Introduction to the Volcanology

Abstract

The main volcanological concept is shown and expressed so that any volcano can be understood easily. Volcanic products are listed and explained in plain language from lava flow to various pyroclastic products. The volcanic products have been explained schematically and their textural, field relationships characteristics are highlighted. The origin of magma within the interior of the Earth is also explained and the link between mantle and crust has been shown. The relationship among crust, mantle, and core has been highlighted embracing the source-to-surface model. An updated explanation of the Pyroclastic Density Currents (PDC) has been done to perceive their danger. Some of the most successful Volcanology books have been used. This will help the students, with a passion for Volcanology, to understand the principles of Volcanology.

1. Introduction

The chapter summarizes the main concept in Volcanology with an overview that will help to understand other chapters presented in this Volcanology Book. This chapter has been chosen, in particular, for under-graduate people who want to deepen their knowledge in Volcanology, producing several avenues that can help them to develop their own research interests—from volcanic geology to forecast volcanic eruption [1, 2, 3, 4, 5, 6, 7, 8, 9, 10, 11, 12, 13]. Some concepts about the constitution of the Earth must be described to better understand the volcanology base. We know that the Earth, broadly speaking, is stratified with a temperature around 6000°C at its center, as at the surface of the sun. The stratification gives some cloud on the production and genesis of the magma. The issue is very much complex but worth to investigate for the sake of clarity about magma generation. The Earth is stratified not only for temperature but also for density, minerals, and physics state as liquid versus solid. In fact, it is the fluid zone of the Earth (asthenosphere versus outer core), the areas much advocated for partial melting: The asthenosphere is the denser, weaker layer beneath the lithospheric mantle. It lies between about 100 and 410 km beneath the earth’s surface. The temperature and pressure of the asthenosphere are so high that rocks soften and partly melt, becoming semi-molten. Earth’s outer core is a liquid layer about 2400 km thick and composed of mostly iron and nickel that lies above the earth’s solid inner core and below its mantle. Its outer boundary lies 2890 km beneath Earth’s surface. Unlike the inner (or solid) core, the outer core is liquid (Figure 1). Hence, the Earth is made of several layers—(a) crust, mantle, inner and outer core. (b) Lithosphere, asthenosphere, mesosphere to core (Figure 2). The asthenosphere must be the most productive melt layer within the mantle structure.

Figure 1.

Earth with a window inside the planet (author’s collection).

Figure 2.

Plot of the P and S wave velocity, and density versus depth from the surface to the core. Upper mantle with a profile of S wave velocity (V_s) and the seismic discontinuity (redrawn after Schmincke [14]).

2. Mantle structure

The section of crust and upper mantle with a profile of the S wave velocity (V_s, Figure 2) recognizes some physical characteristics that help to understand the dynamic of magma in the mantle. Probably, much of the magma that reaches the crust is produced in the asthenosphere, although deeper mantle origin can be envisaged from tomography data (410 and 660 km). Although, it is not clear if some magma can be produced in the outer core and/or in the lower mantle. Helffrich and Wood [15] affirm that seismological images of the Earth’s mantle reveal three distinct changes in velocity structure, at depths of 410, 660, and 2700 km. The first two as said are best explained by mineral phase transformations (Figure 2), whereas the third D″ layer (2700 km) probably reacts due to change in chemical composition and thermal structure. In addition, tomographic images of cold slabs in the lower mantle, and the occurrence of small-scale heterogeneities in the lower mantle all indicate that subducted material penetrates the deep mantle, implying whole-mantle convection. In contrast, geochemical analyses of the basaltic products of mantle melting are frequently used to infer that mantle convection is layered, with the deeper mantle largely isolated from the upper mantle. The geochemical, seismological, and heat-flow data are all consistent with whole-mantle convection provided that the observed heterogeneities are remnants of recycled oceanic and continental, respectively, of mantle volume (Figure 3). The convective cells are the engine of the plate tectonic for whole-mantle versus upper-mantle cells (Figure 3). Figure 4 is the best guess for a model of mantle recently published [16, 17].

Figure 3.

(a): Earth slide highlighting the convective cycle on all the mantle (author’s collection). (b): Earth slide highlighting the mantle with 2 convective cycles—Upper mantle and lower mantle (author’s collection).

Figure 4.

Mantle slice with a characteristic chemical and physical heterogeneity (modified from Helffrich and Wood [15]). The blue blob is residual slabs and/or metasomatized mantle. The pink blobs are identified with the D″ layer.

3. The crust

The Earth is divided into at least 12 main plates of the oceanic or continental type, delimited by distension and compression margins and in some cases by transform faults (Figure 5, [18]). Figure 6 shows the distribution of active volcanism in diverse geodynamic settings (subduction, rift, and hotspot). Crustal end-members are of fundamental importance to understanding the magma evolution on the earth’s surface, especially for the continental crust. Lately, we have enjoyed the model of Steve Sparks and collaborator (deep and hot intrusion zone [20, 21, 22]). The model clarifies how the volcanic rocks were originated either from metasomatized mantle source with the classical signature and/or by crustal contamination. A detailed explanation of trace elements and isotopes of volcanic systems come anyway, can be in their hand [23]. The continental crust of Wedepohl [19] (Figure 6) still can be taken as the best example to use and to perceive the Conrad and Moho discontinuities as the region where the basaltic magma intrudes, stagnate, and evolve through the surface. The bulk continental crust of Wedepohl [19] has a tonalitic composition with distinctly higher concentrations of incompatible elements. A dioritic bulk crust was suggested by Taylor and Mclennan [24] in contrast to Wedepohl [19]. If we make a section from New Zealand to South America (i.e., Chile), we can observe five types of tectonic margin—(a) subduction, oceanic-oceanic plates, (b) hotspot, (c) oceanic rift, (d) subduction oceanic-continental plates, (e) continental rift (Figure 7, [18]; a, b, c, d, e are partial melting zone with the production of basaltic magmas). Although other lithotypes are also present, we can surely say that partial melting could be the Holy Grail of Igneous Petrology.

Figure 5.

Plate tectonic with larger and smaller plates and with the location of all the volcanic activity (modified from Schmincke [14]).

Figure 6.

Crust section (modified from Wedepohl [19]).

Figure 7.

Plate tectonic section from New Zealand to Chile. Modified from Bosellini [18].

4. Partial melting into the asthenosphere

The rise of the S wave velocity through the lower mantle suggests that little is the melting of the mantle, anyhow, little compared to the asthenosphere. Also, because of the variation of temperature, the pressure and water content are little, and however, not as happens in the asthenosphere. On the contrary, the S wave velocity decrease in the asthenosphere from 70 to 250 km. The velocity still keeps lower numbers until 410 km, as is seen in Figure 3. Hence, the magma can be produced, as well in the upper mesosphere, till 410 km. In this part of the mantle (from 70 to 410 km), the rocks are partly melted (Figure 8). So, partial melting is the most important process acting in the upper mantle for the production of magma [14]. In other words, the lower mantle could melt when it participates in the convective cycles. Partial melting can be advocated each time there is granular compaction that will cause the production of a magmatic liquid that buoyantly rises through the crust till erupting (Figure 8). The rise of magma can have a basalt or broadly much-evolved composition (rhyolitic) depending on how much interaction with the crust will have. However, basalts are compositions that help to model the geochemical processes in the rock origins, so are very important. On the other hand, rocks with an evolved composition help to decipher the processes that occur in the crust—partial melting versus crystal fractionation, assimilation, and other minor processes, such as mixing and/or mingling (Figure 9) [25, 26, 27, 28].

Figure 8.

Partial melting scheme. Magma is formed by the compaction of grains. The liquid is present between grains and is formed squeezing such material and so the magma aggregate and buoyantly rise through channels to the surface (author’s collection).

Figure 9.

Rise of basaltic and rhyolitic magma through the crust. Modified from Schmincke [14].

5. Chemical and mineralogical characteristic of diverse magma in diverse tectonic environment

The Petrographic Classification of magma erupting in diverse tectonic regimes (Figure 10)—(1) Distention environment of thin oceanic crust (≈10 km) with the formation of basaltic magmas (e.g., Island and Atlantic rift). (2) Tectonic regime of hot-spot into an internal oceanic plate with the production of basaltic magma (e.g., Hawaii Island and Emperor submerged chain). (3–4) Compressive environment (subduction regime) with insular and continental arcs with the production of rocks with an evolved composition (felsic, andesitic, and rhyodacitic) (e.g., Japan, New Zealand, Chile, and Alaska). (5) Volcanic activity within a continental plate as a hot-spot with the formation of less-evolved and evolved composition, such as basaltic, andesitic, and rhyodacitic (e.g., Etna volcano). (6) Distention environment within a continental crust with the formation of basaltic, andesitic, and rhyodacitic (e.g., Africa rift valley).

Figure 10.

Diverse magma for the diverse tectonic environment—A simple petrographic outline (modified from Bosellini 17).

6. Effusive magma versus intrusive magma with the corresponding volcanic forms

Two types of volcanism are caused by the magma rise through the mantle and the thick crust (Tables 1 and 2; [4]): Intrusive versus Extrusive volcanism (Table 2). Volatile content is essential for the behavior of the magma that has shaped the Earth’s volcanism and volcanic structures [29]. The subcategories of these two styles of volcanism are plutonic, sub-volcanic, effusive, and explosive (Table 2):

1. Plutonic intrusive volcanism: Batholith and Plutons are large intrusive bodies, of granite composition, which originate from partial melting of the crust. They generally have a length of over 2000 km and a maximum width equal to about 1/10 of the length. They are found in areas of the earth’s crust where rocks have been folded and dislocated to form mountain ranges (orogenic crust), causing the partial melting process of the deepest crustal part of the folded zone. In practice, they constitute the nucleus of the most impressive mountain ranges, such as the Andes (South America): [Achala Batholith, Argentina, Antioquia Batholith, Colombia, Guanambi Batholith, Bahia, Brazil, Parguaza rapakivi granite Batholith, Venezuela and Colombia, Cerro Aspero Batholith, Argentina, Coastal Batholith of Peru, Colangüil Batholith, Argentina, Cordillera Blanca Batholith, Peru, Vicuña Mackenna Batholith, Chile, Elqui-Limarí Batholith, Chile and Argentina, Futrono-Riñihue Batholith, Chile, Coastal Batholith of central Chile, Panguipulli Batholith, Chile, Patagonian Batholith, Chile and Argentina, North Patagonian Batholith, South Patagonian Batholith] and Europe: The Alps: Adamello Batholith. (Figure 11A) (taken from https://en.wikipedia.org/wiki/Batholith). Lopoliths: They are intrusive bodies of smaller dimensions than the batholiths of gabbroic composition and originate from the intratelluric crystallization of basaltic magmas of mantle origin that are localized at the base of the crust (Figure 11B).

2. Sub-volcanic volcanism: Dyke Structure: In the upper crust, from the granitic batholithic bodies you can have magmatic ascents that cool down in conduits of various sizes, giving rise to various types of morphologies of sub-volcanic intrusive bodies that take the name of laccoliths, sills, and dikes. Figure 12 shows a dyke common at the Island of Ponza [30].

3. Effusive volcanism: The low viscous basaltic magma due to its low silica and water content escapes from a superficial magma chamber located in the thin oceanic crust and from a deep fracture with a direct ascent from the mantle on the continental crust (Figure 13). The great lava effusions form the basaltic plateau in a distention environment, such as oceanic and continental rifts (Figure 13). The most important manifestations of effusive volcanism are lava flows. The mobility of the flow depends on the viscosity of the lava: (1) Basaltic lavas have high temperatures (∼ 1100°C) and less silica (<50%), and therefore, they have a low viscosity. Due to these characteristics, they can flow for tens of kilometers. (2) Andesitic lavas have temperatures of 900–800°C and silica (∼ 57%), and therefore, have a high viscosity, which prevents excessive flow so that the lava tends to break as it flows. (3) The rhyolitic lavas have temperatures of about 700°C and silica >60%. They are so viscous that the lava accumulates at the volcanic mouth giving dome shapes or even spire-like extrusions. The lava flow can be defined in different forms—(1) pahoehoe lava: Characteristic surface form of basaltic lava. It is formed by high temperature and low viscosity. Characteristic form of lava poor in silica (Figure 14). (2) AA lava: They are lava morphologies typical of basaltic magmas that give the form of scoria type (Figure 15). They are characterized by a lower temperature and a higher viscosity than the basaltic magmas that give the pahoehoe lavas. Flow lava levée: The AA lava and pahoehoe while flowing build levée for fast cooling of the external parts of the flow (Figure 16). Lava-tube flow: It can happen that even when the surface of a large basaltic flow has solidified, the inner part continues to flow into a channel below the solid surface of the flow. If the lava flows go out from this channel, an empty space is formed that produces a lava tunnel (Figure 17). Block lava flow: A typical form of high viscosity rhyolitic andesitic acid lavas form block lava-like as Paricutin volcano (Mexico) (Figure 18). Cooling structures in subareal lava flows: During the final stages of cooling, the inner part of a lava flow contracts in almost hexagonal-shaped columns acquiring a typical columnar crack (Figure 19). Submarine lava flows: In underwater basaltic flows, the lava cools much faster, and therefore solidifies and forms a bubble of vitreous lava or cushions that temporarily blocks the advance of the flow. After a while, the internal pressure lava breaks the crust, and a new bubble comes out. The process is repeated several times, forming layers of pillows on top of each other (Figure 20).

4. Explosive volcanism: The development of explosive volcanism requires viscous magma with a high content of silica and water. These characteristics are generally acquired by processes of differentiation of basaltic magmas in the magma chamber or by direct melting of the continental crust by basaltic magmas or by the direct mixing with external water. Headings are from Table 2.

Magma
	Basalt	Andesite	Rhyodacite
SiO2	<50%	55–60%	>60%
Viscosity	Low	High	High
Volatile content	Low	High	High

Table 1.

Magma composition.

Magma
	Intrusive volcanism		Extrusive volcanism
	Plutonic	Sub-volcanic	Effusive	Explosive

Table 2.

Type of intrusive/effusive volcanism.

Figure 11.

Scheme of relationship from plutons and Lopolith (author’s collections).

Figure 12.

Dyke at Ponza Island (Italy) [30].

Figure 13.

Lava flow plateau (author’s collection).

Figure 14.

Pahoehoe lava flow (author’s collection).

Figure 15.

Lava flow AA (author’s collection). A solidified flow from the last eruption of Vesuvius (1944 AD).

Figure 16.

Lava flow with levee (author’s collection). Levee grew during a Hawaiian eruption.

Figure 17.

Lavatube flow with a growth stalactite melt (unknown volcano).

Figure 18.

Block lava flow. Modified from Schmincke [14].

Figure 19.

Columnar lava (author’s collection).

Figure 20.

Pillow lava formed under the sea (author’s collection).

7. Explosive volcanism

Explosive volcanism can be classified into two types, as shown in Table 3. The internal and external eruptive dynamics of the volcano consist of two phases: Phase 1: Exsolution of volatiles from the over-saturated magma, and explosive fragmentation of magma with the conversion of thermal energy in kinetic energy. Phase 2: Expulsion of the mixing composed of fragments of solidified magma and gas, and a formation of an eruptive column (Figure 21). The two phases are initially connected and determine the explosive style of the eruption. As follow a diagram of the Classification of explosive eruptions with two volcanological parameters (Explosiveness-Height of the pyroclastic column) (Figure 22). This diagram explains the differences between various types of volcanic eruptions. Diagram is redrawn after Cas and Wright [8].

Types of explosive volcanism
Magmatic volcanism	Hidromagmatic volcanism
Volatiles dissolved in magma are those present at the beginning of its formation.	The magma enters in contact with external water and the magma is enriched in water vapor and other volatiles.
In both cases, if products emitted by an explosive plume are affected by a gravitative deposition, falling deposits form. This type of explosive eruption also takes the name of Fall-Out bedded eruptions. In addition, Pyroclastic density currents [31] can also form.

Table 3.

Types of explosive volcanism.

Figure 21.

Draw that show how magma evolves through a pyroclastic material (exsolution and fragmentation). Modified from Cortini and Scandone [25].

Figure 22.

Diagram of the explosiveness versus the height of the eruptive column. Modified from Cas and Wright [8].

8. Magmatic explosive eruptions that produce fall-out bedded deposits caused by Plinian eruption

To explain such deposits, I have schematically subdivided the magmatic explosive eruptions into two phases and sections—how the magma act from inside to outside volcanoes.

Phase 1: The magmatic column rises with a speed of <0.5 m/s.

The volatiles that dissolves form gas bubbles whose upward movement is slow because it is hindered by the viscosity of the magma. The expansion stops when the ratio between gas and magma is about 3:1 causing fragmentation. At the level of fragmentation, the drastic decrease in viscosity causes the sudden increase in speed that passes from subsonic to supersonic (Figure 23).

Figure 23.

Eruptive Plinian column from the source up to the surface. Modified from Schmincke [14].

Phase 2: The mixture of gases and particles expelled from the crater forms a Plinian eruptive column in which three zones can be recognized, which are as follows:

1. jet expulsion area with initial acceleration, of height 1–2 km

2. high convective area 25–50 km

3. umbrella area (Figure 23)

Figure 24.

(a): Fall-out bedded typical with an enlargement of its constituent (pumice) of a Plinian eruption from Somma-Vesuvius volcano (author’s collection). (b): A model to visualize the spread of the pyroclastic deposits around a volcano. Modified from Cortini and Scandone [25].

A typical fall-out bedded deposit is shown in Figure 24A. This massive fall-out bedded is a characteristic Plinian deposit of the Somma-Vesuvius volcano. The main components are pumice, as seen in Figure 25A [32]. To compute such deposits is not very easy, the measurements of thickness and pumice size called, respectively, isopach and isopleths can be a difficult task [8]. Figure 25B shows a model of isopach calculation seen in three-dimensional space [24]. Each isopach is characteristic on just one thickness that can be followed and extrapolated in the field around a volcano that forms tephra fall-out bedded eruption, many model computations have been published, I just quote some: [33, 34, 35, 36]. Generally, Plinian eruptions are characterized by strong volcanic plumes coupled with vertical columns (Figure 25), although as shown from Figure 25 even a strong volcanic plume can have a characteristic bent-over feature like in a weak plume [36].

Figure 25.

Eruptive columns of Plinian eruptions (taken from the web).

9. Strombolian eruption

Phase 1: The magmatic column rises at a speed of 0.5–1 m/s. The rising speed of the bubbles is greater than that of the magma so very large bubbles form on the surface. At a depth of about 100 m, the fraction of gaseous bubbles reaches 75%, and the magma fragments erupt explosively. The decrease in viscosity produces an increase in speed from subsonic to supersonic (Figure 26).

Figure 26.

Strombolian eruption schematically explained (author’s collection).

Phase 2: The expulsion of materials occurs through a series of explosions with short time intervals (0.1 sec to 1 h). The exit speed of the mixture of gas and fragmented particles is about 200 m/s.

When the explosions occur in short intervals of time, an eruptive column is formed in which the height of the jet does not exceed 200 m and the convective column reaches heights of 5–10 km (Figure 26).

The Strombolian eruption has taken the name from the Stromboli Island part of the Eolian Islands, characterizing such volcanic activity worldwide. The style of the Stromboli eruption has been continuously active in the last 2000 years and it was very important for obsidian trade and as a lighthouse of the Mediterranean Sea. Some examples of Strombolian activity are listed and shown in figures. Eruption column of a Strombolian eruption of Vesuvius, the 1944 AD shown in Figure 27. Scoria layers deposed during a Strombolian activity of Vesuvius in the medieval period are shown in Figure 28 [37].

Figure 27.

Paroxysistic phase of the eruption of 1944 AD (author’s collection). Photo made from allied force during the second world war.

Figure 28.

Medieval deposits from a Vesuvius eruption, a sort of Strombolian activity (author’s collection).

10. Hawaiian eruption

Some eruptions descriptions have been subdivided into phases for the sake of clarity from the inside to outside volcano.

Phase 1: The ascent rate of the bubble-free basaltic magma is >1 m/s. The low viscosity allows the formation of large bubbles that are brought to the surface (Figure 29).

Figure 29.

Hawaian eruption (taken from the web).

Phase 2: The expansion of the large bubbles expels fragments of magma outward from the conduit that forms incandescent lava fountains with a speed of 10–20 m/s reaching heights of 200–500 m (Figure 29). The lava fountains erupt high-temperature scoriae that falling still hot remobilize lava flow without roots (spatter fed lava flow). Lava fountains can also form in craters occupied by a lava lake.

11. Explosive eruptions of the hydromagmatic type that produce fall-out bedded deposits Vulcanian eruption

Vulcanian eruptions are direct explosive eruptions rich in magmatic volatile or any water [38].

Phase 1: The magma rich in volatiles interacts with the groundwater causing rapid vaporization and the relative production of water vapor that increases the explosiveness of the eruption.

Phase 2: The expulsion of the products is characterized through a number of discrete explosions that follow one another in variable intervals between 10 min and 1 h comparable to cannon shots.

The single explosions produce eruptive columns between 5 and 10 km high (Figure 30). A typical example of scoria Vulcanian products is shown in Figure 31.

Figure 30.

Volcanian eruption (author’s collection).

Figure 31.

Scoria deposit caused by a Volcanian eruptions (modified from Schmincke [14]).

12. Surtseyan and Phreatoplinian eruptions

They are explosive eruptions that produce fall deposits deriving from water-magma interaction processes, but pyroclastic density currents are much more abundant [8] (Figure 32A and B). As an example, a Surtseyan eruption is an explosive style of a volcanic eruption that takes place in shallow seas or lakes when rapidly rising and fragmenting hot magma interacts explosively with water and with water-steam-tephra. The eruption style is named after an eruption off the southern coast of Iceland in 1963 that caused the emergence of a new volcanic island, Surtsey (Figure 32B).

Figure 32.

(a): Phreatoplinian eruption; and (b): Surtseyan eruption (modified from Schmincke [14]).

13. Explosive eruption form pyroclastic density currents

These eruptions, unlike the previous ones, are characterized by the development of eruptive dynamics with the main component horizontal sliding on the ground. They are divided into—(a) dense pyroclastic currents (pyroclastic flow), (b) diluted pyroclastic currents (base surge).

Both these eruptive types can be found in association with an explosive eruption from fall or develop independently during special types of explosive eruptions. A typical dense pyroclastic current is shown in Figure 33. A pyroclastic flow is formed by the collapse of an eruptive column (Figure 31). The dense pyroclastic currents can be seen in Figures 32 and 33. Figure 33 also shows the fluidization and turbulence of dense pyroclastic currents. The head of a dense pyroclastic current often is characteristic of an ignimbrite eruption. It can form through a pyroclastic flow formed by the collapse of a volcanic dome (Figure 34). Sometimes, it can be easy to find pyroclastic density currents with a massive structure in Figure 35.

Figure 33.

Pyroclastic flow (modified from Branney and Kokelaar [31]).

Figure 34.

Origins of pyroclastic density currents. (A): short single-surge current derived by momentary collapse from a Plinian eruption column. (B): sustained current is derived from prolonged pyroclastic fountaining. The height of the jet (gas thrust) that feeds the current may vary and is transitional into (C). (C): A sustained current derived from a prolonged low pyroclastic fountaining explosive eruption. (D): current with a single (or multiple) surges derived from lateral blasts initiated by catastrophic decompression of a magmatic and/or hydrothermal system. (E): single-surge current derived from a collapsing lava dome or flow front. Hot rock avalanches generate turbulent density currents. (F): deposit-derived pyroclastic density current caused by gravitational collapse and avalanching of unstable loose ignimbrite. The current may be a single surge or more sustained where the collapse is retrogressive. Most large-volume ignimbrites derive from current types (B) and (C), which may involve periods of quasi-steady flow. Many may include significant components derived from currents of type (F) from Branney and Kokelaar [31]. For this paper, we chose the following three types: Soufriere, Pelee, and Merapi types (modified from Francis [10]).

Figure 35.

Shape of typical pyroclastic dense currents. Pyroclastic dense currents with fluidization and turbulence (modified from Branney and Kokelaar [31]).

14. Diluted pyroclastic currents

Base surge and pyroclastic surge originate from the base of eruptive columns of magmatic and phreatomagmatic eruptions [4, 8, 9, 10]. They move radially and are made up of one turbulent cloud of water vapor and ash at high temperatures (Figures 36 and 37). For all kinds of volcanic forms, we can look to Figure 38 depending on the size of the eruption and the volume emitted. For large eruptions, often after the volcanic activity, there is a collapse of the roof of the magma chamber that brings to a structure called Caldera. A classic example can be the Santorini Caldera, as shown in Figure 39. For a thoughtful explanation of the diverse type of caldera refer to Chris Newhall caldera book [39].

Figure 36.

Head of a pyroclastic dense currents and PDC channeled in a valley. Modified from Schmincke [14].

Figure 37.

Pyroclastic flow deposit with massive structure. Modified from Schmincke [14].

Figure 38.

Morphologies of volcanic structure. Modified from Nemeth and Martin [11].

Figure 39.

Santorini caldera (author’s collection).

15. Conclusion

This chapter has covered all the basis of Volcanology deepens some aspects where recent information has come out from literature. I believe that young students can understand better the subject and they will be much prepared to face other volcanology research. The work goes very easy to read, and the number of Figures is so high that the aspects much close to the Volcanology theory are easily understood. I have added a rich list of references that will help all the people to deal with them and to improve their knowledge in Volcanology.

Acknowledgments

Karoly Nemeth is thanked for the thoughtful review that improved the content of the chapter.