8.4. Our relocations suggest an epicentral location close to the trench for the mainshock. For the second earthquake that was damaging on Nias, we estimate Ms ≈ 7 based on seismograms and phase data from Manila and Osaka. We also identified two Ms ≈ 6 aftershocks within 24 hours of the mainshock. We propose a model of seismic rupture which provides an acceptable fit to our new dataset of tsunami run-up and inundation values from 88 local and far-field locations in the Indian Ocean basin.
## 42 Introduction

Numerous large earthquakes have occurred off the West coast of Sumatra, Indonesia, during the historical and instrumental eras (e.g. Newcomb & McCann, 1987), and in particular, since 2000. These events include a tsunamigenic earthquake on 4 January 1907 near the islands of Simeulue and Nias (Figure 1), which Newcomb & McCann (1987) associated with shaking of such severity that “people on Nias could not stand.” As many as 370 people were killed on Nias and at least 1818 lives were lost on Simeulue (Koloniaal Verslag, 1907-08); 1,205 in Tapah (nl: Tapak) district, 431 in Simeuloeë Rajou, 130 in Salang and 52 in Leuköon on Simeulue (Bataviaasch Nieuwsblad, 12 February 1907). It was feared 1,000 people were killed at Koeladeh on Tapa alone (Utrechts Nieuwsblad, 14 February 1907). This led to the disaster being embodied in myth and legend on the island of Simeulue (e.g. McAdoo et al., 2006; Syafwina, 2014; Rahman et al., 2017). As discussed below, the 1907 tsunami was also recorded in the far field, as far away as La Réunion (Bertho, 1910), which gives the event a clearly anomalous character in the context of the comparatively low “Pasadena” magnitude ($M_{PAS} = 7.6$) assigned to the earthquake by Gutenberg & Richter (1954), and for long, the only available measure of its size.

More recently, Kanamori et al. (2010) conducted an extensive seismological study of the 1907 earthquake based on a number of historical seismograms. While they did not compute a seismic moment through waveform fitting, they measured a surface wave magnitude $M_s = 7.8 \pm 0.25$ and estimated a moment magnitude $M_w = 7.8$ by scaling time domain amplitudes of body and surface waves of the 1907 earthquake to those of nearby modern earthquakes with known moment tensors. These results were obtained in the period range 40 – 50 s, but Kanamori et al. (2010) stress that the source was obviously longer, and thus the moment should be larger at longer periods, suggestive of a "tsunami earthquake".
The term “tsunami earthquake” was first used by Kanamori (1972) to discuss the sources of the 1896 Meiji Sanriku and the 1946 Unimak (Aleutian Islands) earthquakes, both of which resulted in anomalously large tsunamis with respect to their instrumental magnitudes. This type of event can be distinguished based on disproportionate relationships between surface wave magnitude ($M_s$) and seismic moment ($M_0$), from the observation of longer than expected process times despite small rupture areas (Sykes, 1971; Kanamori, 1972; Pelayo & Wiens, 1992; Polet & Kanamori, 2000), from the lack of evidence for abnormal stress drops (Pelayo & Wiens, 1992), and from lower than anticipated macroseismic intensities (Kanamori, 1972; Fukao, 1979; Bourgeois et al., 1999). The longer process times result in red-shifting of the source spectrum, and in inconsistencies between deficient seismic magnitudes measured at short-to-moderate periods (thus relevant to macroseismic effects) and enhanced ultra-long period seismic moments (controlling the generation of tsunamis). The ruptures of tsunami earthquakes have been observed to propagate toward the trench axis (Polet & Kanamori, 2000) on very shallow dipping faults or on splays (Fukao, 1979; Pelayo & Wiens, 1990, 1992) located in weakly coupled regions of aseismic convergence (Pelayo & Wiens, 1990, 1992; Bourgeois et al., 1999), or at the very top of the plate interface under conditions of sediment starvation leading to a jagged rupture (Tanioka et al., 1997). Rupture velocities for such earthquakes are also less than expected for typical earthquakes owing to the low rigidity of materials proximal to the trench axis (Fukao, 1979; Pelayo & Wiens, 1990, 1992; Heinrich et al., 1998; Ihmlé et al., 1998). In addition, tsunami earthquakes can occur as mainshocks, which Okal & Saloor (2017) qualified as “Primary Tsunami earthquakes” (PTEs), or as “Aftershock Tsunami Earthquakes” (ATEs), following a larger, regular megathrust event. The exceptional tsunami of the 2011 Tohoku event may be the result of the combination of a regular megathrust event, and of a lower-frequency rupture of the shallowest
portion of the interface, which might have qualified as an ATE, but for its occurrence only 3
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minutes after the initial nucleation (Satake et al., 2013). Enhanced tsunami generation following
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moderate-to-large earthquakes can occasionally be traced to ancillary phenomena, such as
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triggered submarine landslides (Synolakis et al., 2002) or volcanic processes (Satake & Kanamori,
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1991; Fukao et al., 2018). However, these are generally not labelled as “tsunami earthquakes.”
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With the exception of the 2010 Mentawai event (e.g. Newman et al., 2011a; Hill et al., 2012), no other tsunami earthquakes have been conclusively identified off the Sumatran coast during the modern or historical period. Although Kanamori et al. (2010) concluded in more general terms that the 1907 Sumatra earthquake held all the hallmarks of a tsunami earthquake, key aspects of this event remain unaddressed, including (i) conclusive, quantitative evidence of its nature as a “tsunami earthquake”; and (ii) an estimate of the geometry and slip parameters of the source supporting the reported distribution of the tsunami. In addition, perplexing observations clearly in need of further study include (iii) the anomalously violent ground motions (Newcomb & McCann, 1987; Kanamori et al., 2010) in comparison to other tsunami earthquakes; (iv) the lack of aftershocks; and (v) the lack of land level changes comparable to those identified for other large earthquakes in the Simeulue-Nias region, e.g., Meltzner et al., (2012; 2015).

In this article we employ a multidisciplinary approach to tackle these points. Through a
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scrutiny of original macroseismic reports and the systematic analysis of a number of seismograms,
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we separate the mainshock from a previously unsuspected large aftershock, occurring only ~53
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minutes later. Waveform analysis of digitised seismograms allow the quantification of the
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mainshock at mantle wave periods up to 170 s, and further demonstrate significant differences in
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source spectra between the mainshock and the large ~06:12 GMT aftershock. We also compile a
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110 set of 88 qualitative and instrumental observations of the tsunami in the Indian Ocean basin
111 (including nil reports) and utilise a subset of these to prepare a slip model for the 1907 mainshock.

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113 **Reassessing Macroseismic Intensities and the Discovery of Event II**

114 The earliest macroseismic study of the 1907 Sumatra earthquake by Newcomb & McCann,
115 (1987) is a dual zoned intensity map (Figure 2a) utilising the Modified Mercalli intensity scale
116 (MMI) to distinguish between what they defined as weak (MMI I – IV) and strong shaking (MM
117 V – VII), for what was assumed to be a single event. We use the 1998 European Macroseismic
118 Scale (EMS-98; Grünthal, 1998) to assess intensity with our choice of macroseismic intensity
119 assessment adhering to modern practice (e.g. Toppozada & Real, 1981; Ambraseys & Douglas,
120 2004). The EMS-98 scale is equivalent to MMI (Musson et al., 2010); its merits and its suitability
121 for use outside Europe are discussed by Martin & Hough (2015; 2016). In this and subsequent
122 sections, non-English text appears in parenthesis (bh: Bahasa Indonesia, nl: Dutch; fr: French; de:
123 German) accompanying the modern reformed Indonesian or English spelling, e.g., tsunami (nl:
124 vloedgolf). Our re-assessment of intensity focused on an annual summary of felt earthquakes in
125 the Dutch East Indies (Anonymous, 1909) by the Royal Magnetic and Meteorological Observatory
126 (nl: Koninklijke Magnetisch en Meteorologisch Observatorium, KMMO) that was culled from
127 official correspondence. This was supplemented with further accounts extracted from colonial
128 newspapers published in Indonesia and in the Netherlands (see Appendix A and B).

129 Accurate time keeping is vital to accurately associate reports with individual earthquakes.
130 In this respect, the KMMO summary was invaluable as it printed the standard time in Jakarta
131 (known as Batavia Time or BT), which was 7h 7m ahead of Greenwich Mean Time (GMT), for
seismic disturbances recorded instrumentally at Jakarta (nl: *Batavia*) or Bogor (nl: *Buitenzorg*). It
also published the difference in time at places outside Jakarta with respect to Batavia Time. For
example, BT was 36 minutes ahead of local time at Gunungsitoli (nl: *Goenoeng Sitoli*) on Nias,
and 17 minutes behind local time at Pacitan in central Java. Thus, the local time was GMT + 06:31
at Gunungsitoli, GMT + 07:07 at Jakarta, and GMT + 07:24 at Pacitan.

A surprising result of our macroseismic re-analysis (Tables S1 – S2) was the discovery of
two distinctly separate earthquakes that were felt on Nias within an hour of each other on 4 January
1907 (Figure 2b, 2c). A seismograph in Jakarta (Anonymous, 1909) recorded the first at 12:29 BT
(05:19 GMT; henceforth Event I) and the magnetograph at Bogor, detected the second at 13:24
BT (henceforth Event II). Engelbertus Schröder, the civil administrator (nl: *kontrolleur*) at
Gunungsitoli on Nias (Appendix B) provided the strongest evidence with which to make this
distinction; he described a weak but long duration earthquake at about midday, followed by three
"flood waves" (nl: *vloedgolven*) at Toyolawa, which were later followed at 12:50 local time by a
second violent earthquake at Gunungsitoli (Bataviaasch Nieuwsblad, 22 March 1907). His account
appeared in several newspapers such as the *Soerabaisch Handelsblad* (23 March 1907) and in the
*Nieuws van den Daag voor Nederlands-Indië* (22 March 1907); the latter reported the time of the
second earthquake to be "12:05." We believe this is a typesetting error as the identical account
printed on the same day by the *Bataviaasch Nieuwsblad* (22 March 1907) states "12:50," and
importantly 12:50 is also reported by *Schröder* (1917a) himself. The occurrence of two felt
earthquakes separated by ~53 minutes is also corroborated by independent observers from
Gunungsitoli (De Padanger, 26 January 1907), Natal, and Talu (nl: *Taloe*), and discussed below
based on seismograms. Also, note that *Masó* (1907) interpreted the Manila records as requiring
two events, but pointed out the lack of evidence for Event II at Zikawei (Shanghai) or Europe; we
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believe this to be incorrect, at least with respect to Shanghai, as Event II appears in the Zikawei bulletin (de Moidrey, 1912).

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The weak, long duration ground motions in the near field, and the barely perceptible shaking at larger distances associated with Event I (Figure 2b; Table S1) are characteristic of tsunami earthquakes (e.g., *Kanamori*, 1972). In contrast, the descriptions of shaking and damage at Gunungsitoli from Event II (Figure 2c; Table S2) lead us to believe it was responsible for the violent shaking on Nias that was misconstrued as the mainshock by *Newcomb & McCann* (1987) and flagged as anomalous by *Kanamori et al.* (2010). Countless houses were destroyed on Nias by Event II and even weeks later many people were living under temporary shelters (bh: *pondoks*) put up near their dwellings (*Bataviaasch Nieuwsblad*, 22 March 1907). Unusually, neither earthquake was reported in an official summary of news and affairs for January 1907 from Aceh (nl: *Atjeh*) i.e. “Atjeh-verslag over Januari” (*Java Bode*, 16 April 1907). These earthquakes also do not appear in catalogues of felt earthquakes in Thailand (*Nutalaya et al.*, 1985), nor were they reported from anywhere else in south-east Asia.

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## Location of the Mainshock (Event I)

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The earliest published locations for Event I come from *Turner et al.* (1912) and *Szirtes* (1912a), who used arrival times from over 66 stations. As summarized by *Kanamori et al.* (2010), and not surprising for an event at the beginning of the 20th century, epicentral estimates are poorly constrained and are significantly scattered (Figure 3). *Gutenberg & Richter's* (1954) solution (2°N, 94 ½°E) is on the outer rise and slightly west of the solution (2°N, 95°E) by *Turner et al.*
(1912), but both are very unlikely locations for a tsunami earthquake. By contrast, *Szirtes* (1912a) reports a location by T.H. Staverman (2°N, 96¼°E) in the immediate vicinity of the trench.

In their recent compilation of relocations for early instrumental earthquakes, *Storchak et al.* (2013) locate the event at (1.87°N, 94.21°E), near the *Turner et al.* (1912) and *Gutenberg & Richter* (1954) epicentres, and with a remarkably small error ellipse (semi-axes: 28 km and 16 km). The most recent version (5.0) of the ISC-GEM catalog moves the epicentre North by 61 km, to (2.422°N; 94.258°E); it is noteworthy that the confidence ellipse (again very small at 20 by 17 km) does not intersect its counterpart for *Storchak et al.'s* (2013) solution (**Figure 3**), thus casting doubt on their reliability since the two locations are supposed to be derived from the same algorithm. On the other hand, a systematic grid search of the same dataset confirms a large scatter of possible solutions but proposes a best fitting epicentre at 2.67°N and 95.01°E (D. Di Giacomo, *personal communication*, 2017), 125 km NE and across the plate boundary from *Storchak et al.'s* (2013) solution.

*Kanamori et al.* (2010) relocated Event I based on travel times listed in Gutenberg's personal notepads (*Goodstein et al.*, 1980), with an emphasis on differential *S* – *P* times, to minimize errors due to clock uncertainties. Their solution (2.48°N, 96.11°E; **Figure 3**) locates approximately 60 km closer to the trench than the 2004 and 2005 epicenters, at the seaward limit of the zone of interplate seismicity (*Pesicek et al.*, 2010), this configuration being reminiscent of the 2010 Mentawai tsunami earthquake (*Newman et al.*, 2011a; *Hill et al.*, 2012). However, the 95%-confidence ellipse from *Kanamori et al.* (2010), whose parameters are listed in their Appendix A (*Kanamori et al.*, 2010; p. 371), extends ~500 km, from the outer rise, beyond the *Gutenberg & Richter* (1954) and the *Storchak et al.* (2013) locations, all the way to inland Sumatra.
We performed an independent relocation using the dataset listed by the ISC, and the interactive method of Wysession et al. (1991) which includes a Monte Carlo algorithm injecting Gaussian noise in the data. For an event in the 1900s, we give the noise a standard deviation σG = 12.5 s. Our solution converges to (2.24°N; 94.72°E: O.T. 05:19:13 GMT), a location on the outer rise, in the vicinity of the Storchak et al. (2013) and Gutenberg & Richter (1954) epicenters, but our confidence ellipse extends across the trench and grazes the solution from Kanamori et al. (2010); it is essentially contained inside these authors' ellipse.

While all these relocations have very large uncertainties (we regard the Storchak et al. (2013) confidence ellipse as deceptively small), the emerging general pattern (Figure 3) is that Event I was probably significantly displaced up-dip and trench ward with respect to the main seismogenic zone at the interplate contact, to the general area located between the ISC grid-search solution and that of Staverman, in a tectonic environment reminiscent of that of the Mentawai tsunami earthquake of 25 October 2010 (Newman et al., 2011a; Hill et al., 2012). This is our preferred location for the epicentre of Event I (2.5°N; 95.5°E) and is indicated by a red star on Figure 4. Incidentally, this preferred location lies within a zone of low seismic productivity between 1918 and 2007, and adjacent to, but not within, the aftershock zones of the 2004 and 2005 earthquakes (Engdahl et al., 2007; Pesicek et al., 2010).

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216 **Event II and later aftershocks**

Our study is the first to identify several aftershocks associated with the 1907 Sumatran earthquake. For the newly discovered large earthquake (Event II), we were eager to find seismological evidence with which to quantify, and possibly locate it. Waveforms were far fewer
than for the mainshock, but we identified it on new records from Manila (Masó, 1907) and Shimla (Patterson, 1909), and on the Osaka record scrutinised by Kanamori et al. (2010). Station bulletins from Manila (Masó, 1907), Osaka (Anonymous, 1931), and Shanghai (de Moidrey, 1912) also list it. Anonymous (1909) provides the time a magnetograph at Bogor detected this shock i.e. 13:24 BT (~06:18 GMT) which gives an origin time of 06:15 GMT correcting for travel time to Bogor.
We note, however, that the arrival time of the mainshock on this instrument is unknown. The difference in the arrival times of the P waves from Event II with those of the mainshock at Osaka (Anonymous, 1931), the difference in the arrival times of the S waves from both events at Manila (Masó, 1907) and Shanghai (de Moidrey, 1912), and an inspection of the waveform recorded at Manila (Masó, 1907) suggests a ~53 min interval between the two. We therefore estimate an origin time of ~06:12 GMT which is also supported by the first-hand account by Schröder (1917a) of two shocks ~50 minutes apart.

The paucity of instrumental records prevents an instrumental location for Event II, but an epicentral distance (5420 km) to Osaka has been previously published (Anonymous, 1931). Unfortunately, this datum is of little help, since Osaka is at essentially equal distances (within 10 km) from the preferred locations of Event I and II. Finally, an epicentre in the region of Nias for Event II is further supported by the damage at Gunungsitoli (Table S2) and the destruction of houses elsewhere on Nias (Bataviaasch Nieuwsblad, 22 March 1907). It is also likely that uplift inferred to have occurred in the Hinako islands (Bataviaasch Nieuwsblad, 22 March 1907) and documented from a fossil coral microatoll on southern Nias (Meltzner et al., 2015) were related to Event II, and would suggest an epicentral region in south-central Nias or off its south-central coast.
Two strongly felt aftershocks (Anonymous, 1909) at ~09:55 GMT on 4 January (Figure 2d; Table S3) and ~23:52 GMT on 5 January (Figure 2e; Table S4) could be co-related with sparse regional
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and teleseismic recordings (Levitski, 1907; Pechau, 1907; Geiger, 1909; Szirtes, 1912b); origin
times for both are corrected with respect to travel time to Jakarta from the Simeulue-Nias region.

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In addition to the above, nineteen felt aftershocks were counted on Simeulue on 4 January (Bataviaasch Nieuwsblad, 12 February 1907). Earthquakes were felt daily (Nieuws van Den Dag, 24 January 1907) with rarer aftershocks experienced after 15 January 1907 (De Sumatra Post, 1 February 1907). Earthquakes were also felt at Sinabang on Simeulue at 23:45 LT on 19 January 1907 (Bataviaasch Nieuwsblad, 12 February 1907) and on 8 July 1907 (Het Nieuws van Den Dag, 20 September 1907), while on Nias, aftershocks were felt continuously until 17 January (Bataviaasch Nieuwsblad, 29 January 1907). Aftershocks later than 5 January 1907 were reported only from Nias (Anonymous, 1909). Well-timed felt earthquakes are almost absent from Simeulue except for a shock that was quite heavy (nl: vrij hevig) at Sinabang around midday (~12:42 BT) on 4 January 1907 (Utrechts Nieuwsblad, 14 February 1907) without causing any damage (De Sumatra Post, 1 February 1907). We suspect this shock could be related to Event I or a strong local aftershock, neither of which we can definitively rule out. Another large earthquake was felt in the Banda Aceh region later on 21 November 1907 (Figure 2g), but this was probably unrelated to Event I.

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Conventional Magnitude Estimates

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Estimates of conventional magnitudes for Event I have been offered by Gutenberg & Richter (1954) ($M_{PAS}$ = 7.6), Abe & Noguchi (1983) ($M_s$ = 7.6), and recomputed by the ISC (Storchak et al., 2013) ($M_s$ = 7.8). Duda (1965) proposed a figure of 7.8, the nature of his magnitude being unspecified. Kanamori et al. (2010) estimated $M_s$ by comparing 20 s amplitudes
on records available for 1907 to those of records simulated on historical instruments for modern
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earthquakes with abundant determinations of $M_s$, and concluded the 1907 shock probably had $M_s$
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$= 7.8 \pm 0.25$. A review of amplitude data available from Milne instruments (Levitski, 1909; Turner
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& Milne, 1908a, 1908b; 1909, 1910) in the manner of Abe & Noguchi (1983) and Abe (1988)
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yields a Milne surface-wave magnitude ($M_M$) of $7.9 \pm 0.17$ (1$\sigma$) which, while higher than previous
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estimates of $M_s$ (e.g. Gutenberg & Richter, 1954; Abe & Noguchi, 1983), is in the same range as
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$M_s$ determined by Kanamori et al. (2010).
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In the case of Event II, we recomputed conventional magnitude estimates from those
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reports of ground amplitude for which associations could be made, using established formulae and
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guidelines (Kárník et al., 1962; Geller & Kanamori, 1977; Ambraseys & Douglas, 2000). Our best
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estimate of $M_s$ is 7.1, which represents an average for Manila (6.98) and Osaka (7.17). In the case
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of the aftershocks of 4 January (~09:55 GMT) and of 5 January (~23:41 GMT), we estimated $M_s$
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$= 5.8$ for the first one, and a Milne magnitude ($M_M$) in the range 5.5 – 6.0 for both events; however,
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we emphasise that these values are very poorly constrained, and are associated with uncertainties
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of up to $\pm 0.5$ (1$\sigma$).
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## Reassessment of long-period seismic moment and source slowness
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For Event I, *Kanamori et al.* (2010) also inferred a seismic moment (*M*₀) of about 6 × 10²⁷
dyn*cm for Event I at periods of ~50 s, through a comparison of time domain amplitudes of body
waves at Göttingen with those obtained at the nearby Black Forest Observatory (BFO) during the
earthquake of 2 November 2002. This method, which assumes the same focal geometry for both
events (strike (ϕ) = 297°, dip (δ) = 16°, slip angle (λ) = 73°), is confined to the time domain and,
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as such, tacitly assumes that Event I had a source spectrum similar to that of the reference event,
and hence followed scaling laws. However, those are expected to be violated precisely by tsunami
earthquakes, whose source spectrum is red-shifted towards lower frequencies with respect to a
more traditional source, such as the 2002 event for which we have verified a totally regular energy-
to-moment parameter (Θ = -4.92). For this reason, it is necessary to further explore the source
spectrum of Event I in the frequency domain to constrain its seismic moment at periods typical of
mantle waves.

In addition to the waveforms interpreted by *Kanamori et al.* (2010), we revisited the record
from Göttingen (GTT), and located three previously unused waveforms from Pulkovo, Russia
(Golitsyn, 1908), Shimla, India (*Patterson*, 1909), and Manila, Philippines (*Masó*, 1907). The
Pulkovo record only provides a partial, single component record for Event I, but none for
aftershocks (Golitsyn, 1908) while the Shimla seismogram clipped during Event I (*Patterson*,
1909). The Manila record was written on a Vicentini mechanical seismometer with a period T =
1.4 s, which makes it the only short-period recording available to us of the 1907 Sumatra
earthquakes, and its crucial importance to our study is discussed in a subsequent section.

We use the first passages of Rayleigh ($R_1$) and Love ($G_1$) waves at Göttingen (GTT) for
which precise metadata is available, with records hand-digitized at a sampling rate of 1 s. Our
results are presented **(Figure 4)** in the mantle range of frequencies (6 – 10 mHz), expressed as a
mantle magnitude corrected for focal mechanism ($M_C = \log_{10} M_0 - 20$ with $M_0$ in dyn*cm) in the
formalism of Okal & Talandier (1989). We use a slightly adapted focal mechanism ($\phi = 320°$, $\delta =
15°$; $\lambda = 90°$), rotated only 8° in the formalism of Kagan (1991) from *Kanamori et al.'s* (2010), in
order to reduce the scatter between Love and Rayleigh spectral amplitudes. Note that this new
mechanism is also closer to the geometry of the nearby 2005 Nias megathrust earthquake ($\phi =$
310 333°, δ = 8°; λ = 118°). Our results (Figure 4) suggest an average moment M₀ ≈ 2.5 × 10²⁸ dyn*cm
311 (M_c ≈ 8.2) at periods between 100 s and 160 s; at the largest resolvable period (~170 s) the
312 magnitude estimate approaches M_c ≈ 8.4. This value of the seismic moment is about 4 times larger
313 than proposed by Kanamori et al. (2010) from an estimate of time-domain amplitudes of S waves
314 around 50 s and demonstrates slowness in the source of the 1907 mainshock, in clear agreement
315 with its nature as a “tsunami earthquake.” In addition, Figure 4 documents a growth of moment
316 with period, with a regression slope for M_c vs. frequency of -0.09 logarithmic units per mHz. This
317 number is comparable to values obtained for documented tsunami earthquakes (e.g., Java, 2006:
318 -0.11; Mentawai, 2010: -0.08; Manzanillo, 22 June 1932: -0.14; Hikurangi, 1947: -0.07 and
319 -0.08); by contrast, traditional subduction events feature lower slopes (in absolute value) that do
320 not exceed -0.05 (e.g., Maule, 2010: -0.05; Illapel, 2015: -0.05; Manzanillo, 03 June 1932: -0.01;
321 see Okal & Saloor, 2017). This property provides a quantitative confirmation of the nature of
322 Event I as a tsunami earthquake in conjunction with its anomalously low felt intensities.
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324 Ratio of estimated energies between Events I and II

325 We reproduce on **Figure 5** a copy of the Manila seismic record from *Masó* (1907). The
326 seismogram in the top frame (**Figure 5a**) was written on a Vicentini mechanical seismometer with
327 a period T = 1.4 s, which makes it the only short-period recording available to us of the 1907
328 Sumatra earthquakes. The distance to Manila is Δ = 29°. Even though the orientation of this
329 horizontal seismogram (NNW-SSE or 337 - 157°) is pure transverse (the back-azimuth at Manila
330 being β = 247°), this record clearly shows the *P* waves from two earthquakes separated by ~53
331 min, corresponding to the two shocks identified on the basis of macroseismic data. We aligned the
traces of the two events vertically to ease the comparison of their waveforms (Figures 5a; 5b).

The body waves have a dominant recorded period of ~4 s. The P waves of the mainshock are both much weaker and of longer duration than those of the aftershock. Note also that the mainshock P wavetrain consists of several subevents, lasting a total of about two minutes, while the aftershock waveform features two main packets, each lasting about 3 s. By contrast, we show on the lower frame, a record (Figure 5b) written on a longer-period Omori system with a natural period T = 6.4 s, oriented ENE – WSW, which this time is purely orthoradial. On that record, the amplitudes of the P waves from the two events have become very comparable, and the main shock displays prominent Rayleigh waves featuring a period of ~9 s at their maximum amplitude (probably reflecting the peaked response of the instrument), while the surface waves from the aftershock are much weaker.

In an attempt to quantify these observations, we enlarged the Vicentini records (Figure 5a) and digitized them at a time sampling $\delta t = 0.1$ s. Unfortunately, the Vicentini instrument was undamped (Masó, 1907), which means that it would, at least theoretically, have an infinite response at its natural period. In this context, it would not be possible to formally compute an energy-to-moment ratio and a slowness parameter ($\Theta$) as defined by Newman & Okal (1998). This is further compounded by the fact that the photographic magnification of the figure by Masó (1907) is unknown. Under the circumstances, we used the model of a very weakly damped instrument ($\varepsilon = 1.1$) and were able to obtain a ratio of estimated energies between the mainshock (Event I) and the aftershock (Event II); $E_I^E / E_{II}^E \approx 0.2$. Furthermore, we have obtained an estimate of $M_s \approx 7.1$ for the classical surface-wave magnitude of Event II in the previous section, which would correspond to a maximum moment of $\sim 4 \times 10^{26}$ dyn*cm assuming that it follows scaling laws. However, the strong intensities reported may suggest that its source is blue-shifted towards higher
frequencies, in which case the moment could be lower. Given the long period moment estimated above for the mainshock (2.5 × 10²⁸ dyn*cm) this would suggest a factor of 300 between the energy-to-moment ratios of Events I and II, amounting to a difference of ~2.5 units in their parameters Θ. While this number is not unreasonable, it comes close to the maximum range of Θ values that we have documented (see Okal & Saloor (2017) for an extensive dataset) and could be explained using both a very low value of Θ ≈ −6.5 for the mainshock, and a high value, Θ ≈ −4, for Event II. The former value is within the range of observed Θ for tsunami earthquakes (El Salvador, 2012: −6.42; Hikurangi 1947 (II, 17 May): −6.51 (Okal & Saloor, 2017)), while the latter would be typical of a "snappy" earthquake (Okal & Kirby, 2002) occurring inside the over-riding plate, such as the main normal faulting aftershock of the 2011 Tohoku earthquake (−4.12) or the 1939 Chillán earthquake (−4.04) (Okal & Kirby, 2002; Okal et al., 2016). The jagged nature of the *P* waves of the mainshock at Manila would also be in line with observations during the 1992 Nicaragua event (Polet & Kanamori, 2000) and the resulting model by Tanioka et al. (1997).

368

369 **Indian Ocean Tsunami Observations**

We were able to compile 88 tsunami observations of the tsunami generated by Event I across the Indian Ocean, including 12 for which numerical constraints could be extracted (Table S5; Figure 6; Appendix B). They were obtained from individual accounts (e.g. Schröder, 1917a; b), as well as from reports published in local newspapers. Most of the Dutch news reports were corroborated by later official Dutch correspondence (Openbaar Verbaal, 23 April 1908). First- and second-hand accounts from a few locales in the near and far field are available (e.g. Anonymous, 1909; Visser, 1922), and some were repeated by Solov'ev & Go (1974). We also
377 extracted eyewitness or second-hand accounts appearing in contemporary anthropological (e.g.,
378 Baumwoll, 2008; Rahman et al., 2017) and scientific studies (Yogaswara & Yulianto, 2006;
379 Whitlow, 2008; Fujino et al., 2014).

380 The impact of the tsunami was greatest on the southern coast of Simeulue, and on the
381 northern, and western coast of Nias (Figure 6a, b). Without providing additional location details,
382 news and official reports from Simeulue suggest the tsunami struck after Friday afternoon prayers,
383 the sea flooding inland as much as 600 to 800 m with isolated inundations of 1200 m, or to the
384 foot of the hills, that swept coral boulders and marine fauna inland (Bataviaasch Nieuwsblad, 12
385 February 1907; Utrechts Nieuwsblad, 14 February 1907; Openbaar Verbaal, 23 April 1908;
386 Yogaswara & Yulianto, 2006). Better numerical constraint is forthcoming from Nias (Figure 6a,
387 6c), for example, from Afulu (nl: Afoeloe) where the sea washed inland for over a kilometre (nl:
388 ongeveer een paal<fnref n="1" />) as far as the foot of the hills where it deposited debris that included human
389 remains and large uprooted trees (Bataviaasch Nieuwsblad, 22 March 1907). Schröder (1912a)
390 noted that a part of the bay at Afulu was filled with sand and other material (nl: Vermelding
391 verdient de opvulling van een deel de baai van Afulu met zand en ander materiaal) at the time of
392 his visit. Schröder (1912a) also reports Hulo Uma near Afulu was “robbed of all life” (nl: van alle
393 leven beroofd). Southwest of Toyolawa, only the West coast of the island of Pulau Mausi (nl:
394 Ma’uso or Maoesi) was impacted by the tsunami (Schröder, 1912a). However, to the south the
395 tsunami appears to have rapidly diminished in character at Lagundi (Figure 6b; Bataviaasch
396 Nieuwsblad, 22 March 1907).

<footnote n="1">¹ The old Dutch measure “een paal” used in colonial Indonesia is equivalent to 1507 m (Staring, 1871).</footnote>
The island of Simeulue Cut (nl: *Simeuloeë Tjoet* or *Simaloer Tjoet*) off the southern coast of Simeulue (**Figure 6a**) "was lost" or "disappeared" (nl: *is verdwenen*) with only the hill at the centre of the island above the water (**Figure 6b**), suggestive of extreme tsunami erosion or possibly, subsidence. West of Nias, on Pulau Wunga (**Figure 6b**) the tsunami swept the island from the west (*Schröder*, 1917a), destroying three fourths of it including numerous coconut trees, with debris and victims swept into the lagoon at the centre of the island (*Bataviaasch Nieuwsblad*, 22 March 1907). The height of the tsunami at Pulau Wunga is estimated at between 6 m and 15 m (*Schröder*, 1917a; *De Padanger*, 25 February 1907), and in 2005, modern residents told one of the authors (K.E.S.) that corpses were stranded in coconut trees in 1907. The only photographic evidence of this disaster are from Pulau Wunga (*Schröder*, 1917b; **Figures S1 – S2**); how representative these are of damage elsewhere on the island is unknown, but the similarity of the post-tsunami landscape to that of Pulau Sibigau (*Hill et al.*, 2010) following the 2010 Mentawai tsunami earthquake is striking.

On the Sumatran mainland (**Figure 6a**), vague accounts suggest Kayu Menang (*Aceh Post*, 10 June 2013) and Sibolga (*Visser*, 1931) were possibly flooded. At Guhanaga (formerly Geni) a flood in the Krueng Sabe river (*Aceh Post*, 10 June 2013) was attributed to the tsunami by an eyewitness but we consider this account cautiously as the date associated with it is unreliable. The tsunami was reportedly recorded by a tide gauge in the harbour at Padang (*Anonymous*, 1909) but this record has been lost.

On the Indian Subcontinent in erstwhile British India and her dependencies, tide gauges operated by the Survey of India at Port Blair (Andaman Islands), Apollo Bandar (in Mumbai), Chennai (Madras), and Karachi provided the only instrumental readings from this tsunami (*Erskine*, 1908, 1909; **Figure 6d**), but maregrams were unavailable to us. Similar instruments at
Kidderpore (Kolkata), Yangon (Rangoon, Myanmar), and Aden (Yemen) did not record the disturbance (Erskine, 1908, 1909). In Sri Lanka, it was observed at points on the southern coast in particular at Galle where the harbour filled and emptied seven times at half hour intervals, but did not overflow the jetty (Ceylon Observer, 10 January 1907).

In the Southern Indian Ocean (Figure 6f), the tsunami was observed at Mauritius (Erskine, 1909) and on Rodrigues (The West Australian, 5 January 1907). La Réunion (Bertho, 1910; Sahal et al., 2011), supported by readings on a maregraph, reports an "oscillation" (presumably peak-to-peak) of 1.7 m at Pointe des Galets, about 1/3 (if meant to represent a peak-to-peak amplitude) of the runup observed in 2004 at nearby (and vastly developed) Le Port (Okal et al., 2006a), starting at 16:20 local time (which before 1911 was GMT + 3:42), in perfect agreement with the 7.5 hr propagation time observed in 2004. Bertho (1910) also mentions observations in the Seychelles and on the French islands of Amsterdam and Saint-Paul; however, the remoteness of the latter two islands, whose settlement at the time is not proven beyond doubt, lead us to question the veracity of the corresponding reports. Curiously this event is missing from the French national tsunami database (Lambert & Terrier, 2011).

As for Madagascar, historical tide gauge data available from Service hydrographique et océanographique de la Marine (SHOM) for Diego-Suarez (now Antsiranana; SHOM Shelfmark: SHD R 7JJ2384) have no entries for 3 – 6 January 1907 (I. Garnier-Loussaut, personal communication, 2014). Observations of high and low water levels are missing for January 1907 from Tamatave (now Toamasina), and diurnal visual observations made with a tide staff are only available for 15 February until 2 March 1907 from Baie du Courrier (Y. Ferret, personal communication, 2014). No reports could be found in the several French language newspapers published in Madagascar at the time.
443

444

Tsunami Modelling Methodology

445

446

447

448

449

450

451

452

We chose a subset of 8 locations with reliable geographic co-ordinates on the islands of Simeulue and Nias (Figure 6a, 6b), at which either run-up, wave heights, or inundation distances could be determined (Table S5). These sites extend approximately parallel to the Northern Sunda trench, along strike from Simeulue to Nias, for a distance of about 220 km. Their spatial distribution and their proximity to the Sunda trench (60 – 80 km), along with the effect of directivity (e.g. Miyoshi, 1955; Kajiura, 1972) then allow a causal relationship between field observations and the parameters of the rupture, with each site serving as a control point to estimate and constrain slip on a corresponding segment of a model of the parent seismic source.

453

454

455

456

457

458

459

460

Our tsunami simulations use the Cornell Multi-Grid Coupled Tsunami Model or COMCOT, a fully validated finite-difference algorithm solving the non-linear shallow water approximation of hydrodynamics (Liu et al., 1998), which has been applied to the investigation of both historical and modern tsunamis (e.g. Li et al., 2015; Wang & Liu., 2016). However, the extremely limited nature and quality of the available seismological data precludes the use of commonplace modern methods (e.g. Kikuchi & Kanamori, 1982, 1986, 1991) to constrain the parameters (fault length and width, seismic slip) necessary to compute a field of initial conditions for the simulation of the tsunami.

461

462

463

464

In a recent study, Ebel & Chambers (2016) have documented that seismicity occurring in the vicinity of past large ruptures in Eastern North America and California shows a tendency to concentrate its largest events at the edges of past rupture. They went on to show that this geographical trend can be used to map the rupture area of historical events predating the
development of modern (and especially digital) data necessary to conduct detailed source tomography. In this context, and assuming that Ebel & Chambers' (2016) paradigm can be extended to the case of large subduction interfaces, we use our preferred epicentral location, and catalogues of relocated seismicity before and after the 2004 Sumatra-Andaman and 2005 Nias earthquakes (Engdahl et al., 2007; Pesicek et al., 2010), to define a zone of presently low seismic productivity in the uppermost part of the plate contact, extending roughly 250 km between 94.7°N and 97°N. In this general area, we then envision a number of scenarios for seismic rupture along the shallow up-dip plate interface of the Sunda Trench, which we divide into sixteen elementary sub-faults, themselves regrouped into four segments labelled A to D, this methodology being detailed in the *Supplementary Material*. The sixteen patches, each measuring 30 km × 40 km, are distributed in updip-downdip pairs (**Figure 7a**), their exact geometry following Slab 1.0 (Hayes et al., 2012).

Given the slowness we have documented in the source of the 1907 earthquake, we use slip estimates departing from the classical scaling laws applicable to subduction events (Geller, 1976; Blaser et al., 2010) and consider models with variable rigidity (μ). We begin by assuming that the rupture was limited to the shallowest portion of the interface; conventional crustal material would feature a typical value μ = 3 × 10¹¹ dyn/cm², which using the moment M₀ = 2.5 × 10²⁸ dyn*cm obtained from mantle waves, would result in a slip (Δu) ≈ 7 m on each of the up-dip sub-faults. We then model a tsunami earthquake by varying the rigidity on each segment between 0.7 and 2.0 × 10¹¹ dyne/cm², which increases the slip to between 9 and 24 m (**Figure 7a; Figures S3–S7**) in 3 m increments. These values of μ fall in the range proposed in the context of tsunami earthquakes by a number of investigators (Ide et al., 1993; Satake, 1995; Bilek & Lay, 1999; Geist & Bilek,
487 2001; Hill et al., 2012; Li et al., 2015). For each of these models, we then compare the simulated
tsunami amplitudes to those estimated from documentary material (Figure 6).
488

489 The tsunami impacts documented on Simeulue (Figure 6a) were best replicated by more
490 than 21 m of slip on segment B of the rupture. For example, at Lakubang, the inundation distance
491 is thought to have approached 3 km (Yogaswara & Yulianto, 2006; Figure 7b). At this location,
492 despite slip greater than 21 m on segment B, the inundation remains underpredicted and we suspect
493 the inundation distance reported by Yogaswara & Yulianto (2006) is strongly controlled by local
494 drainage that is known to allow tsunamis to travel much farther inland following the channels of
495 natural and man-made drainage features (e.g. Mori et al., 2011). In the region of Nias to the south
496 (Figure 7c), we find that a slip of 15 m is required on Segment D to reproduce the wave heights
497 reported from Pulau Wunga (Figure 7c). Our constraints on the amount of slip on Segments A and
498 C (Figure 6) are weaker due to the lack of quantitative records from northern Simeulue and the
499 Banyak islands, respectively (Figure 2c). However, we infer that the noticeable tsunami impacts
500 on the west coast of Sumatra (Figure 2c) such as the flooding at Susoh (Bataviaasch Nieuwsblad,
501 12 February 1907), Tapaktuan (Anonymous, 1909), and possibly Kayu Menang (Aceh Post, 10
502 June 2013) are probably indicative of slip of 6 to 9 m on those segments.
503

503 In conclusion, our simulation experiments indicate that the values of run-up and inundation
504 estimated from the documentary material require seismic slips varying from 6 to 9 m on Segments
505 A and C, to as much as 15 m on Segment D, and probably more than 21 m on Segment B. These
506 values are greater, by factors ranging from 1.3 to 5, than expected under scaling laws (e.g. Geller,
507 1976) for an earthquake of M₀ = 2.5 × 10²⁸ dyn*cm. This provides an additional argument for the
508 anomalous properties of the 1907 event, and quantifies its character as a “tsunami earthquake,” as
509 defined in Kanamori's (1972) landmark study.
510
Our preferred rupture model satisfactorily predicts tsunami heights at all but one location i.e. Afulu on Nias (Figure 6b; 7b). We note that the modern village is located near a semi-circular bay (Telok Afulu) opening to the south and sheltered on the west by a small island called Hulo Uma (1.251° N, 97.236° E); further, a review of topographic data (see Data and Resources) indicates steep relief (≥ 20 m) close to the shore, and a drainage channel extending inland for ~1 km from the north-western shore of Telok Afulu. Such features can cause highly localised amplification of tsunami waves (e.g. Shimozono et al., 2014), while a small island can also amplify tsunami waves in its lee, a scenario known as the “Babi effect” after it was first observed and modelled at Pulau Babi (8.43° S; 122.51° E) during the 1992 Flores tsunami (Yeh et al., 1994; Briggs et al., 1995). Localised extreme inundation could also be due to an aerial or sub-aerial landslide (possibly triggered by the snappier Event II), as in the case at Riangkroko during the Flores tsunami (Tsuji et al., 1995; Plafker, 1997). We also cannot rule out the possibility of a separate, local tsunami caused by Event II, especially given the lack of arrival times in the accounts of the tsunami at Afulu.

**Figure 8** contours the maximum amplitudes simulated in deep water (i.e., not including the interaction of the wave with coastlines) in the far field, throughout the Indian Ocean basin. As expected, our results are generally similar with previous simulations for scenarios of large earthquakes at the Sumatra trench, as presented for example by Okal et al. (2009), the main pattern being the combination of a major lobe of directivity in the direction perpendicular to the strike of the trench (Ben-Menahem and Rosenman, 1972), and of the focusing effect of the shallow bathymetry along the Southwest Indian Ocean Ridge.

Owing to the limited nature of the available far-field reports, it is only possible to compare them qualitatively to our simulations. We first note that reports of the tsunami at Karachi and
Mumbai, on the Western coast of the Indian subcontinent, are generally compatible with the development of a probable edge wave hugging the broad continental shelf present in the area, while the river port sites at Kolkata and Yangon, with no tsunami reported on existing maregraphs, do appear relatively sheltered in our simulations. In the Southwest Indian Ocean, the Mascarene Islands (Seychelles, La Réunion, Rodrigues and Mauritius) all feature deep water amplitudes on the same order (20 to 30 cm) as simulated for the 2004 tsunami, the latter having run up on the islands to between 0.5 and 2.9 m (Okal et al., 2006a). On the other hand, the Mascarene Plateau, arcing from La Réunion to the Seychelles, clearly screens the Mascarene Basin in its lee, thus protecting the Northern half of Madagascar, and explaining the lack of tsunami reports from the maregraph at Diego-Suarez. This would also explain the lack of observations at Tamatave, which was already the main trading port of the island and whose activity was at the time well described in the local press. By contrast, stronger amplitudes, potentially comparable to those in the Mascarene Islands, would be expected on the southern part of Madagascar, the site of the highest run-up values in 2004 (Okal et al., 2006b); the absence or reports in the press may reflect poor communications and possibly an ongoing state of insurgency. Finally, regarding the islands of Amsterdam and Saint-Paul, deep water amplitudes in the range of 0.2 m and an individual simulated value reaching 0.9 m within 5 km of the coast of Saint Paul would give some veracity to the report of an observable wave (Bertho, 1910), but the presence of a settlement at the time on the island of Saint Paul (by any measure, even more remote than Amsterdam), remains unproven.

In conclusion, despite the largely qualitative nature of the available dataset of observations, simulations based on our preferred rupture model correctly predict the main features of the tsunami in the far field.

555
556 **Discussions and Conclusions**

557 The main conclusion of our study is the confirmation and quantification of the slowness of
558 the 1907 mainshock (Event I), which we achieved from a broad range of seismological
559 observations. By computing its seismic moment from the spectral amplitude of surface waves
560 measured at mantle periods (up to 170 s), we obtain a value significantly greater (by a factor of
561 about 4) than previously derived around 50 s by *Kanamori et al.* (2010). Although it must remain
562 qualitative in the absence of instrumental metadata, the comparison of amplitudes and duration of
563 body and surface waves at Manila (**Figure 5**) clearly documents the difference in source properties
564 between Events I and II, hence the slowness of the mainshock. The identification of Event II,
565 occurring only 53 minutes after the mainshock, clearly resolves the paradox of an event holding
566 many hallmarks of a “tsunami earthquake,” but being felt at surprisingly high intensities. We also
567 propose a model of seismic rupture which provides an acceptable fit to a new dataset of tsunami
568 run-up and inundation values from local and far-field locations in the Indian Ocean basin.

569 In the following sections, we provide a perspective on our most important results from a
570 regional and a global perspective.

571

572 *The value of non-instrumental archives*

573 Our research underscores the importance of meticulously collated and analysed non-
574 instrumental evidence that allows the study of historical earthquakes for which instrumental data
575 is lacking or severely limited. The discovery of Event II was an unexpected outcome borne out of
576 our careful scrutiny of macroseismic data that led us to look for seismological evidence to support
577 it. The absence of this event from earthquake catalogues (e.g. *Gutenberg & Richter, 1954*;
Storchak et al., 2013) is very conspicuous but not unusual, as with other recently discovered large early-instrumental earthquakes (Hough et al., 2005); it results from the limited number of seismograms at hand, and the inadequacy (or simple inexistence) of proper algorithms for association of reported phase times, not to mention the adjustment of local times before implementation of standard time zones. Previously untapped historical and scientific sources located by us have also formed the basis of early 20th century descriptions of the disaster (e.g. de Meyier, 1907; Rappard, 1909; van der Linde, 1920). In stark contrast with other natural disasters in the Dutch East Indies such as the 1883 eruption of Krakatau for which detailed official reports were written (e.g. Veerbeck, 1886), locating an official report for the 1907 disaster was futile despite exhaustive efforts on our part. Szirtes (1912a, p. 5) carries cursory mention of a "detailed study" (de: eingehende Untersuchung) by T.H. Staverman, including possibly a study of its epicentral location, but without a complete citation. The scientific bibliography relating to geology in the Dutch East Indies between 1907 and 1912 (e.g., Verbeek, 1912) has no record of this document, and its whereabouts remain a mystery, including at Strasbourg where Siegmund Szirtes was based (L. Rivera, personal communication, 2017).

Written descriptions of historical earthquakes along the Sunda-Java trench (Newcomb & McCann, 1987) are at best fragmentary or incomplete. We speculate that the 16th century disaster on the south coast of Java relating to the Javanese and Sundanese legend of Ratu Kidul and recently interpreted as a tsunami based on surviving indigenous Babads and written Dutch sources (Reid, 2016), could have been another tsunami earthquake, owing to the lack of reported shaking effects preserved in these materials. However, we emphasise that the brief nature of these documents warrants further work to confirm our assertions as these could also be related to traditional
600 earthquakes, seismically or gravitationally triggered turbidity currents, or meteorological
601 phenomena.

602

603 *The 1907 epicentre in the local tectonic context*

604 Our preferred location of the rupture from Event I is trench ward, i.e. up-dip, in a north-
605 east facing re-entrant on the Sunda Megathrust (Franke et al., 2008) between the ruptures of the
606 2004 (Chlieh et al., 2007) and 2005 (Konca et al., 2008) earthquakes. This is in agreement with
607 our conclusion that Event I was a tsunami earthquake, and the documented observations of the
608 near-field tsunami. Near the western terminus of our rupture, seismic reflection surveys (Franke
609 et al., 2008) and 3-D active source tomography (Tang et al., 2013) have identified a ridge of
610 oceanic basement (Figure 1b) that projects into the Wharton Basin, coinciding with a mapped
611 fracture zone (Singh et al., 2011; Jacob et al., 2014). Onshore on Simeulue, this coincides with the
612 Simeulue Saddle (Sieh et al., 2006) which has served as a persistent barrier to rupture in the past
613 (Meltzner et al., 2012). Another feature to the south in the vicinity of the Batu Islands (Natawidjaja
614 et al., 2006; Figure 1b) represents another barrier to rupture (Philibosian et al., 2014) and
615 demarcates an important segment boundary on the Sunda megathrust near the Batu Islands
616 (Meltzner et al., 2015). Further down dip and beneath the Simeulue Saddle (Sieh et al., 2006), this
617 morphological high can be associated with a region of strong coupling beneath central Simeulue
618 (Tsang et al., 2015), and has been inferred as a structural control for local seismicity and the rupture
619 dimensions of modern earthquakes under central Simeulue (Morgan et al., 2017). The dimensions
620 of this morphological feature are debated (Franke et al., 2008; Tang et al., 2013; also see Figure
621 4b in Morgan et al., 2017). Although our best-fit epicentral location lies slightly to the west of this
feature, its uncertainty ellipse overlaps the western ramp of this morphological high. Incidentally, the largest inferred slip in 1907 on the interface locates in the region of this structural feature, a pattern very similar to that of other tsunami earthquakes, that associate both event nucleation and the largest slip with the subduction of subsurface topography along strongly coupled interfaces (e.g. Tanioka et al., 1997; Abercrombie et al., 2001; Newman et al., 2011b; Bell et al., 2014); this remark is also supported by the jagged nature of the Omori waveform at Manila discussed previously. The weaker tsunami amplitudes on southern Nias suggests that the 1907 rupture did not extend to the Batu Islands segment barrier or terminated in the neighbourhood of this southern segment boundary near the likely rupture of a deadly tsunamigenic earthquake near the island of Simuk (Figure 1b) to the South of Nias, on 9 March 1861 (Reiche, 1863). Since our rupture is only constrained by tsunami observations, further research would also be required to refine the western limit of the 1907 rupture in the region of the morphological high identified by Franke et al. (2008) and Tang et al. (2013), in order to determine whether the rupture stopped at, or propagated through this segment boundary.

The value of maximum slip in our rupture model is obtained in the model of a simple planar rupture, but we also recognise the possibility of a complex rupture involving splay faults within the overlying accretionary wedge. Seismic reflection profiles indicate the presence of frontal folds and pop-up structures south (Franke et al., 2008) and west (Singh et al., 2008) of Simeulue. The relation between splay faults and unexpectedly large sea floor displacements has been numerically proven (e.g. Wendt et al., 2009) and proposed for other tsunami earthquakes (e.g. Fan et al., 2017). However, given the uncertainties in the near field associated with both quantifiable numeric observations, i.e., run-up and inundation distances, as well as the largely unknown local arrival times of the tsunami, we restricted this study to a simple planar model.
645

646 *The lack of geodetic signals*

647 Our study has resolved many unaddressed aspects of the 1907 Sumatra earthquake
648 sequence but leaves unanswered the question of whether any subsidence was associated with this
649 sequence. In this context, we note Hodgson's (1934) remark that "the south coast was partially
650 submerged by an earthquake," that could be interpreted as coseismic subsidence (Meltzner et al.,
651 2010; 2015) or as tsunami inundation under the semantics of the 1930's. Reports from the southern
652 coast of Simeulue are limited and none of the available Dutch accounts mention land level changes
653 associated with the 1907 earthquake, except in brief descriptions from Simeulue Cut (Haagsche
654 Courant, 14 February 1907; **Figure 6a**) and the Hinako Islands (Bataviaasch Nieuwsblad, 22
655 March 1907; **Figure 6b**). We modelled subsidence (**Figures S8 – S9**) of ≤ 50 cm on Simeulue Cut
656 and at the south coast of the island of Simeulue due to a dislocation in an elastic half space (Okada,
657 1985), and calculated even smaller values of subsidence (≤ 15 cm) in the Hinako Islands. The
658 accounts at our disposal do not state whether the writer visited the island of Simeulue Cut, viewed
659 it from a distance, interviewed direct witnesses, or simply derived this information second-hand.
660 If taken at face value, the description from Simeulue Cut might be inferred as coseismic subsidence
661 or extreme tsunami erosion. The other perplexing detail comes from the Hinako islands (**Figure
662 6b**) where the seabed was reportedly exposed (uplifted?) for up to 2 hours (Bataviaasch
663 Nieuwsblad, 22 March 1907). Coral microatolls have been well studied on Simeulue and have
664 contributed to our understanding of the rupture extents of great medieval earthquakes (Meltzner et
665 al., 2015) as they are natural long-term indicators of changes in relative sea-level, but these are
666 often subject to preservation issues that can undermine temporal completeness (Meltzner et al.,
667 2012). In this regard, and pertinent to our study, Meltzner et al. (2012) investigated a coral
microatoll at Ujung Lambajo (LBJ; **Figure 1b**) at a distance of ~6 km northeast of Simeulue Cut but found that the paleo-record at this site only extended to 1955. Much further south, *Meltzner et al.* (2015) record the death of a microatoll (LAG-3B; **Figure 1b**) near Lagundri on Nias and suggest this could either be related to the 1907 earthquake or had “nothing to do with tectonics or relative sea level (RSL) change.” Unlike the locations of many tsunami earthquakes that are devoid of offshore islands, Simeulue and Nias lie within 100 km of the trench axis and therefore increase the likelihood that the coseismic signature of land level changes from tsunami earthquakes could be recorded in the paleo-record of microatolls. It is also entirely plausible that the inferred land level changes on Simeulue Cut and in the Hinako Islands are indicative of a rupture that extended further downdip than in our first-order model. Also, they could have been, both or independently, associated with the aftershock, or with post-seismic after-slip; the alleged Hinako uplift (*Bataviaasch Nieuwsblad*, 22 March 1907) could have occurred concurrently with the death of a coral microatoll (LAG-3B) previously documented by *Meltzner et al.* (2015), and thus, could possibly represent coseismic uplift from Event II. Bearing in mind the uncertainties in all of the above, we do not incorporate these in our rupture model but suggest that further work is required in this unique physiographic environment to uncover the paleo-record of previous tsunami earthquakes near Simeulue.

685

686

## *The 1907 Tsunami Earthquake from a Global Perspective*

Owing to its size and location, the identification of the 1907 North Sumatra event as a tsunami earthquake has a number of implications which must be discussed in a global perspective.

689

**Table 1** compiles characteristics of 21 known such events, including estimates of seismic moments
and, when available, of the slowness parameter Θ introduced by Newman & Okal (1998). At M₀ = 2.5 × 10²⁸ dyn*cm, the 1907 earthquake ranks as the third largest documented “tsunami earthquake,” after the 2004 Sumatra-Andaman earthquake (which can be considered a tsunami earthquake, in view of the clear slowness in its source (Choy & Boatwright, 2007; Okal, 2011)), and the 1946 Aleutian Islands earthquake. Note that the list in **Table 1** may not be exhaustive since the origin of some historical destructive waves remains debated, such as the 16th century tsunami associated with the legend of Ratu Kidul on the south coast of Java predating Dutch colonization (Reid, 2016), and the more recent tsunami of 26 February 1902 in El Salvador (Cruz & Wyss, 1983; Ambraseys & Adams, 1996). Some evidence would also characterise the 1700 Cascadia megathrust event as a tsunami earthquake (Obermeier & Dickenson, 2000; Okal, 2011). As they are felt at deceptively low levels not conducive to the perception of an appropriate level of threat, tsunami earthquakes form a particularly vicious kind of a natural hazard, and their mitigation constitutes one of the biggest challenges facing today's tsunami community. Paramount among remaining questions of primary importance is the relationship between PTEs and ATEs as defined by Okal & Saloor (2017) with traditional megathrust earthquake namely: (i) Does the distribution of PTEs bear a regional signal? (ii) Does the occurrence of PTEs along a subduction zone rule out regular (and possibly larger) megathrust earthquakes? (iii) Can a subduction zone be considered immune to tsunami earthquakes, either PTEs or ATEs?”

A systematic study by Okal & Newman (2001) in the regions of the three 1990s PTEs (Nicaragua, 1992; Java, 1994; and northern Peru, 1996) found no systematic trend for slowness in background seismicity but identified an older event in northern Peru in 1960 as a PTE, 340 km from the 1996 Chimbote source, thus suggesting a lateral coherence in the up-dip properties of the interface governing the occurrence of PTEs. Following their study, the 2006 Java earthquake
provided a confirmation of this trend, occurring 580 km from the 1994 epicentre. Similarly, the 2012 El Salvador PTE occurred only 175 km from the site of the 1992 Nicaragua earthquake; should the 1902 event be confirmed as a PTE (Cruz & Wyss, 1983; Ambraseys & Adams, 1996), then it would earmark a 350 km section of the Central American subduction zone as prone to PTEs. In this context, the 1907 event provides some limited new insight. It was indeed followed, 98 years later, by a slow megathrust earthquake, but the origin of their slow character may not be comparable: based on its location near the trench, the 1907 event's slowness was rooted in its shallow source, whereas in 2004, it emanated from a relatively slow and jagged propagation along its exceptionally long fault length (Ishii et al., 2005; de Groot-Hedlin, 2005; Tolstoy & Bohnenstiehl, 2005). More comparable are the events of 1907 and 2010, which both took place along the shallowest, probably structurally similar, portions of the plate interface, and were separated laterally by 900 km. However, the 2010 Mentawai earthquake was clearly an ATE of the 2007 Bengkulu megathrust event, while the 1907 earthquake was Primary: there is no concrete evidence anywhere in the historical record for an immediately preceding mainshock which would have been significantly larger (Newcomb & McCann, 1987).

Whether or not the occurrence of PTEs along a subduction zone rule out regular (and possibly larger) megathrust earthquakes is particularly critical along the south coast of Java where two PTEs featuring very similar source characteristics took place in 1994 and 2006, and where no truly gigantic earthquake is known in the admittedly short historical record (Newcomb & McCann, 1987). Our assertion would also apply to the Hikurangi subduction zone, off the North Island of New Zealand, where the historical record is even shorter and where the only large inter-plate thrust earthquakes since 1917 are the 1947 tsunami earthquakes (Doser & Webb, 2003; Okal & Saloor, 2017) although Cochran et al. (2006) inferred prehistoric tsunamis from an analysis of wetland
sediment cores. Our assertion would also extend conceivably to Cascadia if indeed the 1700 event
had a slow source. In this context, the proximity of the 1907 and 2005 epicentres would suggest a
negative answer, namely, that PTEs and regular megathrust events may occur along neighbouring
subduction segments (or even along the same one at different depths along the interface). Note that
this conclusion might be supported by the case of the 1946 Aleutian PTE, in comparison to the
nearby 1938 and 1957 megathrust earthquakes which did not exhibit source slowness, the former
even generating a deceptively small tsunami (Stover & Coffman, 1993); however, this analogy
remains tentative in the presence of the Shumagin gap between the 1946 and 1938 ruptures, where
no large events are known since at least 1787 (Davies et al., 1981), which leaves the question of
its potential for a PTE wide open until the next large earthquake fills it.

Tsunami earthquakes have not (or not yet) been documented in a number of subduction
zones having produced large tsunamigenic earthquakes during the era of instrumental seismology
(e.g., Southern Peru, Northern Chile, Makran), or during pre-instrumental, historical times (e.g.,
Hellenic Arc, Ryukyu). A small-scale study in Vanuatu by Okal & Saloor (2017) documents a
lateral, along-strike, variation of the slowness parameter Θ, in conjunction with the subduction of
submarine structures, on a scale of only a few hundred km. The underestimation of the magnitudes
of great earthquakes prior to instrumentation (Hough, 2013), the largely random nature of the
fragmentation of rupture at subduction zones (Ando, 1975), and the gross temporal under sampling
of tectonic cycles by modern seismology which resulted in unexpectedly large events (2004
Sumatra, 2011 Tohoku), has led to the demise of otherwise promising paradigms (Stein & Okal,
2007; McCaffrey, 2007), and thus the only prudent answer to this question must remain negative.
The 1907 earthquake provides no new insight, however, on whether any subduction zone can be
considered immune to tsunami earthquakes, either PTEs or ATEs.
759

760

Hazard Mitigation from Future Tsunami Earthquakes

761

762

763

764

765

766

In the absence of definitive evidence of aseismic slip, tremor, or slow-slip earthquakes up-dip (Feng et al., 2015; Tsang et al., 2015), and based on newly modelled estimates of convergence of 46 mm yr⁻¹ at 1° N, 96° E (Bradley et al., 2017), we believe the rupture patch identified in this study could host a future Mw ≥ 7.7 tsunami earthquake, up-dip and trench ward between the rupture patches of the 2004 and 2005 earthquake with dimensions and physical properties similar to our model (Figure 7a), and with lower than anticipated ground motions to elicit self-evacuation.

767

768

769

770

771

772

773

774

775

776

777

778

779

We recall that on Simeulue, the local legend of the *s'mong* associates strong earthquakes with tsunami hazard and elicits self-evacuation to high ground upon feeling strong shaking (McAdoo et al., 2006; Baumwoll, 2008; Syafwina, 2014), which turned out particularly effective in preventing heavy loss of life on Simeulue during the 2004 Sumatra-Andaman tsunami (McAdoo et al., 2006). It has been proposed (e.g. Syafwina, 2004) that the *s'mong* is based on memories from 1907 preserved in verse or in stories passed down by older generations (we note, however, that *Yogaswara & Yulianto* (2006) record versions of the legend, both with and without severe shaking). If so, the exact role played, in the *s'mong* legend, by the weak shaking due to Event I, and the presumably stronger one from Event II, remains unclear; at any rate, shaking from Event II, estimated to lag Event I by ~53 minutes, may not have occurred before the tsunami attacked Simeulue; this is further suggested by E. Schröder's account of the second shock felt on Nias (located farther from Event I than Simeulue) *after* flooding by the tsunami (*Bataviaasch Nieuwsblad*, 22 March 1907).
While the *s'mong* tradition may have saved many lives in 2004 and 2005, it is clear that is not adapted to the particular challenge of tsunami earthquakes, as illustrated by the case of the 2010 Mentawai event, 600 km Southeast of Simeulue (**Figure 3**), a typical ATE following the 2007 Bengkulu megathrust earthquake. As detailed by *Hill et al.* (2012), the earthquake was felt only "gently" or even not felt at all at some locations on the Pagai islands. Yet, the tsunami attained run-up heights of 17 m and killed more than 500 people. In this case, local villagers did not immediately evacuate as the shaking was weaker than during the seismologically much larger 2007 Bengkulu mainshock (*Borrero et al.*, 2009) and during the 2009 intraplate Padang earthquake (which occurred deeper in the slab, and under the island of Sumatra (*McCloskey et al.*, 2010), both of which did not produce significant tsunamis in the Mentawai islands.

In this context, some interviewed witnesses reported that the 2010 shaking on the Mentawai Islands was both weak and long, lasting as much as "several minutes" (*Hill et al.*, 2012). This is an important observation, since source slowness is expected to both decrease high-frequency ground motion (and hence the level of shaking) and increase the duration of seismic wavetrains; the comparison of these two properties has been earmarked by *Convers & Newman* (2013) and *Okal* (2013) as a means of identifying tsunami earthquakes from the analysis of first-arriving *P* waveshapes. Indeed, during the 2010 Mentawai event, *Newman et al.* (2011a) identified the slow character of the event only 17 minutes after the origin time using a prototype of their method, but unfortunately there existed no means to relay this information in due time to the populations at risk.

Accordingly, we end by re-iterating the recommendation made by *Hill et al.* (2012, p. 19) to include the duration of felt shaking as a warning for spontaneous evacuation, in a sense to implement a variant of *Convers & Newman's* (2013) algorithm based on human perception. This
constitutes a formidable challenge, since our experience gathered from a large number of post-tsunami surveys (e.g. Synolakis & Okal, 2005) is that the perception of relatively short time durations (seconds, minutes) by lay individuals in a situation of emergency and panic is even less accurate than that of distance.

807

## Data and Resources

We present a list of all newspaper titles and dates consulted including libraries and online databases (free and subscription based) where these are currently available (Appendix A) with transcriptions of selected newspaper reports in Appendix B. Figures were created using freely available Generic Mapping Tools (Wessel and Smith, 1998) and QGIS software (http://download.qgis.org/). Bathymetry was digitised for Nias and Simeulue from AMS T503 Series maps (1:250,000; sheets NA 47-5, NA 47-8, NA 47-9, NA 47-13, NA 47-14, NA 46-8) available via the Perry-Castañeda Map Collection at the University of Texas at Austin (http://www.lib.utexas.edu/maps/ams/indonesia/, last accessed July 2017). We also used the General Bathymetric Chart of the Oceans (GEBCO) digital bathymetry dataset (http://www.gebco.net/, last accessed in October 2014). Digital topographic contours (1:5000) are available from Badan Informasi Geospasial (http://tanahair.indonesia.go.id; last accessed 27 November 2017). The topographic data used in the tsunami simulations is bare-ground Shuttle Radar Topography Mission (SRTM) data downloaded from https://data.bris.ac.uk/data/dataset/10tv0p32gizt01nh9edcjzd6wa (last accessed, April 2018). For some key locations including Lakubang, Lukon, and Simeulue Cut, we replace the SRTM data with 4-m resolution digital terrain data purchased from the Nippon Telegraph and Telephone
825 (NTT) Data Corporation. Topographic maps (1:50,000) prepared by the 653rd Topographic
826 Engineering Battalion of the United States Air Force for Simeulue (Series HIND 605 Sheet 5;
827 NLA shelf mark: MAP G8082.S5 s50) and by the British War Office for Nias (Series HIND 614
828 Sheet 9; NLA shelf mark: MAP G8082.N5 s50) were procured from the National Library of
829 Australia (NLA).

830

831 Acknowledgements

832 This article has benefitted from interactions with Alicia Schriker, Andrew Reid, Aron
833 Meltzner, Christina Widiwijayanti, Lauriane Chardot, Hiroo Kanamori, Priyamveda Nanjundiah,
834 Qiang Qiu, Euan Smith, and Shengji Wei. We would like to thank H.P. Shukla, Isabelle Garnier-
835 Loussaut, Susanne Sargeant, Vineet Gahalaut, and Yann Ferret for assistance rendered in our
836 efforts to locate additional seismograms and station bulletins for the earthquake, and maregrams.
837 Hiroo Kanamori provided Japanese seismograms from Hongo and Osaka. Robert Engdahl
838 provided relocated seismicity from *Engdahl et al.* (2007) and Alexandre Sahal provided French
839 newspaper reports from La Réunion. Çağıl Karakaş, Iwan Hermawan, Koen van Noten, and Mari
840 Hamahashi assisted with, or verified selected translations. Aron Meltzner provided field notes and
841 photos from the 2005–2007 field season. We also acknowledge the services rendered by the staff
842 at libraries in The Hague (*Koninklijke Bibliotheek*), Berkeley (*National Information Service for
843 Earthquake Engineering, University of California, Berkeley*), Canberra (*National Library of
844 Australia*), Jakarta (*UNESCO*), London (*The British Library*), Los Angeles (*University of
845 California, Los Angeles*), Paris (*Bibliothèque nationale de France*), and Singapore (*National
846 Library of Singapore*). This study, including archival work in The Hague by SSM, was supported
847 by the National Research Foundation Singapore and the Singapore Ministry of Education under
848 the Research Centres of Excellence initiative. EO was partially supported at Northwestern
849 University by the National Science Foundation, under subcontract from the University of
850 Pittsburgh's Hazards SEES Grant Number OCE-1331463. This paper is a contribution to IGCP
851 Project 639 Sea-Level Changes from Minutes to Millennia. This work comprises Earth
852 Observatory of Singapore contribution number XX.

853

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1330 Wiens, D.A., D.C. DeMets, R.G. Gordon, S. Stein, D.F. Argus, J.F. Engeln, P.R. Lundgren, D. Quible, C.A. Stein, S.A. Weinstein, and D.F. Woods (1985), A diffuse plate boundary model for Indian Ocean tectonics, *Geophysical Research Letters*, **12**, pp. 429 – 432.

1333 Wysession, M.E., E.A. Okal, and K.L. Miller (1991), Intraplate seismicity of the Pacific Basin, 1913 – 1988, *Pure and Applied Geophysics*, **135**, pp. 261 – 359.

1335 Yogaswara, H. and E. Yulianto (2006), Pengetahuan local tentang tsunami pada masyarakat Simeulue, Nanggroe Aceh Darussalam, Assessing and recognising community preparedness in natural disasters in Indonesia, LIPI – UNESCO/ISDR, 69 pp.
1338 **Table 1:** List of Documented Tsunami Earthquakes since 1896.

1339 * PTE: Primary Tsunami Earthquake; ATE: Aftershock Tsunami Earthquake (Okal & Saloor, 2017); L: Tsunami enhanced by ancillary landslide.

<table><thead><tr><th>#</th><th>Date<br>D M Y</th><th>Regions</th><th>Type<br>*</th><th>Moment<br>(10<sup>27</sup> dyn*cm)</th><th>Θ</th><th>Notes</th></tr></thead><tbody><tr><td>01</td><td>15 June 1896</td><td>Meiji Sanriku, Japan †</td><td>PTE</td><td>≈12</td><td>-NA-</td><td>Tanioka & Satake (1996)</td></tr><tr><td><b>02</b></td><td><b>04 January 1907</b></td><td><b>Northern Sumatra</b></td><td><b>PTE</b></td><td><b>≈25</b></td><td><b>-6.5 ?</b></td><td><b>This study</b></td></tr><tr><td>03</td><td>13 April 1923</td><td>Kamchatka</td><td>ATE (L)</td><td>1.2</td><td></td><td>Salaree & Okal (2018)</td></tr><tr><td>04</td><td>23 June 1932</td><td>Manzanillo, Mexico</td><td>ATE</td><td>5.2</td><td>-6.18</td><td>Okal & Borrero (2011)</td></tr><tr><td>05</td><td>21 July 1934</td><td>Santa Cruz Islands</td><td>ATE</td><td>10</td><td>-6.10</td><td>Okal & Saloor (2017)</td></tr><tr><td>06</td><td>01 April 1946</td><td>Aleutian Islands †</td><td>PTE (L)</td><td>85</td><td>-7.03</td><td>López & Okal (2007)</td></tr><tr><td>07</td><td>25 March 1947</td><td>Gisborne, New Zealand</td><td>PTE</td><td>4</td><td>-5.94</td><td>Okal & Saloor (2017)</td></tr><tr><td>08</td><td>17 May 1947</td><td>Gisborne, New Zealand</td><td>‡</td><td>3</td><td>-6.51</td><td>Okal & Saloor (2017)</td></tr><tr><td>09</td><td>20 November 1960</td><td>Northern Peru</td><td>PTE</td><td>2.7</td><td>-6.13</td><td>Okal & Newman (2001)</td></tr><tr><td>10</td><td>20 October 1963</td><td>Kuril Islands</td><td>ATE</td><td>7.5</td><td>-6.42</td><td>Fukao (1979)</td></tr><tr><td>11</td><td>13 August 1965</td><td>Vanuatu</td><td>ATE</td><td>2</td><td>-5.88</td><td>Okal & Saloor (2017)</td></tr><tr><td>12</td><td>10 June 1975</td><td>Nemuro-Oki, Japan</td><td>ATE</td><td>0.8</td><td>-6.43</td><td>Fukao (1979)</td></tr><tr><td>13</td><td>19 December 1982</td><td>Tonga</td><td>PTE</td><td>2.0</td><td>-5.76</td><td>Newman & Okal (1998)</td></tr><tr><td>14</td><td>02 September 1992</td><td>Masachapa, Nicaragua</td><td>PTE</td><td>3.4</td><td>-6.30</td><td>Newman & Okal (1998)</td></tr></tbody></table>
<table><tr><td>15</td><td>02 June 1994</td><td>South of Central Java</td><td>PTE</td><td>5.3</td><td>-6.01</td><td>Newman & Okal (1998)</td></tr><tr><td>16</td><td>21 February 1996</td><td>Chimbote, Peru</td><td>PTE</td><td>2.2</td><td>-5.94</td><td>Newman & Okal (1998)</td></tr><tr><td>17</td><td>26 December 2004</td><td>Sumatra-Andaman</td><td>PTE</td><td>1200</td><td>-6.40</td><td>Stein & Okal (2007); Choy & Boatwright (2007)</td></tr><tr><td>18</td><td>17 July 2006</td><td>South of East Java</td><td>PTE</td><td>4.6</td><td>-6.13</td><td>Saloor & Okal (2018)</td></tr><tr><td>19</td><td>25 October 2010</td><td>Mentawai Islands</td><td>ATE</td><td>6.8</td><td>-6.22</td><td>Saloor & Okal (2018)</td></tr><tr><td>20</td><td>27 August 2012</td><td>El Salvador</td><td>PTE</td><td>1.3</td><td>-6.42</td><td>Saloor & Okal (2018)</td></tr><tr><td>21</td><td>06 February 2013</td><td>Santa Cruz Islands</td><td>PTE</td><td>9.4</td><td>-5.94</td><td>Okal & Saloor (2017)</td></tr></table>

† Charter events from *Kanamori* (1972); ‡ While #08 is smaller than its predecessor (#07), the moment ratio (3/4) is too close to qualify the second event as an aftershock of #07.
<figure><img src="image_2.png" /></figure>

1342

1343

**Figure 1:** Generalized tectonic map of the Sunda megathrust in Indonesia. First-order rupture of the 1907 earthquake (this study) indicated by a filled red polygon. Rupture areas for the 1797, 1833, 1861, 1994, 2004, 2006, 2007, and 2010 earthquakes are also shown (Bilek & Engdahl, 2007; Chlieh et al., 2007; Konca et al., 2006; Bilek & Engdahl, 2007; Konca et al., 2008; Hill et al., 2010). Location of the possible historical Ratu Kidul tsunami earthquake discussed in the text is labelled and appended by question marks. Stars represent selected recent earthquakes. Inset box shows the islands of Simeulue and Nias with the first-order rupture of the 1907 earthquake (this study) indicated by a filled grey polygon, red star representing the epicenter of the 1907 earthquake (this study), and an orange triangle indicating the approximate location of Event II. Outlines of the 2004 (dotted-dashed) and 2005 (dotted) lines are also displayed. Location of coral microatolls LBJ and LAG-3B from Meltzner et al. (2015) are also shown.

1355
<figure><img src="image_3.png" /><figcaption>**Figure 2:** Intensity map (A) from Newcomb & McCann (1987) in comparison with intensities determined by our study for the mainshock at 05:19 GMT (B), and the largest aftershock at ~06:12 GMT (C). White boxes marked with an “F” indicate that an earthquake was felt but macroseismic data were insufficient to assign an intensity.</figcaption></figure>

1356

1357

1358

1359

1360

1361

1362

1363

1364

1365
<figure><img src="image_4.png" /></figure>

1366

**Figure 3:** Relocations of the 1907 Sumatra event. The red dot shows our solution, with associated Monte Carlo ellipse. The green inverted triangle is the epicenter from *Kanamori et al.* (2010), with the confidence ellipse described in their Appendix shown in green. The brown upward triangle is *Gutenberg and Richter's* (1954) epicenter, and the orange square T. Staveman's (*Szirtes*, 1912). The blue diamonds (with respective confidence ellipses) are the ISC solution by *Storchak et al.* (2013; solid, bright), and from the ISC-GEM (5.0) catalog (outlined, light blue); note that the ellipses do not intersect, suggesting that they are deceptively small. The yellow diamond is the
1374 best estimate from a grid search (*D. Di Giacomo*, pers. comm., 2017) and the preferred location is shown as the purple bull's eye symbol. The light brown triangle is *Turner et al.'s* (1912) location.

1375
1376 The black stars show the epicenters of the mega events of 2004 (Sumatra-Andaman, "S-A"), 2005 (Nias, "N") and 2007 (Bengkulu, "B"), and the yellow star ("M") the 2010 Mentawai tsunami earthquake. Background seismicity (1970–2015; h < 100 km; M ≥ 5) shown as light grey dots.

1378
1379 Isobaths are at 500–m intervals (green at 500 m; dark blue at 4000 m and deeper).

1380

1381

1382

1383

1384

1385

1386

1387

1388

1389
<figure><img src="image_5.png" /><figcaption>**Figure 4:** Spectral amplitudes of Rayleigh and Love waves at Göttingen, interpreted as corrected mantle magnitudes *M*<sub>c</sub> (Okal & Talandier, 1989), computed in the geometry φ = 320°, δ = 15°; λ = 90°, in the frequency band 6–10 mHz. The black dashed lines and yellow band show the average value and standard deviation of the full dataset, while the purple dashed line shows its best linear regression. Note the strong increase of moment with period.</figcaption></figure>
MAN

VICENTINI NNW-SSE

T = 1.2 s

04 JAN 1907

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

MAN

OMORI ENE-WSW

T = 6.4 s

04 JAN 1907

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

1397

1398 **Figure 5:** Seismic recordings at Manila (distance 29°; back-azimuth 247°), reproduced from *Maso* (1907), clearly showing the two events, separated by ~53 minutes. (a) Short-period Vicentini seismograph; (b) Omori seismograph. In both instances, the records have been offset to align corresponding phases vertically. Time marks are at 30-s intervals, with original times given as local p.m. times (GMT +8). Standard times (GMT) are overprinted in dark blue. These records illustrate dramatically the different characteristics of the two events.

1404
<figure><img src="image_8.png" /><figcaption>**Figure 6:** Tsunami observations on Simeulue (A), Nias (B), northern Sumatra (C), and in the Indian Ocean basin (D) including Sri Lanka (E) and La Réunion (F). White circles show 76 locations where the tsunami was observed, and black circles indicate sites used to forward model the tsunami. Red circles mark questionable or false records. White triangles show tide gauges where readings were available, and red triangles show tide gauges in operation in 1907 where data was either unavailable, or the tsunami was unrecorded, or records were incomplete.</figcaption></figure>

1405

1406

1407

1408

1409

1410

1411

1412
<figure><img src="image_9.png" /><figcaption>**Figure 7:** Hypothetical rupture geometry and slip parameters for the 1907 mainshock (a) estimated in this study (7A) off Simeulue (5A) and Nias (5B). Green triangles represent locations where the tsunami was reported (Table S5) and red triangles represent locations used in the tsunami model.</figcaption></figure>

1413

1414

1415

1416

1417
04 JAN 1907

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

**Figure 8:** Maximum tsunami amplitude simulated in deep water using our preferred model of rupture for the 1907 mainshock, whose epicenter is shown as the large star. Locations with tsunami effects either instrumentally recorded, or reported are shown as black bull's eye symbols (Ka:
1423 Karachi; Mu: Mumbai; Ch: Chennai; SL: Sri Lanka; PB: Port Blair; Ro: Rodrigues; Se: Seychelles;

1424 Ma: Mauritius; Re: La Réunion). Instrumented locations not reporting the event are shown as open

1425 triangles (Ad: Aden; Ko: Kolkata; Ya: Yangon; An: Antsiranana/Diego Suarez). The port of

1426 Toamasina (Tamatave), Madagascar (To), discussed in the text, is shown as an inverted triangle.

1427 Amsterdam (Am) and Saint-Paul (SP) islands, sites of questionable reports, are shown as grey

1428 centered squares.