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Structural Geology

Merapi structural history is divided into four Periods: Ancient, Middle, Recent and Modern Merapi, as revealed by field studies and geochronological data (Camus et al., 2000; Berthommier et al., 1990). The Ancient Period may have begun around 40 000 yr BP and lasted until 14 000 yr BP when the Middle Period begun. It lies on the basalt and basaltic older structure (Pre-Merapi). Auto-brecciated lava flows, St. Vincent-type pyroclastic flows and lahar deposits compose this structure. Saint Helens-type Edifice collapse took place during the Middle Merapi period which eroded the upper part of the Ancient Merapi structure. A new structure was rebuilt through the next eruptive episodes. The Middle Merapi structure was covered by the deposits of the Recent Merapi eruptions which began around 2200 yr BP. The summit part of Merapi is dominantly composed by the lava of the Modern Merapi eruptions (after 1786)

Seismic studies

Some geophysical surveys were conducted to study the subsurface structure of Merapi. Based on the distribution of Volcano-tectonic (VT) events, Ratdomopurbo (1995) hypothesized a magma chamber located at depth >5 km and a secondary magma chamber located at 1.5 – 2.5 km depth. The absence of VT events at those two zones was interpreted as ductile zones with high temperature that might correspond to magma chambers (Fig. 1.3).
Structural sounding were carried out using active artificial seismic sources in the framework of a cooperation between Volcanological Survey of Indonesia (VSI) and GeoForschungsZentrum Potsdam (GFZ) (Wegler et al., 1999; Wegler and Luhr, 2001; Maercklin, 1999; Wegler 2006). Three 3 km long seismic profiles, each consisting of up to 30 three-component seismometers with an interstation distance of 100 m, were built up in an altitude range between 1000 and 2000 m above sea level (Fig. 1.4).
Fig. 1.4. Location of seismic lines, source points, and mapped fracture zones. Profiles shown as dotted lines are not interpreted. Two circles mark the extend of possible weakness zone (Maercklin et al., 2000).
The observed seismograms show a spindle-like amplitude increase after the direct P phase. This shape of envelope can be explained by the diffusion model. According to this model there are so many strong inhomogeneities that the direct wave can be neglected and all energy is concentrated in multiple scattered waves.
As a result of the inversion using the diffusion model, in the frequency range of 4–20 Hz used in this study, the scattering attenuation is at least one order of magnitude larger than the intrinsic absorption . The mean free path of S waves is as low as 100 m (Wegler and Luhr, 2001).
A low velocity layer above 300 m of depth was suggested where the velocities range from 0.7 to 3.4 km/s (Riedel et al., 1999). The south flank likely has lower velocity than the north-east flank which can be attributed to the different compactness of the different structural age (Riedel et al., 1999; Wegler and Luhr, 2001). In addition, according to Maercklin (1999), there are seismic reflectors which could be explained with a simple two dimensional model based on the ray theory, and it was shown that they are caused by open fissure zones.
A larger scale tomographic study including Merapi volcanic complex was done by Koulakov et al (2007; 2009) using body waves arrival times from tectonic earthquakes. In the crust beneath the middle part of central Java, north to Merapi and Lawu volcanoes, a large and very intense anomaly was observed with a velocity decrease of up to 30% and 35% for P and S models, respectively. Inside this anomaly, E-W orientation of fast velocity takes place that is probably caused by regional extension stress regime along the N-S direction. In vertical section beneath this anomaly, faster horizontal velocities were observed that might be explained by layering of sediments and/or penetration of quasi-horizontal lenses with molten magma (Fig. 1.5) (Koulakov et al 2009; Luehr et al., 2013).
Those tomographical studies could not confirm the existence of the aseismic zone suggested by Ratdomopurbo and Poupinet (1995; 2000) as a shallow magma chamber at depth between 1.5 and 2.5 km. Due to the very scattering behaviour of the shallow layer, the seismic waves recorded during the active source experiment (Wegler and Luhr, 2001) didn’t contain much information of the deeper zone.
Eruptive precursors including seismic activity were reported by Ratdomopurbo and Poupinet (2000) regarding the activity from 1982 to 1995. There were two main eruptive cycles during this period, from 1984 to 1986 and from 1992 to 1994. Both cycles were preceded by VT earthquakes, whereas only the eruption of 1992 was preceded by tremor activity from more than 1 year before. A detailed chronology of pre-eruptive and eruptive activity of June 2006 was presented by Ratdomopurbo et al. (2013). The early precursors occurred in the middle of the year 2005 with seismic activity, increase in deformation, and possible increase of SO2. The short term precursors were marked by an increase in number of VT, MP, and LF as well as deformation in the beginning of 2006.
Focal mechanisms of VT events recorded in 2000 – 2001 were estimated by Hidayati et al. (2008) using both polarity and amplitude of P-wave first motions. VTA and most deep VTB events are of normal-fault types, whereas VTB events located close to the surface yield both reverse and normal fault solutions. Hidayat et al. (2000; 2002) studied very long period (VLP) events that occurred in 1998 and determined that these have periods of 6 – 7 s, display similar waveforms from event to event, and are coeval with MP or LF earthquakes. They carried out moment tensor inversion of the waveforms and proposed a source model consistent with a dipping crack located at about 100 m depth under the 1998 dome. They suggested a source process involving the sudden release of pressurized gas through the crack over a time span of about 6 s. No VLP events were observed during the active periods of 2001 and 2006, whereas a significant number of VLP events were observed in 2010 prior to and during the eruption (Jousset et al., 2013).
Several works were dedicated to study the elastic property changes of the medium using seismic data. Based on the cross spectral method applied on the codas of similar events, Ratdomopurbo and Poupinet (1995) observed an increase of shear wave velocity of 1.2 % before the 1992 eruption.
Using the data of the active source experiment (Wegler et al., 1999; Wegler and Luhr, 2001; Maercklin et al, 1999; Wegler et al., 2006), Wegler et al. (2006) inferred temporal changes in the elastic properties of the edifice of Merapi before the 1998 eruption. They observed a total increase of seismic velocity of up to 0.23% in 2 weeks until 3 days after the eruption. Later, Sens-Schonfelder and Wegler (2006) reported a strong seasonal effect as a function of the Ground Water Level (GWL) of the velocity variation obtained from ambient noise with up to1 % of annual variation.

Other geophysical methods


EDM measurements were conducted for the period of 1988 – 1994 on a summit trilateration network (Young et al., 2000). Cross-crater strain rates accelerated from less than 3 x 106/day between 1988 and 1990 to more than 11 x 106/day just prior to the January 1992 activity, representing a general, asymmetric extension of the summit during high-level conduit pressurization. During the effusive lava extrusion, strain decreased below the background level of less than 2 x 106/day. EDM measurements between lower flank and crater benchmarks during 4 years before the 1992 eruption revealed a long term displacements as high as 1m/year.
Later, in the period between November 1996 and March 1997, other deformation experiments were established using tiltmeters and GPS equipments (Beauducel and Cornet, 1999). An interpretation using a three-dimensional elastic model based on the mixed boundary element method and a near-neighbor Monte Carlo inversion lead to a suggestion of a magma chamber at the depth of 8 km below the summit and 2 km to the east of it. The estimated volume attributed to this magma chamber is about 11 x 106 m3.
Within the Indonesia – German joint research project MERAPI, four tiltmeter stations were installed on the flank during 1995 – 1997 (Rebscher et al., 2000; Westerhaus et al., 2008). In spite of the absence of strong volcano-induced tilt anomalies, rapid, step-like drift changes were detected with amplitudes of 15 to 80 µrad which are generally related to the alternation of wet and dry seasons. Finite-Element-Modelling showed that sign and amplitude of these perturbations are compatible with a pressure source located 1.2 km below the summit with radius of 1.7 km which is consistent with aseismic zone revealed by hypocenter distribution. These perturbations are interpreted as the effect of annual input of meteoric water to the pressure within deeper parts of the hydrothermal system on the central vent of Merapi (Westerhaus et al., 2008).

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Geoelectric measurements

Based on DC resistivity survey, Friedel et al. (2000) developed resistivity models for the north, west, and south flanks with depth of investigation between 600 and 1000 m. For the high conductivity zones appearing in the West and South, a hypothesis was brought forward that the anomalies are caused by meteoric water penetrating highly permeable layers of volcanic deposits to great depth where it influences the extent of hydrothermal zones. In August 2000 a permanent SP and temperature monitoring station was established at the fumaroles field Woro. Correlations between SP, ground temperature anomalies, MP events and the appearance of lava dome during 2001 activity were observed (Friedel et al., 2004). Many SP anomalies and gas temperature coincided with the occurrence of Ultra Long Period (ULP) seismic events which are interpreted as the effects of gas emissions (Byrdina et al., 2003; Richter et al., 2004).
Muller and Haak (2004) derived a 3-D model of the electrical conductivity structure of Merapi volcano from magnetotelluric (MT) sounding and geomagnetic induction vectors (Fig. 1.6). The final model consists of two 3-D structures in the volcanic edifice, i.e. a central conductor (D) and a second conductor lying 5 km to the southwest of summit (E). The high conductivity material is probably hot saline water as suggested by position and lateral extent of the high conductivity material. Another conductive layer at the depth of 3.5 – 5.3 km (C) is attributed to a very porous regional layer containing seawater or fluid of comparable conductivity (Rittel et al., 1998; Muller and Haak, 2004).


Gravity measurements have been applied at Merapi both for mapping and monitoring purposes. The first gravity measurements around Merapi have been carried out by Yokoyama in 1970 (Yokoyama et al., 1970), Untung and Sato (1978). The Bouguer anomaly pattern shows that Merapi is characterized by a low anomaly centred on the summit area. A two-dimensional gravity model indicates that the material on the summit of Merapi area has a density contrast of –900 kg/m3. Based on the model, there are three density values of Merapi: 2600 kg/m3 around the foot, 1800 kg/m3 in the mid-body, and 1600 kg/m3 in the summit area. This model is calculated by assuming a mean regional density value of 2500 kg/m3. The rocks in the summit areas are probably composed of sand, tuff, lava fragments and lava dome (Wahyudi, 1986; Sidik, 1987; Arsadi, et al., 1995b).
Continuous gravity monitoring and microgravity surveys including accurate levelling on Merapi volcano have been carried out by French-Indonesian teams during 1993 – 1995 (Jousset, 1996). Significant variations in gravity were observed for the period 1993-1994. The gravity increased by 100 to 400 µgal which was explained by the change of neighbouring topography due to the growth of the dome during the considered period. Residual drift of continuous data showed correlation with LF events and the occurrence of pyroclastic flows (Jousset, 1996).

Merapi seismic network

Historical Review

Monitoring volcanoes in Indonesia began in 1920 with the establishment of the Dinas Penjagaan Gunungapi by the Dutch East India Company (Dutch: Vereenigde Oost-Indische Compagnie, VOC). This establishment is a response to the eruption of G. Kelut in previous year which caused more than 5000 deaths. Shortly after, observation posts were established including at G. Merapi. In 1924 the first seismic station was installed at G. Merapi, with a Wichert type seismometer. It is an entirely mechanical seismometer, made in Gottingen (Germany). It is essentially an inverted pendulum, which records both components of horizontal motion on rolls of smoked paper. It weights 1000 kg, and has a natural period of 8 seconds. Damping is provided by two air-pistons on the top of the instrument. The pendulum is centered by placing a series of small weights on top of the main mass. This seismometer is no longer in operation, but is visible in BPPTK (Balai Penyelidikan dan Pengembangan Teknologi Kegunungapian). This seismometer was installed at 14 km west of the summit (Neumann van Padang, 1933). He observed an increasing seismic activity before the eruption of 1930. In 1968, Shimozuru performed seismic observations by installing seismographs at about 10 km south of the summit. Seismic signals were recorded on magnetic tapes (Shimozuru et al., 1969). Using 6 months of observation, he proposed the first classification of events of Merapi and their associated physical process. He suggested that multiphase (MP) events are related with dome growth.
Along with the development of monitoring technology, monitoring system of G. Merapi was also improved. In 1982 telemetry system began to be implemented. In cooperation between USGS and VSI (Volcanological Survey of Indonesia, former of CVGHM, Centre of Volcanology and Geological Hazard Mitigation), 6 short period seismic stations were installed around G. Merapi whose data were transmitted directly to Yogyakarta using VHF telemetry system (Koyanagi and Kojima, 1984). In 2 January 1991 the seismic network was digitized by using a Data Translation board, a dedicated PC, a stabilized power supply and the PCEQ – IAVCEI software (J.-L. Got, pers. comm.), in the frame of the cooperation with the French foreign office. In August 1994, 6 more short period seismic stations were installed and digitized, among which the 3-component summit station, in the frame of the cooperation with the French CNRS (J.-L. Got, pers. comm.). Then in 1994 Broadband seismometers started to be used in Merapi with the cooperation between the VSI and the GeoForschungsZentrum Potsdam (Beisser et al., 1996).

Table of contents :

1.1 Background
1.2 Previous studies
1.2.1 Structural geology
1.2.2 Seismic studies
1.2.3 Other geophyisical methods Deformation Geoelectric Measurements Gravimetry
1.3 Merapi seismic network
1.3.1 Historical review
1.3.2 Recent seismic network
1.3.3 Instrumental problems
1.4 Main features of Merapi seismic events
1.5 Thesis Structure
2.1 Introduction
2.1 Data and method
2.3 Seismic chronology during 2010 crisis
2.4 Discussion
2.5 Conclusion
3.1 Introduction
3.2 Data and method
3.2.1 Absolute and uncertainty estimation
3.2.2 Relative location using double difference method
3.3 Results
3.3.1 Absolute locations
3.3.2 Relative locations
3.3.3 Depths versus arrival time difference models
3.3.4 Temporal evolution of the hypocenter distribution
3.4 Discussion
3.4.1 Aseismic zone in Merapi eddifice
3.4.2 Magma migration
3.5 Conclusions and perspective
4.1 Introduction
4.2 Data and method
4.2.1 RSAM and modified RSAM (MRSAM)
4.2.2 Hindsight eruption forecasting
4.3 Results
4.3.1 RSAM and MRSAM
4.3.2 Eruption forecasting
4.4 Discussion
4.5 Conclusions
5.1 Introduction
5.2 Data and method
5.2.1 Recursive event detections
5.2.2 Extraction of families
5.3 Results and discussions
5.3.1 Event detections
5.3.2 Families of events
5.4 Conclusions and perspectives
6.1 Introduction
6.2 Data and method
6.2.1 Velocity variations in the coda of multiplets Multiplets data Doublet method Stretching method
6.2.2 Velovity variations using noise correlation Noise data Data selection and time synchronization Corrections of rain effects 2D location of velocity perturbations
6.3 Results
6.3.1 Velocity variations obtained from multiplets
6.3.2 Velovity variations obtained from noise correlation
6.4 Discussions
6.4.1 Comparison of the methods used
6.4.2 Velocity changes prior to the eruption
6.4.3 Comparison with earlier studies at Merapi
6.4.4 AVV and tectonic events
6.5 Conclusions and perspectives


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