Keywords

12.1 Introduction

Penang Island (Figs. 12.1 and 12.2) is one of the fastest developing regions of Malaysia, with growth especially in tourism and hi-tech manufacturing sectors (Lateh et al. 2011). With increasing development and concomitant population growth has come the increased risk of natural hazards, resulting in a renewed focus on the recognition and mitigation of such hazards on Penang. For example, there have been recent studies on the effects of seismicity on built structures (Fadzli 2007), on tsunami hazard on the west (Colbourne 2005) and north (Jahromi 2009) coasts, and on landslide hazards in Penang (Ahmad et al. 2006; Lee and Pradhan 2006; Pradhan and Saro 2010; Lateh et al. 2011). The Batu Feringgi area on the NW coast of Penang has one of the highest concentrations of tourist beach hotels and related infrastructure on the island, and is thus critical to the tourist economy of Penang and Malaysia.

Fig. 12.1
figure 1

The location of Penang Island, Malaysia (enclosed in oval), in relation to the stable Sunda Shelf, and the tectonically active Sumatra-Andaman subduction zone. The yellow star represents the epicentre of the great 26 December 2004 Sumatran earthquake, which then propagated northwards for 900 km along the Sumatra-Andaman fault (Lay et al. 2005; Neetu et al. 2005), and generated the great Indian Ocean tsunami of 26 December 2004

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Fig. 12.2
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Google Earth satellite view of Penang Island, with Batu Feringgi located at its northernmost end

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Batu Feringgi beach was affected, like most of the coastline of the entire region, by the Great Indian Ocean tsunami generated by the great Sumatra-Andaman earthquake of 26th December 2004 (Lay et al. 2005, Fig. 12.3; Roy and Ismail 2006; Roy et al. 2007). However, unlike areas further north in Thailand, which were severely affected by the tsunami (Goto et al. 2010a), the Batu Feringgi area suffered little infrastructural damage, even though the area was inundated by large waves up to 3.3 m high (Roy and Ismail 2006; Roy et al. 2007) which flooded beach hotel grounds, and about a dozen or more people were reported as having been killed. The author discovered large imbricated granite boulders on the NE end of the tourist beach in the Batu Feringgi area. Since seaward-dipping imbricated boulders have been used as evidence for tsunami activity in other areas (Bryant et al. 1992; Young et al. 1996; Master 2011) (Fig. 12.4), it was decided to investigate the Batu Feringgi imbricated boulders further, and to study their relevance to tsunami hazards in NW Penang. In particular, it was important to know if such imbricate boulders were the result of the 26th December 2004 tsunami, or earlier tsunamis, or some other process, such as storm activity, which has recently been proposed as a source of much boulder movement along coastlines (Williams and Hall 2004; Switzer and Burston 2010; Etienne and Paris 2010; Cox et al. 2012), since both tsunamis and storms can deposit imbricated boulders.

Fig. 12.3
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Tsunami propagation model at time of 2.0 h after the great Sumatra-Andaman earthquake of 26 December 2004 (From Lay et al. 2005). The red colours are peaks and blue colours are troughs in the tsunami waves. Note the location of Penang Island (circle with P), in the Sunda Shelf, where much smaller tsunami waves were experienced, compared to the Bay of Bengal and the open Indian Ocean

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Fig. 12.4
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(a) Imbricated rock slabs, thought to have been produced by ancient tsunamis. (a) Imbricated slabs, situated 32 m above sea level, at Mermaids Inlet, Jervis Bay, New South Wales, Australia. Photo courtesy of Ted Bryant, University of Woollongong. (b) Imbricated granite slabs, Clifton Beach, Cape Town, South Africa (After Master 2011). Measuring tape is 1 m long

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12.2 Geological Setting

Penang Island is situated off the west coast of Peninsular Malaysia, and is composed of rocks similar to those which constitute the mainland (Kwan et al. 1992; Ong 1993). The rocks of the Malay peninsula have been divided into three tectonic domains, the granitic West Coast and East Coast provinces (Bignell and Snelling 1977), separated by the thin medial Central Province. These are now called the Western, Central and Eastern Belts of Malaysia (Tate et al. 2008). The Western and Central Belts are separated by the Bentong-Raub suture zone, representing a collision zone between the two provinces (Hutchison 1977). The Western Belt is part of the North-Thailand-West Malaya Main Range province, which has mainly S-type biotite granites and abundant tin mineralization of Mid to Late Triassic age (Bignell and Snelling 1977; Liew and McCulloch 1985; Azman 2000; Searle et al. 2012). The Central and Eastern Belts (or East Malaya province) consist predominantly of Permian-Triassic I-type hornblende-biotite granites (Searle et al. 2012).

Three episodes of granite emplacement at 307 ± 8 Ma, 251 ± 7 Ma and 211 ± 2 Ma have been suggested for Penang and the NW Main Range province, based on Rb-Sr dating (Bignell and Snelling 1977; Kwan et al. 1992). However, all the granites have been affected by a hydrothermal conductive convection system related to the late Triassic intrusions (Kwan et al. 1992). Recent U-Pb zircon dating of granites from near Kuala Lumpur has yielded ages of 215 ± 7 to 210 ± 7 Ma (Searle et al. 2012).

Penang Island is underlain by two main granite plutons, the North Penang Pluton, composed of biotite-orthoclase granite, and the South Penang Pluton, composed of biotite-muscovite-microcline granite (Ong 1993). The North Penang Pluton has been divided into three major units known as the Tanjung Bungah granite, the Feringgi granite, and the Muka Head microgranite (Ong 1993). The granites on Penang have been dated at 215 ± 6 Ma and 209 ± 2 Ma (Liew and McCulloch 1985). The biotite and muscovite granites of Penang Island have been extensively weathered (Geological Survey of Malaysia 1992; Ahmad et al. 2006). Fission track work on Penang granites has shown that the island suffered a major episode of tilting and uplift in the Oligocene/Miocene (Kwan et al. 1992), since when there has been uplift and erosion of a formerly buried granite body. The western and southern parts of the island have a thin veneer of late Cenozoic sedimentary cover, and there are thick soils of Pleistocene to Holocene age (Geological Survey of Malaysia 1992; Ahmad et al. 2006). At the present day, the Malaysian Peninsula and Penang Island are part of the stable and aseismic Sunda Shelf. The closest tsunamigenic fault zone is the Sumatran-Andaman subduction zone reverse fault along the plate boundary between the Indo-Australian Plate and the SE Asian Plate. The island has rugged topography in its interior, with a highest elevation of 735 m, and a narrow coastal zone where the beaches are, with the elevation rising rapidly to over 100 m within a few hundred metres from the coastline. The beaches on Penang are situated in those few places where the elevation is less than 15 m (Fig. 12.5).

Fig. 12.5
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Digital elevation model for NW Penang Island, after Anon (2011), modelled from ASTER satellite data obtained from GDS (2011). The beaches are found in areas with average elevation less than 15 m (shown green), along the west and north coasts. The locality of the imbricated boulders on Batu Feringgi beach is in the extreme north of the island, indicated with a red oval

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The Batu Feringgi area of Penang Island is underlain by the Feringgi biotite granite (Ong 1993). Petrographically, this granite is composed of quartz, orthoclase, microcline, plagioclase, biotite, and minor muscovite (Ahmad et al. 2006). Some varieties are porphyritic, with large alkali feldspar phenocrysts; other varieties contain large pockets of tourmaline.

12.3 Imbricate Boulders on Batu Feringgi Beach

The locality of the imbricate boulders is at 5°28′51.77″N, 100°15′72″E, in the Batu Feringgi district, northwest Penang, at the northeast end of the long sandy beach of Batu Feringgi, at the start of the rocky headland (having an elevation of up to 30 m) that forms the northernmost point of Penang Island (Figs. 12.5 and 12.6). Here loose granite boulders are arranged landward of a granitic outcrop which contains steep northward-dipping joints (Fig. 12.7).

Fig. 12.6
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Google Earth satellite view of northern Penang Island, showing Batu Feringgi Beach and the locality of the imbricated granite boulders discussed in this study

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Fig. 12.7
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Loose boulders (numbered) at the northeast end of Batu Feringgi Beach, which are banked up against a rocky jointed granite headland. Note sitting person (on left hand side) for scale

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There are three boulders (labeled 1, 2 and 3, Fig. 12.8) which are arranged in an imbricate fashion, with steep northward (seaward) dips of 45° 65° and 70° respectively. The boulders are tabular in shape, and are roughly rectangular or in the shape of parallelepipeds with rounded corners. Although their exact shapes and hence masses are not known, a rectangular box derived from their axial dimensions a, b and c (with a > b > c) would indicate the maximum masses of the boulders, which are calculated as 2.3, 1.1 and 2.4 tonnes for boulders 1, 2 and 3 respectively (Fig. 12.8, Table 12.1). The shapes of the boulders were dictated by the presence of joints in the coastal outcrops, which represent an uplifted and exhumed tropically-weathered granite-tor landscape with residual soils (Ahmad et al. 2006). Boulders that have moved recently can be recognized from the occurrence of striae, crush zones, fresh scars or percussive marks (Etienne and Paris 2010). Such features generally occur when boulders move over a rock platform or limestone terrace, which is absent at Batu Feringgi beach. No such features were found on the imbricated boulders in the present study, indicating that they were probably not of recent origin, but had been subjected to marine erosion for long enough to have obliterated any trace of such damage resulting from transport. The absence of adhering barnacles, limpets or seaweed attachment scars rules out the recent transport of the boulders from the intertidal zone, and they may have been moved only a few metres.

Table 12.1 Physical parameters for imbricated boulders
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Fig. 12.8
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Imbricate arrangement of three boulders (1, 2 and 3), which are analysed in this study. The large granite mass behind boulder 1 is an outcrop, rather than a loose boulder

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12.4 Hydrodynamic Calculations

Boulders transported by moving water can be moved by three different locomotion modes- sliding, rolling, and saltation, and different hydrodynamic factors are involved in each mode. The hydrodynamic equations for the initiation of boulder transport were originally obtained by Nott (1997, 2003), and are called the Nott equations. These equations have been used widely to model boulder transport by storms and tsunamis (Mastronuzzi and Sansò 2000; Kennedy et al. 2007; Scicchitano et al. 2007; Imamura et al. 2008; Etienne and Paris 2010; Switzer and Burston 2010; Goto et al. 2010b; Costa et al. 2011). Nandasena et al. (2011) modified the Nott equations, based partly on the work of Voropayev et al. (2003) and Nanayama and Shigeno (2006), and produced a new set of hydrodynamic equations governing boulder transport initiation.

The conditions for the initiation of transport of a submerged or subaerial boulder under sliding mode are given by the modified Nott equation (Nandasena et al. 2011), as follows:

$$ {\mathbf{u}}^{\mathbf{2}}\ge \mathbf{2}\left(\left[{\boldsymbol{\uprho}}_{\mathbf{s}}/{\boldsymbol{\uprho}}_{\mathbf{w}}\right]-\mathbf{1}\right)\mathbf{gc}\left({\boldsymbol{\upmu}}_{\mathbf{s}}\mathbf{cos}\kern0.15em \boldsymbol{\uptheta} +\mathbf{sin}\kern0.15em \boldsymbol{\uptheta} \right)/\left[{\mathbf{C}}_{\mathbf{d}}\left(\mathbf{c}/\mathbf{b}\right)+{\boldsymbol{\upmu}}_{\mathbf{s}}{\mathbf{C}}_{\mathbf{l}}\right] $$
(12.1)

where

u = fluid velocity required to initiate boulder movement

ρs = density of boulder

ρw = density of fluid

g = acceleration due to gravity

c = length of shortest axis of boulder

θ = angle of slope

b = length of intermediate axis of boulder

Cd = coefficient of drag

μs = coefficient of static friction

Cl = coefficient of lift

In order to produce imbrication of several boulders, the mode of transport has to be rolling/overturning. Initiation of boulder transport under rolling mode requires higher water velocities than the initiation of boulder movement under sliding conditions. For the case of rolling or overturning boulder movement initiation, the modified Nott equation (Nandasena et al. 2011) is as follows:

$$ {\mathbf{u}}^{\mathbf{2}}\ge \mathbf{2}\left(\left[{\boldsymbol{\uprho}}_{\mathbf{s}}/{\boldsymbol{\uprho}}_{\mathbf{w}}\right]-\mathbf{1}\right)\mathbf{gc}\left(\mathbf{cos}\kern0.15em \boldsymbol{\uptheta} +\left[\mathbf{c}/\mathbf{b}\right]\mathbf{sin}\kern0.15em \boldsymbol{\uptheta} \right)/\left[{\mathbf{C}}_{\mathbf{d}}\left({\mathbf{c}}^{\mathbf{2}}/{\mathbf{b}}^{\mathbf{2}}\right)+{\mathbf{C}}_{\mathbf{l}}\right] $$
(12.2)

According to Nandasena et al. (2011), initial transport of a submerged or subaerial boulder will be in saltation mode when:

$$ {\mathbf{u}}^{\mathbf{2}}\ge \mathbf{2}\left(\left[{\boldsymbol{\uprho}}_{\mathbf{s}}/{\boldsymbol{\uprho}}_{\mathbf{w}}\right]-\mathbf{1}\right)\mathbf{gc}\kern0.5em \mathbf{cos}\kern0.15em \boldsymbol{\uptheta} /{\mathbf{C}}_{\mathbf{l}} $$
(12.3)

For joint-bounded boulders, movement is initiated when

$$ {\mathbf{u}}^{\mathbf{2}}\ge \mathbf{2}\left(\left[{\boldsymbol{\uprho}}_{\mathbf{s}}/{\boldsymbol{\uprho}}_{\mathbf{w}}\right]-\mathbf{1}\right)\mathbf{gc}\left(\mathbf{cos}\kern0.15em \boldsymbol{\uptheta} +{\boldsymbol{\upmu}}_{\mathbf{s}}\mathbf{sin}\kern0.15em \boldsymbol{\uptheta} \right)/{\mathbf{C}}_{\mathbf{l}} $$
(12.4)

Following Noormets et al. (2004) and Nandasena et al. (2011), the fluid density is that of seawater, taken as 1.02 g/cm3; the acceleration due to gravity is 9.81 m/s2; and the coefficients of drag (Cd), static friction (μs) and lift (Cl) are taken as 1.95, 0.7 and 0.178, respectively. The mean density of the Batu Feringgi granite is 2.71 g/cm3 (Ahmad et al. 2006). The slope on the Teluk Bahang beach is very flat, and is taken as 2°. Because the offshore topography is unknown, additional calculations were made for a slope of 5°. The physical parameters for the imbricated boulders are given in Table 12.1, and the results of the calculations are given in Table 12.2.

Table 12.2 Calculated flow velocities of seawater needed to initiate boulder movement (umin) under different transport modes (sliding, rolling/overturning and saltating), for loose boulders in submerged and subaerial settings only, at slopes of θ = 2° and θ = 5°
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12.5 Discussion

Because the boulders are imbricated, their transport mode must have been rolling/overturning. The imbrication is produced by bedload transport, and boulder to boulder interaction, and the imbricated boulders dip upstream (in this case, offshore), in the direction from which the currents were coming. For each slope angle (θ = 2° and θ = 5°), the highest velocity needed to initiate movement of any of the three boulders is taken as the minimum velocity of flow that would have moved all three boulders together (in order to produce the imbricate stacking observed). Since the imbricated boulders are stacked among jointed outcrops of granites, calculations were also made for the minimum velocities required to initiate boulder movement in joint-bounded mode, in addition to the calculations made for the case of free-standing boulders (Table 12.2).

The calculated minimum velocities required to initiate rolling/overturning of the imbricated boulders range from 6.07 (θ = 2°) to 6.12 (θ = 5°) m/s for rolling mode, to 9.39 (θ = 2°) and 9.53 (θ = 5°) m/s for joint-bounded mode. It could be argued that the particular values of the coefficients of lift (Cl), and static friction (μs) used in the equations of Nandasena et al. (2011) may not be valid for this case. So a sensitivity analysis was made, in which the coefficients of static friction (μs) and of lift (Cl) were varied by a factor of two (separate calculations were done using half and twice the values used by Nandasena et al. 2011, which originate from Noormets et al. 2004). As an example of these calculations, the umin value for boulder 1 at θ = 5°, under joint-bounded conditions, which was calculated as 9.23 m/s using the μs value of 0.7 from Noormets et al. (2004) (Table 12.2), yielded umin values of 9.08 and 9.48 for μs values of 0.35 and 1.4 respectively. So varying the coefficient of static friction (μs) by a factor of two results only in a ±3 % change in the values of minimum velocities required to initiate boulder movement. The effect of changing the coefficient of lift (Cl) is much larger, yielding umin values of 6.52 m/s for a Cl values of 0.356, and 16.04 m/s for a Cl value of 0.059, compared to the Noormets et al. (2004) value of 0.178 used in Table 12.2. If the density of seawater is changed from the 1.02 g/cm3 used in the calculations (Table 12.2) to a value of 1.5 g/cm3, to allow for a turbulently mixed muddy ocean, then the umin value for boulder 1 falls from 9.23 to 6.48 m/s. The sensitivity analysis shows that even if many of the coefficients used in the calculations presented in Table 12.2 were varied by a factor of 2, the results would still yield minimum velocities in the range 6–9.5 m/s in order to initiate the movement of the imbricated boulders under free standing and joint-bounded conditions. These velocities are higher than the maximum velocities experienced at this specific locality during the great 26 December 2004 Indian Ocean tsunami, the largest known tsunami in recorded history.

A number of video recordings of the incoming tsunami waves on 26th December 2004 were captured by tourists on the north coast of Penang Island, and are available for viewing on YouTube (Videos 12.1, 12.2, and 12.3). Analysis of Video 12.1, taken at Tanjung Bungah some 3 km east of Batu Feringgi, and which includes footage of tsunami waves passing by the mosque called Masjid Terapung, which is built on the coast on a promontory at right angles to the main trend of the coastline, has allowed for relatively accurate measurement of wave velocities. The wave passes by the Masjid Terapung, which has dimensions of 77 m (obtained from Google Earth), in 15 s, which gives an open-water incoming wave velocity (before run-up), of about 5.1 m/s. The video also shows that as the waves reached the shoreline, the velocities decreased rapidly in shallow water (although the actual velocities could not be quantified from the video because of a lack of measurable landmarks). The measured wave heights recorded at Tanjung Bungah were 3.0 m (Roy and Ismail 2006; Roy et al. 2007). Since the velocity of tsunami waves is equal to the square root of the depth multiplied by the acceleration due to gravity (9.81 m/s2), it means that as the waves approach the shore (i.e., during run-up) their velocities decrease exponentially, and their amplitudes increase (e.g., Okal 1988). Another video (Video 12.2), taken on Batu Feringgi beach, shows a narrow shoreline of about 10 m, which was traversed by the first incoming tsunami wave in 3 s, giving a wave velocity of about 3.3 m/s. The water brought in by the first wave was about 0.5 m deep, which would have been just sufficient to cover the imbricated boulders. A third video, taken from a hotel window, shows tsunami waves coming in parallel to a sea-wall, and moving at a rate of about one car-length per second, or about 3 m/s. Videos 12.2 and 12.3 do not record the largest waves to have hit Batu Feringgi, since the highest measured wave heights there were recorded as 3.3 m above mean sea level (Roy and Ismail 2006; Roy et al. 2007).

The run-up wave heights and wavelengths of tsunami waves hitting a shoreline are strongly dependent on local conditions. The best studied tsunami is the 11 March 2011 Japanese Tohoku tsunami, for which the instrumentally-recorded data set contains the most comprehensive collection of coastal and deep-ocean observations ever made during a single tsunami event. From this event, more than 350 coastal water level records show regions of signal attenuation and amplification due to refraction, reflection, and bathymetric focusing (Eble et al. 2012).

Similarly, during storm events, wave behavior is strongly influenced by local conditions. According to Regnauld et al. (2004), bathymetric complexities exercise significant controls on wave refraction patterns during storms, and have resulted in significant variability in coastal response during storms in western France. Another factor of great importance is the storm surge, resulting from the pile up of wind-driven waters from the continental shelves onto the coastal areas (Coch 1994; Flather 2003). Bathymetric factors such as water depth and slope are thus expected to control the velocities, wavelengths and propagation directions of waves on Batu Feringgi beach during storms, as well as during tsunami events.

Penang Island is less than 5.5° from the Equator, and hence is generally thought to be relatively free from the effects of major tropical storms or hurricanes, which, for reasons to do with global atmospheric vorticity, are mainly confined in subequatorial belts between 10° and 30° north and south of the Equator (McBride 1995) (Fig. 12.9). However, on 26th December 2001, typhoon Vamei was formed at 1.4°N in the southern South China Sea in what was previously deemed a cyclone-free equatorial zone, strengthened quickly, and made landfall along the southeast coast of Malaysia 60 km northeast of Singapore, and dissipated over Sumatra on 28th December 2001 (Fig. 12.10). The wind speeds were calculated to be as high as 36 m/s in the southern South China Sea, and would have generated wind waves with wave heights as high as 7.5 m on the southeast coast of Malaysia (Mohammed et al. 2012). An analysis by Chang et al. (2003) suggested that the return frequency for such a typhoon is approximately 1 in 400 years, which could also be taken as the return frequency for similar typhoons or tropical storms affecting Penang Island.

Fig. 12.9
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Map of global tropical cyclone tracks recorded from 1985 to 2005. Compiled by Nilfanion (2006), with background image from NASA

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Fig. 12.10
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Close-up view of SE Asia region in global map of tropical cyclone tracks from 1985 to 2005 (From Nilfanion 2006). The track in the centre of the image is from typhoon Vamei, 26–28 December 2001 (Chang et al. 2003)

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The question of whether boulders are moved by tsunami waves or storm waves is still a hotly debated issue, because the processes are quite similar, and it is difficult to unequivocally prove one mechanism of formation to the exclusion of the other (Williams and Hall 2004; Switzer and Burston 2010; Etienne and Paris 2010; Cox et al. 2012; Lau et al. 2012; Seshachalam et al. 2012). The Batu Feringgi beach appears to be fortuitously situated far enough from the seismogenic and tsunamigenic Sumatra-Andaman fault zone to be on the fringe of the area most prone to devastating effects of megatsunamis generated by the largest earthquakes generated in that region. Yet it contains boulders that were moved by waves with a minimum calculated velocity which was much higher than experienced in that particular location during the largest tsunami event ever recorded. It is possible that some previous pre-historic tsunami could have moved these boulders, but the chances are rather small, since the return frequency for tsunamigenic mega earthquakes, similar to or larger than the magnitude 9.2 earthquake of 26 December 2004 along the Sumatra-Andaman subduction zone, is of the order of millennia (McCaffrey 2008). The imbricated boulders, located at sea level, would be unlikely to have escaped erosion by attrition in this high-energy marine environment over such a length of time. The tectonic situation of Penang Island in the stable Sunda Straits also means that it is not prone to tsunamis generated by landslides on the continental shelf, most of which are triggered by seismic activity (Ward 2001; Nadim et al. 2005). In this case, tsunamis can effectively be ruled out as the causative agents, and it leaves the agency of rare (but not unknown) tropical storms, such as Vamei of 2001, as the most likely cause of the boulder movements on Batu Feringgi.

12.6 Conclusions

The Malaysian Peninsula and Penang Island are part of the stable and aseismic Sunda Shelf. The closest tsunamigenic fault zone is the Sumatran-Andaman subduction-zone reverse fault along the plate boundary between the Indo-Australian Plate and the SE Asian Plate. The 26 December 2004 Sumatra-Andaman earthquake produced a tsunami that flooded the Batu Feringgi area of north Penang Island, but did not result in appreciable damage to infrastructure. The measured wave run-ups were up to 3.3 m above mean sea-level on Batu Feringgi, and the incoming wave velocities were in the range of 3–5 m/s, based on video analysis. However, the minimum velocities required to initiate movement of the imbricate boulders from Batu Feringgi were calculated to be in the range of 6–9.5 m/s. Hence it is concluded that the imbricated boulders on Batu Feringgi beach are the result of tropical storm activity in the past, rather than from the recent tsunami. If the magnitude 9.2 Sumatra-Andaman earthquake of 26 December 2004 (one of the strongest ever recorded) could not produce a sizeable tsunami on the north coast of Penang, then it is unlikely that any recent tsunami could have produced the imbricated boulders at Batu Feringgi, leaving tropical storms as the most likely culprit. The north coast of Penang (including Batu Feringgi beach) is thus regarded as safe from the hazard of damaging tsunamis resulting from mega-earthquakes in the Sumatra-Andaman subduction zone, but the area is prone to rare tropical storm damage, with wave velocities exceeding 6–9.5 m/s.